Atlas of the Textural Patterns of Ore Minerals and Metallogenic Processes [Reprint 2010 ed.] 9783110895506, 9783110136395

Table of contents :
Preface
Part I: The Textural Patterns of Ore Minerals and Their Genetic Significance
Chapter 1. Introductory Concepts
Chapter 2. Consideration of the Concepts of Paragenesis, Temperature Determination and Replacement, Based on Edward's Book
Chapter 3. Inductive Versus Deductive Approach in the Interpretation of Textures
Chapter 4. Mantle and Lower Crust Derivatives
Chapter 5. Replacement Patterns and Processes
Chapter 6. Replacement Versus Ex-Solutions
Chapter 7. Symplectites
Chapter 8. Crystalloblastesis
Chapter 9. Zonal Growths
Chapter 10. Epitaxis – Epitactic Growths
Chapter 11. Inclusions
Chapter 12. Colloform Structures (Gel Structures)
Chapter 13. Sphaeroidal Structures and Textures
Chapter 14. Tectonic Effects
Chapter 15. Weathering and Alteration of Ore Minerals
Chapter 16. Leaching, Diffusion and Element Concentration
Chapter 17. Penetrability (Wegsamkeit)
Part II: Consideration of Hypotheses and Theories on Metallogeny (Study Cases)
Chapter 18. Global Tectonics and Metallogeny
Chapter 19. Differentiation and Metallogeny
Chapter 20. Metallogeny Related to Ultrabasics
Chapter 21. Granites/Pegmatites and Related Metallogeny
Chapter 22. Granitization - Anatexis
Chapter 23. Metallogeny Related to Granodiorites-Monzonites
Chapter 24. Metallogeny Related to Porphyries
Chapter 25. Skarns-Pyrometasomatic Metallogeny (and Superimposed Metallogeny)
Chapter 26. Pneumatolytic to Hydrothermal-Hypothermal
Chapter 27. Controversies – Various Aspects of Metallogeny
Chapter 28. The Witwatersrand Controversy
Chapter 29. The Broken Hill Controversy
Chapter 30. Mount Isa Controversy
Chapter 31. The Role of Brines in Metallogeny (The Tennessee Valley-Type of Deposits)
Chapter 32. The Role of Brines and the Mixed Fluids Hypothesis
Chapter 33. Lateral Segregation Processes
Chapter 34. Volcanogenic (Volcano-Sedimentary) Deposits
Chapter 35. Consideration of Certain Aspects of Banded Iron Formations (BIFs) with Emphasis on Precambrian BIFs
Chapter 36. Fluid Inclusions
Chapter 37. Some Aspects of the Role of Fluids in Metamorphogenic Ores
Chapter 38. Sulphur in Metallogeny
Chapter 39. Study Cases of Isotopes and Their Significance in Metallogeny
Chapter 40. Mass-Replacement of Rocks by Ores and Palaeo-Karst-Type Deposits
Chapter 41. Hypogene, Supergene and Oxidation Mineralizations
Chapter 42. Some Aspects of Manganese Mineral Formation - Transformation - Alteration - Oxidation and in General Mn-Mobilization/Remobilization
Chapter 43. The Significance of Leaching and Diffusion Processes in Ore Formation
Chapter 44. Redistribution – Mobilization – Remobilization
Chapter 45. Zonal Distribution of Elements and Minerals
Chapter 46. Source and Recipient Geoenvironments of Mineralization
Part III: On the Distribution of Elements and Ore Parageneses. The Empirical Laws of Element Segregation-Concentration in Ores
Chapter 47. The Empirical "Laws" of Element Segregation/Crystallochemistry/Isotope Chemistry Versus Genesis of Ores – State of the Art
Chapter 48. Segregation of Elements in Accordance with Their Interrelationships to Form Mineral Association-Parageneses
Chapter 49. Common (Joint) Segregation of Elements
Chapter 50. Hydrothermal and Pegmatitic Element Segregation to Form U-Parageneses
Chapter 51. Superimposed Paragenesis (Element Segregation/Distribution Processes)
Chapter 52. Ti, V, Cr – Their Interrelationships and Antipathies
Chapter 53. The Te, Se, Bi, Au, Ag Element Segregation/Distribution (in Paragenetic Associations)
Chapter 54. Realgar, Orpiment – Cinnabar – Metacinnabar Parageneses
Chapter 55. A Special Case of Non-Ferrous Metal Mineralization in Evaporites
Chapter 56. The Segregation (Distribution) of Sn, Mo and W to Form Concentrations or Ore Deposits
Chapter 57. Special Cases of Element Segregation/Distribution
Chapter 58. Element Segregation/Distribution in the Manganese Parageneses
Chapter 59. Trace Elements in Sulfides (Compatible with a Joint Segregation of Elements in Accordance with the Empirical "Laws" of Element Interrelations)
Chapter 60. Study Cases of Agents of Metal Transportation
Chapter 61. Goldschmidt's 'Laws of Element Distribution' and the Empirical "Laws" of Interrelated Element Segregation (Metallic Element Concentration)
Chapter 62. Geoenvironment – Mobilization – Remobilization (Redistribution of Elements)
Chapter 63. Conclusions of Part III
Illustrations
References
Author Index
Subject Index to the Text Part
Subject Index to the Illustrations

Citation preview

Atlas of the Textural Patterns of Ore Minerals and Metallogenic Processes

Atlas of the Textural Patterns of Ore Minerals and Metallogenic Processes Stylianos-Savvas P. Augustithis

W DE

G

de Gruyter · Berlin · New York 1995

Professor Dr. rer. nat. Stylianos-Savvas P. Augustithis Foreign Fellow of the Academy of Natural Sciences of the Russian Federation Laboratory of Textural Analysis, Athens, Greece Laboratory of Textural Analysis, China University of Geosciences, Beijing, China International Scientific Committee, Member of the Laboratory of Crust-Mantle Constitution, Recycling and Dynamics, Ministry of Geology and Mineral Resources, Wuhan, China

With 926 figures and 10 tables

Library of Congress

Cataloging-in-Publication-Data

Augustithis, S. S. Atlas of the textural patterns of ore minerals and metallogenic processes / Stylianos-Savvas P. Augustithis. p. cm. Includes bibliographical references a n d index. ISBN 3-11-013639-2 l . O r e s . 2. Petrofabric analysis. I. Title. QE390.A88 1994 669-dc20 94-41197

CI Ρ

Die Deutsche

Bibliothek

—

Cataloging-in-Publication-Data

Augustithis, Stylianos Savvas: Atlas of the textual patterns of ore minerals and metallogenic processes / Stylianos-Savvas P. Augustithis. — Berlin ; New York : de Gruyter, 1995 ISBN 3-11-013639-2

® Printed on acid-free paper which falls within the guidelines of the ANSI to ensure permanence and durability. © Copyright 1995 by Walter de Gruyter & Co., D-10785 Berlin All rights reserved, including those of translation into foreign languages. No part of this book may be reproduced or transmitted in any form or by any means, electronic or mechanical, including photocopy, recording or any information storage and retrieval system, without permission in writing from the publisher. Printing: Karl Gerike GmbH, Berlin. - Binding: Lüderitz & Bauer GmbH, Berlin. - Cover Design: Hansbernd Lindemann, Berlin. Printed in Germany.

Preface

More than thirty years ago, microprobe analysis introduced the possibility of analyzing quantitatively in situ small surfaces of polished sections, rendering obsolete many of the optical measurements and other data that had been presented in many standard ore-mineralogy textbooks, and thus a new approach to the subject was called for. In addition, it should also be pointed out that whereas many excellent books are available on the mineralogy of ore minerals (see page 218) and a thorough treatment of mineral species is presented in their special sections, it is believed that as microprobe analytical facilities develop and are widely applied, the identification and compositional variation of ore minerals will become more and more a matter of microprobe analysis and x-ray studies. In contrast to the rather "stereotyped" treatment of textures in conjunction with ore minerals, the author believes that a comparative study (i. e., by the application of the principles of comparative anatomy) of textural patterns is also useful, particularly in following the development and evolution of characteristic textural patterns of ore minerals and thus new perspectives in the understanding of the ore genesis are opened up. This is one of the aims of the present effort. The volume presents extensive studies of specific ore occurrences and incorporates geochemical and isotopic studies, thus presenting a most sophisticated and up-to-date treatment of many problems of ore mineralogy and ore geochemistry. Part I (The Textural Patterns of Ore Minerals and Their Genetic Significance) attempts to present the most common and significant textural patterns of ore minerals and their genetic significance. Particular emphasis is given to the presentation of replacements, symplectic intergrowths and generally textures resulting due to the great reactiveness of ore minerals (see List of Contents). In addition, textural analysis introduces novel interpretations concerning many oremineral intergrowths which have been differently interpreted up to now. The author preferred the presentation of the ore-textures in black and white half-tone photographs and the labeling of the depicted minerals to the use of coloured photographs, (one of the reasons being that it is not always easy to reproduce the exact colours of an ore mineral, under reflected light, by a printed coloured photomicrograph). Furthermore, Part I is based on the polished-section collection of the Laboratory of Textural Analysis, which incorporates polished-sections made in the Mineralogical-Petrographical Institute of Heidelberg University, where the author was a research

student under the guidance of Professor Ramdohr in the years 1954-55 and later in 1961-62 as a post-doctorate researcher. However, most of the samples were collected by the author over a period of 40 years on many field trips, visits to mines and expeditions in many parts of the world. Many samples made available to the author from other collections and sources are also included. The photomicrographs and the interpretations presented [with the exception of Figs. 263, 567(b), 613, 622, 707 and 710, which are based on observations by Ramdohr (I960)] represent the original work of the author which was carried out at Heidelberg University (1954-55 and 1961-62), at the Ministry of Mines (Ethiopian Government) in the period 1956-61 and 1963-67, in the National Technical University of Athens, Greece, in the period 1969-84 (when the author was Professor and Head of the Department of Mineralogy, Petrography and Geology) and during the last decade at the Laboratory of Textural Analysis (Athens, Greece). The author has dedicated Part I to the memory of his teacher Paul Ramdohr, without implying that his interpretations would have been approved by him. Part II (Consideration of Hypotheses and Theories on Metallogeny - Study Cases) includes some of the most significant interpretations, hypotheses and theories which have been presented in the standard international literature, and in papers, abstracts, translations and personal communications made available to the author by a great number of colleagues, each of whom he thanks. He has tried to record some of the hypotheses and theories that have dominated metallogeny in the last half century, where the roots of most of the modern theories are to be traced, as well as their unfolding with time, and has not hesitated to present his own personal interpretations or hypotheses. In his attempt to implement this task, the author acknowledges his cooperation with Elsevier Publishing Company, in his position as a member of the Editorial Board of Chemical Geology during the period 1964-81 and also with Theophrastus Publications where he acted as Scientific Editor on a number of international volumes on economic geology. Furthermore he is indebted to ICSOBA (International Committee for the Study of Bauxites, Alumina and Aluminium), as President of which, in the years 1978-83, he gained valuable experience particularly concerning the weathering mantel of the earth. Also, most valuable experience has proved to be his period as Councillor for Africa in the IAGOD in the years 1966-68 concerning African ore V

occurrences. Needless to say that it is a very difficult task to present the plethora of diverse and often contradictory hypotheses and theories put forward in the last decades, and the author did not hesitate to make extensive reference to the literature and use quotations. Acknowledging the work of colleagues either mentioned or quoted (in addition to the full acknowledgment made by the pertinent references), the author has dedicated Part II to "those who toil the hard soil", clearly expressing his gratitude for their immense contribution. Part III (On the Distribution of Elements and OreParageneses - The Empirical "Laws" of Element Segregation-Concentration in Ores) is a reconsideration of the "laws" of element distribution (respectively segregation of metallic elements in mineral paragenesis) and is rather based on the empirical "laws" of element interrelationships, according to the periodic system. Whilst acknowledging the great contribution of V. M. Goldschmidt and his followers on the subject of element distribution, he has dedicated Part III to Dimitri

VI

Ivanovitch Mendelejeff and Lothar Meyer, and to their more modern followers, whose work has been of fundamental importance to the author in pointing out that the interrelationship of elements in accordance with the empirical "laws" of the periodic system was a main (though not the exclusive) factor for their common segregation in mineral paragenesis. Extensive consideration of the other factors involved in the transportation, derivation and segregation of elements to build mineral parageneses, is also included in Part III. The significance of the interrelationships of elements in accordance with the empirical laws of the periodic system for the common segregation of relatively rare elements to build mineral parageneses was first supported by the author in 1964 ("Geochemical and oremicroscopic studies of hydrothermal and pegmatitic primary uranium parageneses") and has been elaborated since then. Most up-to-date isotope studies and information on specific elements and their derivation have been incorporated and a great number of mineral parageneses (study cases) are considered.

Contents

Preface

V

Part I: The Textural Patterns of Ore Minerals and Their Genetic Significance Chapter 1 Introductory Concepts Chapter 2 Consideration of the Concepts of Paragenesis, Temperature Determination and Replacement, Based on Edward's Book Chapter 3 Inductive Versus Deductive Approach in the Interpretation of Textures Chapter 4 Mantle and Lower Crust Derivatives Chapter 5 Replacement Patterns and Processes Chapter 6 Replacement Versus Ex-Solutions

Chapter 14 Tectonic Effects

94

Chapter 15 Weathering and Alteration of Ore Minerals . . .

98

Chapter 16Diffusion and Element Leaching, Concentration

103

Chapter 17 Penetrability (Wegsamkeit)

105

1

8

Part Π: Consideration of Hypotheses and Theories on Metallogeny (Study Cases) 14 Chapter 18 Global Tectonics and Metallogeny

107

Chapter 19 Differentiation and Metallogeny

Ill

Chapter 20 Metallogeny Related to Ultrabasics

114

17

23

41 Chapter 21

Chapter 7 Symplectites Chapter 8 Crystalloblastesis

52

Granites/Pegmatites and Related Metallogeny . . 118 Chapter 22 - Anatexis Granitization

123

126

64

Chapter 9 Zonal Growths

68

Chapter 23 Metallogeny Related to GranodioritesMonzonites

Chapter 10 Epitaxis - Epitactic Growths

71

Chapter 24 Metallogeny Related to Porphyries

127

Chapter 11 Inclusions

74

Chapter 25 Skarns-Pyrometasomatic Metallogeny (and Superimposed Metallogeny)

129

Chapter 12 Colloform Structures (Gel Structures)

77

Chapter 26 Pneumatolytic to Hydrothermal-Hypothermal . . 139

Chapter 13 Sphaeroidal Structures and Textures

85

Chapter 27 Controversies - Various Aspects of Metallogeny

143 VII

Chapter 28 The Witwatersrand Controversy

Chapter 44 157 Redistribution - Mobilization - Remobilization . 211

Chapter 29 The Broken Hill Controversy Chapter 30 Mount Isa Controversy

159

Chapter 45 Zonal Distribution of Elements and Minerals . . 213

161 Chapter 46

Chapter 31 The Role of Brines in Metallogeny (The Tennessee Valley-Type of Deposits)

163

Chapter 32 The Role of Brines and the Mixed Fluids Hypothesis

165

Chapter 33 Lateral Segregation Processes

167

Source and Recipient Geoenvironments of Mineralization Part ΙΠ: On the Distribution of Elements and Ore Parageneses. The Empirical Laws of Element Segregation-Concentration in Ores

Chapter 47 The Empirical "Laws" of Element Segregation/Crystallochemistry/Isotope Chemistry Versus Genesis of Ores - State of the A r t . . . . 219

Volcanogenic (Volcano-Sedimentary) Deposits . 169

Chapter 48 Segregation of Elements in Accordance with Their Interrelationships to Form Mineral Association-Parageneses

Chapter 35

Chapter 49

Chapter 34

Consideration of Certain Aspects of Banded Iron Formations (BIFs) with Emphasis on Precambrian BIFs Chapter 36 Fluid Inclusions

Common (Joint) Segregation of E l e m e n t s . . . .

180

Chapter 38 Sulphur in Metallogeny

188

191

Chapter 40 Mass-Replacement of Rocks by Ores and PalaeoKarst-Type Deposits 195

199

Chapter 42 Some Aspects of Manganese Mineral Formation Transformation - Alteration - Oxidation and in General M n - M o b i l i z a t i o n / R e m o b i l i z a t i o n . . . . 204 Chapter 43 The Significance of Leaching and Diffusion Processes in Ore Formation VIII

225

Chapter 50

183

Chapter 41 Hypogene, Supergene and Oxidation Mineralizations

221

173

Chapter 37 Some Aspects of the Role of Fluids in Metamorphogenic Ores

Chapter 39 Study Cases of Isotopes and Their Significance in Metallogeny

215

208

Hydrothermal and Pegmatitic Element Segregation to Form U-Parageneses Chapter 51 Superimposed Paragenesis (Element Segregation/Distribution Processes)

230

237

Chapter 52 Ti, V, Cr - Their Interrelationships and Antipathies

239

Chapter 53 The Te, Se, Bi, Au, Ag Element Segregation/Distribution (in Paragenetic Associations)

242

Chapter 54 Realgar, Orpiment - Cinnabar - Metacinnabar Parageneses

244

Chapter 55 A Special Case of Non-Ferrous Metal Mineralization in Evaporites

246

Chapter 56 The Segregation (Distribution) of Sn, Mo and W to Form Concentrations or Ore Deposits . . . .

247

Chapter 57 Special Cases of Element Segregation/Distribution

251

Chapter 58 Element Segregation/Distribution in the Manganese Parageneses Chapter 59 Trace Elements in Sulfides (Compatible with a Joint Segregation of Elements in Accordance with the Empirical "Laws" of Element Interrelations)

255

Chapter 62 Geoenvironment - Mobilization - Immobilization (Redistribution of Elements)

267

Chapter 63 Conclusions of Part III

273

Illustrations

275

Chapter 60

References

589

Study Cases of Agents of Metal Transportation . 260

Author Index

617

Chapter 61

Subject Index to the Text Part

625

Subject Index to the Illustrations

654

Goldschmidt's 'Laws of Element Distribution' and the Empirical "Laws" of Interrelated Element Segregation (Metallic Element Concentration)

258

262

IX

Parti

The Textural Patterns of Ore Minerals and Their Genetic Significance

Parti, In memoriam Paul Ramdohr

Chapter 1

Introductory Concepts

(a) Element Segregation and

Paragenesis

The understanding of the concentration of elements into paragenetic associations, which might be genetically related or superimposed and reactive paragenetic associations, is possible by examining the interrelationships of the elements which segregate under the operation of complex processes. The common segregation of relatively rare elements of the earth's crust into paragenetic associations (i. e. deposits) has been explained by the author, Augustithis (1964, 1967, 1979), on the basis of the interrelationships of these elements in accordance with the empirical laws of the periodic table. Also of importance is the grouping together of these elements into groups that exhibit interrelationships according to the empirical laws of the periodic system. Such groups, recognized by Augustithis (1964), correspond to hydrothermal paragenetic associations, group A = U, (Th), Pb, As, Sb, Bi, Se, Te, (Mo, Sn) and group Β = Fe, Ni, Co, Cu, Zn, Ag, Au and for the pegmatitic uranium parageneses group C = U, (Th), Zr, Hf, Ti, Nb, Ta1 and the REE. It is interesting that in considering the metallogeny of different granitic types in South East Asia (China), Xu Keqin et al. (1982) recognized that certain granitic rock types are characterized by groups of elements of crustal derivation in contrast to other granitic rock types that are characterized by a group of metallogenic elements of mantle derivation. Here again a relationship exists between the elements of the groups in accordance with their interrelationships in the periodic table. It is interesting to note that groups A and Β recognized by Augustithis (1964) also indicate a derivation difference in the sense that the elements of group Β (Fe, Ni, Co, Cu, Zn, Ag, and Au) are mainly derivatives or remobilized derivatives of mantle (or of basaltic derivation by leaching) whereas elements of group A and C (recognized by Augustithis, 1964) are mainly of crustal derivation. The above considerations can furnish an interpretation as to the observed types of metallogeny related to different granites or respectively to granites by graniti1

Mo, W and Sn are also lithophile elements and often occur in pegmatites (see Chapters 56 and 57).

zation involving mainly crust derived materials and to granitization products involving also mantle derived or remobilized materials (mantle derived granites of Xu Keqin which in reality might be granitization products involving mantle derived material where mantle derivation is not a prerequisite. Drescher-Kaden (1961, 1969) and Artus (1959) have already presented such granitic rock types from the Alps). In contradistinction to the recognition by Xu Keqin et al. (1982) of granites with distinct metallogeny, i. e., granites characterized by Cu, Co, Ni, Cr and granites of metallogeny (derivative of crustal mobilization) Augustithis (1982) has presented cases of apogranites containing elements (metallogeny) of crustal derivation and elements of remobilized mantle derivation. Our comprehension of the occurring together of elements that are nonabundant in the earth's crust in concentrations (deposits) is mainly due to segregation of these elements and to the laws or empirical laws governing their common segregation. In particular, in the case of the hydrothermal vein deposits which are supposed to be derivatives of granitic intrusions, it is interesting to note that elements making up the parageneses belong mainly to the groups A and B. As pointed out, most of the elements of group Β could be considered mantle derivatives or recycled mantle derivatives. In contradistinction the elements of group A, as well as the elements of group C (characteristic of the pegmatitic parageneses), consist mainly of crustalderivation elements2. To understand these interrelations the following alternative interpretations are tentatively suggested: (i) The metallic elements building the hydrothermal perigranitic vein deposits are derivatives of initial progenitor materials which have been granitized and consist of crustal and mantle (recycled) materials. As a consequence of this, both crustal and mantle metallic elements could segregate to form hydrothermal veins. (ii) The granite intrusion mobilized lower crustmantle material due either to deep seated granitization or due to derivation from depth (mantle derived granites with an increase in Cu, Co, Ni trace element content).

2

Isotope geochemistry is a useful tool in determining the derivation of elements (see Chapter 39).

1

(iii) The derivation of elements making up so-called hydrothermal vein deposits are derivatives of the granitized material and of the perigranitic country rock material, initially by lateral segregation and solution transportation. Thus, the paragenesis of metallic vein deposits depends on the presence and the availability of the metallic elements in the Bereich of the granitization field and on the lateral or mobilization of elements peripheral to the granitization Bereich, i. e., the perigranitic country rock environment. Depending on the availability of metallic elements which depend on the distribution of metallic elements in initial sediments or metasediments and other rock types subjected to granitization processes and on segregation processes that are responsible for the concentration of the metallic elements into vein deposits, we have the different hydrothermal paragenesis. Within the segregation processes of fundamental significance is the common segregation of elements which depends on the interrelationships of available elements on the empirical laws of the periodic system. As already suggested by Augustithis in 1964, the recognition of the A, Β and C metallic element groups making up hydrothermal and pegmatitic mineral pangeneses, depends mainly on the empirical laws of the periodic system which govern the chemical interrelationships of the elements. In addition to the common segregation of elements for the formation of the mineral paragenesis are the potentialities of crystal structure formation which depends on the interrelationships between crystal lattice, atomic radii, valencies, etc. The segregation of elements is brought about by solutions and the supply of solutions could extend throughout time, in the sense that not all the common segregated elements are simultaneously in solution, i. e., supply of elements in solution to the spaces (Bereich) in which crystallization takes place, can occur at different periods of time. Under sufficient concentration of elements in the solution and under appropriate physico-chemical conditions prevailing in the environment or microenvironment in which crystallization takes place, the elements in solution crystallize out and a crystal lattice is formed with the potentiality of incorporating elements in solid-solution or incorporate elements by substitution of cations. Which mineral will prevail or survive depends on the field of stability of each crystallization, on the reactions and interreactions between the minerals formed and their reaction to the country rock; in addition, a crystal formed in the space where crystallization takes place is subject to dissolution with the changing of the composition of the solutions which can bring about substitution or partial substitution of a mineral formed by subsequent crystallization from the same solutions as the composition of the solutions changes. The supply of additional solutions of different composition in the same space (Bereich of crystallization) can also bring about disso2

lution of already-formed minerals and their partial or complete replacement. The understanding of a paragenetic mineral association inevitably involves consideration of the factors and conditions mentioned above. It is hoped that crystallization sequence of ore mineral assemblages could be deciphered on the basis of textural analysis, as it can be done on the basis of comparative anatomy of textures. Considering the abundance and distribution of elements in the crust mainly on the basis of abundancetables of elements, Goldschmidt (1954) and Wederpohl (1967), and furthermore considering their distribution on the basis of the Verteilungsgesetz der Elemente, the question arises of what the initial distribution of elements was in the star-crust of the earth (the lithic era of the earth). Perhaps a possible extrapolation with the abundance and distribution of elements in extraterrestrial bodies (e. g., the moon) might be of interest. In spite of the fact that our knowledge of the distribution of elements in the moon's crust is very limited, it can be postulated, on the basis of the predominance of gabbroic-peridotitic rock types, that magnesium-rich rocks predominate. Comparative studies show a relatively greater abundance of Ti in lunar rocks when compared to terrestrial basalts. If we assume that the initial earth's crust was comparable to the moon's surface, two fundamental questions arise: (i) Under which processes did the initial gabbroicperidotitic crust change into earth's crust in which granitic and gneissic rock types are very abundant? (ii) Can we relate the changes from a gabbroic-peridotitic crust to a more acid-type with an availability of Mg source which was considered necessary by V. M. Goldschmidt - (as was communicated to the author by Goldschmidt's friend Drescher-Kaden)? In 1973, Augustithis in his Geological spiral, starting from a peridotitic-gabbroic initial earth crust had already suggested the following: In accordance with the concept of the geological spiral the following developments can be suggested: (i) The initial star crust of the earth in extrapolation with the moon's crust, may have been peridotitic-gabbroic in composition, comparable to the gabbroic-peridotitic rocks of the moon. (ii) The formation of atmosphere-hydrosphere around the earth's lithosphere can be seen as the commencement of the geological history of the earth, the beginning of geomorphological and geological cycles and the concurrent operation of exogenetic and endogenetic processes. (iii) The peridotitic-gabbroic initial earth's crust, due to chemical weathering, grado-separation of elements, differential leaching and geochemical mobilization of elements may have produced acid and intermediate residuals. The sum total of these processes perhaps re-

suited in the formation of acid to intermediate anchisediments. ( i v ) The anchi-sediments in accordance with the principles of uniformitarianism would be subjected to the processes of the geological cycle: denudation, aggregation, metamorphism, ultrametamorphism (granitization, gabbroitization and ultrabasic rock formation). Parallel to and concurrently with the geological cycle, while also as a part of it in the realm of geodynamic events, outflows of huge quantities of basaltic materials took place, products of shearing and fusion of the earth's upper mantle, as well as mantle diapirism. Thus, the geological spiral provides a working hypothesis for extensive element mobilization and recycling. Water, too, has played a fundamental part in the formation of the earth's crust, bearing in mind its role in the unfolding of the geological spiral and the phreatic cycles of Vernadsky and Grigorieff s consideration of the role of water in the formation of the crust. Due to the unfolding of the geological spiral an upper crust layer thus differentiated through derivatives of the initial lithic era's crust in which the lithophile elements were abundant (in addition to Al, Si and the metallic A and C elements, already mentioned). In contradistinction to the M g element of the initial gabbroic-peridotitic crust, or due to the upper mantle diapirism, metallic elements of the Β group are associated.

primarily from depth by ascending solutions and been deposited in spaces (veins, pore spaces, etc.), the supergene were considered to be secondary mobilizations due to the weathering of primary metallic concentrations which moved from the surface downwards and enriched the alteration zone, forming the supergene impregnation zone. Particularly for supergene mineral assemblages and deposits, it is believed that mobilization and even remobilization of the elements by low temperature solutions has taken place. Also at the beginning of this century it was proposed that ore deposits (space fillings, veins, pore spaces, etc.) were due to lateral secretion or lateral segregation by solutions in the realm of metamorphism or by leaching. The idea of lateral secretion (segregation) unfolded further along the following axes of thinking. (i) Lateral secretion (segregation) occurs during metamorphism-ultrametamorphism. A most elucidating example of space and vein filling by lateral secretion was put forward by Drescher-Kaden in 1969 who

erals were referred to as being hypogene or supergene in origin meant that metallic element mobilizations by solutions were considered to be not only the mechanism of transportation of the metallic elements but also indicated their derivation.

demonstrated that quartz veins and space fillings in folded chlorite schist were originated from lateral secretion due to the subsequent tectonic effects on the chlorite schist. Similarly Augustithis (1985) explained cases of calcite veins in marbles as being lateral secretions of the marbles in which the carbonate veins occur. Both Drescher-Kaden (1948, 1969 and 1974) and Augustithis (1962, 1973) considered pegmatitic and aplitic veins as exudation products of granites and gneisses under the process of ultrametamorphism. Already in 1908, Holmquist had considered pegmatites as metasomatic products due to granitization an interpretation which could be considered as pioneering to the exudation hypothesis of pegmatites and aplites. The often observed telescoping of pegmatitic veins in depth is perhaps a corollary to the exudation interpretation. One of the most significant cases of lateral secretion was the novel interpretation put forward by Goldschmidt and his friend Drescher (see Goldschmidt, 1954) on the derivation of tourmaline in granites and pegmatitic veins by lateral secretion and mobilization of Β from the adjacent sediments, in the sense that Β was picked up by the granite from the adjacent sediments. The following characteristic quotation is from Goldschmidt, 1954: "An investigation by the author and his friend Prof. F. K. Drescher-Kaden, then in Clausthal, on the boron in shales, hornfels and granites of the Harzburg district, showed that the amount of boron decreases from the shales to the hornfels of the contact zone and the granite takes up boron from the sediments - quite a reversal from established views in petrology." Comparable studies of the Piona pegmatites in the Alps by Augustithis (1973) showed that the tourmaline in the pegmatites was perpendicular to the contact wall of the pegmatite with the gneiss and the boron was a lateral derivative from the gneiss-schist, an exudation product of which was also the pegmatite.

Whereas hypogene ore mineral associations (concentrations) were considered to have originated

(ii) In contradiction to the lateral segregation during metamorphism-ultrametamorphism lateral segregation

Depending on the composition of the geoenvironment in which lateral and selective segregation processes would operate, the composition of vein or pegmatitic deposits will be characterized by the predominance of A and/or C groups or of A and Β groups. As mentioned, in many granitic metallogenic regions, elements of the A and C groups predominate where in other cases elements of the Β group (porphyry type of metallogeny) are most predominant. An exceptional case is the metallogeny of the Abu Dabbab (Egypt) apogranite in which, in addition to the C group elements (granitophile), a superimposed metallogeny derivative due to leaching of intercontinental rift basalts is also present resulting in a reactive paragenesis, see Augustithis (1982).

(b) Remobilization and Recycling of Metallic Elements By the beginning of the twentieth century it was realized that mobilization of elements is a process of great metallogenic significance. The very fact that ore min-

3

also takes place under considerably lower temperature conditions. Augustithis has described cases of magnesite vein formation due to lateral segregation in serpentinized dunities from Yubdo, Ethiopia. All transitions were shown of marginal alteration of forsteritic olivine to magnesite: mobilization of olivine alteration (magnesite) to magnesite veinlets and finally magnesite vein formation due to the unification of the veinlets. According to Augustithis (1965), weathering of olivines due to hydration, with the participation of carbonic acid (C0 2 ) of the air, results in the formation of magnesite. The formation of magnesite is seen as part of the gradoseparation processes suggested by Pieruccini (1962). In the decade of the seventies, several vein deposits were explained by the Canadian School of Petrology as the result of lateral leaching and mobilization by lower temperature conditions of elements from basalts which were supposed to be buried in rifts. Tooms (1976) has suggested that the manganese vein deposits adjacent to the Great Rift were due to element leaching and mobilization from basaltic flows in the Great Rift. Also, Augustithis (1982) explained the paragenesis of the Abu Dabbab apogranites as the result and reaction (superimposed mineral paragenesis) of the granitophile C group element mineral paragenesis characteristic of the granite-gneissic Precambrian rocks of the region and of manganese mobilized by leaching of basalts buried in the Great Rift. In this case the manganese presence in the apogranite is due to lateral secretion due to leaching and mobilization of manganese from basalts most probably buried in the rift, i. e:, a lateral segregation process. Recently, Vgenopoulos (1960) has described magnetite, rutile and ilmenite in quartz veins as being the result of lateral secretion and leaching of adjacent rocks.

(c) Mobilization ments (General

and Remobilization Concepts)

of Ele-

The wide concept of metasomatism includes many special processes that have attained separate status within the broad concept of the alteration of rocks. Such processes are differential leaching, palingenetic associations, element mobilization, recycling of elements, mobilization and remobilization. All the above mentioned processes are contributing factors in the formation of ore mineral parageneses and metallic mineral concentrations. In the present chapter only some selected cases will be presented, not to exhaust the subject, but simply to furnish some examples. (i) Palingenic veinlets of cassiterite, due to the mobilization or remobilization of Sn, most probably hypothermal or contact-zone mineral formation, are mobilized and deposited as cassiterite transecting colloform pitchblende. Cases of this have been presented by Ramdohr (1960) and by Augustithis (1964). The veinlets of cassiterite transecting the colloform pitchblende 4

could be considered as remobilization in the "epithermal stage" of mineralization. (ii) An interesting case of element mobilization and its resultant textures is the often described mobilization of interspersed lead (galena), radiogenic in origin, into synaeresis cracks of the pitchblende in which radiogenic lead occurs as fine interspersed granules. The Pb mobilization in this case was mainly, but not necessarily, restricted within the same mineral. Lead, particularly as galena, exhibits an impressive mobilization capacity either as veinlets or as replacements of a wide gamut of previously formed minerals. In this connection, is the work of Köppel and Scholl (1983) of significance. These authors presented cases of remobilized lead as it was evidenced on the basis of lead isotope investigations in their paper entitled "Bleiisotope und Remobilisation von Erzlagerstätten". (iii) A most interesting case of element mobilization is described by Ottemann and Augustithis (1967) where PGE in the dunite ring complex of Yubdo, Western Ethiopia, (ultrabasic ring complex) occurring as sperrylite in chromospinels which survived birbiritization, was altered and mobilized in the Iateritic covers of dunite and birbirite. Platinum nuggets have been formed by element segregation (accretion), often having altered chromite as the nucleus. Many platinoid group elements have been segregated within the ferroplatin nuggets and interesting but complex textures have been formed, described originally by Ottemann and Augustithis (1967). (iv) In contrast to the cases considered so far, where element mobilization is understandable due to the element kinetics, chromites which are generally considered to be very resistant to alteration processes, are also subjected to alteration and weathering processes, see Augustithis (1962, 1962, 1980) and as a result of element leaching and element mobilization of initial chromites, cases are described by Hutton (1942) of small chromite grains found associated with fuchsite. Most probably remobilization of Cr is responsible for this mineral formation.

(d) Tectonic Mobilization - Polyphase zation

Mobili-

In contrast to the mobilization of the elements by solutions where the main medium of element transportation is water and the processes involved were mainly solubility processes, special cases may exist where mechanical-tectonic mobilization takes place where the medium is a solid with greatly increased plasticity. In this case, in contrast to element mobilization in solution, en masse transportation of fragments takes place in the greatly plastic medium. It should be mentioned though that the same material which can be affected and dissolved in water, can under different circumstances and geoenvironmental conditions be mechanically transported in a plastic medium.

Studies on forsterite crystal plasticity have shown that forsterite-rich mantle dunites can provide the plastic medium in which chromite Schlieren bodies are fractured and mobilized as boutinage structures in a forsteric plastic medium. Perhaps the well-known example of interbanded chromite and anorthosite of the Dwars River, Transvaal, is another example of chromite plastically-tectonically mobilized in or with anorthosite (see Chapter 14). As we have already mentioned in the case of element mobilization by solutions, Cr from an initial chromite might be geochemically mobilized in solution and could be reprecipitated as chromite grains associated with fuchsite. Mantle formed chromite can be further remobilized as xenocryst in basalts which transverse an ultrabasic complex. Augustithis (1965) has described chromite xenocrysts (grain-sized inclusion) in the Yubdo area of Western Ethiopia. Here again we have a mobilization or remobilization of chromite in the melt phase whereby a margin of reaction in the xenocryst is noticeable; this is not a decoloration margin but actually magnetite is formed as the result of a chromospinel reacting with the basaltic melt. Chromite thus provides us with examples of element mobilization by water solutions, mobilization of chromites in melts and chromite mobilization tectono-mechanically in a crystalloplatic medium.

(e) Derivation Environments (Main Types of Source Environments) Source environment and geoenvironmental conditions prevail in the petrogenetic "field" (Bereich) of metallic elements derivation. (i) According to some estimates, the earth's core is considered to consist of Ni and Fe, a possible differentiate from the molten star state or phase - which is also compatible with the big bang hypothesis. (ii) Mantle consists predominantly of dunitic and peridotitic types with a dominant mineral phase of forsterite rich olivines. The magnesium element is predominant. Furthermore, forsterite predominant in olivines incorporates fyalitic molecules, and in addition there is substitution and incorporation of Ni in the forsterite lattice supposedly replacing the Mg cation. As already mentioned, there is a strong interrelationship of Ni, Fe and Co, and the platinum group of elements (as belonging to the same group of the periodic system). (iii) The protolytic (gabbroic) parental layer of basalts is basic/ultrabasic in composition and thus is a parental geoenvironment in addition to the mantle characteristic elements also for Ti. In this respect, of significance is the antipathy or incompatibility of Cr and Ti. (iv) The lower crust is an anorthosite-rich environment in which plastically mobilized chromite can oc-

cur and the anorthosite chromite banding of the Dwars river can take place. (v) Lithophile elements, mainly the C group of elements, are often associated with granite-gneisses and with their pegmatitic exudation products. The lithophile group of elements are treated extensively in Chapters 50, 56 and 57. An example of an environment of lithophile element generation-mobilization is given by Yuan Kuirong and Xang Xinyi (1985) in their contribution entitled "The tectonic environment and tin metallogeny of granites in South China". The following synoptical extract is quoted from their work: "Endogenic tin metallogeny of the granites in South China is closely related to the migration and tectonic differentiation of the post-Proterozoic tectonic environment of the granites. With the southwestward lateral migration of orogenic granites on the initially tin enriched Proterozoic continental crust, the "reverse overlap" of the north-western anorogenic granites upon the older granite zone appeared. It thereby caused the enrichment by differentiation of crustal source syntexis granites in the locally thickened zones and the differential sedimentation controlled by the structurally superimposed unwrapping region. This engendered the decisive conditions in the formation of the world's largest granite-endogenic tin metallogenic zone. The ore source zone formed by the differential sedimentation-enrichment process and its superimposition with the differentiation-enrichment process of S-type metallogenic granites represents the basic metallogenic environment model of endogenic tin in granites." (vi) The weathering cover environment (a complex weathering mantle of many diverse rock types) and often the result of extensive element recycling, is itself of great significance as a source of secondarily derived ores.

( f ) Recipient

Geoenvironments

It has already been mentioned that the mantle geoenvironment is characterized by the abundance of Mg and as in the predominant mineral forsteric olivine, Mg can be substituted by Ni (also the interrelationship of the elements of the same group of the periodic table, namely Ni, Fe, Co and the PGE as belonging to the subgroup of the VIII family, is responsible that these metallic elements are abundant in the mantle environment). As a consequence of the above-mentioned interrelationships, it is no surprise that these metallic elements do occur as ore concentrations in the mantle and its derivatives that are moved or mobilized, in one way or another in the upper crust. Dunitic diapirs or obducted upper mantle are characterized by the presence of chromite deposits. Thus, we have the parental geoenvironment just as the recipient environment, in the sense that Cr, Fe and Co are mobilized and concentrated as metallic concentrations within that parental geoenvironment. 5

Similarly the PGE form minerals or mineral concentrations within the parental mantle geoenvironment, in which case the parental geoenvironment and the recipient geoenvironment are again identical. However, cases are reported where the elements Cr, Ni and PGE are mobilized far outside their parental geoenvironment, e. g., the reported association of chromite grains in fuchsite, where both the chromite grains formed and the Cr of the fuchsite represent Cr mobilizations outside their parental geoenvironment. PGM associated with sulfides are reported by Genkin (1959) and Ramdohr (1960). In the present effort such cases are considered to be mobilizations or remobilizations outside their parental geoenvironment. The mobilization has taken place by solutions. The formation of nuggets consisting of a number of PGM in lateritic covers of birbirites and dunites was treated in a number of publications by Augustithis (1967, 1979). As already mentioned, the Β group of elements of hydrothermal perigranitic deposits have been considered to be mantle derivatives or mantle recycled material which were involved in the granitization process either as parts of pregranitic material (remnants of which are basic xenoliths in granites) or as recycled derivatives of the initial gabbroic earth's crust in accordance with an unfolding spiral. In contradistinction to the mantle geoenvironment and to the fact that it acted both as parental and recipient geoenvironment, the lower crust and in particular the protolytic gabbroic geoenvironment have also been considered as a parental geoenvironment of metallic mineral concentrations. In particular, one must consider that basaltic rocks are mainly mobilized derivatives of the protolytic layer (as experimental petrofabrics have shown), and furthermore has to take into consideration that Mn, Cu, and Zn deposits may be leached derivatives of inter-rift buried basaltic flows, as for example the Mn superimposed mineral pangenesis in the Abu Dabbab apogranites. Considering that basalts are the most abundant rocks in the lithosphere elements, leaching from basalts will be a potential source for the concentration of many deposits characterized by the Β group elements. Perhaps the most significant parental rock of many of the metallic elements comprising the Β group was the initial gabbroic-basaltic earth crust. The recycling of elements in accordance with the unfolding spiral could result in the concentration of these elements in metallic concentrations or disseminated metallic occurrences. In contrast to the parental rocks of mantle and lower crust origin, the derivation of the lithophile group of elements, their abundance and distribution could be explained on the basis of the following considerations: The lithophile elements are differentiates of the unfolding spiral which resulted in the differentiation of the initial basic earth crust to gneissic granitic environments and provide or release a Mg source. Concurrently with this differentiation, there was or there occurred a concentration-segregation of lithophile ele6

ments in the gneissic granitic geoenvironment, or rather a preference of the lithophilic elements to associate with the granitic-gneissic environment. As granitization took place and as basic-in-composition pregranitic material was assimilated, a basic front release occurred which, together with a Mg and Fe release, resulted in the concentration of other metallic elements such as a characteristic metallogeny associated with skarns and in particular Cu, Zn and Mn. The skarn metallogeny, strictly speaking, does not belong to or represents a lithophilic element concentration. In contradistinction the concentration of the pegmatitic mineral parageneses and in particular the concentration of the C group of elements in pegmatites can be seen as segregations associated with the exudation of pegmatitic material from granites and gneisses. As suggested by Augustithis (1990) skarns and the metallic vein hydrothermal concentrations represent metamorphic-metasomatic differentiations in the field of metamorphism-granitization. The fact that the hydrothermal vein deposits contain both mantle-derived and lithophile elements is the result of the wide spectrum, as far as composition and derivation is concerned, of the granitized materials. Furthermore, during the unfolding of the geological spiral (Augustithis, 1973) recycling of the elements of the initial earth crust and a preferential segregation of the so-called "lithophile elements" took place in the differentiated rather more acid part of the crust formed; (the crust that had been formed as a result of the unfolding of the geological spiral). It comes as no surprise then that as a result of granitization-metamorphism and the mentioned differentiation under granitization-exudation, material (pegmatite formation) and element segregation (as veinform bodies) occurred not only of derivatives of the granitized material but of a wide field of metamorphosed and ultrametamorphosed rocks, i. e., in the wide field of metamorphism including granite emplacement (see Lyell, 1830; DrescherKaden, 1982; and Augustithis, 1973; 1985; 1990). In considering the recycling of elements one should not ignore the significance of the geological processes, namely, weathering, aggregation-sedimentation metamorphism. Also one has to see the hydrothermal vein concentrations associated with perigranatic metallogeny as mobilized or remobilized material of complex derivation and of a source, in addition to the granitized material itself. The wide country rock-region may in many cases be a significant contributor as well. Considering the above interrelationship it is difficult to draw a distinct demarcation line between parental and recipient geoenvironment. However, the point must be made that due to element recycling and element mobilization or remobilization, metallic concentration can take place in a recipient rock. The genesis and derivation of the metallic elements might be difficult to relate or trace though to a particular parent rock, which might exist or have existed at some distance.

In this connection it should be pointed out that the element mobilization or remobilization and recycling in a wide sense are attributed mainly to the great mobility (kinetics and solubility of the metallic elements) by "hydrothermal solutions" (solutions of relatively high temperature, of non-magmatic derivation) and of fluids of variable derivation (see Augustithis, 1990; Chapter 9).

(g) Mechanism of Mobilization gering Mobilization)

(Events Trig-

In considering element mobilization or remobilization and the recycling of elements one should look into the mechanisms which can bring about these mobilizations. One is inevitably forced to look at the operation of processes which cause significant developments in the earth's crust and which might be initiated, in cases, deeper than the lithosphere. Thus, inevitably the question of geotectonics or global tectonics and the forces and the mechanism that bring about these changes must be raised. The subject of geotectonics is beyond the scope of the present effort and only brief reference will be made to some of the current hypotheses that are related to the subject of material mobilization. The geosynclinal hypothesis and its merger with the plate tectonic hypothesis has provided the theoretical basis for crustal mobilizations and has supplied theoretical background for tectonic and tectonogenic changes in the earth's crust. Related to the geotectonic events are tectonothermal phenomena such as fusion, melting along geotectonic lines and volcanism-related to deep fissures of the earth's crust. Related directly or indirectly to the mobilization of the earth's crust due to mass movements, we have metallic element mobilization. Thus, accompanying geotectonic mobilization there is mobilization of metallic elements and their concentration. The plate tectonic hypothesis provides the basis for several metallogenic interpretations and hypotheses and for more than a decade was the basis for most metallogenic explanations (see Chapter 18). As a corollary to the significance of plate tectonics in bringing about element mobilization and metallic element concentration is the hypothesis of Kutina (1986) who suggested that there is a direct relationship between plate junctions and the main metallogenic provinces of regions of the earth's crust suggesting a

derivation of the metallic elements from the deep mantle source. In contrast to plate tectonics, several alternative hypotheses have been suggested. Geotectonic events such as expansion of the earth-hypothesis and many alternative interpretations have been put forward in direct contradiction to plate tectonics (see Critical Aspects of the Plate Tectonics Theory, edited by Beloussov et al., 1990). In understanding the mobilization and concentration of elements in the earth's crust, the shear fault lines or major faults in the earth's crust, some extending to great depths, are of interest. Along such lines fusion can take place resulting in melt mobilization either as volcanic lines or porphyry ores. The geotectonic events taking place in the earth's crust may actually be the expression of mantle mobilizations and even of the earth's core relative relation to the mantle (influenced by extraterrestrial forces?). As a result of mantle movements there is mantle mobilization either as mantle obduction or most likely due to the great crystal plasticity of forsteritic olivine/mantle diapirism. Mantle mobilization as cold or hot diapirisms could result in extensive element mobilization and metallic element mobilization or remobilization. One can trace the hypothesis of major crust changes and movements of material due to the influence of surface water contained in the rock masses to the phreatic cycles of Vernadsky. In further consideration of the phreatic cycles' hypothesis of Vernadsky, relating it to the changes of the theoretical unfolding spiral, a hypothetical mechanism is thus available for the understanding of all fluid movements which modern isotope geochemistry tries to classify as derivatives of deep magmatic source or as derivatives of the Bereich of metamorphism, or else, it ascribes to them a surface (meteoric) derivation. One is inevitably compelled to consider all these magmatogenic derivations of water or fluids in contradistinction to extraterrestrial bodies where so far there is an absence of [OH]. The fundamental question which is so hard to answer is whether all water in the earth's crust is of atmospheric derivation. The role of water as a medium for mobilization, transportation and concentration of metals goes back as far as the era of Aristotle. Ever since then it has been, to a greater or lesser extent, a major topic and is extensively treated in the present effort as a medium for metallic element transportation and concentration.

7

Chapter 2

Consideration of the Concepts of Paragenesis, Temperature Determination and Replacement, Based on Edwards' Book

One of the earliest pioneering books is Textures of the Ore Minerals and their Significance by A. B. Edwards (1960) who devotes an extensive part to discussing different aspects of the concept of paragenesis. Edwards' definition of paragenesis, together with an example given by him, is quoted as follows: "To determine the paragenesis of an ore, it is necessary to establish for each pair of minerals present whether they were deposited simultaneously or whether the period of deposition of one overlapped that of the other or whether one was deposited after the other. Frequently the minerals in an ore associate in groups so that an age relationship can be established between the members of an individual group and between the groups as a whole, e. g„ the cassiterite ores at Renison Bell, Tasmania: five groups or stages in deposition can be recognized in the paragenesis as follows: 1. Cassiterite, wolfram, tourmaline, Topaz and quartz (the quartz continuing). 2. Pyrite, arsenopyrite, gold. 3. Pyrrhotite, chalcopyrite, stannite, native bismuth. 4. Manganese-iron-magnesium carbonate, sphalerite, galena, tetrahedrite, fine-grained pyrite and marcasite (both derived from the breakdown of earlier pyrrhotite), a minor second generation of arsenopyrite (associated with the tetrahedrite), microscopic second generation of pyrrhotite (as ex-solution bodies in the sphalerite), jamesonite, canfieldite, pyrargyrite, franckeite, fluorite. 5. Carbonates." However, the concept of paragenesis (mineral association in the broad sense) has a long history as Petrovskaya (1986) reports: "Mineral associations can certainly be grouped with the most ancient objects of observations: common occurrence of one or another mineral substance was undoubtedly noticed by man in his first attempts to use and search for his necessary minerals. Still, the period of scientific comprehension of the noticed facts comprised only an insignificantly small part of a very long history of their accumulation. Likewise, the duration of this period seems to be small in the general evolution of the science of minerals. This situation may be res8

ponsible for the fact that concepts of the nature of mineral associations are still insufficiently clear." Less than two centuries divide us from the time when Romes de Lille in France, Severgin in Russia, Vezrer in Germany, Jemson in Britain and other scientists at the end of the 18th - beginning of the 19th century independently became aware of the considerable significance of mineral associations constantly observed in rocks and ores. Later such a conclusion was clearly formulated by Breithaupt (1848) who, as it is known, suggested the concept of "mineral paragenesis" to denote "their more or less pronounced ability for common occurrence". In contradistinction to the rather descriptive definition of the term paragenesis as outlined by Edwards, Udubasa (1986), considering the common occurrence of minerals in ores and rocks, attempts to present a definition of paragenesis based on the concept of equilibria. Udubasa states the following: "Many relationships observed among opaque minerals occurring in ores and (as accessories) in rocks show that the modes of apparition of minerals hardly fit (sometimes) any definition of "mineral paragenesis". This is why the mineral association (MAT) includes all the minerals found together, suggesting the complexity of geological processes generating them (e. g., remobilization). The mineral paragenesis (MP) is the most stable and wide-spread form of apparition of minerals formed under equilibrium conditions and/or later reequilibrated under the conditions of continuous change of PTF parameters. The mineral assemblage (MAG) contains minerals having common boundaries or appearing as incompletely developed MP due to special environment conditions such as variation of the deposition space, influence of the host rocks, sharp changes of the PTX in a given geological body, etc." Augustithis (1964, 1967, 1979, 1982, 1990) also presents different examples of mineral parageneses where the geochemical interrelationship of the elements comprising the mineral associations of paragenesis is emphasized. It is further supported by him that elements of relatively rare abundance in the earth's crust can segregate in paragenetic associations due to the interrelationship of the elements in accordance to the

empirical "laws" of the periodic system; however, the gathering together of metallic elements and the formation of crystal structures depends on other factors as well which are discussed in detail in Part III. Another aspect of the concept paragenesis is the simultaneous and successive deposition-interpretation presented by Edwards. The following extracts are characteristic of his orthodox views, especially concerning simultaneous deposition (formation): "Simultaneous deposition: Ex-solution intergrowths provide the best evidence of simultaneous deposition but in many ores cooling has been sufficiently slow to allow the complete unmixing and segregation of minerals precipitated in solid solution. Minerals crystallizing simultaneously but not in solid solution also develop mutual boundaries. Successive deposition: Successive deposition is proved if one mineral can be shown to occur as veins (other than antecedent or segregation veins) transecting the other mineral or to replace it or to be molded on it. Crustification, if present, will often reveal the sequence of deposition but if the crustification is an alternative banding of two or more minerals, it more probably indicates alternative fractional crystallization and seems best regarded as a special variant of simultaneous or overlapping crystallization. It is not possible to conceive solutions radically changing their composition for the deposition of each layer. Rather, the succession is to be interpreted as caused by delicately changing conditions of saturation." Whereas experimental mineralogy has often synthesized intergrowths (patterns) comparable or simulating ex-solution (e. g., ex-solution bodies of chalcopyrite in sphalerite), equally convincing is the textural evidence where marginal chalcopyrite has extensions attaining "ex-solutions" forms in the blende. Also, galena veinlets transecting sphalerite are often in turn transected or associated with "ex-solution"-like bodies of chalcopyrite. Moreover, ex-solution bodies of chalcopyrite have protuberances which develop into veinlets extending into and transecting the sphalerite itself with which they are supposed to be in simultaneous intergrowth. The reader is referred to the section on "Symplectites" (Chapter. 7). where a plethora of convincing textural patterns are provided of the replacement nature of a number of ex-solution bodies in different host minerals. Another aspect of Edwards' concept of paragenesis to be discussed is based on van't Hoffs law, that is, in a cooling system combinations take place which evolve heat; the combinations that let off the greatest amount of heat occur first. Furthermore it is assumed by Edwards that this holds during the cooling of a mineralizing solution; on this basis Edwards concludes that the "Comparison of the paragenesis found in a variety of hydrothermal deposits has shown that practically the same paragenesis is obtained in pyrometasomatic (contact metamorphic) hydrothermal, mesothermal and epithermal deposits. This sequence can be

generalized as follows, the minerals in the different groups being placed in their approximate order of deposition, the earliest minerals to form being cited first: 1. Magnetite, ilmenite, chromite, haematite. 2. Cassiterite, tantalite, wolfram, molybdenite. 3. Pyrrhotite, pentlandite, löllingite, arsenopyrite, pyrite, cobalt and nickel arsenides. 4. Chalcopyrite, sphalerite, (interchangeable), bornite. 5. Tetrahedrite, galena, lead sulfosalts, silver sulfosalts, native bismuth and bismuthinite, tellurides, stibnite, cinnabar. The transparent gangue minerals introduced also tend to show an ordered sequence: (i) Quartz, tourmaline, topaz. (ii) Siderite (often manganiferous) fluorite, calcite, barite, chalcedony (epithermal ores). The sequence corresponds in general with the orders of magnitude of the free energies of formation of the oxides and sulfides of the metal, the earlier formed oxides and sulfides having the greater free energies." In spite of the fact that van't Hoffs law might be applicable theoretically to cooling mineralizing solutions, our concepts of a mineralizing solution has been greatly changed over the last years. As isotope geochemistry shows, the derivation of fluids transporting metallic elements can be variable (meteoric, magmatic, metamorphic, or a mixture of all). This of course does not exempt them from van't Hoffs law. However, the fact that solutions can be changing in composition due to changes in the supply of material at the source and due to the fact that they may dissolve different substances in their passage through diverse rocks, may cause changes in their composition. More important though is that solutions with different composition may result in the dissolution or partial dissolution of already deposited minerals and in mineral replacement. (These changes and the resultant textural patterns are extensively discussed in the present volume). Furthermore, in contradistinction to the processes outlined by Edwards concerning the deposition of hydrothermal, pyrometasomatic, mesothermal and epithermal "deposition", the chapters of this volume on pyrometasomatic skarn, pneumatolytic-hydrothermal deposits present the complexity of processes involved in the formation of these deposits. Another aspect of Edwards' concept of paragenesis that deserves consideration is paragenesis and crystal structure. According to him: "Various relationships are a reflection of the relationship between paragenesis and crystal structure. The sulfide of the most abundant metal will serve as host or solvent to the others which will precipitate from it. The order of unmixing varies with the concentration of the various metals in solution, the temperature of their order, disorder, transformations and the rate of cooling, and since the early unmixing is into partial solid solutions which tend to 9

segregate and then unmix further, most complex sequences of deposition will result. According to the composition of the mineralizing fluids, deposition at this stage may be either simple or complex. In some ores a single mineral is deposited, as in some chalcopyrite ores, or the blende group minerals are deposited one after another, in a clear sequence. In others, however, a complex solid solution or group of partial solid solutions, is deposited which will unmix according to its composition to form sphalerite, pyrrhotite, chalcopyrite, stannite and tetrahedrite or tennantite. The sulfide of the most abundant metal will serve as host or solvent to the other which will precipitate from it. These sulfide minerals formed a series of extensive mutual solid solutions. Their role as host (solvent) or as precipitate (solute) varied locally in the ore according to the local concentrations of Cu, Sb, As, Zn, Fe, Ag and S, but from what is known of the temperature relations of these various solid solution series, all were deposited at temperatures above 400° C and probably above 475°C." Undoubtedly Edwards attempts to establish a principle in his paragenesis-crystal structure; namely, that the sulfide of the most abundant metal will serve as host or solvent to the others which will precipitate from it. Also on the basis of this principle, he attempts to establish some rules for a sequence of mineral deposition (crystallization) from solutions and, what is a more difficult task, to establish sequences of ex-solution formations. Here again the basic principle of Edwards might be sound but certain aspects of the role of sulfur in metallogeny and phenomenology of "exsolution bodies" need to be briefly discussed. The role of sulfur and basic concepts of its geochemistry are briefly considered in Chapters 38 and 39 (Sulfur in Metallogeny, and in Study Cases of Isotopes and their Significance in Metallogeny - (a) Sulfur Isotope Studies). Whereas sulfides can be formed from melts, solutions (fluids) and by bacterial action and furthermore, whereas the derivation of sulfur can be complex, as Sisotope studies support, Edwards' concept, "the sulfide of the most abundant metal will serve as a host or solvent to the others", is rather of relative significance since its derivation and supplies can be variable. As Cheney and Lange (1967) show in their contribution, "evidence, permissive and direct, is reviewed of the possibility of forming Sudbury-type ores by sulfurization, the introduction of country rock sulfur into the still hot intrusions". Whether this process is compatible with the concept of Edwards is difficult to elaborate. Equally difficult is it to explain how compatible with Edwards' concept, "sulfides of the most abundant metal", is Vinogradov and Grinenko's (1966) determination that "isotopic composition of sulfur in the sulfides of the Noril'sk copper-nickel ores support that 3050% of the total sulfur content of the intrusives was assimilated from sedimentary CaS0 4 ". Furthermore, and this is important, sulfur isotope studies indicate vari10

able derivation for sulfur in different vein deposits. Thus, the supply of sulfur can be of variable source and therefore variable supplies might be rendered available in the solutions involved in metallogeny. Edwards' concept of solid-solutions, ex-solution formation and sequence (which is widely accepted and experimentally supported) however is basically incompatible with textural observations on symplectic intergrowths as presented in Chapter 7 (Symplectites) where rather complex replacement processes are considered to be responsible for the so-called "ex-solution bodies" presented. Another interpretation of Edwards that must be further considered is where he states: "the sequence of crystallization of a paragenetic association is based on the solid-solution and ex-solved phases relation of the minerals involved". In support of his explanation, Edwards presents, as a study case, the "telescoped cassiterite-sulfide lodes in Eastern Australia", and states the following: "In these ores arsenopyrite and pyrite were the earliest sulfides to be formed. Pyrrhotite if abundant was the next to crystallize, followed by chalcopyrite, sphalerite and stannite. The solid solution relationships between these four minerals, commonly observed within the limits of a single polished section were: Host

Precipitate

Pyrrhotite (now marcasite) Stannite, chalcopyrite, sphalerite Chalcopyrite Pyrrhotite, sphalerite, stannite Sphalerite Chalcopyrite, stannite, pyrrhotite Stannite Chalcopyrite, sphalerite

Furthermore it appears that the relative concentrations of the various solute metals in the solid solutions controlled their order of unmixing in any given crystal. The temperature ranges for unmixing of the several components overlap because all the solute components commonly unmix simultaneously at some stage of cooling." However, the problem is that one should see the order of crystallization and the crystallization sequences from the point of view of multiple supplies and also from the vantage point of substitution-replacement. Considering the paragenetic associations from the point of view of a sequence of crystallization from one supply and under the aprioristic principle of ex-solutions separating under falling temperature, all this complies more with a deductive approach than on the basis of mineral to mineral relation as exact observations reveal. As a corollary to these comments on the above explanation of Edwards are the textural patterns presented in Chapters 5, 7 and 8, respectively

(Replacement Patterns and Processes, Symplectites, and Crystalloblastesis). Edwards' great versatility however must be admired when, in addition to the interpretation furnished of successive crystallization from a single solution, he introduced the concept of successive waves of solutions. Edwards explained the concept of successive waves as follows: "In some complex lead zinc ores, such as those of Captain's Rat, Roseberry, Tas, Mount Isa, etc., there is ample evidence that the bulk of the copper - as chalcopyrite - was deposited independently of lead and zinc, though more or less contemporaneously with them. The copper concentrations in these fields are irregularly distributed through the lead zinc ores and commonly occur marginal to, or even outside the lead zinc ore bodies proper. There is some evidence that the lead and zinc, also, were introduced in successive waves, the lead being the later. The influx of such large volumes of mineralizing fluids must have continued over a definite, though probably a short period of time, and have been spread through probably a considerable volume of rocks. The mineralizing fluids can be pictured as moving through zones in which the rock was temporarily in a state of tension, with fractures, cleavage planes and crystal boundaries providing the channel-ways." In support of the successive waves of solutions and their deposition are the textural patterns presented in the relevant Chapters (Replacement Patterns and Processes, and Successive Replacement of Successive Crystallization Phases). Also supportive of the concept of successive waves are the multimetallic mineralizations particularly those related to polymetamorphism and the cases of alternating prograde and retrograde metamorphism. Another subject treated by Edwards (1960) is the concentration factor (reduction of available sulfur) in considering "diversification" in the composition of solutions (fluids) and their composition (paragenesis). Edwards proposes the following: "A factor not often apparent is the reduction in the amount of available sulfur in the mineralizing fluids as a result of its reaction with iron of ferromagnesian minerals in the country rock to form pyrite. A progressive decline of the ratio of sulfur to introduced metals as mineralization progresses is probably responsible for the increase of metal content, relative to sulfur, of the later formed minerals in a series like pyrite-chalcopyrite-bomite-chalcocite and in the formation of complex sulfosalts in the late stage of mineralization. The role of sulfur in influencing the composition of the ore paragenesis is multiple and certainly deserves special attention in providing interpretations, particularly in interpreting the intergrowths of minerals such as pyrite-chalcopyrite-bomite-chalcocite-covellite native copper, etc., where the textural interpretation

should also be seen in terms of fluids with variable sulfur content." In addition to Edwards' interpretation which specifically explains the series pyrite-chalcopyrite-bornitechalcocite and the formation of complex sulfosalts as a result of reduction of sulfur in the series (which might probably be applicable to the hypogene mineralization) the reverse trend of enrichment in Cu (copper-rich zone) and decrease in it is though characteristic of supergene mineralization where leaching, transportation and deposition (often replacements) take place within the weathering zone of the upper-crust. Resurgence is also described, in addition to the successive waves of mineralization where repeated mineralization has occurred, giving rise to two or more complete generations of ore minerals. Edwards, when considering the possibility of such multigeneration deposits, summarizes the argument as follows: "For some ores however, it is claimed that repeated mineralization has occurred, giving rise to two or more complete generations of ore minerals, each showing a more or less normal sequence of deposition. Schneiderhöhn has sought to explain such resurgences in terms of the heating and cooling history of igneous intrusions and the surrounding rocks. In the lead-zinc deposit of Mount Isa, some geologists recognized three distinct generations of ore minerals each showing a normal sequence of deposition and claimed that there was a distinct time interval between each period of mineralization." In contradistinction, Edwards proposes that the three generations of sulfide deposition could be explained equally well on the evidence presented as a single period of mineralization with concomitant and continuing shearing of the ore bodies, larger and larger fractures being developed as the originally plastic shales were rendered more and more brittle by impregnation with sulfides. It should be noted, however, that both Schneiderhöhn and Edwards accept as the source of such multigeneration deposits (e. g., Mount Isa, Queensland, Australia) fluids or solutions related to igneous intrusives. Considering however the Mount Isa deposit, opinions have greatly changed and there is a controversy between syngenetic and epigenetic sedimentary (see Chapter 30, Mount Isa Controversy). To quote only one of the arguments put forward, the following extract from Murray (1961) is presented here: "Previous theories on the mode of ore emplacement, involving at least four periods of deposition, are too complex; early pyrite is probably syngenetic, and the main Pb-Zn ore bodies epigenetic, were introduced through actual emplacement between the beds rather than by replacement of them." Reading today the section of Edwards on temperature determination and considering that modern literature is mainly restricted to temperature data based on liquid-gas inclusions and their homogenization and to a lesser extent on the homogenization of ex-solution 11

bodies, one tends to accept that such methods and procedures (e. g„ typomorphism, inversion point, monotropic substances, recrystallization) under the heading of temperature determination are almost a forgotten state of the art which can be found in special textbooks of mineralogy. Also, Edwards' synoptical description of geothermometry, as recorded here: "the pyrite geothermometer is based on the postulations that (1) crystals of any given electronically conducting mineral species deposited at a high temperature are more perfect and have a more positive thermo-electric potential than crystals of the same mineral deposited at low temperatures and, (2) that the degree of crystal perfection and the thermo-electric potential vary continuously between any given limits of temperature of deposition", is hardly ever used in the sense described. Experimental and especially synthetic mineralogy have greatly substituted these procedures and the approach of extrapolating experimental results to natural cases furnish data on the geothermometry and geobarometry of minerals and the formation of mineral assemblages. However, the concept of temperature of ore deposits as stated by Edwards, namely that: "orthomagmatic deposits are formed at temperatures above 600" C, pegmatitic deposits at about 570-600° C, pyrometasomatic deposits generally below 600° C, and that epithermal deposits are formed between 50-200° C, mesothermal between 175-300° C and hypothermal between 300-600° C", is somehow acceptable, though more precise data are usually presented on the basis of homogenization of liquid-gas inclusions and ex-solution bodies. The same applies to the temperatures for high, intermediate and low temperature minerals, and to the range of temperature and range of cooling of minerals (see Edwards' table, 1960). Experimental mineralogy and particularly synthetic mineralogy have substituted the old methods and procedures of determining temperatures of mineral formation, and have provided a plethora of data which are quoted in different paragenetic associations presented in this volume. Perhaps one of the most significant statements of Edwards in his monograph Textures of the Ore Minerals and their Significance is that "hardly one of the known textures or structures taken separately is a safe criterion against metasomatism-replacement". Since in the present volume a detailed treatment of the replacement textural patterns is presented in Chapters 3, 5 and 6 (Inductive Versus Deductive Approach in the Interpretation of Textures, Replacement Patterns and Processes, and Replacement Versus Ex-Solutions) certain principles of replacement by Edwards will be quoted as a corollary to the significance given by the author of the present volume to replacement textures and processes. Edwards' definition of replacement can be traced to a number of statements on the subject which are quoted (for maintaining the exact sense attributed to the topic by Edwards). He states the following:

12

"The dissolving of one mineral group and the simultaneous deposition of another mineral or group of minerals in its place is a common feature (process) in ore minerals. Solution (dissolution) and deposition proceed concurrently without the intervening development of appreciable open spaces, the substitution often involves no change in volume. Supergene replacement processes leading to secondary sulfide enrichment are sometimes more dramatic and of greater economic significance but hypogene replacement of sulfide minerals is a major process in the deposition of any complex sulfide ore. A more soluble mineral is always replaced by a less soluble mineral". Another significant statement of Edwards is about diffusion. He characteristically states that: "diffusion plays almost as important a role in replacement as it does in ex-solutions. As in solid solutions, the process is one of substitution. The substituted atom may be retained in the new mineral (metasome) or may diffuse into a region of increasing mobility, culminating in a solution and thus be removed from the locality. Both cations (metal atoms) and anions (chiefly sulfur atoms) may be replaced. The replacement may be selective, one member of the intergrowth being more readily replaced either on account of favourable chemical composition or, of favourable atomic structure (pronounced crystallographic planes). Even within a single mineral, a particular zone may be preferentially replaced owing to some favourable variation from stoichiometric proportions". An extensive consideration of diffusion is presented in conjunction with "ex-solution"-like bodies, interpreted as due to replacement as discussed in Chapters 3, 6 and 7 of the present volume (entitled: Inductive Versus Deductive Approach in the Interpretation of Textures, Replacement Versus Ex-Solutions, and Symplectites). Edwards quotes and summarizes Schouten's (1934) experiment of diffusion replacement as follows: "this demonstrates how a solution which percolates through a sulfide ore will change in composition according to the composition of the minerals it transverses to accomplish a series of replacements. Such a terrain of replacement can take place in one direction. However, the replacement will be of the more soluble minerals by less soluble minerals". Schouten's experiment is very important since it can explain the repeated and multiple replacement patterns so often observed microscopically in ores. Edwards describes certain patterns of replacement emphasizing the often observed selectiveness of the process in the following way: "Replacement, especially supergene replacement, commonly develops along grain boundaries when the metasome grows as a narrow rim around the host. It may resemble a texture formed by complete unmixing of the solid solution, namely that in which diffusion has enabled the exsolution bodies to aggregate into narrow elongated grains in the grain boundaries of the original host. But

in some instances penetration of the host has developed along crystallographic planes which intensifies the resemblance to ex-solution textures". However, the simulation of "ex-solution" textures by replacement, as presented by Edwards, is an understatement since as presented in the present volume most of the so-called "ex-solutions" are actually replacement textures. Another pattern quoted by Edwards, in addition to the cases described of preferential replacement of certain minerals by others, is core replacement (Atoll type) and zone replacement. Ample description of core, zonal and interzonal replacement are presented in the present volume when discussing the textural patterns. In conjunction with replacement patterns simulating ex-solutions, a series of patterns due to replacement simulating eutectic intergrowths are described by Edwards who, considering them in comparison to true eutectic intergrowths, states the following: "The ore minerals most prone to crystallize, forming melts in nature are the oxides, magnetite, rutile, ilmenite and haematite and these minerals also tend to form eutectic intergrowths. The eutectic and eutectoid intergrowths are of two general types, (a) the so-called graphic intergrowths and (b) the lattice intergrowths (resembling ex-solutions). These pseudo-unmixing or pseudo-eutectic textures also fall into two groups, (i)

graphic and subgraphic intergrowths in which the replacement has progressed more or less independently of the crystallographic directions of the minerals undergoing replacements and (ii) lattice intergrowths along preferred crystallographic planes of the host. Graphic textures resulting from replacements can frequently be recognized by their association with veins of the metasome transversing the host. If the replacement is arrested in mid-stage, a bornite chalcocite intergrowth remains which consists of triangular remnants of bornite in the interstices of the lattice of chalcocite lamellae." The eutectoid-type textural patterns and their formation are discussed in Chapter 7 (Symplectites, Myrmekitoids and Graphic Symplectites) as well as in the section: Breakdown Symplectites, where a plethora of textural patterns supportive of replacement graphic and myrmekitic pattern is presented. Concluding the discussion on Edwards' masterpiece, Textures of the Ore Minerals and their Significance, I would like once more to recall his statement "hardly one of the known textures and structures taken separately is a safe criterion against metasomatism replacement" and to add Drescher-Kaden's comment (pers. comm.) that ore minerals are very reactive and therefore their textures are very complex.

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Chapter 3

Inductive Versus Deductive Approach in the Interpretation of Textures

(a) Solid Solutions, Ex-solutions - Deductive Approach A deductive approach to the entire spectrum of the phenomenology of ore mineral textures can be based on: (i) The synthesis of ore minerals from melts and solutions; (ii) The concept of the continuous isomorphic crystallization series and particularly on solid solutions which in turn depend on the following factors: - Atomic size of the solute and the solvent; - Valency of the atoms involved; - Crystal structure; (iii) The extrapolation of experimental results of cooling or annealing of synthetic metallic compounds to natural ore minerals and particularly to native metal compounds. (iv) The experimental results of homogenization and homogenization rates of ex-solution bodies of natural ore samples.

(b) Consideration of the Deductive Principles Concerning Solid Solutions and Diffusion If we assume that the solute is present in the solvent as a result of crystallization from melts or a precipitation from solutions, the attainment of the state of solid solution would depend on the following factors, provided sufficiently high temperatures prevailed: (i) The relative size of the metal atom; (ii) The valency of the two metals; (iii) Their crystal structure. Therefore, the most favorable conditions for the formation of a wide range of solid solutions are that the solvent and the solute are of nearly the same size, do not form intermediate compounds and have the same valency. Furthermore, the factors which favour solid solution and the factors that favour diffusion appear to be a contradiction to the role of some of the factors favouring solid solutions and diffusion, respectively. In contradistinction to the factors favouring solid solutions the following favour diffusion (ex-solution): - Low solubility of the solute; 14

- Larger difference of melting points between solute and solvent; - Large difference of atomic radii of solute and solvent; - Increasing separation of solute and solvent in the periodic table (a most constant factor) which actually is the valency difference. Comparing the factors favouring ex-solution with the factors favouring solid solution it can be seen that whereas large differences in atomic radii of solute and solvent encourage diffusion, solid solution in contrast is favoured when solute and solvent atoms are of nearly the same size. Also, separation of the solute and the solvent in the periodic table is a factor favouring diffusion whereas solid solution is aided when solute and solvent have the same valency. It appears that these contradictions complicate further the problem of solid solutions - ex-solutions. Another interpretation concerning the problem is that the crystal structure is in a disordered state at high temperatures and thus can accommodate the solute atoms. As cooling takes place and the crystal structure passes from the disordered to the ordered state the solute atoms ex-solve or diffuse out. The unmixing of solid solutions takes place by the diffusion of the solute atoms or ions through the lattice of the solvent substance. However, an indisputable fact remains, that of the speed at which unmixing can occur as demonstrated by experimental studies (see Edwards, 1960). Solid solutions of bornite and chalcopyrite cooled from 600° C to room temperature in 5 minutes, unmix to some degree in this time interval. Also another feature of unmixing, revealed by experimental studies, is the speed with which ex-solution bodies will segregate to the grain boundaries of the host mineral if the specimens are annealed or cooled slowly. The above interpretation however is difficult to reconcile with the observations that ex-solutions persist as intergrowths in ore deposits in which the cooling process or annealing lasted perhaps thousands of years. Considering the fact that with diffusion (ex-solutions) atomic migration takes place on cooling and also that rapid migration of ex-solutions takes place at mineral boundaries, somehow incompatible with observations in natural materials, it follows that in nature many of the ex-solution bodies observed may be prod-

ucts of other complex processes, such as replacementsubstitution or substitution-diffusion.

(c) Inductive Approach in Considering Ore Textures In contrast to the deductive interpretation of ore mineral textures, different interpretations can be furnished for most ore mineral textures on the basis of an inductive approach. The following basic concepts support an inductive approach for the interpretation of ore textures: (i) "Hardly one of the known textures or structures in ore specimens considered separately, is a safe criterion against metasomatic replacement." This statement by Edwards (1960) confesses the brutal truth that replacement is a process that can explain most of the ore structures and textures which up to now have been interpreted as products of ex-solutions and eutectic crystallization, etc. However, it should be emphasized that despite the above statement, Edwards did not pursue this concept to the extent that it could have been pursued or applied in explaining ore textures and structures. He made no effort to explain textural patterns that simulated ex-solutions and eutectics to furnish an alternative interpretation on the basis of replacement. It should be mentioned though that despite Edwards' chapter on replacement which shows excellent examples of the process, his abidance to the orthodox concepts of interpreting ex-solutions as unmixing and eutectic textures as simultaneous crystallizations deprived him from the possibility to consider these patterns with the necessary objectivity and to give him the opportunity to provide alternative interpretations based on exact observations. This criticism also applies, in my opinion, to the other classical ore microscopists, and in particular to Ramdohr who with no doubt observed most of the discrepancies between the orthodox interpretations and that presented by phenomenology. Perhaps my great teacher did not want to be an iconoclast, or perhaps did not want to confuse all of us by saying all he knew. (I am convinced that he had observed most of the mentioned discrepancies.) Perhaps he did not think they were important. The statement by Edwards that hardly one of the known textures or structures in ore specimens considered separately, is a safe criterion against metasomatism-replacement leads us to the task of having to consider every textural pattern (taking into consideration all possible variants of a pattern) on its own merit and perhaps understand textural patterns on the basis of the principles of comparative anatomy, a principle that has proved to be most useful in biosciences. (ii) Could textural analysis furnish evidence and interpretations that are contrary to the theory of solid solutions, eutectic crystallization and in general, to the interpretations that are currently accepted? If a pattern is established as being (on support of textural evi-

dence) the product of replacement on the basis of comparative analysis could this interpretation of the pattern be widely applicable or acceptable? As the evidence proved in this Atlas supports, many of the textures which were previously considered to be undoubtedly the result of ex-solutions or eutectic crystallization are shown to be due to replacement and substitution processes. It is therefore possible that once these patterns have been established as being the result of replacement this interpretation could become widely acceptable. It is thus important to present as many study cases as possible of each pattern in order to provide evidence of its replacement origin. Similarly, could concepts, such as diffusion, solid solutions, continuous isomorphic series, diffusion rate, etc., find an application in the replacement hypothesis of some of the most common ore textural patterns? To what extent are these concepts (diffusion, solid solutions, isomorphic continuous series) incompatible with the evidence of replacement? (iii) Substitution in terms of removal and replacement without leaving gaps is considered by Edwards as a significant process in ore textures. (iv) Therefore the question arises whether the substitution (removal and replacement) processes of Edwards have far greater significance and application than recognized up to now? It is possible that a wider spectrum of textural patterns could be interpreted on this basis than has been so far. As the evidence provided in the Atlas shows, the process is of fundamental importance to the ore microscopy. (v) Wave substitution or successive replacement is another process of great significance since multiple and successive replacements commonly take place in almost all ore deposits. As the textural evidence supports, successive replacements may simultaneously take place in the sense that many separate successive replacement series can occur simultaneously. (vi) Factors significant for unmixing, diffusion and ex-solution formation are also of importance to replacement processes and reactions. In particular, the following factors need to be mentioned: - Substitution in solid solutions which is influenced by the relative size of ions, the valencies, the electronegative effect and the structure, also plays a significant role in replacement. - Temperature and pressure suitable for unmixing and affecting substitution, are also effective. - Solid diffusion and diffusion of solutions are also significant for replacement. - There is also a relationship and interdependence between atomic substitution and mineral substitution. (vii) Synantetic and symplectic patterns were previously considered to be of evidence for eutectic simultaneous crystallization, in ore mineralogy and in particular in petrology. Graphic and myrmekitic-like forms were considered to be an indication for eutectic crystallization. However, the extensive evidence pre15

sented in this Atlas supports that most of the symplectic and synantetic textural patterns in ore mineralogy are the result of replacement. In this connection it should be emphasized that synantetic and symplectic textures in rocks are also infiltration and replacement textures (see Augustithis, 1973, 1978, 1979, 1982, 1985, 1990 and 1993). Regarding incompatibility between theory and observation, Edwards himself (1960) emphasized serious incompatibilities between hypothesis and experiment. Specifically, he pointed out that "another feature of unmixing, revealed by experimental studies, is the rapidity with which ex-solution bodies will segregate to the grain boundaries of the

16

host mineral if the specimens are annealed or cooled slowly... It is difficult to reconcile these observations with the persistence of ex-solution intergrowths in ore deposits in which the cooling processes or annealing lasted perhaps thousands of years". These observations sustain that the extrapolation between hypotheses based on physico-chemistry and natural phenomena might after all be inapplicable, in the sense that the wrong physico-chemical explanation is applied while other physico-chemical explanations would have been more appropriate. In this particular case the quoted physico-chemistry of replacement processes would be more applicable.

Chapter 4

Mantle and Lower Crust Derivatives

When the evolution of our thinking regarding mantle is considered it should be stated that in the decade of the 1960s after a chimera that had lasted several years, a re-examination of the concepts started concerning the mantle, its composition and compositional variations, its state and behaviour and above all, to what extent it is implicated or involved in the evolution of the crust. The migration of continents was a fascinating thought and early concepts had already been aquainted in the 15th century (Italian concept based on the outline of the continents). However, it was the presentation of the continental drift theory by Wegener in 1924 that triggered the reconsideration of the fixed-continent concept. The plate tectonic theory, incorporating elements of the continental drift theory, the geosynclinal theory and the convection current hypothesis, unfolded in the late sixties and early seventies as the dominant geotectonic hypothesis. The hypothesis soon attained the status of a theory and for many it is now a fact. Opposition to the acceptance of the plate tectonic hypothesis soon gave in and during the seventies and the greatest part of the eighties it constituted an almost unchallenged global interpretation. However, in the late eighties and early nineties, opposition and reconsideration of the plate tectonic theory began to be reinvigorated, particularly due to the publication of the "Critical Aspects of the Plate Tectonic Theory" (Theophrastus Pubs., 1990). One of the main theories that contradicted the plate tectonics was expansion of the earth, the proponents of which had long fought the plate tectonic theory. Apart from its possible detrimental influence as a whole on geosciences, the plate tectonic hypothesis nevertheless had several positive points and introduced ideas in particular that involved the mantle in the evolution of the crust. The obduction of lower crust-upper mantle and ocean spreading introduced ideas that promoted the concept that mantle is not very far away from crustal events and processes. However, credit must be given to many national and international mantle projects that started in the early sixties and seventies. These projects (e. g., "Upper Mantle Project, 1969, and "Das Unternehmen Erdmantel", 1972) showed the importance of the mantle and the mantle movements in influencing the stability of the crust and

furthermore projected the idea that many ultrabasic bodies were actually diapiric mantle derivatives. As a consequence of the mantle projects the derivation of basalts was also re-examined. It should be mentioned though that the theories on the derivation of basaltic trends from a substratum (Kennedy, 1933) introduced the concept of a deep-lying layer which was the mother layer for basaltic flows on the earth's surface. However, the most fundamental step in distinguishing mantle from crust was Mochorovitch's discontinuity. It must be stressed though that the projects on mantle focused on the significance of the mantle, on its properties and on the fundamental fact that any events in the mantle profoundly influence both the crust and the earth's surface. Parallel to the evolution in thinking on geotectonics and of equal importance, hypotheses and interpretations were put forward by petrologists and petrographers. Thus, the magmatic differentiation by Bowen (1928) and Niggli introduced a model of interpretation not only for the different so-called igneous rock types, and as applied particularly by Buddington (1943) furnished interpretations not only for the layered ultrabasic bodies but also for the greatest part of the profile starting from the lower crust to the upper crust of the earth. On the basis of comparing different layered bodies Buddington believed that there is a systematic variation based on the differentiation hypothesis and he recognized (from bottom to top): 1. A gabbroid or noritic facies (a chilled facies); 2. Ultramafic rocks; 3. Norite or olivine gabbro; 4. Anorthosite with ilmenite magnetite layers; 5. Quartz or granophyric bearing augite gabbro; 6. Granophyre or granite. It is interesting to note that Buddington drew attention to the parallelism between layering of the earth's crust and the layering of the stratiformed sheet of igneous complexes (in the sense that the lower crust/upper mantle is also considered). It must be emphasized that despite the fact that magmatic differentiation introduced an interpretation for the genesis of the main igneous rock types and influenced the thinking for at least half a century, it actually was a very big set-back in the evolution of understanding rock types. The derivation of granitic rock 17

types in particular as a magmatic differentiation of the basaltic magma possibly hindered as unprecedented obstacle the development of further thinking in this field for several years. The generation of granitic rocks from sediments and the evolution of great masses of granites from initial sedimentary rocks, e. g., in South East Asia is supported by the granitization theory and even the many anatectic interpretations such as palingenetic origin of granites, anatectic granites and the more recent equivalent of anatectic granites, the Sgranites, argue against the derivation of granites from the basaltic melt by differentiation. In addition, the derivation of the layered bodies themselves from a basaltic parental melt has been challenged not only by Augustithis (1979), but also by the fact that a lot of the ultramafic and ultrabasic rocks are considered obducted lower crust/upper mantle and not as basic differentiates of basaltic melts. Mantle projects all over the globe showed that the mantle is nearer (in terms of involvement) to the crust than previously believed. Thus, ultrabasic rocks such as peridotites, dunites, lherzolites, etc., are believed to be either mantle obducted or diapiric mobilizations of the upper mantle. Also in this case, the differentiation derivation of ultrabasic rocks from the parental basaltic magma or reservoir is outdated and hardly provides a basis for modern petrographic interpretations. In contrast to the hypothesis of magmatic differentiation which prevented the evolution of petrography, contributions in the early sixties on the derivation of olivine bombs in basalt were believed to be of restricted interest at that time. Some groups introduced the idea that olivine bombs were due to magmatic early crystallization of the basaltic melt and therefore not of global significance. Despite messages being sent by Ernst that olivine bombs are "Briefe von der Tiefe" (letters coming from great depth), their significance was not widely accepted. Augustithis (1972, 1978) confirmed Ernst's suggestion (1962, 1963), that olivine bombs actually were upper mantle fragments in basalts and basalts are derivatives of the protolytic layer (gabbroic in composition) of the lower crust. Augustithis (1978) included a chapter in his Atlas of the Textural Patterns of Basalts and their Significance entitled "Mantle petrography (the petrography of mantle fragments brought up by basalts)". In the present effort though, an attempt will be made to show some of the most significant patterns of ore minerals and their intergrowths (mainly spinels) and to compare them with other similar patterns from dunitic complexes also believed to be diapiric mantle intrusions. Augustithis (1979), in his Atlas of the Textural Patterns of Ultrabasic Rocks, introduced the idea that "Schlieren chromites" were actually tectonically fragmented chromite (tectonogranular, see Chapter 14) and that they were tectonically mobilized in the crystalloplastic forsterite-rich olivines of the dunite complexes (mantle, diapiric in derivation). Also chromite 18

lenses occur in dunites and the leopard ore bodies in serpentinized dunites have been interpreted as boutinage (large and small scale) of initial chromite segregations in the earth's mantle. The tectonogranular chromite as well as the lenticular chromite bodies and their distribution in the dunitic massifs is not due to magmatic segregation or magmatic gravitation-differentiation but due to prototectonics in the visco-plastic forsterite-rich mantle. The chromite crystals with epitactic magnetite found in diapiric ring intrusions are again mantle derivatives due to diapirism of the viscoplastic dunites of the ring complexes. In addition, the interbanded chromite-anorthosite bodies, particularly the type of the Dwars River occurrence in the Transvaal, are also due to crystalloplastic mobilization of chromite and anorthosite (see Chapter 14). On the other hand, the horizontal bands of chromite of the Bushvelt complex and in particular their occurrence with dunite or pyroxenites could represent mantle derivatives due to tectonoplastic behaviour of the host rocks and of the chromite itself (under conditions of great pressure or perhaps pressure release), or could possibly be of ultrametamorphic origin (see Chapter 27, section (h)). In this connection it should be mentioned that the banded iron layered deposits involve many processes and they represent iron segregations that possibly occurred under entirely different geoenvironmental conditions such as those prevailing during early Precambrian. Perhaps there was a transgression or transitional geoenvironment between the lithic era and the beginning of the geological spiral that started to unfold. The banded iron formation and the banded chromospinel occurrences, whether within ultrabasic complexes or related to other geoenvironmental conditions, perhaps should be seen as iron and chrome segregations in an entirely different geoenvironment than that prevailing today. In this connection, the unfolding of the geological spiral is of interest as suggested by the author (Augustithis, 1973) which attempts to explain the transformation of the initial earth's crust from a basic ultrabasic composition to the rather acid crust of today on the basis of non-uniformitarianism interpretations. The debasification processes were evolved with the unfolding of the geological spiral but in the initial phases a geoenvironment more basic in composition predominated in which accumulation of iron and chromospinels may have been possible either by surface reworking (due to exogenetic processes) or by internal mobilization processes. Perhaps the formation of banded chromospinels consisting of rounded chromite sphaeroids with metamorphic garnets, uvarovites, might reflect such conditions prevalent in the early stages of the unfolding of the geological spiral. In an attempt to show some textural patterns indicative of petrogenetic conditions prevailing in the mantle (product of solidification of the initial star "earth") patterns will be presented exhibiting the typical intergrowths between spinels and Fe-Mg-rich silicates

(olivines, bonzites, etc.) that comprise typical olivine bombs found in basalt. The olivine bombs lherzolitic in composition found in basalts are believed to represent a typical sample or, should we say, a common type of mantle material since many olivine bombs from different basaltic flows of occurrences widely apart (Ethiopia, Eifel, etc.) consist of lherzolitic material (see Augustithis, 1978, 1979). In addition, comparison of these textural patterns will be made with lherzolites which are believed to be either obducted upper mantle or mantle diapirs. Fig. 1 shows olivine with a deformation lamella, indicating that deformation was effective on the forsterite-rich olivines which were in a visco-plastic state. However, as shown in Fig. 1, the olivine was sufficiently solid to sustain deformation effects such as the formation of the deformation lamellae. It is interesting though that spinels are intergrown with the olivine or follow the margins of the deformation lamellae. Fig. 2 shows a graphic-like spinel surrounded by a crushed zone of bronzite. The spinel is also enriched marginally in iron. The bronzite-rich mantle was sufficiently solid to sustain deformation effects such as the crushed zone developed between the harder spinel and the softer bronzite (which, as pointed out, must have been in the solid state). By comparing the patterns of spinels in intergrowth with olivines and bronzite, as exhibited in olivine bombs (mantle derivatives), with the patterns of spinels in lherzolites, either obducted or diapiric mantle, the patterns of spinel intergrowth are comparable and commensurable. Therefore, there is little doubt that these spinel intergrowths with the Mg-Fe silicates represent typical mantle patterns. Furthermore, the graphic pattern of these spinels should be seen as skeletal crystals that did not attain their idiomorphism. Considering that dunites are mainly the host rock for chromospinels and that pyroxenites are more often the host for magnetite, this observation led to the suggestion that the Mg-rich ultrabasics have greater affinity geochemically with the Cr-rich representatives of the spinel group and that in contrast pyroxenitic ultrabasics (considerably richer in Fe than the dunitic-peridotitic types) have a greater geochemical affinity with Fe spinels. In this connection, it should not be forgotten though that in most of the olivine bombs (lherzolitic in composition mantle fragments) found in basalts, spinels predominate and often spinels (M 2 M0 4 ) with M2 are predominantly represented by Mg. It should be stressed that lherzolitic mantle fragments (olivine bombs) in basalts are rich in olivines and bronzite (see Augustithis, 1973, 1979) and thus, the limited presence of chromospinel should be understood on the basis that the lherzolitic mantle fragments in basalts are derived from the mantle that is rich in olivines and bronzite. In contrast, ring intrusions have a central body dunitic in composition (see Augustithis, 1965, 1979) and are surrounded by pyroxenite. A typical example is the ring

dunite-pyroxenite intrusion of Yubdo, Ethiopia. It should be noted that the diapir is considered to be a mantle diapirically mobilized (apart from the disposition of the intrusive character of the ring complex, this is supported by the fact that forsteritic-rich olivines are also the main mineral components). The presence of chromospinels with epitactic magnetite and the relative abundance of sperrylite support a deep mantle derivation in contrast to the mantle lherzolitic in composition that is represented as mantle fragments in the basalts (olivine bombs). Another observation to note is that with the diapiric mobilization of the dunitic in composition mantle of the Yubdo ring complex pyroxenite was also mobilized. This can be interpreted as the main dunitic mantle acting as the diapiric-mobilized crystalloplasticallyactive body which contained also mobilized sections of the lower crust/upper mantle (represented by the ring peripheral body of pyroxenite) of pyroxenites (less plastic). Thus, the ring intrusive (diapiric) body of Yubdo represents centrally a dunitic deep seated mantle and peripherically also mobilized by the diapirism pyroxenite. The pyroxenite, as mentioned, might represent the lower crust/upper mantle. Thus, there are representatives of the lherzolitic mantle, rich in forsteritic olivine, bronzite and spinels (rich in Mg); dunitic diapirs and dunitic "intrusive complexes" (again diapiric or obducted) being rich in olivines and chromospinels; and finally there are diapiric ring complexes with forsterite-rich olivines and chromospinels with epitactic magnetite. (In the case of diapiric mantle derivatives which are perhaps mobilized from great depth, we usually observe more platinoids enriched or present, as in the case of the Skiros chromite with epitactic magnetite.) However, lherzolitic fragments in basalts rarely contain spinels which are closer to magnetite in composition. In this connection it needs to be pointed out that basalt transversing dunitic complexes (as in the case of Yubdo basalts) may pick up chromite xenocrysts. Spinel xenocrysts also occur in basalts (Mg-rich spinels, derivatives of a lherzolitic mantle). It should be noted though that despite the predominance of lherzolitic in composition fragments in basalts from different parts of the world and despite the fact that lherzolitic mantle might be predominant and widespread, our knowledge (petrographically) is nevertheless limited concerning the composition of the mantle. On the other hand, the great abundance of dunitic rocks - and considering that dunites are also diapiric or obducted mantle (mantle/lower crust) - suggests that another important mantle type is that of dunitic composition. As mentioned, spinels are most common in lherzolitic fragments in basalts (olivine bombs) and chromospinels are most abundant in dunitic diapirs or obducted dunites, Augustithis also presented cases of dunitic diapirs or obducted body in which spinels predominate (pleonaste?) in contrast to most of the dunitic bodies which contain spinels. However, examples of 19

ferro- and even chromospinels in olivine bombs (lherzolitic in composition) should not be excluded. Fig. 3 shows graphic skeletal spinels in serpentinized dunites (diapiric or obducted in origin) from Konitsa, Northern Greece. In contrast, most tectonically mobilized diapiric bodies such as those from Vourinos, Greece, in addition to lenticular chromite bodies (large scale boutinage as mentioned), also show leopard chromite ores and Schlieren chromite which is mobilized tectonogranular chromite bands as mentioned (see Chapter 14). Fig. 4 shows a Schlieren chromite band from the Xerolivado chromite occurrence, Vourinos. In opposition to Fig. 4 which displays the Schlieren chromite exhibiting tectonogranular character and later formed cataclastic effects, Fig. 5 shows idiomorphic chromite, perhaps a later generation than the tectonogranular Schlieren chromite with which it co-exists. Often the idiomorphic chromite is surrounded by later crystallized antigorite or the antigorite acts as a rebinding material. As previously stated, the chromite possibly represents a later crystallization stage than olivine in which it might be included and zonally incorporated, occasionally exhibiting corroded outlines. In contrast, Fig. 6 shows chromite with a band of olivines, some of which exhibit idiomorphic outlines. In this case the olivine is zonally incorporated in the chromite suggesting that some of the olivine crystallization occurred simultaneously with the chromite. These intergrowths represent phases in the solidification of the mantle and they could be seen as processes of fundamental significance in the crystallization sequence of the minerals comprising parts of the mantle. As already referred to, in the dunitic ring complex of Yubdo the main ore minerals are idiomorphic chromites with epitactic magnetite and in cases exhibiting sperrylite as inclusions in the chromite or in the epitactic magnetite (see Figs. 7, 8 and 9). The relationship between PGM and chromite is very complex and often platin occurs as inclusions in the chromite. However, cases are reported and observed where the platin is mobilized along the cracks of the chromite. Chromite is the most abundant spinel in dunites; however, in addition to the spinel (pleonaste) cases of idiomorphic magnetite in serpentinized dunites also exist (see Fig. 10). Perhaps banded chromite sphaericules, associated with uvarovites (and in some cases having chalcedony), should also be considered as mantle derivatives or as reworked chromites in which case though the chalcedony occurred later. Fig. 11 shows chromite sphaericules with uvarovite and chalcedony (later formed and interspersed into the uvarovite). In contrast, Fig. 12 displays sphaericules of chromite and chromite exhibiting crystal faces. Such patterns comprising chromite associated with uvarovite and interbanded with uvarovite bands perhaps represent a formation of chromite deposition in an environment characteristic for the transitional period of the 20

lithic era and the beginning of the unfolding of the geological spiral. In contrast to the peridotitic dunitic rocks another series of rocks metamorphic or ultrametamorphic in origin comprise pyroxenites, gabbros (paragabbros) and paranorites (see Augustithis, 1979). Textural patterns will be presented of the layered complex of Skaergaard, showing intergrowths of the main rock silicates with iron oxides (from iron-rich bands of the complex) and in particular, magnetite or titanomagnetite. Despite these rocks are not strictlyspeaking "mantle", they are nevertheless derivatives either of the mantle or of the lithic crust. They occur in layered ultrabasic complexes representing iron-rich bands (comparable to the chromite-rich bands in the Bushveld Complex), some of which show the already described sphaericule textures (see Figs. 11 and 12) and which again might represent formations manifesting the geoenvironmental conditions that prevailed in the transition period between the lithic crust and the unfolding of the geological spiral. The Skaergaard layered complex has been considered to be a basaltic intrusion which, due to magmatic differentiation and gravitational settling of crystals, resulted in a complex of layered rocks, believed to be igneous in origin (see Wager and Deer, 1939). In opposition, Augustithis (1979) considers the layered bodies to be complexes where mantle mobilization (diapiric or obducted - in the case of Bushveld) and ultrametamorphism of reworked derivatives either of the mantle or of the initial lithic crust were accumulated under geoenvironmental conditions that prevailed in the transition period between the lithic crust and the unfolding of the geological spiral. This was a vast period of time in which iron and chrome derivatives of the lithic crust or mantle were available and where the conditions might have been different. These conditions contradict the principle of uniformitarianism. The accumulations thus formed under subsequent ultrametamorphism resulted in the iron- and chromite-rich bands found in the layered ultrametamorphic bodies. Other complex formations are the banded iron formations (Precambrian age) which are entirely different to chromite bands in ultrabasics and therefore not comparable (see Chapter 14). They have been formed under different geoenvironmental conditions than those prevailing nowadays (see Precambrian Iron Formations (1987), and Ancient Banded Iron Formations (1990), Theophrastus Pubs., Chapter 35). In the current attempt only some selected patterns will be presented illustrating the relation of iron oxides and silicates in gabbroic and pyroxenitic rock types mainly from Skaergaard. Fig. 13 shows pyroxenes and feldspars corroded and partly engulfed by later magnetite. Fig 14 also shows olivines, mica and plagioclases enclosed by magnetite which occupies the intergranular spaces and which is most probably a poikiloblast (enclosing, corroding and partly replacing the enclosed feldspars and mafic min-

erals). These patterns contradict the interpretation that differentiation and gravitational settling of the crystals (minerals) has taken place since magnetite, considerably heavier than olivines, pyroxenes and feldspars, did not settle as distinct crystals and does not form a distinct magnetite band (bands consisting of magnetite crystals). However, it is a later crystalloblast-poikiloblast, surrounding and corroding the previously formed "metasedimentary-metamorphic" mineral components which it enclosed. Additional patterns show plagioclase corroded and replaced by later magnetite which partly encloses the feldspar (see Figs. 15 and 16). In other cases a pyroxene rim is formed between the plagioclase and the magnetite which might represent a reaction margin (or infiltration of solutions forming a new pyroxene growth, see Figs. 15 and 16). Furthermore, examples are presented where later magnetite infiltrates and replaces feldspars. This case also shows that the magnetite does not represent a distinct crystal formation which was settled gravitationally from the basalt melt due to magmatic differentiation. The patterns presented so far from the relatively iron-rich bands of the Skaergaard Complex indicate that the magnetite is either a crystalloblastic-poikiloblastic phase or may infiltrate and replace the pyroxenes (see Figs. 17 and 18). In other cases within the Skaergaard Complex magnetite might exhibit rounded or developed crystal forms and is enclosed by later crystalloblastic pyroxene (sometimes a new crystallization pattern of pyroxene might develop between the crystalloblastic pyroxene and the magnetite, either as a result of solution "infiltration" or as reaction margin, see Fig. 19). Magnetite in theralites is sometimes corroded, rounded and enclosed by later blastic biotite in which magnetite granules are interspersed (see Fig. 20). These patterns are also incompatible with the gravitative settling of crystals differentiated and separated by the magmatic differentiation of an initial basic intrusion. Often the magnetite exhibits complex sympletic intergrowths either with hypersthene or other silicates (Figs. 21 and 22) in which case the magnetite does not represent an ex-solved phase or an eutectic crystallization with the pyroxenes but a pattern most probably due to solution-infiltration and replacement. As far back as 1979, Augustithis pointed out that dunitic rocks are mainly characterized by the presence of chromospinel and that pyroxenitic rocks with magnetite or titanomagnetite exist where in addition to the prevalence of Fe spinels titanium minerals are also present (ilmenite). Cases exist (Rodiani, Greece) where Ti (in form of oriented lamellae and grains of rutile) is abundant with the chromite (see Fig. 295 and the description there). These geochemical mineralogical incompatibilities of Ti minerals with chromite (with the exceptions mentioned) are better understood consider-

ing that while dunites are rich in Mg (with limited fyalite crystalline molecules in the forsterite-rich olivines), pyroxenitic-gabbroic rocks are comparatively richer in Fe and Al. Also taking into consideration the geochemical relationship of Ti with Fe and Al, the preference of Ti minerals to associate with the magnetite of the pyroxenitic and gabbroic rocks is understood. The opinion is widespread among petrologists that basalts are mantle derivatives. However, experimental studies of petrofabrics by Augustithis and Kostakis (1980) showed that lherzolitic mantle (olivine bombs) and other rocks, believed to be direct or indirect derivatives of the mantle on melting and subsequent cooling, gave mineralogical compositions and textures that were greatly different from basalts. In contradistinction, a gabbroic rock found in the olivine basalts of the Ethiopian rift similarly treated gave patterns comparable and commensurable to basaltic rocks. This gabbroic rock from the volcanoes of the Ethiopian rift is believed to be a representative of the protolytic layer (between mantle and crust). The fact that many basalts contain olivine bombs as xenoliths indicates that fragments of the mantle were involved in the friction zone that melted the protolyte and produced the initial basaltic melt which on ascending assimilated parts of the crust (Augustithis, 1978). Thus, it is not surprising that mantle fragments are present in basalts and that in cases spinels of the mantle occur as freed or partly freed xenocrysts in the basaltic groundmass. Fig. 23 shows a spinel of an olivine bomb enclosed and affected by the basaltic melt (Fig. 23 and the description there). In other instances chromospinels, parts of olivine bombs in basalt, react with the basaltic melt and a magnetitic margin is produced on the chromospinels. Both the magnetitic margin and the chromospinel show "myrmekitic intergrowths" as a result of the synantetic reaction of chromospinel-magnetite with the basaltic melt (groundmass), see Fig. 24 and 25. In contrast, when a basalt conduit transversed a dunitic complex as it is the case with basalts partly covering the ultrabasic complex of Yubdo (Augustithis, 1965), it is possible to find chromite xenocrysts in the basalt exhibiting a reaction rim magnetitic in composition. In opposition, titanium-rich magnetites as phenocrysts might be present due to the solidification of the basaltic melt (Fig. 26), or magnetite partly idiomorphic as inclusions in the pyroxene phenocrysts (Fig. 27). In some cases though, magnetite segregations may occur within the basaltic groundmass, in others idiomorphic apatites are associated with these magnetite segregations (concentrations) and sometimes the magnetite is even symplectically intergrown with the apatite (Figs. 28 and 29, respectively). Considering the ore minerals that are derived from mantle depth, the "rounded" ilmenites are of interest that occur in kimberlites (Figs. 30 and 31) which show the presence of Mg in this ilmenite. These ilmenites are believed to be mantle derivatives, however, it is 21

difficult to reconcile the fact that Ti tends to be more enriched in the pyroxenitic rock-types than in the dunitic lherzolitic that are more typical for the mantle. In

22

this connection, it should not be forgotten that Ti and Al are light elements and as such would be expected to occur in the upper part of the earth's layering.

Chapter 5

(a) Replacements

Replacement Patterns and Processes

(general)

In an attempt to present the phenomenology of replacements it is of fundamental importance to show typical textural patterns as well as to bring up cases which can be considered as examples demonstrating how the replacement process takes place. Quoting Edwards (1960) once more: "Hardly one of the known textures or structures separately seen is a safe criterion against metasomatism-replacement", it is necessary to consider whether there is a convergence of replacement phenomenology and apparent similar textural patterns that simulate replacement being, however, the result of other processes, such as ex-solutions, eutectic crystallizations, etc.. As already mentioned by the author, cases of apparent convergence phenomenology such as graphic-like symplectites in rhyolites and in lunar samples as well as graphic intergrowths in pegmatites and granites proved to be the result of different processes and not to represent true convergence (Augustithis, 1973, 1982). Likewise, in the case of ore mineral textural patterns similar studies and considerations must be made in order to clarify whether convergence of phenomenology took place, meaning, whether similar or apparently identical textural patterns can be produced either by replacement, by ex-solving or by eutectic crystallization. No a priori conclusions can be drawn to decide whether similar or identical textural patterns represent convergence phenomenology of fundamentally different processes. It is therefore important to consider a wide gamut of textures belonging to a textural pattern before concluding on its genesis and crystallization conditions. Thus, in addition to the principle that each case merits its own consideration, the interpretation of a pattern must be approached inductively, i. e., by considering as many cases as possible. The presentation of replacement phenomenology is a difficult task since replacement textural patterns are not aberrant cases but represent the most abundant textural patterns in most paragenetic associations. The understanding of replacements and intergrowths can help to decipher the crystallization sequence and thus genetic conclusions can be attained. Since replacement patterns involve most of the mineral species and innumerable cases theoretically exist,

making a systematic presentation impossible, it is perhaps feasible to consider some of the most common replacement intergrowths and, if possible, bilateral cases. In contrast to the replacement interpretations provided in this volume, Stanton (1966) claims that many of the textures said to be indicative of replacement may be interpreted as the result of the achievement of minimum interfacial free energy during natural growth. However, the author believes that it should be left to the reader to judge whether the textural patterns presented in this volume represent replacements, or are the result of the achievement of minimum interfacial free energy during natural growth.

(b) Replacement Patterns Involving Pyrrhotite, Ilmenite and Titanite Pyrrhotite is believed to be a high temperature mineral formation and is often considered to be a magmatic or hypothermal crystallization. FeS is often subjected to alteration processes and an interesting case of pyrrhotite replacement by magnetite is illustrated in Fig. 32. The magnetite marginally replaces the pyrrhotite along cracks of the sulfide as well. In nature, this replacement process is quite complex and involves a breakdown of the pyrrhotite lattice and a release of S prior to the oxidation of Fe. Whether the replacement process takes place at lower temperatures than the crystallization of pyrrhotite would require, or whether sulfur is released and the Fe oxidation takes place at relatively low temperatures is difficult to determine in natural cases. Furthermore, O'Meara (1961) reported replacement of pyrrhotite by lamellar marcasite and magnetite, or pyrite and (hypogene) carbonate, followed by deposition of chalcopyrite and further pyrrhotite and by assemblages of lower temperature sulfides and sulfosalts. In contradistinction to Fig. 32, Fig. 33 shows the replacement of pyrrhotite following a silicate veinlet partly transecting the pyrrhotite. The magnetite marginally replaces the pyrrhotite in the silicate veinlet and often forms a myrmekitic-like intergrowth with the pyrrhotite. In this case, replacement of pyrrhotite by magnetite involves the release of sulfur and the oxidation of Fe. More complex replacement cases are pre23

sented by Polferov and Suslova (1970), where pyrrhotite, pentlandite, chalcopyrite and pyrite are replaced by magnetite during autoserpentinization and subsequent carbonitization. In contradistinction to the above mentioned examples, magnetite and titanite can marginally replace ilmenite along cracks as well, as demonstrated in Fig. 34 and in particular, in Fig. 35. It shows clearly that marginal magnetite is extending and replacing the partly engulfed ilmenite. In this case, the release of Ti (dissolved) also took place and perhaps reduction instead of oxidation of Fe occurred, making the replacement of ilmenite FeTiOs by magnetite Fe 3 0 4 possible. The fact that ilmenite is also replaced by titanite (CaTiSis) complicates the replacement processes of ilmenite further, particularly when a concurrent ilmenite-magnetite and ilmenite-titanite replacement occurred. The replacement of ilmenite in addition to Fe and Ti also involves Si and Ca. Perhaps a dissolution of the ilmenite took place prior to the crystallization of the tinanite. Besides the cases of the ilmenite replacement described, ilmenite with haematite (?) ex-solution lamellae are replaced by magnetite with the magnetite bodies oriented parallel to the haematite lamellae in the ilmenite. The reverse case of ilmenite replacing magnetite is also often observed. However, this process is very complex since it involves spinel formation following the contact reaction of the ilmenite replacing the magnetite, as demonstrated in Fig. 36. As a consequence to the replacement of titanomagnetite by ilmenite, titanomagnetite with oriented lamellae of spinel is marginally replaced by ilmenite (Fig. 37). The formation of symplectic spinel follows the contact of ilmenite with the magnetite. It should be pointed out that this pattern becomes even more complex since it also involves chalcopyrite veinlets transecting the ilmenite. The spinel lamellae are oriented parallel to the [100] plane of the magnetite. Additionally, titanomagnetite was observed with a replacement margin of ilmenite which in turn has a margin of spinel with some extensions into the ilmenite. The replacement contact of magnetite by ilmenite in this case is also followed by spinel often simulating myrmekitic patterns.

(c) Replacement tute)

Patterns (Pyrite as a Substi-

In contrast to magnetite replacing pyrrhotite and ilmenite, magnetite itself is often replaced by sulfides and an extensive phenomenology of magnetite replaced by pyrite is exhibited. Fig. 38 shows pyrite megablast replacing magnetite. The restricted magnetite relics in the pyrite often simulate a network. In addition to this network type, examples of magnetite replaced along crystallographic directions by pyrite are also found. Occasionally, maghemite is also present as veinlets along silicates. 24

In contradistinction to the patterns of magnetite replaced by pyrite megablasts, cases of tectonogranular magnetite replaced by pyrite occupying intergranular spaces are exhibited (see Figs. 39, 40 and 41). Another interesting type of magnetite replacement by pyrite is shown in Fig. 42. Atoll-type magnetite replacement by pyrite is common and does not represent pyrite with marginal magnetite. There are also cases where pyrite seems to occupy and replace spaces or silicates within the magnetite; such patterns are rather rare and could represent pyrite fillings in spaces of magnetite (Fig. 43). In addition to the rather simple examples of magnetite replaced by pyrite discussed so far, more complex patterns are found when the same magnetite crystal is replaced by myrmekitic-like pyrite or by maghemite (Fig. 44). Replacement patterns are often very complicated as can be understood from the distribution of relic structures. Fig. 45 shows magnetite initially associated with ilmenite, replaced by later pyrite. Relics of the ilmenite are restricted to follow the [100] face of the pyrite. Despite being a rarely seen phenomenon, the incorporation of the relics in crystallographic directions of the new growth by replacement is theoretically conceivable, in the sense that the relics can be accommodated in the interzonal spaces of a mineral phase replacing a pre-existing one. In opposition to cases where relics might be incorporated along crystallographic directions of the replacement-growth, crystallographically oriented ilmenite lamellae in magnetite are observed. The magnetite itself might be replaced by later pyrite in which case ilmenite lamellae partly protrude into the pyrite and relics of the replaced ilmenite are present in the sulfide (Fig. 46). Other replacement patterns of magnetite with ilmenite lamellae are shown in Fig. 47. Pyrite replaced titanomagnetite and as a result, the ilmenite lamella is left in relic form in the pyrite. The pyrite also replaced the interlamellar magnetite. The most interesting patterns of enargite replacement by pyrite are shown in Fig. 48. Considering the general formula of enargite Cu3AsS4 its replacement by pyrite involves the dissolution of Cu and As and the stoichiometric movement of Fe in the Bereich of the pyrite formation. More dubious patterns though of enargite-pyrite replacement are exhibited in Fig. 49. Most probably these patterns represent incomplete development of a pyrite idioblast rather than enargite invading the pyrite. Concerning the replacement of sulfides, textural patterns of berthierite (FeSb2S4) replaced by pyrite (FeS2) are of interest. They could be understood either by the removal of Sb and the reorganization of the lattice and possibly some S removal, or, more likely, by dissolution of the berthierite prior to the precipitation of pyrite (Fig. 50). Additional complex patterns of berthierite replacement by pyrite are illustrated in Fig. 51. Here, berthierite is marginally corroded and replaced by pyrite, also indicating

atoll-type replacement of berthierite by pyrite. In other instances the pyrite replacing berthierite attains idioblastic shapes (Fig. 52). Present textural studies indicated that a wide range of minerals can be replaced by pyrite. The replacement might occur after dissolution of a pre-existing mineral and the occupation of the space might take place without leaving voids. In other cases, dissolution and replacement of only some of the elements take place. The role of sulfur is hard to trace when one sulfide is replaced by another, since it is difficult to prove whether the sulfur remained in the framework in which element substitutions happened, or whether the dissolution of the sulfide involved the removal of the sulfur of the replaced mineral prior to its substitution by another sulfide. Cassiterite, Sn0 2 , is another mineral that can be replaced by pyrite. In this case, dissolution of the cassiterite most probably occurred prior to its substitution by pyrite (see Fig. 53). Low temperature pyrite (Wasserkies) often replaces skeletal pyrargyrite (Ag 3 SbS 3 ); in this case S is present in both minerals. Fig. 54 shows skeletal pyrargyrite partly replaced by pyrite. Similarly, pyraryrite shows partial replacement by Wasserkies pyrite.

(d) Replacements of Magnetite by Chalcopyrite, Pyrrhotite, Enargite and Boulangerite Besides the replacement of titanomagnetite by pyrite, replacement by chalcopyrite and pyrrhotite can also occur in which case the oriented ilmenite lamellae are left as relics in the sulfide. Apparently, there is a preferential substitution of magnetite by the sulfides in comparison to the ilmenite lamellae resisting the replacement. Fig. 55 shows titanomagnetite replaced by chalcopyrite and pyrrhotite with oriented ilmenite lamellae preserved as relics. Similarly, ilmenite bodies are left as relics of initial titanomagnetite replaced by pyrrhotite (see Fig. 56). In both cases, Fe is present in the initial magnetite and in the substitute minerals, i. e., chalcopyrite and pyrrhotite. Again it is difficult to trace what happened to the iron of the incipient magnetite and whether any of it is present in the newlyformed sulfides. In opposition to the described examples, magnetite can be replaced by silicates associated with enargite (CU3ASS4). Here, complete dissolution of the magnetite must have preceded its replacement by the silicates and the enargite (Fig. 57). In addition, concerning the replacement of magnetite by chalcopyrite, the following observations are of particular interest. Fig. 58 shows chalcopyrite with extensions attaining veinform shapes and replacing an adjacent magnetite. The pattern of the atoll replacement of magnetite by chalcopyrite is very abundant (Fig. 59). In some instances though the patterns of chalcopyrite replacing magnetite indicate a more complex pre-his-

tory, in the sense that the magnetite was derived from initial iron hydro-oxides (hydrogels) which were transformed-metamorphosed into the magnetite. Selective replacement of the initial gel bands of the magnetite were then replaced by chalcopyrite. In the above mentioned cases magnetite (Fe 3 0 4 ) was replaced by chalcopyrite (CuFeS 2 ) which most probably involved dissolution or partial dissolution of the magnetite in the way that Fe was maintained in the microenvironment to which Cu and S were added. It is also possible that complete dissolution occurred prior to the precipitation from solutions of chalcopyrite. Of course, replacement of elements and element mobilization processes should not be excluded from the discussion. Further discussion of replacement processes involving magnetite, leads to the replacement by boulangerite (Pb 5 Sb 4 S n ). It is possible that boulangerite occupies either void spaces of the magnetite or replaces silicates which are in intergrowth with the magnetite (Fig. 60).

(e) Replacement by Copper

Minerals

Most interesting and in cases very complex replacement patterns are exhibited by bornite (Cu 5 FeS 4 ) which can replace a wide variety of minerals. Fig 61 shows a complex myrmekitic-like patterns of rutile in intergrowth with silicates marginal to chalcopyrite, invaded and replaced by bornite (restricted replacement by chalcopyrite also occurred). The replacement of rutile (Ti0 2 ) by bornite involves the complete dissolution of the rutile and the introduction into the theoretical void formed of a precipitate of bornite solutions. The way the process takes place is not yet clarified, however, no voids seem to be left. Correspondingly, Fig. 62 illustrates additional examples of the rutile-silicate symplectite that was replaced by bornite. A very common case is also the replacement of pyrite (FeS2) by bornite (Cu5FeS4) which theoretically involves the introduction of Cu, although the stoichiometric balance of the complex might be very complex. As the pattern shown in Fig. 63 might indicate, the pyrite was probably partly dissolved prior to the formation of the bornite. It may though represent another case of mineral dissolution and precipitation of a new mineral phase from solutions. It is difficult to say whether the dissolution occurred concurrently with the precipitation of the new mineral phase or whether a void intervened. When discussing the often characteristic replacement sequence in a supergene zone or a copper enrichment sequence, the following patterns indicating replacement need to be further considered. Fig 64 shows chalcopyrite replacing bornite, the bornite in turn being replaced by covellite (CuS) and a network of fine anastomosing covellite veinlets extending into 25

the bornite. In the covellite-mass neo-crystallization of chalcocite and even chalcopyrite is also present. It seems that the following replacement and neo-crystallizations are the reason for the formation of the textural patterns illustrated in Fig. 64: bornite was replaced by chalcopyrite and also by covellite and neo-crystallizations of chalcocite and chalcopyrite took place. Another interesting textural pattern indicating a sequence of copper enriched minerals is shown in Fig. 65 where pyrite with idiomorphic tendency is surrounded by copper minerals with bornite being formed first and a symplectic intergrowth with chalcopyrite. The chalcopyrite most probably replaces the bornite and laterformed chalcocite (Cu2S) replaces pyrite , bornite and chalcopyrite. In addition to the patterns discussed above, cases are illustrated where bornite (Cu5FeS4) is replaced by stromeyerite (AgCuS), a replacement involving the mobilization of Ag (see Fig. 66). The sequence of the copper replacement series can indeed be variable. Fig. 67 shows enargite (Cu3AsS4) replaced by chalcopyrite (CuFeS2) which involves the removal of As and the introduction of Fe. However, probably it is not just an element migration but actually might involve the dissolution of enargite and the introduction of chalcopyrite. The processes might not involve any remaining void spaces. Most interesting rhythmical replacement textural patterns are shown in Figs. 68 and 69. Here tennantite (Cu3SbS3) is replaced by chalcocite (Cu2S) with a cuprite (Cu 2 0) margin. Two alternative explanations are put forward for the rhythmical and consecutive replacement of tennantite by chalcocite and cuprite: (i) Tennantite is corroded and dissolved and a colloform banded chalcocite margin is formed enclosing the tennantite. The chalcocite is followed by colloform in origin cuprite, precipitated from the same solution as the chalcocite. (ii) Sb is leached out of the tennantite and as a result of subsequent stoichiometric, and perhaps volume readjustment chalcocite is formed. Due to the subsequent S leaching out and oxidation, cuprite is formed. As demonstrated in Fig. 69, corroded outlines of tennantite are exhibited and extensions of the marginal chalcocite grow as veinlets into the tennantite suggesting dissolution of the tennantite and its subsequent replacement by later chalcocite forming solutions of which cuprite was also precipitated. Often complex patterns occur as a result of replacement and infiltration processes. Fig. 70 shows pentlandite (Fe, Ni)S partly surrounded and corroded by chalcopyrite (CuFeS2) which extends and replaces the pentlandite along cleavage directions. It is tentatively suggested that pentlandite was dissolved and replaced by chalcopyrite. The alternative interpretation is that such replacements can take place by migration of elements, meaning that Ni removal and copper introduction are within the scope of possible processes that could happen. However, the corrosion of the pentlan26

dite and the infiltration of chalcopyrite along the cleavage of the pentlandite are processes that took place as the textural pattern of Fig. 70 indicates. Complex patterns are also signified where chalcopyrite encloses, corrodes and replaces pyrite and simultaneously extends as infiltration, replacement structure into an adjacent sphalerite. The chalcopyrite infiltrations in the sphalerite, symplectic in form (simulating ex-solutions of chalcopyrite in the sphalerite), are shown in Fig. 71. The fact that actually many of the so-called ex-solutions of chalcopyrite in sphalerite are replacement textures of sphalerite by chalcopyrite is supported by a series of observations (see Chapter 6). In particular Fig. 72 shows sphalerite marginally replaced by chalcopyrite and extensions of it attaining "ex-solution" forms in the sphalerite. The most characteristic replacement textural patterns of this type are shown in Fig. 73 where veinlets of chalcopyrite transect the sphalerite and also, chalcopyrite replacing sphalerite is exhibited. As already presented, chalcopyrite displays impressive patterns of replacement by chalcocite. Fig. 74 shows chalcopyrite enclosed and corroded by chalcocite with veinlets of it transecting the chalcopyrite. Considering the formula of chalcopyrite (CuFeS2) and that of chalcocite (Cu2S) removal of Fe and relative enrichment of Cu could stoichiometrically explain the replacement, of course the reorganization of the lattice is necessary. However, in this case too, dissolution of chalcopyrite and subsequent precipitation of chalcocite is a possible alternative. In contradistinction to the cases described, chalcocite veinlets following the intergranular between idiomorphic pyrite and chalcopyrite often extend into the latter as ramified veinlets, most often independently of the presence of pyrite (Fig. 75). Comparable to the cases of tennantite replaced by rhythmically banded chalcocite and cuprite margins, examples are exhibited of chalcopyrite corroded, invaded and surrounded by neodigenite followed by brown iron (Fig. 76). Considering the formula of chalcopyrite (CuFeS2) and that of neodigenite (Cu2S), removal of iron from the chalcopyrite, the resultant relative enrichment of Cu and additionally the reorganization of the crystal lattice might explain the chalcopyrite replacement by neodigenite. It should be noted that the brown iron precipitation may be partly due to the leaching out of Fe from the chalcopyrite. Rhythmical interbanding of neodigenite and brown iron took place, as the textural patterns in Fig. 77 indicate in particular, and possibly a colloidal phase was also involved in the unfolding of the process (see Fig. 78). As another example of replacement of copper minerals, Fig. 79 shows enargite (Cu3AsS4) in intergrowth with silicates replaced by haematite (Fe 2 0 3 ). An atolltype replacement is also indicated. The replacement of enargite by haematite involves dissolution of the enargite and replacement by haematite-forming solutions.

In contradistinction to the above-mentioned case of dissolution of enargite and substitution in the void formed by haematite, examples of one copper mineral replacing another in accordance to the tendency of the substitute to be enriched in copper, are shown in Figs. 85 and 86. (However, the mineral substitutions and new mineral growths involved complicate the matter as will be discussed later.) Fig. 80 shows chalcopyrite (CuFeS2) marginally replaced by covellite (CuS). In contradistinction though to the cases where the substitute is enriched in Cu, Fig. 81 shows chalcocite (Cu2S) replaced by covellite where relative migration (perhaps leaching of Cu) took place. Comparable replacement patterns of neodigenite (Cu2S) by covellite (CuS) are exhibited in Fig. 82. Covellite in addition to replacing other copper minerals sometimes exhibits patterns with non-copper sulfides replaced. A very rare case is pyrite replaced by covellite. Galena is also occasionally replaced by covellite, as shown in Fig. 83. Additional to the replacement by covellite, galena can be replaced by Sbcontaining minerals such as falkmanite (Pb 5 Sb 4 S n ). (This replacement pattern is mentioned here to compare and contrast the behavior of Sb and Cu in the case of galena replacement.) The replacement of galena by covellite and falkmanite is a complex process. In the first case, Pb is removed and copper introduced, in the second case the introduction of Sb occurred with lattice re-adjustment (reorganization of the lattice). As a further example for possible replacement of ore minerals by covellite, Fig. 84 shows boulangerite (Pb 5 Sb 4 S n ) being replaced by covellite (CuS). In contrast to the concept that replacement of copper minerals by other copper minerals leads to copper increase, complicated examples are illustrated in Fig. 85 and 86: Fig. 85 shows bornite (Cu2FeS4) replaced by covellite which also extends as a fine network into the bornite. Within the covellite though neo-crystallizations of chalcocite (Cu2S) and chalcopyrite (CuFeS2) took place, thereby replacing the covellite. This in tum means topological enrichment of Cu in covellite to produce chalcocite and in the case of neo-crystallization of chalcopyrite, Fe introduction and reorganization of the lattice. In addition to the above mentioned replacement neocrystallizations presented in Fig. 85 and 86, cases of replacement neo-crystallization of non-copper minerals can take place in copper sulfides. Fig. 87 shows a nucleus of pyrite corroded and replaced by chalcocite. In the marginal parts of the chalcocite, neo-crystallizations occurred. Considering the formula of chalcocite (Cu2S) and that of pyrite (FeS2), removal of Cu and introduction of Fe must have taken place topologically in the Bereich of pyrite neo-crystallization in the chalcocite. With the replacement of copper minerals, complex mobilization processes can take place. As Fig. 88

shows, cuprite (Cu 2 0) is replaced by malachite exhibiting colloform structure. Considering the formula of malachite (Cu 2 [(0H) 2 /C0 3 ] and of cuprite (Cu20), the replacement involves dissolution of cuprite and precipitation from colloidal solutions of malachite since colloform sphaeroids are found in the cuprite as replacement structures. Textural patterns indicating marginal replacement of cuprite or cuprite crystals replaced by malachite are also abundant (see Figs. 89 and 90). Comparable patterns to the Cu-oxide (cuprite) replacement by malachite are found with Cu-sulfides replaced by malachite. Considering the formula of chalcocite (Cu2S) and that of malachite (Cu 2 [(0H) 2 /C0 3 ]) removal of sulfur, perhaps dissolution of chalcocite and precipitation of malachite from solutions occurred. However, it should be pointed out that in these complex processes Cu is present both in the dissolved and the substitute mineral. Considering the replacement of ore minerals, and in particular the copper minerals' textural pattern, Fig. 91 shows an ultimate phase as far as copper enrichment is concerned in the substitute mineral since the copperrich oxide cuprite is replaced by native copper. Many complex replacement processes and textures are exhibited where tetrahedrite (Cu6Sb2S7) marginally replaces bornite (Cu2FeS4), removing Fe and introducing Sb with stoichiometric re-adjustment and reorganization of the lattice. Beyond it, other minerals are often replaced by tetrahedrite. Figs. 92 and 93 show niccolite (NiAs) replaced by tetrahedrite, showing that dissolution of the niccolite took place prior to the precipitation of tetrahedrite. In some ways comparable is the replacement of pyrite (FeS2) by tetrahedrite where dissolution of pyrite preceded the precipitation of tetrahedrite. However in this case, S is a common element in both the replaced and the substitute mineral. Besides the patterns already presented where tetrahedrite replaces bomite and niccolite, additional observations show tetrahedrite (Cu8Sb2S7) replacing germanite [(Cu3(Ge,Fe)S4)]. Considering the general chemical formulae and the textural patterns as shown in Fig. 94, germanite is enclosed, corroded and replaced by tetrahedrite, thus supporting the interpretation that dissolution of germanite was followed by precipitation of tetrahedrite. Very complex patterns of replacement are illustrated in Fig. 95 which shows renierite (approximate formula (Cu,Fe,Ge,Zn)S) included, corroded and replaced by germanite (Cu3(Ge,Fe)S4) which in turn is enclosed and corroded by tennantite (Cu3SbS3). Furthermore, as a late phase crystallization, galena (PbS) replaces both the germanite and the tennantite. Such complex patterns most probably involve dissolution of some minerals and precipitation of new ones from solutions, although element migrations as an alternative or supplemental process should not be excluded categorically. 27

In addition to the pattern shown in Fig. 95, Fig. 96 shows sphalerite (ZnS) marginally replaced by germanite (Cu 3 (Ge,Fe)S 4 ) and with renierite replacing the sphalerite. It should be noted that veinlets of germanite and renierite transect the sphalerite. As the last phase of crystallization in the pattern illustrated in Fig. 96, tennantite replaces the sphalerite as well as the renierite and the germanite. The fact that veinlets of germanite and renierite transect the sphalerite support the interpretation that dissolution of sphalerite preceded its replacement by germanite and renierite. Similarly, the replacement of renierite by tennantite is also due to dissolution and subsequent precipitation from solutions. As already pointed out tennantite (Cu3SbS3) is replaced by galena (PbS). Also as relic structure in the galena, preserved renierite with a margin of germanite is present. This pattern should be understood in conjunction with the pattern illustrated in Fig. 95. Concerning the replacement of sphalerite by tetrahedrite, Fig. 97 shows sphalerite (ZnS) corroded and replaced by tetrahedrite (Cu 8 Sb 2 S 7 ) in which case dissolution of sphalerite has most probably taken place prior to the precipitation from solutions of tetrahedrite.

( f ) Replacement

Patterns of Pyrite

In replacement processes pyrite is often found to replace magnetite. Pyrite is a Durchläufer, a mineral formed under most variable conditions. Thus, it is not surprising that pyrite itself is subjected to complex replacement processes and exhibits a very complex phenomenology. Fig. 98 shows pyrite invaded and replaced by sphalerite. The sphalerite invades the pyrite along cracks and extends from its margins inwards. Comparing the formula of pyrite (FeS2) and of sphalerite (ZnS) it is interesting that theoretically (or stoichiometrically) the pattern could be explained by a substitution of Fe by Zn. However, it is very likely that the pattern was caused by dissolution of pyrite and the introduction of sphalerite involving the processes of dissolution of a pre-existing mineral and its substitution by another without leaving any voids. However, the most complex patterns of pyrite replacement are exhibited when pyrite is enclosed, corroded, invaded and replaced by copper minerals. Fig. 99 illustrates pyrite replaced along its [100] cleavage by chalcopyrite. The formulae of pyrite (FeS2) and of chalcopyrite (CuFeS2) suggest, by stoichiometric balancing, that the introduction of Cu is necessary. In this case, dissolution of pyrite along its cleavage and introduction of solutions out of which chalcopyrite precipitated, is an even more plausible interpretation. There are also examples where pyrite is enclosed or partly enclosed by chalcopyrite and then corroded and replaced by the later chalcopyrite (Fig. 100). 28

Sometimes the pyrite contains chalcopyrite bodies which simulate inclusions or ex-solutions (Fig. 99). In this case however, the chalcopyrite is the cause of replacement of pre-existing pyrite. Additional observations indicate veinlets of enargite (CU3ASS4) replacing pyrite (FeS2). The pyrite exhibits corroded margins (indentations). The replacement of pyrite by enargite is perhaps understandable by a process of dissolution of the pyrite and the introduction of solutions out which enargite was precipitated (Fig. 101).

Very impressive replacement patterns of pyrite by millerite are shown in Figs. 102, 103 and 104. Fig. 102 shows pyrite invaded and replaced by millerite (NiS with little Co in place of Ni). Similarly, pyrite is corroded and replaced by millerite which often extends into the intergranular of the pyrite (Fig. 102). In some instances extensive pyrite replacement is exhibited, sometimes with pyrite relics left in the millerite (Fig. 105). Comparing the formula of pyrite with that of millerite, replacement of Fe by Ni took place while S most probably was not dissolved. Pyrite is also often enclosed, invaded and replaced by neodigenite. Fig. 106 shows pyrite surrounded and invaded by neodigenite. The replacement involves the dissolution of pyrite and the precipitation of neodigenite from the solutions that invaded the voids formed. However, it should be emphasized that no void spaces are maintained. Similarly, Fig. 107 shows pyrite, often idiomorphs, partly corroded, invaded and replaced by neodigenite. Figs. 108-110 show impressive examples of pyrite replaced by chalcocite (CuS2). Pyrite is surrounded by chalcocite and indicates all transitions from marginal replacement to complete substitution by chalcocite. The pyrite often exhibits the typical corroded outlines and in the case of extensive replacement, minute relics of pyrite might be left in the chalcocite. The pattern most likely indicates dissolution of the pyrite and its subsequent substitution by chalcocite. The alternative interpretation that there was an exchange of Fe and Cu with the necessary re-adjustments of volume, should not be entirely disregarded. In contradistinction to Fig. 108, Fig. 109 shows pyrite transected by veinlets of chalcocite with chalcocite surrounding and corroding the pyrite as well. Often microfractured pyrite is invaded and replaced by chalcocite along the microcracks of the pyrite. As can be seen in Fig. 110, corrosion and replacement of pyrite took place. In opposition to replacement along microfractures, comparable patterns are exhibited when pyrite is surrounded, corroded and replaced along its cleavage directions (see Fig. 111-113). Pyrite can be replaced by molybdenite (MoS 2 ) where the prismatic molybdenite, radiating from a center outside the pyrite extends into the pyrite, corroding and replacing it (Fig. 114). In addition, pyrite is corroded and replaced by molybdenite with pyrite relics left in it (Fig. 115).

Replacement of pyrite by another sulfide is shown in Fig. 116 where realgar forms a margin corroding and replacing the pyrite. Here too, pyrite (FeS2) is replaced by another sulfide (AsS). In this case too, pyrite was corroded and dissolved prior to the precipitation of realgar. Similarly, Fig. 117 shows pyrite tectonically fractured with realgar occupying the cracks and occasionally a fragment of pyrite is enclosed, corroded and replaced by the later realgar. In cases, the fractured pyrite is replaced along the cracks by orpiment (As2S3). Sinkovec (1960) described cinnabar (HgS) enclosing pyrite and replacing sphalerite but it is cut by veinlets of the later sphalerite. The replacement of pyrite by molybdenite, realgar and orpiment poses the question whether in such cases the replacement is achieved by dissolution or migration of the cations of Fe and their respective substitution by Mo or As, of course with the corresponding volume re-adjustment. Here it has to be assumed that the sulfur remained as the framework (background) where the cation exchange occurred. As mentioned, pyrite is often replaced by chalcocite. In Fig. 118 both pyrite (FeS2) and enargite (Cu2AsS4) are replaced by chalcocite (Cu2S). The replacement of pyrite by chalcocite was discussed earlier in this volume. The replacement of enargite by chalcocite either represents a dissolution of the enargite and precipitation of chalcocite from later solutions, or it could be interpreted as the main cause of migration or leaching out of As with, in this case, a possible re-adjustment of volume and lattice.

(g) Replacement

Patterns of Sphalerite

Characteristic symplectites of pyrite with sphalerite are often exhibited where crystalloblastic pyrite replaces the sphalerite relic structures that simulate ex-solution bodies in pyrite. Comparing the formulae of sphalerite (ZnS) and pyrite (FeS2), two alternative interpretations of the process could be suggested: 1. Sphalerite was dissolved and replaced by crystalloblastic pyrite, that means, a dissolution of blende preceded the blastic pyrite formation, or 2. Zn was substituted by Fe (with corresponding stoichiometric and volume re-adjustments). With this alternative interpretation, substitution and migration of elements in the lattice should theoretically be considered. In addition to the replacement of sphalerite by chalcopyrite and pyrite, sphalerite can be replaced by a number of other minerals and several textural patterns will be presented. Fig. 119 shows sphalerite marginally replaced by marcasite (FeS2). Here too, the sphalerite is dissolved and replaced by solutions. Marcasite is also precipitated from solutions, independent from sphalerite. In contradistinction to the marginal replacement of sphalerite by marcasite, examples of

sphalerite corroded, included and reduced to relic size are illustrated in Fig. 120. In contrast to the marginal type of sphalerite replacement, complex atoll-type replacement textural patterns of sphalerite (ZnS) by safflorite (CoAs2) are shown in Fig. 121. The safflorite consists of radiating and star-shaped twins often protruding into and replacing the sphalerite. Comparing the formulae of sphalerite and safflorite, the replacement processes involve a removal of Zn and S and the introduction of Co and As. However, it is more likely that dissolution of the sphalerite preceded the replacement by safflorite, i. e., dissolution and subsequent precipitation from solutions of the safflorite. Complex patterns of replacement are observed when sphalerite is replaced by bornite (Cu s FeS 4 ) and chalcocite (Cu2S). Such complex, multiple replacement processes and the resulting textures could perhaps be interpreted as dissolution of sphalerite and precipitation of bomite and chalcocite from solutions following the sphalerite dissolution. In Fig. 122-124 supplementary complex replacement patterns are exhibited. Fig. 122 shows the replacement of sphalerite by bornite and chalcocite. Also, pyrite with idioblastic tendencies is present. Here, the replacement of the sphalerite by bornite and chalcocite is also interpreted as a cause of dissolution of the sphalerite and subsequent precipitation of bornite and chalcocite from solutions; possibly bomite preceded the chalcocite (Fig. 123). In an attempt to show the wide gamut of sphalerite replacement, examples are presented where sphalerite (ZnS) is replaced by arsenopyrite (FeAsS) which replaces the zincblende marginally and often exhibits a tendency for idiomorphism.

(h) Galena as a Substitute (Replacements by Galena and Other Lead-Containing Minerals) Galena is a mineral exhibiting complex replacement patterns because it is late-crystallizing and affects preexisting minerals which it often replaces as a result. Galena often marginally replaces pyrite which is enclosed, corroded and replaced by PbS. Fig. 125 shows galena marginal to pyrite, corroding and replacing it. Additionally, extensions of galena into pyrite can be observed which simulate inclusions, but actually are just protuberances of galena in the pyrite, appearing as inclusions due to the section of the intergrowth. In contrast to the marginal galena replacing pyrite, idiomorphic pyrite is enclosed, corroded and replaced by galena (Fig. 126). Another example is galena exploiting the cleavage and possibly crystallizing in the outlines of the pyrite; the resultant replacement texture is idiomorphic in shape with pyrite partly enclosed in the later galena. Interesting selective replacement textures of pyrite replacement by galena are shown in Fig. 127 where 29

partial zonal replacement occurred. Cases are also observed where marginal galena replacing pyrite is relatively restricted to the outer zone of the pyrite (Fig. 128). However, in other instances the pyrite is transected by fine galena veinlets often replacing the pyrite along cracks. Considering the composition of pyrite (FeS2) and galena (PbS), the replacement involves the removal of Fe and the introduction of Pb. However, here it is possible that the dissolution of pyrite preceded the precipitation of galena. The dissolution and precipitation might represent advancing fronts of solutions. Replacement patterns of pyrite and pyrrhotite (FeS) by galena are exhibited in Figs. 129 and 130. The replacement of pyrrhotite occurred along crystallographic directions of the FeS as a result of galena simulating oriented intergrowths in the pyrrhotite. Comparing the composition of pyrrhotite and galena, removal of Fe and introduction of Pb occurred. Here again though, dissolution of the iron sulfide was prior to the precipitation of galena. Fig. 131 shows arsenopyrite (FeAsS) marginally replaced by galena (PbS) in which case (and in many instances of replacement) the limit reached was marked off by the crystal faces of the arsenopyrite. Replacement here too took place by dissolution of the arsenopyrite and precipitation of galena from solutions. Fig. 132 shows arsenopyrite corroded and replaced by galena with galena infiltration-replacements in the arsenopyrite, suggesting again dissolution of arsenopyrite and precipitation of galena in the void spaces. Lawrence (1962) reports that owyheeite (Pb 5 Ag 2 Sb 6 S 14 ) is relatively abundant in arsenopyriterich samples from Wongabah Mine, characteristically forming aggregates of fibrous, bladed crystals with a fluidal texture, and replacing arsenopyrite and pyrite to some extent. Acicular crystals of owyheeite occur as replacive inclusions in arsenopyrite and rarely in tetrahedrite (Cu 3 SbS 3 or Cu 3 SbS 4 ). Galena exhibits characteristic replacement patterns with copper minerals. As examples, the replacement of chalcopyrite (CuFeS 2 ) and cubanite (CuFe2S3) by galena are put forward. As Fig. 133 shows, chalcopyrite is invaded, corroded and replaced by galena. Similarly, Fig. 134 shows cubanite replaced by galena, which follows crystallographic directions of the cubanite, resulting in a parallel-oriented intergrowth of cubanite and galena. Perhaps most impressive and genetically significant are the replacements of sphalerite (ZnS) by galena. Fig. 135 shows sphalerite in intergrowth with silicate, replaced marginally by galena which attains symplectic intergrowth simulating ex-solution bodies or rather myrmekitic-like intergrowths of galena and sphalerite. In contradistinction, Fig. 136 shows galena replacing sphalerite and the pattern simulates an eutectoid (see Chapter 2). The most typical patterns are found when galena has extensions to the adjacent sphalerite as shown in Fig. 30

137. Occasionally, the extensions of an adjacent galena assume symplectite-like forms in the neighbouring sphalerite. These texture patterns resemble ex-solution bodies of galena in sphalerite. Fig. 138 and 139 show galena replacing sphalerite and the intergrowth resembles ex-solutions of galena in the sphalerite. Replacement veinlets consisting of galena transecting sphalerite are shown in Figs. 140 and 141. Sometimes galena forms a thin margin on silicates in intergrowth with the sphalerite. As can be seen in Fig. 140, the fine galena margins on the silicates are extensions of an adjacent galena. It should be noted that the galena replacement consists either of fine margins of the silicates in intergrowth with sphalerite or as very fine galena veinlets in the sphalerite independent of the silicate (Fig. 141). When the formulae of sphalerite (ZnS) and galena (PbS) are compared the replacement process involves either the substitution of Zn by Pb or most likely, dissolution of sphalerite and introduction of PbS. Complex replacements are observed when galena replaces sphalerite and tennantite (Cu 3 SbS 3 ) as can be seen in Fig. 142. However, the replacement of tennantite by galena involves dissolution of the tennantite and the introduction of galena as precipitate from solutions (Fig. 143). Fig. 144 and 145 show neodigenite (Cu2S) replaced by galena. Sometimes atoll-type replacement of germanite [Cu3(Ge,Fe)S4] by galena also occurred. In the case of the replacement of neodigenite by galena, Cu was removed and Pb introduced. In opposition, in the case of germanite replaced by galena, Cu, Fe and Ge were removed and Pb introduced. In both cases stoichiometric re-adjustments with the appropriate volume changes did possibly happen in addition to lattice reorganization. The alternative interpretation dissolution of both neodigenite and germanite and introduction of galena by solution - should not be discredited since such dissolution and replacement processes by solutions occupying the voids are common and widely accepted. It is possible that dissolution and occupation by new minerals of the voids resume a front replacement process, in the sense that the dissolution of one mineral is followed by the precipitation of the substitute. Replacement processes of pyrite and sphalerite by jamesonite (Pb 4 FeSb 5 S 14 ) are more complex processes as far as element mobilization is concerned. As Figs. 146 and 147 show, dissolution of pyrite and precipitation of jamesonite are the most probable processes. Similarly, Fig. 148 shows selective zonal dissolution of sphalerite and precipitation of jamesonite took place in the way indicated. Additional examples of sphalerite replacement by jamesonite are illustrated in Fig. 149 where the jamesonite invaded and replaced the sphalerite. In this case, too, dissolution and precipitation occurred. Additional to the impressive textural patterns obtained by replacement of sphalerite by chalcopyrite,

galena and jamesonite, other sulfides such as gratonite (Pb 9 As 4 S 15 ) can also replace colloform sphalerite (see Fig. 150). Here also dissolution of the sphalerite and precipitation of gratonite are suggested as possible alternative to element substitution. Replacements by anglesite (PbS0 4 ) are very important. Fig. 151 shows skeletal crystals of pyrite replaced by anglesite. Dissolution of the pyrite (FeS2) and precipitation of anglesite took place. The textural replacement patterns are most complex when anglesite replaces both safflorite (CoAs2) and rammelsbergite (NiAs2). Here the dissolution of the safflorite and the rammelsbergite occurred prior to the precipitation of anglesite from solutions (Fig. 152). Evidence is perhaps the pattern in Fig. 153 that dissolution precedes precipitation of the substitute: carbonates included in niccolite are dissolved and the voids so formed are partly occupied by anglesite which in this case, has selectively replaced the carbonate. A more advanced phase of the process is shown in Fig. 154 where anglesite replaced the carbonate almost completely. In addition to the replacement patterns described, ullmannite (NiSbS) is replaced by anglesite (see Figs. 155 and 156).

(i) Replacements by Silver- and Gold- Containing Minerals (With Reference Also to Cobalt Mineral Replacement) Silver minerals often belong to the last phases of crystallization and may exhibit complex replacement textural patterns. Argentite (Ag2S) may replace sphalerite (ZnS) in which case migration and substitution of the metallic element Zn by Ag is possible. Another possibility is the dissolution of sphalerite and precipitation of the silver mineral (Fig. 157). Pyrargyrite (Ag3SbS3) can replace both ZnS and FeS 2 . Most probably though, dissolution of sphalerite and pyrite took place together with precipitation from solutions of pyrargyrite. The hypothesis of element substitution is not to be disregarded as an alternative (Figs. 158 and 159). Additional observations show sphalerite surrounded and replaced by marcasite (FeS2) with argentite surrounding and replacing marcasite (in some instances the argentite replaces both the marcasite and the sphalerite). Concerning the replacement of marcasite by argentite Fe is removed and Ag is introduced. In contradistinction to the hypothesis of element migration or substitution with the prerequisite of lattice and volume re-adjustment, the alternative hypothesis of dissolution involves the dissolution of marcasite and the precipitation of argentite into the voids. Argentite often replaces cobalt minerals and interesting textural patterns may evolve when cobaltite (CoAsS) is replaced by argentite (Ag2S). Here Co and As are removed and Ag is introduced (with the necessary volume and lattice re-adjustments). An alternative

possibility is the dissolution of cobaltite and precipitation of argentite (Fig. 160). In some cases the argentite forms marginal replacements to skeletal structures of cobaltite. Complex replacement textures may be exhibited where sphalerite is replaced by marcasite and where argentite replaces both the sphalerite and the marcasite. Argentite may also partly replace cobaltite crystals which were originally included in the sphalerite (Fig. 161). Supplementary textural patterns exhibiting cobaltite replaced by silver minerals are shown in Figs. 162 and 163. In particular, Fig. 162 shows cobaltite (CoAsS) surrounded, corroded and replaced by argentite (Ag2S) and similarly proustite (Ag 3 AsS 3 ) replacing cobaltite is shown in Fig. 163. In contrast to examples where argentite is the substitute mineral some very complex patterns are found where argentite is replaced by wire silver (Fig. 164). Fig. 165 shows argentite replaced marginally by native silver. In such cases removal of S (leaching out of sulfur) might account for the textural patterns shown. However, the alternative interpretation - dissolution of argentite and precipitation of silver minerals - should not be disregarded completely. Another group of minerals exhibiting complex patterns of replacement as the result of complicated replacement processes are the silver group minerals. Fig. 166 shows proustite (Ag 3 AsS 3 ) replaced by pyrargyrite (Ag3SbS3) where dissolution of proustite and precipitation of pyrargyrite took place. However, element substitution as alternative interpretation should not be excluded. In addition to the replacement of argentite by silver, cases of even more complex patterns are also common. Fig. 167 shows argentite being replaced by silver and Au tellurite. (The relationship of Au tellurite and silver is uncertain in this particular pattern.) Complex patterns of replacement are also found when argentite is partly replaced by Au tellurite and gold. Fig. 168 shows argentite replaced by Au tellurite and transected by a net of fine veinlets of native gold. Dissolution of argentite and its replacement by Au tellurite and native gold is tentatively suggested. Similarly, dissolution of tellurite and replacement by gold is also proposed. As Fig. 169 shows, native gold might replace silver minerals and complex replacement patterns are exhibited between silver minerals and gold. Fig. 169 shows argentite replaced by native silver and native gold. In the case of the gold replacing argentite most probably dissolution of argentite preceded the precipitation of gold from solutions. In Fig. 170, argentite is marginally replaced by native gold. As a corollary to the replacement of argentite by Au tellurites shown in Fig. 168, Fig. 171 shows argentite replaced by petzite (Ag 3 AuTe 2 ). In addition, the textural pattern shows marginal replacement of the argentite by petzite extending as veinlets into the argentite, 31

rendering the interpretation of dissolution of the argentite and subsequent precipitation of petzite the most plausible. In contradistinction to the cases of argentite replaced by tellurites, petzite is replaced by sylvanite (another gold tellurite (AuAgTe 4 )). Such replacement patterns might be better explained by element migration and reorganization of the lattice, possibly also with some removal of Ag (Fig. 172). In addition, petrographic descriptions are given by Callow and Worley (1965) of tellurite minerals that include petzite, hessite, altaite, calaverite, sylvanite and coloradoite. According to the authors, these minerals completely fill small vugs and minute fractures in the veins and replace gold and most of the sulfides present. Concerning the replacement of cobalt minerals, in particular cobaltite by argentite and proustite, it is considered theoretically possible for the dissolution of cobaltite and the precipitation of argentite to exist. Similar dissolution of cobaltite is proposed for the replacement of cobaltite (CoAsS) by niccolite (NiAs). Figs. 173-175 illustrate cobaltites enclosed or partly enclosed, corroded and replaced by niccolite, supporting the dissolution of cobaltite and the precipitation of niccolite from solutions.

( j ) Replacements Patterns of Cassiterite and Stannite and Replacements Involving W Minerals Stannite and Cassiterite often exhibit mutual replacement patterns that means, cassiterite (Sn0 2 ) is replaced by later stannite (Cu 2 FeSnS 4 ). Cases though of the reverse, cassiterite replacing stannite, are also abundant. Fig. 176 shows cassiterite enclosed, corroded and replaced by stannite. Often the cassiterite is preferentially replaced zonally by the stannite. The replacement patterns in the case exhibited in Fig. 176 are interpreted as the result of dissolution of cassiterite and precipitation of stannite from solutions. Similarly, Fig. 177 shows cassiterite invaded and replaced by stannite. Cases where cassiterite is replaced by stannite and haematite are also common (Fig. 178). An interesting example of cassiterite replaced by tetragonal (prismatic) stannite 3 is shown in Fig. 179. Furthermore, Noväk et al. (1962) reports cassiterite replaced by pyrrhotite via stannite. In contrast to the cases of cassiterite replaced by stannite, cases are illustrated (Fig. 180) where the stannite is enclosed or partly enclosed by cassiterite with veinlets of cassiterite extending into the stannite. Similarly, cassiterite replaces stannite with cassiterite extensions into the stannite assuming lamellar form (Fig. 181). 3

Slatkine (1966) supports that in a second phase the stanniferous solutions penetrated the earlier pegmatite and replaced the feldspar (sericitization also occurred). 32

In addition to the mutual replacement patterns of cassiterite-stannite, cassiterite replaces wolframite as Fig. 182 shows. If the general formula of wolframite [(Fe, Mn)W0 4 ] is compared with that of cassiterite, (Sn0 2 ) it can be seen that replacement of wolframite by cassiterite must involve dissolution of the wolframite and subsequent precipitation of the cassiterite from solutions. Concerning the Sn-containing minerals and their replacement patterns, interesting intergrowths are observed of teallite (PbSnS2) replacing pyrite (FeS2), which in turn partly replaces the stannite. In the case of pyrite replacing stannite, it is most probable that dissolution of the stannite preceded the crystallization of the pyrite and also pyrite was dissolved prior to crystallization of the teallite (Fig. 183). A very interesting intergrowth of blastic pyrite replacing stannite and one of teallite replacing stannite is shown in Fig. 184. As a corollary to Fig. 184, Fig. 185 shows extensive replacement of stannite by teallite. Considering the composition of stannite (Cu 2 FeSnS 4 ) and teallite, removal of Cu and Fe and introduction of Pb has taken place in the case of stannite replacement by teallite. However, dissolution of stannite has most probably taken place prior to precipitation from solutions of teallite. As already mentioned, stannite is often replaced by pyrite and interesting replacement intergrowths are shown in Fig. 186. Pyrite (crystalloblastic) also replaces cassiterite (Sn0 2 ) as a result of cassiterite dissolution prior to the precipitation of pyrite from solutions (Fig. 187). Other possibilities of Sn mineral replacement shows cassiterite being replaced by sphalerite (ZnS) where dissolution of the cassiterite has taken place prior to the precipitation of sphalerite from solutions (Fig. 188). In particular Fig. 189 shows that the sphalerite enclosed, corroded and replaced the cassiterite. Additionally to sphalerite, pyrite also replaces the cassiterite. Comparing the formulae of these two minerals suggests that dissolution of the oxide and precipitation of the pyrite has taken place by solutions (Fig. 190). As far as the cassiterite replacement is concerned, additionally to the sulfides, stannite, sphalerite and pyrite, cases are exhibited where cassiterite is replaced by antimonite (Sb2S3). Here, the dissolution of the cassiterite has probably preceded the precipitation of antimonite from solutions. Concerning the replacement of scheelite by wolframite and vice versa, the following cases are quoted: Neiva et al. (1957) described pseudomorphs of ferberite after scheelite in the W veins of Covas (Serra d'Arga) and Lagoaca (Moncorvo). Brodin and Dymkova (1966) report replacement of scheelite by hübnerite which was shown to give a 12-13% shrinkage of the anionic framework; such shrinkage fractures indicate the diffusive nature of the replacement. Ac-

cording to them, similar shrinkage on replacement was observed also in sulfides, oxides and particularly in wolframites. In contrast to the replacement of scheelite by wolframite, Orlov (1962) described the replacement of wolframite by scheelite which is accompanied by deposition of pyrite. Orlov supposed that a fairly high P. H. S. was necessary if scheelite is formed in place of wolframite where hydrothermal solutions contain Ca and Fe ions. In contradistinction to the replacement of scheelite by wolframite and vice versa, Zuev (1959) reports that schellite and manganiferous wolframite containing 20.9% MnW0 4 are found to be mutually replacive, giving rise to alternate zones of these minerals.

(k) Replacements Minerals

Patterns Involving Bi or Bi

Taking into consideration that native Bi and Bi minerals usually characterize low temperature hydrothermal paragenesis it is interesting to present some characteristic patterns of native bismuth as a substitute and furthermore to consider mutual replacement of Bi minerals. Fig. 191 shows native Bi replacing haematite (Fe 2 0 3 ). The replacement pattern resembles that of an atoll-type. Similarly, Fig. 192 shows native bismuth replacing lamellar aggregates of haematite, the substitute often enclosing relics of the haematite. In both patterns dissolution of haematite occurred prior to the precipitation of the native Bi from solutions. In contradistinction to the patterns described where native Bi replaces haematite, a more complex pattern is illustrated in Fig. 193 where tetrahedrite (Cu3SbS4) is replaced by native Bi. Similarly, tetrahedrite is replaced by wittichenite (Cu 3 BiS 3 ). Considering the replacement of tetrahedrite by native Bi, dissolution of the former has taken place prior to the precipitation of Bi from solutions. Analogously, the replacement of tetrahedrite could perhaps be explained by the dissolution of the tetrahedrite and subsequent precipitation of wittichenite from solutions. However in this case, migration of the elements could provide an alternative interpretation in the sense that Sb may have been removed and Bi introduced with the appropriate change in volume and reorganization of the lattice. In the same pattern shown in Fig. 193, it is possible to interpret the replacement of wittichenite by native Bi as a result of dissolution of wittichenite and precipitation of Bi. The alternative interpretation, namely leaching out of Cu and S from wittichenite and enrichment by added Bi, is also theoretically possible. However, such complex processes that can happen in nature are difficult to interpret on the basis of extrapolation with experimental results. Lawrence (1963) reported that aikinite and native Bi have undergone replacement by wittichenite.

Apart from the patterns presented in Fig. 193, more complex cases and processes are presented in the intergrowths in Fig. 194 where tetrahedrite is replaced by Bi and wittichenite, and additionally emplectite (CuBiS2) extensively replaces the wittichenite. This could be accounted for by leaching out of some S and Cu and by volume adjustment and lattice reorganization. In contrast to the intergrowths where Bi mainly is the substitute mineral, cases are presented of complex intergrowths where native Bi is marginally replaced and invaded by chloanthite ((NiCo)As 3 ). Such patterns hint to a possible dissolution of Bi and to precipitation of chloanthite from solutions (Figs. 195 and 196). In contradistinction to the replacement intergrowths where chloanthite was marginal to native Bi and where extensions of marginal chloanthite in the Bi could be followed, cases of ring structures of chloanthite replacing Bi are shown in Fig. 197. The relationship of chloanthite and native Bi is indeed very complex and even dubious since there are also cases where Bi is mobilized, bursting through marginal chloanthite and extending into the adjacent gangue. In addition to the case of Bi remobilization or mobilization as shown in the intergrowth in Fig. 198, cases of atoll-type replacement of niccolite (NiAs) by native Bi are presented in Fig. 199. Such cases could be explained as a result of dissolution of niccolite and precipitation of native Bi from solutions. As a consequence to the atoll-type replacement, cases are described of Bi replacing niccolite in a banded manner, i. e., as a result of the replacement, Bi bands are present in the niccolite (see Fig. 200). Intergrowths with very complex replacement patterns can be seen in Figs. 201 and 202 where niccolite with a margin of chloanthite has marginal native Bi with extensions transecting the chloanthite margin and extending into the niccolite itself, thus replacing both the chloanthite and the niccolite. Considering the composition of chloanthite ((NiCo)As 3 ) and niccolite (NiAs) and that of the substitute Bi it is more likely that dissolution of chloanthite and niccolite took place prior to their substitution by Bi. Yet more complex patterns of cobalt minerals replaced by Bi are shown in Figs. 203 and 204, where cobaltite with a banded margin of safflorite is replaced by Bi. Dissolution of both cobaltite and safflorite took place together with precipitation of Bi from subsequent solutions. As already mentioned, native Bi can replace the Bicontaining minerals wittichenite and emplectite. Similarly, bismuthinite (Bi2S3) is replaced by native Bi. Perhaps all these cases could be interpreted as the result of S leaching and enrichment of the solutions by Bi. As described, Bi and Bi-containing minerals are often the substitute of a wide range of minerals. However, as already mentioned, Bi may in tum be replaced by chloanthite and sphalerite. As a corollary to the re33

placement of Bi, Fig. 205 shows Bi replaced by chloanthite and sphalerite. It is difficult to understand sphalerite replacing Bi, although in this case, the replacement probably involves low temperature sphalerite formation. Theoretically, such patterns could be understood in terms of palingenetic sphalerite, in the sense that sphalerite was remobilized. In addition to the replacement of native Bi by lamellar sphalerite (Fig. 205), Figs. 206 and 207 show native Bi corroded and replaced by adjacent sphalerite with lamellar extensions into the Bi. Occasionally the intergrowths of Bi and sphalerite are so complex that they are difficult to explain. Fig. 208 shows that mutual replacement of Bi by sphalerite and vice versa has taken place, with chloanthite marginally replacing the Bi. Such cases of mutual replacement of Bi and sphalerite are perhaps best explained as due to subsequent remobilizations. Some more interesting replacement patterns involving Bi minerals are shown in the cases of pitchblende replaced by native Bi. The very complex patterns involved are presented in the section Replacement Involving Pitchblende. Additionally some relatively rare patterns are brought forward which show bismuthinite (Bi2S3) replaced by millerite (NiS) which transects and replaces the Bi sulfide (Fig. 209). Fig. 210 shows marginal millerite replacing the bismuthinite and sending extensions into it. Such textural patterns could be understood as due to dissolution of bismuthinite and precipitation of millerite from solutions, or by substitution of Bi by Ni and lattice re-adjustment. However as Fig. 209 shows, an indentation margin is observed between millerite and bismuthinite which more likely supports the dissolution and replacement interpretation of bismuthinite by millerite. Fig. 211 shows millerite replacing hauchecomite (Ni, Co, Bi)4(S,Sb)2, illustrating another intricate pattern of a Bi-containing mineral replaced by a nickelcontaining mineral.

(I) Some Representative Textural Patterns of Se Mineral Replacements Considering that selenium minerals are usually late phase crystallizations, complicated mutual replacement patterns are often exhibited. Fig. 212 shows trogtalite (CoSe) replaced by clausthalite (PbSe). Such textural patterns could theoretically be interpreted as the result of element removal and introduction, in the sense that Co is removed and Pb introduced. The alternative interpretation of dissolution of trogtalite and precipitation of clausthalite from solutions is another possibility. In contradistinction to the pattern shown in Fig. 212, Fig. 213 shows an even more complex pattern in which trogtalite is replaced, both by hastite (CoSe2) and clausthalite. Taking into consideration the formulae of 34

trogtalite and hastite, it is possible that additional Se and reorganization of the lattice could result in the replacement of trogtalite by hastite. Replacement of trogtalite and hastite by clausthalite would involve the removal of Co and volume and lattice re-adjustments. However, the alternative interpretation, namely dissolution of trogtalite and hastite and precipitation of clausthalite from solutions, is also plausible and should not be disregarded. Replacement of hastite/clausthalite is often dubious though because the hastite margins of replaced clausthalite have also hastite extensions in the clausthalite (Fig. 214). Mutual replacement patterns of hastite/clausthalite are illustrated in Fig. 215. Comparing the formulae of hastite (CoSe2) and clausthalite (PbSe) either replacement or mutual replacement of these minerals involve the mobilization of Co and Pb.

(m) Replacement

Involving

Pitchblende

Despite being a late phase formation mineral in the hydrothermal series (epithermal-telethermal), pitchblende often exhibits intricate intergrowth patterns where pitchblende replaces a number of minerals. However, it is far more interesting that pitchblende itself is replaced by a series of minerals. As Fig. 216 shows pitchblende (U0 2 ) interzonally invades and replaces smaltite ((NiCo)As3). Such textural patterns involve the dissolution of smaltite and the precipitation from solutions (colloidal) of pitchblende. Similarly, fine-zoned smaltite shows interzonal replacement by pitchblende (Fig. 217). In contrast to the patterns exhibiting pitchblende as a replacement, pitchblende is often substituted by other minerals, Fig. 218 illustrates native silver marginally replaced by pitchblende. Most probably, the dissolution of the silver occurred prior to the precipitation of the pitchblende from colloidal solutions. in addition to the common case of replacement of pitchblende by radiogenic galena (lead), bismuthinite (Bi2S3) also replaces pitchblende along synaeresis cracks. Considering the replacement of pitchblende by bismuthinite, it is most likely that dissolution of the uraninite has taken place along with precipitation of bismuthinite into its synaeresis cracks extending form them into the pitchblende. Initial, deformed pitchblende sphaeroids, are centrally invaded and replaced by arsenopyrite (FeAsS). Also in this case, dissolution of the uraninite occurred prior to precipitation from solutions of the arsenopyrite. Deformed initial sphaeroidal pitchblende is also invaded along cracks by tetrahedrite (Cu3SbS4) and either precipitation of the tetrahedrite occurred along the deformation cracks of the pitchblende or dissolution of the uraninite occurred prior to the precipitation of tetrahedrite (Fig. 219). Comparable replacement patterns of pitchblende replaced along a system of fractures or cracks by mar-

casite are shown in Fig. 220. Here, too, dissolution of the pitchblende took place prior to precipitation of marcasite. In addition to the cases described where deformed pitchblende is replaced along the deformation cracks, most important textural patterns exhibiting pitchblende replaced along deformation cracks are shown in a series of figures. Fig. 221 shows pitchblende replaced by native Bi and, in particular, Fig. 222 shows a relic of pitchblende left in the Bi which has centrally replaced deformed pitchblende. However, impressive patterns are also found when deformed pitchblende is replaced along deformation cracks, in which though, an intact pitchblende sphaeroid is preserved (Fig. 223). Cases are reported by Taylor et al. (1966) of unaltered dump specimens containing coffinite replacing pitchblende, and according to Ortlepp (1962), uranothorite originated through replacement in situ of relatively thorium-rich grains by siliceous solutions.

(n) Cell Wood Structure Replaced Uraninite and Other Minerals

by

Replacement of organisms and plants by later minerals is a common process and extensive literature is available. In the present effort just a few examples of cell wood structure replaced by uraninite and pyrite will be presented as examples of how the processes take place and what textures could result. Fossilization of an organism itself is a complex replacement process which involves the substitution of organic matter by solutions resulting either in amorphous mineral matter being formed such as opal, or resulting in crystalline mineral formation by fluids or melts as in the case of the fossil woods on Lesbos Island in Greece. In contrast, fossils themselves might be replaced by ore minerals. A common case is the replacement of fossils by marcasite and pyrite. As Fig. 224 shows, the cell wood structure is replaced by uraninite and marcasite or pyrite. In other instances the cell wood structure is replaced by uraninite (pitchblende) or chalcopyrite and pyrite. Tschanz (1958) described chalcocite replacing wood and forming nodules that contain small variable amounts of pyrite, bornite, covellite and uraninite. De Wiesse (1957) described the replacement of cell walls by pyrite and lumina by chalcocite which is followed by more complete replacement during which bornite, chalcopyrite, tetrahedrite-tennantite, sphalerite, galena and covellite develop. In many instances of cell wood structure replacement relic intermediate organic matter coal substance might be preserved. The consequence of cell wood structure replacement is fossil replacement by ore minerals. Kraume (1960) reports that in the Pb-Zn-bearing Rammelsberg deposit near Goslar, Upper Harz, Germany, which is a sedimentary, submarine-hydrothermal formation in the lower Middle Devonian, grains of

blende occur in sandstones in addition to the main ore bodies, and fossils are replaced by chalcopyrite.

(o) Iron Oxides and Hydro-Oxides tutes

as Substi-

Brown iron can substitute many minerals indicating most complex patterns. The best known case is the formation of limonite pseudomorphs after pyrite. However, limonite can replace a number of minerals that do not necessarily contain iron. Fig. 225 shows native Bi replaced marginally by brown iron. Similarly, native Bi is replaced by brown iron resulting in the formation of bands of brown iron in the Bi. In such cases dissolution of Bi and precipitation from solutions containing limonite (brown iron) has taken place (Fig. 225). In contrast to the replacements where the replaced mineral did not contain Fe, very impressive replacement structures result when pyrrhotite (FeS) is replaced by haematite and lepidocrocite resulting in a box wood structure, characteristic of the replacement of pyrrhotite. However, many intermediate phases intervene between the replacement of pyrrhotite and the formation of the box wood structure. A somewhat more advanced phase of replacement of pyrrhotite mainly by lepidocrocite is illustrated in Fig. 226. The alteration of iron-containing minerals by haematite martitization is a complex process and impressive textural patterns may result (see Chapter 15). In addition to martitization due to alteration of magnetite at low temperatures, cases of martitization and replacement of magnetite at elevated temperatures are also impressive. These will be discussed in Chapter 7. However, complex minerals can also be replaced by haematite. Fig. 227 shows martitized magnetite included as a relic in haematite. The same photomicrograph shows davidite relics included in the haematite. Apparently, complete replacement of the davidite by haematite took place. As already discussed, titanomagnetite might be selectively replaced by pyrite and as a result lamellae of ilmenite might be left as relics in the pyrite (Fig. 228). Considering the composition of magnetite (Fe 3 0 4 ) and of pyrite (FeS2), it is most likely that the dissolution of the magnetite occurred prior to the precipitation from solutions of the pyrite. As a corollary to these processes, Fig. 229 shows titanomagnetite after the dissolution of the magnetite with the skeleton of ilmenite lamellae left as undissolved relics. In contradistinction to cases of magnetite replacement, magnetite is often the substitute of a number of minerals. As mentioned in Chapter 5, magnetite replaces pyrrhotite and ilmenite, although the most common case is the replacement of haematite by magnetite. Such cases are often due to induced reduction either by relatively elevated temperatures or by the influence of organic matter. Fig. 230 shows haematite lamellae replaced due to induced 35

reduction to magnetite. The saw tooth texture of laths of magnetite is similarly formed by induced reduction from haematite. In addition to the induced reduction of haematite to magnetite, comparable patterns of lievrite to magnetite also exist (Fig. 231). Often the replacement patterns are an indication of a very complex history of processes and transformations. As the textural pattern in Fig. 232 shows, initial colloform iron hydro-oxides were changed to colloform structures consisting of elongated haematites, the aggregates though maintained the initial colloform pattern of the iron hydro-oxides before they changed to magnetite. As a result of the replacement, initial carbonate (relic colloform structure preserved) has been changed to iron oxides and finally to magnetite (Fig. 233). In addition to the typical patterns due to induced reduction of haematite and hydro-oxides of iron, cases of zoned carbonates replaced by magnetite are also possible (Fig. 234). The cases of replacement considered so far involved minerals that could be derived either by leaching out of S and subsequent oxidation (as in the case of pyrrhotite replaced by magnetite) or by inductive reduction (as the transformations of haematite to magnetite). In contrast, galena can be replaced by blastic magnetite as illustrated in Fig. 235.

(p) Successive

Replacements

Paragenetic associations of a mineral deposit often represent the result of successive replacements, in the sense that minerals crystallized first are replaced by further subsequent crystallizations. It is not surprising that the present state of an ore accumulation might represent a momentum in the unfolding of the spiral changes to which a geological body is destined to be submitted. This is true when it is said that an ore body is formed at a certain geological period (by this it is not referred to its pre-history), meaning that the elements and particularly the metallic elements that comprised it have existed in the earth perhaps in combination with other elements. The concept of the recycling of elements is perhaps most important to trace not only the derivation of an ore accumulation but also to understand the processes involved in its formation. In addition to the recycling of elements in the broad sense, as outlined, mobilization and remobilization of elements often take place even within the same ore body and within the time span of its formation. Taking also into consideration that not all the elements comprising an ore deposit are mobilized at one time and simultaneously, successive mobilization processes in the sense of successive replacements are most important. The supply of material forming an ore body might be due to successive supplies of solutions (which might themselves be differentiated and out of each supply a series of crystallizations or overlapping crystallization might take place). Thus, the determination of the suc36

cessive mineral replacements might bestow on us the crystallization history of an ore accumulation and help us to decipher its derivation and history. Replacement patterns presented in the previous section might actually represent only instances or cases of replacements, many of them possibly coexisting in an ore accumulation. Thus, in an ore body many replacement patterns may coexist. The determination though, of a sequence of replacement, or the order of successive replacement is, as outlined, most important to the understanding of the crystallization history of a deposit. An interesting example demonstrating the significance of the replacement sequence is the Chang Poy ore deposit of South Eastern China. Cassiterite is replaced by sphalerite and sphalerite is replaced by pyrite, often crystalloblastic, which in turn replaces the cassiterite and the niccolite. Niccolite replaces sphalerite. Pyrite replaces pyrrhotite, and marcasite also results as the alteration product of pyrrhotite. Most interesting is the replacement by antimonite of cassiterite, sphalerite and pyrite. Thus, successive replacement and even concurrent events of replacement may occur in an ore accumulation. Comparable replacement processes take place in almost all mineral deposits and paragenetic associations and these processes are probably most significant in determining the patterns observed in ore deposits. These textural patterns challenge the concept of consecutive crystallization of ore minerals from a single supply of solutions. As the textural patterns of ore minerals demonstrate, the replacement processes are not aberrant but comprise the dominant genetic patterns of ore deposits (as demonstrated by considering the following sections of this Atlas). In accordance with the general principles outlined above concerning successive replacement, some examples will be presented here, demonstrating that in addition to the fundamental concepts concerning successive replacement as a mechanism to determine the evolutionary history of an ore accumulation, textural patterns also support that many intergrowths are actually due to successive replacement. Fig. 236 shows magnetite successively replaced by sphalerite and chalcopyrite. Considering the formulae of the minerals involved, magnetite (Fe 3 0 4 ) is replaced by sphalerite (ZnS). This replacement involves dissolution of magnetite and precipitation from subsequent solutions of sphalerite. In tum chalcopyrite (CuFe2S) replaces both magnetite and sphalerite; also in this case, dissolution of both magnetite and sphalerite and precipitation of chalcopyrite took place. The alternative interpretation on element migration, however, should not be excluded entirely. Additional cases of successive replacement patterns are shown in Fig. 237 where pyrite (FeS2) is successively replaced by galena (PbS) and chalcopyrite (CuFeS2). Also in this case, the intergrowth is due to dissolution or perhaps successive dis-

solution of pyrite and of galena as well and the precipitation from solution in the first instance of galena in the voids (dissolved) of pyrite and chalcopyrite in the voids of pyrite and galena, respectively. Fig. 238 shows pyrite (FeS2) successively replaced by galena and sphalerite. Here again dissolution of pyrite and precipitation from solutions of galena was followed by dissolution of pyrite and perhaps of galena and precipitation from subsequent solutions of sphalerite. In contradistinction to the normal cases of replacement where galena usually replaces chalcopyrite and sphalerite, the patterns presented in Figs. 237 and 238 show successive replacement of galena preceding chalcopyrite and sphalerite in the successive sequence. Successive replacements are also presented in Fig. 239 where arsenopyrite (FeAsS) is replaced by galena and Au. Considering the composition of the minerals arsenopyrite and of the substitutes galena and native gold, most likely the arsenopyrite was dissolved and in the voids a subsequent precipitation of galena and later on gold took place. It is possible that the gold has replaced the galena as well. Cases are reported where cobaltite (CoAsS) is replaced by chalcopyrite and galena. Most probably the cobaltite was dissolved prior to the precipitation from subsequent solutions of chalcopyrite and galena. The fact that the replacement assumes veinlet forms, favours the dissolution of the mineral to be replaced and the precipitation of the substitutes from subsequent solutions. In addition to ore minerals, gangue minerals might be subjected to multiple and successive replacements. Fig. 240 shows quartz with well-formed crystal faces replaced by bornite (Cu5FeS4) and tetrahedrite (Cu3SbS2). Also small scale replacement of quartz by pyrite has taken place. Impressive multiple or successive replacement patterns are observed where sphalerite is replaced by biastic pyrite and also transected by later veinlets of gangue with fine margins of galena. In addition to the veinlets, more massive replacement of the sphalerite by galena took place. Fig. 241 shows chalcocite (Cu2S) transected by a veinlet of malachite (Cu 2 [(0H) 2 /C0 3 ], in which native Cu has also crystallized. Here also dissolution of chalcocite proceeded the precipitation of malachite and Cu. A common pattern of copper minerals showing successive replacement is found where pyrite is replaced by bornite and bornite by chalcocite which also replaces adjacent chalcopyrite (Fig. 242). Variations of these patterns are shown in Fig. 243 where pyrite is replaced by chalcopyrite and the chalcopyrite is replaced by chalcocite. Comparable to the patterns illustrated in Fig. 242, Fig. 244 shows bornite additionally to pyrite replaced by chalcopyrite. Both the chalcopyrite and the bornite are transected by veinlets of chalcocite. For the above mentioned cases it is most likely that migration of elements has played a major role in the replacement rather than dissolution and precipitation

from solutions. However, these processes should not be excluded. The interpretation of complex intergrowths particularly in ore microscopy is often dubious and difficult. However, by examining many variants of the pattern, particularly in the polished section, tentative interpretations can be put forward. The pattern shown in Fig. 245 is interpreted to show replacement of tetrahedrite by pyrite while both the pyrite and the tetrahedrite are transected and replaced by chalcocite. In such cases it is possible that both dissolution and subsequent precipitation from solutions as well as element migration have had a part in the formation of the intergrowth. Complex multiple or successive replacement patterns are not rare and many variants exist. Fig. 246 shows enargite (Cu3AsS4) replaced by fine and coarse grained chalcopyrite. The enargite is replaced by covellite and later haematite. Another complex pattern involving mainly copper minerals is shown in Fig. 247: bornite is replaced by chalcopyrite and both are transected by veinlets of chalcocite. Pyrite is also present showing atoll-type replacement by chalcopyrite. Element migration, dissolution and precipitation are alternative interpretations which could be applicable to explain these complex successive replacement patterns. A rather rare case is illustrated in Fig. 248 that shows magnetite (Fe 3 0 4 ) intergrown with a silicate which is replaced by chalcopyrite which in turn is enclosed, corroded and replaced by a pyrite megablast. Dissolution and precipitation from solutions is tentatively suggested as the mechanism which produced this pattern. Other multiple or successive replacement patterns presented here involve Bi or Bi minerals. Fig. 249 shows native Bi replaced and invaded by sphalerite and chloanthite [(NiCo)As3], Dissolution of Bi and subsequent precipitation from solutions of sphalerite and chloanthite minerals most probably took place. In contradistinction to Fig. 249 where Bi was replaced, Fig. 250 shows sphalerite replaced by bismuthinite (Bi2S3) and in turn replaced by native Bi. The replacement of sphalerite by bismuthinite should be interpreted as a case of dissolution of sphalerite and precipitation of Bi from subsequent solutions. It should be noted that in the patterns of Figs. 237 and 238 galena was first in the successive replacements in contrast to the more common case where galena often replaces both chalcopyrite and sphalerite. Arsenopyrite (FeAsS) is successively replaced by galena and native gold. Here again dissolution of arsenopyrite and successive precipitation from solutions of galena and gold have taken place. An instance of multiple replacement is reported in the case where cobaltite is replaced by chalcopyrite and galena. In this case, dissolution of the cobaltite has most probably taken place and precipitation from solutions of chalcopyrite and galena occurred. The veinform replacement of cobaltite by chalcopyrite supports 37

the interpretation of dissolution of cobaltite along cracks and precipitation from subsequent solutions of chalcopyrite. Somehow comparable patterns are also found in the case of cobaltite replacement by galena. Cases of successive replacement involving silicates are occasionally observed. Fig. 240 shows quartz exhibiting well-formed crystal faces successively replaced by bornite and tetrahedrite. Dissolution of the quartz preceded its replacement by bornite and tetrahedrite from solutions. Small scale replacement of quartz by pyrite took also place. A different type of successive replacement is illustrated in the patterns in Fig. 251 where sphalerite is replaced by crystalloblastic pyrite and by veinlets of gangue with fine margins of galena. The blastic replacement of sphalerite by pyrite most probably necessitates dissolution of the sphalerite and crystallization of the pyrite from solutions. Similarly, gangue-galena veinlets are due to the replacement of sphalerite along cracks, by dissolution of the sphalerite and precipitation from solutions of the gangue and the galena. As already discussed, copper minerals exhibit complex and mutual replacement processes. Cases of successive replacement are also abundant. Fig. 241 shows chalcocite transected by a replacement protuberance (veinlet) of malachite (Cu 2 [(0H) 2 /C0 3 ]) which is centrally occupied by native copper. Dissolution and replacement of the chalcocite took place prior to the precipitation from solutions of the malachite. Replacement of malachite by native Cu metal might be due to lateral secretion of copper from the malachite. Cu-containing minerals often show complex successive replacement and the successive replacement sequence is often indicated by the trend of enrichment in copper. Fig. 244 shows pyrite replaced by bornite which in tum is replaced by chalcopyrite (relics of bornite are maintained in the chalcopyrite). Furthermore, both the bornite and the chalcopyrite are then replaced by chalcocite. During this sequence the Cu content was progressively increased, in the sense that bornite was replaced by chalcopyrite which is relatively richer in Cu and chalcopyrite was replaced by chalcocite which has a higher copper content than chalcopyrite. Progressive enrichment in Cu of the successive replacement of substitute copper minerals is endorsed by the patterns shown in Figs. 244 and 247 where pyrite is replaced by chalcopyrite and where chalcopyrite replaced pre-existing bornite. Both the bornite and the chalcopyrite are replaced by chalcocite. These patterns are comparable to the patterns shown in Fig. 242. Fig. 247, however, shows atoll-type replacement of pyrite by chalcopyrite. Considering the chemical composition of the minerals (Figs. 242, 244 and 247), it is more likely that pyrite was dissolved prior to the precipitation from solutions of chalcopyrite. In contradistinction, the replacement of bornite by chalcopyrite could be interpreted as either the result of dissolution of the bornite and precipitation of chalcopyrite by solutions 38

or as due to the migration of elements, meaning that there was a relative increase in copper and, of course, the necessary volume re-adjustment and lattice reorganization took place. The formation of chalcocite veinlets transecting both the bornite and the chalcopyrite could be interpreted as the result of leaching out of Fe and as mobilization of Cu enriched solutions. Comparable patterns to bornite replaced by chalcopyrite or bornite and chalcopyrite replaced by chalcocite are those exhibiting enargite (Cu3AsS4) replaced by chalcopyrite or covellite (CuS). Haematite also replaces both enargite and covellite. Considering the chemical composition of these minerals in the case of enargite replaced by chalcopyrite, As was removed and Fe introduced. In the case of the replacement of enargite by covellite As was removed. Such patterns could be interpreted as the result of migration of elements or as due to dissolution of the mineral that was replaced and precipitation of the substitute from solutions into possible voids, formed due to dissolution. The replacement of enargite and covellite by haematite is most likely the result of the dissolution of the enargite and covellite and the precipitation of haematite from solutions. In contrast to these patterns though, where the pyrite was replaced mainly by chalcopyrite, Fig. 245 shows tetrahedrite replaced by blastic pyrite and both the pyrite and the tetrahedrite are replaced by chalcocite. The replacement of tetrahedrite by pyrite could be interpreted as the consequence of the dissolution of the tetrahedrite and the formation of the blastic pyrite from solutions. The formation of chalcocite veinlets transecting the pyrite is probably due to dissolution of pyrite and precipitation of chalcocite from solutions. The replacement process of tetrahedrite by chalcocite is perhaps comparable. Further considering successive replacements, silicate intergrown with magnetite is replaced by chalcopyrite and the entire structure is enclosed, rounded and corroded by later blastic pyrite. Some interesting examples of multiple replacement patterns are illustrated in Figs. 249 and 250. Fig. 249 shows native Bi replaced by sphalerite and chloanthite ((NiCo)Asj). The replacement of Bi by sphalerite and chloanthite is rather to be understood as the consequence of dissolution of the native metal and the subsequent precipitation from solutions of sphalerite and chloanthite. In contrast to the replacement of Bi by sphalerite, Fig. 250 shows sphalerite replaced by bismuthinite (Bi2S3) which in turn is replaced by native Bi. The replacement of sphalerite by bismuthinite could be interpreted as the result of dissolution of the sphalerite and the precipitation of bismuthinite from solutions, similarly sphalerite is replaced by Bi. Concerning the relationship of bismuthinite and Bi, it is perhaps possible for the bismuthinite to replace the native metal in this case.

(q) Successive Replacement or Successive Crystallization Phases Successive replacements are the corollary to replacement processes. They incorporate elements of the wave replacement process of Edwards (1960) while at the same time they could also be the result of differentiation and successive precipitation of a single solution supply. The time difference of each crystallization phase and the difference in crystallization times of consecutive crystallization phases are very difficult to determine exactly and precise chronological data are difficult to obtain. It should be considered that additional to the normal sequence of crystallization belonging to a paragenetic association, cases exist where superimposed or reactive paragenetic formations are found which may result in minerals that are associated but of different age. Augustithis (1982) in his studies on the mineral parageneses of the Abu Dabbab area (Red Sea region, Egypt) showed that a lithophilic (crust) mineral paragenesis is associated with granitization formation of the apogranite, Precambrian in age, and a superimposed Tertiary paragenetic association (in cases reacting with the Precambrian paragenesis). Here and in cases of paragenetic associations formed due to polymorphism, it will be difficult to use the term successive replacement or crystallization since the association of the minerals may be due to an independent geological process. The Precambrian association of the granitophile elements of the Abu Dabbab apogranite, for example, is due to mobilization of elements during the Precambrian granitization-apogranite formation whereas the Tertiary manganese paragenesis is the result of lateral segregation and leaching of buried basalt in the rift. Based on this theoretical consideration and bearing in mind that often the patterns exhibited in the case of successive crystallization are due to repeated or multiple replacement processes, an attempt will be made to illustrate some of these patterns. Fig. 252 shows pitchblende including a first generation of galena which exhibits Maltese cross structure (see Fig. 253 in particular). The Maltese cross structure of galena is the result of pre-pitchblende galena formation which is corroded and replaced by the later uraninite. As Fig. 252 shows, a later generation of galena marginal to the pitchblende is formed replacing the uraninite. Thus, the pattern shows successive crystallization and replacement, represented by the Maltese cross structure galena included in the later formed pitchblende which is followed by the marginal galena generation. In addition, successive crystallization is exhibited in Figs. 254 and 255 where smaltite is followed by a marginal growth of colloform pitchblende which is followed by an overgrowth of smaltite (see Fig. 254 in particular). Thus, the successive crystallization pattern is represented by smaltite followed by pitchblende

(colloform and appearing as interzonal in the smaltite) and finally by smaltite overgrown on the pitchblende. The relationship between uraninite and gold may resume successive crystallization patterns in the sense that gold may be enclosed in the uraninite (Fig. 256) while simultaneous gold overgrowth on the pitchblende or gold with silicate veinlets may transect the uraninite (Fig. 257). Another complex pattern involving pitchblende which again is interpreted as due to successive formation is shown in Fig. 258. Colloform pitchblende exhibiting sphaeroidal banded structure is replaced and corroded by a symplectite consisting of bomite and bismuthinite which in turn is corroded and engulfed by a later pitchblende shell (latest layer of the pitchblende sphaeroids). Thus, considering the patterns shown in Fig. 258, first there are the gel sphaeroids of pitchblende corroded and replaced by the bornite/bismuthinite symplectite (see Chapter 12) which then in turn are replaced and partly surrounded by a later shell (zonal growth) of pitchblende. In addition to the successive uraninite/gold patterns already mentioned, cases of pyrite exist transected by veinlets of different sulfides in which at last successive crystallization of gold is exhibited, partly replacing the sulfides of the veinlets and the pyrite host. Fig. 259 shows pyrite transected by a veinlet consisting of silicate (quartz), chalcopyrite and gold which has replaced the chalcopyrite of the veinlet and the pyrite host. The successive crystallization/replacement is thus the pyrite host transected and replaced by a veinlet of silicate and chalcopyrite to which belongs also gold which in addition to replacing the chalcopyrite of the veinlet also replaces independently the host pyrite as well. Fig. 260 and 261 show pyrite replaced successively by chalcopyrite which, as well as the host pyrite, is replaced by the later gold (often a component of the veinlets transecting the pyrite). Additional observations show a fracture system of pyrite occupied by later gold which has also replaced the pyrite host. A later veinlet consisting of sphalerite, chalcopyrite and gold and associated with the fracture system in the pyrite, transected the pyrite which is successively replaced by the sphalerite, chalcopyrite and gold of the veinlet. The gold of the veinlet is related to the gold occupying the fractures of the pyrite and forms the last phase of the successive crystallization series, specifically: pyrite, veinlet consisting of sphalerite and chalcopyrite, and gold, component of the veinlet replacing the chalcopyrite and sphalerite of the veinlet and the host pyrite (also as extension of the veinlet, gold replaces the pyrite). Besides these patterns of pyrite replaced by veinlets consisting of silicates, sphalerite, chalcopyrite and gold, successive replacement is found where pyrite is replaced by chalcopyrite which in turn is replaced by gold which also replaces the pyrite marginally (Fig. 262). 39

Furthermore, cobalt minerals are associated with colloform pitchblende (originally colloform sphaeroids) which have been deformed subsequently and fractured tectonically and which are invaded by pyrite and subsequent tetrahedrite (Fig. 263). This pattern shows how complex the crystallization sequence of an ore can be in the sense that deformation can intervene between the successive crystallization sequences. Cases are presented where successive replacement might be related to polymetamorphism in the sense that chromite (chromospinels) might be included, corroded and partly replaced by later crystalloblastic garnet which is replaced by later chalcopyrite (Figs. 264 and 265). As shown in Fig. 265 in particular, chalcopyrite replaced the chromospinel that is included in the garnet. The chalcopyrite solutions worked their way through the contact chromite/garnet and replaced both. In contrast, Fig. 266 shows garnets corroded and invaded by chalcopyrite which surrounds the garnet and sends extensions into the garnet. Another complex multiple pattern - jacobsite and braunite - is exhibited in Fig. 267. Braunite-1 is en-

40

closed and replaced by jacobsite which in turn is enclosed and surrounded by braunite-2. Fig. 268 shows again braunite-1 enclosed by jacobsite which in turn is surrounded and corroded by braunite-2. Such patterns are more dubious and could either be due to successive replacement or to overlapping crystallization or are due to the possibility that braunite crystallization was more extended in time. The pattern could perhaps represent simultaneous crystallization of braunite and jacobsite. However, in this case it is difficult to understand the pattern shown in Fig. 269 which clearly shows a veinform jacobsite transecting braunite. This veinform replacement of braunite supports that at least some of the braunite was formed before the jacobsite by which it is replaced in veinform. The patterns presented of successive or multiple replacement show the complexity and the difficulties in trying to understand the crystallization sequence in an ore. It is essential to establish the sequence of each pair of minerals in a paragenesis before a crystallization sequence can be tentatively suggested.

Chapter 6

Replacement Versus Ex-Solutions

(a) General Considerations of Ex-Solved Bodies and Replacements Simulating Ex-Solutions (e. g., Magnetite-Ilmenite, MagnetiteSpinel, Chromite-Spinel, Chromite-Rutile Inter growths) For decades the solid solutions/ex-solutions hypothesis has been predominant in ore mineralogy and the patterns of oriented inclusions in hosts have been considered as evidence of ex-solutions. In addition to the observations of oriented intergrowths, experimental mineralogy has provided the theoretical support for the exsolution hypothesis. Edwards (1960) in considering the state of solid solution and the formation of ex-solutions by unmixing described three essentially different types of solid solutions and concerning each types he states the following: 1. Substitution solid solutions in which the atoms of the solute replace atoms of the solvent, furthermore in metals substitutional solid solution is affected by the following factors: (i) The relative size of the two metals; (ii) The valencies of the two metals; (iii) Their tendency to form stable intermediate compounds (the electronegative valency effects); (iv) Their crystal structure. Concerning the electronegative valency effects it is stated additionally: The tendency to form compounds is stronger the more electropositive the one element and the more electronegative the other. Similar valency between solute and solvent atoms favours wide solid solution when they are of different valency. In conclusion Edwards states: "The most favourable conditions for the formation of a wide range of solid solutions are therefore that the solvent and the solute atoms are nearly the same size, do not form intermediate compounds and have the same valency." When in addition they have the same crystal structure a continuous solid solution series may occur. 2. Interstitial solid solutions. The solute atoms are dispersed in solvent atoms, so that they are introduced into the lattice. In addition to the solvent atoms, the solute atoms shall be of distinct smaller atomic radius. Interstitial solid solutions may be expected in some non-stoichiometric minerals.

3. Proxy solid solutions. In certain non-stoichiometric minerals a third type of solid solutions occurs which is designated as proxy solid solution. In these minerals a small number of metal atoms is pictured as occupying positions normally occupied by sulfur or arsenic atoms, so that the minerals concerned have a metal content slightly in excess to their ideal composition. In addition to the basic theoretical considerations of solid solutions/ex-solutions by Edwards, it should be recalled that Edwards emphasized that none of the textural or structural patterns by themselves can stand up to replacement explanation. This indeed indicates how close the phenomenology of replacements is to that of supposed ex-solutions and eutectic crystallization. Schwartz (1931, 1942) considered it necessary to put forward a number of criteria that would define and distinguish ex-solutions from replacements. Furthermore when discussing replacement textures simulating exsolution bodies it will be necessary to refer to these criteria. Replacement patterns simulating ex-solutions will be discussed by themselves. It is therefore useful in the present consideration of replacement "ex-solution bodies" to quote the criteria as proposed by Schwartz. Criteria for the recognition Schwartz:

of ex-solution

textures

by

1. Interpretation of the textures must be based on the visible relationships of the minerals concerned. 2. Minute inclusions distributed generally in the crystallographic directions of the host mineral (solvent) so that the orientation of the intergrowth varies with the orientation of the crystal (with the exception of emulsion ex-solutions). 3. The inclusions (lamellar, blade, plate rod-like emulsions, rounded blebs) have sharp smooth boundaries and are generally single crystals. 4. The inclusions show an even or seriate distribution through the host. 5. If the solute mineral occurs outside the host (solvent) mineral it must occur as more or less interstitial grains. 6. Where the ex-solution bodies form a connected lattice work they show no widening at the contact with one another. 41

7. The composition and crystal structures of the two minerals involved must be such that a solid solution between them is feasible at high temperatures. 8. There must be a general absence of textures indicating replacements between them. 9. Ex-solutions do not occur in veinform. Considering the textural patterns which are supposed to be due to ex-solutions the above mentioned criteria will be used as part of the evidence that most of these patterns are actually formed by replacement. Replacement processes and products can approach substitution ex-solutions. This is further supported by the definition of the latter in the sense that in substitution ex-solutions the solvent atoms can replace those of the solvent. Moreover, these replacement textures in addition to the element replacement processes, diffusion and element mobility will also be discussed. As already mentioned in consideration of the socalled ex-solutions in the sense that they represent unmixing from the solid-solutions and the possible alternative interpretations, a number of case studies can be quoted. Some of the most representative examples have been selected, e. g., the case of magnetite (Fe 3 0 4 ) - ilmenite (FeTi0 3 ) which deserves consideration since it is often quoted as a classical example of solid solution and of ex-solutions of ilmenite which occur as elongated lamellae parallel to their [0001] direction and oriented in the [111] direction of the magnetite. According to the orthodox interpretation, titanium can enter into solid solutions of magnetite at high temperatures forming a homogenous mineral which is slightly anisotropic. Substitution of Ti 4 or Ti 3 for Fe 2 in the magnetite lattice is possible and perhaps involves distortion resulting in a weakly anisotropic solid solution of magnetite-ilmenite. Edwards believes that unmixing is developed at relatively high temperatures when the rate of diffusion is high, so that the unmixing proceeds rapidly. Edwards (1960) states: "The orientation of the ilmenite lamellae in the magnetite is attributed to the sharing of oxygen planes. In magnetite every third and seventh [111] plane consists of oxygen only, while in ilmenite every third [0001] plane consists of oxygen ions only. The spacing of the oxygen ions in these planes is such that if the planes are superimposed on one another, the oxygen position in the two planes practically coincide." Thus, according to the orthodox interpretation ilmenite lamellae in magnetite are basically due to unmixing at relatively high temperatures and are the result of diffusion of Ti which has replaced Fe in the lattice of the magnetite. It is interesting on one hand to see how compatible the phenomenology of ilmenite lamellar bodies in magnetite is with the criteria set proposed by Schwartz concerning ex-solutions, and on the other hand, to consider possibilities of ilmenite lamellae formation in magnetite in nature with relatively lowered tempera42

ture conditions prevailing than those that are supposed by the unmixing hypothesis. In contrast to the ex-solution interpretation, Fig. 270 shows magnetite transected by a veinlet consisting partly of silicate and ilmenite with extensions of the veinlet-ilmenite as oriented lamellae in the magnetite. Considering the criteria set by Schwartz that ex-solutions do not occur in veinform, it is clear that in this case the veinform ilmenite and its extensions as oriented ilmenite lamellae in the magnetite are not ex-solutions but are the result of replacement. Fig. 271 and 272 show oriented lamellae of ilmenite in magnetite extending beyond their intergrowth with the magnetite and occurring together with silicate or pyrite. Considering the statement by Edwards that with slow cooling the migration of Ti from the solid solution of ilmenite-magnetite can move outside the magnetite and occupy the intergranular of the magnetite as granular grain, the present observations (Figs. 271 and 272) do not comply with this interpretation since the ilmenite lamellae that extend outside the magnetite are not granular and do not occupy the intergranular of the magnetite grains. The alternative interpretation presented here is that these ilmenites are replacements that start outside the magnetite and extend into it following the directions of crystal penetrability such as the [111] direction of the magnetite. Additional observations incompatible with the criteria of Schwartz concerning ex-solutions are oriented lamellae of ilmenite following the [111] direction of the magnetite at their intersections either widening or replacing extensively the magnetite (Figs. 273 and 274). In contrast to the usual case of a single mineral phase "unmixing" from the host, cases of multiple exsolution phases are observed and they should be considered from the point of view whether they actually are the result of multiple unmixing or the result of multiple or successive replacement. Schwartz stated for distinguishing ex-solutions: "The inclusions (lamellar, blade, plate rod-like emulsions, rounded blebs) have sharp smooth boundaries and are generally single crystals." Edwards goes even further, stating: "Commonly the initial unmixing is into partial solid solutions which segregate by solid diffusion, and then continue to unmix in their turn with further slow cooling... A further complication is introduced where more than one mineral is precipitated from a single host... The rates of unmixing of the several precipitate minerals, and of their segregation vary with temperature and concentration so that though one solute mineral may begin to precipitate before another, the unmixing of the latter may be completed first." There is no actual disagreement between the criterion of Schwartz and the consideration of Edwards in the sense that a single ex-solution body must be a single crystal and that in the same host there can be more than one ex-solved mineral. However, observations might further periplex the matter since cases are pre-

sented where the sequence of unmixing of the ex-solutions and the fact that they can exhibit intergrowths (thus, the ex-solution body is not a single crystal but an intergrowth of two or more kinds of minerals), renders successive ex-solutions from a single host in cases problematic but also in many cases uncertain. Fig. 275 shows ilmenite lamella following the [111] direction of the host magnetite while in the same magnetite host, spinel bodies are also following the general direction of [100] of the host. It should be pointed out that the spinel bodies follow the contact (margin) of the ilmenite with the magnetite and extend into the ilmenite replacing it. Ulvospinel is also present as an additional ex-solution phase in the single magnetite host. According to the successive ex-solution interpretation from this single illustrated magnetite, the first to be ex-solved is the ilmenite, followed by ulvite and finally spinel. However in this case, the position of the ulvite in the succession of the supposed successive exsolutions is not certain. In contrast to the ex-solution explanation, it is suggested that it is a case of successive replacements in the sense that ilmenite replaced the magnetite. Similarly the spinel replaced both the magnetite and the ilmenite into which it extends from its contact with the magnetite inwards. Whereas the ex-solution temperature of ulvite from magnetite is often given to be about 600° C it is most likely that tablecloth structures of ulvite in magnetite following the [100] direction of the host might also be due to replacement as another titaniferous phase (Fe 2 Ti0 4 ). The fact that these ex-solutions might disappear on heating of the magnetite specimen is more likely due to the diffusion of the elements into the lattice of magnetite replacing it and in this case, the Fe cations of the host. Considering the composition of ulvite (Fe 2 Ti0 4 ) and ilmenite (FeTi0 3 ), ulvite could theoretically be formed from ilmenite with an excess of FeO, or, also theoretically, it could break down to an intergrowth of magnetite (Fe 3 0 4 ) and ilmenite according the hypothetical reaction 3Fe 2 Ti0 4 + Ο = F e 3 0 4 + 3TiFe0 3 . In accordance to the hypothesis of solid solution/ex-solutions it is possible that in titanomagnetites ulvite could also be ex-solved. The alternative interpretation that it may be a successive replacement of magnetite by ilmenite, spinel and ulvite, is supported by the intergrowths presented in Figs. 275 and 276. In contradistinction, Fig. 277 shows the coexistence of ulvite and spinel in the same magnetite sample (as that exhibiting the pattern shown in Fig. 275). The presence of ulvite and spinel (hercynite FeAl 2 0 4 ) could similarly be interpreted as due to successive exsolutions. In contrast, spinels are often considered to be replacements of magnetite as a number of figures will support. Fig. 278 shows ulvite in magnetite and spinel following the cleavage [100] of the magnetite. Additional evidence of spinel replacing magnetite is shown in Fig. 279 where ilmenite lens-shaped bodies and spinel replace the magnetite. As a corollary of the

pattern shown in Fig. 275, Fig. 280 shows ilmenite lamellae oriented in the [111] direction of the magnetite with spinel dots marginal to the lamellar ilmenite with the magnetite. Also a spinel rounded body replaces the magnetite. Often though, the ilmenite is oriented in the [111] direction of the magnetite and spinel in the [100] direction of the same magnetite grain (Fig. 281). Such patterns could be interpreted on the basis of successive replacement. This is supported by the pattern shown in Fig. 282 which shows that the spinel lamellae in the magnetite consist of separate individuals exhibiting symplectic form and that marginal spinel is also replacing the magnetite. It should be pointed out that spinel is also present in the same magnetite following the [100] direction of the host. The symplectic form of the spinel individuals in the magnetite are further illustrated by Figs. 283 and 284. A very complex pattern of spinel marginal to ilmenite lamellae in a magnetite host is shown in Fig. 285: ilmenite lamellae in the magnetite with marginal symplectic spinel and with fine interbanded spinel lamellae with the ilmenite lamellae. Such patterns would support multiple and successive replacement of ilmenite replacing the magnetite, and spinel replacing both the magnetite and the ilmenite. In support of the replacement interpretation Fig. 286 shows titanomagnetite with ilmenite and spinel lamellae following the [100] direction of the magnetite host. This figure also shows chalcopyrite replacing magnetite along the same direction as one of the spinel lamellae. Epitactic spinel is marginal to the magnetite (which is partly replaced by the later chalcopyrite). The coexistence of oriented lamellae of spinel and their continuation filling cracks of the host magnetite is illustrated in Fig. 287. Fig. 288 shows spinel marginally replacing a magnetite grain and sending prolongations into the host magnetite attaining lamellar character. In contradistinction, Fig. 289 shows lamellar spinel following a crack oriented in the [100] direction of the host magnetite and fine interspersed spinels are also present in the magnetite with a tendency to disappear as the lamellar spinel in the magnetite is approached. In contrast to the usual case of ilmenite lamellae oriented in the [111] direction of the host magnetite and the orientation of spinel lamellae in the [100] direction of the same host magnetite, ilmenite and spinel bodies are associated with the magnetite exhibiting unusual patterns. Fig. 290 shows magnetite with a pattern of curved fractures which are occupied by spinel and ilmenite. The ilmenite is often marginal to the spinel and in cases large spinel bodies are present in spaces in between the magnetite (see Fig. 291). Such patterns can probably be explained as the result of ex-solution of both ilmenite and spinel from the magnetite in the sense that the ex-solved phases occupy the intergranular of the magnetite: in this specific case the curved cracks (fractures) of the magnetite. However, as Fig. 290 shows in particular, the spinel is 43

not an intergranular phase but an independent crystallization, that means, the ex-solved spinel could not attain such dimensions (i. e., not as much spinel could be in solid solution in the initial host of magnetite neither could as much Al be as solute in the magnetite). An additional point supporting the interpretation that spinel and ilmenite replace the magnetite is that often spinel veinlets extend from larger bodies of spinel into the adjacent magnetite attaining veinform character, consisting usually of central spinel and marginal ilmenite (Figs. 291 and 292). In cases, the ilmenite and the spinel follow very fine cracks in the host magnetite (Fig. 292). The patterns presented show that what is referred to as ex-solutions in magnetite could represent replacement structures. Also, the conditions of formation of lamellar intergrowths of ilmenite in magnetite might be different from those postulated, on the basis of unmixing upon temperature decrease of a solid solution. Another problem to be considered in addition to the genesis are the geoenvironmental conditions under which such lamellar bodies of ilmenite can be formed in magnetite. Augustithis (1964) showed that magnetite octahedral crystalloblasts (idioblasts) contain lamellae of ilmenite oriented in the [111] direction of the host; spinels are also present in the magnetite either marginal to the ilmenite lamellae or oriented in the [100] direction of the host. It needs to be emphasized that the blastic magnetite octahedra occur in leuchtenbergite in which velonoblastic anthophyllite also occurs. Considering that the formation temperature of leuchtenbergite is thought to be about 200-250° C, it becomes questionable that these lamellar ilmenites (Figs. 293 and 294) are formed under the theoretical conditions postulated for unmixing of ilmenite in solid solution in magnetite by temperature decrease at about 600° C. The presence of blastic magnetite in leuchtenbergite (chlorite) and the very fact that the idioblastic magnetite contains oriented lamellae of ilmenite with marginal spinel bodies clearly contradict the physicochemical conditions that were supposed to prevail during unmixing. The presented patterns of magnetite with oriented lamellae of ilmenite, spinel and ulvite support an alternative interpretation to the ex-solution hypothesis and as pointed out, indicate that a complex replacement process could produce the patterns presented. In contrast to the magnetite another spinel, chromite FeCr 2 0 4 or rather [(Fe, Mg)(Cr, Al, Fe) 2 0 4 ], could contain complex patterns of lamellae oriented in the host chromite. Fig. 295 shows fine rutile (Ti0 2 ) lamellae oriented in the [111] direction of the host chromite. As Fig. 296 shows, often in addition to the rutile lamellae, spinel (hercynite FeAl 2 0 4 ) lamellae occur oriented again in the [111] direction of the host chromite. Cases are observed where the nitile and the spinel lamellae are next to each other or occupy the 44

same octahedral cleavage space (Fig. 296). Additional to the lamellar rutile bodies present, rounded rutile bodies may also be observed in the same chromospinel (Fig. 297). There is usually an antipathy or better an "incompatibility" between Cr and Ti in ultrabasic rocks (see Augustithis, 1979). It is geochemically unusual to find Ti grains (rutile) and rutile lamellae in the chromospinel. However, the fact that spinel and rutile lamellae occur in the same cleavage plane of the host chromite, might indicate that the Ti presence in the chromospinel is related to Al in the chromospinel and that the migration of Ti to form lamellar rutile might be related to the mobilization of Al and Fe which formed the spinel lamellae in the chromite host (Fig. 296).

(b) Sphalerite as Host of Chalcopyrite, Stannite and Galena - also Star-Shaped Bodies in Chalcopyrite Among many others, Edwards (1960) considers sphalerite (ZnS) as the solvent in the lattice in which many solutes can be incorporated. In particular, he points out that in sphalerite up to 17% iron can substitute Zn. He also emphasizes the favourable dimensions of the atomic diameters of the two metals: Fe = 2.48 Ä and Zn = 2.66 Ä (a difference of only 6.8%). Edwards supports that iron-rich sphalerite has the tendency to unmix on cooling, giving rise to sphalerite and pyrrhotite. Furthermore, he believes that a number of sulfosalts of copper, lead and silver are isostructural (isomorphous) and form complete or extensive solid solutions at normal temperatures. The orthodox interpretation of "ex-solutions" further supports that a solid solution within a number of sulfide minerals is further promoted at high temperatures at which disorder in their crystal structures takes place. Edwards states specifically: "When in a state of disorder, i. e., at high temperatures, extensive solid solution is possible between those minerals (e. g., chalcocite, digenite, chalcopyrite, stannite, argentite) which have similar structure and in which the S atoms or their equivalents are more or less equally spaced. Thus, extensive mutual solid solution is possible between ZnS, FeS, CuFe2, Cu 2 FeSnS 4 , Cu 5 FeS 4 ." Since the possibility of solid solutions of sphalerite and chalcopyrite is considered to be a classical case by the proponents of the "unmixing hypothesis", further discussion of the possibility of chalcopyrite in solid solution with the sphalerite is quoted by Edwards: "Chalcopyrite and sphalerite have a very similar structure, in which the Zn atoms are replaced by alternative Cu and Fe atoms, but the unit cell of chalcopyrite is twice as large as that of sphalerite. The difference in size of the Cu and Fe atoms compared with the Zn atoms causes the tetrahedral arrangement of the metal atoms in CuFeS2 to be slightly irregular, so that

CuFeS 2 is tetragonal but with its angles very close to those of the cubic system." The above mentioned interpretations are furnished in support of the hypothesis that chalcopyrite intergrowths in sphalerite are due to unmixing of the solid solutions of sphalerite and chalcopyrite. In contrast to these rather plausible crystallochemical explanations it will be of great significance to observe the textures of chalcopyrite in intergrowth with the sphalerite host and to see if there is any serious discrepancy between the solid solution hypothesis and the phenomenology observed. Figs. 298 and 299 show sphalerite marginally replaced by chalcopyrite with extensions of the latter attaining intergrowth forms with the sphalerite and simulating the often described patterns of chalcopyrite exsolutions in the sphalerite. The marginal chalcopyrite is not an intergranular phase, in no way does it represent an ex-solved intergranular phase but is an independent precipitated phase from solutions after sphalerite which is corroded, replaced and infiltrated by the chalcopyrite. Furthermore, considering the criteria of Schwartz, in particular that ex-solutions do not occur in veinform, a series of observations clearly shows that chalcopyrite bodies previously described as ex-solutions, in reality have extensions clearly assuming veinform character. Fig. 300 shows chalcopyrite "ex-solution bodies" having extensions veinform in the host sphalerite. Similarly Fig. 301 shows veinlets of chalcopyrite topologically attaining "ex-solution" form and character as it is particularly shown when "exsolution" bodies themselves are following the cracks of the veinlets (see also Fig. 302). Cases though are observed where the chalcopyrite follows rehealed fracture lines in the sphalerite (Fig. 303). Such cases are difficult to understand since it can be argued that unmixing of the solid solution of chalcopyrite and sphalerite occurred after the fracturing and rehealing of the sphalerite. (However, even in these cases it must be assumed that the sphalerite was sufficiently solid to enable fracturing to take place somewhat doubtful at the temperatures at which unmixing theoretically occurs). The pattern in Fig. 304 shows the consequence of replacement of sphalerite by chalcopyrite following the cracks of the sphalerite host. Here cracks of sphalerite are clearly invaded and followed by chalcopyrite that has also replaced the ZnS. In the same figure pyrite marginally replaces the sphalerite. Additional observations show veinlets of chalcopyrite following fine cracks of sphalerite (Fig. 305). In cases, the fine veinlets "transgress" into a perfectly oriented pattern of chalcopyrite following the [100] direction of the sphalerite host. Considering the patterns exhibiting fractures, cracks and very fine cracks as "pathways" for the infiltration and replacement of sphalerite by chalcopyrite and the very fact that the chalcopyrite "infiltrated" along these fractures and attained typical forms previously de-

scribed as "ex-solution bodies" of chalcopyrite in sphalerite, it is thus necessary to reconsider the entire gamut of patterns of chalcopyrite present in sphalerite not as "ex-solutions" but as replacements. The often quoted examples of oriented chalcopyrite bodies in the sphalerite as typical examples of ex-solutions should also be reconsidered on the basis that they might represent replacements along penetrability directions of the sphalerite host (Fig. 306). Whereas according to the hypothesis of ex-solution granular bodies of the solute, intergranular to the host are often believed to be due to unmixing which resulted in the "segregation" of the solute as a marginal phase, the above mentioned observations clearly show marginal chalcopyrite replacing sphalerite marginally or along intergranular spaces. Fig. 307 shows chalcopyrite following and replacing sphalerite along intergranular spaces. Furthermore, chalcopyrite occurring as a replacement intergrowth with sphalerite is also marginally replacing ZnS (Figs. 308 and 309). Figs. 310-313 present another textural pattern in support of the replacement interpretation of the socalled "ex-solution bodies" of chalcopyrite in sphalerite. Figs. 310 and 312 show a triangular fracture system in sphalerite occupied in part by chalcopyrite and by galena. The chalcopyrite "ex-solution" bodies are genetically related to a fracture system in the sphalerite and the chalcopyrite partly attain veinform character according to the criteria of Schwartz. This supports non-ex-solution origin in this case for the chalcopyrite in intergrowth with the ZnS. Fig. 310 and, in particular, Fig. 313 show that a fracture line in the sphalerite is occupied also by galena which is in symplectic intergrowth with chalcopyrite of the so-called "ex-solution form". Considering an additional criterion of Schwartz, namely that ex-solutions are single mineral (crystal) bodies, it is clear that the intergrowth of chalcopyrite and galena is incompatible with this criterion of Schwartz. The fact that chalcopyrite and galena follow fracture lines of the sphalerite and the very fact that the socalled "ex-solutions" of chalcopyrite are in symplectic intergrowth, supports the interpretation that both the chalcopyrite bodies and the galena are actually metasomatic fillings, i. e., replacements of the sphalerite. As a corollary to the replacement interpretation of chalcopyrite forming so-called "ex-solution" bodies in sphalerite and in support of the patterns presented in Figs. 310 and 313, there are cases of multiple replacements attaining symplectic character. Fig. 314 shows sphalerite with chalcopyrite and galena intergrowths of the type commonly referred to as "ex-solutions" of chalcopyrite in ZnS. Also similar symplectic intergrowths of galena occur in the sphalerite, in cases the galena following interzonal or intergranular boundaries of the sphalerite. It should be emphasized that the symplectic galena intergrowth in the sphalerite as well as the galena of the intergranular boundaries are both exhibiting replacement intergrowth with the chalcopy45

rite. This pattern again contradicts the set of criteria put forward by Schwartz since if these intergrowths were due to unmixing they should have been single minerals. However, Fig. 314 shows both galena and chalcopyrite in intergrowth and the two in turn with sphalerite, thus supporting a multiple replacement pattern. In addition to the patterns described, cases of chalcopyrite and galena intergrowths of the type referred to as "ex-solutions" exist as separate replacement intergrowths in the same sphalerite (Fig. 315). Another important point concerning the so-called "ex-solutions" of chalcopyrite in sphalerite and vice versa is the percentage of solvent and solute that can be incorporated in the solvent. As stated in the hypothesis of ex-solutions, sphalerite and chalcopyrite form mutual solid solutions with limitations of how much of one phase can be solved in the other. In contradistinction, according to the replacement interpretation such limits are not necessarily a limitation to the extent one mineral can replace the other. The series of patterns presented in Figs. 316-319 clearly shows all transitions of proportional replacement. Fig. 318 shows sphalerite extensively replaced by chalcopyrite with relics of the sphalerite left in the chalcopyrite. Also in the same figure, sphalerite is shown with limited replacement by chalcopyrite. Fig. 319 shows relics of sphalerite in chalcopyrite and represents an extreme case of replacement of sphalerite by chalcopyrite. Such patterns were previously considered as sphalerite exsolutions in the chalcopyrite. However, as the transitions illustrated in Figs. 316-319 clearly support, they are actually replacements of sphalerite by chalcopyrite. In contradistinction, Figs. 316-319 show all transitions of sphalerite from slight replacement by chalcopyrite to advanced replacement of sphalerite by chalcopyrite with only relics of sphalerite preserved in the chalcopyrite. Interesting cases of "ex-solution" bodies of chalcopyrite related to zonal chalcopyrite exist in sphalerite (Fig. 320). Other instances of fine chalcopyrite bodies, zonally distributed in sphalerite, are shown in Figs. 321 and 322. In contradistinction, Figs. 323 and 324 again show the interrelationship between fine zonal chalcopyrite and chalcopyrite "ex-solution-like" bodies in the same sphalerite. In contrast, Fig. 325 shows a sphalerite "core", free of chalcopyrite bodies with an external "zonal" part with fine interspersed chalcopyrite. An "emulsion"-like type of distribution of fine chalcopyrite in blende which is most probably due to separation of chalcopyrite from blende. When considering the phenomenology of fine intergrowths of chalcopyrite "ex-solutions" in sphalerite, particularly from the point of view of ex-solutions versus replacement, interzonal chalcopyrite transgressing to "ex-solution"-shaped bodies as well as fine bodies (often dot- or bleb-shaped) of chalcopyrite interspersed in the sphalerite occur also as interzonal bodies. Fig. 320 shows interzonal chalcopyrite often with transitions to "ex-solution" bodies in sphalerite. The fact that 46

interzonal chalcopyrite in sphalerite attains patterns previously undoubtedly characterized as "ex-solution" deserve further consideration. (i) If the interzonal chalcopyrite and its transitions attaining "ex-solution" forms are actually due to unmixing from the solvent sphalerite (host), the exsolved bodies should be granular following the intergranular of the sphalerite. However, in this case, the chalcopyrite is not restricted to the interzonal spaces only. (ii) If, strictly speaking, the chalcopyrite was an alternating crystallization phase with the sphalerite in the sense that there were successive "zones" of precipitation of sphalerite and chalcopyrite, it would be difficult to reconcile the fact that some of the interzonal chalcopyrite attains "ex-solution" intergrowth forms with the sphalerite. In particular, this is exhibited in the patterns shown in Figs. 323 and 324. (iii) Fine interspersed chalcopyrite (fine dots) occurs in the sphalerite host and in some cases shows a concentration along interzonal spaces of the sphalerite (Fig. 321). This should be considered on one hand in conjunction with the patterns presented in Figs. 323 and 324, and on the other hand in conjunction with the patterns shown in Fig. 325. Fig. 325 shows sphalerite central zones free from ex-solutions and outer zones with emulsion-like bodies of chalcopyrite. As the patterns discussed in (i) and (ii) of this section particularly show, the chalcopyrite has replaced the sphalerite along interzonal penetrability directions, however, not being restricted along these spaces. In contradistinction, the fine dispersed chalcopyrite dots and the emulsoid bodies are difficult to interpret, although in this case a replacement interpretation should not be excluded. As already mentioned, Edwards pointed out that a number of sulfides occur as ex-solutions in sphalerite. Among them he specifically mentioned stannite (Cu2FeSnS4). In contrast to the interpretations by Edwards and others, present observations show stannite as small dots or blebs following a zonal arrangement in the sphalerite. As a consequence of the interzonal replacement patterns of chalcopyrite in sphalerite, in this case the interzonal stannite is also most probably a replacement growth (Fig. 326). Fig. 327 shows stannite replacing sphalerite. The intergrowth often attains symplectic patterns simulating the so-called "ex-solution growths". In addition, Fig. 328 shows stannite replacing sphalerite with protuberances of the marginal stannite attaining symplectic forms in the replaced ZnS. In addition to the patterns described presenting zonal replacement of the sphalerite by stannite, Fig. 326 shows a combination of zonal and symplectic replacement of sphalerite by stannite. This pattern shows that in addition to the penetrability of the stannite replacing solutions along the penetrability zonal direction of the sphalerite, within this zone a symplectic pattern was

formed by replacement which simulates the typical "ex-solution" intergrowths of stannite in sphalerite. Furthermore, discussing the multiple replacement patterns (or multiple "ex-solutions", as most ore microscopists believe), Figs. 135-137 show galena in addition to chalcopyrite as a replacement intergrowth of the sphalerite. Similarly, Fig. 329 shows both chalcopyrite and galena replacing the sphalerite and attaining forms typical of so-called "ex-solutions" of chalcopyrite and galena in ZnS. As a corollary to the replacement nature of "ex-solution"-type intergrowths of galena in sphalerite, Fig. 330 shows galena replacing sphalerite with protuberances which extend along a fracture of the sphalerite and attain typical "ex-solution" symplectic forms. In this chapter on sphalerite replaced by pyrite a number of impressive patterns have been presented (see Chapter 5). Also in the case of fine pyrite Staub (scattered dots) in sphalerite, a replacement process is tentatively suggested as Figs. 331 and 332 indicate. In addition to the patterns of zonal replacement of sphalerite by chalcopyrite and stannite already described, pyrite interzonal replacements of sphalerite also occur, and additional to the zonal pyrite, symplectic replacement structures of symplectic pyrite in sphalerite are also indicated (Fig. 333). As already mentioned, Edwards stated that significant amounts of Fe can be in solid solution in the lattice of sphalerite and as cooling takes place, pyrrhotite can be ex-solved. On the basis of the hypothesis of solid solutions it is possible that pyrite could be formed similarly. In contrast to these interpretations, pyrite replaces sphalerite as the presented patterns clearly indicate (Fig. 334). As also mentioned, in addition to zonal replacement of sphalerite by pyrite and of pyrite Staub in sphalerite (Fig. 331), patterns exhibiting symplectic pyrite in intergrowth with sphalerite are shown in Fig. 334. Here pyrite attains patterns simulating "ex-solutions". The replacement of sphalerite by pyrite could involve the replacement of the Zn atoms by Fe and reorganization of the lattice in addition to volume re-adjustment since there is an excess of S in the pyrite (FeS2) in comparison to sphalerite (ZnS). An alternative interpretation would be dissolution of sphalerite and precipitation of pyrite from solutions. Besides pyrite replacing sphalerite, arsenopyrite (FeAsS) also replaces sphalerite resulting in symplectic intergrowths where the arsenopyrite clearly follows cracks of the ZnS (Fig. 335). In contradistinction to the "ex-solution" patterns of chalcopyrite in sphalerite which in the present effort are interpreted as replacements (see p. 45-46), Fig. 336 shows X-shaped bodies of sphalerite in chalcopyrite. These bodies have been described as star-shaped bodies (Sternchen) and they have been interpreted as exsolutions of sphalerite in chalcopyrite. The fact that these X-shaped bodies are oriented following crystal penetrabilities of the host chalcopyrite and often occur

in trends in the chalcopyrite rather suggest the replacement of chalcopyrite by sphalerite. In contrast to the X-shaped bodies presented in Fig. 336, cases are described of sphalerite replacement by chalcopyrite where relics of the ZnS attain forms resembling "ex-solved" sphalerite in chalcopyrite. In some of these cases it was possible to trace all transitions from sphalerite enclosed in the chalcopyrite to relics of sphalerite in the chalcopyrite identical to the patterns shown in Fig. 337. In addition to the "stars" of sphalerite, Flood (1964) reports sphalerite forming stars in chalcopyrite and cubanite in the copper-zinc mineralization in Trolldalen, Lofoten, N. Norway. Manecki (1965) presented electron-probe results for Zn, Fe, Cu, Sb and S for a traverse across a chalcopyrite with a star-shaped sphalerite inclusion.

(c) Study Cases of Intergrowths Ex-Solutions

Simulating

(i) Cross-shaped sphalerite in pyrite. In continuation of the patterns discussed exhibiting X-shaped sphalerite as replacements of chalcopyrite and relic patterns of sphalerite in the chalcopyrite, cross-shaped patterns (stars) of sphalerite are present in pyrite. It should be pointed out that occasionally these sphalerite stars in pyrite centrally contain chalcopyrite (see schematic diagram, Fig. 338). Two alternative interpretations are put forward: 1. The sphalerite stars in pyrite are ex-solutions in the sense that Zn was a solute in the solvent (pyrite) and with decreasing temperature sphalerite was unmixed, or 2. Sphalerite with chalcopyrite was replaced by the later and often blastic in nature pyrite and the sphalerite "stars" (or more precisely cross-shaped bodies) represent relics which are restricted in the [100] direction of the pyrite host. Such an interpretation explains both the sphalerite star-shaped bodies and their occasional association with chalcopyrite. (ii) Pentlandite - pyrrhotite. In considering the possibilities of solid solutions and ex-solved phases, Edwards and others believe that pentlandite [(Fe, Ni)S] or [(Fe, Ni)9Sg] and pyrrhotite (FeS) provide another case of solid solutions, that pentlandite will dissolve in pyrrhotite to the extent of 40% at temperatures between 425 and 450° C. Furthermore, they believe that pentlandite unmixes upon slow cooling and that it tends to diffuse rapidly into the grain boundaries of the pyrrhotite. Thus, this interpretation of Edwards attempts to explain the presence of pentlandite rims at the margins of the pyrrhotite. As a corollary to Edwards' exsolution interpretation of pentlandite in pyrrhotite, Durazzo and Taylor (1982) support "that experimentation in synthetic systems is essential for a correct interpretation of ore mineral textures such as the pyrrhotite-pentlandite (po-pn) intergrowths of Ni-sulfide deposits". In addition they support that "the cooling 47

rate experiments indicate that massive pn develops by ex-solution between 610°C and about 250° C during slow cooling from temperatures above the mss-pn (monosulfide solid solutions - pentlandite) solvus; coarsened pentlandite forms along basal planes in the monosulfide solid solutions matrix between about 250°C and 150°C; pentlandite "flames" result at 150°C or slightly below". However, as Durazzo and Taylor point out "crossing pentlandite lamellae, obtained isothermally at high degrees of super-saturation and virtually unknown in ores, cannot form during slow cooling". Kaneda et al. (1986) in their paper entitled "Stability of pentlandite in the Fe-Ni-Co-S System", in support of Edwards' interpretation state: "Pentlandite has a wide compositional range in the Fe-Ni-Co-S system and the Fe-Ni-S system. The metal to S atomic ratio is approximately 9:8". The Co content of pentlandite varies from the Co free to nearly a Co9Sg composition. Additionally, Kaneda et al. (1986) support that "pentlandite forms a complete solid solution between (Fe, Ni)9 ± ocSg and Co 9 ± ocSg in the 600-300° C temperature range". It is also supported by them that the pentlandite solid solution decomposes into two fields towards the (Fe, Ni)9Sg and Co 9 S g members at 200° C. The above considerations are supportive of the unmixing interpretation of pentlandite from pyrrhotite, however, a replacement interpretation should not be excluded. In contrast to the orthodox interpretation that the pentlandite is an ex-solution phase of the initial pyrrhotite-pentlandite solid solution, Fig. 339 shows pentlandite flame-shaped bodies radiating from cracks or veinlets in the pyrrhotite occupied with silicates into the pyrrhotite host. Additionally a granular pentlandite (an independently crystallized grain) sends prolongations into a veinlet fracture of the pyrrhotite occupied with gangue. Fig. 340 also shows pyrrhotite with flame pentlandite bodies extending inwards from the margins of the pyrrhotite, and in addition a flame-shaped pentlandite occurs independently in the pyrrhotite. In addition, in this pattern pentlandite occurs marginally to a silicate veinform extension into the pyrrhotite. As the patterns in Figs. 340 and 341 show, flame-shaped pentlandites occur in the central part of the pyrrhotite and cases are shown (Fig. 342) where marginal pentlandite has flame-like extensions in the pyrrhotite. Sometimes the pentlandite is related to cracks or crack-fillings of the pyrrhotite. As already mentioned, the "ex-solutions" of pentlandite in pyrrhotite have been explained by Edwards and others as the unmixing of the solid solution of pentlandite and pyrrhotite. In contrast, the patterns presented support that an alternative interpretation - dissolution of pyrrhotite and precipitation of pentlandite from solutions - is equally plausible. It should be mentioned though, that substitution of Fe by Ni could also result in the formation of pentlandite from pyrrhotite followed by volume and lattice readjustment. (iii) Zinkite as host and as intergrowth in franklinite. Besides the replacement in which ZnS was involved, 48

most impressive replacement patterns are obtained involving ZnO. Fig. 343 shows zinkite (ZnO) replaced by fine lamellar haematite (Fe 2 0 3 ). Here too, contradictory interpretations can be put forward. Either Fe replaces Zn as a result of element migration with the necessary lattice rearrangement, or zinkite has been dissolved and haematite precipitated from subsequent solutions. Considering further the intergrowths of Znoxides the pattern showing zinkite invading and replacing franklinite (ZnFe 2 0 4 ) is of interest (Fig. 344). It should be mentioned that in the same figure "ex-solution" simulating bodies are present in the franklinite. The pattern shown in Fig. 344 is more likely interpreted as due to dissolution of the franklinite and precipitation of zinkite. The alternative interpretation, leaching of Fe and re-adjustment of the lattice results in zinkite, is also plausible. In contradistinction to these two interpretations, the possibility of solid solution of zinkite in franklinite should not be excluded entirely. Another interesting pattern showing the relationship of zinkite and franklinite is illustrated in Fig. 345, where rounded franklinite is enclosed in zinkite. Interspersed fine zinkites in the franklinite also occur, often attaining crystalline form and being oriented in the franklinite. Figs. 346 and 347 show zinkite exhibiting crystalline outlines and following a trend-like pattern (directions) in the franklinite host. As a corollary to the replacement interpretation of franklinite by zinkite, Fig. 348 shows zinkite replacing franklinite in veinform and zinkite following the [111] direction of the franklinite host. The fact that the veinform zinkite has extensions following the [111] direction of the franklinite indicates that both the veinform and the oriented zinkite in the franklinite are due to replacement.

(d) On the Intergrowths of Ilmenite-Haematite Spinels. Considerations of the Solid Solutions of Ilmenite-Haematite The orthodox views consider the solid solutions of ilmenite-haematite as the best understood cases of solid solutions. Edwards (1960) considering the ionic radii of the cations involved and the crystal structure of ilmenite (FeTiOj) and haematite (Fe 2 0 3 ) based on a combination of deductive thinking and experimental data, supports the following: - Ilmenite-haematite forms a continuous solid solution series at temperatures somewhat above 600° C. - On slow cooling it unmixes into two solid solutions: (a) ferriferous ilmenite, the final product contains 6% Fe 2 0 3 and (b) titaniferous haematite, final product of which contains 10% Ti0 2 . - The capacity of these two minerals to form extensive solid solutions is readily understood from comparison of their atomic structure (crystal structure). - In the unit cell of haematite the oxygen atoms are arranged in approximate hexagonal closest packing, in such a way that each Fe 3+ falls between 6 oxygen at-

oms arranged octahedrally. Again according to Edwards, the ilmenite structure is almost identical with the haematite structure but the titanium atoms replacing half of the iron atoms in a strictly ordered sequence. Magnesium, if present, occupies the position of the iron atom. The unit cell of haematite is a = 5.42 Ä and the unit cell of ilmenite a = 5.40 Ä. - The oxide minerals differ from metals in that their component elements are present in the ionized radii state. The ionic radii of the two atoms involved in this substitution are Fe 3+ = 0.67 Ä and Ti 3+ = 0.69 Ä, and they are sufficiently close to cause little distortion of the lattice. On the basis of this deductive thinking it is believed that ilmenite and haematite form a continuous solid solution series which unmixes at temperatures below 600° C. Edwards and the other orthodox opinion holders consider the following often observed textural patterns as a consequence to the outlined hypothesis of solid solutions of ilmenite-haematite. - The haematite bodies are all elongated with their long axes parallel to one another, and to the [0001] direction of the ilmenite, and in the spaces between these rows, further rows of smaller ex-solution bodies are contained. - Irregular shapes of the haematite bodies, however, indicate that they have grown in situ by absorbing exsolving haematite from adjacent ilmenite solid solution. - The ilmenite adjacent to the large haematite bodies is free from even the minutest particles of haematite, indicating that it has been drained by solid diffusion of its content of ex-solvable haematite. The oriented intergrowths resulting from unmixing appear to reflect the presence in both crystal structures of pronounced oxygen planes (O-planes) parallel to their [0001] direction. As put forward by the hypothesis of the ilmenitehaematite solid solution, certain criteria are prerequisite when referring to a pattern of ilmenite-haematite intergrowth as ex-solutions. Fundamental deviations from these criteria would render the pattern problematic when attempting to consider it as ex-solution. In particular, it was stated that the haematite exsolutions are usually elongated bodies with their long axes parallel to one another and to the [0001] direction of the ilmenite (in contrast, the irregular bodies are believed to be in situ growths). In contrast to this criterion, Fig. 349 shows ilmenite (FeTi0 3 ) with curved fractures (cracks) occupied with spinel [hercynite (FeAs 2 0 4 )] and haematite (Fe 2 0 3 ). It should be noted that the haematite itself contains minute ilmenite elongated bodies parallel to the elongation direction of the haematite (these bodies are often referred to as second generation ex-solutions). If the fillings of a fracture by spinel and haematite are considered as a veinform pattern, the haematite and the spinels occurring in this

particular ilmenite pose the question whether they are ex-solutions or subsequent mobilizations. As a corollary to the intergrowth presented above (Fig. 349), Fig. 350 again shows ilmenite with a curved boundary in the main mass of ilmenite. The boundaries of this curved ilmenite and the fracture line in the main mass of the ilmenite are occupied either by isolated spinel or haematite which often extend from these boundaries into the main mass of the ilmenite. The fact that both spinel and haematite occur associated (in a fracture line of the ilmenite) contradicts Schwartz's set of criteria that ex-solutions are single bodies and do not occur in veinform. Additional observations indicate that ilmenite with a pattern of cracks shows spinels extending from one of the cracks into the adjacent ilmenite while haematite partly extends along that crack and extends from it into the adjacent ilmenite (Fig. 351). In the same photomicrograph haematite bodies irregular in distribution are also shown (again including minute bodies of ilmenite). The pattern of these haematite bodies could be better understood as following healed cracks in the ilmenite than as being oriented in accordance to the criterion that the elongated axis of haematite follows the [0001] direction of the ilmenite. As a consequence to the problematic relation of haematite (the so-called ex-solution bodies of haematite in ilmenite) are the cases where the haematite and the spinel present follow curved patterns of fractures in the ilmenite (Fig. 352). It was argued that when a large haematite ex-solution body occurs in ilmenite there are no small haematite bodies adjacent to the large ex-solution in the ilmenite, in the sense that all the haematite in solid solution has been drained to form the larger haematite exsolution in the ilmenite. However, in contrast to this argument, Fig. 353 shows ilmenite with fine haematite bodies (?ex-solution lamellae) which is completely free of haematite as a major crack is approached in the ilmenite. In contrast to the often described haematite-ilmenite ex-solution bodies which could represent a continuous haematite-ilmenite series and which are formed by unmixing at lower temperatures than 600°C, Fig. 354 shows ilmenite with oriented lamellae and spinel in a low grade leuchtenbergite schist indicating that at low metamorphic conditions metasomatic or topometasomatic processes could lead to the formation of ilmenite with (? "ex-solved") haematite and spinel lamellar bodies.

(e) Ti-Mineral Intergrowth Patterns ("Ex-Solutions") in Metamorphic-Metasomatic Rocks Besides the described ilmenite with haematite and spinel lamellar intergrowths occurring in leuchtenbergite-chlorite schist, complex intergrowth patterns of lamellar Ti mineral intergrowths occur in metamor49

phic-metasomatic ore bodies that rather reflect topological conditions than are in agreement with the broad principles of the solid solutions/ex-solutions hypothesis. Considering the textural patterns involved in the transformation of bauxite to emery, Fig. 355 shows pre-corundum haematite with ilmenite lamellar intergrowths. Both the haematite and the ilmenite were subjected to corrosion prior to being included in the later crystalloblastic corundum. This haematite-ilmenite intergrowth has taken place at an earlier phase of the transformation of bauxite to emery. In contrast, intergranular to the corundum crystalloblasts, elongated initial haematite with oriented lamellae of ilmenite occurs. It is more interesting though, that the initial elongated (prismatic) haematite with the ilmenite lamellar intergrowths is replaced by ilmenite (see Fig. 356). As a result of the replacement of haematite-ilmenite by ilmenite, relics of haematite are preserved in the later grown ilmenite which also contains rutile. Topologically in the emery, complex patterns are exhibited where lamellar Ti minerals are in intergrowth with iron oxides. Fig. 357 shows haematite with rutile lamellae oriented parallel to the [h h 2h 1] direction of the haematite. The haematite, perhaps due to further metamorphism, is replaced by magnetite and as a result the rutile lamellae initially associated with haematite are now partly included in the magnetite (as relics of the replacement of haematite by reduction to magnetite). In opposition to the rather simple cases of the rock transformations mentioned (i. e., leuchtenbergite schist and emery) complex patterns of Ti containing minerals lamellar and in intergrowth with other minerals (which also may contain titanium) are often found in davidite mineral paragenetic association. Fig. 358 shows davidite [(AB 3 (0, OH) 7 with A = divalent Fe, trivalent Ce, tetravalent U and also Ca; Β = tetravalent Ti, trivalent Fe and trivalent V. The U content varies from 2 to 25%o] with oriented lamellar bodies of haematite and with fine lamellae of ilmenite parallely oriented to the long direction of the haematite. The haematite lamellar bodies most probably represent replacements of the davidite. Very complex patterns of intergrowth of davidite with ilmenite-haematite and haematite-ilmenite as well as with intergrowths of these minerals with rutile are shown in Fig. 359. As shown by arrow "h", complex replacement patterns of ilmenite by haematite are exhibited. Whereas in this paragenesis there appears to be no "so-called ex-solution" relation of the davidite with the ilmenite-haematite or with rutile, in contradistinction Fig. 360 shows davidite with granular aggregate of haematite with lamellar ilmenite as scattered inclusions, partly replaced by davidite in which they are included. Fig. 361 shows a large rutile grain corroded and partly replaced by later davidite. It should be mentioned that these haematite-ilmenite bodies and 50

the rutiles in the davidite were previously considered as ex-solution bodies formed in the davidite by unmixing. In contradistinction to the patterns discussed, Fig. 362 shows davidite with rutile bodies believed to be formed by unmixing of the davidite under tectonic influence (following a microtectonic line in the davidite). Due to the fact that rutile is scattered outside the tectonic microzone of the davidite, it could also be a replacement pattern.

( f ) Replacement Intergrowths Resembling ExSolution Patterns In contrast to the patterns discussed (replacement versus ex-solutions), a series of patterns will be presented where complex replacement processes are involved and where often the mineral relationships do not allow an interpretation based on the concept of solid solutions and ex-solutions. Fig. 363 shows crystalline pegmatitic uraninite (UO z ) in intergrowth with crystalline columbite (a mixture of niobite ((Fe, Mn)Nb 2 0 6 ) and tantalite (Fe, Mn)Ta 2 0 6 ) where the uraninite replaces the columbite. Excellent patterns of epitactic uraninite on columbite are also described by Strunz (1961). Further patterns of replacement intergrowths simulating ex-solutions are the cases of galena (PbS) replacing pyrrhotite (FeS) in which case dissolution of the pyrrhotite took place rather than migration of Fe out of the pyrrhotite and introduction of Pb. Additional patterns of this intergrowth were discussed in Chapter 5. A very complex replacement is magnetite (Fe 3 0 4 ) in intergrowths with olivine (Mg 2 FeSi0 4 ), in which case unmixing of a fyalite-rich molecule could theoretically result in magnetite ex-solutions. If such a process could take place, fyalite would rather be formed in forsterite. As Fig. 364 shows, additionally to the fine lamellae of magnetite in the olivine, magnetite following a crack of the silicate is also exhibited, thus implying replacement of the olivine by magnetite, forming melts or ?solutions. Besides these complex replacement patterns, Bi is replaced by lamellae of sphalerite (ZnS) clearly oriented in the Bi (Fig. 365). However, typical replacement patterns of Bi by sphalerite were discussed in Chapter 4. In addition to the case of native Bi being replaced by sphalerite and as a corollary to the cases where a native metal is involved in the replacement processes, Fig. 366 shows cuprite (Cu 2 0) where native Cu has been formed topologically due to reduction rather than ex-solution. In an even more complicated way native Ag is formed topologically in pyrargyrite (Ag3SbS3) possibly by leaching out of Sb and S (Fig. 367).

An alternative interpretation could be "topological" dissolution of pyrargyrite and the precipitation from ?colloidal solutions of native silver. This interpretation is supported by the "zonal" distribution of the silver in

the pyrargyrite as well as by the pattern of the native silver (now crystalline due to the "Kristallisationsfreudigkeit").

51

Chapter 7

(a) Oriented

Symplectites

Symplectites

Discussing the intergrowth patterns common in ore minerals, a different type of oriented symplectites is of significance besides the replacement patterns and the replacement patterns that were previously considered as "ex-solutions". Most of theses symplectites are products of replacement or mutual replacement of the minerals intergrown. Crystalline uraninite (U0 2 ) is intergrown with skeletal crystals of siegenite (with the general formula (R2R23S4) where R 2 = Fe, Ni, Co, Cu and R 3 = Co, Ni, Cr). The intergrowth simulates a simultaneous crystallization intergrowth. However, another interpretation is that the siegenite has replaced the uraninite in the sense that crystalline uraninite was dissolved and siegenite precipitated from solutions. The possibility that element migration played a major role is most unlikely (meaning that diffusion in the solid state of the elements occurred comprising siegenite and that the U of the uraninite was leached out). On the basis of the above given interpretations it is most likely that dissolutions of the uraninite occurred and as a result of replacement a skeletal crystal of siegenite developed (the siegenite represents an incompletely developed siegenite idiomorphic growth or even a type of idioblast). As a corollary to the replacement nature of the siegenite in the uraninite, Fig. 368 shows a developed siegenite crystal replacing and maintaining uraninite as interzonal relics within the siegenite. In contrast to the possibility of consecutive crystallization zones of siegenite and uraninite, the replacement interpretation is supported by another skeletal growth of siegenite in intergrowth with uraninite (Figs. 369 and 370). Fig. 371 shows veinform siegenite transecting the uraninite and sending protuberances following the crystal faces direction of the uranium mineral. As a further consequence to the replacement explanation, Fig. 372 shows siegenite partly in complex intergrowth with uraninite and with siegenite protuberances attaining veinform extension in the uraninite. Also Fig. 373 shows a veinform type of replacement of uraninite by siegenite, which, partly as a crack filling and partly as a replacement, substitutes the uranium mineral. As already described, columbite is replaced by uraninite by dissolving of columbite and precipitating uran52

inite from solutions. Fig. 374 shows columbite replaced by uraninite which attains partly well-developed crystal faces. In contradistinction, Fig. 375 shows uraninite replacing columbite in which case the uraninite is oriented within the host. Similarly, Fig. 376 shows uraninite replacing columbite; the uraninite substitutes are often oriented parallel. Comparing the chemical composition of the columbite which can be understood as a mixture of niobite ((Fe,Mn)Nb 2 0 6 ) and tantalite ( ( F e . M ^ T a ^ ) , with that of uraninite (U0 2 ) - despite there being an interrelationship in accordance to the periodic system between some of the elements comprising the columbite and U - the dissolution of the columbite took most probably place before the precipitation of the uraninite. In spite of the mentioned interrelationship in accordance with the empirical laws of the periodic system between elements comprising columbite and U, it is most unlikely that sufficient U or uraninite can exist in solid solution in a system of a continuous solid solution series of niobite and tantalite. Despite the furnished interpretation, such a system (solid solution of uraninite and niobite-tantalite) should theoretically not be excluded and should perhaps be investigated further. As in the case of columbite replaced by uraninite, which in addition to the characteristic replacement patterns and to the oriented infiltration replacement patterns (which simulate ex-solutions) and which also occurs as epitactic uraninite on columbite, similar epitactic and oriented replacements of jacobsite (MnFe 2 0 4 ) occur on braunite [3(Mn,Fe) 2 0 3 , MnSi0 3 ], see Fig. 377. Comparing the composition of these two minerals, jacobsite may be derived from braunite either by element migration (?leaching out of Si) and reorganization of the lattice and volume re-adjustments, or, in this case too, by dissolution of braunite and precipitation from solutions of jacobsite. Both the epitactic patterns of jacobsite and the oriented symplectic intergrowths could be understood as replacements in this case. The oriented symplectites presented so far were intergrowths of minerals that were somehow paragenetically interrelated and had geochemical interrelationships, as it was the case of siegenite in symplectic in-

tergrowth with uraninite, columbite and uraninite, braunite and jacobsite. In contradistinction, symplectites occur between minerals that are paragenetically not very closely interrelated, i. e., titanite/magnetite, ilmenite/pyrrhotite and sphalerite/gangue (quartz). On the other hand, complex symplectic intergrowths occur related to complex replacements such as in the case of symplectic sphalerite in a paragenesis of tetrahedrite, chalcopyrite and gudmundite as well as impressive patterns of galena in symplectic intergrowth replacing sphalerite where blastic pyrite might also be present. Fig. 378 shows magnetite in oriented symplectic intergrowth with titanite. As their chemical composition shows, these oriented symplectic intergrowths are not simultaneous crystallizations under eutectic conditions but they represent replacement of the titanite (CaTi[0Si0 4 ]) by magnetite (Fe 3 0 4 ). It is suggested that the titanite was dissolved prior to the precipitation of magnetite (most probably by solutions). Another type of pattern which might be the result of complex replacement processes is shown in Figs. 379 and 380. Fig. 379 shows pyrrhotite partly altered to marcasite (bird's eye structure is shown) with oriented lamellae of ilmenite (FeTi0 3 ). It is most likely that the original intergrowth was that of ilmenite with magnetite where the magnetite was replaced by pyrrhotite (FeS), thus, the lamellae of ilmenite represent relics of the initial titanomagnetite. As Fig. 380 shows, the ilmenite lamellae are present as oriented intergrowths both in the pyrrhotite and the adjacent silicate. Symplectic intergrowths of sphalerite with gangue (quartz) often simulate graphic-like intergrowths of the sphalerite with the quartz (Figs. 381 and 382). Other complex symplectites of quartz and sphalerite are shown in Fig. 383 where galena is also symplectically intergrown with the sphalerite. In contrast to these rather simplified sphalerite symplectites, sphalerite may assume graphic-like symplectic patterns with chalcopyrite in a paragenesis where blastic gudmundite also replaces the chalcopyrite (Fig. 384). Another complex replacement pattern of graphiclike galena replacing the sphalerite and partly the blastic pyrite is shown in Fig. 385. When considering these oriented symplectic intergrowths which in cases simulate ex-solved phases or simultaneous eutectic crystallizations it must be recalled that in most cases they are intergrowths of mineral phases, incompatible with the concepts of solid solutions, or continuous isomorphous crystallization series. The patterns depicted represent rather complex replacement processes and as such they are in accordance to comparable graphic-like intergrowths of quartz/feldspar which have also been interpreted to represent replacement processes, see Drescher-Kaden (1948, 1969), and Augustithis (1962, 1973, 1985). In addition to the symplectic intergrowths of sphalerite where a graphic-like pattern is exhibited and in which case the sphalerite is considered to represent

skeletal crystals in intergrowth with the quartz and where it is believed the sphalerite did not attain a fully developed crystalline outline, comparable graphic intergrowths of chalcopyrite and pyrrhotite are shown in Fig. 386. Graphic intergrowths of ore minerals with quartz are common. This is supported by a number of additional observations. Fig. 387 exhibits graphic symplectite of bornite in intergrowth with quartz; the bornite is partly replaced by chalcopyrite. Furthermore, complex replacement patterns of silicates (quartz) with galena and chalcopyrite are illustrated in Fig. 388, again exhibiting a graphic-like intergrowth pattern. Also skeletal Bi crystals in intergrowth with silicate exist and Fig. 389 shows a graphic Bi crystal in intergrowth with quartz, where the Bi crystal exhibits some well-developed crystal faces. In contrast, Fig. 390 shows graphic or micrographic intergrowth with quartz where the native Bi did not attain well-developed crystal faces. Considering the graphic or micrographic intergrowths of ZnS, FeS, CuFeS 2 , Cu 2 FeS 4 , PbS and Bi with quartz (Si0 2 ) it should be pointed out that no eutectic crystallization of each or any of the compounds with S i 0 2 is possible and, as already emphasized, these graphic intergrowths with quartz represent replacement intergrowths where a fully developed form failed to be attained and skeletal crystals are exhibited. In addition to the graphic-like patterns of ore minerals intergrown with quartz as mentioned above, symplectites with quartz are rather rarely found but nevertheless graphic-like intergrown: gold/quartz (Fig. 391), sylvanite/quartz (Fig. 392), petzite/quartz (Figs. 393 and 394) exist. Here too, neither Au, AuAgTe2 nor Ag3AuTe2 form eutectic crystallization with S i 0 2 and, as mentioned, are considered to represent replacement skeletal patterns with the quartz, where the ore minerals failed to attain fully developed crystalline forms. Comparable graphic-like replacement patterns are shown where clausthalite PbSe is intergrown with calcite CaC0 3 (Figs. 395 and 396). In this case again no eutectic crystallization system was found between PbSe and CaC0 3 and the exhibited patterns are skeletal crystals of clausthalite in intergrowth with calcite. Very impressive graphic-like intergrowths of uraninite with microcline are shown in Figs. 397-401, where all transitions are exhibited between graphiclike intergrowths of uraninite with microcline and idiomorphic uraninite in the microcline of the same polished section. Furthermore as this series shows, between graphic-myrmekitic symplectite (Fig. 397) and idiomorphic uraninite (Fig. 401), intermediate graphic intergrowths and skeletal forms of uraninite in intergrowth with microcline are exhibited. Pyrite (FeS 2 ) skeletal growths often occur in intergrowth with sphalerite (ZnS) and all transitions are exhibited between skeletal crystals to idiomorphic (or idioblastic) pyrite. The skeletal pyrite occasionally exhibits typical graphic-like patterns which are not due to 53

simultaneous crystallization of FeS 2 and ZnS. The two alternative interpretations often put forward - substitution of Zn by Fe and reconstitution of the lattice with possible volume re-adjustment - may lead to the replacement of sphalerite by pyrite. The other alternative (dissolution of sphalerite and precipitation of pyrite from solutions) is equally plausible. The transition from skeletal pyrite to idiomorphic pyrite is exhibited in Figs. 402 and 403. In contradistinction to the graphic-like intergrowths of pyrite and sphalerite there are patterns exhibiting skeletal graphic-like intergrowths of pyrite and graphic-like intergrowths of chalcopyrite, both in symplectic intergrowth with sphalerite. As Fig. 404 shows in addition to the graphic pyrite and chalcopyrite, pyrite and chalcopyrite are also symplectically intergrown in a background of sphalerite. As the cases of graphic uraninite in microcline and graphic skeletal pyrite and chalcopyrite in sphalerite suggest, often the tendency to idiomorphism by replacement occurs, and all intermediate phases from the graphic-myrmekitic to graphic skeletal and finally to idiomorphic idioblastic forms can be observed. Examples are the cases of well-developed uraninite idiomorphs and pyrite idioblasts shown in Figs. 401 and 403, respectively. Considering the wide spectrum of pyrite graphic/skeletal growths in addition to the patterns exhibited, cases are shown of intergranular (actually intercrystalline) pyrargyrite (Ag3SbS3) with proustite relics (Ag 3 AsS 3 ) intergrown with later Wasserkies pyrite, exhibiting skeletal intergrowths assuming graphic forms and with transitions to idiomorphic pyrite crystals (Fig. 405). In contradistinction to the symplectic patterns discussed and, in particular, to gold/silicate intergrowths where skeletal gold occurred in silicates, cases are presented where lamellae of prismatic silicates are in intergrowth with tellurides and native gold or with native gold only, see Figs. 406 and 407, respectively. Such complex intergrowths may be explained as the result of later crystallization of the gold occupying the interspaces between the silicate laths or elongated prismatic crystals. No case is due to eutectic simultaneous crystallization of the gold minerals and the silicates. As already described, haematite with oriented lamellae of rutile occurs in the Naxos emery (a pattern which is quite common in other occurrences as well) and which indicates the relationship of rutile (Ti0 2 ) and haematite (Fe 2 0 3 ). In contrast, Figs. 408 and 409 show large grains of rutile associated with davidite in which complex myrmekitic-like symplectites of rutile and haematite occur. The following alternative interpretations are put forward in an attempt to explain these complex myrmekitoid haematite-rutile intergrowths: (i) Haematite was dissolved in the rutile and by unmixing gave rise to these irregular in orientation and complex in pattern myrmekitoid haematites in the ru54

tile. This is in accordance with the solid solutions' hypothesis. However, it should be emphasized that the pattern produced does not resemble or approach lamellar orientation of haematite in the rutile as would be expected by the unmixing hypothesis. (ii) Metasomatic replacement of the rutile took place by later haematite in the sense that solutions dissolved the rutile and subsequent precipitation of haematite from solutions occurred. (iii) Another alternative is that haematite/rutile formed a eutectoid which by simultaneous crystallization would result in a complex symplectite. However, it should be pointed out that the patterns which have been produced do not resemble typical graphic or micrographic patterns so characteristic of theoretical simultaneous crystallization. In contrast to the complex and rather uncertain in origin myrmekitoid rutile-haematite symplectites, metasomatic replacement patterns of rutile in intergrowth with wiikite are illustrated in Fig. 410 where in addition to the symplectic rutile idioblastic later rutile is also shown. Considering the approximate composition of wiikite [(Y, Na, Ca, U)(Nb, Ta, Ti, Fe) 2 (0, OH) 2 ] and that of rutile (Ti0 2 ), it is obvious that Ti is common in both minerals. However, it is difficult to imagine that most of the other elements included in the wiikite are leached out and that Ti contributed to the formation of the blastic and symplectic rutile associated with the wiikite (Fig. 410). It is more likely that the dissolution of wiikite occurred prior to the precipitation from solutions of the rutile but it should not be ruled out that some of the titanium might be derived from the wiikite. The symplectites discussed so far - genetic interpretations were difficult in many cases and admittedly dubious in certain cases - nevertheless allow to furnish explanations where replacement processes played a role in the formation of the presented symplectic patterns. As a corollary to the replacement interpretation of most symplectites mentioned so far is the case of neodigenite (Cu2S) replaced by covellite (CuS) where covellite corrodes, replaces and infiltrates into the neodigenite resulting in a symplectite of neodigenite and covellite. Comparing their composition, the replacement could be understood stoichiometrically by leaching out of Cu and re-adjusting the lattice and the volume. In contrast to these stoichiometric balancing interpretations, it is more probable that dissolution of the neodigenite occurred prior to the precipitation from solutions of covellite. As Figs. 411-413 show, corrosion of the neodigenite occurred. This is evidenced by the relic of neodigenite in the main mass of covellite; evidence of replacement and infiltration of the covellite forming soutions is provided by the extension of the covellite from the main marginal mass of covellite to the symplectic covellite in intergrowth with the neodigenite. As especially Fig. 413 shows, the infiltra-

tion of covellite forming solutions has exploited penetrability directions of the neodigenite host.

(b) Symplectic Intergrowths due to Superheating - Mainly under Volcanic Action or High Temperature Fluids In contrast to the replacement symplectite patterns formed under relatively low temperature conditions, cases of symplectites formed under elevated temperatures exist and several examples are mentioned in literature. In the present effort, study cases will be presented: one referring to skeletal and symplectic intergrowths formed under crystallization of high temperature fluids, the other referring to the case of superheating of magnetite or titanomagnetite under the influence of volcanic conditions. Fig. 414 shows intercrystalline patterns consisting of symplectites and isolated crystallization of sphalerite (ZnS), bravoite (Ni-pyrite, in cases with Co) and occasionally of galena (PbS). Fig. 414 shows intercrystalline patterns (i. e., occupying the spaces between crystals and crystal grains for a more elaborate discussion on this subject see Schachner-Korn, 1960) consisting of sphalerite sphaeroids and sphalerite symplectites intergrown with galena or bravoite. Scattered sphalerite and bravoite occur in this intercrystalline pattern. In some cases, in the central part of the sphalerite sphaeroids, bravoite occurs. In other instances though, it is symplectic with galena (Fig. 415). In these ring and "feather" structures sphalerite occurs in symplectic intergrowth with gangue (Fig. 416). In great detail Ramdohr (1960) and Frenzel (1953) studied the superheating of titanomagnetite under the effects of volcanic activity and the transformation of original titanomagnetites into haematite (Fe 2 0 3 ) and pseudobrookite (Fe 2 TiO s ). According to Ramdohr, the magnetite is transformed into haematite and the ilmenite (TiFe0 3 ) into pseudobrookite. A series of observations is introduced to show the formation of symplectic patterns formed under superheating. Fig. 417 shows idiomorphic magnetite in the central part of which, due to superheating, symplectic haematite has been formed. Also symplectic with the central part of the magnetite and between the intergrowths of haematite with magnetite, pseudobrookite occurs. Often the pseudobrookite symplectites do not follow the original oriented pattern of the ilmenite in the original magnetite (titanomagnetite). In the marginal parts of the magnetite, oriented maghemite is also present. As in particular Fig. 418 shows, the pseudobrookite is only irregularly symplectic with the magnetite and in some cases follows the interspaces between the haematite. The pattern of the haematite does not follow nor agree with the oriented pattern of heat-martitization, although the two processes, especially that of heatmartitization, and the symplectic haematite shown in

Figs. 417-419 occur basically by superheating, and where oxidation of the magnetite is assumed to have taken place. Fig. 420 shows that pseudobrookite follows the oriented direction within the magnetite and might, after all, be a replacement product of the original ilmenite, since the pseudobrookite is also oriented in the [111] direction of the magnetite as it is supposed to be in the orientation of the ilmenite in the original titanomagnetite. The heat haematite seems to follow the magnetite areas between the original lamellar ilmenite. If the process had advanced it would have resulted in oriented pseudobrookites in haematite. The symplectic patterns of heat haematite in the magnetite and the irregular in orientation symplectic pseudobrookites indicate, however, that by the heat transformation the original oriented pattern of ilmenite and magnetite was not followed strictly by the pseudobrookite (Figs. 417 and 418). Furthermore, the heat haematite symplectites in the magnetite could be seen as transitory forms in the complete replacement of magnetite due to haematite formation by heating. In contrast to the symplectites due to superheating (volcanic action), spinellids unmixed intergrowths of chromian spinels sensu stricto and magnetite, or chromian magnetite are reported to occur by Eales et al. (1988) in the Natal-Namaqua Mobile Belt (S. Africa) in metamorphic rocks that according to them attained temperatures sufficiently high (upper amphibolite facies) for the formation of homogenous Al-CrFe 3+ Ti spinel phases with compositions not matched in slowly cooled igneous rocks. In addition to the interpretation provided by Eales et al., the reader is referred to Chapter 6, where chromites with rutiles and spinel oriented bodies were discussed.

(c) Myrmekitic and Graphic Symplectites The term myrmekitic texture has been applied to describe the symplectic intergrowth of worm-like quartz bodies in plagioclase which is in contact with K-feldspar. Becke and a number of researchers explained myrmekite as the result of reaction of older K-feldspar with younger plagioclase. In contrast to this interpretation, mainly Drescher-Kaden (1948, 1969), Augustithis (1962, 1973), and Collins (1987) explained myrmekites as the result of synantetic reaction of older plagioclase in contact with younger K-feldspar. In contradistinction to the myrmekites in petrography, in ore microscopy many ore minerals exhibit complex myrmekitic or myrmekitic-like intergrowths which, due to the great reactivity of the ore minerals, are often difficult to explain. In the present effort, many of these symplectic intergrowths will be presented and tentative attempts will be made to provide interpretations. Magnetite in contact with skarn marbles has prolongations which attain myrmekitic symplectic inter55

growths of the magnetite with the marble (Fig. 421). Such patterns are difficult to interpret and two opposing alternatives are given below: (i) According to the magmatic granitic hypothesis due to the granitic magmatic intrusion, magnetite melts infiltrated and replaced the initial limestones which have been metamorphosed due to contact metamorphism into marbles. The magnetite is metasomatically introduced under pyrothermal conditions (contact metamorphic) derived from the granite either in melts or in fluids. (ii) Skarns and especially magnetite are part of the basic front which is released due to granitization. The basic front is released from the granitized mass and the country rocks (mainly) due to the assimilation of basic xenoliths or pregranitic basic rocks which have been granitized. It should be mentioned that in 1947 Reynolds had already introduced the concept of basic front release from granites due to assimilation of basic xenoliths by the granite. The basic front mainly consists of Mg and Fe released from granitized initial sediments. Thus, the myrmekitic intergrowth of magnetite is interpreted as due to the basic front solutions attacking the marble. In addition, the pattern in Fig. 422 shows prolongations of magnetite into the marble with extensions attaining myrmekitic symplectic intergrowth with the marble. Additionally to the magnetite, maghemite exhibits symplectic intergrowth with the marble. In contrast to the "pyrometasomatic symplectic patterns" of magnetite with marbles already mentioned, very complex patterns of bornite (Cu2FeS4) with bismuthinite (Bi 2 S 3 ) exist. The bismuthinite most probably replaces the bornite in the sense that the bornite was dissolved prior to the precipitation of the bismuthinite (or in cases emplectite) from solutions. However, the pattern is more complex since the bismuthinite/bornite intergrowth seems to replace gel pitchblende. The relationship between gel pitchblende and this myrmekitic pattern though is dubious since gel pitchblende engulfed and surrounded a protuberance of the intergrowth (which in turn sends veinlet-like extension or ?relic into the gel pitchblende, see Fig. 423). Furthermore, Fig. 424 again shows a symplectite of bornite with bismuthinite. However, in this case the bismuthinite marginally replaces the bornite and sends protuberances into the bornite. The bismuthinite intergrowths with bornite often follow crystallographic directions of the host in addition to the presence of worm-like bodies of bismuthinite in the bomite host. The oriented bismuthinite lamellae in the bornite are not ex-solutions since bornite and bismuthinite do not form a system of mutual solid solutions. As already pointed out, they are the result of the replacement of the bornite in the sense that they are prolongations of marginal bismuthinite replacing the bornite. In this connection, it should be emphasized that the textural association of bornite with pitchblende and the fact that bornite probably replaces the uraninite rather 56

supports a late phase bornite formation in the telethermal phase. In contrast to the cases described where the bornite was replaced by bismuthinite, cases are observed where the bornite is in symplectic intergrowth with rutile. As Fig. 62 shows, rutile in intergrowths with silicate is partly replaced by bornite. Sphalerite often exhibits symplectic intergrowths with gangue (silicates) and often is symplectically replaced by pyrrhotite (Fig. 425). Considering the chemical composition of sphalerite (ZnS) and of pyrrhotite (FeS) - despite the fact that they could form a mutual solid solution - the pattern presented in Fig. 425 represents most probably a replacement of sphalerite by pyrrhotite, either in the sense that sphalerite was dissolved and pyrrhotite precipitated from solutions or that element mobilization resulted in which case the Zn is replaced by Fe. As a corollary to the symplectic graphic pattern of sphalerite with silicates, Fig. 426 also shows graphic symplectic sphalerite and pyrrhotite with silicates (quartz). When discussing the replacement of sphalerite by galena a wide spectrum of patterns has been presented to illustrate the process. The series of Figs. 427-429 shows sphalerite in graphic intergrowth with galena in which galena replaces ZnS. In contrast to the typical graphic-like patterns of galena and sphalerite (Figs. 427 and 428), Fig. 429 shows rounded forms of sphalerite (exhibiting graphic patterns) enclosed and partly engulfed by later galena. These graphic-like structures of galena in intergrowth with sphalerite, or the reverse case, of sphalerite in galena are, as emphasized, due to replacement of sphalerite by galena. Occasionally germanite exhibits complex patterns of replacement simulating the typical graphic intergrowths where germanite (Cu 3 (Ge, Fe)S 4 ) can be replaced by skeletal crystals of pyrite (FeS2), see Figs. 430 and 431, and also by galena (PbS). Sometimes the patterns exhibited indicate complex successive replacement where the germanite is replaced by galena, which in turn is transected by a later veinlet of tennantite (Cu3SbS3), Fig. 432. As the chemical composition of germanite, galena and tennantite shows, the patterns are rather to be interpreted as due to successive replacement than as due to ex-solutions since it is difficult to imagine solid solutions of germanite with these minerals. When discussing the so-called "ex-solutions" of chalcopyrite and galena in sphalerite it could be seen that replacement was favoured over "ex-solutions" from solid solutions. Furthermore, considering the graphic and micrographic pattern of sphalerite with silicates and the graphic intergrowths of galena and chalcopyrite in sphalerite, evidence of replacement supports that these graphic symplectites are also due to substitution rather than the result of eutectic crystallization. In particular Fig. 433 shows galena bodies in sphalerite transgressing to graphic intergrowths of galena in the sphalerite. Similarly, a veinlet transecting

sphalerite attains forms typically micrographic, again supporting that these are due to replacement and not to eutectic crystallization of simultaneously crystallized mineral phases (Fig. 434). As already pointed out, the myrmekitic symplectites of galena in intergrowth with sphalerite are not "ex-solutions" and similarly the micrographic galena intergrowths in sphalerite are due to replacement. A corollary to the replacement nature of the myrmekitic or graphic-like intergrowth of galena with sphalerite is shown in Fig. 435. There marginal galena is extending and replacing adjacent sphalerite attaining myrmekitic or in cases graphic-like forms. In cases patterns are shown exhibiting myrmekiticmicrographic-like intergrowths of galena and chalcopyrite in sphalerite (Fig. 436). The fact that the micrographic bodies of galena and chalcopyrite are sometimes intergrown support replacement over "ex-solution" or eutectic crystallization since as Schwartz pointed out, ex-solutions are single mineral bodies and do not consist of different intergrown mineral phases. As discussed in the case of the so-called "ex-solutions" of chalcopyrite in sphalerite (Chapter 5) and also in the case of micrographic intergrowth of chalcopyrite with sphalerite, evidence is presented of transitions of grain chalcopyrite (independent crystallization) into protuberances of chalcopyrite attaining graphic or myrmekitic intergrowth patterns (Fig. 437). In addition, a corollary to the replacement patterns of symplectic micrographic or myrmekitic-like bodies of galena and chalcopyrite in sphalerite is the case where veinlets of chalcopyrite or galena transect the sphalerite and show transitions in symplectites (Fig. 438). In contrast to the symplectites that are due to infiltration and replacement, examples are shown where symplectites are relics of replacement. Fig. 439 shows boraite (Cu 2 FeS 4 ) as well as relics of bomite replacement by chalcopyrite (CuFeS2). Fig. 440 shows bornite relics in chalcopyrite and graphic-like intergrowths of bornite in the silicates. As it was the case with galena and chalcopyrite which sent prolongations into the adjacent sphalerite, similarly chalcopyrite sends extensions into the adjacent enargite which attain myrmekitic-like intergrowth patterns (Fig. 441). In addition, fine symplectites of chalcopyrite occur in enargite as well as a crack filling of chalcopyrite again in the enargite (Fig. 442). Furthermore, Fig. 443 shows symplectic intergrowths of chalcopyrite in enargite where the intergrowth attains a graphic- or myrmekitic-like form of chalcopyrite intergrown with enargite. Comparing the chemical composition of enargite (CU3ASS4) and chalcopyrite (CuFeS2), replacement of As by Fe and reorganization of the lattice with volume re-adjustment could be interpreted as the replacement of enargite by chalcopyrite. This could be achieved either by element migration or most likely by dissolu-

tion of enargite and precipitation from solutions of chalcopyrite. When looking at symplectic patterns, most impressive intergrowths of bornite (Cu 2 FeS 4 ) and chalcocite (Cu2S) are common and complex patterns are often exhibited. Admittedly, the interpretation of the patterns is often dubious, nevertheless, cases are presented where evidence supports a replacement of the bornite by the later chalcocite. Fig. 444 shows symplectic chalcocite with bornite. Often the chalcocite tends to follow directions within the bornite. The pattern is tentatively interpreted as replacement of bomite by the later chalcocite. The replacement process means a relative enrichment in Cu and the removal of Fe. The increase in copper is in agreement with the general trend of copper increase with the supergene sequence of copper minerals (the tendency is for bornite to be replaced by chalcopyrite and chalcopyrite by chalcocite, see Chapter 5). Fig. 445 shows chalcocite sending protuberances into the bornite and attaining a symplectic form. Such cases support the replacement of bornite by chalcocite. Whether the replacement takes place by element migration or by dissolution of bornite and precipitation of chalcocite from solutions is difficult to decide for all cases. Fig. 445 shows extensions of chalcocite in the adjacent bornite attaining veinform character and supporting a replacement interpretation. Such patterns would be difficult to explain as ex-solutions or as due to simultaneous eutectic crystallization. There are cases, however, where the symplectites are very complex. As Fig. 446 shows, despite the symplectic chalcocite sending vein-like extensions into the adjacent bornite, it could be argued that the bornite also replaces the chalcocite (in the sense that the bornite sends protuberances into the chalcocite). In contrast, the most convincing evidence that these myrmekiticlike symplectites are due to the replacement of bornite by chalcocite, is provided by the pattern shown in Fig. 447, which shows marginal chalcocite replacing bornite with an extension into the bornite attaining myrmekitic form (see also diagram, Fig. 448). Evidence of the replacement nature of the myrmekitic symplectites of bornite with chalcocite are provided by the patterns in Figs. 449 and 450. Fig. 449 shows chalcocite enclosing relics of bornite attaining myrmekitic-like intergrowth with the chalcocite which also sends extensions into the adjacent bornite. Similarly, Fig. 450 shows chalcocite with the myrmekitic-like relics of bornite in the chalcocite and chalcocite extension veinform into the bornite. The fact that the same chalcocite encloses relics in myrmekitic symplectic intergrowth with the bornite and simultaneously sends extensions into the adjacent bornite supports the myrmekitic-like intergrowths of the bornite as being replacement relics of the bornite by chalcocite. Fig. 451 shows myrmekitic-like relics of bornite transgressing into ring bornite relic-forms and thus, it simultaneously presents atoll-type replacement of bor57

nite in addition to the myrmekitic (worm-like) relic forms of the bornite in the later chalcocite. Further evidence that the myrmekitic-symplectites of bornite with chalcocite are due to the replacement of the bornite by chalcocite and that the bornite wormlike bodies are relics of the processes is supported by a series of patterns exhibited in Figs. 452-454. Especially the patterns in Figs. 452 and 453 show bornite enclosed and partly replaced by chalcocite and at the same time chalcocite extending and invading the bornite along penetrability directions resulting in myrmekitic intergrowths of bornite and chalcocite. Similarly, transitions are shown of bomites enclosed, corroded and replaced by chalcocite. However, due to a more advanced phase of replacement, parts of bornite are represented in the chalcocite as worm-like relics (resulting in myrmekitic symplectic intergrowth of bornite and chalcocite). Thus, the transitions of bornite to symplectic bornite can be traced as bornite is subjected to a more advanced phase of replacement by chalcocite (compare Figs. 453 and 455). On the basis of these comparative studies it is tentatively suggested that typical myrmekitic symplectites of bornite with chalcocite are the result of the replacement of bornite by chalcocite and that the bornite worm-like bodies are actually relics. Adding to the presented textural patterns of myrmekitic intergrowth of ore minerals, ore microscopic studies by Schmidt-Eisenlohr (1966) report chalcocite, bornite and other minerals in myrmekitic intergrowths in breccia pipes from the Cu ore deposit of Messina in the Transvaal (S. Africa). Saager (1968) also describes myrmekitic intergrowths of linnaeite, gold and pyrrhotite in a detrital grain from the Basal Reef (Witwatersrand). Cases of bornite replaced by haematite are also observed where again the myrmekitic-like bodies are relics of the replacement of the bornite by the haematite. Comparable myrmekitic intergrowths of pyrrhotite and galena are shown in Figs. 456 and 457. They are again due to the replacement of pyrrhotite by galena. In addition to the myrmekitic intergrowths presented above, Brodin (1960) in his contribution "A myrmekitic intergrowth of galena and chalcocite" describes a graphic myrmekitic intergrowth of chalcocite-galena in the ores of Ken-Shanyk, Tatas Alatau, Russia, believed to result from the replacement of cubic (now orthorhombic) chalcocite by galena. According to Brodin the process involved an exchange of cations between solid and liquid phases which left the ionic framework of each grain unaltered. Furthermore, when discussing myrmekitic symplectites intergrowths those are of interest which involve Bi minerals. In particular as Fig. 458 shows, tetrahedrite (Cu g Sb 2 S 7 ) is in an atoll-type manner replaced by wittichenite (Cu 3 BiS 3 ) which in turn is symplectically replaced by native Bi which replaces the tetrahedrite as well. Considering the composition of these minerals, the replacement of tetrahedrite by wittichenite mainly involves the removal of Sb and the in58

troduction of Bi with the necessary volume and lattice re-adjustments. In opposition, the replacement of wittichenite by native Bi involves the removal of Cu and S. However, the entire pattern of symplectic replacements might be better explained by assuming dissolution of tetrahedrite and precipitation of wittichenite by subsequent solutions and consequent dissolution of tetrahedrite and wittichenite and precipitation from solutions of native Bi. The pattern of symplectic intergrowth shown in Figs. 459 and 460 (actually a detail of Fig. 458) could be interpreted in a similar way,. Comparable to the pattern in Fig. 458, Fig. 461 shows relics of tetrahedrite (attaining myrmekitic symplectic form) which is replaced by wittichenite. The wittichenite and the tetrahedrite are symplectically replaced by native Bi. In addition, as Fig. 461 shows, impressive symplectic replacements of wittichenite by native Bi resemble an atoll-type replacement pattern in which the wittichenite is substituted by the native Bi. Another impressive symplectic intergrowth pattern is shown where chloanthite [(Ni,Co)As3] replaces native Bi. Occasionally, the chloanthite in addition to the symplectic intergrowth with Bi is also marginal to the native metal (Fig. 462). As an explanation for this pattern, it is suggested that Bi was dissolved and chloanthite was precipitated subsequently from solutions. As already mentioned, chalcopyrite (CuFeS2) symplectically replaces bornite (Cu 2 FeS 4 ), see Figs. 463 and 464. (Veinlets of chalcocite transect both the bornite and the chalcopyrite.) The replacement pattern suggests a relative increase in Cu as far as the replacement series of bornite-chalcopyrite-chalcocite is concerned. However, the relationship of bornite-chalcopyrite and their intergrowth patterns can be very complex. They have been studied both texturally and experimentally by a fair number of researchers in the past: Van der Veen (1925), Wandke (1925), Newhouse (1928), Gruner (1929), Ray (1930), Schwartz (1930), Sugaki (1951, 1951, 1955, 1965), Ramdohr (1960), Edwards (1960), Morimoto et al. (1960), Prouvost (1960), Brett (1962, 1963, 1964), Ziebold and Ogilvie (1963), Yund and Kullerud (1966), Tufar (1967), Sillitoe and Clark (1969), Barton (1970), Kullerud (1971), MacLean et al. (1972), Birchenall (1973), Cabri (1973), Verhoeven (1975), Durazzo and Taylor (1978), Samal and Gilevich (1978). Concerning experimental work on the relationship of bornite-chalcopyrite, of particular interest are some of the conclusions presented by Durazzo and Taylor (1982) in their contribution entitled "Experimental exsolution textures in the system bornite-chalcopyrite: genetic implications concerning natural ores". According to them, "textural interpretation of the ore mineral assemblages, such as bornite-chalcopyrite (bn-cpy) intergrowths, should be based on definite experimentation originate by: (1) simultaneous precipitation; (2) ex-solution during slow cooling from the above solvus; and (3) metamorphism to temperatures above about

250° C. Widmanstätten textures are not compatible with slow cooling but indicate: (1) low-temperature replacement of bornite; or (2) ex-solution of chalcopyrite lamellae from anomalous bornite heated to around 200-250°C during mild metamorphism." Considering the experimental work and the conclusions of Durazzo and Taylor (1982), despite the title of the contribution referring to "Experimental ex-solution textures" as mentioned, the Widmanstätten textures of chalcopyrite in bornite are interpreted as also representing replacement of bornite by chalcopyrite. In contrast to the experimental work which supported replacement of bornite by chalcopyrite, Amcoff (1988) in his contribution entitled: "Experimental replacement of chalcopyrite by bornite: textural and chemical changes during a solid-state process", states that chalcopyrite was reacted with covellite and with chalcocite, respectively, between 200 and 500° C and an ensuing solid-state replacement of chalcopyrite by bornite was observed. Considering the significance of Amcoffs experimental work - oriented bornite bodies (lamellae) were produced by replacement in the chalcopyrite - it is believed that it is necessary to quote some extracts from Amcoffs paper as a corollary to the author's "replacement - ex-solution" interpretation of comparable intergrowths in many ore minerals. Thus, according to Amcoff, it is stated: "the relatively oxidizing conditions of the reaction chalcopyrite+covellite result in "massive replacement", lacking structural control, where bomite and pyrite form complex intergrowth textures in chalcopyrite. Bornite nucleates around growing pyrite aggregate because of the release of copper and a decrease in volume. Diffusion of sulfur along grain boundaries and fractures largely controls the textural development. Reaction under the relatively reducing conditions involving chalcopyrite+chalcocite results in replacement of chalcopyrite in the sequence where chalcopyrite is replaced by bornite, below about 355°C, and by intermediate solid solution (ISS) and later bornite, above 355°C. The textural development, changing from replacement, apparently uninfluenced by directional properties in the host, to semi-oriented replacement, is structurally controlled. This suggests that the process is governed by diffusion of copper and iron through a sulfur framework. It is suggested that the observed formation of oriented bornite lamellae in chalcopyrite and in ISS during the chalcopyrite+chalcocite reaction may be explained by "replacement ex-solution" at constant temperature." Another common symplectic pattern is the intergrowth of bornite (Cu2FeS4) with chalcopyrite (CuFeS2) which replaces the bornite and often results in myrmekitoid graphic-like textures of bornite associated with chalcopyrite. Considering the formulae of bornite and chalcopyrite, the introduction of Cu and lattice reorganization could account for the replacement. An alternative interpretation is dissolution of bornite and introduction of chalcopyrite precipitated

from subsequent solutions. The replacement of bornite by chalcopyrite and the symplectic patterns produced are shown in Figs. 463-465. An additional pattern of symplectite products of replacement processes is myrmekitoid pyrite replacing magnetite which, in turn, is replaced by haematite and as a result myrmekitoid pyrite is partly engulfed by haematite (Fig. 466). The replacement of magnetite by haematite may result in a pattern where only relics of magnetite are left in the haematite and where the pyrite originally associated with the magnetite, is now maintained as myrmekitoid or graphic-like structure in the haematite (Figs. 467 and 468). Comparing the formulae of magnetite (Fe 3 0 4 ) and pyrite (FeS2) an initial symplectite of magnetite and pyrite is produced most probably by dissolution of magnetite and precipitation of pyrite from subsequent solutions. In opposition, the replacement of magnetite by haematite (Fe 2 0 3 ) is due to oxidation. As a result of this replacement the pyrite, originally associated with the magnetite, is now associated with the haematite. Very impressive symplectites are produced when clausthalite (PbSe) is replaced by covellite (CuS). Such patterns are probably best understood by assuming dissolution of clausthalite and precipitation of covellite from subsequent solutions (Fig. 469 and its detailed view, Fig. 470). In contrast to the myrmekitoid and graphic-like symplectites which are the result of replacement and substitution, impressive symplectites are also produced by direct crystallization either of fluids or hypothermal solutions impregnating sediments or volcanics. Fig. 471 shows sphalerite (ZnS) myrmekitoid in form in intergrowth with bravoite ((FeS2) with Ni and Co) and occasionally with galena (PbS). The complexity of these intergrowths is further illustrated in Figs. 472474.

(d) Breakdown

Symplectites

Ramdohr mentioned that a breakdown symplectite occurs which, according to the studies of Quensel et al. (1937) and Wretblad, is allemontite consisting of arsenic and antimony-arsenic (stibarsen) as allemontite-III, or of antimony-arsenic (stibarsen) and antimonite as allemontite-I, and finally the compound antimony-arsen (stibarsen, AsSb) can occur alone as allemontite-II. Edwards also mentioned that the natural intergrowths consist of either (i) a mixture of the Sb-phase and the AsSb-phase (allemontite-I) or (ii) a mixture of the As-phase and the AsSb-phase (allemontite-III). Less commonly, according to Edwards, the homogenous AsSb-phase occurs alone (allemontite-II). Very impressive textural patterns of breakdown symplectite (allemontite) are shown in Figs. 475-477, where additionally to arsenic (black) and stibarsen (white) a mixed crystalline phase of the two is also present. In contrast, most typical breakdown symplec59

tites of arsenic (black) and stibarsen (white) are exhibited in the series of Figs. 478-480. It should be emphasized that the patterns presented are not eutectic crystallizations but as mentioned by Ramdohr Zerfallstruktur breakdown structures or rather breakdown symplectites. As a corollary to the breakdown structure interpretation, Edwards presents a sphaeroidal structure (Fig. 481) which, according to Edwards, shows "zonal" texture of allemontite-III from Pribram, showing rhythmic distribution of AsSb-phase (black) in As-phase (white), etched with K^S. Such a textural pattern could be interpreted as representing colloform zoning or perhaps a breakdown structure of a gel sphaeroid. Nysten (1986) presented most interesting myrmekitic symplectites in which native gold is intergrown with stibnite. The intergrowth has been interpreted as a breakdown product of autostibnite at low temperatures. Hawley and Haw (1957) and in particular Hawley (1961) also describe pseudo-eutectic intergrowths involving niccolite, chalcopyrite, pyrrhotite and maucherite developed mainly by replacement of gersdorffite, the latter participating essentially as a component of the two-or-three-phase aggregates. According to Hawley, niccolite-pyrrhotite and maucherite-pyrrhotite intergrowths may form by the breakdown of ferroan gersdorffite or by normal replacement of pyrrhotite by niccolite. The conditions for the formation of these intergrowths are an increase in temperature and a decrease in sulfur (and arsenic) vapor pressure. The intergrowths are explained as late-stage alterations. In contrast to the breakdown symplectites described so far, Amcoff and Figueiredo (1990) presented "Mechanisms of retrograde changes in oxide minerals from the Proterozoic Serrote da Laje deposit, Ν. E. Brazil." According to them, the deposit is situated in a mafic-ultramafic layered sill and oxidation and cooling leading to successively decreasing diffusion rates resulted in disequilibrium on the microscale. Amcoff and Figueiredo recorded the following cooling and oxidation phases, particularly for pleonaste: (i) rapid change in composition between coarse grains in a granoblastic magnetite host, indicating metamorphic peak conditions; (ii) coarse lamellae in magnetite, indicating commencement of ex-solution, and (iii) composite pleonaste-ilmenite lamellae in magnetite, indicating oxidation ex-solution. In addition to the symplectites of covellite replacing clausthalite as shown in Fig. 470, malachite (Cu[(0H) 2 /C0 3 ]) marginally replaces cuprite (Cu 2 0) and sends extensions into the cuprite attaining myrmekitoid structures. The symplectites are most probably formed by dissolution of cuprite and precipitation from subsequent solutions of malachite, although it is possible that Cu which is present in the cuprite was only partly removed and has contributed to the formation of malachite (Fig. 482). 60

In contrast to the case presented in Fig. 482, Fig. 483 shows chalcocite (Cu2S) invaded along cracks and replaced by cuprite (Cu 2 0). The replacement in this case was probably due to solutions enriched in Cu in comparison to the host chalcocite.

(e) Relic

Symplectites

In contradistinction to the oriented relic magnetite preserved as relics when magnetite was replaced by biastic pyrite, myrmekitoid symplectites may also result where initial magnetite (Fe 3 0 4 ) is replaced by crystalloblastic pyrite (FeS2) (Fig. 484). Comparable myrmekitoid relic structures to those of magnetite in pyrite are found in stannite (Cu2FeSnS4) marginally replaced by cassiterite (Sn0 2 ), where myrmekitoid structures of stannite are left as relics in cassiterite (Fig. 485) which, as mentioned, has marginally replaced the stannite, resulting in a myrmekiticlike stannite/cassiterite symplectite.

( f ) Ore Minerals

- Silicates

-

Symplectites

Symplectic intergrowths of ore minerals with silicates are of great significance. In this connection it is worth commenting on the relative attention paid to the gangue minerals by ore microscopy, and to the opaque minerals by petrography of thin sections of rocks, respectively. The ore microscopist considers what he refers to as gangue as synonymous with something not interesting to him, and similarly the thin section petrographer considers opaque minerals as something not really of concern to him. (Of course there are exception but the exceptions do not disprove the rule.) The reason is that the ore microscopist is so specialized in his subject that the associations and nature of ore minerals are of primary importance to him. On the other hand, the petrographer is so involved in rock-forming minerals and their textures that he prefers to ignore the insignificant (to him) ore minerals present in the thin section. Even those experienced on both ore minerals and rock petrography usually change attitude depending on whether they are examining polished sections of ores or thin sections. However, ore deposits are aberrant concentrations of metallic elements in a geoenvironment consisting of large entities of rocks it is not surprising that the rock environment usually represents either the parental material for the segregated ore accumulations, or in the extreme case, the geoenvironment in which the ore was accommodated (see Chapter 1) If the geoenvironment is the parental material of the ore accumulation, the relationship of the ore minerals and the gangue will be of great significance and in this case, the intergrowths of the ore minerals and gangue are of genetic importance, hence, the significance of the ore mineral gangue symplectites. If, on the other

band, the ore geoenvironment of the ore accumulation is the recipient environment in which the ore mineral was accommodated, both the reaction of the rock environment with the ore and the reactions and synantectic intergrowths (including the symplectites) of ore mineral and gangue are of importance. However, it should be noted that ore minerals and gangue are not genetically separated. In most cases, gangue and ore minerals form an integral part of the ore deposit and they might belong to the same paragenetic association or the same mineral assemblage since, after all, transparent minerals may consist of metallic elements. On the basis of these considerations an attempt will be made to present textural patterns of ore minerals in intergrowths with gangue and each pattern will be discussed on its own. Fig. 486 shows chromospinel with a decoloration margin with myrmekitic symplectites of silicates and chromite and its decoloration margin which is magnetite in this case. It should be noted that the myrmekitic symplectite is not restricted to the decolorized margin of the chromospinel but myrmekitic symplectites also occur with the chromite itself. The textural patterns presented in Figs. 486 and 487 actually represent chromospinels of an olivine bomb (mantle derivative) which has been picked up by the basalt. The chromospinel is thus a xenocryst in the basalt and the decoloration margin (magnetite) of the chromite is the result of the reaction of the chromospinel with the basaltic melt. Similarly the myrmekitic symplectite with the chromospinel and the magnetite is a synantetic reaction of the chromite/magnetitic margin with the basaltic melts which have corroded and infiltrated the chromite and magnetite, resulting in a myrmekitic symplectite of chromite/magnetite and the silicates infiltrated into the ore minerals after dissolution occurred. It should be noted that especially the magnetite margin shows extensive corrosion and dissolution phenomena which occur prior to the infiltration of the basaltic melts into the dissolution channels. These channels were produced on the chromospinel and magnetite as a result of their reaction with the basaltic melt. The above described symplectic pattern clearly represents the reaction of xenocryst (mechanically transported mineral phase) in the melt geoenvironment which is genetically not the parent material for chromospinel. Thus, the basaltic melt is the recipient geoenvironment of the chromite xenocryst and as such, the symplectite produced is the result of the reaction of "transported" chromospinel and its recipient geoenvironment. In contrast, serpentine/chromite symplectites may be formed as the result of the reaction between chromite and serpentine, which is an alteration product of the dunite and which represents the parental geoenvironment in which the chromite was formed. Fig. 488 shows chromite with myrmekitic serpentine symplectically intergrown with the chromite (which is decolor-

ized, see Augustithis, 1960). In this particular pattern, marginal serpentine extends along cracks and dissolution channels into the chromite and results in a chromite/serpentine symplectic myrmekite (see Fig. 489). As Augustithis pointed out, the decoloration of the chromite is due to element leaching as a result of the reaction of chromite with serpentine formation under the serpentinization of the dunite. Fig. 490 also shows serpentine extending from the margins of the chromite inwards to the host spinel and following its dissolution channels of the chromite. The infiltration of the serpentine in this case too, attains myrmekitoid shape (in intergrowth with the chromite). In contrast, Fig. 491 shows serpentine myrmekites following penetrability directions of the chromite host, resulting in an oriented pattern of serpentine with the chromite. Further examination of the intergrowths of serpentine with the chromite (Fig. 492) shows fractured chromite invaded by serpentine following the fractures and extending into the chromospinel attaining myrmekitic forms symplectic with the chromite. Fig. 493 shows granular chromite peripheral to massive chromite (the granular chromite is most probably due to tectonic influence) and with serpentine occupying the intergranular between the chromite and transgressing into serpentine, myrmekitic in shape, in intergrowth with the massive chromite. In this connection it should be mentioned that the transition of Wiederverkittungs serpentine to myrmekitic-symplectite with the chromite is also observed (see Fig. 494). As mentioned, the myrmekitic symplectic serpentine in intergrowth with the chromite is the result of chromite dissolution and "mobilization" of the serpentine into the channels formed by corrosion and dissolution of the chromospinel (Fig. 490). The evidence that marginal, intergranular and Wiederverkittungs serpentine extends into the chromite and attains myrmekitic forms as well as the fact that corrosion and dissolution of the chromite occurred prior to the "infiltration" and mobilization of the serpentine into the dissolution channels of the chromite, contradict the often furnished interpretation that the chromite was simultaneously crystallized with olivine which subsequently was serpentinized. Olivine being incorporated into chromite and zonal distribution of olivine occurring in chromite (the olivine grains can exhibit corroded outlines) will be discussed in Chapter 9. Sulfides are another group of minerals showing very complex symplectic intergrowths with the gangue minerals. Fig. 495 shows sphalerite corroded and infiltrated by gangue attaining myrmekitic symplectic forms with the sphalerite. In contrast to the myrmekitic patterns, chalcocite may be in intergrowth with gangue intergranular between pyrite grains which indicate corrosion and replacement by the later chalcocite (Fig. 496). Some very impressive replacement intergrowths develop when pyrite with marginal neodigenite is infiltrated and replaced by later gangue, and partly by the 61

trated and replaced by later gangue, and partly by the neodigenite (Fig. 497). In opposition to the myrmekitic patterns of gangue (silicates) in intergrowth with pyrite, cases are exhibited of pyrite invaded by quartz which follows the cleavage patterns of the pyrite host (Fig. 498). Similarly, silicates (quartz) are marginal to myrmekitic intergrowth with pyrrhotite and the silicates simultaneously follow the cleavage patterns of the pyrrhotite host supporting, here too, the relationship of myrmekitic symplectic silicates and silicates oriented within the host (Fig. 499). An interesting case of gangue-ore minerals intergrowth is shown in Fig. 500. Zonal cobaltite is infiltrated and replaced by silicates which follow the interzonal spaces of the cobaltite crystals. The symplectites of ore and gangue (silicates) discussed so far clearly show infiltration and replacement of the ores by later melts or solutions which result in symplectic intergrowths of the silicate with the ore minerals. In addition to gangue (quartz) replacing pyrite where the silicate followed the cleavage of the sulfide, pyrite is also replaced by needles of tourmaline which may intersect one another (Fig. 501). In contrast to cases where silicate invades the ore minerals, it is also common that silicates intergrown with ore minerals represent part of the recipient environment into which the ore minerals have infiltrated and were precipitated from solutions. Fig. 502 shows an original mica schist invaded and "impregnated" by ore forming solutions as a result of which mica flakes, often following the original orientation of the mica schist, are included in later formed sphalerite which is partly replaced by chalcopyrite. In addition to mica schist components included as oriented relics in ore minerals, cases are shown where galena encloses and replaces hornblendes to some extend (Fig. 503). Figs. 504 and 505 show another example of a recipient rock invaded and replaced by ore forming solutions which had resulted in the precipitation of antimonite intergranular and occasionally replacing the silicates. In contrast to the symplectites with the silicate of the gangue representing the recipient geoenvironment in which the ore mineral was precipitated from solutions or was mobilized, interesting intergrowths of silicates (quartz) paragenetically associated with the ore mineral are shown in the following figures. Fig. 506 shows idiomorphic quartz included as a whole or galena following the intergranular between well-developed quartz crystals (Figs. 506, 507). Similar patterns of quartz often exhibiting idiomorphic shapes enclosed in galena and bournonite are illustrated in Fig. 508. There are also reports of idiomorphic quartz being indented due to corrosion and included in native gold (Fig. 509). 62

In addition, sylvinite includes fine idiomorphic quartz and has invaded aggregates consisting of wellcrystallized quartz, often extending into the intergranular of the well-developed quartz crystal aggregate. In other instances, marginal to the sylvinite, argentite also extends into the intergranular of the quartz aggregate (Fig. 510). The intergrowths of silicates paragenetically associated with ore minerals can attain very complex patterns. Figs. 511 and 512 show gangue minerals enclosed in sphalerite where the sphalerite has also interzonally infiltrated or is associated with the gangue, the zonal-in-distribution sphalerite exhibits myrmekiticlike forms. As especially Fig. 511 shows the myrmekitic in shape sphalerite associated with the gangue sometimes represents replacement by the sphalerite, since sphalerite is extending into the included gangue. Occasionally well-developed quartz crystals are included and replaced by sphalerite and often a skeletal outline of the quartz is preserved in the sphalerite (Fig. 513). As pointed out, quartz paragenetically associated as a synchronous formation with the ore minerals and exhibiting idiomorphism may be included in a variety of ore minerals. Fig. 514, for example, shows idiomorphic quartz included in tetrahedrite and with pyrite, occasionally occupying the intergranular between the quartz and the other sulfides. Skeletal in form symplectites may result where ore minerals occupy the intercrystalline-intergranular spaces between calcite paragenetically associated with pyrargyrite which also attains skeletal shape and in some cases encloses and partly replaces calcite (gangue). It should be noted that Wasserkies pyrite is formed in the pyrargyrite. Comparable patterns are formed where skeletal pyrargyrite occupies the spaces between well-formed quartz and calcite gangue. Fig. 515 shows pyrargyrite in replacement intergrowth with later argentite, attaining skeletal forms and occupying the intercrystalline spaces between well-formed gangue components. As already mentioned, complex symplectic intergrowths may result where the recipient geoenvironment to the precipitation of the ore solutions is a metamorphic rock. In addition to Rammelsberg's schist, "impregnated" with sphalerite and copper minerals, cases are exhibited where a garnetiferous rock is transected and replaced by pyrite and sphalerite. Fig. 516 shows garnet invaded and replaced by sphalerite and the garnet corroded and replaced by later sulfide. Fig. 517 shows sphalerite replacing the garnet and veins of sphalerite and pyrite transecting the garnetiferous rock. It should be noted that the sphalerite veinlets transect the garnets sending protuberances into the latter which clearly indicates replacement of the garnets by the sphalerite resulting in sphalerite/pyrite symplectites with the garnet.

In contrast to the metamorphic recipient rock types (schists, hornblendite, gametiferous metamorphics) cases are described of an aggregate consisting of interpenetrating fine prismatic silicates which in cases

are associated with pyrite relics and which are invaded and replaced by stromeyerite often occupying the intergranular between the fine prismatic silicates,

63

Chapter 8

Crystalloblastesis

The term blastese was first introduced by Becke (1904) to describe the sprouting of a crystal in a rock which remained in the solid state during metamorphism. The expression blastese, crystalloblastesis, has been applied widely to describe crystal growths mainly of metamorphic rocks and it has been used in this sense for the rest of the century. The term crystalloblastesis has been greatly differentiated and attained a structural-textural descriptive meaning. By adding a Greek or Latin word as prefix, a terminology has developed which describes the wide spectrum of textural patterns mainly in metamorphics. Thus, a number of terms has been introduced such as idioblastic, hypidioblastic, allotrioblastic, xenoblastic, granoblastic, granonematoblastic, granolepidoblastic, nematoblastic, velonoblastic, porphyroblastic, megablastic, panidioblastic, radially-fibroblastic, cumuloporphyroblastic, and many more (Augustithis, 1990). In contrast to the use of the term crystalloblastesis, the terms derived as differentiations of it have been used to describe textures in metamorphic rocks in the strict sense of metamorphism. In this connection, Sander (1930) widely applied "parablastic growths" to patterns of blastic growths formed concurrently with deformation. Whereas the term crystalloblastesis was acceptable to describe feldspar growths in gneisses and in other metamorphics it was for the first time introduced by Drescher-Kaden in a series of studies starting 1926 and continuing till 1982. It was applied as comparable and commensurable to the crystalloblastic (feldspar) growths in metamorphics, and also to granites. However, it is in his book "Feldspat-Quarz-Reaktionsgefüge der Granite und Gneise", (1948), that crystalloblastic growths have been introduced as a concept in a plutonic rock as the granites were considered to be. Drescher-Kaden's revolutionary concept was strongly supported by Erdmannsdörffer in his paper "Die Rolle der Endoblastese in Granit" published in 1950. Augustithis in his doctoral thesis in 1956 (under supervision of Prof. Drescher-Kaden), entitled "Über Blastese in Gesteinen unterschiedlicher Genese, Granit, Metamorphit (Smirgel), Basalt", introduced the term Tekoblastese (tecoblastesis) to describe crystalloblastic growths in basalt where the plagioclase "phenocrysts" were of lower anorthosite content than the groundmass feldspars and where the components of 64

the basaltic groundmass were corroded, included and greatly assimilated by the later-formed plagioclase tecoblasts. Again Augustithis (1979) introduced the term "crystalloblastic" in his book "Atlas of the Textural Patterns of Basic and Ultrabasic Rocks and their Genetic Significance" to describe crystal growths in many basic and ultrabasic rocks such as para-gabbros, paranoides, anorthosites and in many ultrametamorphic in origin basic and ultrabasic rocks. In contrast to "rock" petrography, ore petrography has accepted blastic growths of ore minerals particularly in metamorphics from the very beginning of the introduction of "Blastese" by Becke in 1904. The wellknown octahedra of magnetite in chlorite schist were among the earliest accepted cases of crystalloblastic growths in metamorphic rocks. Betechtin et al. (1957) in their textbook on mineralogy furnished impressive examples of magnetite octahedra metacrystals (crystalloblasts) in chlorite schists. Ramdohr often used the term crystalloblast in his early effort in 1920 to describe especially pyrite crystalloblastic growths in ore mineral accumulations (deposits). Despite the use of the term of crystalloblastesis by Ramdohr (1960), and the attempts of Vokes to present several studies on the blastic growth of pyrites, the term crystalloblastesis has been relatively restricted, despite its acceptance, in ore microscopy, the reason being that there is an inherent difficulty in applying it to ore deposits most of which represent metasomatic concentrations or segregations and, as Daly had already emphasized in 1917, ore deposits associated with granitic intrusions are metasomatic mobilizations. Many ore deposits are metasomatic in origin, where ore mineral growths are due to metasomatic mobilizations, i. e., often introduction of solutions and material into an environment which is, broadly speaking, in the solid state. Thus, there is some difficulty in distinguishing between crystallization from metasomatic solutions, replacement and crystalloblastic growths in senso strictu. Considering the definition of Becke that crystalloblastic growths are the sprouting of crystals while the rock was in the solid state, there is a close relationship between a growth of an ore mineral by replacement (where a pre-existing mineral phase or mineral aggregate is engulfed and assimilated by a new growth due to replacement), and a crystalloblastic

growth in the strict sense of the term. The concept of the sprouting of a mineral could similarly be considered to represent the formation of a mineral under replacement processes. In spite of these considerations, there is ample room for the use of crystalloblastic growths in senso strictu in the ore petrography, too. An attempt will be made to present certain crystalloblastic patterns (in senso strictu) and if necessary to relate them to replacement and metasomatic processes. Fig. 518 shows chromite grains corroded and enclosed in a crystalloblastic growth of garnet which in turn is surrounded and enclosed by later-formed chalcopyrite (precipitation from solutions). It should be noted that the chalcopyrite has partly invaded and replaced the garnet and has also replaced the chromite grains that are either enclosed or associated with the garnet crystalloblasts. In this case, there is a typical garnet crystalloblastic growth and simultaneous replacement of the garnet and chromite by chalcopyrite. In contrast, Fig. 519 shows prismatic gudmundite crystalloblasts extending marginally from a silicate into tetrahedrite margined by chalcopyrite. Often the gudmundite crystalloblastesis is restricted in the chalcopyrite margin or, in cases, occurring in tetrahedrite, it happens that it has a chalcopyrite margin. Comparing the composition of tetrahedrite (Cu8Sb2S7) and that of its margin, in this particular case of chalcopyrite (CuFeS2), it is possible that the tetrahedrite is replaced by chalcopyrite. In this case either dissolution of the tetrahedrite occurred and precipitation of the chalcopyrite from solutions or Sb was removed and Fe introduced with the necessary volume re-adjustment and lattice reorganization. Considering further the composition of crystalloblastic gudmundite (FeSbS) it is possible that idioblastic gudmundite formed as a result of the replacement reaction of tetrahedrite/chalcopyrite whereby the removed Sb and the introduced Fe in a background of S resulted in the blastic growth of gudmundite (in the marginal chalcopyrite replacing tetrahedrite and in the tetrahedrite in which case the gudmindite has a margin of chalcopyrite, see Fig. 519 and especially Fig. 520). Considering further the relationship of replacement and crystalloblastic growths, Fig. 521 shows crystalloblastic safflorite (CoAs2) blastically cutting through the zoning of smaltite crystals [(Ni,Co)As3], Comparing the composition of smaltite and safflorite which blastically cuts through the smaltite zoning replacing it, it is possible that the blastic growth is due to the removal of Ni and the resultant development of a new growth where volume and lattice re-adjustment took place. However, the blastic growth of safflorite in smaltite is essentially a replacement process despite the texturally blastic pattern shown. In contrast to the safflorite crystalloblasts described transecting the zoning of smaltite, the common case of safflorite star-twins could be seen as safflorite crystalloblastesis in the gangue (Fig. 522). In contradistinction safflorite stars (twins) (CoAs2) may occur as crys-

talloblastic replacement growths in rammelsbergite (NiAs2) in which case Ni is substituted by Co resulting in a replacement blastic growth of safflorite in rammelsbergite (Figs. 523 and 524). Despite being replacements of rammelsbergite, the safflorite stars show the characteristics of idioblastic growths as textural patterns prove. Wolframite ((Fe,Mn)W0 4 ) often occurs as crystalloblastic growths transecting marcasite (FeS2), as shown in Figs. 525 and 526. In cases though, the wolframite might be associated with scheelite as a "cavity filling" in marcasite. The formation of wolframite in marcasite is seen as the result of blastic wolframite growth in which dissolution of the marcasite might have taken place before the blastic crystallization of the wolframite. However, as the presence of marcasite relics in the wolframite crystalloblast show (Figs. 525, 526) the latter is most probably an idioblastic growth, in contrast to the case exhibited in Fig. 527, where the wolframite and the associated scheelite as cavity fillings might represent a precipitation from solutions in a cavity space formed by dissolution of the marcasite. Complex patterns of sphalerite (ZnS) replacement by AsFeS are common. Idioblastic arsenopyrite including relics of replaced sphalerite are exhibited (Fig. 528). In contradistinction, subidioblastic arsenopyrite might include relics of sphalerite in a sphalerite background in which chalcopyrite symplectites occur with the ZnS. However, a clear case of arsenopyrite idioblast cutting through a sphalerite band is shown in Fig. 529 and in this case the sphalerite, engulfed by the idioblastic arsenopyrite, shows corroded outlines. A common pattern of arsenopyrite often preserving relics of sphalerite and consisting of a radiating arsenopyrite crystalloblast is shown in Figs. 530 and 531. In cases a "dentritic" or "feather-like" pattern of arsenopyrite crystalloblasts associated with gangue is shown (Fig. 532), demonstrating the complexity of crystalloblastic growth patterns. The relationship between blastesis and replacement was already emphasized as well as the fact that many blastic growths due to replacement, exhibiting crystal sprouting, occurred by dissolution of the pre-existing minerals and by precipitation of the substitute from solutions. However, crystalloblastic growths in sensu stricto have been described and presenting crystalloblastic growths, the differentiation between replacement crystalloblast and crystalloblast in sensu stricto will be mentioned. A case of replacement and crystal development (a type of parablastic crystalloblastesis) is described by Menyaylova (1967). Here, replacement of the propylitic minerals and the development of cinnabar (HgS) phenocrysts in ferromagnesian carbonates, dickite and quartz is restricted to the tectonically active zones. A series of subidioblasts to idioblasts of haematite associated with selenium minerals is presented in Figs. 533-536. In particular Fig. 533 shows a haematite 65

subidioblast transecting native selenium minerals. Fig. 534 and 535 show haematite subidioblasts developed in clausthalite. Specifically Fig. 535 shows the interpenetration of the haematite crystalloblasts enclosing clausthalite relics developed in clausthalite. In contrast, Fig. 536 shows haematite crystalloblast transecting the contact of clausthalite and calcite. Haematite is a typical crystalloblastic growth which consists of elongated prismatic subidioblasts and encloses relics of enargite and furthermore exhibits an interpenetration pattern developing in a background of enargite (Fig. 537). In certain cases a haematite crystalloblast sends out radiating prismatic haematites (Fig. 538) and patterns exhibiting interpenetration of elongated prismatic (in section) haematites are shown. Perhaps one of the most impressive typical (in sensu stricto) crystalloblastic growths is the development of perfect magnetite octahedra in leuchtenbergite schist from Hada Budussa, Adola District, Ethiopia (described by Augustithis in 1964). As mentioned, the magnetite includes oriented ilmenite bodies margined by spinels (see Chapter 6). In contrast to the perfect magnetite octahedra ilmenite crystalloblasts with oriented lamellae were found in the leuchtenbergite schist from Hada Budassa. However, xenomorphic ilmenite crystalloblasts and ilmenite crystalloblasts engulfing and containing relics of silicates are also exhibited (Fig. 539). As emphasized (Chapter 5), pyrite often replaces other minerals either by element exchange or by dissolution of minerals and precipitation of pyrite from solutions. Furthermore considering the crystalloblastesis of pyrite, cases of pyrite blastic growths due to replacement and pyrite crystalloblastesis sensu stricto will be discussed. Fig. 540 shows niccolite (NiAs) replaced by pyrite (FeS2) xenoblasts where dissolution of the niccolite occurred prior to the precipitation in the "voids" of pyrite from solutions. Comparable cases of pyrite crystalloblastesis exhibiting relics of colloform structure are shown to replace and transect marcasite (Fig. 541). In connection with crystalloblastic growths due to replacement, a series of xenoblastic patterns of pyrite will be presented which shows the wide range of minerals that can be replaced blastically by pyrite. Fig. 542 shows magnetite, a derivative of induced reduction of haematite (originally colloform iron hydro-oxides) replaced by pyrite, dissolution of the magnetite took place prior to precipitation of pyrite from solutions. Very complex replacement patterns of pyrite crystalloblast enclosing, corroding and substituting sphalerite are shown in Figs. 543 and 544. In Fig. 543 pyrite is shown with symplectic relics of sphalerite marginal to the pyrite and prolongations of the pyrite in symplectic replacement (intergrowth) with the sphalerite. In Fig. 544 though, fragmented sphalerite and silicates enclosed and replaced by later xenoblastic pyrite are depicted. In addition, xenoblastic pyrite shows relics of sphalerite in the later pyrite.

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Instances of subidioblastic pyrite crystalloblastically replacing granular pyrrhotite are shown in Fig. 545. It should be noted that relics of pyrrhotite are present in the crystalloblastic pyrite. Considering the composition of pyrrhotite (FeS) and pyrite (FeS2) stoichiometric introduction of sulfur might account for this pattern but the alternative interpretation - dissolution of the pyrrhotite and precipitation of the pyrite from solutions should not be ruled out. However, more complex replacement processes have been involved in the case of subidioblastic pyrite growth in a background of chalcocite in which the subidioblastic pyrite poikiloblast maintains relics of the chalcocite. When the composition of chalcocite (Cu2S) and pyrite is compared, in the case of the blastic growth-replacement processes that resulted in the patterns shown in Fig. 546, removal of Cu and introduction of Fe as well as volume and lattice re-adjustment could lead to the formation of the poikiloblastic pyrite with relics of chalcocite in a microenvironment in which chalcocite predominates (Fig. 546). In contrast to xenoblastic and subidioblastic pyrite, blastic and replacement processes could perhaps be responsible for the patterns shown. Typical patterns of idioblastic pyrite growths, in the strict sense of the term blastic, are common. Fig. 547 shows idioblastic pyrite displacing chalcopyrite lamellae in intergrowth with galena. The bending of the chalcopyrite lamella is due to the crystalloblastic growth of the idioblastic pyrite. In addition, idioblastic pyrite is shown in a background of a bornite/chalcocite microenvironment in which the pyrites have grown (Fig. 548). Wedge-shaped pyrite crystalloblasts transecting bornite replaced by stromeyerite can be seen in Fig. 549. Fig. 550 shows an idiomorphic pyrite crystalloblast transecting a pattern of bornite replaced by stromeyerite. Galena is also enclosed in the pyrite crystalloblast. In contrast to xenomorphic (xenoblastic) pyrite replacing magnetite idioblastic pyrite crystalloblasts are shown, corroding and replacing included magnetite while maintaining magnetite relics in the intergranular (intercrystalline) spaces (Fig. 551). Comparing the composition of magnetite (Fe 3 0 4 ) with pyrite (FeS2) it is clear that dissolution of the magnetite must have taken place concurrently or prior to the crystalloblastesis of pyrite. In contrast to the xenomorphic crystalloblasts of pyrite (Figs. 543, 544) most impressive patterns of idioblastic pyrite replacing sphalerite are shown. Fig. 552 shows a band of sphalerite associated with silicates transected by a later idioblastic pyrite in which relics of the initial sphalerite band are maintained as relics in the pyrite. Pyrite idioblasts also show corroded and partly replaced sphalerite grains zonally included in the pyrite crystalloblast (Fig. 553). Fig. 554 again shows corroded and assimilated relics of sphalerite associated with silicates in the marginal area of the pyrite crystal-

loblast. However, sphalerite relics are maintained across a pyrite subidioblast (Fig. 555). The treatment of relic sphalerite in a pyrite idioblast can result in complicated patterns. The corroded sphalerite sometimes assumes symplectic forms in the later pyrite due to replacement of the sphalerite by the pyrite. Occasionally though, whereas sphalerite relics can remain in the pyrite, pigments-size sphalerite relics would be rather restricted in an outer zone of the pyrite (Fig. 556). Such a zonal distribution of the pigment remaining in the pyrite might be due to initial zonal incorporation of the sphalerite inclusions by the growing pyrite crystalloblast or they could be autocathartically pushed into the marginal zone of the pyrite. When the composition of the sphalerite (ZnS) is compared to pyrite (FeS2), removal of Zn and introduction of Fe with volume and lattice re-adjustments could theoretically account for the replacement of the sphalerite by the crystalloblastic growth (in sensu stricto) of the pyrite. As pointed out and as many textural patterns show, in many cases of crystalloblastesis complex replacement processes are dominant in ore petrography, and since replacement processes are inherent to metasomatism and most ore body concentrations are due to metasomatic processes, the relationship of crystalloblastesis and replacement is of great importance for the understanding of the patterns of ore minerals. In the present section only a few selected examples of blastic growths were presented. However, many of these replacement processes discussed in Chapter 5

could be considered to represent crystalloblastic growths since replacement is inherent to metasomatism. Kalliokoski (1965) described characteristics of metamorphosed ores including granoblastic textures and porphyroblastic growths of some sulfides from North America. Concluding the synoptical deliberation of crystalloblastesis and considering the significance of crystalloblastic pyrite both as the most common growth in ores and as a model for studying the process, some extracts will be presented from the contribution of Vokes and Craig (1991) on the subject which could also be most useful in interpreting similar structures elsewhere. According to them "the Gressli Caledonian stratabound pyrite-sphalerite-(chalcopyrite) ore exhibits microstructures indicating unusually high mobility and replacive activity of the base-metal sulfides and gangue at a late stage in its metamorphic history." Furthermore, Vokes and Craig support that "the dominant pyrite shows a granoblastic, often weakly directed fabric in which subhedral grains 0.5 to 2 mm in size, partly showing a polygonal foam texture, dominate over euhedral metablasts. Sphalerite and lesser chalcopyrite, with mainly quartz gangue, are present in varying proportions as matrix.... older, poikilitically enclosed sphalerite within the pyrite metablasts is partly replaced by the invading chalcopyrite. Both inclusionand matrix-sphalerite show, in places, abundant chalcopyrite and pyrrhotite "disease". All the sulfides are in places markedly replaced by later quartz."

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Chapter 9

Zonal Growths

The growth of a mineral from solutions depends on the supply of solutions by precipitation (deposition) of which a crystal growth develops, meaning that it starts from a cern (Keim) crystallization and layers are added. Thus, crystal growth is accomplished by adding additional zones from the supplies from outside the crystal. (In contrast, the growth of organic matter depends on the assimilation of substances internally by an organism.) Of course this is true for most organisms and microorganisms. However, the reversibility of viruses to the crystalline state makes things more complex in such cases. Thus, since the crystal growth (perhaps with the exception of certain blastic growths) depends on the added supplies of material for its development, any variation in the composition of the supplying material will be reflected in the composition of its zones of growth. In some cases however, the added zones might not differ significantly in composition. Nevertheless, due to the discontinuity in supply of material this might lead to a physical discontinuity and despite nondifference in composition, slightly different physical behaviour might be observed. As a corollary to the textural patterns showing crystal zonal growth, Frenzel and Ottemann (1967) reported zoned pyrite from Nukundamu, Fiji, with up to 10% Cu by weight in solid solution. Damley and Killingworth (1962), using a microanalyzer, traced variations in the Co content of pyrite from Chibuluma, Zambia. Comparable microanalytical studies by Aubert et al. (1964) show that gold is not uniformly distributed in the arsenopyrite from Chatelet (Creuse) but occurs in zones with up to 1.2% Au. According to them, these gold-rich zones do not correspond to any apparent optical or structural features in the arsenopyrite. Rather elaborate studies from Cannon et al. (1963) of a crystal of lead sulfide from Picher, Oklahoma, show significant differences in isotopic composition of lead in successive growth zones. As they report, lead isotope ratios in the parent ore fluid evidently changed with time during crystal growth. They also maintain that the growth history of this crystal could be used to interpret the genesis of the Mississippi Valley Pb-Zn deposits. Fig. 557 shows a zonal growth of galena which is expressed by the arrangement of the [100] cleavage 68

and its behaviour on polishing. Thus, the arrangement of the cleavage pattern of galena indicates a zonal growth pattern which might not reflect composition differences but is mainly due to the orientation of each successive zone of growth. Fig. 558 again shows the zonal pattern of galena as indicated by the cleavage (due to polishing). In contrast to zonal patterns depending on the orientation of the growth zones, in most cases the zonal growth actually depends on slight variations in composition of the supplied zones. Fig. 559 shows zonal growth of cobaltite. Considering that in the composition of cobaltite (CoAsS) small quantities of Fe can be incorporated, it is possible that zonal variations may rely on the fluctuation of the Fe content. Smaltite with the standard composition of ((Ni,Co)As3) may show composition variations in As between 2.9-3.0. This variation in As content may in turn result in a variation of the composition of the zonal growth which might not be obvious at normal polished surface. Fig. 560 shows smaltite with variations in the behaviour of the zonal growth on Anlaufen, the polished surface of the sample being exposed to air for more than 30 years. (Thus, a natural etching took place on the polished smaltite surface.) Fig. 560 shows finer and broader zones of growth and also a zig-zag and curved twin intergrowth "plane". Fig. 561 shows finer and broader zones of growth of the smaltite crystal. The patterns of zonal growth of smaltite may vary greatly and as Fig. 561 shows, the growth of the external zones of an individual crystal may be interrupted by the growth of another zonal individual. Considering the composition of the zones of smaltite in addition to the inadequacy in the Fe content (which might reach 5%) differences in the composition due to Ni and Co are often determined by microprobe analysis. Another cobalt mineral exhibiting impressive zonal growth and variations in the composition of zones is safflorite. The standard composition of safflorite is given as (CoAs2) with different quantities of Fe incorporated; also small quantities of Ni and Bi may be present. Thus, the variation in composition of the safflorite zones may be attributed to differences in Fe and also in the trace elements, Ni and Bi. However, differences in the main elements Co and As can also influence the zonal composition.

Figs. 562 and 563 show safflorite zoning (the polished surface was exposed to the air for about 30 years). The zonal patterns of safflorite may indicate change in form and habitus. The change in form and habitus as the zoned crystal develops is a common phenomenon and excellent examples of bravoite are shown by Ramdohr (1960). In contrast to the direct formation of safflorite by precipitation from solution, patterns of safflorite exhibiting complex zonal growths and, in cases, interbanded with gangue (zones) are often shown. Figs. 564 and 565 show zonal safflorite in cases interbanded with gangue; most probably the pattern is a derivative of the transformation of gels to crystallization (the cobalt mineral shows readiness to crystallize Kristallisationsfreudigkeit). In contrast to the normal zonal growth often exhibited by zircon, its metamictic derivative malacon shows alteration zonal structures in addition to the normal zonation which assume a colloform pattern and are secondarily imposed on the malacon due to water absorption in the transformation from zircon to malacon (metamictic changes, see Augustithis, 1964). Fig. 566 shows malacon exhibiting normal zoning due to growth and the colloform zonal pattern superimposed. In contrast, Fig. 567a. shows malacon with zonal growth while at the same time it exhibits a change in form (and habitus). In this connection it should be pointed out that Ramdohr presented excellent patterns of bravoite exhibiting zonal growth. The white central zones are poor in Ni and approach the composition of pyrite, whereas the outer zones show an increasing Ni content (see photomicrograph based on Ramdohr's contribution, Fig. 567b.). The zonal growth of the bravoite shows a change in form and habitus of the crystal during its development. In addition to bravoite, excellent patterns of tetrahedrite exhibiting change in form and habitus are shown in Fig. 568 with zoned central tetrahedrite (theoretical formula (Cu3SbS4)) and which may have excessive S ([Cu 8 SbS 7 ]). The variation in the content of S as well as the variation in Sb can greatly influence the zonal composition of tetrahedrite. In addition, differences in form and habitus are also exhibited (Fig. 568). In contrast to the differences in composition of the zones of a zoned mineral, a zonal pattern may be exhibited due to the successive or alternative precipitation of different minerals in which the one is interzonally associated with the main mineral. Figs. 569 and 570 show cuprite (Cu 2 0) with interzonally associated malachite. Fig. 571 shows cuprite margined by malachite enclosed in larger cuprite in which case a type of zoning is considered to have taken place. Cassiterite also exhibits impressive zonal growth (zonal structure) which might represent differences in the pore space (which in some cases is due to differences in the presence of "ex-solutions") and perhaps in other cases it is due to differences in the Fe content or

the presence of traces of Nb-Ta. In contradistinction, profound zonal structure differences might be due to differences in the orientation of zoned individuals and changes in form and habitus perhaps play a significant role. Fig. 572 shows cassiterite exhibiting zonal structure and change in form. Fig. 573 shows zoning in which case no change in form of habitus occurred. Fig. 574 shows alternating zones due to differences mainly in the orientation of the growth zones. Zoned cassiterite from the W-Sn deposit of Sokhret Allal (Zaer granite, Central Morocco) has been analyzed by electron microprobe by Giuliani (1987). According to him "the mineral zoning is related to variable Fe, Ti and Nb contents: the darker zones are enriched in Ti and Fe whereas the lighter zones are purer." Similar studies on wolframite by Nakashima et al. (1986) of a fresh euhedral crystal from the Kaneuchi Mine, Kyoto Prefecture, shows only slight compositional variation of less than 3 mole% MnW0 4 in the form of oscillatory zoning. Campbell and Petersen (1988) demonstrated that wolframite crystals from San Cristobal (Peru) show complex compositional zoning within individual crystals, e. g., one sample contained two crystals with contrasting zoning: one crystal had a high-Fe core and a high-Fe rim, whereas the other had a Fe-poor core and a high-Fe rim. It was suggested that these two crystals were formed at different moments and that the wolframite composition did not change monotonically with time. As mentioned, bravoite shows excellent zonal structures due to composition variation and changes in form and habitus are often reported. Schachner-Korn (1982) reports the zoning in bravoite is due to changing contents of Ni and Co in the ore-forming solutions and that sometimes it is possible to recognize habit changes developed during the growth of the crystal. As a corollary to the zonal growth of bravoite which is mainly due to the Ni and to a lesser degree to the Co content, is the zoning of pyrite. Microprobe studies determined the zonal distribution of a number of elements in pyrite, in addition to those already mentioned. In addition to Ni and Co, gold and platinum elements could be present in trace quantities in pyrite and any zonal inequality in the trace element distribution might indicate a difference in supplied material. However, an impressive difference in the form and habitus of the zonal structure of pyrite is often exhibited (as is the case with bravoite). Fig. 575 shows zonal structure of pyrite exhibiting changes in form and habitus. Another ore mineral with zoning is chromite and, as Fig. 6 shows, developed small crystals of olivine are zonally incorporated in a chromite crystal grain from Rodiani, Greece. Similarly, chromite crystal grains from South Africa show silicates (olivines) corroded and zonally incorporated within chromite (Figs. 606 and 607). In contradistinction to the chromite zoning exhibited on the basis of inclusions, Augustithis (1960) showed de69

coloration margins of chromite often with myrmekitic later serpentine intergrowths. Papunen and Idman (1982) reported that chromites have two kinds of zoning, called the iron-trend and the aluminum-trend, correspondingly. They stated "the iron-trend of zoning was formed in deuteric reactions of early chromites or chromite crystals were mantled by secondary magnetite in serpentinization". As already mentioned, safflorite as a derivative of gels showed impressive zonal patterns and simultaneously exhibits differences in form and habitus. An ad-

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ditional case of derivation from gels (transformation of colloid to crystalline) is found in Fig. 576, a zonal structure of metacinnabar is shown derived from gels. Such zonal patterns of elongated crystallizations from gel sphaeroids indicate that in addition to variations in the composition of initial gels, possible differences in the composition of the crystalloid phase may also take place in the transformation phase from gel to crystalline (that is, zonal growth might take place during the development of the elongated crystals).

Chapter 10

Epitaxis - Epitactic Growths

Concerning the phenomenon of epitaxis it is necessary to examine the limitations of the concept in respect to overgrowths on one hand and zonal growths on the other. It may be important to consider the relation of epitaxis with symplectite intergrowths: a mineral which is epitactic on another mineral might extend into it as symplectic intergrowth or, independent to the epitactic growth, the mineral might exist as an inclusion or, in cases, as symplectites resembling "ex-solutions". In ore microscopy epitactic growths are common and a number of cases are recorded. In the present effort some selected cases will be presented with the aim to show typical examples as well as to speculate on the relationship of epitaxis, overgrowths and zonal growths. Fig. 577 shows chromite with an epitactic growth of magnetite. Here, the magnetite is a margin on the crystal face of the chromite. In contradistinction, chromite often corroded or rounded might have an "overgrowth" of magnetite which may attain idiomorphic shape (see overgrowth of blastic magnetite on a corroded or rounded chromite nucleus, Kozani, Greece; Fig. 578). In contrast, epitactic magnetite may occur as a growth on a "corroded" chromite outline (Fig. 579). In some cases, chromite shows multiple epitaxis, that is, the chromite is surrounded by magnetite which in turn is surrounded by a margin of sulfide (pyrite?). It should also be noted that an overgrowth of a sulfide took place on the magnetite grain. It is now the question whether consecutive epitactic growths of magnetite occurred on chromite and sulfide (pyrite) on magnetite. It should be emphasized though that both the chromite and the magnetite have a spinel lattice and epitactic growths are understandable, in which case the surrounding pyrite on the magnetite would simply be an overgrowth. In addition to the common case of epitactic magnetite on chromite, Strunz (1961) presented very impressive cases of uraninite epitactic on columbite (Fig. 580). An attempt will be made to further examine the relationship of uraninite and columbite of the same occurrence (Hagendorf pegmatite). Fig. 581 shows uraninite growing on the columbite with the uraninite not having any extensions into the columbite. Fig. 582 shows uraninite grown on the columbite and well-developed faces of uraninite are exhibited. In contrast, in Fig. 583

uraninite is partly engulfed by the columbite. It should be noted that in this case, too, the uraninite exhibits well-developed crystal faces as it extends out of the columbite. Comparing the patterns in Figs. 583 and 584, the latter shows uraninite partly engulfed by columbite, which, however, in contrast to the uraninite in Fig. 583 is idiomorphic. Fig. 585 shows a growth of uraninite on columbite with the crystal faces developed on the outer side of the uraninite only. It should be pointed out though, that in the columbite an uraninite inclusion exists and developed faces of the epitactic uraninite and those of the inclusion (also with developed crystal faces) are similarly oriented (compare the inclusion with the epitactic growth, Fig. 585). The case of uraninite epitactic on columbite should be further discussed in order to determine whether epitaxis bears any relation to replacement. In contrast to the uraninite representing an overgrowth or the idiomorphic uraninite partly engulfed by columbite (Figs. 582 and 584, respectively), Fig. 586 shows epitactic uraninite with a root extension into the columbite. In this connection, it should be stressed that as early as 1948 Drescher-Kaden already showed that epitactic growths of quartz on K-feldspar may have root-like extensions into the feldspar (see diagrams from DrescherKaden (1948) "Die Feldspat-Quarz-Reaktionsgefüge der Granite and Gneise und ihre genetische Bedeutung", Figs. 89 and 105). Drescher-Kaden explained the pattern as the result of the replacement of the feldspar by the graphic quartz and considered the quartz epitaxis as due to replacement and overgrowth of the quartz on the feldspar. As a corollary to the later growth of the uraninite on the columbite are the patterns which show epitactic growth of uraninite having root-like extensions of uraninite replacing the columbite and, as extensively discussed in the section about symplectites, the replacement textures of uraninite whereby columbite is replaced by the later uraninite assuming "ex-solution" forms. Other cases of epitactic growths include magnetite epitactically overgrown by ilmenite. Fig. 587 shows magnetite grains epitactically overgrown by ilmenite, in cases, the ilmenite forms fine margins of the magnetite. 71

In contradistinction to the ilmenite margins epitactic on magnetite, magnetite can exhibit spinel (hercynite) epitactic on the magnetite and simultaneously spinel intergrowths are found in the magnetite, in cases marginal to oriented ilmenite. Further it should be noted that the epitactic spinel has extensions into the magnetite. Often the epitactic spinel shows intergrowth with chalcopyrite. It needs to be mentioned though, that the epitactic spinel in cases shows "prismatic " forms (Figs. 588 and 589). Additional intergrowth patters show titanomagnetite with spinel oriented bodies surrounded by an ilmenite margin. The contact of the magnetite with the ilmenite is symplectic and is followed by myrmekitic symplectic spinel. It should be pointed out though, that marginal to the ilmenite (which occurs as a margin to the magnetite) is epitactic spinel, occasionally surrounded by chalcopyrite or in intergrowth with it (Fig. 590). Possibly the most impressive epitactic growths exhibited are the symplectic-epitactic growths of jacobsite on braunite (Figs. 591, 592). Furthermore, symplectic epitactic jacobsite in intergrowth and epitactic on the braunite is shown in Figs. 593 and 594. However, cases exist where the epitactic overgrowth has corroded and surrounded the mineral on which it shows epitaxis. Fig. 595 shows sphalerite partly enclosed and corroded by pyrrhotite which also exhibits a tendency to attaining crystalline (developed crystal) faces. Similarly pyrrhotite forms epitactic margins on sphalerite (Fig. 596). Arsenopyrite is also marginally corroding sphalerite and simultaneously exhibits well-formed crystalline outlines (i. e., externally the arsenopyrite develops crystal faces), see Figs. 597 and 598. Pyrite in cases surrounds magnetite with which it is also symplectically intergrown and simultaneously exhibits well-developed contours (Fig. 599). In addition, cases are exhibited of pyrite epitactic on sphalerite where along with the corrosion of the sphalerite, the pyrite shows well-formed crystalline outlines (Fig. 600). In contrast to the examples described so far, where the epitactic mineral has overgrown on another ore mineral, cases are presented where epitactic bournonite has overgrown silicates (?quartz) and tetrahedrite (Fig. 601). Rare and complex epitactic patterns are also found with bornite overgrown on rutile (which in cases is symplectic with silicates). In addition to the wellformed bornite (exhibiting a tendency to develop crystal faces), bornite is also in symplectic intergrowth with rutile (Fig. 602). In the following, certain cases of overgrowths will be mentioned that can hardly be considered to represent epitaxis. Figs. 603 and 604 show pyrrhotite with a margin of albandite which simulates epitaxis since albandite also occurs as inclusions in the pyrrhotite. It should be mentioned though, that the crystal faces or crystalline 72

outlines of albandite are not well-developed, and the pattern is rather a case of an overgrowth. Similarly pyrite might form an overgrowth on gersdorffite and as Fig. 605 shows, pyrite extends beyond the gersdorffite and colloform patterns of pyrite are also exhibited. This case does not represent an epitactic growth of pyrite on gersdorffite either. In contrast to these overgrowths, magnetite is surrounded by a margin of malacon (epitactic on the magnetite) and with a well-developed zoned malacon epitactic on the marginal malacon. Rutile is also epitactic on the malacon (see Fig. 619 and Chapter 11). The relationship of epitactic growth and overgrowths, and especially the phenomena of inclusions and overgrowths (zoning), are further considered in Chapters 9 and 11. Considering the relationships of the minerals that are epitactically overgrown on the host (on the mineral on which the epitaxis took place) from the above mentioned cases, only the case of the epitaxis of magnetite on chromite shows crystal-lattice relationships. Chromite crystallizes in the cubic system and in the 4/m 3 2/m class and belongs to the space group Fd3m and Ζ = 4. Furthermore, the lattice constant of synthetic chromite is a = 8.344. In comparison, magnetite crystallizes in the cubic system, in the class 4/m 3 2/m and belongs to the space group Fd3m and Ζ = 8. Its lattice constant is a = 8.374. Comparing the crystal data of chromite and magnetite the only conspicuous difference is that the lattice constant is slightly larger in the case of magnetite. The comparison clearly supports the epitactic relationship of magnetite and chromite. However, it seems that the crystal system, class, space group and lattice constants are not a prerequisite for the epitactic relationship of minerals. As mentioned, Strunz presented an impressive case of uraninite epitactic on columbite. A comparison of the crystal data of uraninite and columbite does not show similarities or approximate values. To be more specific, uraninite crystallizes in the cubic system and in the 4/m 3 2/m class. It belongs to the Fm3m space group and its lattice constant is a = 5.368-5.555. In contrast, columbite crystallizes in the orthorhombic system and in the 2/m 2/m 2/m class. It belongs to the space group Pcan. However, both uraninite and columbite have the same Ζ values (4). Comparing the crystal data of uraninite and columbite shows that there is no close relationship but uraninite is undoubtedly epitactically grown on the columbite (Fig. 580; and also the textural patterns that exhibit the epitactic growth of the uraninite on the columbite). Furthermore, the comparison of the crystal data of ilmenite-magnetite, jacobsite-braunite, sphalerite-pyrrhotite, arsenopyrite-sphalerite, pyrite-magnetite, bournonite-tetrahedrite and bornite-rutile do not show close relationship. These textural patterns with epitactic characteristics support the concept that their rela-

tionship is based more on the possibility that the epitactic mineral replaces and overgrows the host mineral. In such cases the prerequisite of related crystal lattice

and the other crystal data (crystal class, space group and lattice constants) are not essential for the epitactic relationships of minerals.

73

Chapter 11

Inclusions

In the section about epitaxis, it was mentioned that the relationship of epitactic growths and inclusions should be further considered from the point of view of inclusions. Also of importance in understanding epitactic growths is the relationship of zoning and epitactic growth. Epitactic growths should also be discussed from the point of view of overgrowths and thus, the relationship of overgrowths and epitactic growths needs to be examined as well. In the effort to comprehend the relationship of the included and the host mineral an attempt will be made to represent just a few examples indicating that in some cases genetic or geochemical interrelationships exist between the included mineral and its host. Fig. 606 shows a chromite grain centrally enclosing a partly corroded olivine and at the same time zonally enclosing olivines, reduced in size and indicating more advanced corrosion. Considering that olivine ((Mg,Fe) 2 Si0 4 ) is a mineral rich in Mg and Fe, the geochemical interrelationship of Mg, Fe, Ni and Cr should not be disregarded. However, more important is the crystallization sequence in mantle rocks (or their derivatives). Olivine and chromite are crystallizations of melts and depending on the composition of the melt either chromite crystallizes first, followed by Mg-Fe silicates, or, the silicates are found to crystallize first and then the chromite. As Fig. 606 shows, the silicates crystallized first (occurring included in the core of the chromite) followed by a chromite layer (zone), and subsequently by silicate crystallization, interzonally between successive zones of chromite crystallization. Fig. 607 shows silicates zonally incorporated in a chromite crystal (subidiomorphic) and furthermore in contrast to the patterns shown in Fig. 606, Fig. 608 shows chromite including silicates, in cases marginally engulfing them, suggesting a later crystallization for the outer chromite (engulfing the silicate). Cases are exhibited where chromite with "ex-solution" of spinel and rutile include rounded rutile grains (Fig. 609). Considering that there is an antipathy of Cr to Ti (despite their interrelationship in accordance to the horizontal interrelationship in the periodic table). In this particular case, the fact that rutile lamellae occur oriented in the [111] direction of the chromite, suggests that there is a geochemical interrelationship of the included mineral (rutile) and the host chromite. 74

The interrelationship is not restricted to hosts derived from magmatic crystallization (as shown by the mantle derived chromite-olivine and rutile-chromite examples already discussed) but also to nugget formation developed at considerably lower temperature. Fig. 610 shows chromite included in a Pt-Fe nugget. Osmiridium crystals are also zonally incorporated. In addition, complex in composition PGMs are included in the peripheral part of a Pt-Fe nugget grown in laterite (see Chapter 13 and Fig. 611). In both cases, the chromite grain and the PGMs are geochemically related to ferroplatin. A rather complex case indicating the relationship of uraninite (U0 2 ) included in columbite, consisting of a mixed crystal of niobite [(Fe,Mn)Nb 2 0 6 ] and tantalite [(Fe,Mn)Ta206] is presented in the series of patterns shown in Figs. 612-614. Fig. 612 shows uraninite exhibiting some developed crystal faces included in columbite. In contradistinction, Fig. 613 shows an uraninite crystal with developed crystal faces included in columbite. However, in both cases there appears to be an intergrowth of uraninite and columbite. Fig. 614 shows uraninite included in columbite, exhibiting well-developed crystal faces (in addition to corroded outlines) and also epitactic uraninite on the columbite is shown. It is very interesting that the crystal faces of the included uraninite and the ones of the epitactic uraninite have the same orientation (see also Chapter 10), suggesting that the included uraninite was also epitactically related to the columbite while the latter was zonally developed. Uraninite included and epitactic on the columbite as well as the fact that the developed crystal faces of the included and the epitactic uraninite are similarly oriented, show that there is a relationship between inclusion and epitaxis in this case. Considering also that U is geochemically interrelated with Nb and Ta, in addition to the crystallographic interrelationship of epitactic growth, and the host it seems that there is also a geochemical interrelationship between the columbite and the included uraninite. The fact that uraninite occurs as an inclusion in the columbite and simultaneously as an epitactic growth on the same host would suggest that inclusion and epitaxis could be related, that is, that the epitactic uraninite is an overgrowth on the columbite (the relationship might be of successive generations of

crystallizations; uraninite inclusion, columbite, epitactic uraninite). Figs. 615 and 616 show hawleyite (a cadmium mineral) included in wurzite/sphalerite (see Augustithis and Vgenopoulos, 1980). It should be pointed out that also in this case there is a geochemical interrelationship between the host sphalerite/wurzite and the included Cd mineral hawleyite. Cd and Zn are related as subgroup elements of the II family in the periodic system. Further examples supporting that included minerals and their hosts may (in some case only) show geochemical interrelationships is provided by the case of ilmenite (TiFe0 3 ) with haematite oriented bodies where ilmenite is surrounded (included) by laterformed titanite (CaTi(0Si0 4 )). Also in this case the ilmenite and the surrounding titanite are geochemically related as indicated by the abundance of Ti in both. A rather interesting case indicating geochemical relationship between host and included mineral is shown in Fig. 617. Well-developed zoned crystals of cobaltite (CoAsS) are surrounded by safflorite (CoAs2). The safflorite might represent a later differentiation of the same solutions which first formed cobaltite. Similarly, a fragment (or resorbed) zonal smaltite ((Ni,Co)As3) can be surrounded by later formed safflorite and in this case there is also a strong geochemical interrelationship between the included and the later formed mineral which surrounds the smaltite. Another example of the interrelationship of host and included mineral or central mineral growth and marginal phase is shown in Fig. 577. Chromite indicating a crystalline outline is epitactically surrounded by later formed magnetite. In contrast, marginal magnetite surrounding a chromite xenocryst is formed when a chromite crystal grain is found in a basalt flow as xenocryst (see Augustithis, 1965). In opposition, corroded and resorbed chromite may be surrounded by later formed blastic magnetite in slightly metamorphosed pisoliticoolitic lateritic ores which are actually derivatives of ultrabasics. As discussed in the case of epitactic growth of magnetite on chromite (see Chapter 10 and Fig. 578), there is a crystallographical and geochemical interrelationship between chromite and magnetite. A rather rare case is colloform pitchblende (U0 2 ) included in coffinite (USi0 4 ) as shown in Fig. 618. It should be pointed out here that the colloform pitchblende is interbanded with coffinite indicating the textural interrelationship of these two minerals in this case. It should also be mentioned that whereas the pitchblende is amorphous (despite included uraninite crystallites), it is surrounded by tetragonal well-crystallized coffinite. Thus, whereas there is hardly any crystallographical interrelationship between pitchblende and coffinite, both are uranium minerals and geochemically related. Uraninite (pitchblende) is included in zonal malacon (metamictic zircon) exhibiting yet another case where the included mineral and the later formed host are geo-

chemically interrelated because of the geochemical relationship of Zr and U. It should also be emphasized that the zircon may contain radioactive elements as well. In contrast, cases may be found where malacon was most probably tectonically pushed into magnetite (see Augustithis, 1964). In this case the relationship of the included mineral to the host may just be the effect of mechanical action (tectonic effect). However, cases are presented from the same occurrence where magnetite is epitactically surrounded by malacons (Figs. 619, 620). In contrast to the tectonically mobilized malacon in magnetite, cases are shown where crystalline magnetite is included in zonal malacon which in tum sends extensions into the outer magnetite in which it is included (Fig. 621). The case of malacon in uraninite or magnetite indeed is problematic. Augustithis (1964) put forward alternative interpretations to explain the malacon included in the uraninite (Fig. 622) and the cracks related to the malacon: "Selected cases of uraninite show its structural relationships with other metallic and nonmetallic minerals in the paragenesis. Microscopic observations (from the Faraday Mine) show the structural interrelationship between uraninite and zircon (malacon) and magnetite. In the particular case discussed below, the structural interrelationship of zircon and uraninite is difficult to interpret and several alternative suggestions exist. Fig. 622 shows malacon built up of zones in uraninite. Cracks can be seen radiating from the malacon into the uraninite and these are often referred to as radioactive cracks (radioactive Sprengungen). The following considerations regarding the uraninite-malacon interrelationship and these cracks are advanced: A) Fig. 622 shows cracks produced by zircon radiating into a substance that is more radioactive, namely uraninite. B) It is difficult to imagine that the emanation of rays would result in the production of such cracks. C) Fig. 622 shows that the radioactive cracks tend to follow certain directions in the uraninite and do not radiate with equal intensity at all points around the malacon. In order to explain the difficulties inherent to the given explanation of these cracks, the following alternative suggestions are put forward: (i) Due to the metamictic alteration of zircon to malacon, the zircon incorporated water and consequently there was an increase in volume which subsequently acted as a mechanical force producing cracks. (ii) The zircons (malacons) possibly represent later growths and due to the force of the crystallization they formed cracks in the surrounding minerals." From the above mentioned examples it follows that neither epitactic growths nor inclusions have a geochemical and mineralogical interrelationship as a prerequisite in the sense that a mineral may be surrounded by a younger mineral independent of their composi75

tions. This is true to a lesser degree for epitactic growths (see Chapter 10). Additional cases of included minerals with only a remote geochemical relationship between the included mineral and the host are illustrated in Fig. 623 where gold is enclosed in uraninite (Augustithis, 1964, emphasized the paragenetic relationship of gold and U). However, the case of an association between gold and uraninite is better understood as a paragenetic interrelationship. Yet it should not be underestimated since radioactivity has often been used for the exploration of gold. Another case where the same host mineral, in this case pyrrhotite, may include or surround two kinds of minerals (pentlandite (FeNiS) and silicate), indicates

76

that whereas pyrrhotite is strongly paragenetically and geochemically interrelated with pentlandite, the pyrrhotite host is probably not strongly related to the silicate inclusion even if the latter contains some Fe and Mg. Philpotts (1961) reports that pyrrhotite, pentlandite and chalcopyrite have replaced pyroxene and olivine but their alteration products, serpentine and amphibolite, were unfavourable for replacement. These cases where the included mineral is geochemically or paragenetically interrelated to the host mineral (the mineral surrounding or including the inclusion), do not satisfy the establishment of laws or rules. They should only be seen as tendencies in the chaotic relationship of included minerals and their hosts (which may be any later formed minerals).

Chapter 12

Colloform Structures (Gel Structures)

(a) General The significance of the colloidal state of matter is stated expressively by Wolfgang Ostwald (1935) as follows (translated from the original): We call the discontinuities between macroscopic and atomic structure as colloidal discontinuities, the dimensions between wavelength and atomic diameter as the colloidal dimensions, and the state of this intermediate discontinuity as the colloidal condition of the matter. The colloidal science is the study of the physical-chemical peculiarities of this state of the intermediate discontinuity. It follows from the so-called "dimension-definition" of the colloidal science that this state can be basically attained by any body, so that the colloidal condition is a general state of matter. The quotation above is a justification of the great abundance in geological formations of forms and structures derived from the colloidal state of matter. Furthermore, it is not surprising that an extensive list of mineral substances are recorded to have been derived from the colloidal state and thus showing colloform structures. Only selected cases of colloform structures will be presented that will serve to establish colloform patterns as patterns only (in the sense that similar patterns exist in a great variety of mineral substances which will not be discussed here). Considering the general concept of colloidal chemistry, among numerous profound questions that could be raised, is undoubtedly the fundamental question of form: why do gel structures (colloform structures) tend to attain sphaerical form or forms that tend to approach sphaerical shape? It should be emphasized that this question and its answer are inherent to the fundamentals of the colloidal state. Ostwald described three ways in which transition can be achieved from macroscopic discontinuity to atomic structure. The colloidal dimensions are considered as "intermediate" between macroscopic discontinuity and atomic structure. As stated by Hardy (see Ostwald, 1935) this transition through the colloidal discontinuity is understandable by considering the matter in the boundary state (Grenzschicht Zustand). (The dimensions of the boundary state are within the conditions and dimensions of the colloidal state.)

Other ways in which the transition from macroscopic discontinuity to atomic structure can be achieved is by difformation or by dispersion. Ostwald recognized a one-dimensional lamellar dispersive system, a two-dimensional fibrillar dispersive system and a three-dimensional corpuscular dispersive system. Particularly by understanding difformation and dispersion processes an attempt can be made to understand sphaeroidal structures (products of coagulation of colloids). Furthermore, in discussing difformation Ostwald made reference to the sphere: by a given volume of a sphaerical body has the smallest specific surface (surface/volume) namely in case of volume 1 cm3 = 4.836. Any deformation of the sphere will result in larger values. An equally large cube has the specific surface of 6.00, a tetrahedron 7.20. Considering difformation in achieving colloidal dimensions, Ostwald states: by shaping one cm3 into a film with average colloidal dimensions (ΙΟπιμ), a circular disc is produced with 564.2 cm radius and a surface of 200 m2 (this example is a case of one-dimensional difformation). In the case of two-dimensional difformation, by drawing a cm3 into a fiber of 10 πιμ diameter, a fiber length of 12.7 million km is obtained with a surface of 400 m2. It is understood that film and fiber can lead from the macroscopic discontinuity to atomic structure. However, the results in the case of dispersion are more impressive. By choosing average colloidal dimensions (10 πιμ) a cm3 can result in a surface of 200 m2 total with lamellar colloidal splitting (Zerteilung), 400 m2 total with fibrous splitting and 600 m2 with corpuscular splitting. The case of the molecular disperse system is more impressive: the splitting of a cm3 into particles (particle diameter 0.1 πιμ = 1 Ä) will result in a surface of 60,000 m 2 total. Considering the examples quoted above, the colloidal state is a transitional phase of the macroscopic discontinuity to the atomic structure. Both difformation and dispersion can result in the colloidal state. Colloidal solutions in the various mineral-forming processes are significant, as is particularly evident from the colloform structures exhibited by mineral formations, the colloidal state has not been an "aberrant" condition in the rock-forming processes but a state of great importance. 77

As discussed by Ostwald (1935), the colloidal condition is a state of matter. "It follows from the so-called "dimension-definition" of the colloidal science that this state can be basically attained by any body, so that the colloidal condition is a general state of matter." The transition from macroscopic discontinuity to atomic structure over the colloidal state and the reverse process, namely the combining of atoms to form macroscopic discontinuity, is achieved either by building "macro molecules" (Staudinger, 1920), the polymerization of which could lead to crystallization, or by the formation of colloidal particles (agglutinations of colloidal dimensions), the coagulation of which could result in colloform structures of gel structures. Both difformation and dispersion means the manifold increase of specific surface. In the reverse process, the formation of the macro-structural discontinuity from colloidal particles (e. g., 1 Ä dimension) will mean the maximum possible reduction of specific surface which is achieved by the formation of sphaeroidal forms, or approximately sphaeroidal. Thus, it can be concluded that the maximum possible difformation and dispersion means the maximum possible increase of specific surface of the total colloidal particles. The reverse process will mean the maximum possible reduction of specific surface which is achieved by the formation of a sphaerical body. Considering the reverse process to difformation-dispersion, namely the coagulation of the dispersive system to form a gel or a colloform structure, it is necessary to discuss coagulation and in general the mechanism of precipitation of the colloidal systems. Ultramicroscopic observations indicate that colloid particles (Kolloidteilchen) exist in a continuous irregular (Brownian) movement. An electrical charge (Ladung) prevents the colloid particles (which, owing to the Brownian movement, approach each other) from being united to form aggregates. If the electrical charge ceases to operate, the particles will tend to possess reduced surface area by uniting, so that they will form coarsely dispersed aggregates. The continuation of this tendency to possess reduced surface area results in the formation of gel sphaeroids, i. e., bodies with maximum reduced specific surface. Thus, the understanding of sphaeroidal gel-structures is possible by considering the reverse process of difformation and dispersion. Since colloform structures are particularly important in ore mineralogy, an attempt will be made to discuss the main causes which bring about coagulation of gels. (i) Addition of a differently charged electrolyte. In accordance with the theory of colloidal conditions the colloid particles, being suspended in liquid or gas, are electrically charged (Wandern im elektrischen Feld). It is considered that metal oxides and hydro-oxides are usually positively charged. In accordance with the theory of coagulation, the negatively and positively charged colloidal particles require cations and anions of the added electrolyte, respectively, for coagulation 78

to occur. Consequently, one of the causes for gel-formation is the addition of a negatively charged electrolyte. In metallic mineral paragenesis (e. g., pitchblende paragenesis) such conditions can occur in nature by adding other sols to a sol, the particles of which are differently charges, so that coagulation can occur by mixing colloid solutions. (ii) Temperature changes of colloid solutions. As it is the case with crystallization of hydrothermal mineral processes, cooling is also of significance to the coagulation of colloid solutions, especially in hydrothermal veinform deposits. Nevertheless, temperature changes in general influence coagulation conditions in colloidal solutions. Ore microscopic observations indicate the importance of gel (colloform) structures in hydrothermal paragenesis. These structures are regarded as being due to colloidal precipitation, and although it is stated in literature that they represent relatively low temperature conditions of formation, they often form the oldest minerals present. In this connection, it should be mentioned that colloform structures are observed in cements in which they are regarded as high temperature formations (Ottemann, pers. comm.). Studies by Augustithis on volcanic rocks showed that colloform structures or colloform-like structures can be formed at the final phase of the consolidation of volcanic melts. Augustithis (1978) described gelhyaloid structures in basalts and other feldspars exhibiting colloform-like patterns that are considered to be formed from colloids. In order to understand colloform structures in hydrothermal mineral parageneses and to put the genesis of gel structures into perspective, a discussion of colloidal solutions and of structures formed after their coagulation is essential. An attempt to explain the theoretical aspects of the problem is tentatively put forward. The recognition of crystalloid and colloidal substances by Graham (1860) led to special examination of crystalloid and colloidal conditions in general. Nowadays it is supposed that under suitable conditions all substances are able to form colloid solutions - there is no limitation to crystalloid and colloid substances but to crystalloid and colloid conditions in general. Furthermore, the size of the substances is of significance for the understanding of colloidal conditions, since dissolved substances within the limits of 1-2 μ are in the solution state. With the increase in size of the particles to 100-500 μ, there is a transition of the solution to suspension. In general, in dispersed systems the following grades are recognized: coarse dispersed, colloidal dispersed and molecular dispersed. Thus, the formation of colloidal solutions for the coarsely dispersed systems is effected by dispersion, whereas for the molecularly dispersed substances existing in true solution, it is effected by condensation. The aforementioned theoretical concept can help to explain many of the colloform patterns in ore petrog-

raphy. The colloidal conditions from which many gel minerals are formed is variable, however, gel pitchblende and many other gel minerals are believed to originate from condensation rather than from dispersion. Such systems have the property of reversibility of gel to sol. For a gel structure which changed into a crystalline structure rearrangement of the atoms must be assumed. This can take place rather in the state of a sol than in that of a gel. Therefore, the property of reversibility of gel to sol is important for the transformation of gel to crystalloid. Consequently, the transformation or partial transformation of gel to sol will result in the formation of true solutions, in which rearrangement of the atoms can take place and a crystalline structure can be produced. The transformation of gel to crystalloid state is also important for the understanding of the initial gel patterns transformed to the crystalline state. The transformation of colloform to crystalline structures depends on the reversibility of gel to sol and, in turn, better characterizes the gels derived from condensation rather than from dispersion. In contrast to this type of transformation and apart from the size of the substance in solution, the capacity and readiness of a substance to crystallize is a phenomenon which Ramdohr refers to as Kristallisationsfreudigkeit (readiness to crystallize). In this connection it should be mentioned that whereas most minerals can exhibit both crystalline and colloform structures certain minerals almost always tend to crystallize due to their Kristallisationsfreudigkeit. PbS is an example for such a substance which exhibits colloform patterns only very rarely. Furthermore, it should be mentioned that ore microscopic observations indicate that both crystalloid and colloid conditions may exist for the same substance. Fig. 666 shows crystals and gel pitchblende in direct vicinity to each other. For their coexistence two alternative explanations are suggested: (i) The "crystalloid" (crystalline) phase was formed first and subsequently the "colloid", so that the crystalline and colloidal solutions were of different age. (ii) The crystalline and gel pitchblende, or other mineral substances, were formed simultaneously. It can be argued, however, that it is not a question of age difference of the crystalline and the colloid phase (since the interrelationship of these structures suggests an approximately simultaneous formation), but of the size of the dissolved particles as the cause for their coexistence. As the studies on replacement structures, ex-solutions versus replacement, symplectites, etc., (see previous Chapters) have shown, replacement in which diffusion processes were of fundamental importance for the patterns have been discussed under the relevant headings. One fundamental difference between these replacement-diffusion and the colloform patterns and structures is that the colloid conditions are characterized, in contrast to crystalloid conditions, by a smaller

capacity to diffusion (geringeres Diffusionsvermögen) which is due to the size of the particles (Teilchengröße) of the dissolved substances under colloid conditions. Thus, there is a fundamental difference between replacement-diffusion patterns and gel (colloform) structures. Based on the general principles of colloidal physicochemistry as outlined in the general and introductory part of this chapter on colloform structures, it is necessary to present some of the most characteristic gel derived structures with emphasis on colloform structures encountered in ore metallography. From the great plethora of gel derived structures such as kidney forms, botryoidal structures, globular and mammilary structures, only one example of a mammilary-botryoidal structure is presented in this section (Fig. 624). However, emphasis is given to the microstructural derivatives of colloid solutions and a wide spectrum of colloform patterns of pitchblende will be used as a colloform derived mineral which provides an example for a variety of textures and processes. Fig. 625 shows pitchblende as pore space filling of sandstone. The pitchblende microsphaeroid derivatives of colloid solutions occupy pore spaces between the sand grains. It is evident that the coagulation of the colloidal solutions took place under low temperature conditions in the sandstone. In addition to the size variation of gel derived structures and microstructures as demonstrated by comparing Figs. 624 and 625, pitchblende furnishes examples of composition variation of the different shell of colloform structures. Fig. 626 shows gel pitchblende exhibiting variation in colour of the shells building up the colloform structure. The light coloured bands of the gel pitchblende indicate variation of composition which are due to differences in the colloidal solutions which coagulated to form the gel structure. In contrast, the darker marginal zone is due to oxidation and is subsequent to the coagulation processes responsible for the formation of the gel structure. Also Figs. 627-629 show the variation in colour of the different shells (bands) building up the colloform pitchblende structures. The variation in colour of the different bands is the result of differences in the composition of the colloidal solutions. In addition, also differences in the very microstructure of the shells building up the gel pitchblende sphaeroids is shown in Fig. 629. It shows broad bands and intricate, complex-in-structure fine bands. In cases pitchblende may show variation in composition exhibited by differences in colour due to different generations of coagulated material. As Fig. 630 shows, a first generation banded colloform pitchblende (darker in colour) has been fractured and surrounded by a later pitchblende generation forming a thin margin on the first generation fragments. Fig. 631 shows a fragment 79

of generation 1 partly surrounded by a different in colour banded generation 2 pitchblende formation. In contrast to the usual patterns of pitchblende sphaeroidal structures where a marginal darker alteration zone can be formed by oxidation (Fig. 626), cases are exhibited of gel pitchblende sphaeroids where the central part is altered (darker due to oxidation of the uraninite). As a corollary to the possibility that oxidation of central gel layers can take place, Fig. 632 shows gel pitchblende structures indicating dissolution of the central part and of a layer whereas the external gel layer remains unaffected.

(b) Relic Gel Structures As mentioned, the transition of colloform to crystalline structures is achieved mainly by the reversibility of gel to sol and possibly from sol to a solution phase out of which crystallization can take place. The reversibility from gel to sol is greater in the case of gels formed by condensation than of gels from dispersion. It is believed that a number of gels such as pitchblende, sphalerite, cassiterite, and others, are formed by condensation and therefore, they are reversible to sol, and that crystallization can take place from the state of solution. Considering further that the transformation from colloform to crystalline can occur by preserving relic gel structures, the formation from gel to sol and the formation of a solution phase may take place in the form of an advancing front. As a result, the transformation of colloform to crystalline can be achieved while the initial gel banding is preserved as a relic structure. Fig. 633 shows initial colloform pitchblende transformed to crystalline uraninite where the initial gel banding is preserved as pore space banding in the crystalline uraninite. In addition to the gel structure that can be preserved in the crystalline state, Fig. 634 shows a structure which is believed to be a relic from an initial colloform pitchblende. Petrographic studies on the preservation of relic gel structures in quartz and other silicate minerals which were originally considered gels (quartz was gel chalcedony) were presented by Levicki (1955), Augustithis (1973), Drescher-Kaden (1969). Fig. 635 shows uraninite with a preserved structure which is believed to be a relic from an initial colloform pitchblende. It should be mentioned that although the structures presented in Figs. 634 and 635 are rather rare, nevertheless they are most probably relics from a gel phase. Common gel relic structure is also shown by manganite (MnO(OH)), see Figs. 636 and 637. It is assumed that manganite gels are formed by condensation and that they are easily reversible to sol. For the sol a solution phase easily permits crystallization of manganite, again due to its readiness to crystallize. 80

As mentioned, depending on the reversibility of gel to sol, a solution phase may be formed out of which crystallization of the substance exhibiting a colloform texture can take place. The transformation of gel to crystalline structures is believed by some researchers to be the transformation from an amorphous state to an ordered crystalline state. However, the hypothesis that the disordered state is regarded a stable form should not be ignored. The transformation from colloform structure to crystalline often results in the preservation of gel relic structures in the crystalline phase. Considering gel pitchblende to be the result of condensation rather than dispersion it is most likely that gels formed by condensation are more reversible to sol than those formed by dispersion. It is thus possible that gel pitchblende is reversible to sol and a solution phase or front of solutions might result out of which a crystalline state can be formed. It is thus possible that gel pitchblende is transformed to crystalline uraninite. As the transformation is achieved, relics of the initial gel phase may be preserved in the crystalline phase. Fig. 633 shows crystalline uraninite in which fine initial gel banding is preserved. Gel relic patterns may also be preserved in niccolite (NiAs) indicating again that a transformation of the gel niccolite phase to the crystalline niccolite took place. Niccolite in addition to pitchblende may have a gel phase formed by condensation. In this case the gels are easily reversible to sol and thus a solution phase can be formed from which crystalline niccolite is produced. The undeniable presence of gel relics in the niccolite (Fig. 638) supports the transformation of gel to crystalline niccolite, which is most common and most stable. Sphalerite (ZnS) gels are also believed to be formed by condensation and are thus reversible to sols. Out of the solution phase crystalline sphalerite is formed which might exhibit gel relic structures. It should though be mentioned that typical gel sphalerite is common and the substance is not characterized by a great readiness to crystallize since colloform sphalerite is often found. Figs. 639 and 640 show gel sphalerite associated with marcasite either forming a margin to the marcasite, see Fig. 639, or being interbanded with it, see Fig. 640.

(c) Gel Phase

Interbanded

In addition to gel sphalerite, interbanding with marcasite often takes place. The gel phases are either interbanded or occur as gel phase interzonal to a crystalline phase or a fine crystalline phase may even be interspersed in the colloform phase. Fig. 641 shows colloform sphalerite (iSchallenblende) interbanded with gel gratonite (Pb9As4S15). It should be mentioned that despite its Kristallisationsfreudigkeit gratonite exhibits colloform structure here and is interbanded with sphalerite (ZnS).

It is most likely that these interbanded gel phases have been formed by coagulation from the same solution phase. Another interpretation would be the alternation of colloidal solution phases. A rather rare case of gel phase interbanding is illustrated in Fig. 642 which shows fine colloform interbanding of galena (PbS) and anglesite (PbS04). The comparison of the chemical compositions of galena and anglesite shows that derivation of these interbanded gel phases from the same colloidal solutions is most probable. In contradistinction to the interbanded gel phases described, colloform malachite may occur as an interzonal phase in cuprite (Cu 2 0) and in this case as a derivation of the crystalline cuprite. The colloform malachite may have been formed from the same solutions or their formation may be due to alternation of solution phases (Fig. 643). Whereas cuprite and malachite are copper-rich minerals and their derivation from copper-rich solutions is explainable, the case of smaltite ((Ni,Co)As3) with interzanal colloform pitchblende is more problematic in the sense that the pattern might represent an interruption in the crystallization of smaltite from solutions and the intervening coagulation in the crystallization and formation of the interzonal or interbanded pitchblende. It should be mentioned though that there must have been a change in the nature of the solutions which resulted in the crystallization of smaltite and the coagulation from colloidal solutions of pitchblende. In contrast to the interbanding or interzonal relation of a crystalline and a gel phase, cases are observed where an interspersed fine crystalline phase is included in a gel phase. Figs. 644 and 645 show small gratonite crystals (in cases intergrown with hutchinsonite) included in a gel phase exhibiting gel banding of hutchinsonite ((Tl,Pb)S{Ag,Cu)2S-5As2S3). Considering the composition of gratonite (Pb9As4S15) it is possible that crystallization of the gratonite crystalline phase occurs from the same solutions as well as the coagulation-formation of the hutchinsonite from colloidal solution phase. In this connection, it should be mentioned that gratonite is characterized by a great readiness to crystallize (Kristallisationsfreudigkeit). Concerning the intergrowth and interrelationship of gel and crystalline phases, the case of galena (PbS) related texturally with plattnerite (Pb02) is most interesting (Figs. 646 and 647). In this connection it should be emphasized that galena occurs only rarely as colloform (Fig. 642) and that it is characterized by a readiness to crystallization. It is possible that plattnerite has been formed from colloidal solutions and that galena, due to its great readiness to crystallization, did crystallize from the same solution. This is supported by the fact that galena and plattnerite are essentially lead minerals.

(d) Interbanding

of Colloform Mineral

Phases

In addition to malachite as an interzonal colloform band in cuprite, cases of rhythmical interbanding of tenorite (CuO) and chrysocolla (CujHjSijOjiOH)^ are shown in a series of figures. Fig. 648 shows chrysocolla interbanded with tenorite exhibiting rhythmical colloform banding with the tendency for the outer zones to be richer (broader bands) of tenorite. In contrast, Fig. 649 shows complex colloform interbanding of chrysocolla and tenorite with the outer zone being rich in chrysocolla. As Figs. 650 and 651 show, the thickness of the respective tenorite and chrysocolla zones varies, especially in the pattern in Fig. 561, which shows a central colloform chrysocolla mass followed by a margin of tenorite. Fig. 652 depicts chrysocolla with a tenorite external band indicating indentationlike structures in its relationship with the chrysocolla. In the contact skarn of the Seriphos granite (Seriphos Island, Greece) complex rhythmically interbanded lievrite [CaFe2Fe3(OH)Si2Og] and brown iron colloform patterns are found (Fig. 653), and similarly complex lievrite-haematite (Fe 2 0 3 ) rhythmically banded colloform structures also exist in the same skarn bodies (Fig. 654). The colloform rhythmical interbanding of lievrite and haematite is most probably a derivative from the same colloidal solutions. It is thus possible to have a rhythmical interbanding of two gel mineral phases from the same colloidal solutions when relationships (chemical-geochemical) exist as their formulae indicate.

(e) Complex Gel Patterns Exhibiting Rhythmical Banding and Interbanding of Gels Considering the possibilities of gel patterns some complex gel patterns of manganese minerals and brown iron and other cases are presented here, which indicate that intricate processes are involved in the generation of colloform and banded colloform textures and especially their transformation to crystalline aggregates (often radiating). Figs. 655-659 and descriptions present and discuss manganese patterns. Studying these interbanded colloform phases further it is interesting to note the interbanded pyrite, chalcopyrite, uraninite structures which are derivatives from colloidal solutions. Fig. 660 shows gel pyrite interbanded with chalcopyrite (also a derivative of colloidal solutions) with a "colloform" margin to the gel pyrite of uraninite, originally colloform, which has subsequently been converted to crystalline uraninite. Additional to the rhythmically interbanded brown iron lievrite colloform pattern, cases are exhibited of brown iron rhythmically banded structures which are derivatives of the alterations and solutions of a complex ilmenite with haematite (oriented bodies of haematite in the ilmenite, Fig. 661). Comparable rhythmically banded brown iron colloform structures 81

occur again due to alteration and solution of an ilmenite-haematite granular pattern with which the colloform structure is associated (Fig. 662). The formation of brown iron rhythmically banded colloform patterns either in enclaves formed by the solution of iron-rich minerals (see Figs. 661 and 662) or, in cases as needle iron, often rhythmically banded with haematite (Figs. 663-665), result in impressive banded colloform structures surrounding clastic chromite grains (Fig. 664), or aggregates of oolites or pisolites in secondary Ni-Cr lateric iron deposits indicating that iron-rich solutions -derivatives due to alteration of initial iron-rich minerals - often result in the formation of complex colloform structures.

( f ) The Interrelationship line Phases

of Gel and Crystal-

With respect to the relationship of gel and crystalline phases, only a few examples were presented of transformations of gel structures to the crystalline state. The significance of the gel to sol reversibility was mentioned, which in turn depends on whether the gel was formed by condensation or by dispersion. Perhaps the coexistence of gel and crystalline states for the same substance also depends on whether the gel phase is formed by condensation or dispersion. As pointed out, gels formed by condensation are reversible, thus it is possible that a crystalline phase precipitates from the same solution or from a differentiated fraction of it while from the same solution a gel phase can be produced by condensation. While gels formed by condensation are more reversible to sols it is possible that from these sols solutions can be generated out of which crystallization takes place. Fig. 666 shows the simultaneous existence of gel pitchblende and crystalline uraninite. Such textural patterns could result from the same solutions or a true solution will give rise to the crystalline uraninite while a colloidal solution by condensation will form the gel pitchblende. In this case there are perhaps two different solutions available at different time intervals. In contrast, cases may exist where small crystals of uraninite segregate into a marginal phase of a cobalt mineral or form aggregates assuming gel form (Figs. 667 and 668). However, there is a kind of incompatibility inherent to the coexistence of crystalline and gel uraninite adjacent to each other, nevertheless, such patterns do exist. An additional example of the coexistence of a crystalline and a gel phase in the same sample is shown in Fig. 669. A gel pyrite is surrounded by a border of pyrite crystals. In this case it is possible that the gel phase was formed first and that the crystalline margin represents a later formation.

82

(g) Textural Patterns of Crystallite Aggregates Either as Derivatives from Colloidal Solutions or Due to the Transformation of Initial Gels A great complexity of patterns exists which in fact are derivatives of colloidal solutions either in form of gel structures - some have been discussed already - or of fibrous or microcrystalline aggregates derived by the transformation of initial gels (the general part of this section presents a theoretical overview of the processes and phases involved). As a result of the coagulation of colloidal solutions, gel sphaeroids can be formed. On subsequent transformation they result in radiating fibrous crystallites building up the sphaeroid. Fig. 670 shows part of a psilomelane sphaeroid [BaMn 2 Mn 6 4 0 16 (0H) 4 ] followed by a shell of braunite [3(Mn,Fe) 2 0 3 MnSi0 3 ] in which elongated needle-like crystallites of braunite are formed perpendicular to the braunite shell layer (Fig. 671). A wide spectrum of textural patterns can result from subsequent crystallization or rather transformation of initial gel psilomelane structures. Fig. 672 shows a dendritic pattern consisting of fine psilomelane. A rather more compact pattern of "dendritic" in appearance psilomelane is shown in Figs. 673 and 674. Comparably complex dendritic forms can be produced from initial gel phases. As Fig. 675 shows, native silver, dendritic in pattern, may be overgrown with chalcocite. The dendritic textural pattern of native silver, a crystalline derivative of initial gel, is further illustrated in Figs. 676 and 677. Similarly chalcocite might be surrounded by a brown iron dendritic pattern and in this case, a derivative of an initial gel phase, formed either by direct crystallization from "colloidal solutions" or by transformation of initial gel structures to crystalloids assuming a dendritic pattern (Figs. 678 and 679). Furthermore, as a result of the transformation of initial gel structures complex patterns may result where prismatic microcrystals may be interspersed or oriented within an initial gel sphaeroid. Fig. 680 shows lepidocrocite (FejOj lHjO), a gel phase in which prismatic lepidocrocite crystallites and haematite are formed. In contrast to other gel sphaeroids of lepidocrocite a central aggregate of lepidocrocite crystals might be formed with a layer of gel lepidocrocite in which gel haematite is also formed. This layer is followed externally by a more compact lepidocrocite gel layer (Fig. 681). Fig. 682 shows a dendritic pattern of lepidocrocite crystallites formed by the transformation of an initial lepidocrocite gel phase. Similarly, the series of Figs. 683-686 shows lepidocrocite crystallites dendritic in pattern either as central parts of lepidocrocite gel sphaeroids or showing crystalline lepidocrocite together with gel lepidocrocite and haematite in which gel patterns may exist as well. There can be complex

intergrowths between the gel lepidocrocite phase and the gel haematite. In the general part of this Chapter an attempt was made to correlate the transformation of gel sphaeroids to sphaeroids consisting of radiating elongated crystals or fibriolitic crystallites (their crystallization starting from the center of the sphaeroid). From the above discussion it is apparent that the general principles of the relationship of gel sphaeroids and their subsequent transformation to sphaeroids consisting of radiating fibriolitic crystallites or of elongated crystals find wide application and perhaps can provide the explanation to most of the cases of sphaeroids consisting of radiating crystallites. In addition to the cases already discussed, Fig. 687 shows radiating metacinnabar (HgS) elongated crystals radiating from the center of the sphaeroid a part of which is shown in Fig. 687. However, within the metacinnabar sphaeroid, metacinnabar crystallites exhibiting interpenetration intergrowths, interpenetration textures and zonal growths are shown in Figs. 688-690. Comparable gel-derived sphaeroids which have been transformed subsequently to sphaeroids consisting of radiating hausmannite (Mn 3 0 4 ) crystals are shown in Fig. 691. Also in this case, the elongated hausmannite crystals radiate from the center of the sphaeroid which is also a derivative from gels. In opposition to the mentioned sphaeroids where elongated crystals radiate out of the center of the sphaeroid, cases are described of prismatic crystals extending from the periphery of the sphaeroid inwards. Fig. 692 shows a sphaeroid consisting of a central part of malachite in which elongated brown iron crystallites radiate inwards from the periphery. These sphaeroids, too, are derivatives of complex gels with subsequent transformation to brown iron crystallites associated with malachite. Levicki (1955) described colloform cassiterite (Sn0 2 ) and several references exist about "wood cassiterite" which can be regarded as a type of colloform cassiterite. Present observations show an aggregate of radially arranged cassiterite crystallites simulating a sphaeroid in shape (Fig. 693). In contrast, the cassiterite aggregates may be more diffuse in form (Fig. 694). The ones presented in Figs. 693 and 694 are most probably derivatives of initial cassiterite gels. On the other hand instances exist of radiating elongated cassiterite (a part of the cassiterite sphaeroid) associated with stannite which in turn is replaced by later pyrite (Fig. 695). It is believed that here also the cassiterite sphaeroid, consisting of radiating elongated cassiterite crystals, is a derivative of gel cassiterite due to the subsequent transformation of gel cassiterite into a sphaeroid consisting of radiating cassiterite crystals.

(h) Special Cases of "Colloform Patterns" due to Their Readiness to Crystallize (Kristallisationsfreudigkeit) In contrast to the typical gel banded structures and to the colloform patterns described so far a selection of rather untypical colloform patterns or cases of minerals that rarely exhibit colloform patterns are presented here. Fig. 696 shows a colloform marcasite (FeS2) including and replacing sphalerite (ZnS). This colloform pattern common in marcasites is exhibited also in the case of galenobismuthinite (PbBi2S4) which rarely shows colloform patterns perhaps due to its tendency to crystallize. It should be emphasized though that the colloform pattern shown in Fig. 697 actually represents a relic phase due to the transformation of the greater part of the galenobismuthinite to the crystalline state. Two additional, rather disputable colloform patterns are exhibited in Figs. 698 and 699 which show niccolite (NiAs) associated with gangue (?carbonates) forming the central part of the colloform "sphaeroid", which is followed by a broad initial gel layer of rammelsbergite (NiAs2) or pararammelsbergite (NiAs2) margined by a thin rim of safflorite (CoAs2). This again is believed to be a derivative from the initial gel sphaeroids but it should be stressed that this layered structure - despite the exhibited colloform pattern - consists of a crystalline phase derived from the transformation of the initial gels to crystalline phases. As in the case of galena and galenobismuthinite, also niccolite, rammelsbergite and pararammelsbergite show great tendency to crystallize. The result is that colloform patterns are rarely preserved and often they are precipitated out of solutions in the crystalline state.

(i) Colloform Structures Secondarily due to Alteration

Formed

In contrast to the colloform patterns formed from colloidal solutions cases are observed where colloform patterns are due to alteration of a pre-existing crystalline mineral as a result of intracrystalline infiltration of solutions causing solution front mobilization within an existing crystalline mineral phase. Fig. 700 shows a malacon which is a metamictic phase of zircon. Due to the metamictic state, the zircon lattice is broken down and possibly absorption of water by the malacon took place. It is conceivable that the zonal growth of the zircon is preserved in the metamictic malacon and that in addition secondary colloformlike structures are produced, in some cases independent of the preserved zoning. Such colloform patterns could be produced by the operation of the solution front which moved within the malacon and resulted in element mobilization (see also Fig. 701). Comparable secondary gel patterns can be produced in pyrrhotite (FeS) which, due to alteration, may 83

change to marcasite exhibiting colloform sphaeroidal patterns (Fig. 702). Equivalent alteration colloform patterns are also produced in davidite as a result of advancing alteration fronts which might result in element leaching out and forming secondary minute rutile grains (Fig. 703).

( j ) Synaeresis Gels when subjected to drying processes shrink and the developing shrinkage cracks are known as synaeresis cracks. In some cases the synaeresis cracks assume hexagonal patterns, in others they are sphaeroidal and more often they start from the center of gel sphaeroids and radiate outwards. Sometimes the synaeresis cracks act as channels or spaces in which mobilization of mineral substances can take place. In other instances the synaeresis cracks act as channel ways for more pronounced oxidation or alteration of the colloform structure in which they occur. Cases where coffinite has been mobilized and introduced into the uraninite exhibiting hexagonal or cyclical in cross-section synaeresis cracks are exhibited in Figs. 704 and 705. Fig. 706 shows a combination of hexagonal and cyclical in cross-section synaeresis cracks which are partly occupied by coffinite or by radiogenic lead (galena). As mentioned, some of these cracks might act as oxidation channels of the uraninite in which they occur. The most common kind of synaeresis cracks, however, developed in drying gel pitchblende sphaeroids is that of the radiating type as in Fig. 707 which also shows the presence of some radiogenic lead. In contrast to the type of oxidation of pitchblende marginal to the gels or following cracks, synaeresis

84

cracks may also develop in pitchblende independently to the oxidation portions of the pitchblende (darker in colour) due to formed U0 3 . As mentioned, the pattern of the synaeresis cracks bears no direct relation to the differences in the composition of the pitchblende in the shells (zones of growth) of the gel sphaeroids (Fig. 708). Some pitchblende though exhibits synaeresis cracks showing darker in colour pitchblende between those cracks (see arrow "p" in Fig. 709). As mentioned, often in synaeresis cracks of pitchblende radiogenic lead is laterally segregated or mobilized. The fillings of the synaeresis cracks of pitchblende can be occupied by a variety of minerals. Fig. 710 shows pitchblende small gel structures associated with smaltite and safflorite and with some arsenopyrite as the filling of the synaeresis of the pitchblende. An interesting case is exhibited in Fig. 711 showing cyclical in section synaeresis cracks in pitchblende occupied by later mobilized clausthalite. In addition, Fig. 712 shows clausthalite occupying synaeresis cracks in pitchblende as well as interzonal (inter-shell) spaces. In Fig. 713 clausthalite occupies inter-shell spaces of the pitchblende as well as dissolved central parts of the gel pitchblende. In the same pattern covellite occupies synaeresis cracks and inter-shell spaces (Figs. 712 and 713 as well). As already discussed, mobilization of solutions along the synaeresis cracks of pitchblende is common and variable. Fig. 714 shows bismuthinite occupying the synaeresis cracks of pitchblende and in contradistinction, synaeresis cracks in pitchblende occasionally show multiple deposition. The margins of the synaeresis cracks are occupied by galena whereas in the central part tetrahedrite is formed (see Fig. 715 and its detail Fig. 716).

Chapter 13

(a) General - Manganese

Sphaeroidal Structures and Textures

Nodules

Besides the colloform patterns (see Chapter 12) which exhibit sphaeroidal structures and which mainly result from gel coagulation, there are also other important sphaeroidal patterns in ore minerals. Manganese nodules, platinum (gold nodules) nuggets and oolitic pisolitic textures and structures in bauxites are some selected examples of sphaeroidal structures and textures which will be discussed briefly. Augustithis (1982) extensively studied sphaeroidal structures and textures as documented in his Atlas of Sphaeroidal Textures and Structures and their Genetic Significance. Manganese nodules have been the subject of many studies for the last two decades and extensive treatments are available, some references are quoted here. This topic was also discussed by Augustithis (1982). Following are some selected exemplifications on manganese nodules. Fig. 717 shows a manganese nodule from the Pacific Ocean near Hawaii, and as the section of it shows, the nodules are sometimes made of a non-metallic nucleus with a metallic cover (Fig. 718). An attempt will be made to present textural patterns from some of the nodules studied in order to demonstrate the complexity of the structures exhibited. Overall however, it could be concluded that some patterns are colloform or colloform-like. Fig. 719 shows an altered fragment of rock, most probably basaltic in composition surrounded by fine bands of manganese oxides (ramsdellite (Mn0 2 )) alternating with fine transparent mineral bands. Similarly Fig. 720 shows banded "colloform manganese" oxides and (hydroxides) with alternating transparent mineral bands. In addition, interspersed metallic grains are present in the matrix. Also in this case, the pattern exhibited is colloform or colloform-like, indicating that colloidal solutions might have played a role in the formation of the nodules. In opposition, cases are also considered that resemble typical "cauliflower structures" found in stromatolites (see Augustithis, 1982). Fig. 721 shows a pattern that could either represent complex colloform structure or a section of a "cauliflower structure" of a stromatolite. Augustithis (1982) presented detailed microanalytical studies of these nodules and an extensive review of the genesis of the nodules is also included. (To avoid repetition, please see Chapters 17 and 18 of the Atlas of

Sphaeroidal Textures and Structures and their Significance, Augustithis, 1982).

Genetic

(b) Platinum Nuggets and the Controversy over Platinum Nugget Formation Platinum nuggets were discovered by Russian mining engineers as placers in the Yubdo river (Wollaga, W. Ethiopia) at the end of the nineteenth century. One of the earliest studies on the ultrabasic complex of Yubdo was carried out by Duparc et al. (1927), who described a new rock type occurring as a cover on the dunite and named it birbirite after the small river Birbir (a tributary of the Yubdo river). Augustithis (1965) studied the ring complex of Yubdo and presented a detailed study of the rock types, dunite, pyroxenites and birbirites. In contrast to previous view points that considered Pt in the complex to be associated with olivine, he found sperrylite as an inclusion in the chromite, associated with epitactic magnetite on the idiomorphic chromite grains in the dunite. In his paper Mineralogical and geochemical studies of the platiniferous dunite, birbirite-pyroxenite complex of Yubdo (Birbir) Ethiopia (1965), he emphasized the derivation of birbirite as being from the dunite by leaching out the Mg from the olivines and by alteration of the chromite, while the epitactic magnetite changed into iron hydroxides. However, it should be stressed that in the alteration of dunite to birbirite the sperrylite that was originally included in the chromite of the dunite, or occurred in the epitactic magnetite, resisted the birbiritization processes. Augustithis (1965) pointed out that "the comparison of dunite and birbirite reveals that the latter contains structures, intact or partly altered, that are beyond doubt of dunitic origin. Also, structural studies showed that besides the Mg leaching out partial Fe and Cr mobilization occurred. In contrast to these mobilized components, Pt in the form of sperrylite has hardly been mobilized. This should be seen as a result of the great resistance of sperrylite towards alteration processes". In this way, he drew attention to the resistance of sperrylite that overall remained unaltered in the birbirite. This is based on the fact that sperrylite is found in birbirite [(an alteration derivative of dunite) consisting of greatly altered chromite, chalcedony (a derivative of 85

olivines due to Mg leaching out) and iron hydroxides. Also, sperrylite grains were found in the birbirite.] The fact that sperrylite resisted the birbiritization process is supported by the existence of sperrylite in the birbirite. In his paper, Augustithis also showed that the Pt in the dunite was mainly in the form of sperrylite and that sperrylite was preserved in the birbirite. Thus, it is important that the source of the Pt minerals was found in the form of sperrylite which by alteration contributed to the formation of the Pt nuggets in the lateritic cover (Fig. 722). In a later paper, On the phenomenology and geochemistry of different leaching and agglutination processes (1967), Augustithis used the term agglutination as synonymous to accretion and refers to the segregation of the platinoid elements to form nuggets in the Yubdo lateritic cover derived from the alteration of the underlying platiniferous dunite. In this paper he pointed out the interrelationship of the platinoid elements building the nuggets in accordance with the empirical laws of element interrelations according to the periodic system. In a further paper, together with Ottemann (1967), Geochemistry and origin of platinum nuggets in lateritic covers from ultrabasic rocks and birbirites of W. Ethiopia, it was the first time 4 in history of geochemistry and mineralogy of PGE and PGM that platinum elements were considered to be mobilized at low temperatures (those prevailing under lateritization processes - impossible at the time) which could form platinum nuggets (consisting of ferroplatin and gold) containing several PGM such as osmiridium, roseite and others. This interpretation led to controversy. Although the authors were criticized and strongly opposed by Cabri and his associates, Cabri et al. (1981), studying the Pt nuggets from Yubdo, Ethiopia, admitted that some of the nuggets had scalloped and embayed surfaces. They furnished additional details on the mineralogy of the Yubdo Pt nuggets. Also Bowles (1986) states the following: "The possibility of platinum group element mobility at ground-water temperatures has been considered (Fuchs and Rose, 1974; Westland, 1981; Hodge et al., 1985) and it was suggested by Ottemann and Augustithis (1967) for the origin of the angular Pt-Fe alloy nuggets found in laterites and birbirites at Yubdo (Joubdo, Youbdo, Yubda, or Joubda) in W. Ethiopia." Ottemann and Augustithis (1967) and particularly Augustithis (1979) supported that the nuggets in the laterite were formed from the sperrylite that was included originally in the ultrabasics and in the birbirite. In contrast to the argument that Augustithis supported the non-solubility of Pt under lateritization conditions, the section on Evidence of the growth of platinum nuggets in the lateritic cover (Ottemann and Augustithis, 1967) will be cited: ' With the exception of osmiridium which was probably recrystallized in the Witwatersrand Conglomerate, see Koen,

(1964). 86

"The following is the main evidence for growth of Pt nuggets in the lateritic soil: (i) Ore microscopic examination has failed to show platinum grains, of the size of the nuggets, in the ultrabasic rocks; in fact, no native platinum has been observed so far. Thus, derivation of the platinum nuggets from platinum grains of the ultrabasics is almost totally excluded. Perhaps the primary Pt carriers are sperrylite and other platinoid minerals - so far unidentified - containing Pt or platinoid elements. (ii) The angular shape and protuberances of the Pt nuggets (which show no rounding due to transportation) can be taken as evidence of suggested growth in the lateritic soil. The eluvial nature of the lateritic cover, in addition to the complete lack of rounding of the Pt nuggets, suggests that they have been formed by growth during the alteration of the ultrabasic rocks or perhaps during lateritization. (iii) Platinum nuggets enclose chromite grains identical to those of the ultrabasics. It should be pointed out that whereas some of the chromite grains are unaltered, others show alteration and corrosion margins (Figs. 723 and 724). (iv) The presence of altered chromite grains as inclusions in the nuggets supports the view that the nuggets' growth has taken place in an environment of disintegration of the ultrabasics or of the birbirite (it is here supported that alteration of birbirite took place under lateritization), where, nevertheless, chromite was a relatively resistant phase. (v) As already mentioned, platinum nuggets have a coating which, with the aid of ore microscopy and microprobe analysis, is shown to consist mainly of iron oxides. In addition, ore microscopic observations show that such a coating also exists around chromite grains which are included in the ferroplatin nuggets (Fig. 725). The coating around platinum nuggets (Fig. 723) could be formed during lateritization in a similar manner as the coating surrounding the chromite grains included in the nuggets. Thus, the fact that the ferroplatin itself encloses chromite grains with such an iron oxide coating suggests that it was probably formed under lateritic conditions. (vi) Electron scanning showed that gold is present in ferroplatin. Furthermore, geological (field) studies proved the presence of auriferous quartz veins transversing the Yubdo complex. It is possible that Au present in the ferroplatin was freed from the veins and incorporated in the lattice of the ferroplatin during the alteration process. Another explanation for the presence of gold in the ferroplatin could be its derivation from the primary rock. This would imply that gold was separated from the rock under alteration through colloidal dispersion and was later incorporated in the ferroplatin through auto-aggregation as reported by Goni et al. (1967). This study which was concerned with the mobilization and aggregation of gold may also serve to explain the formation of the platinum nuggets."

The quotation shows that the authors support a derivation of the PGE element building the nuggets by alteration of the original sperrylite and other possible PGE that existed in the ultrabasic rock dunite (and also in its derivative, the birbirite, since sperrylite resisted the alteration processes of birbiritization and it was also found in the birbirite (see Augustithis, 1965). The paper by Ottemann and Augustithis Geochemistry and origin of platinum nuggets in lateritic covers from ultrabasic rocks and birbirites of W. Ethiopia, started, as mentioned, a controversy that was divided in two phases: in the beginning (the polemic phase) most of the opponents of this interpretation opposed the view that PGE elements can be mobilized under low temperature conditions by alteration. In this connection, a section from the Atlas of the textural patterns of basic and ultrabasic rocks and their genetic significance (Augustithis, 1979) will be cited with the title: "Low temperature mobilization of the Pt group elements in lateritic covers (synoptical discussion)": "Lateritization of platiniferous dunite and the low temperature mobilization of the Pt group elements to form nuggets (consisting of a complex platinum mineral paragenesis) has been described by Ottemann and Augustithis (1967) in lateritic covers of Yubdo, W. Ethiopia dunitic intrusion. In contrast to this hypothesis of element agglutination (accretion) proposed by Ottemann and Augustithis (1967) and Augustithis (1967) for the formation of Pt nuggets in the lateritic eluvial cover of Yubdo (W. Ethiopia) dunite, Cabri and Harris (1975) suggested that the nuggets are primary growths in the dunite and that they are liberated in the eluvial lateritic cover after the disintegration of the dunite. According to Cabri and Harris, the absence of ferroplatin in the size of nuggets (Fig. 722) in the unaltered dunite is attributed to the platinum grains being far more dispersed in the unaltered dunite. However, studies by Duparc et al. (1927) and Augustithis (1967) failed to show the presence of primary Pt group minerals of the size of the nuggets discovered in the eluvial cover. In fact, Duparc believed Pt to be in the lattice of olivine. Fire assay of the dunite, despite the absence of nuggets in the crushed rock, showed the presence of Pt (more than 1 ppm), suggesting that Pt group minerals (perhaps more than sperrylite) are present in the unaltered dunite of Yubdo. In the early seventies the U.N. special fund (Ethiopian Mineral Survey - Preliminary Report on Magnetic and Geochemical Surveys over the Yubdo Ultrabasic Complex, Yubdo Wallaga (1969)) started an extensive investigation by drilling for primary Pt in the unaltered dunite of Yubdo. No ferroplatin of the size of the nuggets was found in the primary unaltered dunite. Pt nuggets of the size and appearance as illustrated in Fig. 722, are abundant only in the eluvial cover of the dunite." Cabri and Harris (1975) also discussed the paper by Ottemann and Augustithis (1967) regarding the presence of altered and unaltered chromite in the ferro-

platin nuggets and suggested that the Pt could grow around the chromite grains during the crystallization in the dunite. However, it should be noted and it is emphasized in the "Atlas", that the ferroplatin nuggets contain altered chromite grains, i. e., chromite grains showing distinctly chemical alteration during birbiritization or lateritization (Figs. 723, 724) prior to their engulfment by the later ferroplatin. In support to the accretion hypothesis, Cousins and Kinloch (1976) stated the following: "Physical evidence for the state of accretion ("element-agglutination "of Ottemann and Augustithis, 1967) is, in the opinion of the authors, evident in the rounded and sculpted form of alluvial and especially platinoids, in the size of the grains (six or seven times larger than Merensky reef platinoid), and in the zoning evident in microscopic and electron probe studies of grains (Stumpfl, 1974; Cabri and Harris, 1975; Barrass, 1974; and this paper)". Thus, the polemic phase ended more or less in favor of the interpretation of the nugget formation under alteration processes of ultrabasic rocks, about 20 years after the publication of the papers (Augustithis, 1965; Ottemann and Augustithis, 1967; Augustithis, 1967)s, and Bowles (1986) concluding in his paper: "Theoretical studies have shown the platinum elements to be mobile in latentes under very acid, chloride-rich conditions with a high Eh. The textural evidence obtained from a study of the platinum group minerals from Guma Water, Sierra Leone, indicates that platinum, iridium, osmium and ruthenium can be taken into solution under the conditions found in a lateritic environment, transported and subsequently deposited under different conditions. The change in conditions may be brought about by a change in the path taken by the fluid circulating within the latente or by entry of the fluid into a different horizon or region of the laterite. When deposition occurs, large crystals of platinum group minerals can develop, showing welldeveloped crystallographic features. Variations in conditions of deposition may lead to resolution or etching of the surface of the crystal, whereas a return to condi5 (i) Mineralogical and geochemical studies of the platiniferous dunite, birbirite-pyoxenite complex of Yubdo, (Birbir) Ethiopia, Augustithis, 1965: Sperrylite was found associated with chromite and magnetite in the dunite and it resisted the alteration of dunite into birbirite since sperrylite is found in the birbirite. (ii) Geochemistry and origin of platimum nuggests and lateritic covers from ultrabasic rocks and birbirites of W. Ethiopia, Ottemann and Augustithis, 1967: supporting the generation of Pt nuggets in the laterite cover derived from the dunite underneath; (iii) On the phenomenology and geochemistry of different leaching and agglutination processes, Augustithis, 1967: the platinoid element segregation to form Pt nuggests was explained to be due to the interrelationship of these elements according to the empirical laws of the periodic system.

87

tions favoring deposition permits continued growth of the crystal over the etched surface. The platinum group minerals recovered from Yubdo, Ethiopia, show similarities with those from Guma Water and may be another prominent example of the proposed process which is analogous to the solution, transport, and development of gold within the laterite. It has been shown that organic colloids aid in the solution and transport of gold, and it may well be that these are significant in the movement of platinum group elements. Formation of gold nuggets by this process has been offered as an explanation for the larger size of gold nuggets retrieved from the secondary environment rather than from primary source rock and it is suggested that the larger size of alluvial and eluvial platinum group minerals is also due to growth in the secondary environment. The important economic significance of this work is that near-surface pockets high in platinum group minerals may be found at appropriate horizons in laterites overlying basic or ultrabasic intrusions."

(c) Ore Microscopic and Microprobe Studies of the Ρt Nuggets of Yubdo, Wollaga, W. Ethiopia The presence of Pt nuggets in lateritic eluvial covering the dunitic and partly birbiritic rocks of the ultrabasic complex of Yubdo is a fact that attracts the attention of geologists and especially of geochemists and mineralogists. The prevailing concept that the PGE are not dissolvable at low temperatures which prevail with weathering and alteration of rocks and the fact that ferroplatin and other platinoid minerals such as osmiridium are believed to be crystallizations from dunitic magmas at high temperatures, render any explanation doubtful and in fact suspicious that is based on the concept of the Pt nuggets being formed in lateritic covers. This was the attitude Ottemann and Augustithis (1967) encountered some 25 years ago. The interpretation that Pt nuggets can be formed in low temperature solutions is supported by the work of Westland (1981) and Hodge et al. (1985) who showed physicochemically that Pt and the Pt group of elements are dissolvable (i. e., solutions are possible) at low temperature conditions. Ottemann and Augustithis had contended that the ferroplatin nuggets of Yubdo also contained appreciable amounts of gold. In this connection it should be pointed out that the similarity of Pt nuggets and gold nuggets was emphasized and reference was made to the origin of gold nuggets at low temperatures as explained in the papers of Goni et al. (1967), Machairas (1963), and others. An attempt will be made to present the phenomenology of the Pt nuggets from Yubdo based on ore microscopy and it is referred to the microprobe studies made on these nuggets by Ottemann and Augustithis, (1967). Fig. 722 shows an average-sized Pt nugget 88

coated with a layer of Fe-Mn oxides. It should be noted that the platinum nugget is sculpted and its shape suggests that it has been formed within the laterite. Both the size of the nugget (it is 6-7 times the size of the Pt minerals of the Merensky reef), and the sculpted surface of the nuggets have been considered by Cousins and Kinloch (1976) to represent growth in the laterite by accretion. Fig. 723 shows, as mentioned, a polished section of the auriferous ferroplatin enclosing corroded and altered chromite grains (derivatives from the altered dunite) which may have acted as nuclei for the precipitation of low temperature PGE solutions in the lateritization environment. It should also be emphasized that the nugget shows zonal structure, as indicated by the distribution (zonal) of osmiridium crystals in the ferroplatin. In this connection it needs to be pointed out that the zonal growth of the ferroplatin has been used as an indication by Cousins and Kinloch for the accretion origin of the nuggets. However, the osmiridium crystals have been formed within the lateritization environment. This again contradicts the orthodox view that osmiridium would only crystallize at very high temperatures during the consolidation of the dunitic magma. Fig. 724 (a detail of Fig. 723) shows the corrosion and alteration of the chromite grain included in the auriferous ferroplatin and the presence of an idiomorphic (?idioblastic) osmiridium crystal. Fig. 725 shows a nugget where the auriferous ferroplatin keeps several chromite grains together, some of them are altered. In some cases an Fe-Mn oxide margin surrounds chromite grains as indicated in Fig. 725 which is comparable to the Fe-Mn margin surrounding the ferroplatin nugget shown in Fig. 723. As Fig. 725 and its detail Fig. 726 show, the Fe-Mn margin has been formed surrounding the chromite grain (also an adjacent chromite grain in the ferroplatin shows extensive corrosion) and can perhaps be considered to have been formed before the ferroplatin (it should be recalled that most of the Fe-Mn margins are later than the ferroplatin since they form margins (coatings) of the nuggets (see Fig. 722). However, the relationship between ferroplatin and the Fe-Mn margin is not always simple. Fig. 727 shows ferroplatin with a Fe-Mn margin which in turn is surrounded by later mobilized ferroplatin (some of the ferroplatin is included though in the margin). In contradistinction to the cases where the Fe-Mn margin surrounding chromite grains is enclosed by ferroplatin, Figs. 728 and 729 show Fe-Mn oxides (margins) with extensions following a crack of the ferroplatin, clearly indicating that the Fe-Mn margin is a post-ferroplatin formation to a great extent (without though excluding the possibility that some was formed parallel with the accretion of the auriferous ferroplatin, see Figs. 725 and 726). Additional evidence is provided by the pattern in Fig. 730 which shows an altered chromite grain included in the ferroplatin which in turn is surrounded by an Fe-Mn margin sending a protuber-

ance into the ferroplatin. An interesting case is presented in Fig. 731 showing osmiridium crystals in the peripheral part of the ferroplatin nugget which is surrounded by the later Fe-Mn oxide margin. It should be pointed out though that extensions of the margin have exploited the space between the osmiridium and ferroplatin and which partly replaced the osmiridium (see arrow "m", Fig. 731). Further patterns are shown expressing the complexity of the relationship between ferroplatin and the FeMn margin. Fig. 732 shows ferroplatin with a margin of Fe-Mn oxides which is mobilized along a crack in the ferroplatin; at the same time, ferroplatin is mobilized and forms a thin margin on the Fe-Mn oxides (margin of the Pt nugget). Considering the relationship of ferroplatin and the other minerals included in the nugget, the cases shown in Figs. 733 and 734 are of interest, depicting elongated fibriolitic silicates (perhaps antigoritic asbestos) included in the ferroplatin. Figs. 733 and 734 also show chromite with epitactic magnetite (martitized) occurring in the ferroplatin. In this connection it should be pointed out that these chromites have not passed the birbiritization alteration phase since magnetite would have been altered to iron hydroxides. As shown in Fig. 734, ferroplatin replaces the epitactic magnetite margins of the chromite. The alteration of the epitactic magnetite supports a later replacement of the magnetite by the ferroplatin.

(d) Platinoid Minerals Grown in the Ρt Nuggets That Have Been Formed by Accretion As already mentioned, there is a tendency for the osmiridium crystals to be zonally arranged within the ferroplatin nugget (Fig. 723). Occasionally, the osmiridium crystals are in the peripheral parts of the nuggets as shown in Figs. 724 and 731 (in this case some replacement of the osmiridium by the Fe-Mn oxide margin took place). In this connection it should be stressed that other platinoid minerals occur in the peripheral part of the nugget as exhibited in Figs. 735 and 736 (there is a general view of the marginal part of a ferroplatin nugget). The following is an attempt to present some of the platinoid minerals that occur in the Pt nuggets of Yubdo. In addition to osmiridium (Figs. 723 and 724), microprobe studies (Ottemann and Augustithis, 1967) proved the presence of a number of minerals, some of which might have been new minerals (the attempt to name an (Os, Ir)S mineral "roseite" was not accepted because of insufficient quantitative and crystallographic data). Nevertheless, a number of minerals of unique composition were described at that time. In addition to osmiridium, "roseite" with the composition (Os, Ir)S was determined on the basis of microprobe studies as well as minerals "a" and "b" have been discovered (Fig. 737 and description there). Mineral

"a" consists of nickel, sulfur, palladium, rhodium and iron, and mineral "b" is distinguished from "a" by a higher Ni content and additional Co. Mineral "c" is present associated with "a" and "b". It contains Rh, Pd and Pt (however no Ni, Fe or Co), see Fig. 738 and the description there. As already mentioned, these minerals often occur in the peripheral part of Pt nuggets as shown in Fig. 735. Furthermore of particular interest are intergrowths patterns of these minerals and their relationship to fibriolitic asbestos antigorite which is also present in ferroplatin nuggets. Fig. 739 shows a well-developed osmiridium surrounded by a margin of (Os,Ir)S-roseite. There are other cases presented which support the replacement of silicates (fibriolitic antigorite asbestos) by (Os,Ir)S (Figs. 740 and 741). The fact that in ferroplatin nuggets fibriolitic silicates (asbestos) are included, supports a post-alteration growth of the ferroplatin in the sense that antigorite is formed by alteration of dunite. Also, the fact that (Os, Ir)S replaces the fibriolitic silicates is additional evidence for the growth of ferroplatin nuggets and the included platinoid minerals after the alteration of the dunite and the formation of asbestos. On the basis of the textural analysis presented and considering the aforementioned "controversy", accretion of the PGE has taken place which resulted in the formation of auriferous ferroplatin nuggets in which osmiridium is zonally included. In addition, in the peripheral part of the nuggets a number of other platinoid minerals also occur. In this connection it should be mentioned that this is perhaps a major breakthrough, not only in mineralogy and geochemistry of the PGM but also in ore metallography in general, since it demonstrates that minerals previously believed to be formed at very high temperatures can be formed at very low temperatures under special geoenvironmental conditions and by accretion or element agglutination processes. In this connection, the author would like to emphasize that Ramdohr (1960), Genkin (1959), and Stumpf! (1962) have supported the possibility of hydrothermal Pt mineral formation. In addition, recent studies by Styles and Gunn (1991) show that "mineralogical studies of PGM are important to establish the processes associated with their formation, particularly now that the mobility of these elements is widely accepted". It should be emphasized that when Ottemann and Augustithis (1967) explained Pt nuggets in the Yubdo Iaterite as growth formed at Iateritization conditions (nowadays widely accepted), they were called all sorts of "names" by the scientific community. Makovicky et al. (1986) and Wu et al. (1991) demonstrated experimentally the solubility of PGE in hydrothermal solution. Furthermore, now there is no doubt that PGE are soluble under low temperature solution conditions (within the temperature range of Iateritization). 89

(e) Oolitic and Pisolitic Textures and Structures It should be emphasized that in contrast to the manganese nodules and the nuggets of platinum, the oolitic and pisolitic textures observed in bauxites and laterites are of quite different origin. As far as the oolitic textures in karstic bauxites are concerned, in addition to variation in composition of the oolitic layers due to differences in the supply of material, the greatest differences in composition can be produced by leaching, leading to formations of parts of the bauxite (including oolites) different in composition. In 1978 Augustithis et al., in contrast to the above mentioned prevailing views, had already contended that the difference in the oolites (iron-rich layers and iron-depleted layers) is due to differential leaching out of iron, and for that reason, too, portions of the bauxite were depleted of iron. From the plethora of textures and structures of bauxitic and lateritic oolites and pisolites presented by the author in his Atlas of sphaeroidal textures and structures and their significance, (1982), only a few examples will be brought up to present some textures pertinent to leaching out of iron to form oolitic and pisolitic layers of some of these textures in bauxites. Fig. 742 shows bauxite with an iron-rich compact mass and bauxitic oolites differing in their iron content. Iron-rich and iron-poor layers are exhibited in the same oolite while in cases, the central oolite is irondepleted and shows iron-rich outer layers. The reverse is also indicated. In addition, portions of the bauxite (with oolitic structures showing a different degree of layers depleted in iron) are shown as leached patches. The white (iron-poor oolitic layers and iron-depleted patches of the bauxite) are due to leaching out of iron subsequent to the oolite formation (Figs. 742-744). Very interesting oolitic textures may be exhibited where alternating iron-rich and iron-depleted layering is not the result of differences in the supplied material but of differential leaching of the different layers of the bauxitic oolites (Figs. 745-747). There are also instances of complex patterns of oolites where more than one oolitic generation comprises the "oolitic structure". Fig. 748 shows a small oolite enclosed in a later-formed larger oolite. Similarly, Fig. 749 shows three smaller oolites enclosed by a larger layered oolite. In contrast to the oolites, pisolites might develop in bauxites and, in cases, show that they are formed as diffusion ring structures (due to iron leaching and deposition). In some cases the initial oolites are interspersed in the pisolitic structure which actually represents a diffusion ring structure. This subject is discussed more in detail by Augustithis (1982). In contrast to the pisolites with a comparable diffusion ring structure (Figs. 750, 751), cases exist of pisolitic bauxites where the pisolites are gibbsitic in composition and relatively iron-rich in a gibbsitic 90

matrix (Figs. 752, 753). As especially Fig. 752 shows, secondary kaolinite is formed in the pisolite. Apart from the oolites and pisolites described above, cases exist where most complex structures are exhibited. Fig. 754 shows a diffusion sphaeroid (comparable to the diffusion ring structures) in which case though the rhythmical banding exhibited by the diffusion sphaeroid could be attributed to the operation of colloidal solutions. The formation of oolitic structures in karstic bauxites and the formation of pisolitic structures in lateritic bauxites again is treated in detail in the Atlas (Augustithis, 1982). In contrast to the oolitic structures of karstic bauxites, iron oolites in the Fe-Ni-rich laterites (Larymna type, see Augustithis, 1962) represent a restricted in abundance phase in which case oolites with or without a chromite nucleus might be present. Fig. 755 shows oolitic textures in the Larymna laterite, some with a chromite or silicate nucleus, others without nucleus. As indicated in Fig. 755, deposition of iron-rich layers might have precipitated around a chromite fragment. Similarly, Fig. 756 shows a larger chromite nucleus in an iron oolite. In this case though the chromite nucleus is corroded and indented by the later iron oolite and its effects. Augustithis showed that the oolitic phase is restricted in its copiousness in the Larymna ore and evidence is abundant of resorption of the oolites and the formation of a haematite phase, which is subsequently hydrated leaving pisolitic relics in the hydrohaematite matrix (background mass). A consequence of the complex formation processes of pisolitic structures is the pisolite in Fig. 757 which shows a compact haematite mass comprising the pisolite (which in turn has been formed by resorption of the oolitic phase and where relics of the chromite are present in the pisolites). Marginally the pisolite is affected by hydration. Also, the hydrohaematite matrix is formed by hydration of the haematite phase, relics of which are the pisolites. Under reduction conditions, clastic chromite grains with a pisolitic or oolitic iron-rich ore may act as nucleus for the blastic development of magnetite which might partly transverse the oolitic structure and therefore constitutes blastically-grown magnetite epitactically overgrown on corroded chromite (Fig. 758). In this connection though, it should be mentioned that the blastically formed magnetite might be epitactically formed (under lateritization conditions) on the clastically derived chromite nucleous (see also Chapter 10). In connection with the formation of sphaeroidal structures it should be mentioned that phosphorites may show phosphatic ooids which may exhibit different degrees of deferrification, see Fig. 759. In contradistinction, iron ooids may exhibit fine banding due to the difference in the composition of the different layers (see Fig. 760, and Augustithis, 1982).

( f ) Organogenic

Sphaericules

In contrast to the sphaeroidal textures and structures discussed, organogenic sphaeroids and sphaericules should also be considered. Despite these structures not being very abundant or dominant features, nevertheless their significance should not be underestimated, particularly since their role in sulfide mineral building is not fully understood or investigated. Fig. 761 shows a pyrite sphaericule in sphalerite. The sphaericule most likely represents sulfidized bacteria (aggregations of pyrite granules). However, its relationship to the sphalerite is uncertain since in this case, no sphalerite sphaericules are indicated. In contradistinction, Fig. 762 shows sphalerite sphaericules associated wiht chalcopyrite. Whereas the sphalerite sphaericules most probably represent sulfidized bacteria, there is no evidence of such development in the case of associated chalcopyrite. Additional examples where sulfidized bacteria result in pyrite sphaericules associated with copper minerals are shown in Fig. 763, where an aggregate of pyrite sphaericules is partly replaced and enclosed by covellite (CuS). The relationship between sulfidized bacteria resulting in pyrite sphaericules and other sulfides is an open question. Fig. 764 shows pyrite sphaericules representing sulfidized bacteria associated with galena. The question is whether galena is formed by sulfidized bacteria or whether it represents a later replacement of pyrite. Furthermore, cases of pyrite sphaericules associated with sphalerite and gangue are not rare, supporting the view that the bacterial action does not necessarily have a microenvironment as a prerequisite consisting entirely of sulfides (Fig. 765). Dahanayake and Krumbein (1985) endorse on the basis of optical and SEM observations in their contribution "Ultrastructure of a microbial mat-generated phosphorite", that the mineralization is concentrated on ooids and microooids. The coated grains occur within microbial mats. Microbial mats represent the formational environment of the ooids and oncoids. Furthermore, "both the coated grains and mats exhibit similar filamentous micro-organisms". In addition "Rogenpyrit" is considered by Fabricius (1962) to represent colonies of sulfidized bacteria. Rickard (1970) described framboidal pyrite where the sphaeroids are on the average a few μ in diameter and only in exceptions as large as 100 μ and contain microcrystallites smaller than 1 μ in size. The internal structure of the framboids is explained as a function of the formation processes of pyrite and of its surface and crystal chemistry of the microcrysts. The textures are explained as involving the replacement or filling of organic globules and gaseous vacuoles. In contrast, Paunen (1966) explained framboidal pyrite as pyritized remains of organic plants, and Kato (1967) showed that the framboidal sphaericules of pyrite replace diatoms. Tong (1977) described framboidal

aggregates as 1-100 μιη in diameter in several stratabound Cu ores of southern China. In contrast to the organogenic sphaeroid mentioned, Sawlowicz (1990) reported finely dispersed, primary copper sulfide mineralization from the zones of the Cu-S paragenesis in the Zechstein copper-bearing shale from the Fore-Sudetic Monocline, Poland. The mineralization is represented according to him by "small sphaericules and framboids (5 to 10 μπι in size) and framboidal clusters of digenite, chalcocite and minor covellite, dispersed throughout the whole shale but especially abundant in the organic rich laminae". Sawlowicz continues "the textures, chemistry and some different features from pyrite framboids suggest that the copper sulfides in Fe-poor zones of the Kupferschiefer are primary precipitates". Additional cases of non-organogenic sphaeroids are described by Binda et al. (1985) from the proterozoic Siyeh Formation of Alberta, Canada. The ooids are composed of concentric layers of pyrite and chalcopyrite, and covellite is present in several of the ooids. Binda et al. suggest that the ooids grew during early diagenesis of the arenite in which they occur, and the source of metals were black argillaceous sediments above and below the arenite. In contrast to the organogenic (bacterial action) of the framboidal pyrite, Ostwald and England (1977) showed that pyrite framboids and their recrystallizations occur in cryptocrystalline silica veins and amygdales in weathered andesites. The framboids and the chalcedony showed textural evidence of colloidal origin. Taylor (1982) described a sequence of oxidate facies sediments with a cupriferous massive sulfide deposit in the Solomon Islands. A siliceous sinter bed within the sediments contains magnesioferrite sphaeroids having a distinctive framboid texture. Taylor explains the formation of the magnesioferrite framboid texture as due to the coagulation of magnetic iron hydroxide gel particles due to magnetic attraction but facilitated by the presence of a strong electrolyte. Taylor also proposed, by analogy, a simple mechanism to account for framboid pyrite formation. The following quotation presents synoptically the mechanism: "This mechanism requires primary iron sulfide particles to be attracted to one another because of the ferrimagnetic properties of a precursor FeS polymorph or alternatively by Van der Waals forces accentuated by the presence of charged ions in a strong electrolyte. Ordering of resultant microcrysts is mainly a close-packing effect which produces robust aggregates resistant to deformation during subsequent diagenesis". A review of framboidal pyrite is included in the Atlas of the sphaeroidal textures and structures (Augustithis, 1982) and the reader is referred to this review for a more detailed description of framboidal structures and to the special literature quoted there.

91

(g) Volcanogenic-Hydrothermal

Sphaeroids

In contrast to the cases described so far (manganese nodules, Pt nuggets, oolitic and pisolitic structures in bauxites, latentes and phosphites), there are ?hydrothermal "volcanogenic" sphaericules from Mechernich (see also Chapter 7). Fig. 766 shows a hydrothermal complex sphaericule structure consisting of a sphaleritic sphaeroidal nucleus, followed by bravoite (Schachner-Korn, 1960). This sphaericule is overgrown by another sphaericule again comprised of a sphalerite nucleus and followed by a "layer" of bravoite with which the galena is associated having an external layer of sphalerite. Fig. 767 shows a sphaeroid consisting centrally of sphalerite sphaeroids surrounded by bravoite with an external shell of sphalerite. Outside the sphaeroid, galena is associated with sphalerite and bravoite. The above mentioned illustrations suggest that in some of the sphaeroids there is a central "sphaeroidal" structure or there are aggregates of sphalerite followed by a shell of bravoite (with which the galena might be associated) which in turn is followed by an external shell (layer) of sphalerite. However, the outlined structure and sequence of the shell (layers) of these sphaeroids may differ greatly from the cases presented in Figs. 766 and 767. As Fig. 768 shows, a ring structure comprises the sphaeroid and the shell (or layer) is composed of sphalerite and galena. The central part and the surroundings of the sphaeroid consist of gangue. The concise consideration of some selected examples of sphaeroidal structures in ore microscopy shows the wide spectrum of processes that contribute to the formation of these structures and textures. An attempt was made not to exhaust this vast subject but to present case studies illustrating the complexity of problems encountered, in the effort to explain these structures and textures.

(h) Thucholitic

Sphaeroids

Another "sphaeroidal structure" that deserves discussion is the occurrence and origin of thucholite. Fig. 769 shows a sphaeroidal structure of C (recrystallized) and the presence of uraninite, which in turn is transversed by a system of secondary mobilization of C veinform transversing the uraninite. The origin of these sphaeroidal thucholitic structures has been a subject of intensive controversy whether the C is a derivative of bituminous (plant remains) or oil-drops (fluid hydrocarbons) mobilization. Also, the association of C and U deserves discussion. The classical theory of thucholite formation, namely that the thucholites result from the radioactive polymerization of fluid hydrocarbons, is further supported by the investigations of Hoekstra and Fuchs (1960). According to their explanation, thucholite is formed as 92

follows: "Uraninite and related minerals crystallized in pegmatite about 800 m. y. ago, the uranium was oxidized and taken into solution, part of the uranium was later precipitated by oil carried in aqueous solution, precipitation followed by oxidation and solidification of the oil approximately 300 m. y. ago". Similarly, Welin (1965) concluded that the transport of the carbonaceous matter has occurred in the ground water moving along fractures in the Precambrian bedrock and is deposited down to a depth of several hundred meters. This organic material of plant or animal origin primarily has a composition which varied within certain limits but is essentially represented by asphaltite. Welin further supposed that when these aqueous underground solutions passed through a radioactive mineral "some compounds of the carbonaceous matter were reactive enough to solve U, Th and Pb by some unknown chemical reaction". Furthermore, a fraction of the dissolved uranium has again been deposited with the organic compounds constituting thucholite. In contrast to the classical theory of thucholite formation from radioactive polymerization of fluid hydrocarbons, Snyman (1965) showed that thucholite examined under methylene iodide immersion revealed various rounded, elongated or poorly defined bodies which closely resembled boghead structures of coal. Snyman suggested that thucholite represents highly coalified algae and that it compares with coal on the basis of its content of PbO, P 2 0 5 and T h 0 2 . Also analyses for moisture 6 , ash, volatiles, carbon and the measurement of calorific values compare well with coals. Furthermore, uraninite is observed in roundish thucholite bodies which supports an "organogenic origin" of uraninite by crystallization in cavities in individual algae. Jedwab (1966) suggests that thucholites from various locations typically show structures in reflected light similar to those of the humic constituents of coals and lignites. In this connection it should be emphasized that often thucholites have undergone metamorphic processes which might have obliterated any initial organogenic or organogenically-derived structure (see Fig. 769). In fact, the textural pattern exhibited suggests a recrystallization of both the original, precipitated uraninite and the carbonaceous substance, which due to recrystallization, indicates undulating extinction and is ' "graphitized" (see again Fig. 769). In contrast to the two prevailing hypotheses for the origin of thucholites, biostasy and rhexistasy can be of significance for the formation of rocks of organic matter [Erhart (1967), and in particular, Drescher-Kaden (1969)], which might be present in ultrametamorphic and "plutonic" rocks, actually rocks derived by ultrametamorphism and transformation. In contradistinction, Kranz (1967) considers the participation of or6

Alpha particles bombardment from uraninite is consid-

ered by Snyman to account for the subhydrated character of thucholite.

gaiiic components in the transport of ore metals in hydrothermal solution to be of importance. The existence of organic matter in ultramorphic rocks supports the explanation that organic accumulations existed as syng e n e i c material within rocks in which thucholites occur. The association of C with U is geochemically an accepted fact and many instances of U adsorption by C are known. The adsorption of U by organic matter of

originally organic "colloids", plant or animal in derivation, provides an explanation for thucholite formations in rocks of diverse origin. In this case the carbon is not introduced either in form of fluid hydrocarbons or in solution but it was in situ and acted as a "trap" for the concentration of U from solutions which contained U +6 (which is geochemically very mobile).

93

Chapter 14

(a) Deformability

Tectonic Effects

of Ore

Minerals

From research on solid substances subjected to compressional or tensional forces it is known that the behaviour of a body will depend on its limits of elasticity and plasticity prior to irreversible rupture of the continuity of the body. Whereas the limit of elasticity of a body (in this case of ore bodies) is not of great significance overall, the limits of plastic deformation (the point when plastic deformation starts and when rupture takes place) are of great importance when considering the deformation of an ore body. Niggli introduced the term crystal plasticity back in 1929 and worked on the changing of the behaviour of minerals under diverse conditions. Augustithis (1965) showed that the plastic behaviour and in general the deformability of the main rock forming silicates depend on the type of tetrahedral structure they possess, i. e., whether they consist of isolated tetrahedra, tetrahedral arrangement in chains or sheets or tetrahedra building a three-dimensional scaffold. Thus, the deformability of some of the main rock forming minerals is directly dependent on their crystal structure. A study on crystal structure and deformation reveals that under different conditions minerals behave differently, e. g., quartz which is normally brittle, has been shown to be greatly plastic under diverse conditions (see Augustithis, 1965, 1985). The behaviour of ore minerals under deformation (mainly compressional forces) is also dependent on the structure of the ore mineral (the lattice group and in particular, the crystal behaviour, i. e., the plasticity limits that the crystal possesses). Molybdenite (MoAj) which crystallizes in a typical Schichtengitter (sheet lattice), exhibits impressive plastic deformation expressed by the undulating extinction of the deformed mineral and in cases by the deformation of its shape. Fig. 770 shows molybdenite with half-crossed nicols exhibiting undulating extinction. Figs. 770 and 771 also show molybdenite with undulating extinction (with half crossed nicols). Especially in Fig. 771 the undulating extinction and deformation in shape are noticeable that took place. However, when the limits of plasticity of the molybdenite are exceeded, rupture takes place as visible in Figs. 771 and 772. In contrast to molybdenite the behaviour of chromite (FeCr204) which has a chromospinel lattice is very dif94

ferent and very variable under diverse conditions. In this connection it should be mentioned that whereas there is evidence of brittle behaviour of chromite, data about plastic behaviour are missing. However, indications of possible plastic behaviour of chromite are presumed in cases of chromite bodies exhibiting ptygmatic folding and tectonic interbedding with anorthosite (see Augustithis, 1982, 1983). Very impressive patterns of chromite fracturing are exhibited in Figs. 773-775. Fig. 773 shows chromite micromylonitized. In this case the fracturing of the chromite can be considered to represent the end phase (under certain conditions) of plastic deformation of chromite, i. e., rupture occurred. Similarly, Fig. 774 shows tectonically affected chromite exhibiting curved fracture. In addition to the micromylonitization and fracturing of chromite shown in Figs. 773 and 774, fractured chromite is glued together by secondary mobilization of Wiederverkittungs serpentine (Fig. 775). In certain cases second generation idiomorphic chromite crystals in Schlieren chromite, actually formed after Schlieren chromite which is present in the serpentinized mass, are tectonically fractured. Recrystallized antigorite combines the chromite fragments as well as marginally partly surrounds the chromite (Fig. 776). Besides the curved cataclastic pattern in Fig. 774, cases are exhibited of chromite showing a type of oriented crack pattern (Fig. 777) which in cases can result in the formation of tectonogranular chromite due to further fracturing. A very important case of peripheral chromite fracturing leading to a marginal zone of cataclastic tectonogranular chromite is exhibited in Fig. 778. In addition to the tectonogranular chromite which is marginal to a more massive chromite (Fig. 778), cases are shown where massive chromite transgresses to granular chromite due to fracturing (Fig. 779). In the above mentioned cases the fracturing of chromite (cataclasis or tectonogranular chromite formation) occurred while the chromites were in a more or less plastic environment or microenvironment where olivines with great crystalloplasticity or serpentines were the predominant minerals. In contrast, most complex associations can occur where fractured chromite comes with marble (in a melange type, tectonical mo-

bilization of material, see Fig. 780 and Augustithis, 1979). As indicated, there is a transgression of massive chromite to tectonogranular and granular chromite in crystalloplastic environment. In the same environment due to tectonic influences chromite banding of the tectonogranular chromite can take place and sometimes, orientation may be exhibited between the Schlieren bands of chromite formed and the undulating planes of the deformed olivines (Augustithis, 1979). In contrast to the hypothesis that regards the Schlieren chromite as magmatic gravitational differentiation of a crystallizing dunitic magma, Augustithis suggested that the Schlieren bands in Xerolivado, Greece, represent tectonogranular chromite mobilized in bands. As a corollary to this, Fig. 781 shows a Schlieren chromite band indicating tectonic shape rather than a product due to crystal settling from dunitic magma under differentiation. However, more impressive cases of chromite tectonoplastically associated with silicates are exhibited in the case of chromite interbanded with anorthosite at the Dwars River occurrence, Transvaal, South Africa. In contrast to the magmatic differentiation interpretation of the interbanded chromite/anorthosite occurrence of the Dwars River, Augustithis (1982, 1983, 1985) suggested a plastic tectonic mobilization of the chromite and the anorthosite bands. The fact that the chromite bands in the Dwars River exhibit tectonoplastic deformation is supported by the ptygmatic folding of the chromite bands in anorthosite (Fig. 782). Also, whereas Fig. 783 shows the classical case of chromite and anorthosite, its detail Fig. 784 shows that the chromite is tectonically mobilized (see arrow "a", Fig. 784). Most significant cases illustrating tectonoplastic mobilization of chromite and anorthosite are exhibited in Figs. 785-787. Fig. 785 shows lens-shaped anorthosite, tectonically mobilized within a chromite band from the chromite/anorthosite occurrence at the Dwars River. The anorthosite band shows deformation due to tectonoplastic mobilization of chromite and anorthosite. Fig. 787 shows a rather round anorthosite body again included in a band of chromite. In this case, the presence of anorthosite bodies in the chromite bands is incompatible with magmatic gravitational differentiation of chromite and anorthosite since the presence of lensshaped bodies of anorthosite in the middle of a thick chromite band cannot be explained by magmatic differentiation and gravitational settling. As a corollary to the tectonoplastic mobilization of chromite and anorthosite, the pattern is showing the undulating contact of chromite and anorthosite bands (Fig. 786). Resorption of chromite and chromite mobilization along cracks of the anorthosite are also observed (Fig. 788, see also Chapter 14). Comparing the fracturing of chromite, as indicated in Figs. 773-775, and the tectonoplastic mobilization of chromite and anorthosite bands, it is apparent that

chromite which crystallizes in a spinel lattice and hardly exhibits any crystalloplastic behaviour under tectonic deformation and mobilization in the lower crust environment, can behave plastically. As studies of Augustithis in 1965 and 1973 showed, brittle quartz under certain tectonic conditions may behave as a plastically deformable body. In a similar way the chromite is perhaps plastically mobilized together with the anorthosite and of course, a subsequent recrystallization may have occurred. Considering the behaviour of chromites in dunites and particularly in serpentinized dunites, chromite "potatoes" from Domokos, Greece, may be scattered, interspersed in the serpentinized dunitic mass (Fig. 789). In addition to these tectonically rounded chromite bodies, chromite leopard ores indicate that the pattern is due to tectonic mobilization of the more resistant chromite fragments in a more plastic environment consisting of olivines and serpentine (Fig. 790). In some cases the leopard chromite ore may exhibit all transitions from fractured chromite to corroded and rounded chromite bodies (due to serpentinization) resulting in typical leopard chromite ores, a fact which in this case, too, is considered to be the result of tectonic influence, fracturing and subsequent corrosion by the serpentine (or mobilized serpentine). In other cases, leopard chromite ore shows ellipsoid shaped chromite bodies which could represent resistant shapes (in the sense that the strain ellipsoid represents a deformation process). More recent, comparable studies by Leblance et al. (1981) and particularly by Hock and Friedrich (1985) support the impressive tectonic effects and deformation structures that can be observed in chromite when plastic deformation is mainly caused by mantle flow: "The concordant chromite bodies are often tectonically disrupted and boudined forming strings of pods or fault-controlled pocket-like deposits. With increasing tectonization, chromite shows pull-apart textures and lineations (plastic deformation), shearing, prismatic jointing, brecciation and mylonitization (brittle deformation). Recrystallization of cataclastic chromite occurs on a microscopic scale." Taking into consideration the fact that the deformation of an assemblage of ore minerals will depend on the plastic limits of each mineral kind and on the mutual influence which will be exercised by the behaviour of one mineral on another, the pattern produced will be dependent on the intensity of the deformation force and on the mentioned behaviour of the minerals making up the assemblage. In order to clarify the generality of the above statement, some examples will be furnished showing mutual influence exercised by the behaviour of one mineral on another in the case of a typical pyrite, sphalerite and galena paragenetic association. Fig. 791 shows tectonically crushed pyrite and fractured sphalerite in a background of mobilized galena. In the attempt to interpret the patterns it should be pointed 95

out that pyrite is the most fragile mineral, followed by sphalerite to a lesser extent, in contrast to galena which is more plastically mobile (deformable). Thus, the pattern in Fig. 791 which exhibits fractured pyrite, less fractured sphalerite and plastically mobilized galena occupying the spaces between the fractured minerals and the fracture spaces of the pyrite and sphalerite, can perhaps be understood best. Fig. 792 shows fractured pyrite and sphalerite and heavily plastically mobilized galena. As corollary to the great plastic mobility of galena, Fig. 793 shows sphalerite with fractures along which plastically mobilized galena has moved. It should be emphasized that his pattern despite its resemblance to the replacement pattern, represents plastic mobilization of galena in this case and not precipitation of galena into the dissolution spaces of dissolved sphalerite, as it is often the case with replacement of sphalerite by galena (see Chapter 5). Additional to the deformation effects on the sulfides presented, elaborate studies and descriptions of deformation of massive sulfide ores in Central West Greenland are found by Pedersen (1981). Concerning the microstructures observed, he stated the following: "The microstructures of the massive and the porphyroclastic ore tectonics indicate syntectonic recrystallization under high stress and at high strain rates, corresponding to the thrusting of the ore bodies. The microstructures of the mobilized sulfides show evidence of repeated plastic/cataclastic deformation and recrystallization, corresponding to highly variable strain and strain rate conditions during the mobilization." Besides the fragile behaviour of pyrite, cases of other minerals exhibiting patterns of fragmentation will be considered. Fig. 794 shows cobaltite (CoAsS) fracturing resulting in the formation of rounded granular cobaltite. In contradistinction, plastically deformable minerals such as uraninite can show fragmentation under certain tectonic conditions (Fig. 795). Very impressive fragmentation is shown by pitchblende in gangue between two rammelsbergite crystal grains (Fig. 796). In both cases uraninite behaved like a fragile body. In opposition sphaeroidal gel pitchblende may be deformed and in the fractures native Bi might also be plastically mobilized. Studying further the possibilities of tectonic mobilization of a mineral phase into a pre-existing mineral, Fig. 797 shows malacon most probably tectonically pushed along cracks in uraninite. Despite the fact that this interpretation is disputable (see alternative interpretations, Augustithis, 1964), the pattern in Fig. 798 can be interpreted as malacon occupying the fractures of magnetite, illustrating that some of the malacon is of post-magnetite mobilization. Perhaps the best illustrated example of deformation that resulted in "ex-solving" a mineral phase present in "solid solution" or in mobilization of a fraction of the solution formed by friction can be seen in the pattern in Fig. 799 where davidite is tectonically affected and 96

fine rutiles are formed along the line of tectonic influence of the davidite.

(b) Deformation and Ore Mobilization mobilization

or Re-

The influence of tectonics on rock has been investigated and many profound studies are available covering the wide spectrum from global tectonics to microtectonics. As far as microtectonics are concerned, Sander's Gefügekunde (1930, 1956) stands out as a landmark. In contrast only rather limited specialized publications and treatments exist that deal with the Gefügekunde petrofabrics of the ore minerals. Schachner-Korn (1954) made an effort to introduce petrofabric analysis in ore petrography. In addition to their efforts, many ore studies include research related to the deformation of ore minerals. Additional studies have been made concerning the relation of ore deposits and occurrences on global tectonics and on the relationship of tectonic influence and ore body deformation, and even on the relationship of deformation and ore mobilization or remobilization (see Chapter 18). Furthermore, extensive studies were made concerning the metamorphic origin of ore deposits on one hand and the deformation of ore deposits due to dynamic influences on the other hand. In the present section, some selected patterns will be presented related to deformation effects and mobilization and remobilization of ores. A series of patterns (Figs. 800-805) shows tectonic influence and remobilization of ore materials. Fig. 800 shows a fragment of banded pitchblende surrounded by remobilized pitchblende, demonstrating clearly the effects of tectonic influence on fragmentation of an older pitchblende generation and also the effects on the remobilization of U oxides which resulted in the remobilized (second generation) pitchblende. Figs. 801 and 802 show fragments of first generation pitchblende surrounded by remobilized (second generation) pitchblende. Fig. 803 shows a fragment of older pitchblende partly surrounded by remobilized pitchblende which may also have resorbed the first generation. It should be noted that the remobilized second generation of pitchblende transgresses from smooth colloform bands as it surrounds the first generation and exhibits most complex colloform patterns outwards. In contrast to the two generations of pitchblende considered so far, namely the fragmented and the remobilized pitchblende surrounding it, Fig. 804 shows three generations of pitchblende. Generation 1 banded pitchblende - is fragmented due to tectonic influence and is again partly surrounded by colloform banded pitchblende - generation 2 - which is further fragmented as a result of tectonic influence. It is sur-

rounded by remobilized, generation 3, pitchblende which has partly surrounded the fragments of both generation 1 and 2. The third generation forms a thin rim. It should be noted though that often the different generations of pitchblende show differences in composition, as indicated by the dark and light gray colours. As a corollary to the remobilization of pitchblende and the fact that the remobilization might be triggered by tectonic effects, banded gel pitchblende is fragmented and surrounded by a rim of remobilized second generation pitchblende (Fig. 805). It should be pointed out that the remobilized pitchblende extends and occupies a fracture in the first generation of banded pitchblende. In contrast to the remobilization of pitchblende surrounding the older generations, pitchblende is remobilized in some cases and occupies the interleptonic spaces between silicates (Fig. 806). Very complex intergrowths of silicates and pitchblende exist where uranium oxides have infiltrated into the interleptonic spaces of the silicates and occasionally probably replaced them (Fig. 807). In contrast to the aforementioned remobilization of pitchblende, uraninite itself may be the recipient of palingenic cassiterite. As Figs. 808 and 809 show, fractured uraninite is invaded in veinform by remobilized (palingenic) cassiterite. Early cassiterite, most probably hydrothermal, was palingenetically mobilized along tectonic fractures of deformed uraninite. Considering the palingenic remobilization of cassiterite, it should be pointed out that the relationship of tectonic deformation of pitchblende and mobilization as veinlets of the cassiterite is of particular interest. It is most probable that remobilization of casserite is directly influenced by tectonic events which perhaps generated solutions leading to cassiterite mobilization (or remobilization).

(c) Tectonoplastic Mobilization of Chromite Related to Anorthosite of the Bushveld Complex, Dwars River, Transvaal According to most of the published studies, banded chromite anorthosite deposits from Dwars River, Transvaal, have been interpreted as magmatically formed banded anorthosite chromite layers due to magmatic differentiation. However, the author (Augustithis, 1982, 1983) pointed to some structural relation of chromite bodies associated with anorthosite that contradict the hypothesis that the banded deposits represent undisturbed magmatic layering.

In this section only a few structural patterns will be presented which show that tectonoplastic deformation has played a significant role in the genesis and present structural state of these chromite/anorthosite bodies. Fig. 782 shows fine chromite ptygmatically folded into the anorthosite supporting an intense tectonic deformation of the chromite in the anorthosite. In addition, Fig. 810 shows initial banded chromite disrupted and mobilized in the anorthosite. In contrast to the hypothesis that chromite/anorthosite banding occurred under magmatic layering conditions, banded chromite with gaps in the layering is shown in Fig. 811 and especially in Fig. 812 the lower band is shown as continuing undisturbed while there is an interruption in the layering of one chromite band. However, the most meaningful evidence of tectonoplastic mobilization of chromite bands in the anorthosite is presented in a series of patterns in Chapter 14 ("Deformation and Ore Mobilization or Remobilization"). In this connection, Fig. 785 shows lens-shaped anorthosite bodies in the chromite band and a tectonic deformation of the anorthosite bands associated with the chromite bands. Considering that due to gravitational settling of the chromite and anorthosite components under magmatic differentiation, anorthosite lensshaped bodies cannot be formed in the chromite bands, it is therefore possible that the tectonoplastic mobilization of the chromite in relation to the anorthosite possibly provides a more satisfactory interpretation. Fig. 788 shows chromite bodies often partly resorbed in the anorthosite bands of the Dwars River occurrence. Also remobilized chromite occupies fractures of the anorthosite. Similarly, Fig. 813 shows chromite occupying a fine fracture of the anorthosite and as the detailed figure (Fig. 814) shows, the chromite fracture filling consists of granular chromite crystals. It should be pointed out (again Fig. 814), that no pronounced tectonic fracturing or cataclasis is noticeable in the chromite which is very uncommon in tectonically deformed chromites (see Chapter 14). The en masse mutual mobilization of the anorthosite lenses in the chromite bands and the deformation of the chromite bands themselves most probably took place under tectonoplastic high pressure conditions in the lower crust geoenvironment. Under comparable conditions, the remobilization of the chromite occurred along fractures as well as the resorption and mobilization of chromite bodies in the anorthosite as shown in Fig. 788.

97

Chapter 15

Weathering and Alteration of Ore Minerals

Attempting to differentiate between weathering and alteration of minerals it is not unusual for difficulties to be encountered when trying to introduce definitions. Of course it is very hard to draw the demarcation line between weathering and alteration since alteration includes most of the weathering processes (in particular chemical weathering which is actually the main topic here). On the other hand, if weathering is defined as surface alteration process, it is difficult to define the depth at which these surface processes take place since the percolation of surface water and its alteration effects are difficult to limit. In certain rock types the alteration of the rock (depending on the penetrability of rpcks to surface derived water solutions) may exceed 500 metres. Considering though that surface water might ultimately be the source of all water in the earth's crust, including (OH) in minerals, the influence of water as an alteration agent within the crust is enormous. Vemadsky's phreatic cycles and the significance of water solutions in the evolution of the earth's crust should be considered (pers. comm. with Grigoijew). Considering the above mentioned difficulties, a common presentation of the subject of weathering and alteration of the ore minerals will be attempted by bringing forward case studies and defining the environment as much as possible in which the processes took place. Examples will also be presented showing that the minerals most resistant to weathering, such as the chromites, may be subjected to weathering alteration processes and in cases, the phenomenology of surface weathering of chromite sand grains and of chromite altered due to serpentinization may be comparable and commensurable. Fig. 815 shows a chromite sand grain with a margin of decoloration and with solution (dissolution) channels extending into the outer part of the chromite grain. Comparable decoloration margins of chromite were found in chromites that have been altered due to serpentinization (Augustithis, 1960). They show that chromite decoloration is due to element leaching. Myrmekitic serpentine has intergrown with the chromite particularly in preference with its decoloration margins. Thus, the decoloration margins of the chromite sand grain and of the altered chromites by serpentinization processes are comparable and commensurable. Microprobe analytical studies of chro98

mites and their decoloration margins (Panagos and Ottemann, 1966) showed that the chromite cores are relatively rich in Mg and Al compared to the decoloration margins which in turn are richer in Fe and Cr. (Panagos and Ottemann, however, explained that the chromite decoloration margin occurred either with a late magmatic phase or as penecontemporaneous with the serpentinization.) Further studies on chromite grains in birbirites by Augustithis (1965, 1967) showed that chromites exhibit decoloration margins and dissolution channels of the chromite in an environment of dunite birbirization (leaching of Mg and recrystallization of chalcedony from the residuals of alteration of olivines, see Augustithis, 1965, 1967, 1979). Fig. 816 shows chromite in the birbirite of Yubdo, W. Ethiopia, with decoloration margins and dissolution channels. A review of the decoloration margins of chromite on the basis of the studies mentioned and on additional observations suggests that chromite sand grains (rounded by attrition, see Fig. 815) show decoloration and dissolution channels restricted to decoloration margins. In this case, there is no doubt that the decoloration of a chromite grain and its dissolution channels have been produced by surface weathering processes. Fig. 817 shows chromite from Rodiani, Greece, with decoloration margins and with myrmekitic serpentine preferentially found· in these margins. Augustithis (1960) showed that marginal serpentine extends into the chromite resulting in the intergrowth of serpentine/chromite. As mentioned, studies of Panagos and Ottemann revealed that the core (central chromite free from myrmekitic serpentine) is relatively richer in Mg and Al compared to the decoloration margins which in tum is richer in Fe and Cr. In contrast to late magmatic formation of the decoloration zone, the decoloration of the chromite is the result of serpentinization and leaching of elements from the chromite. The intimate relation of myrmekitic serpentine and its preferential intergrowth with the decoloration margin of the chromite are in opposition to the late magmatic formation of the decoloration margins of the Rodiani chromite and in favour of the formation of the decoloration of the chromite by leaching of elements during serpentinization. Thus, the fact that the decoloration margins of chromite are formed in chromite sand grains (Fig. 815)

produced by surface weathering and the fact that decoloration margins and alteration effects are also found in chromites in birbirites (Fig. 818) support the view that chromites with decoloration margins (Fig. 817) in serpentines from Rodiani, Greece, are not the result of a late magmatic effect but due to element leaching during serpentinization. It should be noted though that chromite xenocrysts caught in basaltic melts show alteration margins of magnetite with which, due to magmatic corrosion and infiltration into the margin of magnetite and chromite, myrmekitic symplectites are produced by the basaltic melts infiltrated into the magnetite/chromite (Figs. 486 and 487). In contrast to the decoloration margins of the chromite, chromite grains show darker coloured margins which extend also into the chromite along cracks (Augustithis, 1962). Microprobe analytical studies of the chromite and its darker margins by Augustithis and Mposkos (1980) showed that a relative decrease in Fe and an increase in Cr took place in the dark alteration margins of clastic chromite grains of the Ni-Fe laterites of Larymna, Locris, Greece (see Figs. 819 and 820). It should be mentioned that the darker alteration margins are formed in chromite grains in an iron-rich microenvironment. As Fig. 819 shows, clastic chromite grains act as a nucleous of a limonitic oolite. The microprobe studies revealed a relative decrease in Fe and an increase in Cr when comparing the chromite grain and its darker margin. Similar dark alteration margins in clastic chromite grains are shown in Fig. 820. In general, the dark coloration margins which are characterized by a decrease in Fe and an increase in Cr compared to the unaltered chromite core indicate that in an Fe-rich microenvironment (limonite, haematite, hydrohaematite) leaching out of iron from the chromites took place. In contrast to the above mentioned cases, extensive microprobe analytical work on the decoloration margins of chromite in the birbirite of Yubdo, Ethiopia, by Augustithis and Mposkos (1980) showed that these margins of chromite exhibit a relative increase in Fe and a decrease in Cr compared to the unaltered chromite (Fig. 816). Comparing the microenvironments of chromites with decoloration margins and of chromites with dark alteration margins, the following fundamental differences are apparent: (i) The chromite in a serpentine environment as a product of Mg leaching out of dunitic olivines showed decoloration alteration margins (Fig. 817). Similarly, the chromite in the birbirite (Fig. 816) showed decoloration chromite margins. In this case, too, a strong leaching out of Mg occurred in the birbiritization of the dunite. As the comparative microprobe analytical work of Augustithis and Mposkos (1980) showed, there was a relative increase in Fe and a decrease in Cr in the decoloration margins of chromite in the birbirite as mentioned. Thus, the microenvironment of Mg leach-

ing out Cr has been leached out of chromite. In this connection, the geochemical relationship of Cr and Mg should not be disregarded. (ii) As already stated, clastic chromite grains in a Ferich microenvironment surrounded by limonite, haematite, hydrohaematite, reveals dark alteration margins in which the content of Cr is relatively increased and in Fe decreased in comparison to the unaltered chromite. Thus, the dark alteration margins are caused by Fe being leached out of the chromite grains which occur in an Fe-rich microenvironment. As a contrast to the decoloration margins discussed, cases of decoloration of chromite closely associated with cracks or cataclastic cracks of it are often observed. Fig. 821 shows chromite with a crack marginal to which alteration has taken place (decoloration accompanied by an increase in the internal reflection of the chromite). As Fig. 822 shows, dark bodies often associated with this type of alteration are formed representing probably Mn or Mg concentrations. In the case of clastic chromite in the Ni-Fe lateritic deposits of Larymna, Greece, chromite alteration is often pronounced resulting in the dissolution of chromospinel and possibly in the formation of Cr oxides. Figs. 823 and 824 show chromites greatly corroded and to a large extent dissolved. Relics of them are closely associated with a dark mass representing perhaps Cr oxides formed by the alteration of the chromite. Occasionally dissolution of the chromite is followed by replacement with haematite (Figs. 825 and 826). Chromite, however, can also be greatly altered when attacked by talc, as in the case of "hydrothermal" alteration of chromite by talc. Fig. 827 shows a chromite grain with an alteration margin due to "hydrothermal metasomatic solutions" related to talc formation within serpentinized dunitic bodies. Very impressive chromite alterations occur when chromite is affected by hydrothermal metasomatic solutions as described by Augustithis (1960) - now it is known that these metasomatic alterations are related to rhondigite formation - (Augustithis, 1979). Fig. 828 shows unaffected chromite (a) passing into chromite alteration stage I (b), while alteration stage I is in contact with alteration stage II (c). In addition, mobilized enclaves consisting of ilmenite, rutiles, anatases, ?perwoskite and carbonates (d) are within the alteration chromite stage I. Furthermore, Fig. 829 shows unaltered chromite (a) in contact with chromite alteration stage I (b). The enclaves consisting of metasomatically mobilized substances of rutile, ilmenite and anatase, are indicated by (d). In addition, Fig. 830 shows chromite alteration stage I (b) transgressing to chromite alteration stage II (c). Enclaves consisting of rutiles, ilmenites and carbonates are also shown (d), see also Fig. 831. In the alteration stage II (c) dark lamellar bodies possibly represent "relics" of the transformation of chromite into stage II (c). In contrast, Fig. 832 shows an enclave in alteration stage I (b) which con99

sists of rutiles, ilmenites and carbonates (see also Fig. 833 and the description there). On the basis of the decoloration margins (Figs. 815817), alteration stages I (b) and alteration stage II (c) indicate changes due to relative increase in Cr leaching which approach a magnetite composition (advanced alteration stage II (c)). These alterations of chromite where Cr leaching took place related to rhondigite formation and to the subsequent formation by metasomatic mobilizations of rutile, perwoskite and ilmenite, are to be seen in one polished section. As a corollary to these processes is the presence of rutile oriented lamellae in the unaltered chromite (see Chapter 6). Magnetite (Fe 3 0 4 ) shows a wide spectrum of alteration patterns mainly due to oxidation and occasionally to dissolutions. Fig. 834 shows magnetite with oxidation alteration to haematite where the alteration follows "solution paths" within the magnetite. Additional alteration effects on magnetite are shown in Fig. 835 with marginal and alteration oxidation following a crack pattern in the magnetite. These alterations are mainly due to oxidation and mostly haematite is formed. In contradistinction to the oxidation alteration of magnetite into haematite which follows margins and solution paths, more oriented alteration patterns result due again to the oxidation of magnetite. In this case maghemite follows mainly octahedral penetrability directions within the magnetite, sometimes initiating from cracks within it (Fig. 835). The oxidation alteration of magnetite most often takes the form of alteration lamellae of haematite following the octahedral directions of the magnetite. In contrast to the super-heated heat martitization (see Chapter 7), magnetite due to oxidation alteration, changes to martite haematite (lamellar haematite) following mainly the [111] direction of the magnetite. Fig. 836 shows martitized magnetite with fine martite lamellae following the [111] direction of the magnetite. However, a noticeable alteration margin of the magnetite is also exhibited indicating that - despite the fact that the oxidation alteration of the magnetite follows the [111] direction of the host - there is a more intense alteration of the magnetite into the haematite marginally. As a corollary to the marginal martitization of magnetite, magnetite is shown with marginal martite (Fig. 837), in which case the martite actually forms a rim surrounding the magnetite which does not follow its [111] direction. In other cases, the martitization is not strictly confined to the [111] direction of the host magnetite and results in the formation of more irregularly shaped haematite bodies than the typical martite lamellae (Fig. 838). However, martitization of magnetite, despite being more intense in cases at the margin of the magnetite, usually follows the [111] direction of the host (Fig. 839). 100

Additionally Fig. 840 shows martite extending into the magnetite from its margins and following the [111] direction of the host magnetite. In cases the martites result in partial alteration of the magnetite into the haematite. A very interesting pattern of magnetite alteration due to martitization is produced and the central unaltered magnetite is delimited by haematite lamellae following the [111] direction of the magnetite. Cases of martitization and simultaneous maghemitization of the magnetite are also common (Fig. 841). Instances also exist where clastic magnetite grains may show zonal martitization resulting in a broad zone of haematite replacing the magnetite (Fig. 842). Ramdohr, too, showed excellent patterns of zonal martitization of magnetite. Since most minerals can be subjected to weathering processes, especially to chemical weathering, the presentation of weathering patterns is an enormous subject which is far beyond the scope of the present Atlas. However, some selected examples will be presented with the aim of showing the type of processes and textures that can be produced. Ferberite (FeW0 4 ) when subjected to alteration due to chemical weathering (percolation of surface water), changes to haematite and often Fe hydroxides are also formed. The alteration of ferberite to haematite involves the leaching out of W and the re-adjustment of the lattice as well as volume changes. Fig. 843 shows ferberite partly changed to haematite. Another mineral most susceptible to alteration processes is pyrrhotite (FeS) which is often changed into marcasite (FeS2) and as a result the well-known bird's eye structure is produced (Fig. 844). The transition of pyrrhotite to marcasite is often characterized by the formation of the Zwischenprodukt (intermediate product). A series of figures illustrates the patterns created during the transformation of pyrrhotite to Zwischenprodukt. Fig. 845 shows pyrrhotite altered to Zwischenprodukt. Fig. 846 shows pyrrhotite relics in a mass of Zwischenprodukt. In contrast to the typical patterns of pyrrhotite alteration to Zwischenprodukt and finally to marcasite, cases are shown where initial pyrrhotite is altered by advancing solution fronts resulting in the formation of "secondary colloform patterns", mainly consisting of fine marcasite. In the marginal area of the colloform pattern, crystalloblastic pyrite is formed (Fig. 847). A rather impressive pattern of pyrrhotite alteration is also the formation of the "box work" structure, consisting of lepidocrocite and haematite. This is believed to be created due to S leaching and subsequent oxidation of Fe (Fig. 848). Comparable to the alteration of pyrrhotite to marcasite (bird's eye structure) are colloform sphaeroids of marcasite formed in pyrrhotite by advancing solution fronts resulting in secondary marcasite sphaeroids in the pyrrhotite or at its margins (Figs. 849 and 850). Chalcopyrite (CuFeS2) when affected by a front of advancing alteration solutions, may also show secon-

darily produced rhythmical colloform banded structure, consisting of alternating bands of brown iron (limonite) and needle lepidocrocite (Fig. 851). Sometimes complex element mobilization processes are related to the alteration of minerals, e. g., in the case of alteration of hausmannite (Mn 3 0 4 ) to psilomelane [BaMn2Mn64Oie(OH)4] whereas the comparison of the chemical compositions shows that the derivation of psilomelane from hausmannite by weathering involves additional elements to those forming the hausmannite (Fig. 852 shows the alteration of hausmannite to psilomelane). In other cases, the alteration-weathering processes may result in the formation of perfect pseudomorphs, e. g., pyrite replaced by limonite. Fig. 853 shows siderite (FeC0 3 ) partly dissolved and partly replaced by a framework (skeleton) of limonite. This is in contrast to cases where relics of the initial siderite are preserved. Cases are also exhibited with all the siderite dissolved and a framework of limonite formed whereby the crystal outline of the siderite is preserved. Considering that brown iron [limonite, goethite (HFe0 2 )] is a more stable form under weathering conditions where oxidation-hydration may prevail, it is not surprising that alteration minerals such as maghemite γ -(Fe 2 0 3 ) may in tum be further altered to brown iron (limonite-geothite). Fig. 854 shows maghemite replaced due to alteration-weathering processes (oxidation-hydration) to brown iron. The alteration of the maghemite took place marginally and along cracks. In contrast to the formation of brown iron as a result of weathering of Fe-containing minerals as already discussed, cases are presented where brown iron is formed by alteration replacement of minerals that do not include iron in their composition. Fig. 855 shows native Bi replaced marginally by brown iron. Another example of a mineral replacement is cassiterite (Sn02), in this case by haematite and brown iron. The composition of cassiterite does not contain Fe which could be liberated during weathering and contribute to the formation of haematite and brown iron. In contrast to the cases of alteration of minerals which do not contain Fe in their composition, impressive alteration patterns result when Fe-rich minerals are subjected to hydration oxidation processes. Figs. 856 and 857 show lievrite [CaFe2Fe3(OH)Si2Og] surrounded, corroded and invaded by brown iron. In cases colloform structure is exhibited and lievrite replaced, resulting in the formation of an anastomosing veinform pattern (Fig. 857). As mentioned, maghemite is formed as an alteration mineral of magnetite (Fig. 835). In contradistinction to the maghemite replacing the magnetite and associated with martite, cases of maghemite marginally replacing magnetite are also found (Fig. 858). Fig. 859 shows maghemite replacing marginally magnetite and pyrite. The alteration patterns presented show that due to oxidation alteration haematite and maghemite may be formed from a number of Fe-containing minerals. By

alteration weathering processes brown iron (limonite, goethite) may replace not only minerals containing iron (such as magnetite, haematite, siderite, pyrrhotite, pyrite, chalcopyrite, lievrite, maghemite) but also minerals that do not contain iron such as Bi and cassiterite. In such cases though, the iron mobilization is from a source outside the altered mineral. Occasionally weathering may result in the marginal dissolution of the minerals affected. Fig. 860 shows sphalerite corroded and indented due to alterationweathering. It should be pointed out that in addition to the indentation of the sphalerite a thin margin of chalcopyrite was also formed, associated with the corroded and marginally affected sphalerite. The Cu was most probably in the sphalerite and it was freed due to its dissolution. Subsequently it was mobilized and it formed chalcopyrite with Fe and S made available by the marginal disintegration of the sphalerite. For the consideration of the weathering alteration of minerals as mentioned only some case studies are presented from the plethora of possibilities. In 1964 Augustithis had recognized that detailed analytical studies were necessary. Attempts will be made to compare at microscale the composition of an unaltered mineral with the analytical results obtained from an altered product or, in some cases, from altered zones. Fig. 861 shows a diagram of a sample of uraninite with alteration zones from Gordonia, S. Africa. The circles correspond with the areas of the polished section analyzed by X-ray fluorescence spectroscopy. Fig. 862 also shows the contact between areas 1 and 2 of the diagram. Fig. 861 and similarly Fig. 863 show the contact of areas 3 and 4 in Fig. 861. Quoting Augustithis (Augustithis and Ottemann, 1964) the following data are the results of the analyzed unaltered uraninite and its alteration margins. Also quoted are the conclusions concerning element leaching due to the alteration of the uraninite from Gordonia. Table la Estimated ratios in the unaltered uraninite

Corresponding ratios in the black margin

U : Th U :: Pb U :Υ Th: Pb

U : : Th U : Pb U: Υ U: Pb

= 10 : 1 = 10: 700° C) is inconsistent with thermochemical considerations. The primary process leading to the formation of magmatic iron ores is most likely the enrichment of iron and oxygen under surface conditions in the presence of water and, directly or indirectly, in communication with atmospheric oxygen.

113

Chapter 20

Metallogeny Related to Ultrabasics

In contrast to the cases described where chromite ore bodies are considered to be products of magmatic differentiation and where, by fractional crystallization, crystals settling occurred, examples are presented where the presence and distribution of chromite within the host rocks or in "recipient" rock (rock unrelated to the genesis of chromite) is due to tectonic mobilization or remobilization. Tectonic mobilization or remobilization is considered responsible for the formation of a number of chromite ore bodies (Schlieren, boutinage in dunite/serpentinites, lens-shaped bodies and banding with anorthosites; see Augustithis, 1986). Chakraborty (1958) described chromite lenses and veins along sheer planes within dunites and peridotites, indicating a structural control in their emplacement. According to Chakraborty, these chromite ore bodies are attributed to "residual liquid injection". Mukheijee (1962) suggests that the chromite ore bodies of Nausahi, Keonjhar district, Orissa, are early magmatic, the chromite accumulation being carried up as autoliths during emplacement of ultramafic host rocks. Furthermore, Watts (1963) in his studies of chromites from Iran states: "that columnar or tabular orebodies have, in many areas, been produced by forceful intrusion of chromite after consolidation of their host rocks". Also according to Watts, the chromite is associated with picotite and less commonly with chromitediopside, chromite-bronzite, uvarovite, kämmererite and fuchsite. Another case of chromite mechanically mobilized is described to originate from Pini, Minas Gerais, Brazil, where chromite in serpentinites is considered to be the result of mechanical injection and substitution in the schist. Filter pressing of magmatically differentiated chromite liquid during a diastrophic phase is the proposed mechanism. Petrascheck (1966) described Schlieren chromite in serpentinite of Ahito Mountain, Togo, which was converted into steep lenses during post-tectonic processes culminating in the metamorphism which determined the form, quality and quantity of the ore. Thayer (1969) in his paper entitled "Gravity differentiation and magmatic re-emplacement of podiform chromite deposits" points out that "to produce the thickness of massive chromite and volumes of dunite and peridotite found in alpine complexes requires very large volumes of magma, or primary magma more ma114

fic than tholeiite, and differentiation trends very different from those of known stratiform deposits". In contrast to Thayer, it is proposed that massive chromites have been formed in the mantle and due to the crystalloplastic behaviour of dunites, are tectonically mobilized during the emplacement of mantle. It is therefore not surprising that relatively small dunitic bodies contain large volumes of chromite, or large massive chromite lenses, e. g., Rodiani, Vourinos, northern Greece (see Augustithis, 1979). In contrast to the chromite bodies and occurrences described, a small lens of chromite ore occurs in hornblendic schists and gneisses near Wankur, Khammam district, Andra Pradesh, India. According to Sharma (1960) quartz occurs with the chromite and some of the marginal chromite grains are shattered. It is suggested that the lens is a metamorphosed fossil placer. Augustithis (1979) described chromite fractured and mobilized with marbles in Drepanon, northern Greece, and also cases of chromite bodies associated with leuchtenbergite (with anthophyllite blasts) in Hadabudussa, Adola district, southern Ethiopia. Metamorphism must have played a role in both cases. Augustithis also emphasized the significance of element geochemistry when explaining the relationship of dunite with Cr and the abundance of Fe oxides (magnetites) with pyroxenites. He stated: "Another important factor for understanding the paragenetic association of chromite dunite, i. e., the association of chrome-spinels with forsterite-rich portions of the mantle could be broadly understood by considering the geochemical interrelations of Mg in the olivines and Ni-Cr (i. e., size of atomic radii Mg +2 = 0.78 Ä and Ni+2 = 0.78 Ä - Ni and Cr are interrelated in accordance with the periodic table). It is therefore probable that the paragenetic association of the different types of spinels with the ultrabasics depends, among other factors, on MgO/FeO ratios". He further suggested that "considering the paragenetic concept: spineltype/ultrabasic-type, it is clear that a distinguishing of spinel paragenesis is possible, namely that chromespinels and chromites in the forsterite-rich dunites and magnetite-titanospinels prevailing in ultrabasic and basic rocks where the FeO is above a certain marginal value in the MgO/FeO ratios". These rather empirically proposed trends of preferential association of Cr (and to some extent Ni) in pri-

mary distribution with the ultrabasics in which the MgO is prevailing, and magnetite spinel, Al- and Tispinels in the wide range of ultrabasics in which FeO is above a certain value range, can provide a tentative explanation for the preference of Ti in pyroxenites and gabbros rather than in dunites. The distribution of Ti in basic and ultrabasic rocks is also of interest, particularly for the attempt to explain the virtual absence of titanomagnetites, Ti-spinels in dunites which, by contrast, are very abundant in gabbroic rocks. However, cases are described by Augustithis (1960, 1979) where Ti (rutile) occurs as "ex-solutions" and in grain-form with chromite grains of Rodiani, Greece. The presence of Ti and Cr together in the ultrabasics of Rodiani is of geochemical significance. The titanium and chrome are related as subgroup elements of the periodic system. They belong to the same horizontal row (Sc, Ti, Cr, Μη, Fe, Co, Ni), although, as it was stressed, Ti is not abundant in the dunite (nor in dunitic mantle) and Ti-spinels are virtually absent from the dunitic types. As pointed out, it was suggested for the elements Ti, V and Cr to exist an "antipathy" (for Ti and Cr), despite the fact that they are related as next to each other in the periodic table. However, study cases will be presented from the international literature elaborating further on the relationship of Fe, Ti, Cr, V and Mn. According to Lister (1966), "magnetite and ilmenite from concordant and discordant Fe-Ti oxide bodies in mafic, anorthositic and syenitic plutons were analyzed for Fe, Ti, Cr, Al, Mg and Μη. Partition of minor elements between magnetite-ilmenite pairs is demonstrated. The preference of V, Cr and Al for magnetite indicates that these elements substitute for Fe 3 + , while the preference of Mg and Mn for ilmenite may be due to more ionic bonding in this mineral. Magnetite formed early in differentiation contains more Cr, V, etc., than magnetite formed later, while ilmenite formed at an early stage is richer in Mg and poorer in Mn that formed later". Elaborating further on the relationship between Fe and V, Willemse (1969) in Bushveld, describes vanadiferous iron ore occurring as plug-like pegmatoid bodies and as seams of magnetite (1.6% V 2 O s on the average is given for the magnetite seams). On the other hand, Ti minerals in the ore include small ilmenite granules, rather rare broad lamellae of ilmenite which ex-solved from magnetite at an early stage, "ulvite" (ulvospinel), and a wide range of intergrowths of dispersed ilmenite ("proto-ilmenite") with magnetite, maghemite or martite. In addition to the antipathetic relationship between Ti and Cr, as already described, Willemse suggests that there is a similar relationship between V 2 0 5 and T i 0 2 of the ore (in Bushveld): the lowermost seam has about 2% V 2 O s and 14% T i 0 2 , the uppermost about 0.3% V 2 0 5 and 18-20% T i 0 2 . As the Ti is mostly in ulvospinel, it is considered that a decrease in oxygen fu-

gacity in the magma determined the increase in Ti content upwards in the layered sequence. In contrast, recent studies of vanadiferous magnetite plugs from Bushveld (see Figs. 907 and 908) by Augustithis show ilmenite and spinel lamellae in the magnetite and the presence of V in the magnetite is perhaps understandable due to the interrelationship of Ti-V in accordance with the empirical laws of the periodic system. Another case presenting the relationship between Fe and Ti is put forward by Fominykh (1963). He reports of the largest ore deposit of the Urals which consists of vanadium-rich titanomagnetite ore localized in gabbros and pyroxenites of the Kochkaner massif. The Ti/Fe relationship is demonstrated on the basis of analytical results which show an increase in T i 0 2 from 0.31% in "barren pyroxenites" to 3.90% in the massive titanomagnetite ores. Concurrently, the Ti/Fe ratio increases from 0.026 to 0.050. In addition, Van Rensburg (1965) described sulphides occurring disseminated through pipe-like bodies of magnetic iron ore occurring above and below the Merensky Reef horizon in the Lydenburg district. The pipe-like bodies occur in the cores of ultramafic pegmatoid bodies surrounded by hortonolite-dunite and diallage pegmatoid. The disseminated sulphides are pyrrhotite, pentlandite, chalcopyrite, cubanite, mackinawite and valleriite, a suite typical for the Merensky Reef. Panagos and Ramdohr described valleriite occurring as a hydrothermal phase paragenetically associated with pyrrhotite, chalcopyrite, pyrite in an ultrabasic containing chromite, garnet (uvarovite) magnetite, olivine and alteration minerals of the ultrabasic. This is a case of superimposed mineral paragenesis on the chromite, magnetite olivine rock of the ultrabasic. A comparable case of chromite grains enclosed in an uvarovite crystalloblast and later replaced by chalcopyrite and sphalerite-chalcopyrite was presented in Chapter 8 together with relevant illustrations. In contradistinction to the cases of superimposed parageneses, Desborough (1966) proposes two alternative interpretations on the basis of the low Zn/Cu and Zn/Ni ratios (and the presence of accessory sphalerite with pyrrhotite and chalcopyrite) to explain the metallogenesis of the basic Sudbury-type intrusion: (i) high temperature sulphurization of a basic magma effecting the extraction of Ni, Fe and Cu from solid silicates thus explaining the low Zn content of the ores (although resulting in the generation of large quantities of magnetite which are normally not present); (ii) dissolution and mobilization of disseminated Ni-Cu-Fe sulphides at 450-750° C, without mobilizing sphalerite or involving the production of large volumes of magnetite. Desborough suggested the feasibility of the operation of both processes. In this case, it is also open to argue whether the Ni, Fe and Cu sulphides are a superimposed paragenesis (the source though of the material 115

might be through remobilization from the complex and its geoenvironment). Woodall and Travis (1970) described the occurrence of nickel deposits in Archaean mafic and ultramafic rocks in western Australia (Kambalda) in the "greenstone" belt. The primary ore, occurring as massive and disseminated mineralization, consists of pyrrhotitepentlandite assemblage with subordinate pyrite (cobaltiferous) and minor chalcopyrite. The ratio given for Ni/Cu is about 10:1. Another case where sulphide mineralization is associated with basic and ultrabasic rocks, is the Thompson-Moak Lake area, Manitoba. The nickel deposits occur in serpentine. The Thompson ore body occurs as an irregular sheet within biotite schist (a serpentine lens lies within the schist). According to Patterson (1963), pyrrhotite and pentlandite are the main sulphide minerals of the ore, chalcopyrite is widespread and gersdorffite is present locally. The ore body contains Co and the PGE. The large, low-grade ore body at Moak Lake is a serpentinized intrusion containing disseminated sulphide minerals, mainly pyrrhotite and pentlandite and some minor massive ores in zones of fracturing. Also, the low-grade nickel deposit at Mystery Lake consists of sulphides, mainly pyrrhotite and pentlandite, disseminated through serpentine. Mo flakes also occur in some places. The metallic elements present in these ultrabasic bodies will be considered on the basis of subgroup elements of the periodic system as well as their relationship in Chapter 49. Eliseev et al. (1961) presented an extensive study on the Cu-Ni ore of Pechenga (Finnish: Petsamo). The Cu-Ni sulphide ores are associated with ultrabasic intrusive sills including tuffaceous, sedimentary and lava series of the Upper Proterozoic age. The richest ores are found at the lower contact of the differentiated serpentine-alkali gabbro complexes, being found disseminated in the igneous rock as well as in the underlayering tectonic breccia and other rocks. Over a hundred such differentiated sills are recorded. According to the evidence and interpretations by Eliseev et al., the sulphide mineralization occurred during the final magmatic phase of autometasomatism and the superimposed hydrothermal phase. The principle minerals are pentlandite, pyrrhotite, chalcopyrite, pyrite, ilmenite and magnetite (in the oxidation zone covellite, violarite, malachite, azurite and others occur). In contrast to the cases reported by Panagos and Ramdohr (1965), where chromite was formed at the ultrabasic rock stage of formation, uvarovite probably represents a metamorphic phase and the sulphides a later, superimposed paragenesis (or later hydrothermal phase). Eliseev et al. (1961) also suggested a probably superimposed sulphide phase (due to ? autometasomatism) in the ultrabasic sills of Petsamo. Grenier (1955) describing the northern zone of the ultrabasic mineralization at the Appalachians of southeastern Quebec, gives a complex mineral assemblage. The northern zone contains both massive and dissemi116

nated sulphides, millerite, violarite, nickelferous pyrite (? bravoite), sphalerite, chalcopyrite and chromite. In this case too, it is possible that the chromite belongs to an ultrabasic stage (phase of crystallization) and the sulphides might represent a superimposed phase or paragenesis. Thalhammer et al. (1986) support that the mineral association at Pevkos, Cyprus, consisting of pyrrhotite, pentlandite, maucherite, chalcopyrite, cubanite, magnetite and valleriite with minor amounts of westerveldite, bomite, neodigenite, covellite and valleriite, and the mineralization at Lakxiatou Maurou, Cyprus, consisting of pyrrhotite, pentlandite, löllingite, chalcopyrite, cubanite and chromite with traces of magnetite, pyrite, maucherite and valleriite, represent complex paragenetic associations. They further support a non-magmatic origin for the sulphides and arsenides, which were deposited during serpentinization. Also in these cases, chromite and perhaps magnetite belongs to the ultrabasic phase of formation (magmatic or mantle diapirism) and the sulphides and arsenides represent a "superimposed" paragenetic association formed due to serpentinization (? autometasomatism). In addition to the cases discussed, where the PGM are associated with ultrabasic rocks and are paragenetically related to chromite or chromospinels, cases of paragenetic association of PGM and sulphides are common. One of the most significant occurrences of PGM is the Merensky Reef of the Bushveld complex, which, as mentioned, is a differentiated layered complex according to the orthodox views. According to Cousins (1969), the Merensky Reef is a regular layer of pyroxenite containing sulphides of Ni, Fe and Cu and thin bands or concentrations of chromite (it contains by far the greatest known ore reserve of plantinoids in the world). The Reef occupies a specific position in a regularly layered sequence of anorthositic, gabbroic and pyroxenitic rocks. Cousins maintains that "the heavy minerals, which appear to have been primarily precipitated under the influence of gravity from solution in the magma, contain the platinoid metals principally as sulphides, but also as ferroplatinum, together with arsenides, tellurides, bismuthides, antimonides and possibly stannides and selenides. Considering the mechanism proposed by Cousins platinoid metals were precipitated under the influence of gravity from solutions in the magma - this could be seen again as a case of superimposed paragenesis (?hydrothermal) in the sense that the chromite might represent the first phase (or paragenesis related to the pyroxenite) and platinoids were (?) autometasomatically-hydrothermally mobilized by solutions at a subsequent phase,. As present textural studies of the Merensky Reef show, PGM are following fractures of chromite and other silicates and support the "superimposed paragenesis" interpretation of the platinoid minerals. Concerning the mobilization or remobilization of PGE, Ottemann and Augustithis put forward that it can

take place under lateric conditions (see Chapter 13). Stampfl (1966) stated that native platinum occurs in a troctolitic gabbro at Congo Dam, Sierra Leone. According to him, titanoferous magnetite and silicates have been replaced by bornite and chalcopyrite, and platinum is found as small grains in the sulphide-rich parts. It is suggested that low temperature solution "hydrothermal" were responsible for the formation of the platinoid minerals. Genkin (1959) and Ramdohr (1960) have also supported the hydrothermal formation for certain platinoid containing minerals (see also Augustithis, 1979). In contradistinction, Piispanen and Tarkian (1984) support that volatile components played a significant role in the solution, transport and final deposition of sulphides and PGM. In this connection, it should be ac-

knowledged that Koen (1964) tentatively suggested that platinoid grains of the Witwatersrand conglomerate showing rounded outlines were formed in situ or that they were inherited intact from the source area, taking into consideration the minimum size which will suffer shape modifications during transportation. In support of the geochemical relationship of the basic rocks (or mafic minerals) with gold, Forster (1960) describes a lens-like ultrabasic igneous complex of the early Pre-Cambrian granite-gneiss of the Low Veld, north-eastern Transvaal, showing layering due to magmatic differentiation. An association of olivine and gold occurs in the foot wall of the complex and because of cataclasis in situ and metamorphism, gold recrystallized and concentrated locally.

117

Chapter 21

Granites/Pegmatites and Related Metallogeny

Orthodox and Unorthodox Views on the Relationship Between Granites and Metallogeny Considering the magmatic hypothesis of granites and the relationship between intrusive granitic bodies and metallogeny, some aspects will be briefly discussed in the present Chapter. (i) The hypothesis of granitic magma (including anatexis - i. e., neo-magmatists' point of view, those who support regeneration (palingenetically) of granitic magma from sediments), has acted as deus ex machina for solving many problems in geology and particularly in explaining perigranitic mineralization (the metallogeny in the granitic aureoles and the surrounding rocks). (ii) The skarn and ore bodies at the contact of the granites with the intruded country rocks have been interpreted as pyrometasomatic-pneumatolytic or thermal metamorphic-metasomatic derivatives of the granitic magma. (iii) Broadly speaking, the following phases have been recognized: the orthomagmatic phase (the main phase of the crystallization) of the granitic mineral components; the pegmatitic crystallization phase; the pneumatolytic-metasomatic phase (in cases preceding or synchronous to the pegmatitic); the hypothermal, mesothermal, hydrothermal, epithermal and telethermal crystallization phase (often each characterized by special mineral paragenesis or mineral assemblages). Whereas most of the study-cases presented in this Chapter are based on the concept of the granitic magma hypothesis, it should be emphasized that the entire concept and certain parts of it have been seriously challenged at different times of the century old granite controversy. Since the presentation of "the other side of the coin", namely the transformist's views, have been shown by a number of authors, only reference to published works will be made here, in case some readers are not familiar with these views and interpretations. The subject of granitization is an enormous topic which is outside the scope of the present effort. Only some of the most significant references on granitization will be quoted here. The orthomagmatic phase and crystallization sequence from a granitic magma has been challenged specifically by Drescher-Kaden (1948) in "Die Feldspat-Quarz-Reaktionsgefilge der Granite und Gneise, 118

und ihre genetische Bedeutung"; Erdmannsdörffer (1950) in "Die Rolle der Endoblastese in Granit"; Drescher-Kaden (1969) in "Granitproblem"; Augustithis (1973) in "Atlas of the textural patterns of granites and their associated rock types"; and Augustithis (1993) in "Atlas of granitization textures and processes". The pegmatitic phase has been interpreted as an exudation phase by Drescher-Kaden (1948) and Augustithis (1962) in "Researches of blastic processes in granitic rocks and late graphic quartz in pegmatites (pegmatoids) from Ethiopia"; and the leucocratic (aplitic) phase by Drescher-Kaden (1974) in "Aplitische Gänge in Graniten und Gneisen" as a hydrothermal metasomatic phase, again as an exudation phase. Here, metasomatic replacements played a most significant role. The granite textures as products of magma crystallization have been challenged by Erdmannsdörffer (1950) in "Die Rolle der Endoblastese in Granit" as well as the transgressive contacts of granites. Skarn bodies have been interpreted as a result of the basic front release by the granites as a result of assimilation of basic xenoliths by the granites. This interpretation was advanced by Reynolds (1947) in "The association of 'basic fronts' with granitization" and later by Augustithis (1973) in "Atlas of the textural patterns of granites, gneisses and associated rock types (Chapter 26)". Hydrothermal metallogeny (mineralization) has been regarded as perigranitic metamorphic-metasomatic mobilization by Augustithis (1990) in "Atlas of metamorphic-metasomatic textures and processes". (Daly, 1917) considered metallogeny mineralization as metasomatic metamorphic mobilization.) Augustithis (1990) in particular regards granite-related metallogeny as products of metamorphic metasomatic processes, a part of which is the granite formation itself. In Chapter 3 (Geochemistry - Isochemical versus Allochemical) he referred to metallogenic processes (including hydrothermal mineralization) as due to metamorphic-metasomatic mobilization and differentiation, involving not only granite but also the wider geo-environment in which granitic "Malakton" was formed and mobilized. Augustithis states the following: "Strange though it may sound, in the case of metamorphic-metasomatic differentiation (graniti-

zation, skarnification perigranitic metallogeny), geotectonic events have perhaps been the ultimate cause - since all these mobilizations are an integral part of geotectonic evolution and require an energy source". Furthermore, as a consequence of the concept of the granitization sensu stricto, the role of granitic magma intrusion as a deus ex machina providing solutions for most metallogenic problems ceased to play a dominant role, thus, the controversy of syngenetic versus epigenetic origin for certain stratabound mineralizations and deposits was initiated and encouraged, which later also involved vein deposits which hitherto had been considered to be products of magmatically derived solutions or fluids. Research on the origin and derivation of fluids in metamorphics and also in metallogeny (or deposits) evolved into a controversy. Based on isotope geochemistry, the derivation of fluids has been considered to be either magmatic or meteoric-connate in origin while cases are recorded in metamorphics where fluids of both derivations (according to some researchers) might have operated. At any rate, the derivation of fluids from "magmatic granites" and granitization through the operation of metamorphic-metasomatic solutions is extensively discussed by Augustithis (1990) in "Atlas of metamorphic-metasomatic textures and processes", Chapter 9: Fluids in Metamorphics. Considering certain elements which can be classified as lithophile-granitophile, e. g„ Mo, Sn, W and U, it should be noted that in the same "granitic complex" and its geoenvironment, they can occur in almost all (or in some) of the metamorphic-metasomatic phases. For example, Augustithis (1971, 1973) reported from the Kimmeria granite in northern Greece that Mo occurs disseminated in the granite, in the skams (cases are also reported of Mo in the pegmatitic phase) and as hydrothermal phase in quartz veins. Thus, we have the element occurring in the granite and its perigranitic geoenvironment. Occasionally some of the characteristic elements belonging to the typical pneumatolytic phase of the granite-perigranitic geoenvironment as Goldschmidt and Drescher-Kaden (in Goldschmidt, 1954) pointed out, have been picked up by the granite from the country rock sediment, as in the case of B. Tungsten is another example that can be present in skarns, in pegmatites and in hydrothermal quartz veins. As often reported from southern China, W occurs in the Devonian country rock sediments and its presence and enrichment in some granites, skarns and veins is believed to be a remobilized derivative from the sediments, due to metamorphic-metasomatic differentiation. Also Sn is a granitophile element which occurs in the metamorphic-metasomatic differentiation phases (granitized material, skarns, pegmatites, hydrothermal veins). It occurs in granites, in the pegmatitic pneuma-

tolytic phase and in hypothermal-mesothermal veins as well as remobilized veinlets transecting uraninite (see Figs. 808 and 809). Uranium while occurring with accessory minerals in granites and as oxides, can be enriched in the pegmatitic phase and in hydrothermal-telethermal veins of the metamorphic-metasomatic differentiation. While U in the granites might be associated with accessory minerals which are often derivatives from granitized sediments themselves due to the great geochemical mobility of m it is not surprising to find it (rarely in skarns) in pegmatites (exudation and lateral segregation products) in granites and the perigranitic geoenvironment, and as mentioned in the "telethermal phase", often as colloform pitchblende. Uraninite impregnation and replacement of sediments is also common (Fig. 625). In contrast to the concepts outlined of metallogeny related to granitization, study cases will be presented, according to orthodox views where the relationship between granites and mineralization (metallogeny) is based on the hypothesis of granitic magma crystallization and derivation of mineralization from the intruded granitic magma. For most economic mineralogists the concept of granitic intrusion is used as deus ex machina for interpreting the different phases of mineralization related to granites. Abdullaev (1954) in "Genetic connection of ore deposits with granitoid intrusions" emphasized the importance of considering not only the physical and chemical factors of the ore forming fluids but also the geological importance of assimilation and hybridism as factors of ore deposition, especially in the environment of the upper crustal zones which are the main region of the action of magmatic emanations and in which the "differential zonality" of ore deposits is formed. From the plethora of magmatic interpretations put forward to explain the relationship between granitic magma and metallogeny, certain study cases will be presented, indicating diverse interpretations. Oftedahl (1958) introduced the hypothesis of exhalative sedimentary ores, involving the precipitation of metals in volcanic gases escaping from a crystallizing magma chamber into the sea. (Gases similar to those known to contain metals in contact pneumatolysis around a granite may, according to Oftedahl, be the source of the orogenic pyrite and magnetite ore bodies in the Scandinavian Caledonites as well as in the Rio Tinto type of deposits, when emitted at the bottom of the sea.) It is argued whether the relatively sudden appearance of iron in the sea, precipitating in a suitable chemical environment as sedimentary iron ore beds, is due to exhalative origin. It is also tentatively suggested that some zinc, lead and copper ores with certain sedimentary features may represent similar precipitations. In addition, Koark (1963) interpreted the Falum type of sulphide deposits in the Precambrian of central 119

Sweden as having been formed by an exhalative sedimentary process during the building of the Leptite Formation in which the deposits are situated. In contrast to Koark, Geijer (1965) argued that in the sulphide province of central Sweden the sulphide mineralization took place during the folding of the leptite series and the concomitant intrusion of the early Svecofennian granites. Matveyenko and Shatolov (1958) discussing the metallogeny of north-eastern USSR, state the following: "This is the largest structural-metallogenic unit on the globe. Gradual displacement of the activity in geological processes, including magmatism, from the north-west to the south-east towards the border of the Pacific Basin is characteristic. A characteristic feature of Mesozoic and Cenozoic mineralization is its relation in time and space to intrusive complexes. Those of granitic composition gave rise primarily to noble metal and rare metal mineralization (gold, tin, tungsten, molybdenum, mercury, etc.). The predominant number of deposits belong to post-magmatic hydrothermal formations of varied genetic types." Another interesting case of metallogeny related to granitic magma is presented by Kloosterman (1967). Cassiterite deposits occur in greisens and stock-works of veins in sub-volcanic granites intrusive to the Precambrian Brazilian complex. In addition, according to Kloosterman, the presence of ring structures, orthoclase, perthites and both granite and porphyries, the absence of tourmaline and the association of biotite granite with topaz greisens and cassiterite, and of hornblende granite with ilmenite and zircon, emphasize all the resemblance of the Rondonian intrusions of the Nigerian younger granites (sub-volcanic granites). A rather contrary case of mineralization is presented by Buryak (1965) in his studies entitled: "Relationship between the mineralization of Precambrian strata and regional metamorphic zoning in the Vintim-Patom uplands". According to him, the mineralization appears related to metamorphic grade so that (1) rocks of the sericite-chlorite subfacies of the green schist facies contain quartz carbonate veins, much pyrite and rich gold ore; (2) rocks of the biotite-chlorite subfacies of the green schist facies contain quartz and quartz carbonate veins, pyrite, some pyrrhotite and some gold; (3) rocks of the epidote-amphibolite facies contain quartz veins, pyrrhotite and little gold and, (4) rocks of the amphibolite facies contain pegmatites, pyrrhotite and chalcopyrite and no gold. These facts suggest a genetic relationship between ore genesis, metamorphism and contemporaneous folding with the emplacement of a synorogenic late Palaeozoic granite. The interpretation that regional metamorphic zoning (related to metallogeny) is due to synorogenic granitic intrusion, is incompatible with the basic concepts of regional metamorphic zoning, as pointed out by Barrow (1893, 1912), where granites might be present in regional metamorphic terrains but are not the cause of regional metamorphism. 120

Lopez et al. (1944) support that the copper deposits of the Aroa district, Estado Yaracuy, northern Venezuela, resulted from the hydrothermal replacement of schistose graphitic and sericitic limestones interbedded with strongly folded calcareous, graphitic and micaceous schists. The mineralization followed lowgrade regional metamorphism of the sediments and a period of intense local silicification and introduction of sericite, zoisite, epidote, tremolite, actinolite, diopside, chlorite, talc, calcite and dolomite. Further deposition of quartz, calcite and lesser amounts of clinozoisite, zoisite, sericite and microcline accompany the successive introduction of pyrite, galena, sphalerite, chalcopyrite and minor bornite. A supergene alteration is also recorded. Lopez et al. suggested that the mineralization was related to granitic intrusion, although no acid igneous rocks are exposed in the area. From the perspective that metallogeny is attributed to granitic intrusion and furthermore, considering that at that period (1944) no modern sophisticated interpretations were available, it is not surprising that a granitic magma deus ex machina has been brought up to provide an interpretation for the origin of mineralization (metallogeny). Concerning the different case studies treating the relationship of granitic magma and metallogeny, Shcherbakov and Perezhogin (1963), in "Geochemical relation between gold mineralization, intrusives and the enclosing rocks in western Siberia" determined a gold concentration in intrusive siliceous igneous rocks, and in a variety of surrounding country rocks, including those of metamorphic, volcanic and sedimentary origins. According to them "granitoids enriched in gold are associated with volcanic rock terrains. The latter contains almost twice as much gold (0.0065 ppm) as the sandstones and shales. The data suggest that auriferous granitic magmas are formed in situ from pre-existing rocks with a relatively high gold content". The model presented of the relationship between granitoids and metallogeny is based on the hypothesis of anatexis (i. e., granitic magma generation by melting). A comparable case of the granite-metallogeny relationship is described by Sofoulis and William (1969) from western Australia. The authors affirm that "acid to intermediate volcanic rocks of Archaean age are intruded by ultramafic and doleritic dykes and also by discrete granite bodies in tension fractures or openings caused by sinistral dragfolds. The granites are arranged en enchelon and have caused remobilization of Cu which was originally dispersed within the volcanics and pyroclastics. The most significant amounts of Cu (1980 ppm) are now localized at contacts between the granite and coarse pyroclastics: Cu also occurs associated with an aplite-dolerite complex". In contradistinction to the relatively simple mineralization assemblages considered so far, very complex mineral assemblages and association of granites and metallogeny are also often reported. Examples of such

cases are mainly included to illustrate the complexity of mineralization related to granitic bodies. Lawrence and Markham (1962), in "A contribution to the study of molybdenite pipes of Kingsgate, N. S. W., with special reference to ore mineralogy" state the following: "The Mo-Bi bearing quartz pipes in the Kingsgate area occur within and show gradations to, irregular intrusive masses of granitoid quartz, which probably represents the extreme magmatic differentiate of highly acidic granites. The pipes may be classified as epi-magmatic plutonic offshoots grading into hydrothermal bodies. The complex relationships of ore minerals include two main assemblages: pyrrhotitechalcopyrite-iron-rich sphalerite in galena, and molybdenite with Pb, Bi and Te-bearing minerals; also coexisting assemblages not necessarily in equilibrium, in the Pb-Bi-S system include bismuthinite-bismuth, galena-bismuth, bismuthinite-galeno-bismuthinite and cosalite-galeno-bisutite-bismuthinite. In the Bi-Te-S system, bismuth-ikunolite, ikunolite-joseite-A, ikunolite-bismuth-joseite-A, ikunolite-joseite-A-bismuthinite, bismuth-A-joseite-B, and bismuth-ikunolitejoseite-A-joseite-B assemblages have been observed. The occurrence of apparently stable pyrite-pyrrhotite and bismuth-bismuthinite assemblages suggests that deposition took place in a system closed with respect to sulphur. Arsenopyrite, (?) gudmundite, cassiterite, wolframite, gold and pyrargyrite occur". The above quoted detailed studies of mineral assemblages related to granite indicate the complexity both geochemically and mineralogically of the mineralization in the particular study case. A rather less complex mineralization (in comparison to the Kingsgate, N. S. W.) involving Mo, is the case of the well-known "Climax molybdenum deposit of Colorado". According to Butler and Vanderwilt (1931), "the Climax molybdenum deposit occurs in altered Precambrian granite and has the shape of a cone, enlarging downwards. The ore mineral is molybdenite in small crystals in veins with quartz and locally, orthoclase gangue. In the ore zone pyrite, fluorite and topaz are widespread and finely crystalline; chalcopyrite, hübnerite and sphalerite are only locally common". In contrast to the opinions which regard aplites as a postmagmatic "aqueous phase", Öhlander (1985) regards aplites as representing the direct differentiation products of the granites which have solidified without the development of a mobile aqueous phase (except locally where small pegmatite segregations have been formed). Specifically Öhlander supports that "as crystallization proceeds, the concentration of Mo in the magma increases, and when the aplites solidify, the Mo that cannot be incorporated in the rock-forming minerals is precipitated as molybdenite. Biotite seems to be the major Mo-carrier among the rock-forming minerals in aplites". As a corollary to the mineralization of Mo, W, Sn and U considered in the case of metamorphic-metasomatic differentiations, some additional mineralization

cases of some of these elements related to granitic intrusions will be further presented. Sharp (1958) in his studies of mineralization in the intrusive rocks in Little Cottonwood Canyon, Utah, described an area of about 2 miles west of Alta, containing scheelite and molybdenite. The minerals occur in zones of jointing and fracturing which form a concentric pattern around an intensely fractured central zone. The more intense fracture zones are mainly within a leuco-quartz monzonite intrusive. Furthermore, Agard et al. (1966) have revealed the presence of Be in metamorphosed limestones at the edge of the Azegour granitic massif. W, Mo and Cu are also present in the mineralized zone, which is also attributed to granitic intrusion in this case. Legraye and Goffinet (1955) describe the case of an assemblage of parallel veins located in metamorphic schists in the vicinity of a nose of granitic gneiss. The veins contain wolframite as the main mineral, with quartz and scheelite, tourmaline, white micas and some accessory minerals. In addition to W, Β is another element mobilized in the present study, indicating the relationship of this type of mineralization to the granites. In another case, Sobotka (1956) reported the occurrence of wolframite in the gold-bearing district of Kasejovice, indicating the genetic connection of the auriferous veins with the granitic and granodioritic pluton of middle Bohemia. As a corollary to the genetic connection of W and Au as pointed out, Geffroy and Laforgue (1957) reported scheelite with native gold and with pyrite, arsenopyrite, sphalerite, chalcopyrite, galena and a Pb-Sn sulphide in an aggregate of quartz, biotite and calcite, as a metalliferous vein in the metamorphic aureole associated with the batholith of Mont d'Ambazac. The complexity of mineral parageneses associated with W is indicated by Victor (1957) from the Burnt Hill wolframite deposit, New Brunswick, Canada, where wolframite ore occurs in a series of quartz-topaz fissure veins cutting across interbedded schists, phyllites and quartzites one mile south of an exposed granite batholith. As Victor reports "early vein minerals include beryl, quartz, wolframite, biotite, topaz, cassiterite, molybdenite, apatite, muscovite and scheelite. Late vein minerals are arsenopyrite, sphalerite, pyrrhotite, pyrite, chalcopyrite, galena and native bismuth". Concerning the metallogeny of Sn, Sainsbury and Hamiliton (1967) support that "the principle lode-tin deposits of the world are closely related to biotite or biotite-muscovite granite or to subvolcanic dacite to rhyolitic rocks along major orogenic belts. In Precambrian rocks these belts are broader than in younger rocks, and most of the deposits are in pegmatite with spodumene, tantalite, columbite and other heavy metals. In post-Cambrian rocks contact metamorphic, pneumatolytic-hydrothermal, subvolcanic or tin-silver, "fumarolic" and disseminated deposits are recognized. 121

The principal deposits are of pneumatolytic-hydrothermal and tin-silver types, and are localized along zones of deep faulting within the broader granitic belts." Furthermore, according to the authors, "within each district, the Sn deposits have been formed in the higher temperature or deeper parts of lodes and grade upwards or outward into Cu and W deposits". Sainsbury and Hamilton also support that the concentration of lodes along deep post-granite faulting and other geological and geochemical relations within individual districts is interpreted as evidence that the deposits were formed by magmatic emanations from depth, from not visible rocks. (The associated metals are all highly volatile and might be carried up subcrustal fractures by solutions rich in CI, F and B.) From the presented study cases it is apparent that there is a genetic relation between granites and Sn metallogeny, but the implication of granitic magma as the mechanism operating everywhere, and a solution to all the problems in the sense of deus ex machina can be seen in the case of the tin deposits at Gierczyn in the Iser Mountains, Lower Silesia. According to Jaskölski and Mochnacka (1959), cassiterite in close association with pyrrhotite, chalcopyrite, blende, galena, arsenopyrite and some accessories, is described from the ore deposits of Gierczyn. The ore bodies occur as small concordant lenses in mica schists which were strongly folded during Caledonian orogenesis. The mineralizations are regarded as caused by hydrothermal emanations from late Variscan intrusions which outcrop about 10 km further south. Another example of magmatic granitic intrusions is reported by Ekwere (1985), in the case of the ring granitic complexes of Banke and Ririwai Young granitic complexes of Northern Nigeria. These complexes are associated with Sn-Nb mineralization and are, according to Ekwere characterized by high Li, F and Rb contents and Rb/Sr ratios and low Ba and Sr contents and Ba/Sr ratios. Another reported case is Tongolo Anorogenic magmatic granitic complex, northern Nigeria. According to Imeokparia (1985), the mineralization is domi-

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nated by Nb-Sb and Sn-W types. Furthermore he states, "the Tongolo biotite granites are characterized by enhanced values of a suite of trace elements (Sn, Nb, W, Zn, Rb, Li, F, Th, Y, U)" which identifies them as "specialized" granites. In addition, according to Imeokparia, "these granite series are marked by what is interpreted as the "lithophile factor" (heavily loaded by Li, Rb, F, Th, Ga, Y, U) dominated by magmatic processes and metallization factors (Nb, Zr, Ca, U, Zn, Li and Sn, W, Rb, F, Th) which are dominated by postmagmatic processes". In addition to the ring granite complexes and the "specialized" granites mentioned, which were described as magmatic intrusions, Ball and Basham (1984) consider the Bosworgey granite cusp as an apical portion of the concealed northern extension of the Tregonning-Godolphin granite ridge (Cornwall). The Bosworgey granite cusp is characterized by unusually high values of Β, Ρ, Mn, Fe, As, Cu, Nb, Ta, Bi, Sn, W, U and S which are largely present as tourmaline, apatite, pyrite, arsenopyrite, chalcopyrite, bismuth, columbite, cassiterite, wolframite and uraninite, and low levels of Zr, Hf, Ti and REE present in zircon, ilmenite and monazite. Furthermore, according to the authors, although the Bosworgey samples show characteristics of "S" type of granites, the accessory mineral assemblages are typical for high temperature lodes (cassiterite, wolframite, arsenopyrite, chalcopyrite) and the assemblage is concluded to be the cusp analogous of hypothermal lodes produced by extreme differentiation and concentration of volatiles. In contrast to the magmatic interpretations provided, Serebryakov (1961) presents cases of different forms of high temperature alkaline autometasomatism in granitoids of certain tin-bearing massifs of the upper course of the Kolyma River. Several stages of metasomatism are recognized. Tin mineralization is associated only with sodium metasomatism at relatively shallow depth during which the related feldspar lattice is preserved. Also in this case, autometasomatism might be seen as part of the granite forming process.

Chapter 22

Granitization - Anatexis

(a) Granitization Related to

Mineralization

Considering the general hypothesis of granitization related to metallogeny presented in the previous Chapter, granitic magma in the sense of deus ex machina cannot, and is not, be used to provide the explanation for many cases of sediment-metamorphic rock related mineralization. The trend to explain certain deposits as syngenetic-sedimentary or sedimentary is perhaps an indirect consequence of the wider influence which has been exercised by the granite controversy. Admittedly, up to now, the supporters of granitization have not provided detailed studies explaining granite related metallogeny as a consequence of granitization sensu stricto. The supporters of anatexis (neomagmatists) also accepted the explanations of magma derived metallogeny, since once a magma generation by anatexis was available, the metallogeny would be in accordance with the scheme of magma consolidation. Thus, a fundamental theoretical gap exists between the supporters of granitization sensu stricto and of anatexis of sediments, since the mechanism of granite related metallogeny is entirely different, despite the acceptance of an ultimate common source and that is the sediments (or more precisely, the precursor rocks). Augustithis (1973, 1993) is a representative of the granitization school, since he provided evidence of a granitization origin of granites exhibiting textural patterns (metamorphogenic/metasomatic)7. After compar7

Perhaps a fair critique regarding Augustithis' views of the significance of textural analysis on the granite problem, is a review by the Australian Mineral Foundation (AMF Information Book Review Series No. 2061) concerning the Atlas of granitization textures and processes (1993), which states: "The 'granite question' of the 1950s has been laid to rest for most igneous petrologists by the acceptance of the compromise recognition of granites and 'granites', identified largely on field evidence as "I type" or "S type", and also accepting the S type granites may under some circumstances be mobilized and emplaced in a manner which makes them indistinguishable from I type. The author and compiler of this magnificent Atlas, an indefatigable producer of petrological atlases, contends that texturally, and in every other respect, granites are granites, and that all are the produce of a granitization process involving the transformation and remobilization of mainly sedimentary and metasedimentary materials. The Atlas is presented in two Parts, each supported by an appropriate text profusely illustrated with black-and-white

ing these patterns with metamorphic-metasomatic patterns in metamorphic rocks sensu stricto, he interpreted perigranitic mineralization (1985, 1993) as a metamorphic-metasomatic differentiation in the sense of a sequence or relationship of the following processes: metamorphism-ultrametamorphism, granitization, perigranitic-metasomatism, skarnification, exudation phase related to granitization (in pegmatites, aplites), "hydrothermal" vein mineralization as a result of metamorphic-metasomatic differentiation and lateral secretion (segregation) of the granitized wider geoenvironment involving sediments. Thus, granite related metallogeny is a derivative of granitized sediments and of country rocks, and the mechanism is not the deus ex machina interpretation of an intruding granitic magma and its consolidation mechanism that provides solutions to all problems. However, it should be emphasized once more that the granitization school has a long way to go before it can provide plausible interpretations for the significance of granitization sensu stricto in perigranitic metallogeny. However, some sporadic contributions have appeared at times, although far too insufficient to revert the trend of thought of granitic magma mineralization to a complex and sophisticated process of granitization related perigranitic mineralization. Furthermore,

photomicrographs (some 800). Part 1 develops the author's views on granite genesis with seven essays: The granite problem and the significance of textures; the bio- and geoevolutionary system (granites as a model of a geo-evolutionary system); granitization (sensu stricto) versus anatexis; rapikivi sphaeroids; orbicular structures; granophyres derived from quartzites - metamorphically-metasomatically; apogranites, their textures and metallogeny. Part 2 is also divided up into brief chapters (26 of them), each concerned to describe and illustrate a particular fact of granitic texture and associated intergrowth. Again the topics cover the granite problem and the importance of textures; the granitization process; crystalloblastesis; tecoblastesis; and so on, covering all manner of crystallization phenomena, intergrowths, rapakivi structures, accessory minerals, tectonic influence on granitic rocks, metasomatic and other alteration processes, together with several chapters addressing differing aspects of pedogenesis, metallogenesis and geochemistry. Although the author is convinced that the overwhelming evidence supports a granitization process rather than a primary granitic magma, he concedes that in some cases either concept could be supported on textural grounds alone ..."

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it will be very difficult to convince economic geologists (specialists on mineralization) to abandon a working hypothesis as the granitic magma intrusion is for a complex, problematical and not yet widely accepted "scheme" of transformation and mobilization (remobilization) of materials as is advocated by the supporters of granitization and this, although granitization has won most of the "battles" concerning the textures, the space problem and the origin of the granites in general. In contrast to the orthodox views - granite intrusion is considered as due to magma solidification and the related mineralization as a derivative of the granitic magma - long ago believers in granitization pointed out that, if granites are products of transformation of sediments/metasediments (or precursor rocks), granite related mineralogy must be a derivative of the granitized sediments or of the wider geoenvironment related to the granite emplacement ("granite" being a part of wider metamorphism). Again, in contrast to the orthodox views regarding the granite derived mineralization as a consequence or phases of magma consolidation (namely: orthomagmatic phase, pneumatolytic-metasomatic (pyrometasomatic), pegmatitic, hypothermal, mesothermal, "hydrothermal", and telethermal), the unorthodox views see granite related metallogeny as part of granitization related mobilization processes and as an ultimate source of the sediments. An attempt will be made here to present only some cases of contributions where a derivation of mineralization is related to the sediments and to the precursor material, transformed in one way or another into metamorphic-ultrametamorphic "granitoids", and even to granitoids produced by anatexis (where the granitic magma scheme is not precisely followed). Marmö (1960) while discussing the origin of ores from the standpoint of ore prospecting, concludes that sediments are the main source of ore forming materials and may also yield the potassium of late-kinematic granites which often accompany the ores. Furthermore, Marmö adds in his contribution: "On the possible genetical relationship between sulphide schists and ores" that in many so-called hydrothermal ores, emplaced under hydrothermal conditions, the ore forming material has been derived from sediments mainly of sulphide-graphite schists. A more elaborate case of mineralization related to granites is again presented by Marmö (1962), in "The molybdenum bearing granite of the Wankatana River, Sierra Leone". According to the author: "disseminated molybdenite occurs in the slightly sheared central portion of a petrologically complex late-kinematic granite occupying the core of a plunging anticline in metasedimentary schists to the north-east of Bumbuna. A homogenous aplitic granite is associated with aplite, porphyroblastic microcline granite, and with lenticular, partially granitized remnants of the country rock". Marmö further points out that although the original in124

trusion was sodic, the metasomatic introduction of microcline has resulted in strongly potassic zones. Furthermore, the molybdenite is concentrated in a potassic, gneissose, aplitic rock, rich in biotite and myrmekite and with accessory fluorite and garnet; tourmaline occurs in nearby pegmatites. The emplacement of granite was accompanied by potassium and minor boron, metasomatism and, during a period of shearing, by introduction of hydrothermal solutions bearing F and Mo. The deposit is compared to similar Mo mineralization at Mätäsvaara and at other localities in Finland. Additional studies by Dechow (1960), in "Geology, sulphur isotopes and the origin of the Heath Steele ore deposits, Newcastle, New Brunswick, Canada" present evidence that "sulphur isotope ratios obtained on sulphide and sulphate specimens from ore, granite, acid and basic volcanics, porphyry, and sediments do not demonstrate any fractionation during hypogene or supergene mineralization". Furthermore, the ratio spread (21.82-22.02) suggests a homogenous source of mineral solutions. Isotope ratios for sulphides in a gneissic granite near the ore bodies correspond closely to those for the ore. Enrichment of ore sulphides in 34 s suggests a source rich in sulphate. An original source bed buried to depths sufficient to effect granitization, reduction of sulphates and mobilization of the resulting sulphides towards favourable loci is proposed to account for the ore bodies. In this connection, it should be pointed out that in many cases when using the term granitization, its connotation is uncertain since in cases it implies only granite formation which is contrary to granitization sensu stricto. De Carvalho (1969), describing the Gondomar Sb and Au mines near Oporto in the Douro River region of Portugal, states that Sb-Au mineralization is concentrated in the north Douro River fault, whereas Pb-ZnAg mineralization predominates in the Castelo de Paiva area to the south. The Sb-Au deposits occur in quartz veins in schists, and are typical hydrothermal veins related to the Hercynian granitization. Another case of mineralization related to granitization is reported by Brotzen (1957). He considers both syngenetic and epigenetic viewpoints and stresses that many of the characteristic features of the ores favour an epigenetic interpretation. The existence of hidden younger granites as ultimate source of the ores, however, is considered less likely than the formation of mantled gneiss domes. According to Brotzen, such a process brings the process of granitization into discussion. From these study cases discussed above, it can be concluded that, with the exception of the mobilized syngenetic sedimentary and the metamorphic-metasomatic ores, the unorthodox petrology remains a nonexploited field by economic geologists.

(b) Anatectic Granitoids and

Mineralization

In contrast to granitization, cases are reported of metallogeny related to anatexis of sediments. Here a temperature increase has resulted in remobilization of sediments and their components. Such views support the operation of melts-fluids and are incompatible with granitization sensu stricto where solution-fluids were of primary significance. Certain cases supportive of mineralization, temperature increase or related to anatexis will be presented for covering the wide gamut of processes put forward as possible interpretations of certain types of mineralization. Sullivan (1957), considering the role of temperature in ore deposition, presented a discussion on the significance of thermal stability ranges of minerals in the solid state in relation to zoning and paragenesis as being more important than relative solubilities in aqueous solution for the determination of the sequence in ore deposition. According to Sullivan, granitization, metamorphism and ore formation may be different expressions for reactions and movements of elements in the earth's crust in response, particularly, to heat, pressure, temperature and chemical gradients. Sullivan thus suggested "a different approach to the subject of ore deposition". However in this connection, it should be stressed that any element mobilization in solid state is related to solid-diffusion, an approach already proposed for granitization by Perrin and Roubault (1937). Granitization in solid state is in contrast to the granitizationmetamorphism concepts of Erdmannsdörffer, Drescher-Kaden and Augustithis, where solution-fluid operations are essential (see Chapter 23). In contrast to Sullivan's interpretation of mineralization "in solid state" palingenesis of granitic melts or anatexis of sediments is proposed by some researchers. Ljunggren (1964), reporting a newly discovered tinbearing region in the province of Western Brazil, states that the Brazilian deposits are characterized by the association of cassiterite with topaz, and the Bolivian by cassiterite and tourmaline. Furthermore, according to Ljunggren, the tin mineralization in Bolivia is related

to the occurrence of large granitic batholiths of the East Cordilleras of Triassic-Jurassic and Tertiary age, whereas the cassiterite-topaz minerals are replacements in Precambrian/Early Palaeozoic granites. Thus, mineralization in the latter area is considered not to be related to the formation of the rocks but is a later event. Genetically, the Bolivian mineralization probably had its origin in Early Palaeozoic Andean geosynclinal sediments, formed by weathering of the Brazilian granites from which large amounts of tin were derived. This tin, included in the hydrolysates of the marine sediments, would then have given rise to a second mineralization, related this time not to topazification but to extensive tourmalization connected with the transformation of these sediments into magmatic granitic melts in Middle Mesozoic to Tertiary time. What is interesting in the mineralization described by Ljunggren, is the recycling of the Sn involving weathering, deposition and anatexis. Another case of mineralization associated with granitic magmas formed in situ is reported by Sherbakov and Perezhogin (1963) from western Siberia. Granitoids enriched in gold are associated with volcanic rock terranes. The latter contain almost twice as much gold as the sandstone and shales. The data suggest that auriferous granitic magmas are formed in situ from pre-existing rocks with a relatively high gold content. Furthermore considering metallogeny related to anatectic granites, Ashley (1984) on basis of field, mineralogical and chemical data, supports the derivation of the Crocker Well (Southern Australia) granitic rocks as products of anatexis during high-grade metamorphism. The sodic granitic rocks produced are largely peraluminous and contain high NajO, low K^O, CaO, Rb, Ba, Sr and ferromagnesian elements and variable but commonly high U, Th, Nb, Ce, Y and F values. Also, as Ashley reports, "significant U-Th mineralization is restricted to fractures and local breccia bodies which contain a mineral assemblage rich in quartz, F-bearing phlogopite and minor fluorapatite, sodic plagioclase, niobian rutile, thorian brannerite, monazite, muscovite, chlorite, tourmaline and fluorite".

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Chapter 23

Metallogeny Related to GranodioritesMonzonites

Besides granites, granodiorites have also been the subject of a similar controversy, namely, whether they are crystallizations from magma or granitization-granodioritization products. Whereas the basic concepts of granodioritization have been included in the treatment of granitization, selected cases of the relationship of granodiorites and metallogeny, based on orthodox views, will be presented here - accepting a priori that they only represent limited study cases. Huttl (1963) reports that "the Getchell Au mine, Potosi mining district, Humboldt Country, Nevada, USA, is at present the second largest mine in the USA worked for gold along. The ore bodies are vein systems developed within a major fault zone cutting Palaeozoic limestones and carbonaceous argillites, intruded by a Mesozoic granodiorite stock. Au is associated with realgar, orpiment and lesser amounts of pyrite, pyrrhotite, arsenopyrite and marcasite." Another case of mineralization associated with granodiorite is reported by Wrucke et al. (1968) from Nevada. Geochemical investigations show two well-defined belts containing anomalous concentrations of Sb, As, Bi, Cd, Cu, Au, Pb, Hg, Mo, Ag, Sn, W and Zn. Strong concentrations of Au occur around a granodiorite stock at Tenabo and on the flank of a magnetic anomaly thought to represent a buried stock of Gold Acres. As mentioned, Sobotka (1956) reports an occurrence of wolframite in the gold-bearing district of Kasejovice, Bohemia, which is considered to testify the close genetic connection of the auriferous veins with the granite and granodioritic pluton of Middle Bohemia. In addition, Hilmy et al. (1968) report that gold mineralization occurs in quartz veins transversing granodiorite rocks: the mineral assemblage includes arsenopyrite, pyrite, niccolite, pyrrhotite, chalcopyrite, marcasite, gold (or electrum) and goethite. Considering metallogeny related igneous intrusions (in accordance with orthodox views), Reed and Elliott (1968) report an occurrence of Pb, Zn and Ag ores from the Alaska range (about 23 miles from Farewell). The deposits are replacement bodies and fissure veins in limestones and fracture fillings in igneous breccia. According to them, the deposits in limestone show a

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close spatial relationship to igneous breccia and granodiorite porphyry. The ore minerals are argentiferous galena and sphalerite. Additional to study cases of the relationship between granodiorite and metallogeny, similar study cases are presented, indicating the relationship of metallogeny and monzonite and also quartz porphyry monzonite. In this connection, it should be mentioned that a more detailed treatment of metallogeny related to the porphyry-type of intrusions (granitic porphyry, granodiorite porphyry and quartz-monzonite porphyry) will be presented in Chapter 24. Lacy (1953) in his studies "Differentiation of igneous rocks and its relation to ore deposition in Central Peru" reports that economic sulphide deposits are postmonzonite porphyry in age, and large pyrite ore bodies are intimately associated with quartz monzonites. According to Lacy "the ore fluids leach alkalis, CaO and MgO from the intrusives and their country rocks and were related to the extensive late-magmatic alteration of the quartz monzonites. Intrusion and differentiation of the quartz monzonites and the mineralization probably occurred during the Lower-Middle Miocene; mineralization in the Julcani and Raura districts may be of Pliocene age". Olson (1966) reports a case of Pb-Ag-Zn mineralization associated with a Laramide (?) quartz monzonite intrusion cutting Pennsylvanian-Permian sediments in the Santa Cruz County, Arizona. Hypogene ores encountered at depths below 360 ft, consisting of galena and sphalerite in a quartz gangue, are located adjacent to fault intersections with limestone. The hypogene ores are predominantly replacement deposits with minor fissure fillings. Furthermore, Wilson (1963) describes a case of mineralization from Yerington mine, Nevada, which has yielded over half a billion pounds of copper. The ores are in quartz monzonite porphyry. The mineralization was accompanied by intense silicification. The primary sulphide minerals are essentially limited to pyrite and chalcopyrite which predominates. Minor bornite and covellite are reported and primary chalcocite has been detected.

Chapter 24

Metallogeny Related to Porphyries

The term "porphyry" was first given to an altered variety of porphyrite (porphyrites lapis) on account of its purple colour, then extended to all rocks containing conspicuous phenocrysts in a fine-grained or aphanitic groundmass. The resulting texture is described as porphyritic. The term porphyry usually implies a hypabyssal rock containing phenocrysts of alkali feldspars. This definition is based on the Glossary of Geology (AGI). Drescher-Kaden emphasized the fundamental textural difference between a granitic porphyry and a granite (see "Die Feldspat-Quarz-Reaktionsgefüge der Granite und Gneise und ihre genetische Bedeutung", Drescher-Kaden (1948) especially Section 5 "Reaktionsgefüge und Korrosionsmerkmale an Einschlüssen basisicher und quarzporphyrischer Gänge, p. 89-97). Another term requiring some definition is "quartzporphyry", i. e., a rock containing phenocrysts of quartz and alkali feldspars, usually orthoclase, with or without mica, in a microcrystalline or cryptocrystalline groundmass. If the phenocrysts are abundant the rock becomes a granite porphyry. (These definitions are based on the AGI). Considering the aforementioned definitions and facts, a number of relevant terms are used in association with metallogeny. In this Chapter a number of study cases will be introduced to work on the metallogeny of this type of deposit, having accepted though the term porphyry in a textural sense as suggested by Drescher-Kaden (1948). According to Di Colbertaldo and Pagnacco (1961), argentiferous galena occurs in red quartz-porphyry restricted to two faults which run practically parallel and straight. The minerals identified include: (i) metallic minerals - galena, pyrite, marcasite, chalcopyrite, pyrrhotite, argentiferous tetrahedrite, arsenopyrite and zincblende; (ii) alteration minerals - goethite, covellite, chalcocite and cerusite; (iii) gangue minerals - quartz, fluorite, baryte, siderite and chlorite; (iv) minerals of uncertain origin - rutile and ilmenite. According to Di Colbertaldo and Pagnacco, the veins were formed in two tectonic episodes, each followed by a phase of mineralization. Considering the gangue minerals, the metallic minerals and the minerals of uncertain origin and moreover, the fact that the veins were formed as a

result of tectonic episodes, melting and mobilization of diverse materials might have taken place. A more complex mineralization related to quartzfeldspar porphyry is described by Threadgold (1958), which follows narrow quartz veins containing gold, stibnite, berthierite and other sulphide minerals. In addition, cases are described of epi-mesothermal hypogene deposition of chalcopyrite and cupriferous pyrite which, according to Randall (1960), is probably related to sanidine-bearing quartz porphyry and andesite porphyry, which have intruded as small stocks in the Upper Jurassic quartzites and limestones of the Hualhuani series in southern Peru. Another case study treating the relationship of metallogeny and quartz porphyry is presented by Horsfield and Bannister (1957). The host rocks of the iron deposit of Eisenerz, Austria, are folded and faulted Silurian-Devonian limestones in a zone between the Northern Chalk Alpine region and the Central Alps; the limestones are interbedded with shales, and overlie upper Silurian "transformed quartz porphyry" introduced along fracture systems (also in this case, the relation between the porphyry and tectonic fractures is emphasized) and produced lenticular, lowmedium grade deposits of siderite and ankerite. In contrast to quartz porphyries that are more acid in composition, study cases of porphyries which are more basic in composition will be presented, which might represent plutonic equivalents, intermediate in composition, which could have been subjected to melting by friction along major faulting. Reed and Elliott (1968) described occurrences of Pb, Zn and Ag ores from the Alaska range (about 23 miles from Farewell). The deposits are replacement bodies and fissure veins in limestone and fracture fillings in igneous breccia. The deposits in limestone show a close spatial relationship to igneous breccia and granodiorite porphyry. The ore bodies are argentiferous galena and sphalerite. In addition, Hawley and Clark (1968) described occurrences of Au and other metals associated with small quartz diorite porphyry stocks. The mineralization is in veins, disseminated and as breccia-pipe type. According to Hawley and Clark, additional to gold, Sb, Cu, Ag and Zn are abundant locally. White et al. (1957) recorded vein and disseminated deposits. The mineralization consists of chalcopyrite, 127

bornite, molybdenite, pyrite, haematite and chalcocite related to Tertiary stocks of quartz diorite and dacite porphyry. It is suggested that some chalcopyrite, tetrahedrite, linnaeite (?) and some galena have been formed by ex-solution in bornite. Also included is quartz diorite porphyry formed by melting due to friction as well as the dacite porphyry which might be intrusives and, in cases, extrusives. Additional study cases of porphyries, more basic in composition than quartz porphyry, were reported by Reeves (1963) from Hillsboro district, New Mexico, where the hypogene mineralization of Au-Ag-Cu is related to quartz monzonite porphyry dikes in andesitic extrusives. The mineralization consists of pyrite, chalcopyrite, galena, sphalerite, bornite, malachite, azurite, cuprite, argentite, chalcocite, gold and silver. Wilson (1963) reported a mineralization mainly consisting of chalcopyrite and pyrite with minor bornite and covellite and with primary chalcocite (reported) related to quartz monzonite porphyry from Yerington mine, Nevada. Among the porphyry type of deposits, the copper porphyry type is of significance. The copper porphyry

128

deposits are deposits (according to Bateman, 1950, p. 486) of disseminated copper minerals in a large body of porphyry. From the plethora of research work on these deposits, only a few case studies were selected in the present work. Swayne and Trask (1960) in their study "Geology of El Salvador" describe a large copper porphyry deposit in northern Chile. The most important primary ore mineral is chalcopyrite. Bornite is commonly observed but sparsely distributed. Molybdenite as fine disseminations and veinlets is an important accessory mineral. Pyrite is widely but unevenly distributed in fine seams, veinlets and disseminations, and enargite and tamatinite are rare. In addition, according Swayne and Trask, the ratio of iron to copper appears abnormally low in comparison to other porphyry copper deposits. Furthermore, Macnamara (1968) described a copper porphyry mineralization at Pnaguana, Bougainville Island. Andesite has been intruded by a series of quartz granodiorite bodies. Both the intruded and the intrusive rocks show extensive mineralization and wallrock alteration. Chalcopyrite is the dominant ore mineral and is associated with bornite and traces of molybdenite.

Chapter 25

Skarns-Pyrometasomatic Metallogeny (and Superimposed Metallogeny)

The genesis and metallogeny of skams is a vast subject which is beyond the scope of the present atlas. However, an approach will be made to theoretically outline the most significant types of Skarns: their genesis and metallogeny, as is attempted in a recent publication (1991) by Theophrastus Pubs., Athens. (i) Geology-mineralogy of skarns - Main provinces of skarn deposits in the world; - Skarns of different geological environments; - Mineralogy and petrology of skarns; - Mg-skarn types; - Ca-skarn types; - Structures/microstructures and textural studies; - Zonal relation in time-space. (ii) Fluid-country rock interactions (mechanism of formation of skarns) - Bimetasomatic skarns; - Contact infiltrations; - Autocreative skarns. (iii) Physico-chemical conditions of skarn formation based on experimental and thermodynamic data - Temperature and pressure; - Physico-chemical calculations and models; - Behaviour of rock forming components (Ca, Mg, Si, etc.); - Behaviour of fluid components (HjO, C0 2 , HCl, HF, etc.); - Mass-transport and component balance; - Inclusions; - Duration of processes. (iv) Skarn formations with relation to magmatic intrusions - Granitoids pegmatites; - Alkaline rocks; - Gabbro-syenite plagiogranites; - Skams of the contacts of lamprophyres. (v) Skarn formation with relation to metasomatic granitoids - Basic front mobilization; - Country rock mobilization; - Relation of xenolith-assimilation to skarn formation; - Skams formed within the granites by mobilization. (vi) Geochemistry of skarns - Major element geochemistry; - Trace element geochemistry; - Isotope geochemistry of skams (oxygen, carbon,

sulphur, etc.); - Phase rule considerations. (vii) Metallogeny of skarns - Space-time relationships of mineral assemblages in different skarn deposits; - Phlogopite, Fe, Β; - Fe, Cu, (Co)Zn, Be; - Proximal Sn, Be-F-W skams; - Distal Sn skams; - Colloform hydrothermal (lievrite-hedenbergite) skams. (viii) Exploration methodologies - Economic evaluation of skams; - Pathfinder elements for Au-bearing skams. The above outline is a theoretical scheme which puts forward the main topics and aspects related to skams. In this connection, it should be mentioned that the subject of skams is presented in the publication Skarns: their genesis and metallogeny (Theophrastus Pubs., 1991) under the headings: Metallogeny of skams; Magnesian skams; Regional consideration of skarn types; Physico-chemical studies of skarn formation and Exploitation of skams. Thus, it covers only some of the topics and aspects of the outline presented above. Some general concepts and different aspects concerning skarn types, intrusives, geotectonic setting, zonation, metallogeny and its variation, magmatic and post-magmatic phases and in addition, the role of fluids are briefly considered as follows: Skams are formed in the aureole and contact zone of an intrusive rock but are not restricted, strictly speaking, to the contact alone. They are caused by the thermal influence of the intrusive (referred to as the metamorphic effect) and usually followed by the metasomatic effects, which can be produced even after the consolidation of the intrusive. The type of skam formed will depend on many factors and an important one is the composition and geotectonic setting of the intruded rocks. Zharikov (1991) mentions that "different skam types (magnesian, calcic, contact-infiltration, bimetasomatic) have been specified, with their geological position and principle genetic parameters of formation and of ore mineralization being indicated. Based on experimental investigation, physico-chemical conditions of the skam formation have been defined, i. e., range of P-T; 129

aMgCl2 and aCaCl2 effects of the C0 2 activity and acidity." Zeng Ming Hua (1991), describing the polymetallic Qibaoshan skam of South China, mentions that the ore deposit is located in the contact zone of granite porphyry and carbonate beds. Furthermore, he supports that "where the intrusive body contacted dolomite and dolomitic limestone, Mg-skarn was formed, and where the intrusive body came in contact with limestone Caskarn was formed. In perpendicular section, Mg-skarn is on the Ca-skarn, both being transitional to each other". Similarly, Wang Shufeng and Zhang Qiling (1991) report that metamorphism caused the formation of calc-silicate marbles and hornfels which were subsequently replaced by skarnoid and ore minerals during metasomatism. "Large bodies of skarn formed in micritic carbonate rock of Devonian age adjacent to a Qianlishan granitic pluton. Skarn formation is divided into magnesium-rich type from dolomite and calciumrich type from marble. Magnesium skarn is formed first and at high temperatures, followed by calcium skarn. It is obvious that bodies of magnesium skarn are smaller than those of calcium skam. Massive calcium skarn is dominated by andradite garnet and zoned from andradite garnet + salite through andradite garnet + idocrase + fluorite, then from idocrase + wollastonite + calcite to wollastonite + calcite. Additional phases involve actinolite, epidote, ilvaite, titanite, quartz, scheelite and so on". As pointed out, in addition to the composition of the intruded rocks, the composition of the intrusive also is of great significance. Casquet and Tornos (1991) discuss additional factors to the composition of the intruded rocks. The following extracts are quoted from their paper: "Skams are widely distributed in all the Hercynian Basement of the Iberian Peninsula, near the contacts of Precambrian and Paleozoic marbles and calcium-rich rocks with gabbros to leucogranites of Hercynian age. They can be classified in four main types: Fe, W, Sn(-W) and polymetallic skarns. The regional distribution matches the zonal division of the Hercynian Mobile Belt as based on stratigraphic, tectonic and magmatic criteria. Skarn type is found to be dependent on magma chemistry and depth of formation; other variables seem to be of minor importance. They control the fluid composition as well as the precipitation mechanisms involved in ore formation. However, local variable such as metal pre-concentrations or oxygen buffering by the protolith may significantly alter the silicate and ore mineralogy. Calcic Fe skarns are the most abundant and they are related to epizonal intermediate to basic rocks of the cafemic association; such rocks occur in subduction related settings or in late wrench-shear environments. If the protolith is Fe-rich enough (amphibolites, diorites, ...) iron skarns can form in many tectonic settings. 130

Calcic W skarns are mostly related to peraluminous granitoids of intracontinental fold belts (W(-Sn) association) or postorogenic uplift regions W(-Mo) and W(-Au) associations). Usually they are related to mesozonal adamellites to leucogranites that do not show significant W-enrichment. When skarns are related to more epizonal bodies, granitic rocks are highly hydrothermalized and W(-Sn) and W(-Mo) veins and greisens are common. Magnesian Sn(-W) skarns are only related to dolomitic marbles in contact with mesozonal, highly evolved, leucogranites of collisional type. In such a tectonic setting, Sn appears as a geochemical imprint in Fe, W and polymetallic skarns. Polymetallic skarns are also scarce and are epizonal and distal equivalents of Sn-W skarns. They have characteristics of Zn-Pb skarns, such as the presence of Mn-rich hedenbergite, abundant sphalerite and C0 2 -poor hyposaline fluid inclusions; the relationship of these skarns with highly evolved granites produces also on garnet and clinopyroxene an imprint typical of W skarns, as well as a polymetallic mineralization with Zn, Cu, Sn and W." Regarding Casquet and Tornos' hypothesis for skarn type formation, it can be noticed that in addition to the composition of the invaded rocks the composition of the intrusives and geotectonic setting were also considered to be of significance. Furthermore, their hypothesis is partly based on the plate tectonics theory, since subduction and collision are considered to be involved in the emplacement of the intrusives. Moreover, Pertsev (1991) supports "that the progressive stage, the magnesian skarn can appear not only at the magma-dolomite contact, but at contacts of hard silicate rocks with dolomite, owing to replacement of either dolomite or any other magnesian rock (.exoskarn) or of a solid silicate rock (endoskarn)." Additional skarn types are reported by Lowell (1991). He records that a "reduced-W" type endoskarn is distributed over a large area adjacent to Late Cretaceous/Early Tertiary granites in the Salcha River region. Lowell further says: "The Salcha endoskarn association is, in many respects, comparable to major "reduced-W type" skarns such as MacTung and CanTung but lacks the high levels of Cu, Zn, Mo and Bi which characterize such deposits. The Salcha skarns, in contrast, are high in W, F, Sn, B, Be and Li which are typically found in greisen alteration associated with highly evolved tin granites". In addition to the skarn types the experimental work of Zaraisky (1991) on bimetasomatic calcareous skarns is considered of great genetic significance, as presented in "Experimental modeling of bimetasomatic calcareous skarn zoning". Zaraisky supports the following: "Contact interaction between granodiorite and marble has been studied experimentally using the author's method for direct modeling of the metasomatic processes. First, granodiorite and marble were crushed to a maximum grain size of 0.007 mm and packed densely one on top of the other into open plati-

num capsule (d = 5 mm, 1 = 50 mm). The contact between the marble and the granodiorite was at half the depth of the capsule. Experiments were run in autoclaves at Ρ = 1 kbar, Τ = 600° C, with solutions of different composition. After the run, the structure and composition of the discrete bimetasomatic zones formed on the contact were analyzed by electron microprobe. The overall thickness of a bimetasomatic column was 2-10 mm, with individual zones ranging from 0.002-0.005 mm to several mm. The contact rocks have been found to interact in all runs but skarntype bimetasomatic zoning developed only in the presence of near-neutral chloride solutions (NaCl, KCl, CaCl2, MgCl2). Active Fe leaching from femic minerals of granodiorite into chloride solutions promotes its accumulation in the contact zone and, as a result, the development of pyroxene-garnet skarn parageneses. Bimetasomatic zone sequences, their mineral composition and material transport within the experimental columns have been found to coincide closely from the Tuijinsky deposits in the Urals. The occurrence of skarns in contact aureoles of intrusions is explained not only by high-temperature conditions of formation but also, in all probability, by the chlorite nature of magmatic fluids." Zaraisky's experimental work, in addition to providing a model for bimetasomatic skarn type formation (zoning), also emphasizes the role of diffusion, leaching and, above all, the role of chloride fluids which according to him are magmatically generated. (It is possible though that his model could be applicable to the basic front hypothesis of skarn formation, see page 132). Additional to Zaraisky's work concerning bimetasomatic calcareous skarn zoning, study cases will be presented, reporting skarn zoning in calcareous, dolomitic and pelitic rocks. Furthermore, in cases, comparisons will be presented between Ca- and Mg-skarn zonation. Since skarn zonation depends on a wide range of factors, the resultant mineralogy and mineralogical sequence can indeed be complex. From the plethora of studies available, only few will be presented, aiming not to exhaust the subject but to present some characteristic cases. In 1991, Newberry reported that "Zoning and mineral composition data from 13 scheelite-bearing skarns in the Sierra Nevada area are employed to show fundamental differences between skarns hosted by relatively pure and impure carbonate lithologies. Skarns formed in relatively pure carbonate show clear-cut mineralogical zoning outward from granite; calciumpoor (subcalcic) garnet + quartz to granite + salite (± scheelite) to idocrase + wollastonite (± scheelite + garnet + pyroxene)... Skams formed in impure carbonate show irregular mineralogical zoning and generally irregular mineral compositional zoning (on both a deposit and thin section scale), although subcalcic garnets do show regular patterns of decreasing calcium content with increasing proximity to granite (deposit

scale) and garnet rim (thin section scale). Calc-silicate compositions vary widely within individual deposits compositional ranges of 10-95% andradite and 5-85% hedenbergite are common." In contradistinction to the Sierra Nevada skarns described by Newberry, Aksyuk (1991) took into consideration the "depth" factor in particular, and reported the following mineralogy and sequences for magnesian and calcareous skarns. "Various facies of magnesian skarns such as hypersthene, diopside, periclase one with mervinite, monticellite and calcareous skarns such as wollastonite, wollastonite-gehlenite, larnite and others are mainly related to their formation depth which decreases from hypersthene to periclase-monticelliie for magnesian skarns, whereas for calcareous ones, from wollastonite to larnite and tilleyite facies. Values of these depths depend on fluid composition changing from essentially C0 2 at an early stage to essentially HjO in late stages and from activities of Ca and Mg in fluid related to the stability of carbonates and C0 2 regime." In contrast to the cases described above, Pertsev (1991) supports the following concerning basic in composition Mg skarns and their exhibited zonation. "Magnesian skarns are a group of high-temperature, contact-metasomatic rocks that differ from common, or "calc" skarns in chemical and mineral composition, which approach those of basic and ultrabasic rocks. They originate in the contact aureole of magmas or cooling igneous bodies, because silica-rich magmas interact with rocks that are under-saturated with respect to silica but rich in magnesia (dolomite, magnesite, brucitite, etc.). This interaction is made with the help of fluids that carry rock-forming components and are responsible for contact metasomatism. Like other types of contact-metasomatic rocks, magnesian skarns exhibit metasomatic zonation. The most common zonation type - metasomatic column - forms at the contact between a granite melt and dolomite, and the following zones arise: granite —> clinopyroxene + bitownite fassaite + spinel —> forsterite + spinel —> calciphyre (calcite + forsterite + spinel) —> dolomite. Podlessky et al. (1991) present the following mineral assemblages and zonation for magnetite-bearing magnesian skarns of Mongolia. "Iron ore bearing magnesian skarn of Mongolia formed by selective magmatic replacement to Precambrian-Cambrian dolomites by Paleozoic granitoid intrusions. Major rock-forming minerals of magnesian skarns of the magmatic stage are clinopyroxene, spinel, forsterite, plagioclase and dolomite. Assemblages of the postmagmatic stage contain phlogopite, clinohumite, xanthophyllite, pargasite, tremolite, brucite, serpentine, chlorite and calcite. Irregular lens-form contact bodies of the magnesian skarns are characterized by zonal structure: granitoid / periskarn (may be absent) / spinel-pyroxene skarn / pyroxene-spinel-forsterite skarn / forsterite calciphyre / dolomite." 131

In contrast to the calcareous and dolomitic rocks which, as mentioned, may show advanced cases of skarnification and zoning, comparably pelitic rocks show metasomatic zoning sequences when skarnified. Lummen and Verkaeren (1991) support the following: "Two different metasomatic zoning sequences are commonly found in skams formed in pelitic rocks, namely: I) schist / biotite zone / amphibole zone / pyroxene zone / garnet zone, and II) schist / biotite zone / pyroxene zone / garnet zone. Sequence I is found in Mg-rich environments whereas sequence II occurs in Fe-rich ones. The iron in solid solution with magnesium is thought to be responsible for the absence of the amphibole zone in sequence II. The process of skarn formation in pelitic rocks has been modeled in the system K p - CaO - FeO - MgO - A1203 - Si0 2 - R , 0 with CaO and K^O as perfectly mobile components." Skam bodies have been considered as pneumatolytic metamorphic bodies formed at the contact of an intrusive magmatic body in calcareous rocks. Goldschmidt (1911) in his classical work Die Kontaktmetamorphose im Kristianagebiet believes that Fe and Mg were added to the calcareous rocks from the intrusive. He also describes the role of CI for the transportation of Fe and thus, his concept of gaseous "emanations" eventually developed into the concept of fluid introduction from the intrusive into the calcareous country rock. The role of fluids in metallogeny will be discussed in Chapters 36, 37 and 60 in the sections "Some aspects of the role of fluids in metamorphic ore". In this context, the change of fluids with a decrease of Τ in skarns should also be considered. Concerning fluids in skarns, Zheng Ming Hua (1991) reports "Hydrogen and oxygen isotopic compositions show that the ore-forming fluid is magmatic water mixed with meteoric water." He furthermore adds, "The analysis of gas-liquid inclusions in quartz shows that the ore-forming fluid contained high salinity, and high CI" content (relative to F" content). Metallic elements might have migrated in chloride complexes". Lowell (1991) also supports that "...the chemistry and petrographic features of moderately altered scapolite-bearing skarns are used to quantitatively document major and trace element mobility and track compositional changes in the metasomatic fluid as a function of falling T. The molecular quantity F + t^O* is used as a "metasomatic index" to show that (1) essentially all major elements were mobilized subsequent to the prograde thermal climax; (2) CaO, P 2 0 5 , Si0 2 /Al 2 0 3> CaO/MgO + FeO and Ca0/Al 2 0 3 increased during retrograde alteration; (3) MgO, FeO, Ti0 2 , Si0 2 , A1203, FeO/MgO and I^O/NajO decreased during retrograde alteration; (4) LREE were depleted relative to HREE during alteration; (5) W, Sn, Be and Β additions accompanied F-metasomatism; and (6) retrograde alteration was accompanied by a volume increase proportional to F + H 2 0 + in the resulting rock. Petrographic relationships, chemical data, and equilibrium con132

straints in the Ca0-Al 2 0 3 -Si0 2 -C0 2 -H 2 0 system are used to show that the initial alteration phase (scapolite) required Τ > 500° C and X C02 > 0.10 at Pt = Pf = 2Kb. Sequential appearance of margarite and vesuvianite + quartz indicate that substantial though not necessarily continuous reductions in both Τ and XCQ2 accompanied the early phase of retrograde alteration. The later appearance of calcite, fluorite and scheelite marks a reversal of the falling XCQ2 trend and signals the arrival of a fluid with greisen traits." Considering the theoretical outline given above and in the attempt to present the main topics and aspects (controversies), it should be pointed out that both the presented outline and the study cases which will follow, favour a magmatic-metasomatic interpretation for the genesis of skams (with the exception of topic No. 5 of the outline). However, the root of the interpretation that skams are magmatic-metasomatic-pneumatolytic products of an intruding magma as it reacts with country rocks (mainly limestone or dolomite) is to be found in "Die Kontaktmetamorphose im Kristanagebiet", Goldschmidt (1911). Nevertheless, despite the predominance of views bases on the concept that skams are magmatic-metasomatic bodies formed from emanation fluids and solutions of a consolidating granitic magma, the author believes that a reconsideration of our views on the genesis of skams is necessary, based on the following points which could form the framework for a new interpretation of the genesis of skams. (i) Are "intrusive" granites magmatic or tectonically mobilized diapirs or "malakton" of ultramorphic origin? (ii) Can skam bodies and their metallogeny be interpreted as basic front mobilizations released under granitization and particularly by the assimilation of pregranitic phase remnants of which are the xenoliths in the granites? (iii) Is the reconsideration of interpretations concerning the magmatic-metasomatic origin of skam bodies possible? The answer to these basic questions is not only complex but each answer is a vast subject (outside the scope of the present book) which has been treated in a number of contributions by the Metasomatic-Transformist School (see p. 118). For question (i) the reader is referred to Chapters 21 and 22 and also to another publication of the author: Atlas of the textural patterns of metamorphosed (transformed and deformed) rocks and their genetic significance, Augustithis (1985). On the basis of research work (Augustithis, 1993) in Atlas of granitization: textures and processes, granites exhibiting these textural patterns (blastic-endoblastic growths), synantectic-symplectic intergrowths and all the metamorphic-metasomatic traits which are non-deuteric but are metamorphic-ultrametamorphic, while intrusive types are diapiric or mobilized ultrametamorphic "malakton" (see Drescher-Kaden, 1982).

For question (ii), the reader is referred to the papers of Reynolds (1947), entitled "The association of basic 'front' with granitization", and "A Front of metasomatic metamorphism in the Dalvadian of Co. Donegal". Also in support of the basic 'front' related to granitization, the reader is referred to Chapters 25 and 26 of Atlas of the textural patterns of granites, gneisses and associated rocks, Augustithis (1973), in which the skarns of the Intrusive Young Granite of Kimmeria, northern Greece, and its metallogeny is extensively discussed. Chapter 4 of the Atlas of the textural patterns of basic and ultrabasic rocks and their genetic significance, Augustithis (1979) is also recommended to the reader. On the basis of these studies, it is suggested that the release of Fe-Mg from intrusive granites might be due to the assimilation of a basic pre-granitic phase of which xenoliths are relics. To assist with question (iii), the reader is referred to Chapter 3 of the Atlas of metamorphic metasomatic textures and processes (Geochemistry - Isochemical versus allochemical; Augustithis, 1990), and also to Chapter 4 (Metamorphic processes) of the Atlas of the textural patterns of metamorphosed (transformed and deformed) rocks and their genetic significance, Augustithis (1985). From these studies, the granitic emplacement and the perigranitic metamorphism and metallogeny including skarn formation are seen as caused by ultrametamorphism-granitization in the realm of metamorphism, an integral process of which was the emplacement of the "malakton". Furthermore, as a result of metasomatic differentiation, (Augustithis, 1990), pegmatites (Augustithis, 1962) and aplites (Drescher-Kaden, 1974) are seen as exudations of the granites and gneisses under ultrametamorphism with the operation of solutions. In support of the involvement of country rocks in the metallogeny of skarns and non-magmatic derivation of metallic components comprising the skarns, and in addition to Oen's interpretation (1968), that Mg and Fe were mainly derived from the country rock, as well as very significant geochemical studies by Tu Guangzhi (1991), support a non-magmatic derivation of Pb-Zn and probably Mn in skarn of north China. Tu Guangzhi states the following: "For skarn deposits, little doubt has arisen as to the magmatic origin of the ore-forming metals. Pb isotopic and mineralogical studies (sulphides and skarn silicates) of two Pb-Zn skarn deposits in north China, however, give strong preference to a non-magmatic local source for Pb and probably also Zn and Mn at least partially. It is postulated that the ore-forming metals were originally enriched but dispersed in the sedimentary formation. Later granitoid intrusion simply acted as a heat source to remobilize and reconcentrate the metals, thus giving rise to a skarn deposit." In addition, Zhu Shangqing et al. (1991) report the following (partly supporting a non-magmatic derivation of some ore minerals in skarns): "The Tongguanshan copper deposit is one of the best known deposits in the lower and middle Yangtze copper and iron metallogenic belt.

According to the occurrences, alterations and mineralogical assemblages, the deposit can be subdivided into three types of mineralizations: The upper mineralization which occurs in the upper part of the contact zone between the limestone and the intrusive body, is considered as a conventional skarn deposit. The middle mineralization occurs in the stratum and forms the major part of the deposit. Ores in this part are laminated, banded, colloform and framboidal, and stratabound. It shows horizontal zoning in an order of magnetite-pyrrhotite-pyrite-kappa-pyrite from the contact zone outward. Of these, pyrite, especially kappa-pyrite is the primary product of sedimentary and diagenetic origin, and magnetite and pyrrhotite are products of sedimentary and diagenetic origin, and magnetite and pyrrhotite are products of metamorphism and/or hydrothermal alteration which had preserved the relic structures after the original pyrite. But post-magmatic hydrothermal derived magnetite, pyrrhotite and pyrite can also be observed in the contact zone, therefore, there are two sorts of pyrite-magnetite series in the middle mineralization." Bearing in mind the above considerations concerning the genesis of skarns and the possible alternative interpretations, a number of study cases based on orthodox views will be presented, less for reasons of their genetic interpretations but for their value in describing mineral assemblages and the chemistry of some of these skarns. In this connection, some study cases will be presented which are in accordance to magma intruding into the country rocks. Piirainen et al. (1967) suggested that primary skarn iron ores are due to liquid immiscibility caused by the reaction of parent magma with dolomitic country rock. According to them, the C 0 2 produced during various skarn forming reactions is the chief cause for the unmixing of silicate and the iron oxide phase. As a corollary to the preferential relation of skarn/dolomites, Shabynin and Zarevich (1967) support that skarns derived from limestone differ in chemical composition and mineral association from skarns derived from dolomites. It appears that ore mineralization tends to be concentrated selectively in skarns derived from dolomites. The selectivity of mineralization (metallogeny) to dolomites during skarnification is perhaps the result of increased Mg mobilization of the basic front to certain limestone layers or masses. It is noticeable when initial carbonate layers later skarnified are interbanded with quartzites (which have not been affected by the basic front skarnification). Lundberg (1967) described sulphide-bearing skarn iron ore of Sahavaara, northern Sweden, as forming a tabular body concordantly enclosed in Precambrian quartzites, phyllites and graphite schists. The ore is associated with a dolomite horizon which it partly replaces. The ore mainly consists of magnetite but pyrrhotite and pyrite are common. 133

Furthermore, regarding skarnification on the basis of a granitic magma intrusion, Bezsmertnaya and Gorzhersky (1958), in the study entitled "Near-ore alteration of polymetallic deposits of the Rudny Altai (USSR)" support that in this particular deposit three thermal metasomatic stages are distinguished: high temperature skams; intermediate temperature chloritolites; sericitolites, listwenites. According to them, this is the result of prograde metamorphism, again based on a magma intrusion. In contrast to the descriptive approach concerning skarn formation and pyrometasomatic deposits, Park (1965), on the basis of experimental studies, attempts to explain a pyrometasomatic magnetite-pyrrhotite-pyrite-Ca, Mg, Fe-silicates deposit at Marmora, Belleville, Ontario, Canada. Thus, the phase geochemistry of the deposit is considered in relation to the experimental data on the Fe-S and Fe-S-0 systems, and according to Park, the ore deposit formed a closed system battered by the magnetite-pyrrhotite-pyrite assemblage which determined the values of PS2 (105 atm.) and P0 2 (10 20 atm.) at the calculated equilibrium temperature of 495°C. It seems rather difficult to reconcile metasomatic mobilizations - an open system - with the concept of an ore deposit formed as a closed system. Considering the different aspects of skarns, the contribution of Stolyarov (1964), "Hypogene anhydrite at the Alekseyevka copper ore deposit, central Urals" is of interest: skarn copper deposits at Alekseyevka contain anhydrite in the deepest part. They must have been formed at comparatively high temperatures, at least in the mesothermal range. Another interesting case is presented by Kalinin (1961) of the iron deposit at Tayat, Eastern Sayan. According to him, the deposit is "a contact metasomatic type", about 700 m from a granodiorite intrusive into arkose and tuffaceous sandstone. Hydrothermal alteration consists, in order of (1) hornfels formation, (2) albitization, (3) scapolitization, (4) magnetite deposition and (5) formation of veins of apatite and diopside. Furthermore according to Kalinin, chemical hypotheses indicate ionic chloride. Complexes require too low a pH for the transport of iron but soluble double salts of NaCl FeCl3 or 2NaCl FeCl2 may be stable under ore forming conditions. The suggested mechanism of Fe transportation attempts to explain the fact that the "contact-metasomatic deposit" is in this case at a distance of 700 m from the granodiorite. In contrast to the skarns described, associated with granites, Shei Kwang-Hong (1965) described a case of iron skarn at the contact zone between ultrabasic rocks and calcareous siltstones. According to him, the skarns are formed of almandine, chrompicolite, scapolite, phlogopite, magnetite (including magnesioferite), serpentine and chlorite. The skarns are believed to be derived from ultrabasic magma. In accordance with Di Colbertaldo et al. (1967), the magnetite deposit of the Aosta Valley, Italy, occurs in a lens of serpentine intercalated between bedded Trias134

sic limestones and Jurassic shales. According to them, the formation of the deposit was related to the intrusion of a gabbro-peridotite magma involving three types of magmatic activity: liquid magmatic segregation, autohydration of ferriferous silicates and metasomatism of calcareous wall rock. The ore is classified as a liquid magmatic-pneumatolytic deposit of Mesozoic age, metamorphosed during Alpine orogeny. In contrast to these basic intrusions and the skarnifications produced, Augustithis (1979), in his studies of basic front mobilization in the aureole of the granite of Seriphos, Greece, reported dolomitization and serpentinization of marbles. In places, magnetite has followed epidote and forsterite crystalloblastesis (the forsterite might be serpentinized) in the marble; also in cases, a synkinematic mobilization of serpentine and magnetite has taken place. Sometimes within the granitic aureole, lievrite-hedenbergite interbanding occurred at Koudouros, Seriphos, where mega-colloform structures of interbanded hedenbergite/lievrite are observed. All these phases and structures are products of the basic front mobilization released from the granite (Augustithis, 1979). From the plethora of studies available on the gamut of metallic elements abundant in skarns, an approach will be made to present selected study cases which refer to some economic elements present in skarns, e. g., W, Sn, etc. Monseur (1956) described phases of mineralization in the Salsigne mine (Ande, France) where successive mineralization of wolfram, scheelite, mispickel, pyrite and chalcopyrite took place. The deposit is considered to be of the transitional pneumatolytic-hydrothermal type. The important point seems to be that the mineralization is considered to be formed successively, i. e., many phases of solution fluids operated. Baumann and Starke (1964) demonstrated on the basis of the hübnerite/ferberite ratio, that several wolframite minerals of different chemical history are placed side by side. According to them, at Pechtelsgriin, secondary ferberites formed by hydrothermal alteration are found in addition to pneumatolytic wolframites with an H/F coefficient > 0.3. The average H/F values of mined samples of wolframite/ferberite characterize the hydrothermal "degree of alteration" of the primary pneumatolytic wolframites. Also in this case, a multistage mobilization of W is believed to have taken place. In contrast, Boludan et al. (1964), on the basis of the H/F coefficient (average of 13 tests: 0.185), the wolframite of the Aue-Lauter Raum, Erzgebirge, is believed to be deep-pneumatolytic. In the following part, study cases will be presented of skarns and pyrometasomatic deposits where in addition to W other metallic elements are of significance. Anon (1963) described a tubular pyrometasomatic ore body developed in folded and faulted skarns, which has been controlled by an anticlinal flexure, a fault, and by a favourable chert horizon. The deposition of

diopside, garnet and quartz apparently followed that of diopside, garnet and quartz and evidently that of scheelite, but preceded the introduction of pyrrhotite, chalcopyrite and minor sphalerite. In addition, reserves are estimated to be > 1.2 million tons averaging 2.47% W 0 3 and negligible (< 0.2%) Mo, Bi, As, Sb, Pb, Sn and P. Nevertheless, the presence of Mo, Bi, As, Sb, Pb, Sn and Ρ indicate that in addition to W, other elements as well have been mobilized. Agard (1965) mentioned that in the pyrometasomatic W-Mo-Cu deposit of Azegour, Haut Atlas, Be is found in the idocrase rock. Thus, in addition to W, Mo-Cu and Be are also mobilized. Furthermore, Gabert and Vinken (1965) recorded that scheelite deposits of Sangdong, South Korea, occur as a flat, somewhat winding tabular ore body, parallel to the bedding of the enclosing Cambrian beds. Following the formation of skarn rocks during a silicate pneumatolytic phase, scheelite-bearing quartz veins were emplaced in the skarn zone during the later stages of mineralization. The scheelite deposit is part of a tin-tungsten province. The significance of this association is widely reported and will be commented in detail in Chapters 56 and 57. An interesting case of W-Cu-Ag ores is reported from the pyrometasomatic deposits of Tem Piute, Nevada. The ores are located within a zoned, garnet-pyroxene skarn adjacent to a quartz monzonite. According to Buseck (1967), within an inner zone of the metamorphic aureole, the characteristic assemblage is pyrrhotite-marcasite-magnetite with local concentrations of molybdenite and ilvaite. An outer mineralized zone is characterized by sphalerite with minor galena, bismuthinite, cosalite, galenobismuthite and native bismuth. Pyrite, chalcopyrite and scheelite are uniformly distributed through the skarns (fluorite constitutes the predominant gangue mineral). In addition to the pyrometasomatic deposit of Tem Piute, Nevada, where Mo is reported, studies by Wu Liren (1991) support that "The magnesioskarn, calcioskarn and postmagmatic acidic leaching alteration caused by these igneous intrusive bodies are well developed there. Only the acidic hydrothermal solutions derived from magmas which have formed the fine-grained pseudoporphyritic monzogranite and granite porphyry, with a high content of Mo (120 ppm) invaded skarns along fissures to form Mo-ore deposits at 240-330° C. Such an acidic hydrothermal solution has invaded the fine-grained pseudoporphyritic monzogranite and granite porphyry along fissures also, to form disseminated and fine-veined Mo-mineralization. Thus, the Yangzjiazhangzi Mo-ore deposits belong to a porphyry-skam type." Moreover, Podlessky et al. (1991), describing the complex paragenesis of skarns in Mongolia, stated: "the skams bear magnetite mineralization, borates (ludwigite, szaibelyite, fluorborite, kotoite), sulphides (mainly sphalerite, pyrrhotite, bomite) and, in places, molybdenite. The boron and sulphide mineralization

formed after magnesian skarns and magnetite mineralization." Annealed monoclinic pyrites, according to Buseck (1967), in the inner-zone ores are tentatively suggested to indicate a formation temperature of 455-510°C. Sulphosalts in the outer mineralized zone were probably formed at less than 235 ± 25°C. Furthermore, according to Buseck the noteworthy features of these deposits are their relatively low temperature of formation and the exceptionally high gradient across the aureole. In contrast to the cases described of scheelite in skarns, or closely associated with it, cases are reported of scheelite and related metallogeny as veins in the metamorphic aureole. Geffroy and Lafforgue (1957) recorded scheelite and native gold with pyrite, arsenopyrite, sphalerite, chalcopyrite and a Pb-Sn sphalerite, chalcopyrite and a Pb-Sn sulphide, in an aggregate of quartz, biotite and calcite as a metalliferous vein in the metamorphic aureole associated with a batholith of Monts d'Ambazac, Haute-Vienne. Gabert and Vinken (1965) described scheelite-bearing quartz veins formed subsequently to a skarn body. The scheelite quartz veins were formed during a pneumatolytic-katathermal transitional stage and the scheelite formation reached its peak during a hypothermal stage. (It should be noted that the fluorescent colour of the mineral scheelite was dependent on the molybdenum content which increased with the temperature formation.) Besides the recorded transitions between skarn pyrometasomatic and hydrothermal scheelite occurrences, Fonteilles and Machairas (1968) report scheelite, accompanied by pyrrhotite and lesser chalcopyrite, sphalerite and other sulphides occurring in garnet vesuvianite-hedenbergite skarns of a granitelimestone contact. The richest ores were formed in a hydrothermal stage that followed the formation of the skarn. In addition to the scheelite formation already mentioned, Wang Shufeng et al. (1991) support the following, concerning the prograde and retrograde formation of scheelite in the Shizhuyuan skarn, Hunan, China: "Patterns of ore deposition were intimately tied to original lithologies and metasomatic types. Some initial scheelite deposition accompanied high temperature metasomatism of marble and concentrated toward the marble replacement front. Most initial scheelite and bismuthinite deposition contemporized with the retrograde alteration of skarn. All the initial scheelite was remobilized and deposited with alteration." Furthermore, Lowell (1991) also describes a case of prograde and retrograde scheelite formation: "In the Salcha River area, Alaska, metamorphic lithologies were intruded by two major granite bodies which produced W-bearing skarns and extensive B-F metasomatism. The W-bearing skarns are almost exclusively restricted to fine-grained aluminous calcsilicate gneisses which during prograde thermal metamorphism, transformed to an assemblage of salitic pyroxene + calcic plagioclase + quartz + calcite. This assemblage inter135

acted with greisen-affiliated aqueous fluids to produce a distinctive retrograde alteration suite composed of scapolite + vesuvianite + scheelite + fluorite. In extreme cases, F-metasomatism destroyed most prograde plagioclase and quartz, eliminated scapolite altogether, and produced rocks with as much as 35% fluorite and 7% scheelite." In contrast to the skarn and to these pyrometasomatic deposits in general in which W was important, some study cases will be quoted which indicate Sndominant mineralization in skarn and in pyrometasomatic deposits overall - according to the magmatic concept. According to Adam (1960), the island of Billiton, Indonesia, forms a part of a belt of tin mineralization that extends from Burma through Malaysia into the Java Sea. The primary lodes consist of pneumatolytic replacement deposits in folded sediments of PermoCarboniferous age, which have been intruded by Cretaceous granites. The deposits are of pyrometasomatic origin with the main lodes composed of skarn rocks consisting of amphibole, pyroxene, andradite, ilvaite and iron sulphides. Cassiterite was never found in association with wolframite, although veins containing these minerals singly may occur in close proximity. The separation of W and Sn metallogeny as reported by Adam is in contrast to the usual association of W and Sn. In addition, Khazov (1967) in his studies of the Kitel Sn deposit in USSR, indicates that the Sn deposit is associated with an anticline north of Lake Ladoga. The core of the anticline consists of a Middle Proterozoic synorogenic intrusion of plagioclase-microcline gneissic granites and associated migmatites. The orebearing skarn contains three types of mineral associations: (1) magnetite with chalcopyrite, sphalerite and cassiterite, (2) magnetite-sphalerite with chalcopyrite and cassiterite and, (3) cassiterite (rarely). Ores poor in Sn and rich in Fe are related to the early gneissic granite intrusion and ores rich in Sn and Zn to a later intrusion of rapakivi granites. Furthermore, Watanabe and Hoshino (1991), considering tin behaviour and its implications for skarn genesis, maintain that "during the later phase of early metasomatic stage, tin was first incorporated into andradite-rich garnet, resulting in the formation of stannian garnet containing up to 4.5 wt% S n 0 2 under conditions of high f 0 2 , alkaline fluids, higher temperatures (> about 350° C), or some combination hereof. Malayaite would be formed somewhat later than stannian garnet, but earlier than cassiterite and other tin-bearing sulphides and sulphosalts under alkaline and possibly lower temperature conditions. During the later hydrothermal stage, tin was incorporated into cassiterite, stannite and tellurium-bearing canfieldite, under conditions of lower f 0 2 , possibly lower fS 2 , acid to intermediate fluids, lower temperature, or some combination thereof." In support of the orthodox metallogenic view in which skarn and pyrometasomatic bodies are due to 136

the influence of magmatic intrusion, Schröcke (1963) stated the following: "Calculations based on reactions important in pneumatolytic deposition of tin ores show that Ti, Si and Al cannot be transported in the magmatic gas phase as TiCl 4 , SiCl 3 and A1C13 because of their unfavourable equilibrium constants and small partial pressures. The partial pressure of SnCl 4 , however, especially at high HCl pressures, reaches an amount sufficient for the deposition of ores." However, in accordance with a pneumatolytic interpretation of cassiterite mineralization, Monseur (1955) discussing the Montredon and Salsigne cassiterite deposit, supported that the former belongs to the pneumatolytic phase while the latter is extended into the hydrothermal. Williams (1958) recorded that the tin-tungsten ores of the Murphy and Shepherd mine, Moina, occur in a series of vertical quartz veins within the contact aureole of a Devonian granite. The vein minerals include wolframite, cassiterite, pyrite, marcasite, magnetite, haematite, pyrrhotite, galena and scheelite together with quartz, fluorite, topaz, micas and laumontite. According to Williams, three stages of mineralization are distinguished which postdate the contact metamorphism. As it was the case with W, Sn also transgresses from skarn to hydrothermal mineralization. On the basis of the study cases presented, magnetite is recorded as an important mineral in skarns and in pyrometasomatic deposits in general. However, the presence of Fe and Mg as important elements is also explainable by unorthodox views, namely by the basic front release due to granitization. In this hypothesis, too, the association of polymetallic mineralization with Mg-Fe basic front mobilization is comprehensible. From the plethora of magnetite skarns only a few cases will be quoted to illustrate diverse mineral parageneses and cases of generation of parageneses (based on the concept of superimposed paragenesis). Leonard and Buddington (1961) described iron ores from the St. Lawrence Country, northwest Adirondacks, New York, where the deposits, all near axes of major synclinal folds, are of two types: (1) magnetite deposits in skarns and (2) magnetite deposits - with or without haematite - in microcline granite gneiss. Less important ore minerals associated with magnetite and haematite include pyrite, pyrrhotite, chalcopyrite, sphalerite, molybdenite, bornite, ilmenite, marcasite (?) chalcocite, covellite, vonsenite, löllingite, graphite and valleriite (?). The deposits are believed to be due to high temperature replacement. In contrast to the magnetite, pyrometasomatic magnetite paragenesis as described by Leonard and Buddington, Banäs (1965) discussed a case of superimposed paragenesis (see Chapter 51) from the Sudetes Mountains, Poland. Magnetite-sulphides-fluorite skarns occur in the Snieznik Klodzki complex and have several zones. The quartzite zone includes native metals, oxides, hydroxides, sulphides, selenides, sulphates and phosphates, among them cosalite, schapba-

chite, miargyrite, tiemannite, berzehanite, naumannite, aramayoite and stromeyerite. Furthermore, according to Banäs, two phases of mineralization are identified: the older skarn magnetite and the younger hydrothermal polymetallic fluorite phases. The magnetite mineralization of the skam type is considered a regional metamorphic development: the older regional skarns were then overprinted by a phase of contact skam. The study case by Banäs shows the complexity of processes. It also shows that multigenetic concepts may, in some cases, provide plausible interpretations. In the following, several additional study cases will be presented of skarn parageneses with superimposed later, mostly hydrothermal or hypothermal parageneses. In many cases it is believed that there is a transition from skarn mineralization to hydrothermal mineralization, where the intrusive might have been the source in both cases. According to the orthodox-magmatic interpretation, the pyrometasomatic mineralization was followed by the hydrothermal (hypothermalhydrothermal) phase. In contrast, as already mentioned, Augustithis (1990) regards the granite, the contact metasomatic and hydrothermal mineralization as the result of metasomatic differentiation by which granitization-granitic "malakton" emplacement took place. The skarns represent basic front release from the granite (the basic front release is due to xenoliths assimilation on granitization). For the skarn formation and the perigranitic vein mineralization, both the granitic emplacement and the wider involved geoenvironment contributed by processes such as lateral segregation and solution/fluid mobilization (Augustithis (1990) discussed the possible mutual interrelationship of fluids granitization, granites fluids supply). Considering the outlined theoretical controversy which is the consequence of the fundamental granite controversy, a number of cases of skam paragenesis with superimposed hydrothermal paragenesis will be further presented, discussing orthodox magmatic interpretations. Mochnacka (1966), studying polymetallic deposits from Lower Silesia, stated the following: "The mineralization in the Wolnosc mine, Kowary, is in the metamorphic granitic mantle of the Karkonosze Mountains, and occurs as skams and lenticles of magnetite ore and later unrelated polymetallic ore deposits in breccia and in carbonate veins. The principle ore in the polymetallic assemblage is pitchblende; it is associated with coffinite, liebigite, arsenopyrite, löllingite, tiemannite, clausthalite, rammelsbergite, tetrahedrite, smaltite and stromeyerite". According to Mochnacka, four stages of primary polymetallic deposits are distinguished: (1) an As-Co-Ni stage, (2) a pitchblende (nasturan) stage, (3) a sulphide-selenide stage, and (4) a carbonate stage. As already indicated, the polymetallic mineralization (represented by the stages mentioned) represents a polymetallic mineralization superimposed on the skam metallogeny.

Another case of superimposed paragenesis is described by Sims et al. (1958) from the Copper King uranium mine, Larimer County, Colorado. Pyrometasomatic sulphide magnetite deposits, consisting of pyrite, sphalerite, chalcopyrite and magnetite and a vein pitchblende deposit, consisting of pitchblende and associated minerals (uraninite, coffinite and U0 3 -rich pitchblende) constitute the two types of mineral deposits. Sims et al. provide age determinations and support a superimposed paragenesis for the pitchblende. Alpha-helium age determinations on the magnetite indicate its late Precambrian age; 206Pb/238U and 207 Pb/235U age determinations on hard pitchblende from the vein indicate an early Cenozoic age. Yudalevich et al. (1966), considering the skam polymetal mineralization in the central part of the Chatkal Range, which is related to the middle phase of evolution of the Chatkal-Kurama structural facies zone, consider it to be older than the rare metal mineralization which could be considered as a superimposed mineralization on the skam-polymetallic phase. (However, also in this case, the polymetallic mineralization might represent a superimposed phase on the skam mineralization.) Popov and Dokov (1963) contemplate that the johannsenite and rhodonite skams in Madan, Rhodopes, Bulgaria, were found before the sulphide mineralization. Several sequences have been distinguished by them, and in this case, the sulphide parageneses are superimposed on the skam. Furthermore, two types of ore deposits are described by Koroleva (1965) in the fault zones of AltynTopkan, USSR, the one being of Pb-Zn skam type, the other of Pb-Zn hydrothermal type. According to Koroleva, the concentrations of silver and bismuth increase considerably towards the end of the hydrothermal stage, and independent Ag and Bi minerals become abundant in late stage deposits. In this case, despite Pb and Zn being present both in skam and as hydrothermal, it may also be considered as a superimposed hydrothermal phase on the skam phase. However, due to multiple stages of mineralization as recorded by Koroleva, transitions might be regarded as probable. In addition to the reported Pb, Zn, Ag association with magnesian skams, manganoan skams (sometimes regarded to be a type of calcic skam) may also contain these elements. Zhao Yiming (1991) supports the following: "The distinguishing features of manganoan skams are: (1) a special manganese-rich calcic silicate association, such as johannsenite, Mnhedenbergite, rhodonite, bustamite, pyroxmangite, Mnilvaite, spessartine, Mn-actinolite, etc.; (2) they occur usually along the lithologic contacts or faults, fractures of carbonate wall rocks distal to intrusive bodies; (3) a relatively lower formation temperature, and infiltrational type of forming mechanism; (4) association always with Pb-Zn(Ag) sulphide mineralization and, therefore, manganoan skams may serve as important indicators when searching for Pb-Zn(Ag) deposits." 137

Selenium and tellurium are reported in skarns of the Altay-Sayan region, and as suggested by Vakhrushev and Dorosh (1966), a different behaviour is exhibited: "Different behaviour of Se and Te has been established in skarn deposits, including skarn-gold ore deposits on the one hand and vein deposits on the other. It is proposed to single out deposits of gold-bearing skarns as a separate genetic type or as an independent skarn-gold ore formation". In addition to Dorosh's proposal that skarn-gold deserves special treatment, Korobeinikov (1991) supports that "although the skarn-formation theory has been reliably worked out by Korzhinsky, Zharikov and their followers, the problems of correlation between skarn-formation and gold mineralization, and especially gold geochemistry in the contact-metamorphic and metasomatic processes, still need investigation. Autometasomatites of productive intrusions also need more attention. These are caused mainly by the vagueness of the following problems: problem of gold sources, forms of relations between mineralization and magmatism, metamorphism, metasomatism, metal fractionation in magmatic and metasomatic processes." Furthermore, Korobeinikov describes the gold-skarn formation geotectonic settings as follows: "Gold-skarn formation deposits occur frequently within carbon-bearing rocks of different crust structures in connection with the formation of syninversion-orogenic granitoid intrusions. The deposits are located in the folded eugeosyncline and, less frequently, miogeosyncline structures in Early and Late Paleozoic and Mesozoic periods of stabilization and tectomagmatic renovation, the structures being saturated with hypabyssal intrusions of contrast series of gabbro-plagiogranite and diverse batholite formations (Kuznetsov, 1964)." Soler and Ayora (1991) present an entirely diverse geoenvironment and tectonic setting for the gold-arsenopyrite skarn from L'Alt Urgell, central Pyrenees, Spain, and report that: "This is the first report of goldbearing arsenopyrite mineralizations in a skarn setting from the western Mediterranean Hercynian belt. They

138

are situated at the SW contact of the Andorra granodiorite intruding a thick pile of Devonian limestones. Skarns are mainly developed through carbonate bedding planes in the areas where the intrusion cuts these planes. Ore reserves are also controlled by Alpine thrust and Neogene faulting." In contrast to the prevailing opinions, Yang Minzhi (1991) indicates that not only Au and Ag but also Pt, Pd, Os and Rh have been enriched in skarn copper deposits. A rather special case of element presence is reported by Cioflica and Vlad (1970). They maintain that "B, Mo and Bi mineralizations are associated with skarns formed from hornfelsed andesitic and lamprophyric dykes and in the adjoining calcareous strata in the Codu nappe (comprising Permian-Jurassic rocks). Magnesian skams with boron minerals (ludwigite, kotoite, szäjbelyite) are genetically related to the andesitic hornfels; calcareous skarns related to the lamprophyric hornfels, carry molybdenite in the grossular, vesuvianite and diopside skarns while Bi minerals occur in the wollastonite skarns formed in the strata next to the dykes." In contrast to the pitchblende superimposed paragenesis on skarns, cases are reported of uranium as part of skarns. Welin (1964) reports that all the uranium mineralizations in Uppland, central Sweden, are connected to iron ores which mostly are of skarn character; their main ore mineral is magnetite. The occurrences are divided into two groups: (1) local uraninite disseminations in skarn ores; and (2) pitchblende, haematite and some simple sulphides in fractures and breccia with gangue of chlorite, calci te and quartz. In contradistinction, Whittle (1960) described an uraniferous skarn from Mary Kathleen, Queensland, Australia, in which fine-grained uraninite is disseminated through orthite-apatite enriched rock in a garnetiferous skarn. According to Whittle, the skarn and ore formation resulted from accessions of late magmatic emanations of a differentiated granodioritic intrusive emplaced in calcareous-magnesian metasediments.

Chapter 26

Pneumatolytic to HydrothermalHypothermal

Considering the hypothesis of granitic magma intrusion, successive phases have been recognized: orthomagmatic crystallization - pyrometasomatic (skamification) - pneumatolytic - pegmatitic - aplitic - hypothermal mesothermal - ?hydrothermal - ?epithermal telethermal. However, in contrast to these phase successions widely accepted for the metallogeny of granite from Cornwall and elsewhere, some of these phases have been detached from the concept of a granitic magma intrusion, e. g., epithermal is related to post-volcanic activities in many cases. Also skamification is often detached from a granitic magma intrusion and consolidation and is often related to porphyries and subvolcanic activities and, in cases, it is described to be related to basic intrusions such as ophiolites (serpentinites, etc.). In addition, the term hydrothermal is sometimes also detached and used in conjunction with volcanism or with remobilization of material under metamorphism. Low temperature "hydrothermal" solution may even be related to low temperature metamorphism or diagenesis. However, in most of the study cases to be presented here "hydrothermal solution" is attributed to a granitic magma intrusion. Whereas for economic geologists, the granitic magma consolidation is used as deus ex machina for solving most of the problems of mineralization (metallogeny), for most petrologists anatexis, granitization sensu stricto and ultrametamorphism-granite emplacement in general is better accepted. This deviation from the orthodox magma hypothesis is the merit of specialists on granitic rocks, who, sooner or later, having first rejected a basaltic magma differentiation of granitic magma, then - what is more important - realized that the huge masses of granitic complexes leave no other interpretation than an involvement of the upper crust. Even those who believe there are "granites" when confronted with the space problem and with the derivation of these huge masses of "mobilized" material, cannot get away from the interpretation that they are S-type granites. Augustithis in the Atlas of granitization textures and processes emphasized that the granites exhibiting the textural patterns presented in the granitization atlas "are products of granitization sensu stricto". Also Augustithis (1973, 1985, 1990, 1993) emphasized on the basis of comparative anatomy and textural analysis

that each granitic textural pattern is caused by a process which operated in the realm of metamorphism-metasomatism. Thus, textures and processes form an integral system - and should be judged on this basis. The author, however, in the second part of the present effort and specifically in the section entitled "Considerations of Hypotheses and Theories on Metallogeny" presents case studies (against his convictions) based on all hypotheses and interpretations in order to confront the student of mineralization with pluralistic information. In particular, as it was the case with skamification where the relationship of basic front and granitization was also mentioned, and also in the case of hydrothermal mineralization, two aspects based on granitizationultrametamorphism should be emphasized: (i) that hydrothermal mineralization may represent solution-fluids mobilization under metamorphism-ultrametamorphism, where lateral secretion and selective element mobilization is possible, and (ii) that in the realm of metamorphism, hydrothermal solutions can be produced due to friction processes (Drescher-Kaden and Heller, 1961). As already mentioned in conjunction with transition phases between skarn-types of mineralization and hydrothermal mineralization, similar transition mineralization phases are reported between pneumatolytic and hydrothermal. Bertolani (1962) reported that in the ore minerals associated with the ophiolites of Corchia (Parmessan Appennines) there is a cobalt mineral akin to the linnaeite group, probably siegenite or carrollite. The presence of this Co mineral together with chalcopyrite and intergrowth of sphalerite-pyrrhotite suggest a pneumatolytic-hypothermal genesis. Foissy (1966) reports that primary gold deposits of the Tsaratanana area, Madagascar, occur as disseminations and vein deposits in amphibole rocks and ferruginous quartzites. The gold mineralization is considered to have originated as a result of pneumatolytic and hydrothermal processes associated with migmatization and granitization of country rock. Monseur (1956) reports that successive phases of mineralization of wolfram, scheelite, mispickel, pyrite and chalcopyrite occurred at the deposit of Slasigne (Aude, France). The deposit is considered by Monseur as the transitional pneumatolytic hydrothermal type. 139

In contradistinction, Tumeaure (1960) considers the tin-silver deposit of Potosi and Oruro and the tin deposit of Liallagna, Huanuni and Morococala as xenothermal high temperature (400-500° C) and shallow depth, temperatures dropping to 100°C in the final stages of mineralization. According to Turneaure, typical minerals in the tin veins are quartz, cassiterite, bismuthinite, pyrrhotite, stannite, marcasite, pyrite, siderite, tealite and franckeite. In the tin veins tetrahedrite is abundant; other lead and silver sulphosalts, alunite and kaolinite are present. As already pointed out, the temperature drops from 400-500° C xenothermal to 100° C in the final stage of mineralization. Furthermore, Taylor (1969) emphasized the relationship of zones of subhorizontal quartz-feldspar pegmatitic veins with later hypomesothermal cassiterite-bearing lodes. Two structural lode types are recognized: those which originate outside the pegmatitic zones and split up or die out on entering them, and the more numerous but minor intra-pegmatitic lodes. In addition, homogenization studies of primary fluid inclusions in quartz, calcite and dolomite in the Hanes Fata Baii Municaceasea, East and West, Bridiser and Staniya deposits indicate temperatures corresponding to an upper mesohypothermal stage. The mean geothermal gradient of the mineralizing solutions, according to Borcos (1966) is estimated to 10°C/100 m and is the same for all studied deposits. Also Biste (1977) recorded a hypothermal Cu-As mineralization related to a Hercynian oblique fault south of Muravera, southeastern Sardinia, which contains a polymetallic sulphide paragenesis of arsenopyrite, chalcopyrite, pyrite, galena and sphalerite. The cases of hypothermal mineralization presented above and the parageneses reported, indicate a relationship with "intrusive bodies" or with faults, thus suggesting a derivation from a temperature source of some sort. In contrast to the study cases of hypothermal mineralization considered, study cases of mesothermal mineralization and parageneses will be presented, not to entirely cover the subject but to present some examples from the plethora of available studies concerning mesothermal and, in general, hydrothermal mineralization and parageneses. An interesting study case is presented by Sheridan (1964): "chalcopyrite, oxidized superficially to brown copper oxide or decomposed to malachite and azurite is associated with grey ore and bornite in a gangue of compact silica. Pyrite and small quantities of gold, silver and tin are also present." According to Sheridan, the mineralization persists to a depth of 2200 ft, where pyrite and tetrahedrite or similar sulphantimonides/sulpharsenides of copper occur. The deposit is unzoned and mesothermal in origin; it is independent of host rock lithology and follows the fault contact between a slate and a sandstone formation. Another case of mesothermal mineralization is reported by O. Gaspar da Cruz (1967) from Cerro do 140

Algard deposit, Beja, Portugal. The mineralization fills a system of fractures developed parallel to the strike of pre-Visean quartzites and slates. The main mineralization consists of pyrite with lesser arsenopyrite and löllingite; chalcopyrite, enargite, stannite and galena are also present. The average Co values of 1000 ppm in the pyrite and 5000 ppm in the arsenopyrite-löllingite are associated with Ni contents of 500 ppm for both assemblages. The deposit is classified as a mesothermal vein deposit. Idriceanu et al. (1965) support that the BuciumIzbita mineralization in the Vulcoi-Garabia eruptive complex includes pyrite and chalcopyrite with subordinate chalcocite, covellite, germanite, enargite and tetrahedrite. Investigation of fluid inclusions in minerals from seven different horizons up to 240 m vertically apart showed that the homogenization temperatures for quartz ranges from 270° C to 300° C, suggesting an upper mesothermal character for the deposit. Furthermore, Idriceanu et al. emphasized a significant parallelism between the concentration change of elements such as Cu and Ge and the frequency and temperature curves: the hydrothermal process is believed to have had a pulsating character. Examples supporting the derivation of hydrothermal solutions from a magmatic source will be quoted presenting orthodox views on hydrothermal mineralization. Fedink and Kusnir (1967) consider a group of ore deposits in the area of Cho Dien as representing the most extensive Pb-Zn mineralization of Southeast Asia. The ores are composed principally of cadmium and galium-bearing sphalerite, argentiferous galena and pyrite. Host rocks are weakly metamorphosed Devonian limestones. Primary hydrothermal veins, bedded layers, stratiform deposits and secondary assimilations filling solution cavities in the limestone are present. Spatial relationships indicate a genetic relationship with granitoids of the Tam Tao massif (Vietnam). Coleman (1957) supports that "the suite of metallic minerals in the extensive shear zone system on the west shore of Yellowknife Bay, Canada, appears to have been formed during three separate periods of mineralization from hydrothermal solutions having a magmatic origin". Pyrite and arsenopyrite were formed during the first period, sphalerite, chalcopyrite and pyrrhotite in the second mineralization. According to Coleman, during the last period at temperatures below 350° C and fairly low pressures, the minerals formed included galena, stibnite, gersdorffite, ullmannite, gudmundite, marcasite, sphalerite, pyrrhotite, chalcopyrite, boulangerite, meneghinite, jamesonite, bournonite, tetrahedrite and berthierite with native lead and antimony. Some of the gold may have been deposited in solid solution in pyrite and arsenopyrite. In some cases, reaction of antimony-bearing solutions with pre-existing gold has given rise to autostibnite.

Stumpf! (1960) also related hydrothermal mineralization to granitic intrusion. According to him, the Bushveld granite produced a fault system which provided the channels for the entry of the ore forming solutions into the host quartzites. Orthoclase, cassiterite, molybdenite and bismuthinite are found, indicating a high temperature hydrothermal paragenesis. In contrast to Stumpfl, who emphasized a definite relationship of fault system and granitic intrusion, Bernard and Dudek (1967) suggest that spatial or temporal relationships between mineral and adjacent plutons are only rarely evident in the case of the hydrothermal mineralization following shear zones and associated with deep-seated tectonic lines of weakness in the massif. Again in connection with tectonic effects, the hydrothermal mineralization is related to brecciated and highly crystalline limestones. The mineralization exhibits a pyrite-sphalerite-galena-tetrahedrite-albandite sequence. Sphalerite is widespread and contains Cd, Fe, Μη, In, Ge and Sn in minor amounts. It is concluded by Lazär (1966) that the mineralization occurred at relatively high temperatures. Lukas (1970) describes lens-shaped interrupted ore veins associated with the youngest system of four fault planes in Schleining, Burgenland. From these joints, apophyses of small ore veins concordant and disconcordant to the planes of the country rock (a low grade phyllitic marble Palaeozoic or Mesozoic) branch off. Arsenopyrite, cinnabar, chalcopyrite, ilmenite, magnetite, pyrite, sphalerite, marcasite and a number of oxidation minerals occur. According to petrographic, petrofabric and geochemical data, the hydrothermal antimony concentrations are controlled by the young discordant joints. The present cases where the hydrothermal mineralization is associated with fault, shear and joint systems, indicate that solutions ascended along these pathways but their origin - fluid-solutions of magmatic derivation or fluid solutions as a result of friction hydrothermal solution generation in a Bereich (field) of metamorphism/ultrametamorphism - is open to question. In the following, cases will be recorded of superimposed hydrothermal parageneses (see also Augustithis, 1982). Salvadori and Zuffardi (1961) describe a case in the Muru-Mannu, Oridda area, where studies on skarn and limestones revealed the presence of various phases of metamorphism, often superimposed on and followed by a mineralizing hydrothermal cycle. In contradistinction, Andrushchuk et al. (1967) describe a massive sulphide in the Kizil Dere, northern Caucasus, in silty shales along a fault. Macro/microscopic investigations show the following parageneses: (1) quartz-pyrrhotite, (2) chalcopyritepyrrhotite including early sphalerites, quartz, cobaltite, löllingite, etc., (3) quartz-sphalerite-chalcopyrite, (4) sphalerite-chalcopyrite-galena, (5) quartz-carbonatepyrite. Cinnabar, realgar, baryte, fluorite and datolite

were deposited hydrothermally (?superimposed) after the main period of ore deposition. The examples quoted above and the cases described support the multiple nature of the hydrothermal processes and the fact that solutions may be generated through a tectonogenic event and not necessarily due to multiple magmatic intrusions. Firsov (1963) reports that "contact metamorphism in the Yana-Kolyma ore-belt, north eastern USSR, was caused by granitoid bodies of Cretaceous age. The primary segregation of gold in quartz veins range in size from parts of a millimeter to several centimeters or more". According to Firsov, as a result of recrystallization, the large segregations have disintegrated into microcrystals with a diameter ranging from 0.005 to 0.001 mm, less commonly to 0.003 mm. Gold crystals in veins formed under mesothermal and epithermal conditions show a peak abundance of [113] crystals. The formation of [210] crystals of gold is possible at higher temperatures. As pointed out, the studies of Firsov suggest mesothermal crystallization for gold crystals exhibiting [111] octahedral forms on a topomorphic basis. In contrast to the cases where solid inclusions homogenization was used for determining temperature conditions of formation, topomorphism is, in the case of Firsov's studies, applied to suggest temperature conditions of formation of the vein deposit. In contradistinction to the mesothermal deposits mentioned, Di Colbertaldo and Omenetto (1962) on the basis of ore and gangue mineral paragenetic associations, consider the metalliferous veins of Sos Enattos (Sardinia) as meso-epithermal veins of Late Hercynian age. According to them, the main ore minerals present are: sphalerite, galena, chalcopyrite, pyrite, linnaeite, and chloanthite-smaltite and the main gangue minerals are siderite, calcite and quartz. In contrast to meso-epithermal or mesothermal deposits, Cevales (1961) describes several deposits from Gran Paradiso, Italy: "all show a sulphosalt mineralization of Pb, Cu, Ag with subordinate Hg and Au. The gangue is usually siderite, rarely quartz. The most common minerals are: chalcopyrite, pyrrhotite, valleriite, galena, pyrite, arsenopyrite, jamesonite, chalcostibite, tetrahedrite, freibergite, sphalerite, covelline, chalcocite, bournonite, native gold". All the deposits from Gran Paradiso are considered to belong to a single metallogenic province of epithermal-subvolcanic character, genetically related to mobilization and lateral segregation of the metalliferous content of the crystalline massif. Another type of epithermal mineralization is described by Tan (1958) from Szehuang-Tzeping area, Taipeihsien, Taiwan. Sooty black cryptocrystalline pyrite and marcasite and sulphur occur in a gangue of quartz, opal and clay minerals on the slope of a Pleistocene volcano: the country rocks are mainly andesite flows with interbedded agglomerate lenses. The shallow-seated occurrence, the abundance of opal, and the representative textures of crustification, brecciation 141

and vugs, indicate that the deposits are epithermal in origin. Furthermore, the presence of diffusion banding, reniform and botryoidal textures suggests a colloidal hydrothermal solution. Another case study of epithermal mineralization and a stage of mineral formation is described by Shchlegov (1962). According to him, the ore deposits of Transbaikal are classified into two groups which show stages of mineralization that follow a definite sequence. The first group incorporates mineral deposits of molybdenum, tungsten, tin, polymetallic character, and gold. The second group includes epithermal deposits as veins of mineralized chalcedonic quartz which bears minerals of tungsten, antimony, mercury, gold and fluorite. It is interesting that both the polymetallic mineralization and the epithermal have geochemical representation of the W-Mo group of elements (interrelated according the empirical laws of the periodic system). Another case study indicating the wide spectrum of geoenvironmental conditions which can result in the formation of epithermal mineralization is described by Zucchetti (1966). He maintains that the Hg deposit of Almad6n in Spain is located in an intensely metamorphosed Silurian-Devonian complex of subvertical argillites and quartzites. The ore is contained in three quartzite bands, slightly below the top of the Silurian series and within parallel basalt sills of probably Tertiary age. According to Zucchetti, the deposit is essentially epithermal, probably telethermal and hypabyssal, the ore being distributed by impregnation. In contrast to epithermal study cases (see also Chapter 26) "hydrothermal" solutions are considered responsible for the formation of a number of ore deposit types. Kraume (1960) considers the Pb-Zn-bearing Rammelsberg deposit near Goslar, upper Harz, as a sedimentary submarine-hydrothermal formation of the Middle-Devonian. According to Kraume, in addition to main ore bodies, grains of blende in sandstone and fossils are replaced by chalcopyrite. In addition, the ore bodies near Bleiberg (Kreuth, Carintha) are described by Schulz (1968) as submarine

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synsedimentary in origin. The ores contain the paragenesis sphalerite, wurtzite, galena, marcasite, pyrite, baryte, fluorite, quartz, calcite, dolomite and clay minerals. The ores are rhythmically layered and occur in the upper 60 m of the Landinian limestone-dolomite series. The ores are derived from low temperature hydrothermal solutions. However, the deposits exhibit general sedimentary features, such as external and internal denudation by mechanical erosion, attack by solutions, metasomatic processes and syndiagenetic ruptual and steady formations. Another case where sedimentary and hydrothermal processes occurred is described by Seguin (1978). According to him the Michipicoten range in Ontario has a dual origin: sedimentary deposition of siderite, pyrite and a little magnetite followed by a hydrothermal phase in the 200-270° C range which produced ankerite, pyrrhotite and other sulphides. However, the origin of a deposit might be uncertain, e. g., whether it is sedimentary (synsedimentary) or hydrothermal (secondary mobilized). As a result of these possible contrary interpretations, a controversy concerning a number of significant deposits has evolved (see Chapter 26). As a corollary to the controversy, a study case will be presented based on the work of Saager (1969), entitled "The relationship of silver and gold in the Basal Reef of the Witwatersrand System, South Africa", concerning a disputed deposit which by some researchers is believed to be synsedimentary and by others epigenetic (hydrothermal). According to Saager, Ag/Au ratios of individual gold particles are constant at about 0.0812 but ore samples give higher values and a large but really systematic variation. This is interpreted to indicate homogenization of detrital gold during transportation and sedimentation so that part of the original Ag content has been redistributed. Also the palaeogeography of the sedimentation and the effects of later metamorphism and pseudohydrothermal activity are considered by Saager.

Chapter 27

Controversies - Various Aspects of Metallogeny

(a) General (with Reference to Pt

Nuggets)

The progress of geoscience has often been marked by the following series of developments: a hypothesis is postulated, it attains the status of a theory, sometimes it evolves into a rule, but hardly ever it does become a law. Often, after reaching the status of a theory, it may attain the universal state, then, very often, it is unable to explain the new challenges that are put forward. As new facts accumulate, it enters a phase of being challenged and doubted and eventually, in many cases, it passes into oblivion. However, this development is not necessarily unavoidable. Cases are known where a hypothesis reached the universal state and when challenged by new facts or interpretations, it was able to respond by modification. Another very common development is that a hypothesis attains theory status and reaches the universal state, but then its basic concept may fail. However, and this is most interesting, whereas the main concept of this theory becomes obsolete, side branches of it or derivative concepts may attain an independent theory status and survive the new challenges - after, of course, being detached from the initial parent hypothesis. An example of such a case is the basaltic magma differentiation hypothesis. It evolved into the magmatic differentiation theory which attained its universal state in which it explained all "igneous" rock types as derivatives of a differentiated initial basaltic magma and subsequently was challenged by new accumulating facts and interpretations. Examples for such challenges were the huge granitic massifs in areas and the absence of equivalent basic and ultrabasic rocks that would justify basaltic magma differentiation; the fact that granitic and gneissic massifs were explained as products of granitization or anatexis and that dunites with associated chromite deposits would require tremendous quantities of initial tholeiitic basaltic magma to be produced. The problems (and many more, e. g., endoblastesis in granites, etc.) resulted in the basaltic magma differentiation hypothesis being seriously challenged. It proved unable to provide explanations and so passed eventually into disuse. Side branches of this hypothesis, however, did survive - e. g., the hypothesis of a more basic-ultrabasic magma which responded to the challenge of inter-

preting chromite ores as differentiation products associated with ultrabasic bodies, such as dunite ores. Another more recent example is the plate tectonic hypothesis which, after merging with the geosynclinal hypothesis, passed to theory status and subsequently reached the universal state since plate tectonics was used to explain not only global tectonics and crustal movement but also provided interpretations to problems, each one being just as complex as plate tectonics itself. Metamorphism was explained as due to the subduction of sediments (subduction metamorphism), the ultrabasic complexes and their related metallogeny as oceanic (lower) crust/upper mantle obduction; the genesis of granitic massifs as due to collision of moving crustal plates and metallogeny as directly related to plate movement and their junctions. Thus, opposition to plate tectonics came from each of these fields. Subduction metamorphism failed to be compatible with all the detailed petrographical and experimental work that had provided information and complex, sophisticated explanations for decades (see Augustithis, 1985, 1990). In addition, ultrabasic complexes are rather interpreted as due to mantle diapirism and huge plateau basaltic flows in the central parts of the tectonic plates (Augustithis, 1978, 1979) are other interpretations and facts incompatible with the plate tectonic hypothesis. The existence of old (e. g., Precambrian) gneissic granitic massifs and their transitions, the endoblastic growths and their metamorphogenic traits (Augustithis, 1973, 1993) and, in general, the textural patterns of granites are incompatible with an interpretation that granitic massifs are formed as consolidation of melts produced due to collision of crustal plates. In contrast to the basaltic magma differentiation theory which failed to meet the challenges mainly because of its central "concept" - a melt of certain composition moving upward, producing all kinds of rocks on differentiation (e. g., dunites, gabbros, granites, rhyolites) - proved to be incompatible with facts and unrealistic. The plate tectonic theory on reaching the universal state, was challenged on all fronts. The global tectonic concept of the plate tectonic hypothesis was not only challenged by the "expansion of the earth" hypothesis but also by a plethora of new hypotheses and interpretations (see Critical Aspects of the 143

Plate Tectonic Theory, Beloussov et al. (eds.), Theophrastus Pubs, Athens, 1990). Thus, and this is perhaps of interest, plate tectonics has failed to provide adequate interpretation in metamorphism, in petrology, in metallogeny and, above all, in interpreting many global tectonic features. Whether plate tectonics will survive future challenges is difficult to say. Furthermore, whereas metallogenic interpretations (based on plate tectonics) were very "fashionable" during the attainment of universal state of the plate tectonic theory, they became less "fashionable" in the late 80s. In an attempt to mention controversies which played a significant role in the development and evolution of geosciences for historical reasons, one should not omit the neptunists and plutonists and how the controversy passed into oblivion due to the progress of geosciences and the differentiation of the science itself made these interpretations look more like the "mythology" stage of geology. An interesting point to think about is how, with the lapse of time (assuming that sciences will progress), geologists will look upon some of the very modern theories that are fashionable today. This is an open question and is left to the reader to decide. Nevertheless, despite the possibilities presented in this book, several controversies will be discussed since they cannot be ignored at the present state of discussion. Certain controversies are of fundamental significance in providing metallogenic interpretations and through the confrontation of ideas and interpretations, it is believed that geosciences will benefit. Some of the controversies are concerned with special topics and will be discussed in the appropriate sections (see platinum nugget controversy, namely, whether platinum nuggets can be formed in lateritic covers of platiniferous ultrabasics, Chapter 13). Other controversies are of greater importance and their implications of wider application. The granite controversy, or as it is often called, the granite problem, is a fundamental issue in the author's opinion and as he supports, the answer is neither "there are granites and granites" nor that granitization is equivalent to anatexis (Augustithis, 1993). In this connection though, it should be emphasized that whereas some 40 years ago the granitic magma was considered as deus ex machina for interpreting vein deposits, in the last decade many interpretations have been advanced, not only challenging a granitic magma or even an epigenetic derivation (a magmatogenic source) but directly relating the derivation of many ore formations - not only stratabound deposits to a sedimentary origin. Is this a case where the granitic magma hypothesis, after reaching its universal state, is seriously challenged by new and accumulating facts (gained by isotope geochemistry, geochronology and modern experimental mineralogy), where it is not only challenged in the main field of petrology but also in its pe144

ripheral hypotheses (related to granitic magma) such as hydrothermal vein metallogeny? Considering that many metallogenic controversies have their roots in the early 60s, an attempt will be made to follow their course by presenting study cases. What is fascinating about controversies in geology is that they maintain the main characteristics of any other controversy and are just as fierce and passionate. Occasionally, such sentences can be found as an answer to an "accusation" that the studied samples were not sufficient or that "I collected a ton of samples and I was afraid to interfere with the production of the mine", or other slogans like "if science makes no sense, it is nonsense". However, despite extreme cases, criticism and controversies provide the "fuel" for more work. In many cases, the best answer to a criticism has proved to be more dispassionate work and the presentation of more facts along with more critical presentation of views. Study cases will be presented from the plethora of controversies which, despite their complexity, in many cases may point to a path in the "jungle" of the metallogenic "interpretations" as so many people call it. An interesting case is the resurgence of an interpretation which was previously defeated in a controversy. In this connection, Lambert (1982) in "Constraints on the genesis of major Australian lead-zinc-silver deposits: from Ramdohr to recent" mentions the following: "Such deposits (high grade stratiform Pb-Zn-Ag ore bodies) were long regarded as having formed by intricate hydrothermal replacement of "favourable" strata. Almost 30 years ago, Professor Ramdohr concluded that this was not the case at Broken Hill and, since then, it has become increasingly accepted that these deposits were generated by sedimentary-exhalative processes. However, in the past five years there has been a resurgance of the epigenetic versus synsedimentary debate, with particular reference to the virtually unmetamorphosed and little deformed McArthur deposit." Considering resurgence, I would like to point out that as the 20th century A. D. draws to a close, it is marked by many historical events, where resurgence of "long-forgotten" beliefs gain once more tremendous impetus, thus re-writing history. Whether this tendency will influence geosciences 8 is uncertain, particularly when the compressing influence of certain currently prevailing hypotheses will be diminished.

(b) Epi-Syngenetic

Ore

Deposits

Dunham (1967) suggested to make a "classification" of "sedimentary" ore deposits by dividing them into those of unequivocally sedimentary origin, those probably 8 Recently Augustithis was harshly accused of resurgence of petrographic methods in the study of granites and metamorphics, where trace elements resulted in the loss of traces of sense.

but not unequivocally sedimentary and deposits possibly not of sedimentary origin but occurring within sedimentary wall rocks. However, the terms unequivocal, equivocal and probable are arbitrary and reflect the need to be considered each on its own merits. In contrast, a number of workers suggest a dual origin. Specifically Seguin (1978) concluded that the Michipicoten range is of dual origin: sedimentary deposition of siderite, pyrite and little magnetite, followed by a hydrothermal phase in the 200-270° C range, which produced ankerite, pyrrhotite, pyrite and other sulphides. In contradistinction, Di Colbertaldo (1969) introduced the concept of "synsedimentary features in epigenetic ores". According to him, the base metal sulphide deposits of Pira Roma in Sardinia and the Raibl in the Alpine Middle Trias are considered and show textural features suggesting a combined epi-syngenetic origin. The ore fluids are postulated as filling fissure veins during their ascent, followed by the production of syngenetic "flats" in limestone after their escape on the sea-bed. In comparison to Di Colbertaldo, King (1966) introduced the concept of epi-syngenetic mineralization. In the English Midlands many mineralized bodies, hitherto considered to be of hydrothermal origin, are believed to be pseudo-neptunian or neo-neptunian "dykes" filled with deposits of epigenetic mineral matter derived from a remobilized earlier syngenetic deposit. Raguin (1967), discussing the mineralization of the Kosaxa mine in Honshu, Akita prefecture, Japan, states the following: "The ore that is being worked, called Kuroko, is a rich, very fine-grained black complex ore containing Cu, Zn and Pb with pyrite and baryte. In Japan, this ore forms accumulations in the Miocene acidic lavas of northwestern Honshu and southern Hokkaido. A layer of foraminiferal sediment at the top of the deposit indicates a sedimentary origin for the Kuroko. On the other hand, a siliceous ore named Keiko, which appears beneath the Kuroko ore, clearly has the qualities of a stock-work type deposit". Monseur (1966) considers the ore of the dolomite deposit of Reocin as well as those from Udias and La Florida (Spain) as probably syngenetic, although some arguments are in favour of a hydrothermal origin. According to him, the problem has not yet been solved. The main minerals are sphalerite, galena, iron sulphides (marcasite, pyrite, melnikovite-pyrite). In contrast to the syngenetic-sedimentary concept, Palas and Scharm (1967) state: "The genetic problems of pyrite polymetallic deposits of the Jesenik Mountains (USSR) are considered as follows: The deposits are regarded as pre-metamorphic, comagmatic and connected with the early geosynclinal spilite-keratophyre formation of the Jesenik Devonian geosyncline. The ore mineralization is largely of an epigenetic nature, although the possibility of a sedimentary origin for some locations must not be neglected".

A very interesting aspect of the problem is presented by Richards (1963): data for Cu, Zn, Pb, Ag, Sn and As, and Cu-Zn-Pb-Ag metal ratios for the discordant hydrothermal ore body of the Conrad mine compared to analogous data from concordant ore bodies of possible exhalative-sedimentary origin show significant differences. The discordant hydrothermal bodies indicate a relatively high mean abundance of each of the elements Cu, Zn, Pb, Ag, a wide spread in Cu:Pb:Zn ratios and a high Ag:Pb ratio. Besides the contributions of the 60s, more recent contributions also support a dual interpretation for stratabound deposits. Jankovic (1982) reports that "the Sb-As-Tl-Ba mineral assemblage, accompanied by PbZn sulphides and minor amounts of silver and gold, was formed in connection with undifferentiated hydrothermal solutions, which discharged at the bottom of the Neogene basin (syngenetic type of ore mineralization) and/or within volcanic-sedimentary series - tuffs and sinter (epigenetic type of ore mineralization). In the syngenetic type of mineralization, the ore minerals were deposited as disseminated colloids and metastable minerals, while in epigenetic mineralization small massive veins and lenses formed. Shadlun (1982) in her contribution: "Ore textures as indicators of formation conditions of mineral paragenesis in different types of stratiform lead-zinc deposits" supports that "There are three types among stratiform deposits of lead-zinc iron disulphide rich ores occurring in sediments (1) stratabound hydrothermal-sedimentary in carbonate-sandy-argillaceous strata, (2) hydrothermal in the zones of interlayer brecciation in carbonate rocks, (3) metamorphosed hydrothermal-sedimentary in metamorphic schists." Furthermore, Dill (1988) also supports a dual interpretation, stating that "in addition to syngenetic endogenous and exogenous fluorite-barite accumulations, a third mode of formation, called "epigenetic unconformity" is proposed to have been the most dominant type of formation for most of the veins of barite and fluorites in the southern part of the Federal Republic of Germany." Sassano and Procyshyn (1988) support in their mineralogical and paragenetic studies of the metallic and nonmetallic constituents of the carbonatehosted cupriferous ores of the Acton Vale, Upton section, Klippen Belt, Quebec, that "several generations of pyrite and chalcopyrite are associated with sphalerite, galena, pyrrhotite, covellite, bornite, idaite, djurleite, malachite, azurite, Fe-oxides, and minor barite." Furthermore, "the deposits are characterized by three discrete mineral assemblages which are composed of syndiagenetic minerals, epigenetic minerals and supergene minerals." Finally, Gruszczyk (1982), considering the genesis of the zinc-lead ore deposits of Upper Silesia, Poland, concludes that "the present picture of the Upper Silesian Zn-Pb ore deposits is the result of several overlapping processes. The deposits can be regarded as polygenetic because both syngenetic 145

and epigenetic processes played a significant role in their formation." In addition to trace element studies, the results of sulphur isotopes of some Central African sulphide deposits are also equivocal according to Dechow and Jensen (1965). Their comprehensive account of the sulphur isotope compositions of 550 sulphides and 10 sulphates from Zambia, Zimbabwe and the Katanga Province of the Congo Republic suggest that neither a simple biogenic origin of the sulphur nor a simple magmatic hydrothermal origin is adequate to explain variations in isotopic composition. When transposing the argument syngenetic-epigenetic to magmatic versus metamorphic sources, an interesting consideration is presented by Hawley (1956). In his review on magmatic and metamorphic processes in which ores might be concentrated, he concludes that both may play a part in the concentration of ore fluids and in the resulting deposits. He furthermore adds that the magmatic interpretations seem much more efficient and likely, and best explain textures and mineral occurrences in the Superior and Grenville sub-provinces of eastern Canada. He also maintains that metamorphism is undoubtedly effective in redistributing former ore deposits and in the formation of some of the more diffuse deposits and smaller secondary hydrothermal deposits in areas where contemporaneous volcanism is missing. A different point of view concerning the syngenetic and epigenetic hypotheses is presented by Brotzen (1957), who considers the copper mineralization of northern Zambia and Katanga, and stresses that many of the characteristic features of the ores favour an epigenetic interpretation. The existence of hidden younger granites as ultimate source of the ores, however, is considered less likely than the formation and action of mantle gneiss domes. It should be emphasized that such an interpretation introduces the processes of granitization. Furthermore according to him, from the geochemical point of view, intrusive greenstones may be related to the ores.

(c) Sedimentary (Syn-Sedimentary) posits

Ore De-

The aim as defined by Haddon F. King (1967) that "seeking an understanding of the role that known biological and physico-chemical processes could have had in the formation of the stratiform ore bodies which are the principle sources of the world's Fe, Cu, Pb and Zn", probably best serves as an introduction to this Section. Considering the significance of biological processes in the formation of sedimentary deposits, Bächtinger (1960) reports: "The copper ores of the Miirtschenalp, Switzerland, occur in chloritic lenses in a Permian siliceous conglomerate, either as diffuse grains or in veins with quartz, dolomite and calcite. Pyrite, chalcopyrite, tetrahedrite, bornite and molybdenite occur in the ore, 146

with stromeyerite, native silver and covelline in the cement. A biogenic influence is suggested in the precipitation of the ores". Katayama (1960) supports that U is concentrated under reducing environment in the presence of HjS formed by biological processes. In contrast to the biological influence, Kitt! (1967) favours sedimentary concentration processes responsible for the formation of Cu deposits. He also maintains that "instead of the concentration of rare and dispersed elements by magmatic processes preconcentration by sedimentary processes is considered more probable in the genesis of ore deposits. For the Cu deposits of Precambrian to Tertiary age in Argentina, preconcentration of Algomic or Leurentinic age with subsequent mobilization of different kinds is assumed. According to Polge (1966), Pb-Zn ores occur in the Jurassic strata (in the region of Melle, Deux Sevres, France) and they include geodes with galena crystals, jasperoids with sphalerite and galena and disseminated sphalerite in dolomitic rocks. Α hydrothermal origin is refuted because the trace amounts of Pb and Zn in detrital and clay layers free of carbonate are relatively high. Marmö (1960), on the other hand, favours "many of the so-called hydrothermal ores, emplaced under hydrothermal conditions, the ore-forming material having been derived from sediments mainly from sulphidegraphite schists". In addition to the biogenic and sedimentary ore deposits sensu stricto which could produce vein deposits on mobilization, several cases of submarine hydrothermal sedimentary deposits are recorded, in cases complex and multiple processes operated for their formation. An example of hydrothermal and submarine sedimentary deposit is put forward by Anon (1969), where the Rammelsberg ore bodies are found concordantly in Middle Devonian shales, under the overturned limb of a Lower Devonian sandstone anticline. According to Anon, the so-called rich ore contains 7% Pb, 18% Zn, 1% Cu, 12% Fe, 20% barite, 4.2 ounces/ton Ag, 0.02 ounces/ton Au. Furthermore, the banded ore found on the footwall of both main ore bodies consists of alternate beds of sulphides and shales and has about half the metal content. Anger (1966), comparing the Norwegian sulphide ores of Sulitjelma, Skorovas, Joma, Lökken, Röros, Folldal and Vigsnes with the sulphide ore of Rammelsberg, Goslar, supports a syngenetic origin, based on 5 34 S and 32 S/ 34 S values in addition to mineralogical evidence. Furthermore, Anger (1963), on basis of comparison considers the Buchaus, Middle-Newfoundland, sulphide deposits as formed by submarine hydrothermal action and emphasizes the similarities of the Buchaus and the Rammelsberg deposits. Rentzsch (1963) supports that the Pb-Zn-Cu ore in Middle Triassic limestones at the river Isker near Sedmocislenici in northwestern Bulgaria are submarine hydrothermal and comparable on the basis of sedimen-

tary textures, structures and rare element contents to the Pb-Zn ore of the Alps and of Upper Silesia. Concerning syngenetic-sedimentary deposits, the work of Nishihara Hironao (1957), entitled "Origin of the "manto" copper deposits of Lower California, Mexico" is very interesting. According to him, previous theories on the origin of the Mexican ore bodies of disseminated chalcocite in tuff sandstone, which include hydrothermal solutions of magmatic origin, hot springs, supergene enrichment and lateral secretion, lack evidence. It is suggested that the copper was transported in solution from surrounding heights into the Pliocene seas and deposited with the sediments in shallow waters. In contradistinction to the syngenetic-sedimentary mechanism suggested by Nishihara, Petrov (1961) proposed a more sophisticated process for the formation of molybdenum in a brown coal deposit of Uzbekistan. He concludes that the sedimentary type of molybdenum mineralization, represented by jordisite-ilsemannite mineralization, occurs in areas of acid effusives with evidence of polymetal and molybdenum mineralization. According to him, molybdenum was concentrated in the coal deposits in a number of ways, the two most important being: (i) by the activity of plants and their subsequent carbonization and (ii) deposition of molybdenum by underground waters in the brown coal body which was a strongly reducing medium. Regarding environment and processes of sedimentation, the "Genesis of Franklin-Sterling, New Jersey ore bodies" by Callahan (1966) is worth to take into consideration. He suggests that the environment and processes of deposition giving rise to a dominantly Fe-ZnMn oxide type mineralization in sediments of the Red Sea apply to the formation of the stratabound Fe-ZnMn oxide-silicate mineralization of the FranklinSterling ore bodies. Furthermore, concerning sulphide deposits of conformable type, Stanton (1958) reports results of several thousand analyses for Cu, Zn and Pb made on specimens from 14 sample deposits all chosen from the conformable type, mineralogically and structurally similar: they have chemical similarities in persistently exhibiting low copper content, poor coordination between copper and zinc, good coordination between lead and zinc, similarity of lead/zinc ratios and marked consistency in the behaviour of the ternary group CuZn-Pb. Popov (1959), in "Regularities in the distribution of cupriferous sandstones in central Kazakhstan" reports that the cupriferous sandstones of the western part of central Kazakhstan and northern Kirgiziä area are closely associated with uniformly red formations occupying a completely determined position. All known ore-bearing cupriferous sandstones in these regions are in zones of Variscan depressions. According to Popov, the copper-bearing formations are closely associated with gypsum and salt and it is concluded that the copper ores are truly sedimentary.

Other cases of syn-sedimentary deposits are discussed by Schulz (1967), who supports that the baryte deposits within the upper Wetterstein limestone (Landinian) of Bleiberg (Gailtal Alps) are of syn-sedimentary origin. Their mineral content consists of fluorite, quartz, calcite, sphalerite, galena, pyrite, marcasite and clay minerals. The baryte layers show gradual transitions to the ordinary limestone and rhythmic laminates as evidence for syn-sedimentary formation. It is considered that the materials for the paragenesis were supplied by submarine hydrothermal solutions. Furthermore, Schulz (1968), considering the same region (i. e., near Bleiberg-Kreuth, Carinthia) maintains that the ore bodies are submarine syn-sedimentary in origin. They contain the parageneses sphalerite, quartz, calcite, dolomite and clay minerals. The ores are rhythmically layered and occur in the upper 60 m of the Landinian limestone-dolerite series. The ores are derived from low temperature hydrothermal solutions. In his study, Schulz considers general sedimentary features such as external and internal denudation by mechanical erosion, attack by solutions, metasomatic processes and syndiagenetic ruptural and steady deformations. A rather special case of mineralization (comparable to molybdenite "adsorbed" by coal) is reported by Andreyev and Chumachenko (1964), in "Reduction of uranium by natural organic substances". They conclude that the reduction of U rv by organic matter such as peat, wood, cellulose, kerogen, etc., has been experimentally demonstrated at 25° C, 1 atm, at various pH values from 2-18 to 8.07. Although the organic matter was not present in solution, but as a solid phase, uranium was withdrawn from solution forming insoluble oxides and hydroxides of U™. Thermodynamic evidence is presented to substantiate their explanation for the occurrence of pitchblende in sedimentary rocks. In the attempt to present some characteristic cases of syngenetic-sedimentary ore types, some cases will be introduced concerning syngenetic pyrite. Wright (1965), in "Syngenetic pyrite associated with a Precambrian iron ore deposit" gives an account of the geology, mineralogy and geochemistry of the pyrite zones lying above the steep Rock Lake iron deposit in northwest Ontario, Canada. These zones contain about 50% pyrite which occurs as elongated lenses, and approximately 90% of this pyrite consists of fragments with colloform and cryptocrystalline textures. This type of pyrite does not replace any minerals or rocks, but it is veined and replaced by later crystalline pyrite, goethite and haematite. Wright also reports that a trace element study reveals a marked paucity in the number and quantity of trace elements present in the zones. Combined with low Co content, low Co/Ni ratios, presence of As in the pyrite and negligible Se suggest a sedimentary rather than hydrothermal origin. Considering further evidence for a possible sedimentary origin of pyrites, Jensen and Whittles (1969) present sulphur isotopic values from Nairn pyrite opencut 147

mine, Southern Australia, which give 824S from -12.8%e to -20.6%c which are corroborative of geological evidence that the deposit is bacteriogenic-syngenetic in origin. As a corollary to the studies reported so far which mainly refer to contributions of the 50s and 60s, more recent studies are presented in support of the hypothesis of sedimentary derivation of stratabound deposits, and also in this case, a wide gamut of sedimentary processes are considered in conjunction to the formation of sedimentary deposits. Chen (1988), in his study of the Mesozoic and Cenozoic sandstone-hosted copper deposits in South China, reports that the variegated rock formation is generally of fluviolacustrine delta facies, shore facies and shore-cauce facies. According to Chen, "the ore bodies are stratiform, occurring in the light-coloured beds of the variegated rock formation. The mineralization comprises chalcocite, bornite, chalcopyrite, native copper, covellite, pyrite and haematite." These copper deposits are believed to have been formed through three processes of enrichment, i. e., weathering, deposition and diagenesis. Venerandi-Pirri and Zuffardi (1982), in opposition to prevailing views which accept a pyrometasomatic or pneumatolytic-hydrothermal genesis for the tin deposit of Monte Valerio, Toscany, Italy, support a syn-sedimentary tin deposition followed by partial supergene remobilization; metamorphic influence is considered minimal. In addition, Preidl and Metzler (1984), interpreting the copper-bearing shales (Kupferschiefer) in the Sudetic Foreland report that "the concentration of copper sulphides occur in lagoonal areas in which an oxygen deficiency was associated with a great amount of organic matter". Furthermore, "the beds formed in the oxidation facies represent the coastal zones of an open sea basin, where the influence of inflow and outflow were considerable." In support of a sedimentary genesis of some stratabound deposits, rhythmites, slump structures and turbidites have been used. Fontbot6 and Amstutz (1982) consider sphalerite rhythmites of the Trzebionka mine, in the stratabound Pb-Zn district of Upper Silesia-Cracow, as being similar to diagenetic crystallization rhythmites of other stratabound deposits in shallow water carbonate facies and therefore, as an argument for a non-epigenetic origin of the area. Zimmermann (1982) considers the presence of barite-bearing large and small slump structures, turbidites and breccias, disrupted bedded barite, later deposited as fine lenses in turbidites, as evidence for a diagenetic formation and thus supports a sedimentary genesis for the stratabound barite deposits of the Stanley shale of Arkansas. In addition, Beran et al. (1985), considering the Moldanobicum (Bohemian Massif), Austria, a W mineralization bound to calc-silicate rocks (in the Bunte Series), now in amphibolite to higher amphibolite fa148

cies, support that "the main constituents of the scheelite-bearing rocks are ferro-salite, meionite-rich scapolite and quartz. The average tungsten content is estimated to be 1500 ppm; no further elements that would be characteristic of an exoskam formation (i. e., epigenetic) are conspicuous... The mineralization seems to be stratabound - contacts with intrusive rocks were not observed." As they point out, the argument is whether the stratabound mineralization is exoskarn or sedimentary formation and they further support that the field evidence is easier to reconcile with a syn-sedimentary origin. Moreover, considering other types of syn-sedimentary deposits a stratabound finely disseminated Ni-Co arsenide/sulphide mineralization in a weakly metamorphosed series of interbedded slates/phyllites, fossiliferous marbles and basic volcanics and volcanoclastics is reported by Kobe (1982) from the Paleozoic of the Yauli Dome, central Peru. He maintains that "the intimate interlayering of these formations on both the macro- and microscales is taken to indicate a synsedimentary origin for the mineralization, related to submarine deposition of lavas, tuffs and their hydrothermal exhalative derivatives among the seafloor sediments". An additional Cu-Pb-Zn-Ag sedimentary deposit type is described by Lehne and Amstutz (1982) from Colquijicra, central Peru, where sedimentary and diagenetic textures are very common and various types of stratification, geopetal textures and slumping structures are considered as evidence for a sedimentary genesis.

(d) Sedimentary Mobilized or Recrystallization)

Ores

(Mobilization

In contradistinction of the syngenetic sedimentary deposits and as a corollary to the syngenetic-sedimentary ore metamorphosed, a number of study cases will also be presented where ore mobilization took place (without excluding metamorphism as a cause). According to Chowdhury et al. (1960), sporadic occurrences of low temperature galena-sphalerite-chalcopyrite mineralization are found in Permo-Carboniferous dolomite bodies over a stretch of 450 miles in the western Himalayas. The deposits are restricted to particular stratigraphic horizons and do not appear to be related to igneous intrusions in the area. It is suggested that the primary sediments in the area carried abnormal concentrations of Pb, Zn and Cu, which were remobilized and redistributed in structurally weak zones. Knight (1957), in "Ore-genesis - the source bed concept" postulates that sulphide ore bodies are mostly derived from sulphides deposited syngenetically at one particular stratigraphical horizon of the sedimentary basin and that the sulphides subsequently migrated differentially under the influence of rise in temperature of the rock environment.

Calembert (1957, 1958) also supports in his study of the metallogeny of Palli6res (Grad, France) that the sulphides had a sedimentary origin but there were important and perhaps multiple migrations. The most general paragenesis is: pyrite and marcasite, blende, galena, pyrite; cerussite, malnikovite, pyrite, tetrahedrite, bournonite also occur. The country rocks are dolomite, sandstones, sandy limestones, conglomerates and marls. Neither the country rock mineralogy nor the paragenetic associations suggest significant metamorphism for the ore mobilization (migration). Guillou (1969) describes an antimonite deposit in the region of La Sierra de Caurel (Spain) and, according to him, the stratigraphical, palaeogeographical, lithographical and tectonic control of the Sb syn-sedimentary (Ashgillian) mineralization is of a volcano-exhalative origin for the metal (Llandeilo) followed by pedogenesis and redistribution during the Ashgillian sedimentary cycle. Later these deposits were mobilized and locally reconcentrated by Hercynian tectonic features. In contrast to cases where mobilization or remobilization of ore occurred without metamorphism of the initial sediments (syngenetic to the ore), cases will be introduced where ore recrystallization and mobilization occur concomitantly with the metamorphism of its geoenvironment (country rock sediments). Jaskölski (1960), in his studies of the geology and paragenesis of the cassiterite deposit in Iser Mountains, Silesia, suggests an original sedimentary concentration of cassiterite and magnetite in pelitic and arkose rocks, which later, after both contact and regional metamorphism, recrystallized as schists and gneisses with macroscopic sulphides and microscopic cassiterite. In addition, Klominsky (1964) described a magnetite and polymetallic deposit east of Klinovec, Bohemia. The wall rock is an amphibolite with features of transition to a skarn. According to Klominsky, the find of pisolitic textures in the magnetite, its topological relationship to amphibolite, and its geological position of the magnetite layer are evidence for the sedimentary origin of this iron ore accumulation. The transformation of initial iron hydroxides (pisolite) to magnetite supports metamorphism. However, the initial sedimentogenic texture has been preserved, despite the transformation of the initial sediments to amphibolite.

(e) Syngenetic-Sedimentary Ores Metamorphosed and Sediment Participation in Sulfides For the behaviour of syngenetically formed sulphides in metamorphism, the contribution of Davis (1969) that presents features supporting the syngenetic hypothesis as the origin of the Kilembe Fe-Cu-Co sulphide deposit and those attributed to metamorphism, is very interesting. Gastil et al. (1960) describe a series of Precambrian strata representing at least three cycles of sedimenta-

tion, orogeny and metamorphism. According to them, this is one of the largest known reserves of sedimentary iron (in the Labrador geosyncline). It should be considered that in most metamorphosed portions, the iron-bearing sediments have been converted to quartzspecular haematite magnetite rock. A detailed description of sedimentary ores metamorphosed from Kambo (West Katanga) is presented by Bartholomd (1962). The mineralization is located in the "S6rie des Mines", mainly composed of more or less siliceous dolomites and fine-grained detrital rocks. The minerals identified are: quartz, dolomite, muscovite, chlorite and sulphides including chalcocite, digenite, covelline, bornite, chalcopyrite, pyrite and carrollite. A study of the mineral assemblages, in particular bornite-digenite, suggests a deposition temperature below 200° C. The equilibrium of the sulphide association is discussed in accordance with the known system Cu-Fe-S, and the textures of the ore are related to the evolution of the solutions which circulated through the rocks after sedimentation and consolidation. Bartholom6 reports porphyroblasts in the sedimentary rocks formed metasomatically. In addition to the cases described, Bogdanov (1962) reports that Proterozoic cupriferous sandstones occur in small troughs along and parallel to a belt 100 km long in the Udokan region (eastern Siberia). They are associated with metamorphosed red beds, metamorphosed salt-bearing deposits and lenses and beds of graphitic sandstone. The copper minerals have a zonal distribution. The ores with contents of up to 50% sulphides have well-preserved sedimentary structures. According to Bogdanov, the sandstones are primary sedimentary deposits which were metamorphosed at conditions of muscovite-chlorite and biotite-chlorite subfacies of the greenschist facies. Sediment participation in sulphides and element mobilization from sediments to sulphides is another aspect that deserves consideration. Vinogradov and Grinenko (1966), in "Isotopic composition of sulphur in the sulphides of the Noril'sk copper-nickel ores and the genesis of the ores" report that 30% to 50% of the total sulphur in the Noril'sk intrusives were assimilated from sedimentary CaS0 4 that amounted to 1.5% of the mass of the intrusives. As a corollary to that, Cheney and Lange (1967) consider the possibility of forming Sudbury-type ores by sulphuritization, the introduction of country rock sulphur into still hot intrusions. Independent of this interpretation, Cheney and Lange discussed isotope data from Palisade sill, Cornwall (USA), Sudbury, Noril'sk and Doryren (Siberia), the Insizwa sill (South Africa), Porcupine (Canada), the Duluth gabbro, and Stillwater complex, with the conclusion that the sulphur of the ores is derived from the country rocks rather than from the intrusions. The significance of these data is perhaps greater than acknowledged at present. They might mean sulphurization, as Cheney and Lange suggested, or they might 149

eventually lead us to other interpretations, such as the one Augustithis (1978) has favoured, e. g., that graphic quartz/feldspar intergrowths exist adjacent to olivines in the Pallisade sill, which in turn could mean far greater contamination from the country rock. Augustithis (1978) specifically states: "Silica migration by "hydrothermal" solutions from the country rock into which the basaltic melts invaded, is the most probable. This would also explain the coexistence of the incompatible phase of olivine in basalt and the adjacent existing granophyric quartz-feldspar intergrowth." As Goldschmidt and Drescher-Kaden (in Goldschmidt, 1954) showed, Β can be picked up from the country rock sediments by the intruding granite. Also, the sulphur isotope data presented by Vinogradov and Grinenko, Cheney and Lange, might mean greater "sedimentary element migration in the intrusives" than hitherto accepted. Whether this will eventually lead to a transformation derivation of some of these intrusives (as in the case of granites), especially with the banded ultrabasic "intrusives", is a thought that should be followed up. In this connection, it should be noted that the author, Augustithis (1979) interpreted the layered banded bodies of Skaergaard as metamorphic-ultrametamorphic. Perhaps the application of isotope geochemistry will be able to solve this controversy.

( f ) Sedimentary mentary)

Exhalative

(?Volcano-Sedi-

As already mentioned in section (a) of this Chapter, Ramdohr's argument that the Broken Hill, Australia, high-grade Pb-Zn-Ag ore deposit is due to exhalative sedimentary processes and not due to hydrothermal replacement processes (epigenetic) paved the way for the unfolding of the hypothesis of exhalative sedimentary ore formation and considerably enhanced the significance of this interpretation and its wider application. Furthermore, the sedimentary exhalative is a controversial interpretation and for the reader, not familiar with the details, to be informed of the pros and cons, some selected study cases will be chosen from the plethora of available publications. Kullerud et al. (1959) support that "well-known data from the Valley of Ten Thousand Smokes (Alaska) and the Cyclades Islands (Greece) show that while some or all of the mineral constituents of certain ore bodies may occur in volcanic gases, the concentrations are many orders of magnitude lower than necessary for extensive ore deposition even from the largest imaginable volumes of volcanic exhalations". In contrast to the aprioristic negative position of Kullerud et al. for volcanic gases as agents of extensive mineralization, Oftedahl (1959) estimates the magma chambers (bodies of granite rock) in the Oslo area to be a maximum 20x20x10 km and refers to Goldschmidt's conclusion that ore deposits along the contacts of these plutonic bodies are the result of en150

richment and transportation of iron, etc., solely by the aid of some kind of volatile activity from a granitic magma chamber (in contrast to these interpretations, see Chapters 25 and 60). Concerning the difficulties of iron enriched in a gas phase, he suggests that a superheated HjO gas may be a potent solvent. Oftedahl expands his hypothesis and interprets the association of porphyries with the ore bodies to be due to the interaction of Na-rich sea water and hot, acidic glass particles. According to Oftedahl, the deposits at Meggen and Rammelsberg are exhalative-sedimentary. In contrast to Oftedahl's exhalative-sedimentary hypothesis, from the geochemist's viewpoint Landergren (1958) does not accept Oftedahl's explanation that the enrichment and transportation of iron took place solely by the agency of volatile activity emanating from a granitic magma chamber. According to Landergren, iron-rich exhalations are possible only if halogen compounds, but no water, are present. Furthermore, the degree of oxidation of Swedish skarn iron ores also makes the exhalative-sedimentary origin improbable. Kautsky (1958), also referring to Oftedahl's explanation of the Huelva (Spain) province deposits, recalling previous work concerning the Peninna-Lazarza ore body and to Rio Tinto, points out that the Huelva ores are genetically not related to ignimbritic tuffs but are associated to aureoles of porphyries. Following the exhalative-sedimentary hypothesis, Geijer (1964), evaluating the existing geochemical data and the accumulated field observations, confirms once again the long established interpretation that the sulphide deposits of the Falum type in the Precambrian of central Sweden, are genetically associated with the intrusion of the synkinematic Svecofennian (early Precambrian) granites. Geijer maintains that no evidence exists in favour of the arguments presented by Koark (1963), according to whom the sulphide mineralization should have occurred during the building of the Leptite Formation (in which the deposits are situated) by an exhalative-sedimentary process. In contrast to Kullerud's position concerning the formation of ore deposits by volcanic emanations, Goryainov et al. (1967) in their study of the sillimanite metamorphic rocks of the Tundra series of PreKarelian age layers of kyanite quartzite associated with ferruginous quartzite, amphibolite and gneiss containing pyrite, galena, sphalerite and chalcopyrite in kyanite-rich partings, support that textural relationships indicate that the ore minerals formed from the original components of the sedimentary rock, which were derived from volcanic exhalations, rather than by metasomatic introduction of material. In conjunction to the sedimentary exhalative deposits and as a corollary to them, Appel (1986), in "Stratabound scheelite in the Archean Melene supracrust belt, western Greenland", reports that a tungsten province exists ca. 300 km long and up to 120 km

wide, with extensive banded amphibolites containing up to 2% W and 0.16 ppm gold. According to Appel, the tungsten occurs as scheelite which is associated with tourmalinites and stratabound tourmaline-rich layers in amphibolites of presumed tuffaceous origin and with an iron formation containing high amounts of tungsten, zinc, copper, lead, molybdenum and tin. Appel further supports that the scheelite is stratabound and of submarine exhalative origin. An even more vivid picture concerning exhalative deposits is presented by Mangan et al. (1984), in "Submarine exhalative gold mineralization at the London-Virginia mine, Buckingham Country, Virginia". Here, finely disseminated gold occurs dispersed with minor amounts of pyrite, sphalerite, chalcopyrite, galena and tennantite in a ferruginous quartzite and quartz-muscovite schist. However, it is interesting that the deposit is believed to have been formed by processes analogous to those currently active in the Atlantis II Deep of the Red Sea. In their interpretation, however, they involve not only exhalative processes sensu stricto but also hypersaline brines, hydrothermal silica deposition, volcanism and metamorphism. In this connection, it is perhaps interesting to quote the relevant passage from their paper: "The deposit is believed to have been formed by processes analogous to those currently active in the Atlantis II Deep of the Red Sea. Silica-rich, hypersaline brines discharged through fractures in the sea floor and ponded in a local basin. Episodic influx of clastic debris and extensive deposition of hydrothermal silica diluted the concentration of sulphides and gold to produce a low-grade, siliceous mineralized zone. Emanation from the exhalative vent was terminated when the basin was capped by a lava flow. Subsequent regional green-schist grade metamorphism has recrystallized the silica into a granular quartzite and produced minor remobilization of the gold and sulphides."

(g) Epigenetic

Deposits

In contrast to the sedimentary syngenetic deposits and the concordant syngenetic types briefly discussed, several stratiform deposits or deposits in sediments are considered to be epithermal. Ford (1969), regarding the stratiform deposits of Derbyshire, England, where deposits of galena, sphalerite, fluorite and baryte occur in Carboniferous limestones, believes that they are epigenetic telethermal and the lateral flow was at least as important as the upward movement of the ore forming fluids. Leroy and Poty (1969) present thermometric studies of fluid inclusions in quartz of the granite-episyenite suite and of the pitchblende ores in fractures and showed that the fluid phases were rich in C 0 2 at high temperatures and low pressure during the deposition of uranium.

Furthermore, Davidson (1962), researching some genetic problems of the Dzhezkazgan copper ore (Kazakshtan), on the basis of Russian work on copper mineralization in Carboniferous shales, sandstones and conglomerates, suggests that the mineralization is probably epigenetically associated with the Hercynian orogeny. Davidson recognized two distinct phases, copper being concentrated during the early phase and lead and zinc in the later phase. Also Cornelius (1967) described a breccia pipe associated with epigenetic mineralization in Mount Morgan, Queensland, Australia. According to him, the Cu-Au-Ag ore body is a pipe breccia and not a replacement deposit in a fault zone. Chalcopyrite, pyrrhotite and pyrite form the bulk of the sulphide minerals, but calaverite, tellurobismuthite and tetradymite have also been identified. In contradistinction, Kantor and Biely (1965) on the basis of lead isotopes consider the Jan Nepomucky deposits near Pila as due to a mineralization related to Tertiary magmatism. The deposit occurs in Triassic limestones close to the contact with andesites. It contains mainly anglesite, cerussite and galena. Isotopic analysis gave 204 Pb 1.370, 206 Pb 24.82. 207 Pb 21.42, 208 Pb 52.39%, indicating a pre-Tertiary age of mineralization, and therefore not related to andesite. Galena from a nearby galena-sphalerite-pyrite deposit in the Moras Valley has 204 Pb 1.35, 206 Pb 25.07. 207 Pb 21.12, 208 Pb 52.46%. This mineralization is probably related to Tertiary magmatism. It should be noted that the geochronological evidence used supports an epigenetic metallization in the Moras Valley due to magmatic derivative solutions. Furthermore, considering different mineralizations related to sedimentary rocks, cases are present where a syngenetic-sedimentary origin is excluded on the basis of exhaustion processes. According to Barnes (1959), the quantitative estimation of Zn, Cu and Fe content of sedimentary rocks near ore deposits of the Hanover district, New Mexico, the northern Mississippi Valley district, and the San Francisco del Oro district, Mexico, indicates that the total metal content of the deposits could not have been concentrated through a process of lateral secretion from surrounding sedimentary rocks. Another case of a possibly epigenetic mineralization is presented by Harder (1963), discussing the origin of itabirites (quartz banded ores). The origin of the banded iron ores is explained by different authors in two ways: (i) the iron is derived from weathering solutions, (ii) the iron is derived from solutions arising in volcanic or magmatic processes. The latter possibility is within a broader epigenetic concept. Accepting a broad concept of epigenetic mineralization, cases are considered where tectonic influences produce either "hydrothermal" solutions due to friction (see Drescher-Kaden and Heller, 1961) or can result in remobilizations. Boyle (1965) reports that the Bathurst-Newcastle PbZn-Cu deposits are interpreted as epigenetic, in con151

trast to recent interpretations of such ores as syngenetic and premetamorphic. It is suggested that the deposits were formed during a late stage of shearing associated with the main tectonic and metamorphic processes which affected the region. Furthermore, Gjelsvik (1960), when discussing the Skorovass pyrite deposit, in the Grong area, Norway, reports that the deposit occurs in a relatively flat-lying strongly metamorphosed series of greenschists, representing submarine pyroclastic rocks and basalts. No important structural control of the ore deposition has been found. Wallrock alteration includes chloritization, silification, carbonization and sericitization and varies considerably in intensity and character. The ore minerals are pyrite, chalcopyrite and blende: they probably have a hydrothermal-metasomatic rather than syngenetic origin. Considering the effects of stress and remobilization of materials, Dunnet and Moore (1970), in "Inhomogeneous strain and the remobilization of ores and minerals" support the following: "The responses of a rock sequence to a differential stress system include a change of shape or strain of the rock mass and the migration of pore fluids. If the fluids carry ores in solution, this can lead to their selective remobilization and redeposition. The fluids will normally migrate from zones of high strain to those of low strain: however, the stress drop associated with late brittle structures may produce hydraulic fractures in zones of high strain which would become channel-ways for migrating fluids". According to Dunnet and Moore, examples of structural localization of Pb-Zn ores in southwest Sardinia suggest that in different localities both high and low strain environments are locally enriched in remobilized ores: these environments may be distinguished by the nature of rock fabric. It is a corollary to the theory of remobilization that the textures of the resulting ores are indistinguishable from those produced by epigenetic fluids. The statement of Dunnet and Moore that remobilization of fluid-solutions may result in textures indistinguishable from textures formed by epigenetic mineralization is supportive of the general concept that after all "lateral secretion" or "segregation due to tectonic effects" may result in mineralization texturally comparable and commensurable to epigenetic mineralization. Concerning epigenetic mineralization, it is infeasible to cover all the possible geological, petrological and mineralization conditions, and for that reason only some additional cases will be shown. Warrington (1965) presented the geology and mineralogy of the Cu and Pb deposits of Alderley Edge, Cheshire, and he considered the surface and subsurface features of the deposits on which he suggested an epigenetic origin for the ores: Cu, Pb and Co are related to the fault system of the area. A high Co/Ni ratio reported for galena is taken to indicate derivation from an acid igneous source. 152

In contradistinction, Korzhinsky (1959), in "The advancing wave of acidic components in ascending solutions and hydrothermal acid-base differentiation" reports that "in rocks altered by the flow of ascending solutions, vertical zoning should give (from bottom to top): a zone of autometasomatic alteration of magmatic rocks; a zone of maximum leaching; and a zone of fading acidity of advancing wave, giving a predominance of precipitation over leaching in telethermal ore veins. Considering U deposits in contrast to sedimentary interpretations, Nash (1968) in his study of uranium deposits in the Jackpile Sandstone, New Mexico, supports that U deposits whose principal ore mineral is coffinite, are due to epigenetic U-bearing fluids, derived either from Upper Jurassic volcanic exhalations or by leaching of arkosic sands, that entered the Jackpile fluvial sediments soon after sedimentation. Abundant humic carbonaceous material in these sediments effect sorption and reduction of U 6+ , thus, coffinite is commonly associated with organic material in the sediments. Furthermore, local control is exhibited by sedimentary structures and permeability variations. In contrast to these cases where epigenetic mineralization is considered to be the sole process, Friedrich (1964) proposed a model of genesis for lead-zinc ore deposits of the eastern Alps where all transitions between epigenetic and syn-sedimentary depositions of ores are possible. According to Friedrich, along steep rupture zones the base of the geosyncline sank heterogeneously. Ore deposits in the form of pipes, veins, etc., originated along germanotype settling joints in the older sediments of the geosyncline. The rest of the orebearing solutions passed upwards and precipitated on the ocean floors, thus forming syngenetic sedimentary ores. The metalliferous reservoir is situated in abyssal (plutonic) depths. In addition to Friedrich's interpretation (1964), according to which there is a relationship between epigenetic metallogeny and deep rupture zones, Eisenlohr et al. (1989), in "Crustal-scale shear zones and their significance to Archean gold mineralization in Western Australia" support that large Archaean epigenetic gold deposits show a broad spatial relationship to regional lineaments in greenstone belts, although in detail, they are situated in subsidiary brittle-ductile fault structures. "Fluids originating from a deep source, follow a complex path and re-equilibrate with different lighologies and with metamorphic fluid during migration to higher crustal levels." It is interesting to note that Eisenlohr et al. in their epithermal metallogeny involve the action of metamorphic fluids as well. Characteristically they support "the multi-source origin and continuous reequilibration of the fluid with crustal rock, which includes granitoid and greenstone-belt lithologies of different ages, is reflected in the diverse isotopic and geochemical signature of the gold deposits". Another case where epigenetic mineralization is attributed to connate brines set in motion by a high geo-

thermal gradient accompanying continental rifting is proposed by Olade and Morton (1985), in which case a relationship between tectonic influences and epigenetic metallogeny is also suggested. Describing the geoenvironment and the type of Pb-Zn mineralization in the southern Benue Trough, Nigeria, they state: "The epigenetic Pb-Zn deposits of the southern Benue Valley (Nigeria) are localized within Cretaceous sediments of an intracontinental rift basin ... and the trace element contents of sphalerite and galena are also consistent with low temperature of formation and epigenetic oriit

gin. Ilani et al. (1990) reports of an epigenetic manganese mineralization occurring within Cretaceous carbonate strata of Israel, which is also structurally controlled. According to them, "in southern Israel, manganese· and iron-bearing waters rising along fault zones precipitated these metals as different redox conditions were encountered. The manganese and associated trace metals are believed to be leached from magmatic bodies by corrosive deep-seated brines and circulating meteoric waters." Derr6 et al. (1982), considering epigenetic metallogeny, describe Sn-W veins occurring in the contact aureole of composite granitic plutons South of Bragan^a in Portugal. They also report thin layers of calcic quartzites with or without scheelite interbedded in the Ordovician and Silurian metamorphic country rocks. The scheelite-bearing calcic quartzites in the contact aureole of the granite are changed into tactites and thus support the interpretation that the scheelite mineralization is related to hydrothermal processes and therefore is not syn-sedimentary but epigenetic. This interpretation is contrary to the submarine exhalative stratabound scheelite occurrence in Greenland, as reported by Appel (1986). (h) Ultrametamorphic-Sedimentogenic Versus Magmatogenic Differentiation tative Layering

Origin Gravi-

Considering the international literature particularly concerning the chromite layers (bands) and the Merensky Reef of the Bushveld complex (e. g., Cameron and Emerson (1959), and Cameron and Desborough (1969)), these were equivocally believed to be magmatic differentiation products either of parental basaltic magma or of an ultrabasic fraction magma. Basically, the interpretations have been in accordance with Bowen's concept of the "evolution of igneous rocks" (1928) and Niggli's magmatic differentiation. Moreover, the layered complexes and especially the chromite-layering (bands) in ultrabasic complexes have been explained as due to magmatic differentiation and fractional crystallization and by gravitative settling of crystals forming layers as suggested by Wager and Deer (1939) for the Skaergaard complex. Even such revolutionary-minded authors as Drescher-

Kaden provided magmatic differentiation interpretations for a layered complex in Kaersut, Greenland, in "Sekretionsdifferentation" (Drescher-Kaden and Krüger, 1927) and also in their paper "Zur Kenntnis des Peridotits von Kaersut (Grönland) und seines Ganggefolges" (1932). There was the unanimous agreement that layering in ultrabasic complexes was due to magmatic differentiation/fractional crystallization and gravitative separation of layers. When Augustithis (1979), in "Atlas of the textural patterns of basic and ultrabasic rocks and their genetic significance" proposed that the Skaergaard and other ultrabasic complexes are ultrametamorphic-metamorphic with the involvement of mantle lower crust, and that the Schlieren banding of Vourinos in Greece was due to prototectonics in a mantle diapir, as the Vourinos complex was interpreted, the international scientific community reacted with severe criticism which at best expressed scepticism. However, despite criticism, Augustithis continued his studies on textural analysis and encouraged by the change of opinion of the international scientific community on previous controversies in which he was also harshly criticized, namely (i) that platinum nuggets have been formed under lateritic conditions (Augustithis, 1965, 1967, 1979; Ottemann and Augustithis, 1967), and (ii) - partly justified - that granites exhibiting metamorphic-metasomatic tracts are products of granitization (Augustithis, 1962, 1973, 1993; see Chapters 13 and 22), he reinstates his previous position concerning the ultrametamorphic/sedimentogenic origin of basic portions of the ultrabasic complexes and, in particular, the chromite banding of Bushveld, with the presentation of the following additional observations and interpretations based on comparative textural analysis. Fig. 909 shows an enclave of clastic chromite grains, some rounded, others irregular, some even recrystallized, enclosed in recrystallized coarse-grained chromite, which forms the main mass of the Wintervelt chromite (band) layer. The Wintervelt chromite band is one of the main (mined) chromite bands in the Bushveld complex. It is about 1.30 m thick and has been followed underground in the mine for more than 1 km. It should be mentioned here that sampling of the band showed the abundant presence of clastic chromite enclaves in the chromite band which consist of coarsegrained recrystallized chromite. The clastic enclaves are often pea-to-small-almond in size, are interspersed in the main chromite band and represent relics of an initial clastic chromite sand layer which was deposited on a weathered mantle or metamorphosed ultrabasic surface in early Precambrian times, overlayered by anchi-sediments. The entire complex has been metamorphosed-ultrametamorphosed and whereas, as already mentioned, the clastic chromite enclaves are pea-to-small-almond sized, smaller enclaves of clastic chromite grains are also interspersed in the main chromite band. Fig. 910 shows a relatively small en153

clave in the main chromite band consisting of recrystallized coarse-grained chromite. Also in this case, the chromite enclave shows angular, rounded and recrystallized clastic chromite grains in the background mass (of initial sedimentogenic matrix, now transformed into plagioclases and pyroxenes). These smaller and larger enclaves of clastic chromite grains represent sedimentogenic relics preserved in the main chromite band which consists of coarse recrystallized chromite (see Fig. 911). A typical textural pattern of the clastic chromite enclaves preserved consists of angular and rounded clastic chromite grains (Fig. 912). (Again, the initial sedimentogenic matrix is transformed due to metamorphism into feldspars and clinopyroxenes.) In cases, the transformed initial sedimentogenic matrix on metamorphism resulted in oriented pyroxene and plagioclases, perhaps following the initial banding of the anchi-sediments that comprised the matrix of the relic enclaves in which clastic chromite grains are also preserved (Fig. 913). These enclaves might represent initial parts rich in matrix. However, if the interpretation of the textural analysis based on the principle of comparative anatomy, is extended to its maximum acceptable consequence, the clastic chromite grains exist as interspersed relics in the platiniferous Merensky Reef as well. Figs. 914 and 915 show clastic chromite grains, often well-rounded, interspersed in the Merensky Reef pyroxene pegmatoid. As especially Fig. 915 shows, rounded and angular clastic chromite grains occur mainly with anorthiterich plagioclases of the Merensky Reef "pegmatoid". Similarly rounded chromite grains are observed included in the plagioclase xenoblast (previously referred to as intercumulus as well as in the pyroxene crystalloblasts, see Fig. 916). As the textural analysis supports, the Merensky Reef of the Bushveld complex is also an ultrametamorphicmetasomatic product (topometasomatic) in a comparable sense that the granitoid pegmatites represent exudation products of granitized material. Similarly the pyroxene-anorthite-rich plagioclase Merensky Reef pegmatoid has been formed as a "reef' following a zone or band in the Bushveld complex in which, due to metamorphism-metasomatism, pyroxene crystalloblasts were formed (often exhibiting idioblastic forms, see Fig. 917) and with intercrystalline ?poikiloblastic anorthite-rich plagioclase xenoblasts (with less well-developed crystalline outlines). It should be emphasized that the rounded and angular chromite grains represent relics of the initial anchisediments, the metamorphism/topometasomatism of which they resulted in the pegmatoid pyroxene megablasts and feldspar poikiloblasts. The rounded and angular chromite aggregates often forming "clusters" or being interspersed in the Merensky Reef pyroxene pegmatoid represent undoubtedly a relic sedimentogenic phase, comparable to palaeosoma. 154

On this occasion, it should be pointed out that no PGM has been observed as an older inclusion in the rounded chromite. The observations recorded so far support that the PGM observed are later mobilizations, perhaps lateral secretions or "hydrothermal" mobilizations from the wider geoenvironment of the Bushveld complex, which is ultrabasic-basic in composition. Fracture fillings of platinoid minerals (PGM) are often observed and recorded in the Merensky Reef (Fig. 918). Considering the clastic chromite grains perceived as relics in the enclaves (relics) of the Wintervelt chromite bands (see Figs. 909 and 910), and the comparable, both in form and size, clastic chromite grains interspersed in the Merensky pyroxene pegmatoid, it is supported that the chromite Bushveld layers represent initially chromite sands derived by weathering of chromite-rich primordial ultrabasics (exposed mantle or initial peridotic in composition earth crust) transformed by metamorphism (recrystallization) into the chromite layers consisting of coarse chromite recrystallized grains, often exhibiting crystalline contours. The observations recorded by Augustithis (1979) of rounded chromite grains and chromite grains exhibiting crystalline outlines, associated with uvarovite (and having later chalcedony) could similarly be interpreted as initial sedimentogenic clastic chromite recrystallized under metamorphic conditions (see Figs, l l a n d 12). In contradistinction to the chromite layers which are rich in clastic chromite enclaves, representing relics, originally the Merensky Reef mainly contained sporadically interspersed clastic chromite grains, relics that are preserved as scattered grains or aggregates of grains, associated with later formed pyroxene and feldspar crystalloblasts. The temperatures involved in the case of the chromite layers (formed by recrystallization of the clastic initial chromite sands) and the Merensky Reef must have been lower than the melting temperature of the chromite grains. Otherwise clastic chromite grains could not persist so abundant neither in the chromite layers nor in the Merensky Reef.

(i) On the Hypothesis of Liquid Unmixing due to Pressure Release Versus Orthodox Petrographical Views on the Formation of Layered Ultrabasic Bodies and Their Significance in Ore Prospecting In contrast to the furnished interpretations concerning the anorthosite-chromite layering of the Dwars River (see Chapter 14), and the initial clastic derivation of the chromite bands at the Wintervelt mine (see Chapter 27, previous section), a hypothesis based on the principle of physics that mixed liquids under great pressure separate when the pressure is released, could be considered as a novel alternative hypothesis for interpreting layered ultrabasic bodies. Concerning the forma-

tion of the layered complexes, the above-mentioned diverse and contradicting hypotheses are briefly outlined as follows: (i) Basaltic magma differentiation and its alternative basic-ultrabasic magma differentiation (these are the basic magmatogenic interpretations - and a vast classical literature is available on support of this hypothesis which is fundamentally based on Bowen's "Evolution of Igneous Rocks" (1928), and Wager and Deers' gravitative crystal settling hypothesis (1939). (ii) Metamorphic-metasomatic derivation of the layered basic and ultrabasic complexes involving metamorphism/mobilization of basic anchi-sediments of early Precambrian and involving mantle and lower crust tectonic mobilizations (e. g., the banded chromite-anorthosite of the Dwars River occurrence). (iii) The third interpretation is a novel hypothesis which is based on physics, supporting that when pressure is released, mixed liquids will separate (unmix). Prof. A. B. Vistelius has already made a suggestion/hinted at this possibility concerning the formation of layered deposits and particularly the anorthositechromite banding (e. g., Dwars River) and the Wintervelt chromite bands associated with norite (also dunite) and pyroxenite due to liquid unmixing by pressure release ~ mixed chromite pyroxenite also exists. (However, it should be pointed out that insufficient information is available for the possibilities of application of this hypothesis.) The hypotheses outlined above and the respective correctness of them is of great practical significance for locating particularly chromite deposits and to a lesser extent Pt deposits (since PGE can be mobilized under magmatic, hydrothermal and even lateritic conditions). According to the basaltic magma differentiation or ultrabasic magma differentiation hypotheses, the chromite layers should only occur in the lower parts of the layered complexes (since chromite is heavier than the silicates). Different modifications of the hypothesis attempt to explain diversities. Concerning the gravitative separation of chromite bodies as an example for the interpretations provided, the following extracts from Cameron and Desborough (1969), and Cameron and Emerson (1959) are quoted: "Chromite occurs in the eastern Bushveld complex in the critical zone, a layered series of pyroxenite, norite and anorthosite with minor gabbro, dunite and harzburgite. The zone extends for 75 miles. The critical zone is divisible into chromitic and non-chromitic intervals, according to the presence or absence of cumulate chromite: most chromitic intervals consist of chromitic dunite and harzburgite. The upward change from predominant pyroxenite to predominant norite and anorthosite, and the upward increase in total Fe and Fe/Mg and Fe/Cr ratios in chromite, Fe/Mg of orthopyroxene and Ab/An of plagioclase all suggest that the critical zone is the result of progressive magmatic differentiation."(Cameron and Desborough, 1969)

Other studies by Cameron and Emerson (1959) on Bushveld support that field and laboratory studies indicate that the layered structure of chromite-bearing zones of Bushveld Complex is essentially the result of gravitative accumulation of early-formed crystals. Irregularities in layering and other features suggest, however, that the magma was at times in motion. However, they also suggest that these are evidence textural features, bronzite-chromite veinlets in xenoliths - to justify the conclusion that chromite in the high-grade chromitite members (e. g., Steelport and Leader Seams) has crystallized in part in situ, following partial resolution. Considering the incompatibility between gravitative layering and chromite bodies occurring both in layered as well as in "intrusive" bodies, where chromite bodies most often are not complying with the hypothesis of gravitative settling due to magmatic differentiation (see especially Chromite-Anorthosite Banding, Augustithis, 1982 and 1983) and considering the objections and criticisms raised in the Atlases by Augustithis (1978, 1979), an unorthodox interpretation is put forward by Augustithis as already pointed out. According to the metamorphic-metasomatic hypothesis, the prospecting for chromite and Pt deposits is more complex, more difficult and more unpredictable than following the "rules" of gravitative crystal settling/differentiation. The behaviour of a rigid body or fragments (as chromite can be) in a visco-elastic geoenvironment as forsterite-rich mantle 9 is, requires an entirely different conception for understanding their distribution. Chromite lenses in dunitic bodies could be comparable to gigantic boutinage structures. Also metamorphic and metasomatic mobilization under tectonic influences as well as prototectonics in mantle material, as some parts of these ultrabasic complexes are, would render the location of chromite ore bodies more complex and other criteria are required. Another basic question is whether the bands of chromite (in cases 1.3 m thick and followed underground for more than I km) could represent initial chromite sand layers derived by weathering of exposed primordial mantle chromite-rich initial "lithic crust" under entirely diverse geoenvironment conditions (incompatible with the uniformitarian concept). The chromite bands and their dunitic sea-bottom as well as the superimposed anchi-sediments would be intensely metamorphosed. Current research work by Augustithis reveals chromite-clastic grain enclaves as lens-shaped small bodies (relics) within the recrystallized chromospinel bands (see previous section (h) and Augustithis, 1994). Another complex problem is whether under mantle/lower crust prototectonics, anorthosite-chromite interbanding and mobilization can take place (see 9

Also an important indication for prospecting for chromite in ultrabasic bodies is that Mg-rich forsterite is in cases associated with chromites. 155

Augustithis, 1982, 1983), e. g., the chromite-anorthosite interbanding in Dwars River, Bushveld complex, where plastic remobilization under mantleAower crust conditions is proposed. If, on the other hand, the third hypothesis "liquid separation on release of pressure" is considered, specific band associations might be of interest for chromite prospecting. For the investigation of such processes, mixed liquids at high pressure should be used for experimental verification of this theoretical concept. It is suggested that mixed liquids of the composition chromite/anorthosite, chromite/dunite and chromite/norite/pyroxenite should be used (if possible?). Thus, the respective theories and hypotheses are of fundamental significance for searching for chromite deposits in the layered ultrabasics as well as in "intrusive" dunite (mantle diapirs). In contradistinction to layered ultrabasic bodies, considering the alpidic type of chromite deposits, which could also exhibit layered bodies, show that chromite occurs either as lenses or as Schlieren bodies

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in dunite, and two alternative interpretations should be considered further: (1) Differentiation of the basaltic or ultrabasic magma, respectively, and by gravitative separation crystal settling takes place forming Schlieren bodies, lenses, pods or "pipes" which might subsequently be tectonically mobilized or have been mobilized in the visco-elastic dunitic or even serpentine parental geoenvironment. (ii) The dunitic bodies are mantle diapirs and as the visco-elastic forsterite-rich body is tectonically mobilized, chromite mantle bodies are differentially mobilized with or within the dunitic or ultrabasic bodies. This, in fact, explains the association of large chromite bodies with relatively small dunite bodies - since the chromite bodies could hardly be explained as magmatic differentiates of the specific small dunite "intrusives" (i. e., they are mantle chromite pieces diapirically mobilized together with the dunite and the size of the chromite bodies is independent of the size of the "intrusive" dunitic bodies).

Chapter 28

The Witwatersrand Controversy

The Witwatersrand System in South Africa is one of the most significant deposits, not only for its Au production (and other metals, e. g., uranium) but also because it has been the subject of a syngenetic-sedimentary versus epigenetic controversy for years. Schematic presentation of the features of the Witwatersrand System is presented in a booklet published by the Chamber of Mines of South Africa, which is also accompanied by a general description of the geology of the Witwatersrand System (Fig. 919). The following is quoted: "The Witwatersrand basin and its gold fields were laid down many millions of years ago. Geologists say that a great lake stretched 250 km form the Kiersdorp area to the present Springs-Nigel region east of Johannesburg. Several rivers, they believe, flowed into this lake from granite mountains to the north and south-west, carrying with them sand, silt and pebbles and fine particles of gold. The prehistoric land was carved and shaped by these rivers, which deposited pebbles and gold in fan-shaped deltas near the shores of the lake. Over aeons there were alternating periods of deposition, of activity and calm. Sometimes fast-flowing rivers and streams brought rock particles from northern mountains to the lake, rolling them out across the braided (laced with gravel and sand bars) deltas, piling them up, layer upon layer. At other times the streams ran slowly and carried only sand and silt with very fine gold. Eventually the lake was silted up, filled with vast sedimentary deposits containing gold-bearing reefs rather like jam layers in a cake. The mud, pebbles, sand and gold were fixed and preserved in the sedimentary deposits which became very hard rock - now known as the Witwatersrand System - in places 7500 metres thick. The earth's formative turmoil continued and the Witwatersrand rock became warped as the nascent formations were thrusted upwards and slipped downwards, sometimes a few centimetres, sometimes thousands of metres. Then volcanic action forced up great dykes of igneous material and lava flowed across the country. Deposits of dolomite rock were laid down later when the area was covered by the sea or at least another great lake. The whole bed of Witwatersrand rock tilted at some stage; it became like a giant saucer with one edge deeply buried and the other thrust near the surface. It was on the up-thrust northern lip of the

saucer that the Witwatersrand gold deposits were discovered in an outcrop in 1886. Sedimentologists agree that the main deltas on the prehistoric lake, along with the river systems and their flood-plains, were the areas of greatest concentration of deposits of gold-bearing particles. Today, these ancient river deposits, now covered by kilometres of gully rock, form the rich gold-bearing reefs of South Africa's "golden arc" where the main gold-mining activity concentrated for more than a century." From the vast number of contributions supporting sedimentary origin of the Witwatersrand System, only some will be quoted in order to elaborate on some significant problems of the mineralization in this system. Saager (1969), in "The relationship of silver and gold in the Basal Reef of the Witwatersrand System, South Africa" supports the following: From the Basal Reef conglomerate horizon of the Upper Witwatersrand System in the Orange Free State, Ag/Au ratios of individual gold particles are constant at about 0.0812 but ore samples give higher values and a large but really systematic variation". According to Saager, they are interpreted to indicate homogenization of detrital gold during transportation and sedimentation so that part of the original Ag content has been redistributed. Also, the sedimentation and the effects of later metamorphism and pseudohydrothermal activity are considered. Back in 1968, Saager had already reported a myrmekitic intergrowth of linnaeite, gold and pyrrhotite which had been observed in a detrital grain from the Basal Reef. The uraninite and the Co-Ni minerals of the Reef might have been derived from primary U deposits of the Chingalore/Wittichen type. The myrmekitic intergrowth suggests the same origin for the uraninite and part of the gold. In addition, according to Saager, molybdenite, stromeyerite, proustite, tennantite, dyscrasite, safflorite and haematite are other minerals reported. In a more recent publication, Saager et al. (1982) compare the pyrite quartz pebble conglomerates from the pre-Witwatersrand with the lower and upper Witwatersrand sediments, which show "far reaching similarities in their mineralogy". According to them, the erratic and low gold and uranium contents of the pre-Witwatersrand conglomerates are explained to a large extent by the inefficiency of the operating sedimentary reworking processes. The detrital chromites of all the conglomerates studied have identi157

cal chemical composition which points to an origin from the same provenance areas. Furthermore, concluding, Saager et al. support that "from this study it is apparent that the Archean pyritic conglomerates are primitive forerunners of the Witwatersrand gold-uranium placers". Koen (1961) supports that detailed measurements and statistical analysis of size distribution of uraninite, zircon and chromite fractions in numerous samples of the Witwatersrand blanket indicate strongly that all these are hydraulically equivalent and were deposited from suspension in water. Furthermore, Koen interpreted nodular precipitation of uraninite on muddy floors of marshed or shallow lakes during intraformational breaks, followed by reworking during periods of renewed wave action. Support for a sedimentary origin of the mineralization of Witwatersrand also comes from Liebenberg (1960), who states that the uraninite and iridosmine are shown to belong to the heavy detrital suite of minerals in the Witwatersrand conglomerates. Liebenberg also points out that particles of gold exhibit "pseudohydrothermal" shaping claimed to be due to metamorphism. In addition, a sympathetic relationship between the amounts of Au, U and Os in the conglomerates show that the gold, uraninite and iridosmine are consanguineous and since the two latter minerals are alluvial, it is claimed that the gold is also alluvial. Concerning the distribution of gold and silver in the Witwatersrand conglomerates, it is believed by Reh (1964) that only by precise investigations and by sampling followed by statistical evaluation of a large number of channel samples it is possible to determine the distribution of gold and silver in the Witwatersrand conglomerates. From such investigations, new facts arise concerning the primary distribution of these elements in the "fossil placers" on one hand, and the redistributional changes caused by metamorphism on the other hand. Additional studies by Krige (1966), regarding the distribution of gold and uranium patterns in the Klerksdorp gold field - the Klerksdorp region is a compact block containing six large gold mines in the Vaal reef on the north-western rim of the Witwatersrand basin - based on a three parameter lognormal frequency distribution pattern, reveals differences between the gold and the uranium distribution. According to Krige, the uranium distribution within the individual mines is less variable than that of gold. The gold distribution is less variable than that of uranium for the region as a whole. Considering the gold and silver distribution and the fineness value of gold from the two Witwatersrand gold mines, Von Rahden (1965), in "Apparent fineness values of gold from two Witwatersrand gold mines" states the following: "Variations in the apparent fine-

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ness values of gold, 1000 Au (Au + Ag) in crushed ore, from two Witwatersrand conglomerates (Ventersdorp Contact Reef and Main Reef) are discussed. It is shown that samples with a high gold content have a high apparent fineness whereas low apparent fineness is more typical of low-grade ores. There is no correlation of apparent fineness with sample depth, intrastope variation being marked". An explanation is given of the observed variations in apparent fineness based on the redistribution of Ag relative to Au during the "solution" stage of the modified placer theory. Additional studies by Hallbauer (1978) on the morphology and other properties of gold particles from the fossil placer gold deposits of Witwatersrand indicate that in some areas primitive plant life Thuchomyces lichenoides on the deltas, on which it was deposited, not only trapped and thus concentrated the fine gold particles but probably also reacted biochemically with the particles to determine the forms in which gold is observed in the carbonaceous remains of the plant life. In consideration of the synoptical studies presented so far, the earlier work and conclusions of Ramdohr (1958) seem to be justified, in the sense that he warned that "no simple explanation for the mineralization of the conglomerate is possible". In his study of the two rival hypotheses "modified placer" and "hydrothermal", Ramdohr concluded that uranpecherz is of detrital origin but has since been modified and partly dispersed in carbonaceous material: the apparent chemical stability of the uranpecherz is ascribed to its youthful, thoroughly crystalline nature at the time of transportation and deposition. In addition, "a large proportion of the pyrite is considered to be of detrital origin, but the remainder is polygenetic. Gold has been crystallized but migrated only a very short distance". In contradistinction to a pronounced metamorphic effect on the Witwatersrand Complex, studies on the metamorphism of the Witwatersrand pyrites by Hallbauer and von Gehlen (1983) conclude that "the influence of metamorphism on the Witwatersrand pyrites can be described as only slight and generally negligible". Present observations, however, show later mobilized pyrite forming a margin on gold pebbles (Fig. 920). In addition, pyrite pebbles often show overgrowths of pyrite exhibiting developed crystal outlines. Furthermore, in contrast to the small gold pebbles, rounded by attrition, native gold can be recrystallized. Fig. 921 shows a recrystallized native gold grain exhibiting a hexagonal outline and at the same time an extension of the gold is replacing an adjacent pyrite. Pyrite relics are preserved in the gold replacing the pyrite. These observations support that recrystallization and replacement processes under low metamorphism took place extensively in the Witwatersrand conglomerate.

Chapter 29

The Broken Hill Controversy

In contrast to the Witwatersrand System where the controversy is between syngenetic sedimentary or epithermal (hydrothermal) for Au, U and some other metals, the Broken Hill deposit in Australia has been interpreted as due to chemical sedimentation, hydrothermal replacement, metamorphic sedimentary, or epigenetic hydrothermal. More recent studies support a submarine exhalative origin and mantle metasomatism for the genesis of the Broken Hill Pb-Zn-Ag deposit. Stanton and Richards (1961), on the basis of essay data, considered the abundance of lead, zinc, copper and silver in each of the four operating mines as a function of lode stratigraphic position. According to them, correlation between elements, with the exception of lead and silver, is not good and does not support the view that each ore lens is characterized by particular metal ratio. These observations are discussed in the light of two prevailing theories of the origin of the Broken Hill lode - hydrothermal replacement and chemical sedimentation. A particular problem is the very low Cu/Zn ratios in the Broken Hill ores. Condon (1959) favours a sedimentary origin for the Broken Hill ore bodies. He believes minor sedimentary structures such as cross-bedding, ripple-marks, scourand-fill, and slumping have been observed in many rocks of the Willyama Series as well as in the Broken Hill ores. These observations support the theory that the ores are originally sedimentary deposits which have undergone little or no movement during metamorphism. In contrast to Condon's interpretation, according to Segnit (1961) in his study of the "Petrology of the zinc lode, New Broken Hill Consolidated Ltd., Broken Hill, New South Wales, Australia", on the basis of petrological studies of rocks from section 62 of the zinc lode, Ν. Β. H. C. upholds that they do not support the concept of large scale metamorphism in the formation of large scale metasomatism in the formation of the Broken Hill lode and associated rocks. Most rock types appear to be of simple metamorphic sedimentary origin. According to Segnit, garnet-quartzites originate by metamorphism of a magnetiferous sediment, sillimanite gneisses are derived from argillaceous beds while the presence of zinc spinel gahnite is indicative of zinc adsorption on clay surfaces during sedimentation. Estimated temperatures of metamorphism are consistent with the lode formation (sphalerite geother-

mometer 600° C). Segnit accepts the view that both rocks and lode have undergone contemporaneous metamorphism. Still well (1959) on the basis of detailed petrological studies carried out at intervals along a line of the lode of Broken Hill ore bodies, considers them to be of epigenetic, hydrothermal origin. The presence of sulphide minerals and manganese garnets in altered dolerite dykes which transgress the Willyama Series rocks indicates that the formation of the lode is post-metamorphism. Furthermore, wallrock alteration is indicated by the frequent replacement of sillimanite by sericite and the garnetization of quartzite by replacement of calcic plagioclase. Stillwell suggests five stages in the formation of the Broken Hill lode: (i) sericitization of sillimanite; (ii) feldspathization leading to the formation of lode pegmatite; (iii) pneumatolytic stage involving introduction of Mn and Fe; (iv) hydrothermal stage with introduction of sulphide minerals; (v) late stage alteration. It should be emphasized that Stillwell believes that these stages have an ultimate magmatic cause. As time passes on, new interpretations are put forward in such a dynamic evolving science as geosciences, especially when it comes to providing new explanations for controversial subjects. Plimer (1984), considering "The role of fluorine in submarine exhalative systems with special reference to Broken Hill, Australia", suggests that "the calcite-fluoritefluorapatite assemblage at Broken Hill indicates that ore deposition was probably from hypersaline fluorinebearing fluids which decreased in pH by base leaching reactions which released Ca 2+ and mixing with seawater promoting the rapid and simultaneous precipitation of calcite, fluorite and fluorapatite as a result of temperature and salinity decrease and pH and [Ca2+] increase". Löttemoser (1991) also emphasized the significance of exhalites for the genesis of the Broken Hill Pb-ZnAg deposits in "Trace element geochemistry of exhalites associated with the Broken Hill Pb-Zn-Ag deposit, Australia." According to him "this volcanosedimentary sequence (the Lower to Middle Proterozoic Willyama Supergroup) contains exhalites, chemical sediments which are intimately associated with the Broken Hill ore bodies. Exhalatives consist of predominantly calcite, feldspars, garnet, gahnite and tourmaline". Furthermore, neutron activation analyses 159

by Löttemoser of the sulphide ores and exhalatives reveal pronounced geochemical changes with distance to the mineralization. "Thus, exhalatives precipitated from the hydrothermal fluids distal to the mineralization scavenged these trace elements (Sc, Th, Hf, Cr and Ta) from sea waters, reflecting either slow accumulation rates or prolonged exposure of the hydrothermal precipitates to sea-water." He continues that "in contrast, low Sc, Th, Hf, Ta and Cr values of sulphide ore and proximal exhalatives are likely the result of rapid accumulation rates and rapid burial of the hydrothermal precipitates in the volcano-sedimentary pile which did not allow scavenging of hydrogenous elements." Plimer (1985), on the other hand, in addition to his submarine exhalative interpretation, proposed that the

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Broken Hill Pb-Zn-Ag deposit is a product of mantle metasomatism. He maintains that "ore deposition occurred during the highest geothermal gradient coeval with an event of bimodal basic-rhyodacitic volcanism". He also suggests that "in the mature stage of rifting, propagation of deep fractures suddenly devolatilized the mantle, releasing C 0 2 and other fluids which, together with basalts, ascended and caused crustal melting to form acid magmas. The resultant ore fluid exhalation and basaltic and rhyodacitic volcanism were therefore coeval." Plimer further concludes that "The composition of the associated basic rocks and younger alkaline rocks, the premetamorphic alteration assemblage, and the ore body chemistry all suggest that the Lower Proterozoic source area for the ore fluids was metasomatized mantle."

Chapter 30

Mount Isa Controversy

Another deposit of controversial origin is the Mount Isa mineralization in Australia. The controversy in this case is also between syngenetic sedimentary and epigenetic. This, however, is an over-simplifying version of the controversy since the discussions allow a wider participation of processes in each case. Murray (1961) reviewed some aspects of the Mount Isa mineralization. The structural evolution of the ore zones is discussed and the conclusions of Blanchard and Hall (Proc. Austral. Inst. Mining and Metall., 1942, 125, 1-60) radically modified; gravitative slumping is not considered to have played a major part in the development of the intense crenulations. According to Murray, previous theories on the mode of ore emplacement, involving at least four periods of deposition, are too complex; early pyrite is probably syngenetic, and the main Pb-Zn ore bodies, if epigenetic, were introduced by actual emplacement between the beds rather than by replacement of them. The balance of evidence suggests, however, that the mineralization is essentially syngenetic and suffered later restricted mobilization during folding and fracturing. Furthermore, the copper mineralization, about which less is known, was introduced into the healed zones of "brecciation". Finlow-Bates et al. (1977), discussing a possible primary phase in the lead-zinc-bearing sediments at Mount Isa, state synoptically the following: "Quantitative mineral data from the Pb-Zn-bearing sediments of Mount Isa, Queensland, Australia, were studied using linear correlation analysis and R-mode cluster analysis. Pyrrhotite was found to be preferentially associated with galena and sphalerite". It is postulated that "during sedimentation, formation of Pb and Zn sulphides depleted an already limited S supply to the point where the field of FeS stability was entered". This interpretation is in contrast with the widely held opinion that pyrrhotite in stratiform ores is formed by metamorphic decomposition of pyrite. They also indicate that an empirical support for the sedimentary formation of pyrite is provided by textural and quantitative data from Mount Isa and other stratiform Pb-Zn deposits. Stanton (1963), applying his methodology of studying whether there is a correlation between mineralization and sedimentary geoenvironment in which it occurs, on the basis of 800 mine assays for Cu, Zn, Pb,

As, Fe, S, CaO, MgO, C0 2 , Si0 2 and A1 2 0 3 studied the abundances and abundance relationships in the Mount Isa copper ores and lead-zinc ores. Furthermore, the relationships between C0 2 , A1 2 0 3 and S, along with other evidence, confirm that both ore types and their host rocks are genetically related, the differences being considered a function of the environment of sedimentation. He interpreted the results as being compatible with a volcanic-sedimentary origin for the Mount Isa sulphide ores. Furthermore, Stanton (1966), on the presumption that the deposition of stratiform sulphide ore deposits may have taken place either by replacement or sedimentation; the first processes involving selective substitution of material in the sediment undergoing replacement and the second simple addition. He further suggests that the genesis of such deposits may be reflected in the bulk chemistry of the non-sulphide constituents immediately associated with the ore. Selected constituents of the host rock (e. g., A1 2 0 3 , Si0 2 , MgO, C0 2 ) have been plotted on a triangular diagram and the mutual variations studied with increasing concentrations of the ore forming elements (S, Fe, Cu, Zn, Pb). On this basis it is supported that the Mount Isa, Rio Tinto and Zardu (Iran) deposits, in which ore forming elements were apparently added to the host rocks, were formed by sedimentary processes. More recent work on the genesis of Mount Isa Pb-Zn deposit is in accordance with the basic dispute of the controversy syngenetic versus epigenetic. Russell et al. (1981) support "that the large sediment-hosted lead + zinc deposits like Mount Isa, McArthur River, Navan, Rammelsberg and Sullivan, form a distinctive group characterized by stratiform, syngenetic sulphide ores that formed in local basins on the sea floor as a result of protracted hydrothermal activity accompanying continental rifting." Doubts as to the sedimentary origin of the Mount Isa deposits were raised by Fisher (1960) who, despite accepting a syngenetic origin, puts forward as his main objection to the syngenetic hypothesis, the lack of a satisfactory mechanism for the deposition of mixed sulphides in a marine environment. In contrast to the syngenetic hypothesis, Cordwell et al. (1963), on the basis of detailed exploration in an area immediately south of Mount Isa, indicate that blocks of "greenstone" have been thrust underneath the southern end of the Mount Isa ore bodies. Furthermore, 161

upward extension of faults originating in the greenstone basement are considered to be the major control for the deposition of silica-dolomite Cu ore bodies. According to Cordwell et al., this new evidence points to an epigenetic origin for the Mount Isa ore. Extensive studies of O'Meara (1961) also support that replacement of pyrrhotite by lamellar marcasite and magnetite, or by pyrite and (hypogene) carbonate, was followed by deposition of chalcopyrite and further pyrrhotine and by assemblages of lower temperature sulphides and sulphosalts. Rare aikinite and berthierite have been identified. O'Meara further points out that in depth, pyrrhotine increases at the expense of pyrite and the Ni content of the ores increases but there is no decrease in sphalerite and galena. In many respect, the mineralization is very similar to that of the adjacent Ag-Pb-Zn ore bodies. It is suggested that the mineralization may be classified as mesothermal and that the occurrence of valleriite and chalcopyrite indicates that the latter mineral was deposited below 225 °C. Another very elaborate study is presented by Solomon (1965), in "Investigations into sulphide min-

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eralization of Mount Isa, Queensland". He claims that Mount Isa deposits show very little evidence compatible with hydrothermal processes of formation, and therefore, a low pressure/temperature origin is proposed. The fundamental problem with such a hypothesis is to find a collector mechanism capable of overcoming the barrier of significant concentration, but it is assumed that sulphate reducing bacteria could accomplish this. Adsorption on clays and cation exchange are ruled out as mechanism because of their inability to overcome this barrier. Solomon adds that the movement of solids by plastic flow is considered and applied to fold zones that reflect differential response to deformation. On the basis of S isotope studies, Solomon indicated an enrichment in the heavier isotope relative to meteoric standard. The isotope spread varies from pyrite (+7 to +3%o 5 3 4 S) to chalcopyrite (+12 to +16%o 5 3 4 S). Solomon interprets these data as indicative of a biogenic origin for the sulphur in the sulphide assemblage.

Chapter 31

The Role of Brines in Metallogeny (The Tennessee Valley-Type of Deposits)

As a result of syngenetic-sedimentary and epigenetic mineralization a new trend developed or rather reached a peak in the late 60s/early 70s. The idea combines elements of the syngenetic-sedimentary hypothesis and elements of the epigenetic, in the sense that connate water trapped in considerable depth after solving salts change to brines. As a consequence of temperature gradient increase, the brine (solving metallic components) also rose, and when ascending along fractures, resulted in vein deposition or, in cases, in stratabound mineralization. However, it should be pointed out that the above mentioned schematic concept has been greatly elaborated and sophisticated interpretations based on this hypothesis have been presented. Davidson (1965), in "A possible mode of origin of stratabound copper ores", suggests synoptically the following: "The hypothesis is presented that stratabound copper originate, where chemical and structural conditions were favourable, from ascending brines which had leached the metal from primary sulphides of magmatic origin". According to Davidson, the brines themselves were derived from evaporite horizons from which they descended under gravity and after reaction with the primary sulphides became re-ascendant on reaching regions of high geothermal gradient. Furthermore, the hypothesis is argued by a wider ranging examination of the geographical association of this type of Cu deposit with the evaporite series and is applied in particular to the Kupferschiefer of Germany and Poland. In his study case, Davidson refers to stratabound copper deposits and not to vein-type, and the mineralization deriving from primary sulphides by leaching. Nevertheless, his interpretation involves brines, leaching from solution of metallic elements and due to an increase in geothermal gradient as ascent of the brines; in his case, instead of vein deposits stratabound mineralization resulted. So in the pioneer study of Davidson, all the basic components of the aforementioned "schematic concept" are first introduced. A more intensive in scope review is the work of K. C. Dunham (1970), in "Mineralization by deep formation water: a review". Dunham presents some aspects of the knowledge of the chemistry of deep waters of sedimentary formations. Hypersaline brines may evolve from connate water of marine origin by osmotic filtration, or from meteoric water by solution of

evaporites in the course of sinking to great depths in sedimentary basins. The chemistry of mineralizing fluids indicates that these could be derived from deep formation waters of enrichment in K, in metals, and possibly in F: the temperature of >300 C indicated could only be attained by formation water if it sank to depths >20 000 feet before rising to the site of the ore deposition, or if igneous intrusions or a much enhanced geothermal gradient provided heat closer to the surface. Furthermore, according to Dunham, the metals may travel in solutions depleted in reduced sulphur, probably as chloride complexes. Sulfide formation may occur when an environment of active anaerobic or non-anaerobic sulphate reduction is reached, or when mixing with sour petroleum fluid occurs. Concluding, Dunham supports that mineralization occurs in relation to limited channels of substantial flow (controlled by tectonics and hydraulic conditions) rather than in static traps. Another elaborate example of mineralization by chlorite-rich brines and their possible role in ore formation is discussed by Bush (1970). According to him, the chemical composition of inclusion brines in Mississippi Valley type Pb-Zn deposits, is of significance for their genetic understanding. The evolution of the Na-Ca-K-chloride brine in arid coastal plain (sabkha) sediments of the Trucial Coast state of Abu Dhabi is described and the similarities and differences between this brine and inclusion brines are considered. Furthermore, the distribution of these ores and of the sabkha deposits are compared and a close correlation is pointed out. The problems of the source of H 2 S and of heat are considered: a mechanism for the inorganic reduction of evaporitic sulphates by hydrocarbons satisfies the shallow depth of origin for the sulphur in the sulphides and produces enough heat to increase the brine temperature to the required level. Bush also considers the possible role of the brines as the origin of the deposits. Tooms and Rugheim (1969), in "Additional metalliferous sediments in the Red Sea", support that analyses for Zn, Cu and Mn in gravity core sediments indicate abnormally high concentration in one of the cores up to 7% Mn, 0.30% Zn, 0.26% Cu. There is a marked correlation between Zn and Mn in all the ores. Furthermore, it is suggested that at least some of the metal in the sediment has been derived from extruded brines 163

via the sea water and, in general, metal content and Mn/Zn ratio in the sediments decreases from the source. Degens and Ross (1969) correlated heavy metal metallogeny with hot brines of the Red Sea. Pohl et al. (1986), considering the metasomatic siderite deposits in North Africa (Maghreb) suggest a new genetic model: essentially hot brines from buried evaporites have reacted with their host rocks during early diageneses, becoming acid and reducing. Thus, "allowing solution and transport of Fe and other cations, the solutions were channeled toward the apical parts of diapirs where they deposited their solute by metasomatic processes in fractured roof carbonates." Another depositional model proposed by them includes pressure and temperature drop as well as mixing with the surface waters. It is applicable to vein-type iron ores and minor barite and fluorite. The model is comparable to Pb-Zn deposit formation in the same area and temporal and spatial relationship between the two models is suggested.

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Dozy (1970) in his geological model for the genesis of the lead-zinc ores of the Mississippi Valley, USA, also supports that elements present in these ores could be derived from pelitic and other sediments, together with part of the connate or formation fluids. It is propounded that water freed from sediments during compacting and diagenesis may not always be able to escape along permeable channels but may occasionally give unstable conditions with fluid pressure exceeding hydrostatic and approaching lithostatic values. Tectonic disturbance may then trigger a sudden settling of these sediments, causing the formation water or "hot brine" to be pushed upwards along a path of least resistance: the liquid may even slightly lift the formations above it. Furthermore, it is suggested that minerals are precipitated as the rising brine meets and mixes with cool water of differing composition near the edge of the basin. It should be emphasized that Dozy's model of mineralization does not require the assumption of magmatic activity nearby.

Chapter 32

The Role of Brines and the Mixed Fluids Hypothesis

The Mississippi Valley-type Ore Bodies, the North Pennine Ore Field and the Abbeytown Pb-Zn Deposit In their contribution "Stratigraphic habit of Mississippi Valley-type ore bodies", Beales and Onasick (1970) support that "the precipitation of these ores took place either from a single-fluid system, or it was associated with the mixing of at least two subsurface fluids". According to them, whatever the transport and precipitation system were, they must have been affected by a series of stratigraphical restraints. It is supported that the stratiform ores were precipitated in voids without metamorphism and generally with little concurrent solution of the host rock. Furthermore, it is contented that ore transportation took place on the flanks of large sedimentary basins and formed part of the normal progressive evolution of such a basin. The mixed fluid hypothesis is considered as accommodating more of the stratigraphical observations than the single-fluid model. Heyl (1969), in "Some aspects of genesis of zinclead-barite-fluorite deposits in the Mississippi Valley", supports that the deposits show noticeable variations in the composition of fluid inclusions, lead isotopes and trace elements, and their wallrocks and structural controls. It is further supported that such variations suggest important genetic differences in the deposits. It is proposed that the solutions that deposited the ores in the main districts were concentrated Na, Ca, Κ chloride brines from adjacent basins heated to about 70160°C by magmatic heat provided by deep-seated magmas. Heyl propounds that in each district the saline brine solutions are thought to have mingled with various amounts of other solutions, such as small magmatic fractions, potassium- or sulphate-rich brines from more distant basins, or local sulphate brines and meteoric waters. Heyl concludes that some Mississippi Valley ore deposits may comprise geologically similar end-members formed by different complex genetic paths, particularly regarding the solutions and fluids concerned. Dozy (1970), in "A geological model for the genesis of the lead-zinc ores of the Mississippi Valley" also suggests, as already mentioned, that "elements present in these ores could be derived from pelitic and other sediments, together with part of the connate or formation fluids. In addition, it is suggested that water freed

from sediments during compacting and diagenesis may not always be able to escape along permeable channels but may occasionally give unstable conditions with fluid pressure exceeding hydrostatic and approaching lithostatic values. Tectonic disturbance may then trigger a sudden settling of these sediments causing the formation water or 'hot brine' to be pushed upwards along a path of least resistance". Minerals are precipitated as the rising brine meets and mixes with cooler water of different composition near the edge of the basin. Dozy's model does not require the assumption of magmatic activity nearby. As it was the case with the Mississippi Valley deposits, as mentioned, fluid mixing was proposed as a mechanism of transportation and deposition, Sawkins (1966), in "Ore genesis in the North Pennine orefields, in the light fluid inclusion studies", suggests fluid (solution) mixing. Fluid inclusion data with a salinity of 20 equiv. wt% NaCl and Na/K ratios between 6.8 and 12.4. Baryte formed at 130°C to less than 50"C from solutions with Na/K ratios between 15.3 and 46.0. Sawkins furthermore provides an explanation of the fluorite-baryte zoning in this ore field involving mixing of juvenile hydrothermal solutions with Ba-rich connate waters. Major ore bodies are also located in or near the zone of mixing and may have resulted from the reduced S of connate waters precipitating the base metals of hydrothermal solutions (implying in this case juvenile solutions). Sawkins further suggests a similarity of fluid inclusions and geological data from the North Pennines and the Mississippi Valley deposits and a similar origin is proposed. As already mentioned, Dunham (1970) in his review concerning mineralization by deep formation waters, suggests that "hypersaline brines may evolve from connate water of marine origin by osmatic filtration, or from meteoric water by solution of evaporites in the course of sinking to great depths in sedimentary basins. The chemistry of mineralizing fluids indicates that these could be derived from deep formation waters by enrichment in K, in metals and possibly in F: the temperatures of >300° C indicated, could only be attained by formation water if these sank to depths greater than 20.000 ft. before rising to the site of ore deposition, or if igneous intrusions or a much enhanced geothermal gradient provided heat near the surface." Dunham further states that the metals may travel in solutions de165

pleted in reduced sulphur, probably as chlorite complexes. Sulfur formation may occur when an environment of active anaerobic or non-aerobic sulphate reduction is reached or when mixing with sour petroleum fluid occurs. Another case exhibiting similarities with the Mississippi Valley-type Pb-Zn deposit (which is due to mixing of fluids), is discussed by Varvill (1959) from Abbeytown lead-zinc mine in Eire. According to him, the ore is restricted to a dolomitized zone in the highly faulted Lower Limestone Group of the Carboniferous Limestone. The deposits are largely controlled by small anticlinal flexures and vertical joint planes in close proximity to a persistent bed of calcareous grit,

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which may have channeled the ore solutions due to the relatively greater porosity. As Varvill points out, there are no well-defined parent fissures and no igneous rocks in the area. The mineralization consists of sphalerite, galena, pyrite and, rarely, of fluorite and baryte, and shows many points of similarity with the Pb-Zn deposit of the Mississippi Valley type, as pointed out. Many of the sedimentogenic ore deposits were previously thought to be hydrothermal-magmatic, in the sense that igneous activity, even when not apparent, was supposed to be involved. Particularly, the deus ex machina assumption of granitic magma was considered as a plausible interpretation.

Chapter 33

Lateral Segregation Processes

Roving through the international literature, the reader comes across study cases on lateral secretion, segregation and accretion leading to ore body formation. Despite the fact that these terms are not synonymous, they share a common process, namely the segregation of dispersed and interspersed elements or compounds to form either small aggregates or nodular (nugget) sized bodies or veinform mineralizations with gangue which might by itself be a product of lateral secretion. Calcite veins in marbles and quartz veins in chlorite schists are examples. Considering lateral secretion processes and thenpossibilities and limitations for the formation of various ore bodies, the process is differently evaluated by different researchers. Both Drescher-Kaden (1948, 1969) and Augustithis (1962, 1973, 1993) consider lateral secretion a significant process. Pegmatites are believed to be lateral segregations (exudations) under conditions of metamorphism and ultrametamorphismgranitization. A number of processes seem to be operating concurrently: (i) increase of solubility in the intercapillar-intergranular (see Drescher-Kaden, 1964); (ii) lateral secretion, diffusion of solutions to loci of "space available" solution gradients; (iii) solution mobility along directions of penetrability; (iv) tectonic mobilization (generations of zones of tectonic disturbances - forming the loci of "space availability"); (v) concurrent with the tectonic mobilization, generation of zones of friction; (vi) generation of hydrothermal solutions due to friction (see Drescher-Kaden and Heller, 1961). Whereas in vein formation processes concurrent operation of most of the processes mentioned takes place, in other cases such as accretion, nodule formation solutions and solution mobility or diffusion might be important. In most cases of open space formation, such as vein formation, free circulation of solutions is prominent. In other cases, vein formation due to substitution is possible and sporadically reported. Both Drescher-Kaden (1969) and Koark (1963) presented aplite formation as products of substitution, due to solutions moving in the intergranular and in cases, in the intracrystalline in

which vein aplites, relics of the substituted rocks, are also preserved. Augustithis (1973, 1993) reported several cases of perthites due to K-feldspar substitution and the processes in microscale are very characteristic and indeed very abundant. As mentioned, both Drescher-Kaden and Augustithis explained pegmatites as exudation products of granites and gneisses which are also comparable and commensurable to the calcite and quartz veins in marbles and chlorite schists, respectively. In the course of studying pegmatite formation, some significant features must be emphasized: (i) Concentration-segregation of rare metals and rare earths is common in pegmatites (see Fersmann, 1934, and Chapters 21 and 50); (ii) Large scale replacement and metasomatic mobilization are significant (see Drescher-Kaden, 1948; Augustithis, 1962, 1973, 1993); (iii) Pegmatites are in cases detached from granites and telescope within the host rock (particularly in gneisses); (iv) Cases are observed in Cornwall, UK, where pegmatites transect (cut across) hydrothermal quartz veins; (v) Transition cases are observed of feldspar-quartz to quartz veins. The idea that pegmatites and aplites are melts derived from consolidating granitic magma after the orthomagmatic phase, is strongly disputed, mainly by Drescher-Kaden (1948, 1969, 1982) and by Augustithis (1962, 1973, 1993). Especially the formation of graphic pegmatites is, without any doubt, due to quartz replacing feldspars (extensive analysis of these processes is given in Drescher-Kaden and Augustithis, see above). Thus, a metasomatic interpretation is in accordance with all the exhibited textural patterns of granites, gneisses, aplites and pegmatites. Augustithis (1962) pointed out that the crystalloblastic sequence is more significant than their classification into granites or gneisses. Metasomatism in which lateral segregation is a participating process, is a most important process in pegmatite formation. Concurrently with the large-scale processes responsible for the main rock forming minerals (feldspars, quartz and mica), metals and other elements not very abundant in the earth's crust are segregated together in 167

an ultrametamorphic geoenvironment to form the characteristic and often complex in composition pegmatitic assemblages parageneses. This fact has already been referred to above. An approach to explain the common segregation of the rare elements in pegmatites has been made by Augustithis (1964) where the relationship of these elements in accordance to the empirical laws of the periodic system was emphasized. However, the common segregation of these rare elements is not the product of a simplified process, it inevitably involves more complex processes in a wider geoenvironmental Bereich (field) where a common behaviour pattern of related elements is possible. In addition to the aforementioned general remarks, certain study cases will be presented in support of lateral segregation and accretion processes. Boyle (1961), discussing the geology, geochemistry and origin of the gold deposits of the Yellowknife district, supports that chemical evidence shows that during metamorphism some of the elements in the greenstones were mobilized and migrated by diffusion into the shear-zone system. Extensive chlorite and chloritecarbonate-sericite schists were produced, and mineralized gold-quartz lenses and veins were formed by precipitation in dilatant zones within the shear zones. In the sediments, various elements were mobilized during metamorphism and migrated into, and were precipitated as quartz lenses in dilatant zones in faults, fractures and drag-folds. Zuffardi (1960), discussing sedimentary ores of the middle Silurian in Sardinia, describes a series of Silurian shales and phyllites containing polymetallic ore lenses which appear to be related to interbedded limestones in metamorphic aureoles around Hercynian granites. A syngenetic origin in a volcanic-free environment is suggested, together with concentration of the ores by lateral secretion. Another rather complex case of lateral secretion is described by Augustithis (1967, in "On the textures and paragenesis of the gold-quartz-tourmaline veins of Ondonoc, western Ethiopia"). He concludes that the gold-quartz-tourmaline veins of Ondonoc occur in granodiorites (products of granodioritization-granitization) and schists, and follow the non-distortion planes of the theoretical strain ellipsoid. Textural prismatic tourmaline (in the veins), a late phase "metasomatic" tourmalinization is recognized. The gold occurs mainly with pyrite (as later replacement following cracks of the sulphide, see Fig. 922), which, in turn, is considered to belong to a late phase with ankerite and prochlorite. Micromylonitization and other tectonic influ-

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ences are recognizable and bodies of country rock, e. g„ schists and tourmalinefels are engulfed and tectonically mobilized within these veins, thus rendering the interpretation of the paragenesis more problematic. In addition, it is probable that the Ondonoc veins represent "hydrothermal exudation" products of the country rocks consisting of granodiorite and schists. The fact that the tourmaline formation was not concluded in a single phase (early crystallization) and the evidence of a late-phase tourmaline metasomatism shows that the sources of boron were not exclusively or necessarily those of pneumatolytic or hydrothermal (epimagmatic) origin. A sedimentary boron source is also supported for granites by Goldschmidt and Drescher-Kaden (see Goldschmidt, 1954) and for pegmatites by Oftedahl (1964), in "On the occurrence and distribution of boron in pegmatites". Similarly, Drescher-Kaden in Augustithis (1973) and Augustithis (1973) support a sedimentary source for the boron and tourmaline formation of the Piona pegmatites near Como, Italy. In contrast to a lateral secretion derivation, or to an exclusively lateral derivation of metallic element in ore deposits, Barnes (1959) supports, as mentioned, that quantitative estimation of Zn, Cu and Fe content of sedimentary rocks near ore deposits of the Hannover district, New Mexico, the northern Mississippi Valley district, New Mexico and the San Francisco del Oro district, Mexico, indicates that the total metal content of the deposits could not have been concentrated by a process of lateral secretion from the surrounding sedimentary rocks. In contradistinction to the lateral secretion, lateral segregation and exudation processes as outlined in this Chapter, an almost comparable process is the accretion or nodule (nuggets) formation in lateritic environments. An interesting case of gold nuggets' formation is presented by Machairas (1963): "three generations of gold crystals are recognized on the nuggets, each generation being characterized by different morphology and composition. The third generation forms small but often idiomorphic octahedral crystals with 1.05% Ag, compared with 6.35% in the second generation. The gold has been concentrated in the oxidation zone and has recrystallized in the "iron hat" in the latentes and also in the clayey alluvial gravel". Goni et al. (1967) also discussed a comparable case of gold accretion (nugget formation) by auto-aggregation. Ottemann and Augustithis (1967) also reported the formation of ferroplatin nuggets containing gold in lateritic covers of dunites and birbirites (see Chapter 13).

Chapter 34

Volcanogenic (Volcano-Sedimentary) Deposits

(a) General According to the hypothesis of basaltic magma differentiation, the different types of volcanic rocks are products of differentiation of the basaltic magma (Bowen, 1928). It is therefore not surprising that metallogeny related to volcanics depends, on one hand, on the basaltic magma, and on the other hand, on the reactions of ascending melts, fluids, solutions and gases released by the basaltic magma and its differentiates with the recipient geoenvironment. The differentiation hypothesis and the related metallogeny has already been discussed in Chapter 19, for the generation of different types of volcanic rocks as direct differentiates of basaltic magma, see Augustithis (1978), in Atlas of the textural patterns of basalts and their genetic significance. In contrast to the differentiation hypothesis, Augustithis (1978), considering that basalts are products of fusion of the protolytic layer (in which the upper mantle is involved) and that the different types of volcanics are due to melts generated by friction of various crust rocks, another important factor comes to fruition for metallogeny related to volcanics and its ultimate derivation in addition to the basaltic magma. The derivation of volcanics depending on melting by friction (fusion) of rocks of varying composition (along the traverse mantle-protolyte-crust which is extensively discussed by Augustithis, 1978, 1979). Thus, the different types of metallogeny related to volcanics do not necessarily depend on a differentiating basaltic magma, but are related to tectono-volcanic events as stipulated in the hypothesis of melting due to friction. Consequently, the different types of metallogeny related to volcanics, namely submarine exhalative, submarine volcanic and epigenetic (telemagmatic) and hydrothermal (thermal springs) are directly related to the broad causes of volcanism as tectono-magmatic events. As mentioned when discussing volcanism, the generation of melts and the resultant volcanic rocks is very important; their composition though depends not only on the nature of the melted rocks but on large scale assimilation and contamination processes to which the ascending melts are subjected or which they have caused (see Augustithis, 1978, Chapter 33, The significance of assimilation in basalts and associated rocks).

Concurrently with the generation of melts from different rocks, mobilization of metallic elements takes place, either in the melts, emanations or as "hydrothermal" epithermal solutions. Thus, there are two possible, perhaps concurrently operating processes, namely mobilization of the metallic elements of the respective melts (i. e., the metallic components of the rocks which have been melted by friction) and the mobilization of metallic components from the wider geoenvironment involved in the generation of melts, passage and geoenvironment of emplacement. The above mentioned theoretical framework may assist to understand the following: (i) That certain parageneses are related to specific volcanics (depending on composition - geoenvironment of melt generation and emplacement). (ii) That characteristic and specific elements (often interrelated in accordance to the empirical laws of the periodic system) are present in a paragenetic assemblage. (iii) That a wide (range) spectrum of element mobilization may take place and that relatively rare elements (often interrelated in accordance to the empirical laws of the periodic system) are concurrently (or interruptedly) segregated which also characterizes specific volcanic terrains and types. The mobilization-concentration (segregation) of elements and the formation of specific mineral assemblages, among other things, depends: (i) On the interrelationship of elements in accordance to the empirical laws of the periodic system. This in turn implies elements with related chemical behaviour can respond similarly to created conditions and can be simultaneously mobilized (see Part III). (ii) On the chemical compound that a metallic element can be carried in melts, fluids, solutions or in gaseous state (see Chapter 60). (iii) On the temperature/pressure conditions which may influence the crystallization and stability of a mineral, i. e., the physico-chemical conditions of the crystallization of a mineral from melts, solutions or gaseous state. (iv) On the capacity of a crystal lattice formed to incorporate (accommodate) an atom of a specific radius, which in turn will determine the type of elements that can be incorporated in a crystal or the elements that could be substituted (see Goldschmidt, 1938, 1954). 169

In connection with the aforementioned factors which influence the mineralization related to a particular volcanic activity, an attempt will be made to present study cases, on one hand, depending on the type of volcanic activity (e. g., submarine volcanic, submarine exhalative - sedimentary exhalative (see Chapter 27), subvolcanic and volcanic), and on the other hand, cases will be presented of mineralizations related to particular types of volcanic rocks, e. g., basalts, andesites, dacites, rhyolites, etc. Kinkel (1966), in "Massive pyrite deposits related to volcanism, and possible methods of emplacement" presents the geological features of the massive pyrite deposits of Cyprus, Rio Tinto (Spain), Bathurst (New Brunswick) and Shasta County (California) and it is suggested that these were genetically related to submarine volcanism. Furthermore, the feasibility of volcanism as a source of sulphur and sulphides is supported by Kinkel by reference to the recently formed deposits of volcanic origin in Japan and Taiwan. The possible mechanism of formation of large volcanically related sulphide deposits and the role of colloids is supported by Kinkel's contributions. Wolff (1966) also supports a submarine volcanic origin for the Cyprus pyrite deposit. According to him, the pyrite ore deposits of Cyprus generally containing about 2.25 to 4.5% Cu, occur in the pillow lavas of the Troodos Mountain massif. The age of the ore deposition is principally Upper Cretaceous with a final phase on the Lower Tertiary and as mentioned, are believed to be formed by submarine volcanic processes. TrepkaBloch (1985) also supports that the Näsliden and Rävliden deposits in the Skellefte field, Sweden, consist of stratiform massive sulphide ores associated with submarine volcanic and clastic rocks. The ores are pretectonic and consequently "they are considered to have been formed syngenetically with the deposition of the host rocks. Banding and interlayering with host sediments are common features. Cu/Zn and Zn/Pb ratios of the ores show stratigraphically and laterally defined trends. Cu/Pb/Zn ratios correspond with those found in other deposits of volcanogenetic origin." Certain oolitic iron stones are considered to be of submarine volcanic origin as well. Schweigart (1965), in "Genesis of the iron ores of the Pretoria Series, South Africa", suggests that these deposits are submarine volcanic in origin. Simic (1963) considers that the manganese mineral deposits of Popovic-Polje in northwestern Bosnia, consisting mainly of psilomelane and pyrolusite and occurring as irregular blocks, interbedded lenses and layers in shales of the diabase-chert series are considered to have originated via submarine eruption of ultrabasic magma. As a corollary to that, Simic indicates that the Mn content of the ore, including the host sediments, is 15% and that Ni, V, Cu, Co and Cr are constantly present. Concerning the already mentioned contribution of Kinkel (1966), it should be emphasized that Kinkel 170

concludes that these deposits have been formed by volcanic emanation operating under submarine conditions. Furthermore, it is suggested that the volcanic emanations enter a near-surface submarine environment of sea-water, wet sediment or wet volcanic debris. According to Kinkel, "metallic ions and colloidal material are peptized by H 2 S, which also reacts with Si0 2 , 0 2 or Hj to produce sulphur. This sulphur then reacts with metallic ions to give colloidal sulphides. The resultant sulphide-sulphur soles are then flocculated or converted to gels in the presence of electrolytes, e. g., NaCl. These flocculates and precipitates can then settle in the absence of strong currents or may be filtered, precipitated, or absorbed by the action of argillaceous tuffs, sediments, or argillitized rocks." German (1964) presents many examples of pyrite deposits of volcanic origin. He proposed the following mechanism: first a fumarolic-solfatara stage followed by a covering mantle of volcanic and sedimentary rocks; introduction of iron exhalations resulting in massive and disseminated pyrite; lastly in the post-caldera period metal-rich solutions may enter different volcanic foci. In addition to these cases of submarine exhalative volcanic deposits, study cases of exhalative sedimentary deposits were presented in Chapter 27. In contradistinction to an exhalative sedimentary interpretation sensu stricto, Bernhard et al. (1982) reconsider the concept "exhalative". They maintain "exhalites have been defined as submarine volcanosedimentary rocks, either detrital or chemical but deposited by hydrothermal fluids (true solutions or suspensions on the ocean floor). Mineralogical similarities between exhalites, on the one hand, and matter formed by wall rock metasomatic alteration of veins and/or open space filling products (i. e., alteriles) on the other, is probably the first general observation on which one can rely. Does it imply a simple chemical deposition from a true solution, in both places, or the reworking of previously deposited alterites within the stock-work and further sedimentation on the sea floor?" It should be stipulated that Bernard et al. (1988), in "The exhalative sediments linked to the volcanic exhalative massive sulphide deposits: a case study of European occurrences", question some geological information concerning South Spain, Cyprus and the High Pyrenees (France) where massive sulphide bodies have been described and interpreted as volcano-sedimentary deposits. Bajwah et al. (1987), considering the volcanic hosted Big Cadia stratabound iron-copper deposit in central western New South Wales, Australia, support a volcanic exhalative origin of the mineralization. The deposit shows "considerable variation in the major elements Mn, Ba, Ag, Pb, Zn, Cd, Se, Co and Ni. The preferential concentration of Co and Cd in pyrite, Zn and Ag in chalcopyrite and Mn in magnetite can be attributed to variations in activities of the ions in the hydrothermal fluid at the time of crystallization of the mineral phases, or in cases such as the concentration of

Co in pyrite, depend on compatible electron spin states between Co2+ and Fe2+". They support that trace element concentrations, especially Co and Ni contents and Co/Ni ratios in pyrite (average Co:Ni = 17.1) substantiate a volcanic exhalative origin for the mineralization. Considering ore deposits formed as a result of volcanic activity, a number of study cases discussing mineralization due to subvolcanic activity and related epigenetic mineralization will be presented. Maucher (1960) considers the ore deposits associated with the Cretaceous-Tertiary igneous activity in the ore province of the eastern Pontides, Turkey, to be in the form of subvolcanic dykes, impregnations and breccias and also as submarine sediments, precipitated from volcanic exhalations or solutions genetically related to dacitic lavas. The minerals consist of pyrite, chalcopyrite, blende, galena, bornite, enargite and boulangerite. Perichaud et al. (1966) also describe a mineral paragenesis associated with microgranite dykes and compare it with similar assemblages in subvolcanic horizons in Bolivia and with Freiberg occurrences (in all these the deposition sequence has been telescoped). The minerals recorded from the Fournial mine, west of Massiac, Clermont-Ferrand, include arsenopyrite, cassiterite, sphalerite, pyrrhotite, stannite, argentiferous stannite, canfieldite, pyrite, marcasite, galena, argentiferous tetrahedrite, freieslebenite, native silver, pyrargyrite, polybasite, miargyrite and argyrodite. As mentioned, this complex paragenesis is considered to be associated subvolcanic microgranites. Another case of mineralization related to subvolcanic activity is reported by Frenzel and Ottemann (1967): "remarkable coarse-grained clear aggregates of idaite and sphalerite-wurtzite, and zoned pyrite with up to 10% Cu by weight in solid solution, occur in sulphide mineral paragenesis at the Cu-Zn deposit of Nukundamu, Fiji. A mesothermal to epithermal Cu-As sulphide deposit of subvolcanic origin shows a distinct transition to siliceous subvolcanic rock containing disseminated pyrite". Abe and Sekine (1963), in their contribution "On the relationship between massive pyritic ore deposits and hydrothermal Cu-Pb-W veins in Okenobe mine, Japan" also report of the occurrence and relationship between subvolcanic hydrothermal copper-lead-zinc-tin-tungsten quartz vein deposits and bedded massive cupriferous pyritic deposits. The mineralization is considered to be subvolcanic in origin. Besides the discussed mineralization related to volcanic and subvolcanic activity (see Chapter 27), late phase mineralizations related to volcanoes and often carriers of rare and precious elements, also deserve to be considered and some study cases are presented in the following. Cohen (1962) concludes that the gold telluride mineralization occurring in the Tavua, Fiji area, occurs in steep, flatly-dipping faults around the periphery of a

late Miocene to Pliocene caldera, and Portnov and Vel'dyaksov (1972), in "Factors in the localization of gold-silvei mineralization in the Karamken deposit of the Oknotsk-Chukotka volcanic belt, USSR" report that the Karamken deposit is located in a caldera with vent facies rocks exposed in stratified tuffs and lavas. Their conclusion is that gold-silver mineralization is confined to zones of adularization. Additional cases of gold-mineralization related to volcanism are presented by Guha et al. (1988). In contrast to a general tendency to group the deposits as vein type with varying compositions, they support a strong influence of regional lithologic and tectonic evolution of the area on gold mineralization patterns. They explain that "a syn-volcanic period of mineralization comprises both the volcanogenic massive sulphide and disseminated mineralization, and the subsequent epithermal mineralization. The latter is also associated with evolving volcanic landforms and synvolcanic intrusions." According to Trurnit et al. (1982), most of the hydrothermal Pb-Zn-Ag mineralization in the PicapampaAija polymetallic mining district in the Caldera Negra, Department of Ancash, Peru, is associated with a Neogene Caldera. They also maintain that structural and rock-mechanic controls also played an important role, however, the deposit is believed to be mainly "hydrothermal volcanogenic". Duda et al. (1977) discuss a complex mineralization of Hg in the Bubnik Hg deposit in the northern part of the Slansk6vrchy Mountains (eastern Slovakia). The mineralization is of stockwork impregnation character in the Osväska neovolcanic complex. Shcheglov (1962) in his contribution on recent deposits of Apapet Springs, Kamchatka, reports that silicified tuff from a spring bears thin films of cinnabar and iron oxides. Spectrographic analyses of the film show Hg, Sb, As, Ba, Sr, Zn, Sn, Li, Ge, Be, Cu, Pb, Ag, Zr, Cr, V, Ti and Ni. According to Shcheglov, the cinnabar deposits are currently being formed from mineralizing solutions associated with the central Kamchatka deep fault. The mentioned elements concurrently mobilized as traces with Hg, and the fact that the mineralization is associated with deep faulting suggest that element mobilization occurred along the entire "traverse" of the fault, e. g., upper mantle, lower crust, upper crust (see also Chapter 18). In contrast to the study cases which refer to volcanic processes: submarine-volcanic, submarine-exhalative, subvolcanic, etc., some other examples will be presented relating mineralization to volcanic rock types. Petrascheck (1968) holds the opinion that the origin of ore deposits of Cu, Zn, etc., in connection with basaltic magmatism, can perhaps be explained by a collector activity of sulphur from the primary basaltic magma, the metals originating from the crust. Velinov (1964) states, "late Senonian basalts and dacitic rocks with subsequent pyrite-chalcopyrite-sphalerite mineralization were followed by post-Laramian monzodiorite and 171

quartz-diorite porphyrites related to which are metasomatic magnetite-pyrite-chalcopyrite and pyrite-chalcopyrite-galena to pyrite-galena-sphalerite mineralization". In cases, the relationship of mineralization to specific volcanism and volcanic rocks is more complex. As studies of the Rhodope massif by Bogdanov et al. (1974) support, the Mn and phosphorite are genetically interrelated formations of latite-andesite-rhyolite. The sphalerite-galena formations in the eastern Rhodopes and Ossogovo are connected with volcanic and hypabyssal intrusives and appear in spatial line with the felsitoid rhyolites. The stibnite mineralizations show spatial and structural links with the short basalt dykes of the same formation, while the fluorite is spatially related to manifestations of Neogene volcanism. Bogdanov et al. conclude that these interrelations are interpreted as indirect proof of magma variations between structural blocks. As far as volcanic rock types and mineralizations are considered, rhyolites seem to be related to diverse ore mineral paragenetic associations. Kulcsär (1970) reports gold and cinnabar occurrences in the eastern margin of the Tokaj range in Hungary. The gold occurs as thin plates, up to 0.3 mm across, in montmorillonite clay dykes in K-metasomatized rhyolite tuff at the Rudabarya-Berg, 3.5 km southwest of Sätovaljauhely, in which a mineralization of Pb, Zn, Ag and As is found. In addition, 7 km southwest, cinnabar occurs as a late mineral in barytized, alunitized, silicified rhyolite and in a siliceous cement to quartzite breccia at Kiräly-Berg. Considering the types of mineralization related with rhyolites, Staatz and Griffiths (1961) report that beryllium occurs with fluorite, quartz, chalcedony, opal, calcite, montmorillonite, and residual quartz, feldspar and biotite in altered Tertiary rhyolite tuff. The beryllium deposits occur along the periphery of the Spar Mountain fluorspar district and probably are related in origin to fluorite-rich solutions derived from rhyolite magma during a more recent period of volcanic activity (nodules in the tuff (0.008-0.54% Be) contain 0.63 to 3.8% Be).

(b) Post-Volcanic Ore Formation (Modern Submarine Formation of Black Smokers and Hydrotherms) Besides the exhalative sedimentary and volcanogenic deposits discussed, a genetically interesting type of mineralization is reported by Holland (1980), entitled: "The black smokers at 21° Ν on the East Pacific Rise", in which he states the following:

172

"Early in the 1970s it was suggested that submarine hot springs should exist, and that their exit temperature might be quite high. A search for such hot springs on the Midatlantic Ridge during the course of the famous expedition in 1974 was unsuccessful. However, warm springs with exit temperatures up to ca. 25 °C were discovered shortly thereafter near the Galapagos Spreading Center, and hot springs with exit temperatures up to 350° C were discovered on the crest of the East Pacific Rise at latitude 21° Ν in the spring of 1979. At 21° Ν the hydrothermal vents have built mounds several tens of meters in length, which are surmounted by chimneys up to 5 meters in height. These chimneys consist of sulphides, sulphates, and minor silicate minerals. Wurtzite, chalcopyrite and pyrite are the most important sulphides; anhydrite is the most important sulphate. Chimney formation seems to be a fairly complicated process that involves the seawater. Extremely rapid precipitation of sulfide and sulfate minerals in the zone of mixing above the chimneys produces the plumes to which the "black smokers" owe their name. Biological activity is intense in the vicinity of the vents. Bacteria which oxidize S = to S0 4 = , Cu+ to Cu2+, and Fe+2 to Fe+3 dominate the base of the food chain. Among the larger animals, fish, crabs and tubeworms are most prominent. The exploration of the new ecological niches around the vents continues to be a rich source of new insights for marine biologists. The vents at 21° Ν are too small to be of economic interest. However, exploration along the East Pacific Rise during the summer of 1980 turned up a number of other hydrothermal vent areas; the most promising area, between ca. 13° and 15° S latitude on the E. P. R. remains to be studied. It is possible that deposits in this area will contain commercially interesting quantities of sulphides, and that they will tum out to be present-day analogues of the copper deposits on Cyprus." It should be pointed out though that Holland supports that the "black smokers" could be the present-day analogues of the copper deposits on Cyprus; however, in addition to the mineralization currently occurring at the bottom of the Red Sea (see Mangan et al.), Lebedev (pers. comm.) reports the association of welldeveloped micro-crystals of wurtzite with sphalerite and pyrrhotite from the Gulf of California. The modern submarine "hydrotherms" support that subcrustal (?mantle) metallic element mobilization of Cu, Zn, Ag and Au might be related to deep fissure systems and perhaps could also account for the formation of some of the deposits referred to as "exhalative type".

Chapter 35

Consideration of Certain Aspects of Banded Iron Formations (BIFs) with Emphasis on Precambrian BIFs

Considering the diversity of opinions and interpretations concerning stratabound deposits (see Chapters 4 and 34), it is logical that BIFs should follow these Chapters. The controversies and diverse interpretations have been presented as part of the stratabound deposits discussion, and some of these aspects are also applicable in the case of BIF (banded iron formations), e. g., sedimentary versus volcanogenic, synsedimentary versus exhalative submarine, the role of bacteria. Appel and LaBerge (1987) emphasized "Precambrian iron formation is an extremely variable rock type that was deposited over along span of early Earth history. Because the rock is variable, and was deposited in a wide variety of environments, many researchers do not agree on a common definition of iron-formation. Some restrict the definition to chert-bearing rocks; some include pyritic slates; others include sulphiderich tuff or graywacke". Since the presentation of an extensive discussion of BIF is outside the scope of this volume, only a few aspects of this subject will be presented with emphasis on the genetic processes involved in the formation of this variable rock. Appel (1987), in "Geochemistry of the early Archean Isua iron formation, West Greenland" records that this formation is unique in age (3.8 b. y. old), and it resembles younger iron formations hosted in basic volcanic rocks. He emphasizes that this iron formation is associated with stratabound copper sulphides and with stratabound scheelite mineralizations. In contrast, Lepp and Goldich (1964) and Lepp (1987) consider the origin of Precambrian iron formations and distinguish them from younger counterparts on the basis of the following chemical characteristics: (i) "With few exceptions Precambrian iron-formations are much larger and contain more Si0 2 and less A1203, P, Ti and trace elements than their younger counterparts. (ii) They are remarkably uniform in iron tenor and overall composition irrespective of age, structure, or association. (iii) Although the various mineral facies have distinctive chemical compositions, the differences are related to the iron content of the dominant iron minerals and not to different depositional environments. (iv) Si0 2 and CaMg(C0 3 ) 2 , the dominant gangue minerals, exhibit an inverse abundance relationship."

In an attempt to suggest model(s) for BIF it is necessary to consider the initial and immediate source material which formed it. Lepp and Goldich (1964) propose a low P 02 that allowed Fe2+ to travel with Si0 2 and other soluables from an adjacent weathered surface (laterite), which is considered as a possible "initial source" material to the sea. Furthermore, considering the worldwide uniformity of BIF composition, a well mixed reservoir (sea water) is proposed by them as the immediate source. As supported by Lepp (1987), "the lateritic weathering model still holds as an explanation for the differing behaviour of Fe and Si0 2 in Precambrian and Phanerozoic sediments with the proviso that the major solubles first became conservative elements of seawater." In further discussion of their model, Lepp states the following: "Our explanation for the uniform 30% iron tenor required that Ca and Fe accumulate in the initial mud in atomic proportion of 1:1. Because the 30% tenor is a primary feature of iron formations I see no reason for abandoning this model until a better explanation for the quantitative chemistry comes along. This requires that silica be considered a replacement product of CaC0 3 . Several petrographic studies support this conclusion." In support of the hypothesis of weathered material forming the "initial source" of BIF, Reimer (1987), in "Weathering as a source of iron in iron-formations: The significance of alumina-enriched paleosols from the Proterozoic of Southern Africa", provides a most convincing study case where weathered andesites could supply the Fe required for the formation of banded iron deposits. "Among the theories regarding the origin of banded iron formations, weathering has always occupied a prominent place as a potential source of the iron. However, a joint preservation of banded iron-formation basins and the source areas of their sediments has not been observed so far and thus no test of the feasibility of this idea was possible. A characteristic process during weathering is the enrichment of alumina in the residual material, accompanied by the loss of up to 60% of the original rock material in solution. Corresponding paleosols with above 20% A1203 have been described from sequences as far back in time as the early Precambrian.... The association between iron-formations and paleosols is especially widespread and well-developed in the 2,1-2,3 Ga-old 173

Transvaal Supergroup of the Kaapvaal Craton. True banded iron-formations and more siliciclastic ironformations occur over wide areas in close stratigraphic contact with alumina-enriched paleosols such as the one developed on the Hekpoort andesites over an area of about 100.000 km 2 . The original uncompacted thickness of the preserved paleosol was up to 10 m, of which about 55-60% were removed during weathering. An iron loss of about 80g/kg of source rocks was calculated. The total mass of iron removed from the paleosol was about 60 χ 10 9 1 whereas the iron ores overlying the Hekpoort andesites away from the paleosol contain 18 χ 109 t Fe 2 O s . The good agreement between the two figures suggests that it is possible to obtain the iron required for the iron-formation directly from weathering solutions". Rai and Paul (1987) also support the "weathering" hypothesis in their studies entitled "Geochemistry of BIF, iron ores and associated lithologies from JamdaKoira Valley of Bihar, India". The following extract outlines the process. "Mature terrestrial paleoweathering of pre-existing basic lavas and associated acid effusives, seem to have provided the main bulk of iron and silica which were precipitated primarily as chemogenic sediments in a shallow intra-continental basin bordering on a land-mass of low relief. The BIF has been classified as a unique representative of Lake Superior-type oxidates, characteristically enriched in titanium but impoverished in all other minor- and traceelements and deposited possibly under very specialized Precambrian sedimentary environmental conditions." Yet another contribution in support of the "weathering" hypothesis is the paper by Murthy (1987), "Origin and cyclothemic pattern in the Precambrian banded iron formation of Donimalai area in Sandur schist belt, Karnataka State, India". Synoptically it is reported that "The BIF was originated probably by the chemical weathering of a mafic rock, and the silica was transported in solution/colloids to a marine environment and deposited due to changes of Eh and pH. Microbiota which was noticed in the BIF and cherts was believed to have played some role in the origin of BIF". It is interesting to point out that microbiota also played a role in the BIF. Besides the significance of weathering in the BIF (cases of weathering of volcanic rocks were considered), the association of BIF with volcanics is perhaps more than coincidental. BIF are often hosted in volcanic rocks. Stratigraphic and geochemical studies of Archaean iron formations and sedimentary host rocks in the Sturgeon-Savant Belt, Ontario, by Shegelski (1978, 1987) suggest that sulphide iron formations accumulated with proximal turbidites in deep water of felsic volcanic centres whereas oxide iron formations were deposited with distal turbidites down paleoslope on a deeper basin plain. Hoppe et al. (1987), considering the Precambrian BIF in the Serra dos Carajäs/Parä State (Brazil) explain that those jaspilites of late Archaean age are exclu174

sively composed of oxide-facies rocks and show a distribution indicating a long wide shallow basin with quite sedimentary conditions. Underlying and probably overlying are thick piles of basic volcanic rocks with affinities for ocean basalts. In contradistinction to Hoppe et al., Gibbs and Wirth (1987) maintain that although the Carajäs iron formations are Archaean and overlie a dominant basaltic sequence, they differ significantly from the Algoma-type iron formation in their apparent continuity over a broad region and in the continental nature of the associated volcanic and clastic sedimentary rocks. Furthermore, Stulchikov (1990) in his study of a banded iron formation within a Precambrian greenstone complex of an old syncline structure of the Ukrainian Shield reports that the iron banded formation is associated with a volcano-sedimentary ophiolitic rock complex. Stulchikov reports "that the iron-banded rocks can be divided into two varieties according to their mineral composition, i. e., a magnetite-silicate-quartz variety in the lower par of the section associated with basic lavas, and a magnetite-carbonate-silicate-quartz one in the upper part of the section associated with keratophyre alumino-silicate formations. The first variety is characterized geochemically by the predominance of ferric iron in magnetite, the second one is ferrous (in siderite and ankerite)." The relationship between BIF and volcanics is also pointed out by Foster and Gilligan (1987), in "Archaean iron formation and gold mineralization in Zimbabwe", where it is supported that iron formations are preferentially associated with volcanic rocks, particularly those of komatitic and tholeiitic affinities. As a corollary to the role played by volcanics in the formation of BIF, and considering the environment of deposition, a study case is presented by Fralick (1987), in "Depositional environment of Archaean iron formation: inferences from layering in sediment and volcanic hosted end members". The following part is quoted, clearly indicating that in addition to sediments, volcanics associated with BIF are of importance: "Layering in Algoma Type, oxide-facies iron formation (I. F.) was studied in the Beardmore-Geraldton Terrane of Superior Province, Canadian Shield. Where the I. F. is associated with sedimentary units it consists of finely microlaminated hematite, magnetite and jasper packages contained within coarsening upward, chemical to clastic dominated successions. Volcanic associated I. F. exhibits thicker and more mineralogically homogeneous banding than the clastic associated variety. The differences in mineralogy, internal structure and thickness of the layering may be related to proximal (volcanic associated I. F.) to distal (clastic associated I. F.) changes in depositional style away from areas of exhaling hydrothermal fluids. I. F. precipitated near vent fields formed interflow rainout deposits associated with volcanic ash layers, whereas distal chemical sedimentation built deposits found where, under normal circumstances, pelagic muds should dominate in

submarine fan successions." Furthermore, Bai Jin and Luo Hui (1987), in "Tectonic control of BIF in the Wutaishan region, Shanxi Province, China", support that the BIF occurring in the Archaean Wutai Group is closely related to mafic volcanics which were the product of intense volcanic activity and evolved from the tholeiitic series of the calc-alkaline series; which after consideration of the chemical composition and laminated structure, is comparable with that of Algoma-type. Volcanism and associated exhalation are, according to them, the controlling factors in composition and depositional facies of the BIF. In contrast to the Algoma-type of BIF described, where sedimentary and volcanic rocks were "accumulated" in tectonic belts, the Lake Superiortype (deposition environment on continental margins) show often lower contents in minor elements. However, both types as Gross (1990) reports, show similar average contents in weight per cent of iron (30), silica (48), calcium (1.6) and magnesium (1.5). It should be pointed out though that as Gross states, "The average contents of minor elements are usually higher in all facies of Algoma type. Strong correlation coefficients were found in all facies for A1203, KjO, NajO, Ti and Zr, and in some pairs of elements in the Cu, Zn, Ni, Co, Cr, Ti and Zr group... Trends in the association and content of elements are similar in the major types of iron-formation but are more clearly defined in Algoma type facies. Differences in composition of oxide facies of different age are attributed to environmental factors in the depositional basins and not to alternative primary sources of constituents." In addition to both the Algoma and the Lake Superior types and geoenvironment of deposition described, as already mentioned, Gibbs and Wirth support that the Carajäs BIF differ significantly from the Algoma-type of iron formations in their apparent continuity over a broad region and in the continental nature of the associated volcanic and clastic sedimentary rocks. Another geoenvironment of BIF is presented by Wilson and Hyndman (1990), in "Tectonic interpretation of the Archaean lithologic package enclosing iron-formation in the Southern Tobacco Root and Northern Ruby Ranges of Southwestern Montana, USA", who suggest that the Archaean lithologies including iron-formation originally resided in a fore-arc basin setting. By way of contrast to the mentioned geoenvironments of banded iron formation, Schieber (1987), in "Small scale sedimentary iron deposits in a MidProterozoic Basin: viability of iron supply by rivers" presents a diverse type of iron formation consisting of extensive horizons of pyritic shales in the MidProterozoic Belt basin geoenvironment of Montana, USA. In particular, Schieber states the following: "These pyritic shale horizons may contain several hundred million tons of iron in pyrite. The source of the iron is problematic, and it is proposed here that it was introduced into a basin by continental runoff. Pyritic shale units were deposited after major regressions (or

pulses of coarse clastic sedimentation), thus suggesting a mode of iron introduction similar to Phanerozoic oolitic ironstones. The sedimentary record of the Beltian sequence indicates semi-arid climate, a hinterland of low relief, and low sedimentation rates. A combination of these parameters with data from recent environments provides physical and chemical constraints on a 'fluvial' model of iron introduction. With such a model the average size of the drainage basin and the amount of introduced iron can be estimated". In addition to the discussion of geoenvironmental conditions of banded iron formation, another aspect, the palaeo-atmospheric conditions in Precambrian times, deserves to be considered briefly. It is often stated that Precambrian BIF must have formed under different conditions in comparison to later times. This concept is incompatible with the "Huttonian principle of Uniformitarianism". However, our references about palaeo-atmospheric conditions at the time of BIF are very difficult to substantiate. In this connection only some sporadic references will be used and the picture presented is indeed fragmental, i. e., utterly incomplete. Despite the inherent difficulties in the study of Archaean and early Proterozoic atmosphere, perhaps the consideration of iron oxides might be useful. LaBerge et al. (1987) presented a model: "... iron formations originated almost entirely as a biological precipitate of fine grained ferric hydrate and, at least part of the silica as siliceous frustules. Diagenetic and metamorphic modifications resulted in dehydration and recrystallization of fine ferrihydrite to form hematite crystals in layers and granules that had little organic matter, and degradation of organic matter in dysoxic muds to generate Fe2+ serving as a source for the development of magnetite and spherical siderite grains. Although hematite and siderite were the major iron minerals of the Early Proterozoic and Archean ironformations, the diagenetic iron silicate minerals coexisted metastably in the sedimentary environment, but appear to have reacted during late diagenesis and lowgrade metamorphism to form widespread magnetite. Magnetite typically nucleated and formed overgrowths on diagenetic hematite crystals. Magnetite and iron silicates, such as minnesotaite and stilpnomelane, appear to be the equilibrium phases of ferrous and ferric iron at elevated temperatures. The proposed biological precipitation may explain: (1) the presence of hematite iron-formations in the Archean when the atmosphere was deficient in oxygen but locally the oceans may not have been; (2) the immense time span during which the major Proterozoic iron-formations accumulated; and (3) the scattered occurrence of small iron-formations that formed long after the "zenith" of iron-formation deposition. Under an oxygenated atmosphere in later times, such small iron-formations apparently developed in localized basins in volcanic settings where abundant ferrous iron had been added from hydrothermal springs." 175

Braterman and Grains-Smith (1987) also put forward a model for iron(III) formation (which essentially refers to the conditions of the palaeo-atmosphere): "Iron(II) near the surface of Precambrian ocean waters would have been directly photo-oxidized by sunlight, precipitating iron(III) and putting hydrogen into the atmosphere. Mechanisms for this reaction are discussed. It seems that production of Η atoms from [Fe(OH)]+(aq) would have been a dominant photoprocess. A model incorporating nine assumptions leads to estimates of amounts of iron that could have been precipitated in this way. Although rough, these estimates suggest that photo-precipitation may well have been a major iron precipitating factor for the BIFs. Since Precambrian ocean surfaces would in any case have been iron depleted, some form of upwelling would appear to have been an essential part of the mechanism of BIF genesis". Foster and Gilligan (1987), considering ferruginous sediment deposition and the change from Fe(II) to Fe(III), propose photosynthetic or ultraviolet photochemical processes. Markov et al. (1990), however, in their attempt to obtain information on the palaeo-atmospheric composition and evolution, in "Distribution and origin of noble gases in magnetites from Precambrian banded iron ores", state the following: "Noble gas isotopic geochemistry has been studied in magnetites from some skarn and carbonatite USSR deposits. Abundances of the 4 He, 3 He, 4 0 Ar, 3 6 Ar, 20 Ne and 132Xe isotopes have been studied to obtain information on the paleo-atmospheric composition and evolution in Precambrian. The research has failed to discover the primary 40 Ar/ 36 Ar ratio; in all cases, this ratio was above atmospheric due to the large input by radiogenic 40 Ar. The step-by-step annealing of samples has shown that in some cases radiogenic 40 Ar had formed in magnetites in situ, whereas magnetite from the Goroblagodatskoye deposit (Urals) skarn had trapped excess Ar. The presence of non-radiogenic noble gases in BIF magnetites suggests that part of these gases had been trapped while the mineral was being formed. Non-radiogenic gas concentrations point to contribution by surface water in the composition of magnetite-forming fluids, while He isotopic composition indicates the absence of any large contribution by non-radiogenic juvenile noble gases. Apparently, non-radiogenic noble gas concentration in Precambrian atmosphere did not differ much from those of modern atmospherics. The isotopic composition of Xe in the samples studied does not differ from atmospheric ones within a reasonable margin of error. Comparison of noble gas isotopic compositions from magnetites of BIFs, and skarn and carbonatite deposits has not identified any reliable geochemical variations in these ore formations. Magnetites from skarn and carbonatite formations contain large amounts of noble gases whose ratios and isotopic compositions reflect the significance of meteoric water for magnetite formation". 176

As already mentioned, sedimentation is an essential process in BIF. In addition to the volcanogenic contributed material, a number of researchers support that normal sedimentation was the fundamental process. In contradistinction, other researchers support chemical precipitation and yet others a combination in which chemical precipitation played a role. Adegbuyi and Olade (1990), in "The Precambrian Itakpe iron deposit in Central Nigeria" conclude that the iron deposit is a high grade cyclic-metamorphosed banded iron deposit of sedimentary (psammitic) origin subjected to intense granite metasomatism. Bronner et al. (1990), in "Geochemistry and knowledge of banded iron formations: West African shield, an example", also point out that the well-developed centimetric banding is related to sedimentary layering and could be more or less emphasized by diagenetic and metamorphic processes. In support of the sedimentary contribution to the formation of high grade iron and manganese ores from Karnataka, India, Murthy (1990) maintains that the BIF forms the protore of the high grade iron ore deposits and it exhibits several primary sedimentary structures. The BIF essentially contains alternate bands of chert and iron oxides. Furthermore, Murthy supports that the chert bands have dispersed iron oxides due to which several colors are imparted to the chert. In contrast to normal sedimentation, support for the chemical precipitation hypothesis emphasizing the uniformity of composition of some BIF over a large extent (see Lepp, 1987) suggests chemical precipitation is the essential mechanism of BIF. McConchie (1990), in "The geology and geochemistry of the Joffre and Whaleback shale members of the Brockman iron formation, Western Australia", reports that the Brockman iron formation is host to substantial reserves of iron ore and is one of the largest and most intensively studied Precambrian BIF in the world. Furthermore, he mentions that there is general agreement that it was largely deposited as a chemical precipitate. The following extract from McConchie's paper provides an overview of the deposit and its formation processes: "The data reveal that despite substantial variations in mineralogical composition across the province there is no significant variation in chemical composition. This observation implies that the recharge of iron and silica to the decompositional environment must have been remarkably uniform and that processes controlling deposition must have operated more or less uniformly over the entire depositional area. Deposition on a submarine platform or shelf is favoured and likely depositional mechanisms are evaluated. The origin of internal structures, other than banding, in the sediment is assessed in relation to the physical behaviour of colloidal precipitates; these structures are concluded to be post-depositional in origin. Diagenetic processes have enhanced some primary structures (including banding) and practically destroyed others

depending on the rate of crystallization and dehydration in each band; there is also subordinate controls." Majumder (1990), in "The Precambrian BIF of Bihar, Orissa" maintains the "Primary sedimentary structures, like current, ripple and mud cracks, indicate that BIF are chemical precipitates in a shallow water environment in which iron oxide coprecipitated with silica forming an incipiently banded rock, whose regular and uniform banded nature was due to diagenetic accentuation as revealed from petrography. The presence of skeletal crystals of magnetite within the silica and iron bands with ripples and load casts in the iron oxide bands, suggest that magnetite was the precursor iron oxide while the gel state of silica is indicated by penecontemporaneous sedimentary structures like slumping and microdessication cracks. Geochemical data also support the primary sedimentary origin of the precursor magnetites". In contradistinction, Cheng Yuqui and Wu Jiashan (1990) report that "the occurrence of two closely related, often interbedding and even interchangeable iron-rich rock types in the Archaean iron formation of the dominant low granulite facies metamorphism near Taiyao, Shanxi, China is unique. The BIF iron ore, containing Fe mostly just below 30%, is composed mainly of quartz and magnetite, and of minor ferrohastingsite and almandine, or ferrosalite and ferrohypersthene, with or without retrogressive tremolite and actinolitic tremolite. The magnetite-rich hornblendic rocks (granulites and amphibolites) are of more complex mineral constitution, and of much lower Fe content (FeO + F e 2 0 3 , 17-19%)". They go on to say: "The iron content of the two iron-rich protolithic types probably have a common volcanic source. But the BIF ore is basically of distal chemical precipitation origin, and the other probably derived from ferruginous, tuffaceous gritty sandstones, is chiefly of mechanical sedimentation origin and formed closer to volcanic activity both in time and space. They thus constitute a bimodel iron-enrichment in a rather shallow, quick yet fluctuationally sinking sedimentary and volcano-sedimentary marine basin. Both types are rather rich in LREE and it is probable that the local Archaean mantle was already depleted in these elements". It is obvious from the above that Cheng Yuqi and Wu Jiashan support a mixed origin for the BIF where chemical precipitation played a significant role. In contrast to the chemical precipitation process and the role of photosynthesis, as already outlined, bacteria have been suggested as being responsible for the biological precipitation of Precambrian banded iron. In an attempt to substantiate inference of bacterial action in banded iron formation, Robbins et al. (1987) extrapolate the catalytic role of living bacteria in the precipitation of iron and silica with the possible function of bacteria of the Proterozoic. More specifically they state: "Among the important catalytic processes for precipitation iron and silica in today's aqueous environment are those performed by iron-bacteria and sili-

ceous algae. Not only do iron-fixing bacteria and silica-fixing algae trip these undersaturated elements out of the water, but in doing so, they leave distinct mineralogical and textural evidence of their former presence. Spherules, magnetite crystals, siderite spherules, chalcedony spherules, and gelatinous granules in these rocks indicate that ancient iron- and silica-fixing bacteria and algae could have played a major role in the precipitation of iron formation". Additional studies by Robbins (1987) attempt to substantiate the existence of bacteria in BIF on the basis of microstructural-morphological studies. In "Appelella Ferrifera, a possible new iron-coated microfossil in the Isua iron-formation, southwestern Greenland", she specifically states the following: "Thin sections and acid residues from the highly metamorphosed Isua Iron-Formation, southwestern Greenland, contain rounded, carbonaceous, iron-coated microstructures. Acid residues of the microstructures under the SEM exhibit external surface morphologies with clavate and murate projections. Such finely-divided wall structures are more typical of organisms rather than minerals. The microstructures are discrete entities, about 5 μηι in size, apparently consisting of adhering individuals similar to microcolonies of modern bacteria. Features characteristic of bacterial fission and budding also are present. These microstructures are interpreted to be microfossils, here named Appelella ferrifera. A. ferrifera also forms aggregates resembling coenobial-type granules. The microfossils therefore may be the remains of a capsular, chemotrophic iron bacterium, similar to the living Siderocapsa." In another paper, Simonson and Lanier (1987) support the existence and role of microbiota in shelf or arenitic iron-formation - "Several different origins have been proposed for the stromatolitic structures found in shelf or arenitic iron-formations. The stromatolites consist mainly of chert and are closely associated with chemical arenites, as well as lesser amounts of rudite and lutite." In addition they believe, in extrapolation to shelf or arenitic iron-formation that the precursors of these sediments were primarily siliceous in composition, and that the single most important factor in the excellent preservation of microbiotas in ironformations was rapid and pervasive silica cementation. Another aspect of BIF that deserves mentioning, is the trend of changes, despite uniformity of composition within a BIF. As already mentioned, Fralick (1987) reports differences in mineralogy, internal structure and thickness of the layering in an Algomatype BIF in the Canadian Shield. The differences are attributed to proximal (volcanic associated I. F.) and to distal (clastic associate I. F.) changes in depositional style. Moreover, as also mentioned, trends of changes are reported by Shegelski (1987), discussing a BIF from the Sturgeon-Savant Greenstone Belt, Ontario, who suggests that sulphide iron formations accumulated with proximal turbitites in deep water or felsic volcanic centres whereas oxide iron-formations were de177

posited with distal turbidites down paleoslope on a deeper basin plain. Overall, BIF are regarded as having uniform composition which though can greatly be changed due to diagenetic and particularly metamorphic metasomatic processes. Considering diagenetic and metamorphic changes, LaBerge et al. (1987) support that "modifications resulted in the dehydrations and recrystallization of fine ferrihydrite to form hematite crystals in layers and granules that had little organic matter, and degradation of organic matter in dysoxic muds to generate Fe +2 serving as a source for the development of magnetite and spherical siderite grains. Although hematite and siderite were the major iron minerals of the Early Proterozoic and Archean iron formations, the diagenetic iron silicate mineral greenalite is common in some granular Early Proterozoic iron-formations. The ferric and ferrous iron minerals coexisted metastably in the sedimentary environment, but appear to have reacted during late diagenesis and low-grade metamorphism to form widespread magnetite. Magnetite typically nucleated and formed overgrowths on diagenetic hematite crystals." The diagenetic changes reported by LaBerge et al. (1987), however, refer to the early phase of banded iron formation. Another interesting case of diagenetic changes is presented by McConchie (1987) in his study of the Brockman Iron Formation in Western Australia: "The regional variation in mineralogy, compared with the constancy in chemical composition, reflects diagenetic/metamorphic modification of sediments that were compositionally uniform across the province at the time of deposition and indicates that diagenesis was largely isochemical, even on the centimetre band scale, except with respect to H 2 0, H + and C0 2 . Major diagenetic reactions are described in relation to mineral paragenesis and the importance of differing band permeabilities and resistance to ion migration is emphasized." In addition to these diagenetic changes, more pronounced changes are due to metamorphism and its accompanying metasomatic changes. From the plethora of known cases of metamorphic effects on BIF only some examples will be presented, indicative of the changes often observed in banded iron formations. Often as a result of metamorphic changes high grade iron ores are formed from the BIF which, as already mentioned, have initially about 30% iron content. Another pronounced metamorphic effect is the resultant metamorphic facies, which depending on the host or associated rocks, are attained according to the degree of metamorphism. As Adegbuyi and Olade (1987) report, "The Itakpe iron ore deposit of Okene area in central Nigeria lies within a polyphase high grade gneissic terrain as a relict of an ancient banded iron formation in the Nigerian crystalline basement of the West African Precambrian shield. The major lithologic units of the highly deformed deposit are Proterozoic biotite hornblende paragenesis (sillimanitic and gar178

netiferous) in which intercalations of ferruginous quartzite bands and concordant granite gneisses are to be found. Occasional concordant amphibolites, hornblende schists and discordant granitic and pegmatitic intrusions of the late Precambrian age are emplaced within the sequence." Considering that the Precambrian BIF are of Archaean or Proterozoic age, it is not surprising that with the lapse of such long periods of time they have been subjected to metamorphic changes. Another study case of BIF being subjected to intense metamorphic change is reported by Mahabaleswar et al. (1987) from Karnataka, India. They also described the metamorphic facies evolved as follows: "The iron formations of high grade region occur as continuous bands and as well as intercalations in the associated gneisses. They show distinct banding. In the thin section the iron formation show schistose and xenoblastic textures. There is rapid variation in the proportion and distribution of the minerals. The essential minerals that are found in these rocks are magnetite, quartz, orthopyroxenes, clinopyroxenes, amphiboles, garnet, pyroxferroite and lamellar pyroxenes. The chemistry of these minerals has revealed that the orthopyroxenes and clinopyroxenes are ferrohypersthenes and ferroaugites, respectively. The garnet belongs to pyralspite series, whereas the amphiboles are represented by grunerite, ferroan pargasite and edenite. The lamellar pyroxenes are Ca-poor and Ca-rich pyroxenes, with Ca-rich and Ca-poor lamellae in them. Pyroxferroite is an Fe-rich triclinic pyroxene. The presence of these silicate minerals is attributed to bulk composition, oxygen fugacity and pressure/temperature conditions of metamorphism. Geothermobarometric estimations give values of Τ at 650°C and Ρ at 7.5-8.0 kb." Sill et al. (1987) present cases of early Precambrian times from Hebei, China, and among other things, state the following: "BIF occurs extensively through the Archaean and Proterozoic of China. BIF from the Archaean high-grade gneiss terrain in Eastern Hebei Province and from a late Archaean terrain in Liaoning Province, northeastern China are discussed. In eastern Hebei, BIF of two ages is found. The older is associated with supracrustal amphibolites of 3.5 Ga, quartzites, metapelites, rare carbonates and quartz-plagioclase biotite gneisses, possibly of metavolcanic origin. The younger BIF is more extensive but there is less associated amphibolite. BIF occurs as enclaves or layers within amphibolite to granulite facies tonalitic to granodioritic gneisses. Oxide iron formation is dominant with quartz + magnetite ± actinolite ± cummingtonite ± orthopyroxene ± clinopyroxene. Carbonate iron formation is rare. The BIF was metamorphosed at pressures of 4.5-6.5 kb and temperatures of 600-750°C. Deformation and metamorphism have completely obscured the original mineralogy, texture and association of the iron formation." Among the metamorphic rock types (and in contrast to the high grade metamorphics) that are often associ-

ated with BIF or in which BIF are hosted, are greenstone-schists or greenstone belts of various derivation (sedimentogenic, metamorphic, volcanic or ophiolitic; see Shegelski (1987), Laajoki and Gehör (1990), and Stulchikov (1990). Laajoki and Gehör (1990), in "The mineralogy and regional metamorphism of the Precambrian iron formations in Finland" refer to greenstone-types associated with BIF in some of the Finnish occurrences. They state the following: "The low grade Kittilä occurrences are characterized by the mineral assemblage of quartz-magnetite-carbonates (mainly manganoan siderite and siderite)-stilpnomenale-iron chlorite with minor minnesotaite and biotite and their metamorphism took place below 350-400° C. All the medium grade occurrences are characterized by abundant amphibole (grunerite and hornblende) and garnet in silicate iron-formations. Reflecting difference in bulk compositions, the Late Archaean Ilomantsi and Early Proterozoic Kainuu occurrences differ from each other in the compositions and amount of carbonates; minor calcite in the former, abundant dolomiteankerite and siderite in the latter. Furthermore, the former are hornblende-rich and their garnet is homogenous almandine, while dominant amphibole in Kainuu is grunerite and garnet shows compositional zoning and is more manganoan. The geothermometers give approx. 550°C and 520°C for the Ilomantsi and Kainuu occurrences, respectively." As already stated, metamorphism may result in high grade iron ore and may also cause complications in ore body shapes by intense metamorphism and deformation and may thicken and impart tabular shapes to them (see Zhan Quisheng, 1987). Bai Jin and Luo Hui (1987) also report the following: "As a result of flexural flow, the BIF-bearing series is flattened in fold limbs, and thickened in fold hinges. Assuming that deformation occurred under plane strain, the strain measurement data indicate that the thickness of the iron ore beds, along with the country rock, was decreased by about 40% in fold limbs, and increased by about 70% in hinges. The superposition of multiple folding caused the iron ore beds to be repeated, which is why a tremendous reserve of great industrial value occurs in a well-defined district." In contrast to these rather "pure deformation" effects, tectonic influences and accompanying substances, mobilization may cause pronounced mineralogical and compositional changes to BIF. Macdonald (1990), in "Banded oxide facies iron formation as a host for gold mineralization", characteristically states: "Banded Oxide-facies Iron Formation (BOIF) is host to gold mineralization in most Archean shields of the world. There has been some controversy as to whether gold mineralization is synsedimentary or whether the metal was introduced during later deformation and hydrothermal alteration. Detailed investigation of BOIFhosted gold deposits in the Pickle Lake area, Geraldtown, and the Porcupine Camp in Northern Ontario indicates that the mineralizing event is coeval with de-

formation of the host rocks. The contrast in physical properties between BOIF and associated rocks commonly results in brittle failure of the BOIF, increasing its permeability, and permitting fluid access to the rocks. Chemical reactions between fluid and the ironrich rock induce sulphidation of iron oxide to form iron-bearing sulphides and promote gold precipitation." It should be pointed out that the synsedimentary and the hydrothermal alteration hypothesis (somewhat comparable to the controversies about auriferous conglomerates) is further enhanced by the study of the Jardine deposit, Montana, which according to Macdonald, has auriferous sediment that may be an integral component of the sedimentary sequence and has been subjected to some later, limited remobilization into structurally prepared sites. In support of the mineralization hyothesis (including that of gold), related to tectonic channeling and the subsequent introduction of solutions, Foster and Gilligan (1987) say "Gold mineralization in the ironformations commonly occurs as sulphide-rich and sometimes carbonate-rich ore shoots within zones of deformation. Brittle deformation was accompanied by the development of quartz-sulphide veins and replacement of stratiform magnetite by iron-sulphides adjacent to the veins. Pyrite, pyrrhotite and arsenopyrite are common and minor quantities of chalcopyrite, galena and sphalerite are usually evident." In contrast to the synsedimentary and hydrothermal alteration hypothesis (related to deformation of the BIFs, see also Chapter 27), Strydom et al. (1987) proposed a submarine exhalative origin for sulphides intimately related with BIF. More specifically they state: "The huge base metal ore deposits at Swartberg (PbZn) and Gamsberg (Zn-Pb) are intimately associated with the iron formation. The stratigraphy of these economically important areas is discussed and brought into context with adjacent outcrop areas. The presence of diamictite or conglomerate units suggests that minor uplift in a clastic shore environment took place during sedimentation. The abundant iron-formation and barite argue for a submarine exhalative origin giving rise to precipitation of base metals under reducing conditions." Furthermore, Gauthier et al. (1987) proposed an exhalative interpretation not only for the mineralization associated with BIF but also for the iron. An extract from their paper indicates that "The iron-formations are magnetitic and may include important amounts of iron sulfides (pyrite, pyrrhotite), carbonates (magnesite, dolomite, calcite), silicates (forsterite, phlogopite) and graphite. The magnetite-graphite assemblage is interpreted as a high-grade metamorphic product of an initially iron-carbonate facies iron-formation. The ironrich sediments are considered to be exhalative deposits emplaced along the margins of a rift basin, consistent with current models for submarine exhalative sediment-hosted base-metal deposits." 179

Chapter 36

Fluid Inclusions

(a) General Fluid inclusions in crystals have been used since Sorby's time as a type of geothermometer. With the progress of geoscience, homogenization of inclusions and studies of the composition of the solutions (fluids) in the inclusions acquired a more determinative role, in the sense that not only crystallization temperatures could be estimated but also pressure conditions prevailing at temperatures of crystallization and, what is more fashionable nowadays, the derivation of mineral forming solutions themselves can be suggested. Perhaps indicative of analytical achievements of fluid inclusion studies, possibly at present, is the procedure and the results outlined in a recent paper by Diamond et al. (1991), "Comparison of SIMS and crush-leach analyses of fluid inclusion cation ratios: application to gold-quartz veins at Brusson, Val d'Ayas, northwestern Italian Alps". They support the widely made assumption that bulk crush-leach analyses of favourable fluid inclusion samples accurately reflect the electrolyte composition of individual fluid inclusions. The affirmation of their conclusion was based on secondary ion mass spectrometric analyses (SIMS) of individual inclusions and the comparison of the analytical results to crush-leach analyses of the same natural (quartz) sample. Furthermore, homogenization of crystal inclusions has been used to investigate pulsations of temperature and pressure as well as multiphase and multistage crystallizations. The fact that different crystal phases in a mineral association exhibits differences in temperature of crystallization, supports, in addition to a gradient in temperature with depth, possible differences of derivation of solutions out of which crystal phases were formed as well. Considering the plethora of homogenization studies of inclusions, conclusions on the solution (fluid)-crystal-formation equilibria can also be derived but what is more important is that the fluids, their nature and derivation can also be determined. Furthermore, considering that crystal inclusions are fluids or gases, their composition and, possibly, their isotopic composition could shed some light on the nature and origin of fluids/solutions out of which crystallization of the minerals comprising mineral associations of a paragenesis took place. In addition, homogenization studies of in180

clusions can elucidate cases of fluid/solution interactions of invading fluids and their interaction with wallrock derived solutions. Considering the significance of replacement processes as revealed by the study of replacement textures (see Chapter 5), the fact that different minerals show different inclusion-homogenization temperatures, is supportive of the different stages of their formation (both in terms of time and derivation). Thus, to conclude the crystallization temperature of an ore mineral assemblage on the basis of homogenization of the inclusions of a single mineral phase and draw conclusions on the derivation of solutions/fluids responsible for the mineralization is an unacceptable over-simplification.

(b) Study Cases (Multiphase, Crystallizations)

Multi-Stage

The general introductory statements above can perhaps be best understood by reference to study cases based on the international literature. Piznyur (1968), in "Pressures during the genesis of the Zhireken copper-molybdenum deposit, Eastern Transbaikal" supports that multiphase inclusions show that mineralization occurred in a temperature range of 580° C to 60° C mainly from concentrated hydrothermal solutions which efferversced as a result of repeated fracturing. Furthermore, the sequence of mineralization took place at progressively lower pressures ranging from 2200 atm in the early stages to 200 atm in the final stage, clearly supporting the interdependence of Τ and P. Tchemokolev (1969), in his thermometric studies of the quartz in the lead-zinc deposits of the region of Madan, Bulgaria, based on more than 200 measurements of the temperature of homogenization of gasliquid inclusion in Gradist6 (deposit) quartz crystals showed that the quartz of the principal ore forming stage was produced between 360° C and 280° C, whereas the quartz of the last low temperature carbonate stage displays temperatures of homogenization of primary inclusions ranging between 276°C and 144°C. Thermometric studies on calcite in the upper Mississippi Valley lead zinc deposits by Ericksen (1965) support that four generations of calcite are present.

Homogenization temperatures of fluid inclusions in the calcites were determined, and indicated a progressive decline in the temperature of calcite deposition; thus, a thermal gradient for calcite crystallization is suggested. Furthermore, Schmidt (1962), in "Temperatures of mineral formation in the Miami-Picher district as indicated by liquid inclusions" based on the visual method of liquid inclusion geothermometry indicates: sphalerite I, 120-85°C; pink dolomite 96-76°C; sphalerite II, 105-83°C; calcite, 68-52°C for the temperatures of mineral formation in the district. According to Schmidt, variation in temperatures suggests two surges of solutions probably channeled by the Miami Trough graben. Yajima and Touray (1970) support that their fluid inclusions study on fluorite vein at Hamman, Achemeche, central Morocco, indicates that early calcite with fluorite crystallized around 200° C in a highly concentrated brine containing about 3% NaCl, but sulphides and quartz mainly crystallized below 100° C. Discussing the aforementioned study cases, the thermometric study on minerals must be seen in conjunction with multi-stage crystallization, pressure/temperature dependency, nature and derivation of solution-fluids as well as considering other geoenvironmental factors, some of which can in turn be understood on the basis of thermometric studies on the homogenization of fluid inclusions (e. g. geothermal gradient). Borcos (1966), in "Some considerations on the determination of the thermodynamic conditions of formation of some hydrothermal veins and deposits of the Metalliferous Mountain region", oh the basis of numerous temperature determinations made by the method of homogenization of primary fluid inclusions, investigated the sequence of deposition of the metalliferous and gangue minerals in the Hanes, Fata Baii, Municaceasea East and West, Bradisar and Staniya deposits. The temperatures correspond to an upper mesohypothermal stage. The mean geothermal gradient of the mineralizing solutions is estimated at 10°C/100 m and is the same for all the deposits. Furthermore, Idriceanu et al. (1965), investigating fluid inclusion in minerals from seven different horizons up to 240 m vertically apart (in the Bucium-Izbita mineralization, in the Vulcoi-Gorabia eruptive complex, Romania) showed that the homogenization temperature for quartz ranges from 270° C to 300° C, suggestive of an upper mesothermal character for the deposit. According to them, the geothermal gradient of the mineralized area is 6.8°C/100 m. In addition, they observed a significant parallelism between the concentration change of elements such as Cu and Ge and the frequency and temperature curves: the hydrothermal process is believed to be of pulsating character.

(c) High and Low Temperatures Determined by Homogenization of Inclusion Suggestive of Different Derivation of the Metallogenie Solutions An example of the relation of high temperatures determined by homogenization of inclusion and the derivation of the mineralizing solutions is furnished by considering a calcite mass in Greenville gneiss of the Basin Property, Faraday Township, Canada, which contains betafite, apatite, zircon, molybdenite, pyrrhotite, fluorite, albite, amphibole and biotite. According to Giblin (1955), temperatures of filling of fluid inclusions in apatite indicate a minimum temperature of formation of 365°C. The relationship between the metallogeny and the gneiss helps to explain the relatively elevated temperatures determined for the homogenization of the apatite inclusions. Furthermore, according to Grushkin (1969), in his study "Composition of gas-liquid inclusion in fluorites from the Khingan tin deposit", the composition of the inclusions was uniform but the concentration of salts was higher in early than in late fluorite. Primary inclusions in green and pink fluorites homogenized at 400435°C and 292-333°C, respectively. The fluorites also contain small crystals of cassiterite, halite, sylvine and chlorite. The relatively high temperature of the solutions and presence of F, CI, C0 2 , ammonium group and nitrates show that volcanic gases from the associated Cretaceous volcanism played an important part in the formation of the hydrothermal solutions. Another case of liquid inclusion homogenization at relatively elevated temperatures is reported by Stolyarov (1964). According to him, skarn copper deposits at Alekseyevka containing anhydrite in the deepest parts must have been formed at comparatively high temperatures (at least in the mesothermal range). Stolyarov reports that gas-liquid inclusions at Kal'mykyr were homogenized at 240° C. Furthermore, fluid inclusion studies of the polymetallic hydrothermal ore deposits of Bolivia by Sugaki et al. (1988) generally indicate that homogenization temperatures and salinities (in mostly quartz and partly sphalerite, cassiterite and barite of the tin polymetallic deposits) are comparatively high for ore deposits formed by cassiterite mineralization; frequently indicating a temperature > 300° C and a salinity higher than 20 equiv. wt% NaCl. Deposits with W-Bi and tourmaline mineralization show results ranging up to 500° C and 56 equiv. wt% NaCl. These features reveal that the hydrothermal fluid related to the Sn-W-Bi mineralization may be of magmatic origin. In contrast to the above cases of relatively high homogenization temperatures of liquid inclusion - derivatives of elevated temperatures - cases of liquid inclusion homogenization at relatively low temperatures are reported. According to Freas (1961), temperatures of vapour bubble disappearance in liquid inclusions within fluo181

rite, sphalerite and quartz indicate a temperature ore deposition over a range of 94° C to 142° C which is considered by Fraes as a reasonable first approximation. Another example where the origin of the metallogeny and the inclusion brines is supportive of relatively low temperatures is recorded by Bush (1970), in "Chlorite-rich brines from Sabkha sediments and their possible role in ore formation". Bush maintains that there is a similarity between the composition of inclusion brines in the Mississippi Valley type Pb-Zn deposits and the Na-Ca-K-chloride brines in arid coastal plain (Sabkha) sediments. Thus, the inclusion brines suggest a sedimentary origin for the mineralization of the Mississippi Valley type of Pb-Zn deposits.

182

(d) Liquid Inclusions Due to Hydrothermal and Low-Temperature Solutions Mixing Groves and Solomon (1969) report a most interesting case of liquid inclusions due to mixing of hydrothermal and low temperature solutions. According to their study, fluid inclusions in quartz and fluorite from a cassiterite-sulphide replacement deposit showed that fluorite of the same generation was deposited over the temperature range of 200-580°C. Fluorite and quartz from later fissure veins were deposited over a range of 170-380°C. Furthermore, "the gross distribution and the observed correlation between temperature of formation, salinity and alkali ratios of the fluids are explained by the mixing of initial, hot, saline hydrothermal solutions with low Na/K and Na/Li ratios with cooler, less saline, relatively Na-rich meteoric and connate waters in conjunction with heat loss by the solutions to the wallrocks".

Chapter 37

Some Aspects of the Role of Fluids in Metamorphogenic Ores

Although certain aspects of metamorphogenic ore deposits and minerals have been treated in different sections of this volume, the need for the treatment of fluids in metamorphic ores under a separate heading has become prominent. Over the last few years, the study of fluids in metamorphism has acquired great significance and a lots of contributions have appeared. Augustithis (1990), however, did not restrict the topic to metamorphic rocks, strictly-speaking, but considered the subject in relation to granitization and emphasized the interrelationship between metamorphics/granitization, in respect to fluid generation. The following are extracts from the "Atlas of metamorphic-metasomatic textures and processes": "An overview of the significance of fluids in metamorphics inevitably involves the consideration of fluids in granites, due to the recent trend to consider granitic rocks metamorphogenic. Another significant point for considering fluids both in metamorphics and granites is that the same fluids might operate in both realms metamorphism-granitization and a separation of the topic is therefore difficult if not impossible. Furthermore, an overview of the genesis and operation of fluids inevitably involves a consideration of the relation of the petrogenesis of metamorphics and granites. The following general consideration attempts historically to consider some of the significant hypotheses and scientific statements put forward. Considering metamorphism, Daly (1917) states 'New crystallization on non-magmatic rock substance is the one basic principle that seems best to express the essential idea shared by Lyell (1833) and most other geologists since 1833 ... That petrographical criterion has its counterpart in the physico-chemical criterion of the alternative definition, namely the proof of the constructive activity of solutions'. This principle has been challenged at different periods in the history of metamorphism. Several hypotheses of magma directly or indirectly involved in metamorphism (regional) have been introduced at times and the most recent "Fluids in Metamorphics" is one of them, since there is a general tendency to assign the derivation of fluids in metamorphics to a magmatic source. In contrast to the attempts to involve 'magma' in metamorphism, there have been several attempts to explain the generation of 'magma', particularly granitic

'magma', as a result of metamorphism or ultrametamorphism and in cases, through the operation of fluids." As Chapter 36 shows, in the last decades extensive studies on gas and fluid inclusions in ore minerals have been carried out, providing background information for the role of fluids in metamorphogenic ores. Furthermore, a corollary to the importance of fluids in understanding not only metamorphogenic ore minerals' concentration and paragenesis but also ore deposits' formation in general, is the concept outlined in Chapter 60, which forms an integral part of the role of fluids in metallogeny. Concerning the role of fluids in metamorphics, which actually implies metamorphogenic mineralization, a very interesting IGCP meeting (no. 291) was initiated by Petrascheck (1991) in which his followers, associates and many other scientists participated. Their initiative to organize this meeting needs to be admired and many interesting papers were presented on metamorphically generated (derived) fluids. However, remembering the many arguments I myself had with Petrascheck (especially at a meeting in Mersa Alam, Egypt, in 1979), when he was a most fervent supporter of vein deposits being magmatogenic (also emphasized in his books, see Petrascheck, 1992 and previous editions), and also the many arguments put forward by Drescher-Kaden in the 50s and 60s in support of granitization and vein deposits being metamorphogenic, together with the refusal of the scientific community at that time to accept these interpretations, I am pleasantly surprised by the change of attitude as expressed in many of the papers at the IGCP meeting. In this connection, Pohl (1991), admitting this change of thinking, should be quoted: "Ideas on the role of regional metamorphism forming new mineral deposits have changed considerably in the last years: formerly, this role was mainly seen as a destructive one (with a few exceptions like graphite). Early tentative speculations that ores may be concentrated by regional metamorphic processes were disregarded. In contrast, today it is observed that increasing numbers of mineral deposits are interpreted to be "metamorphogenic", in many cases without even discussing the reasons for this classification. There is clearly a need for recalling some criteria that may be used in such attributions." 183

However, the change of attitude by some, and especially by Petrascheck, is surprising. Remembering their arguments and opposition during the 50s and 60s and later their acceptance (some of them) of vein deposits generation as a result of regional metamorphism, and forgetting the historical sequence of events and the past in the name of modern techniques, such as advances in isotope-geochemistry and geoscience progress in general, there is need for scepticism that they avoid admitting that some decades before others had advocated these very processes, namely that many vein deposits especially related to metamorphic terrains were metamorphogenic. Also, the claim of many of the more recent original "contributions" (especially when sophisticated instrumental means and "advanced interpretations" were used), must be considered on the basis whether some of their conclusions and interpretations coincide with explanations and ideas provided and used before. The fact that additional evidence in support of explanations is proved does not mean that the explanations are novel. Progress in science and particularly in geoscience has been slow and was made step by step (and in many cases, interpretations of the past proved to be more sound than at times believed to be; see Augustithis (1992), "James Hutton, the founder of modern geology"). With the aim to draw the attention of the reader to the fact that certain explanations on the role of fluids in metamorphics either have their origin in, or are parallel to, the interpretations discussed some years ago, an attempt will be made to present the results and conclusions of some of the papers presented at the IGCP meeting in 1991. The idea of mobilization and remobilization has its roots some decades back (see Chapters 1 and 44), and the relevant section (Chapters 27-35) on sedimentogenic stratabound and vein deposits derived from them. In their interesting contribution "Source of metals in ore forming processes in the Apuane Alps, NW Tuscany, Italy: constraints by Pb-isotope data", Lattanzi et al. (1991) state: "Recent research on mineral deposits of the Apuane Alps highlighted the importance of ore forming processes that are coupled with the greenschist facies metamorphism of the Apenninic (Tertiary) orogeny (Benvenuti et al., 1989). The ore deposits include Pb-Zn-Ag veins which are spatially closely related to tourmalinites (Bottino) and to mostly stratiform pyrite-hematite-barite deposits. Textural and structural evidence argue for a syn- to late-metamorphic origin of the vein mineralization whereas the stratiform deposits (e. g., Pollone and Mte. Arsiccio) were interpreted to be of sedimentary-diagenetic origin. The metamorphic imprint includes both, remobilization effects upon pre-existing ores and formation of new deposits from metal enriched source beds. The tourmaline rich layers within phyllites and porphyriesporphyritic schists are enriched in elements such as Sn, 184

Au, Pb, Sb, Ag, As and W and could thus represent possible source beds especially for the vein deposits of Bottino". It should be noted that their interpretation is compatible with the remobilization of stratabound deposits to form vein deposits. Admittedly, the employment of isotope geochemistry and their geochemical interpretation is a contribution to the idea of remobilization. Another interesting contribution by Hladikovä et al. (1991), "Isotopic characteristics of metamorphosed ore deposits in the Silesicum (Bohemian Massif, Czechoslovakia)" is essentially in accordance to the recrystallization-mobilization of the volcano-sedimentary deposits of the past literature (see Chapter 27, and especially Chapter 34). Quoting their basic idea: "The largest Bohemian base-metal stratabound deposits are situated in the Silesicum (NE margin of the Bohemian Massif) in volcano-sedimentary sequences. Primary ore accumulations which were associated with the Devonian submarine volcanism were remobilized and redeposited during Hercynian polyphase metamorphism of greenschist (locally to amphibolite) facies. Most of the primary textural and structural features have been affected by this metamorphism and the stratabound base metal ore have been mostly recrystallized". It can be seen that their basic concept is compatible with the remobilization of volcano-sedimentary stratabound formations, as mentioned. Also in their case, the authors have employed isotopic geochemistry as proof for their explanation that the deposit has been recrystallized and partly remobilized under metamorphism. Furthermore, recognizing that metamorphism is a more intense agent of changes than processes within diagenesis-sedimentation, their interpretation is basically compatible with similar interpretations of the past concerning recrystallization and vein form deposits formed from volcano-sedimentary formations. Unfortunately, the pioneering contribution of Drescher-Kaden and Heller (1961), "Reibungswärme als Energiequelle hydrothermaler Vorgänge" (friction generated heat as an energy source for hydrothermal processes) has greatly been ignored. The generation of solution phases due to deformation is, however, an idea which has been resurrected with great justification and from recent contributions some examples will be quoted: Benvenuti et al. (1991), in "Synmetamorphic barite and polymetallic veins mineralization in the Apuane Alps, Tuscany, Italy", state the following: "Recent research in the Apuane Alps highlighted the importance of greenschist facies Apenninic metamorphism (Dt = 27 Ma, D2 = 12-8 Ma; Τ = 350-450°C; Ρ = 3-4 Kbar) to the formation of ore deposits by mobilization of elements, at least in part from pre-existing deposits/anomalies. Prominent examples of these phenomena include the synmetamorphic barite (Pollone) and polymetallic (Bottino) vein deposits of SW Apuane Alps. At Bottino, a number of veins occur parallel to the main foliation of the host Paleozoic volcano-sedi-

mentary rocks. Structural analysis indicate that vein formation occurred prior to the end of Dj deformation. At Pollone, barite (-quartz) veins, parallel to almost perpendicular to main foliation, occur through a Middle-Upper Triassic (?) siliciclastic formation. Meso- to microstructural relationships point to a preD 2 (syn-D2?) origin of (part of) these veins." The generation of vein deposits depending on deformation directions, is compatible with the mobilization of materials by solutions generated by tectonic effects and is most likely compatible with the hypothesis of Drescher-Kaden and Heller. The gold-quartz tourmaline veins of Ondonoc in Western Ethiopia are also considered by Augustithis (1967) to be compatible with tectonic genesis. As put forward, the veins occur in granodiorites and schists and follow the non-distortion plane of the theoretical strain ellipsoid (see Fig. 923 and the description there). Textural studies reveal a sequence of crystallization phases. In addition to an early prismatic tourmaline, a late phase "metasomatic" tourmalinization is recognized. The gold occurs mainly with pyrite (Fig. 922) which in tum is considered to belong to a late phase with ankerite and pyrochlore. It is probable that the Ondonoc veins represent "hydrothermal exudation" products of the country rock consisting of granodiorite and schists. Apart from the cases where deformation is related to vein mineralization, examples are reported where, in addition to deformation, brines are also involved. In this connection, as concluded by Hein and Behr (1991), in "Synorogenic ore deposition in the Variscan external belt of central Europe" (where sulphides and gold mineralization are reported), structural and temporal relationships exist between ore deposition and thrust-tectonics as well as geochemical and fluid inclusion data which point to a regime of tectonically driven ore forming solutions (tectonic brines). Furthermore, they state: "Low-salinity C0 2 -rich fluids were expelled along developing channelways from the hinterland (Saxothuringian) into the externides during the initial pulse. Metamorphic dewatering and fluid-rock interaction of deep-seated rock units lead to increased fluid salinities and metals were mobilized from a large scale uniform reservoir during progressive deformation. The ore-forming solutions ascended via deeply developed shear-zones into positions favourable for deposition. Time-temperature relationship of the mineralization fluids indicates that ore deposition represents a stage pertinent to the thrust-orogenic evolution." In contradistinction to the cases described before, Hein and Behr combine in their interpretation deformation and brines, a concept widely used before in the case of ascending brines causing mineralization along deep rift faults. In addition to the instances where the term metamorphic fluids was not further defined as far as derivation is concerned, an interesting case of synorogenic fluids and metallogeny is reported by Nesbitt (1991),

in "Synorogenic fluids in the Canadian Cordillera with implications to the formation of metalliferous deposits". Considering meteoric water and its role, Nesbitt concludes: "The results of the regional fluid studies include: (1) The majority of quartz veins in the southern Canadian Cordillera formed from deeply convected meteoric fluids. (2) The widespread presence of meteoric water in upper to mid-crustal units suggests that the brittle crust was generally close to hydrostatic; however, transient overpressuring to lithostatic values is possible. (3) The depth of penetration of surface fluids is closely tied to regional structural style with the deepest penetration (to lower amphibolite facies) occurring in extensional regimes and the shallowest (mid-greenschist) in thrust regimes." "Studies of mesothermal Au systems indicate that the deposits form from large regional crustal flow systems, the fluids involved in ore formation are isotopically and chemically evolved meteoric water, and areas where metamorphic fluids dominate the non-forming fluids are barren of mesothermal Au mineralization. Pb-Zn-Ag vein systems in the Canadian Cordillera are generally located in continental shelf facies, Proterozoic units. The vein systems have similar isotopic properties to the mesothermal Au systems, but fluid inclusions indicate higher salinities for the Pb-ZnAg vein forming fluids. These systems are believed to have been a product of similar processes to the mesothermal Au systems, but due to their continental margin origin of the regional units, the salinities of the fluids were higher and the style of mineralization was dominated by Pb, Zn and Ag." Nesbitt's concepts, particularly the meteoric water circulation in the crust, is compatible with Vernadsky's phreatic cycle concept. It should also be pointed out that a plethora of studies concerning meteoric water involved in metamorphism has already been presented. The third factor in Nesbitt's interpretation, namely the influence of tectonics and its role in bringing the other two processes into an integral system is also known, in the sense that friction generates heat which could be a source for "hydrothermal" solutions. The significance of isotope geochemistry in providing evidence for the operation of the processes as an integral system should though be credited to Nesbitt. In contrast to the cases presented where the pangeneses involved were relatively simple cases of mobilization, Graeser (1991), in "Sulphosalt minerals from Lengenbach/Switzerland: a product of remobilization process during Alpine metamorphism", presents remobilization processes due to metamorphism resulting in most complex parageneses. According to him, in Binntal County, Valais, a white dolomitic rock of Triassic age that was influenced by Alpine metamorphism, and complex sulphosalts and sulphide mineralization occurs. Graeser states, "In contradistinction to all previous hypotheses which presumed that all substances required for the formation of the sulphosalts were already present in the dolomite 185

prior to Alpine metamorphism, we postulated a larger scale hydrothermal activity from outside the dolomitic rock. It is easily recognizable from field observations that there exist two completely different types of ore minerals in the dolomitic rock: 1. part of the ores (in particular galena, pyrite and sphalerite) occur as stratiform layers in the dolomite and undoubtedly formed syngenetically with the carbonate sediment. 2. sulphosalts and other sulphide minerals like realgar, orpiment, etc., appear mainly in small cavities that are clearly produced by hydrothermal activity. Additionally, from ore microscopic studies, it became obvious that Pb-As-sulphosalts formed by reaction of galena with solutions containing arsenic sulphides, the same was the case with the formation of arsenopyrite from pyrite. All these considerations finally led to the conviction that at least arsenic and presumably thallium, too, did not belong to the original ore mineralization but were introduced from outside into the dolomite at a later moment. After an intensive search for a possible source of these elements, the remnants of a strongly altered ore mineralization consisting of the primary ore minerals chalcopyrite and tennantite was detected in the gneisses about 4-5 km farther to the south. Tennantite turned out to be the almost pure end member of the series tennantite-tetrahedrite, with a very low Sb-content. Within an area of several km2 around this locality, a large number of various, partly new mineral species were detected, all of them containing large amounts of arsenic oxide (e. g., asbecasite and cafarsite with about 60 wt% As 2 0 3 ). Obviously these minerals formed due to a remobilization of arsenic from the primary Cu-As-ore by the activity of hydrothermal solutions during Alpine metamorphism. These Asbearing solutions partly migrated as far as the dolomite sedimentary cover of the gneiss nappe where they reacted with the syngenetic Pb- and Fe-ores in the dolomite. The first Pb-As-sulphosalts formed by this procedure was jordanite, the mineral with the lowest Ascontent among Lengenbach sulphosalts. Due to consequent additional supply of arsenic, jordanite gradually was substituted by minerals with higher As-contents; the precipitation of pure As-sulphides (realgar, orpiment) occurred only after some saturation of the sulphosalts with arsenic. Other elements like thallium and the extremely rare tin (in the mineral erniggliite Tl2SnAs2S5) undoubtedly have the same origin like arsenic as proven by remarkable contents of T1 and Sn in some of the secondary As-oxide minerals in the gneisses". On the basis of the above-quoted most elaborate descriptions by Graeser, the following paragenetic associations can be recognized: (i) Galena, pyrite and sphalerite as syngenetic strataform layers in the carbonate. (ii) Remnants of a strongly altered ore mineralization of primary ore minerals chalcopyrite and tennantite in a gneiss at 4-5 km distance. 186

(iii) Within the area new minerals were detected (e. g., asbecasite and cafarsite). According to Graeser, these minerals were formed due to remobilization mainly of As in hydrothermal solution due to Alpine metamorphism. (iv) These As-bearing solutions partly migrated as far as the sedimentary cover and reacted with Pb and Fe (of the dolomite). Jordanite was formed and with the increase of As, supply minerals with higher As contents were formed and finally realgar and orpiment. Thus, (i) represents the primary paragenetic mineral association (galena, pyrite, sphalerite); (ii) represents the source paragenesis (chalcopyrite, tennantite), supplying the elements which were mobilized; (iii) is the paragenesis containing minerals formed as a result of mobilization (e. g., asbecasite and cafarsite [Ca3(Fe,Ti)3Mn(As04)6-2H20]); and (iv) is the paragenesis characterized by minerals formed from the primary paragenesis (stratabound) containing Fe, Pb, Zn and the introduced solutions (carriers of As) and is the reactive paragenesis with jordanite being formed. Also in this phase, the minerals formed, such as erniggliite (TljSnASjSg), mainly belong to the superimposed paragenesis (perhaps also the realgar and orpiment). A comparable study concerning apogranite mineralization10 has been presented by Augustithis (1982; see also Chapter 51, "Textural and mineralogical studies of the apogranites of Abu Daddab, Eastern Desert, Egypt"): "The apogranitic bodies in the Abu Dabbab area of the Eastern Desert, Egypt, are concordant and discordant with the gneisses and contain small concentrations of Nb, Y, Ta, Zr as well as Cu, Zn, Sn, Ni, Cd, Sb, Ag and (Te). Metamorphic-metasomatic interlocking quartz and feldspar intergrowths as well as idioblastic quartz and poikiloblastic orthoclase support a transformist's explanation for the origin of these apogranites. The Nb, Y, Ta, Zr present are a part of the typical "pegmatitic group of elements" in contrast to the Mn, Fe, Cu, Zn, Sn, Ni, Cd, Sb, Ag, Te characteristic of the "superimposed" intercontinental rift mobilizations (deposits) typical of the Red Sea Rift (buried basalts) The presence of the group Nb, Y, Ya, Ti, Zr (Hf), U, Th can be understood by the availability of these elements in the initial sedimentogenic metamorphic granitization environment (responsible for the genesis of the apogranites) and on their interrelations according to the empirical laws of the periodic system (see Part III)... In contrast, the group of elements Mn (Fe) Zn, Cu, Ni, Cd, Sn, Sb, Ag, Te in the Abu Dabbab apogranites is understandable as elements representing a "superimposed" paragenesis (as a geochemical rather than mineralogical concept, sulphides are also ore10

In the sense that primary, superimposed and reactive parageneses have been formed in both cases.

microscopically observed) which is characteristic for the intercontinental rift type of deposits (very abundant in the adjacent areas of the Red Sea Rift; see Tooms, 1976) and is derived by the leaching of metallic elements from buried intercontinental rift basalts. The special significance of Mn both as later impregnations in the Abu Dabbab apogranites and as independent manganese deposits in the region of the Eastern Desert support the hypothesis of "superimposed paragenesis" for the elements Mn, (Fe) Zn, Cu, Ni, Cd, Sn, Sb, Ag, Te in addition to the true apomagmatic paragenesis represented by the elements Nb, Y, Ta, Ti, Zr, (Hf), U, Th and their corresponding mineralization. As mentioned, in addition to the true apogranitic paragenesis represented by Ta-Nb minerals, zircons and perhaps by cassiterite and the "superimposed paragenesis" Mn, (Fe) Zn, Cu, Ni, Cd, Sn 11 , Sb, Ag, Te, which is represented mainly by ramsdellite (Mn0 2 ), some sulphides (?) and perhaps other minerals of this group, X-ray diffraction analysis by Vgenopoulos de-

termined the presence of nickmanite (MnSn(OH) 3 ), spiroffite (MnTe 3 0 8 ) and manganotantalite (MnTa0 6 ). It is possible that some of these minerals and particularly the manganotantalite represent a "reactive mineral phase" of the tantalium of the true apogranitic paragenesis and of the "superimposed paragenesis". As a corollary to this "reactive phase" (manganotantalite) is the presence of some Mn (determined by microprobe spot analysis) in the Ta-Nb interspersed grains of the apogranitic paragenesis." Considering that many of the interpretations provided concerning the role of fluids in metamorphogenic ores are comparable or compatible with interpretations already provided for explaining metamorphism and metallogenesis of vein formation in stratabound deposits, either new instrumentation has led to certain common interpretations or mimesis of some creative minority's work of the past is followed by the trend of geosciences in the 90s.

11

In conclusion, it should be pointed out that regarding the elements Sn and W and particularly tin, doubts are expressed whether they belong to the element group derived by leaching of intercontinental rift basalts, since, as it is well-known, Sn and W are granitophile elements. However, the fact that nickmanite occurs in the Abu Dabbab apogranites makes perhaps the problem even more complex.

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Chapter 38

Sulphur in Metallogeny

Regarding metallogeny, the role of S is of particular significance since it is a primary constituent in three mineral groups: sulphides, sulphosalts and sulphates. Wederpohl (1967), comparing the crystal chemistry of oxygen and sulphur, states the following: "The outer electrons of sulphur have the S2!34 configuration common to all group VI elements. Elements of this series become progressively more metallic as their atomic number increases. Sulphur has an electronegativity of 2.5 which is moderate, compared with the value of 3.5 assigned to oxygen. Further, in sulphur and subsequent elements, d orbitals are available for hybridization. The crystal chemistry of sulphur, and of the later group VI elements, is therefore quite distinct from that of oxygen". Furthermore, the coordination number of S is extremely variable and it ranges from unity to unusually high values in disordered structures. Characteristically, Goldschmidt (1954), considering the crystal chemistry of sulphur in relation to geochemistry, states: "the crystal chemistry of sulphur is rather complex on account of the different valency states and bond types that occur in various compounds of this element. For a short summary, it is useful to arrange the various types of bonding, beginning with typical anions of divalent sulphur, passing through the particles of semi-metallic and semi-ionic compounds to the covalent bonds of elementary sulphur and then to particles in which sulphur is joined to four oxygens in the sulphate ion. The last may be formally regarded as electropositive sexivalent sulphur ions joined to oxygen anions, but this, of course, does not give fair consideration to the actual covalent character of the sulphate ions". Furthermore, as already pointed out, S is a major constituent in sulphides, sulphosalts and sulphates and despite its relatively small abundance (scarcity) in silicate rocks of the upper lithosphere (as compared with the much greater average abundance of sulphur in meteorites), it is of great importance in ore concentrations. In addition, the geochemistry of sulphur, as stated by Goldschmidt, "is connected with some of the most important processes not only in inorganic nature, but also in the metabolism of all classes of organisms. Among inorganic processes we may mention the separation of an iron sulphide phase in meteorites, as well as calcium sulphide, the formation of sulphides of iron 188

and many other metals from magmas and from aqueous solutions in the hydrosphere, and the formation of sulphates in the zone of weathering. Further, we may mention the concentration of the sulphate ion in the ocean and in evaporation. The importance of sulphur in the chemistry of organisms is indicated not only by the necessity of sulphur compounds in every living cell, in proteins and in many other organic compounds, but also the fact that the metabolism of great and important groups of organisms, the sulphur bacteria, depends on the utilization of sulphur compounds. The metabolism of these organisms is substantially based upon the utilization of differences in thermodynamic potential, or free energy, between various mostly very simple compounds of sulphur, such as sulphates, sulphides, thiosulphates, various hydrides of sulphur and elementary sulphur. The metabolic products of sulphur bacteria can again operate important processes in organic and inorganic nature, such as changes in oxidation-reduction potentials, including local and regional consumption of elementary oxygen, accumulation of hydrogen sulphide, and precipitation of important sulphide deposits of various metals. The geochemistry of sulphur is thus connected with a great variety of problems". However, considering Goldschmidt's concept, further analysis on the basis of more recent findings is necessary concerning the sulphide-oxide level in the deeper interior of the earth (below mantle or between mantle and core) and the sulphide drops (formed by magma immiscibility) from the basaltic, gabbroid or noritic stage of evolution. These drops have the tendency to sink downwards through the liquid silicate melt because of their higher specific gravity, and they often coalesce to larger masses of liquid sulphides. Goldschmidt's concept also includes that "the downward trend of segregated heavy sulphide magmas is the reason for the scarcity of sulphur in the silicate rocks of the upper lithosphere - implying the granitic rocks and also needs further discussion in the light of granitization, the unfolding of the geological spiral and the concept of metasomatic differentiation (perigranitic metallogeny). Goldschmidt's concept of a sulphide-oxide level below mantle is hypothetical and despite the different hypothesis of metallogeny derived from deep-seated

sources (see Chapter 18), it should most likely be excluded as a major source of sulphide metallogeny related to granitic mobilizations, since it is difficult to accept the existence of fracture systems that would allow the ascent of metallogeny through the 800 km or more of the crystalloplastic forsterite dominant mantle. However, mantle derived metals are abundant in the crust (see Part III and Chapter 39). In contradistinction, sulphide formation from the basaltic, gabbroid or noritic stage of evolution by formation of sulphide drops (through magma immiscibility) which is stated by Goldschmidt, will have a tendency to sink and form large sulphide masses (which after crystallization and cooling would consist of pyrrhotite, subordinate pentlandite, chalcopyrite, chalmersite and pyrite) is widely accepted as a possible interpretation by many adherents to the magmatic differentiation hypothesis and many typical deposits have been suggested as having been formed by these processes (see Chapter 19). As examples of such deposits, Goldschmidt considered the formation of Sudbury-type deposits and Bushveld. In contrast to the interpretation of Goldschmidt and its "supporters" 12 , Cheney and Lange (1967), in "Evidence for sulphurization and the origin of some Sudbury-type ores" state the following: "Evidence, permissive and direct, is reviewed of the possibility of forming Sudbury-type ores by sulphurization, the introduction of country rock sulphur into still hot intrusions". According to them, published isotope data for the deposits and related rocks of the Palisade sill, Sudbury, Noril'sk and Dovyren (Siberia), Duluth gabbro, and the Stillwater Complex, support the conclusion that the sulphur of the ores is derived from the country rocks rather than from the intrusions, although the process is not fully understood.

It should be mentioned though that isotope-studies supporting S derivation from mantle source (but not from a sulfide-oxide layer below mantle) are presented by Robinson et al. (1982), " S u l f u r isotopes and the origin of stibnite mineralization in N e w England, Australia". Concerning the derivation of S from mantle, they state: "Stibnite mineralization in the antimony province of N e w England can be divided into Central type ores (veins of stibnite + quartz + berthierite) and Peripheral type ores of stibnite + quartz + native antimony ± berthierite. The 'Central' stibnites have 6 3 4 S C D T values between 0 and -25%o, with a large group at 0 ± 2%o (1σ). They represent precipitation from a limited supply of mantle sulfur and the acquisition of sedimentary sulfur. We consider that the different ore types were produced from distinct ore solutions derived from two immiscible melts. These originated in the deep mantle, were mobilized by tectonic activity and supplied the antimony and most of the sulfur to the ores." Vaasjoki (1985), Johansson and Rickard (1985), Fox et al. (1988), and Vervoort and White (1991), all considering the possible derivation of metals from a mantle source on the basis of isotope studies, support that mantlePb may be present in different types of ore (see also Chapter 39). 12

Furthermore, Vinogradov and Grinenko (1966), in "Isotopic composition of sulphur in the sulphides of the Noril'sk copper-nickel ores and the genesis of the ores" support that 30-50% of the total sulphur in the Noril'sk intrusives was assimilated from sedimentary C a S 0 4 that amounted to 1.5% of the mass of the intrusives. Another aspect of Goldschmidt's hypothesis concerning the distribution of S should also be more critically discussed. When studying the abundance of sulphur in meteorites or in the solar atmosphere, Goldschmidt suggests the following: "The downward trend of segregated heavy sulphide magmas is thus the reason for the scarcity of sulphur in the silicate rocks of the upper lithosphere as compared with the much greater average abundance of sulphur in meteorites or in the solar atmosphere". This statement implies that the sulphur in the silicate rocks of the upper lithosphere moved downward as segregated heavy sulphide magmas resulting in low S content for the upper lithosphere silicate rocks (?the granites). The process is considered to be due to immiscibility - magmatic differentiation. Goldschmidt's concept of granitic rocks was in compliance with a granitic intrusive magma. In contradistinction to the interpretation of sulphur distribution in the upper lithosphere silicates (~ granites) the unfolding of the geological spiral (see Chapter 1) and the debasification, and furthermore the concept of granitization, or even of anatexis, support a rather different interpretation for the abundance of S in the upper lithosphere silicate rocks. Moreover, considering granitic rocks and the abundance of sulphides in granitic skarns as well as in the perigranitic (hydrothermal) vein deposits, one should rather see S mobilization not as a downward trend of segregated heavy sulphide magmas, but perigranitic (around the granite) S mobilization, since often extensive sulphide metallogeny is related to granitic intrusions and their surrounding country rocks. Augustithis (1990) considers the processes - granitization, granite emplacement - skarnification and perigranitic metallogeny - as metasomatic differentiation. There is yet another aspect of Goldschmidt's concept of S distribution and derivation which seems to require reconsideration, namely that exhalative S due to volcanic activity is magmatic in derivation (i. e., coming from great depths), and in this respect opinions vary immensely. Suzuki et al. (1957), in their contribution "On the sulphur deposits of the Akan Sulphur Mine, Hokkaido, Japan" report that the main ore deposit of the Akan sulphur mine is in the central explosion crater of Nakamachineshiri of the volcano Meakan, eastern Hokkaido. It consists of (1) the lower banded ore with 25-30% S, (2) the bluish-green high-grade ore with 5060% S, (3) the upper banded ore and (4) sublimated ore forming the uppermost part of the deposit". Furthermore, "the correlation of the distribution of sulphur deposits in Japan with zones of volcanic activity leads 189

to tbe conclusion that such deposits are found in the region of most calcic magma". In contrast to the above mentioned case of magmatic S, Augustithis (1964, 1978), when researching a volcanic line cutting through the Danakili salt plane (where a depth of more than 2000 m evaporate thickness is reported), considers the S impregnations of a number of the recorded volcanoes along the volcanic line to be due to S emanations released by the dissociation (breakdown) of anhydrite (or gypsum) of the

190

evaporite series transected by the volcanic line (see Fig. 924), and due to the liberation of S and its mobilization as emanation by the volcanic activity. Goldschmidt's hypothesis of a deep-seated oxidesulphide level below the mantle is not disproved by the considered cases, however, the presented study cases suggest that the problems of S derivation cannot be ascribed to a source below the mantle at least for the examples discussed (see also Chapter 39).

Chapter 39

(a) Sulphur Isotope

Study Cases of Isotopes and Their Significance in Metallogeny

Studies

Isotope pattern ratios, isotope fractionation, and in general, isotope analysis have proved to be most useful for understanding the genesis of ore deposits. Stanton (1960), in "The application of sulphur isotope studies in ore genesis theory - a suggested model", discusses a theoretical pattern of 32 S/ 34 S ratios and considers it as an approximation for a systematic work on ore genesis. Rentzsch and Pilot (1966) used galena from the sulphide deposits in the northwestern Balkans for 24 sulphur isotope determinations. According to them, the resulting range of dispersion was shown to be from 6 34 S +10.5 to -16.0%o. They further support that in this respect the galenas are isotopically the heaviest: 6(PbS) > 5(ZnS) = (FeS2). These results were discussed with regard to the genesis of the deposits. Ault and Kulp (1960) support that sulphur isotope ratios indicate that most hydrothermal sulphide deposits are derived from crustal sources rather than from mantle. Furthermore, they pointed out that variations in ratios reflect differences in the source. They also maintain the following: (i) Isotope fractionation is possible in oxidation-reduction reactions, but is unaffected by transportation and crystallization processes. (ii) Specimens showing anomalous lead isotope ratios also display highly variable sulphur isotope ratios whereas those containing normal lead exhibit a restricted ratio range for sulphur. (iii) The isotope fractionation factor between coexisting sulphide and sulphate (formed under equilibrium hydrothermal conditions) can serve as a geothermometer. Cases are known, however, where isotopic studies did not lead to equivocal genetic interpretations. Sulphur isotope compositions of 550 sulphides and 10 sulphates by Dechow and Jensen (1965) from Zambia, Zimbabwe and the Katanga Province, suggest that neither a simple biogenetic origin of the sulphur nor a simple magmatic hydrothermal origin are adequate to explain the variations of isotopic composition. Gavelin et al. (1960), in "Sulphur isotope fractionation in sulphide mineralization" support: (i) Hypogene sulphides deposited at lower temperatures exhibit greater variations in isotopic composition

than those in the same deposit formed at higher temperatures. (ii) The result from approximately 300 isotope analyses on sulphides and sulphates in specimens from England, Sweden, Germany, United States and Sardinia indicate that no correlation is established between temperature stages in hypogene mineralization and isotopic composition, even where metal zoning can be determined. (iii) Significant divergences in isotope ratios in coexisting sulphides and sulphates suggest that some fractionation may take place during hypogene mineralization, particularly at low temperatures (and may be evident within a single specimen). (iv) No fractionation is apparent in supergene oxidation of sulphides to sulphate (if the sulphide contributes sulphur to sulphides redeposited below the zone of oxidation, these may be enriched in the lighter isotope). Regarding metallogeny related to sulphur, the S derivation from sedimentary source is of importance, which inevitably involves the presence of S in the hydrosphere and particularly in the oceans. The following general considerations of Goldschmidt seem of interest: "Sulphur is accumulated to a large extent in the hydrosphere, especially in the ocean, predominantly as the S0 4 2 " ion. The sulphate may be derived directly from volcanic action or from oxidation of other volatile volcanic products, or it may be derived through the oxidative weathering of sulphides in rocks. The amount of sulphate ions in the sea in comparison with the content of sulphides in the average magmatic rock is, however, so much greater, that much of the oceanic sulphate must be derived from gaseous exhalations, and a considerable part of it may be a relic from volatile sulphur compounds of a primordial atmosphere." S isotope studies reveal that "sedimentary" sulphur is an important constituent of many sulphide deposits related to sediments. The following study cases are presented to show the variable type of deposits in which sedimentary S is either a main constituent or occurs in small percentages. As mentioned, Vinogradov and Grinenko (1966) showed that 30-50% of the total sulphur in the Noril'sk intrusives were assimilated from sedimentary CaS0 4 that amounted to 1.5% of the mass of the intrusive. 191

Anger (1966) in his comparative studies of Norwegian and German sulphide deposits, "Die genetischen Zusammenhänge zwischen deutschen und norwegischen Schwefelkies-Lagerstätten unter besonderer Berücksichtigung der Ergebnisse von Schwefelisotopen-Untersuchungen", supports that the Norwegian sulphide ore deposits of Sulitjelma, Skorovas, Joma, Lökken, Röros, Folldal and Vigsnes are reviewed in regard to their geological relations, mineralogy and, in most cases, the textural interrelationships of their minerals and their sulphur isotope characteristics. Anger discusses evidence for a syngenetic origin of these deposits. Similarly, the Rammelsberg ore sulphides are reviewed form the same standpoint and 32 S/ 34 S ratio data are put forward in support of a syngenetic origin of these deposits as well. Fruth and Maucher (1966), in "Spurenelemente und Schwefel-Isotope in Zinkblenden der Blei-Zink-Lagerstätte von Gorno", studied the trace element content and sulphur isotopes in the sphalerite on the basis of extensive ore sampling from the layered lead-zinc deposits of Gorno, northern Italy. The values of trace element content and the isotope values were related to facies changes in adjacent sedimentary rocks, suggesting that deposition of ore minerals was contemporaneous with sedimentation. Considering the derivation of S in the sulphides of the lead-zinc mineralization at Tyndum, Scotland, Pattrick et al. (1983) support that the sulphide sulphur source was probably in the Dalradian metasediments where disseminated pyrite averages +6%o. Seccombe (1990), in "Fluid inclusion and sulphur isotope evidence for syntectonic mineralization at the Elura mine, southeastern Australia", support that "sulphur isotopic compensation (634S) of pyrite, sphalerite, pyrrhotite and galena ranges from 4.712.6%o and indicates a sulphur source from underlying Cobar Supergroup metasediments". Cortecci et al. (1983), in "A sulphur isotope study on pyrite deposits of southern Tuscany, Italy", also support that sedimentary marine sulphate was the ultimate source of sulphur in the ores; and Ripley (1983), considering the sulphur derivation in the sulphides of layered sills in the Deer Lake complex, Minnesota, suggests that the parent magma was not initially saturated with sulphur and that local sulphide concentrations are the result of incorporation of sulphur derived from metasedimentary country rocks. Isotope geochemistry has verified that in many cases sulphur and sulphur in sulphides may be due to sulphur bacteria. Back in 1924, Buchanan had already supported the occurrences of free sulphur in marine muds, presumably produced by the oxidation of H 2 0 by true sulphur bacteria. Similarly, Ljunggren (1960), in "A sulphur mud deposit formed through bacteria transformation of fumarolic sulphide" supports that mud containing 30-60% of amorphic μ-sulphur occurs at the bottom of Lake Ixpaco, 50 km south of Guatemala City. Due to fumarolic activity along the shores of the lake sulphur crystals and pickeringite are deposited. 192

Mud sulphur is formed by sulphur bacteria transforming IIjS into sulphur and sulphuric acid, acidifying the lake water to pH 2.3. The production of S by bacteria as indicated by a plethora of geochemical studies (only two of them are presented here) is further verified, as mentioned, by isotope studies in the case of a number of sulphide deposits where S is believed to be due to bacterial action. Jensen and Whittles (1969), in "Sulphur isotopes of the Nairn pyrite deposits, South Australia" report that sulphur isotopic values from the Nairn pyrite open-cut mine, Southern Australia, give 5 34 S of -12.8%o to -20.6%c, corroborative of geological evidence that the deposit is bacteriogenic-syngenetic in origin. Solomon (1965), in "Investigation into sulphide mineralization at Mount Isa, Queensland" reports that "the Cu-Pb-An-Ag sulphides of the Mount Isa deposits are concentrated in bands that are mostly parallel to the bedding of the enclosing sediments". The Mount Isa deposits show very little evidence compatible with hydrothermal processes of formation and a low pressure/temperature origin is proposed. Furthermore, Solomon states "that the fundamental problem of such a hypothesis is the finding of a collector mechanism capable of overcoming the barrier of significant concentration, but it is assumed that sulphate-reducing bacteria could accomplish this". As a corollary to this interpretation, 39 sulphur isotope analyses of five main sulphide minerals by Solomon show an enrichment in the heavier isotope relative to the meteoric standard. The isotope spread varies from pyrite (+7 to +3 l%o 6 34 S) to chalcopyrite (+12 to +16%c 5 34 S). Furthermore, Solomon interprets these data as indicative of a biogenic origin for the sulphur of the sulphide assemblage. Schroll et al. (1983), in "Sulphur isotope distribution in the Pb-Zn deposit Bleiberg (Carinthia, Austria)" conclude concerning the derivation and distribution of sulphur: "(i) The sulphate minerals of Bleiberg were formed from sea-water sulphate, as well as the sulphide minerals by bacteriogenic reduction. (ii) The S isotope distribution of the syndiagenetic sulphides depends primarily on the facies conditions of the sedimentation. (iii) Diagenesis and redeposition of minerals scarcely alter the sulphur isotope composition. Multistage processes, mostly in an open system with bacterial precipitation, generate increasingly light sulphides. (iv) It is in accordance with the assumption of lowtemperature genesis that sulphur equilibria do not occur during mineralization". These conclusions could perhaps be applicable to other comparable cases. In contrast to the sedimentary and biogenic S in sulphide deposits, other S isotope studies support a magmatogenic origin. Tupper (1960), in "Sulphur isotopes and the origin of the sulphide deposits of the BathurstNewcastle area of New Brunswick", on the basis of sulphur isotope analysis of more than 300 specimens of

host rocks and ore sulphides, indicates two types of deposits - massive and fissure - probably originating from two separate but related sources. The narrow range of 32 S/ 34 S values in the sulphide deposits and homogeneity of lead isotope abundance data support a magmatic hydrothermal origin. Furthermore, Gregory and Robinson (1984), in "Sulphur isotope studies of the Mt. Molloy, Dianne and Ο. K. stratiform sulphide deposits, Hodgkinson Province, North Queensland, Australia", support that "magmatic ore fluid with 5 3 4 S I S around 0%o predominated at the Dianne and Ο. K. deposits where the fluid at Mt. Molloy mixed with seawater to acquire a higher 5 34 S IS ". In addition, Petersen (1965), in his study of the regional geology of the Andes of Central Peru, summarized the general features of about 30 ore deposits representing a great variety of types, all of them are epigenetic and many of them show association with centres of igneous activity. According to Petersen, S-isotope ratios indicate a magmatic origin.

(b) Lead

Isotopes

Lead isotope studies have a wide range of application not only for solving genetic problems but also in determining age relations and, in special cases, rates and growth stages of galena crystals. Cannon et al. (1963), in "Lead isotope variation with growth zoning in a galena crystal" report that "a large crystal of lead sulphide from Picher, Oklahoma, has significant differences in isotopic composition of lead in successive growth zones". Lead isotope ratios in the parent ore fluid evidently changed with time during crystal growth. In contradistinction to crystal growth rates, lead isotope studies are especially applicable for genetic correlation of deposits. Yesikov et al. (1965) report that "isotopic compositions of lead ( 206 Pb/ 204 Pb, 207 Pb/ 204 Pb, 208 Pb/ 204 Pb) are presented for 40 galean specimens from Transbaikalian ore deposits of Middle and Upper Jurassic age. Furthermore, according to them, the general similarity of the slope of the lines of fit on 208/204 - 206/204 diagrams suggest a common or similar source for several types of sulphide-rich deposits which are associated with small intrusives, while the oxide-silicate deposits related to larger granite bodies show a somewhat different trend. In addition, Richards (1967) supports that solid source lead isotope studies show that the ores appear (in general) to be derived from sources very similar in U-Th-Pb relative abundances although Dugald River prospect, Queensland, and Broken Hill, New South Wales, are over 800 miles apart. Other derivation studies by Kantor and Biely (1965) show on the basis of lead isotopic analysis that a deposit containing mainly anglesite, cerussite and galena, occurring in Triassic limestones close to the contact with andesites, gave 204 Pb 1.370%, 206 Pb 24.82%, 207 Pb 21.42% and 208 Pb 52.39%, thus indicating a pre-

Tertiary age of mineralization, and therefore not related to the andesite. In contrast, a galena from a nearby galena-sphalerite-pyrite deposit in the Moras Valley has 204 Pb 1.35%, 206 Pb 25.7%, 207 Pb 21.12% and 208 Pb 52.46%. This mineralization is probably related to Tertiary magmatism. It should be pointed out that lead isotope studies often support remobilization of material (see Chapters 1 and 44). According to Mincheva-Stefanova and Amov (1969), galena from a number of low temperature hydrothermal metasomatic lead-zinc and later lead-copper-silver ore deposits show 206 Pb/ 204 Pb ratios of 18.76, 18.62 and 18.45, the first value being indicative of younger (probably Oligocene) age, the other two of greater age but post-Palaeozoic. The authors conclude that the deposits are formed at the expense of mobilized Palaeozoic ore deposits. Collins et al. (1952), in "Age determination for some uranium deposits in the Canadian Shield" support on the basis of 52 new lead isotope age determinations from localities in the Canadian shield, that the range of ages determined from samples of pitchblende within an area is attributed to the partial dissolution and redeposition (remobilization) after the original precipitation (see also textural patterns of such remobilizations, Chapters 12 and 14). Regarding mobilizations further, Hemes (1965) suggests that a "principal feature of interest is the existence of a major province of J-type anomalous Pb-deposits along the border zone of the Scandinavian mountain chain. In this study, the J-type of Pb is interpreted as normal Pb which became progressively enriched in radiogenic Pb during passage through Precambrian rocks. Furthermore, lead isotope analysis can investigate genetic problems with particular reference to the derivation of lead itself. Cannon et al. (1961), in "The data of lead isotope geology related to problems of ore genesis", report "isotopic variations in ore leads (1280 analyses) and rock leads (199 analyses) are similar. Nearly three-quarters of the ore leads and almost twothirds of the rock leads fall into the ordinary lead category whereas, of leads from non-radioactive ore minerals (102 analyses) in radioactive deposits, less than one-quarter are ordinary lead .... with the exception of Joplin-type leads, most ore leads exhibit geologically reasonable model ages and systematic isotopic variations which indicate that they are products of some common process, whereas U- and J-leads appear to originate from a different process". (For the genesis of radioactive leads, see Augustithis, 1964). Concerning the derivation of lead, a number of contributions support that mantle derived Pb may be present or significant in various deposits. Vaasjoki (1985), in "The Teutonic Bore deposit, Western Australia: a lead study of an ore and its gossan", states "The Teutonic Bore leads plot below average 206 Pb and 207 Pb, suggesting that the lead in the ore contains a significant mantle component". Furthermore, he adds that the 193

above feature of the isotope data is consistent with the idea of a "mantle plume" origin of the Eastern Goldfields greenstone belts. Vervoort and White (1991) also support Pb-mantle derived in volcanogenic massive sulphides in "Archean mantle Pb and Nd isotopic evolution: evidence from Footwall Roacks to the Noranda volcanogenic massive sulphide deposits, Abitibi Belt, Quebec". Johannson and Rickard (1985), in "Some new lead isotope determinations from the Proterozoic sulphide ores of Central Sweden", support that the isotopic composition suggests the existence of pre-Svecokarelian crust in the district and is consistent with exhalative-sedimentary ore formation in an active continental margin environment. However, their isotopic composition of the sulphide ores precludes derivation exclusively from recycled Svecokarelian lead and suggests a substantial lead contribution from a mantle-like source. Another interesting contribution in favour of partial derivation of lead in massive sulphide deposits is presented by Fox et al. (1988), in "Genesis of basalt-hosted massive sulphide deposits from the Trondheim and Sulitjelma districts, Norway: ore lead isotopic concentrations". According to them, "when plotted on 207Pb/204Pb 206 Pb/204Pb diagrams, the data (isotopic analysis results from the deposits) define a linear trend extending from the mantle to the upper crust model growth curves of Doe and Zartman (1979)... This isotopic trend is interpreted as resulting from the mixture of lead from a mantle source with that of an upper-crustal end member."

(c) Cu, Β, Ο and C Isotopes - Consideration Special Cases of Metallogeny

of

In "Natural variations in the abundance ratio and the atomic weight of copper", Shields et al. (1965) report that "the absolute ratio 63Cu/65Cu was determined for

194

106 samples including a variety of copper minerals. No variation was found for chalcopyrite, tetrahedrite, enargite or for the Canyon Diabolo or Bruderheim meteorites; however, secondary copper oxides are enriched in the heavy isotope." Shergina and Kaminskaya (1963) carried out two investigations on variations in the n B/ 10 B ratio. Investigation No. 1 was across the mineralized zone of a high temperature Ag-Sb deposit containing stibnite, galena, sphalerite, bornite, etc., in hydrothermally altered andesites (B is present in solid solution in the plagioclase of the andesite). The concentration of n B is lower in the hydrothermally altered andesite zone than in the unaltered andesites. Investigation No. 2 was carried out along a drill-section through a Cu-Ni sulphide deposit in phyllites. The ore body zone is enriched in "B. According to Shergina and Kamniskaya, fractionation of boron isotopes up to 2.5% appears to occur during hydrothermal alteration processes, ? perhaps some leaching of n B took place. In contrast to the cases presented above, Hamilton (1968), in "Variations in carbon and oxygen isotope ratios as a possible guide to ore" suggest that C and Ο isotopes might be useful as a guide to ore, i. e., that contour maps showing variations in C and Ο isotope ratios might be useful for ore exploration. Possibly related to depositional environment, as syngenetic ore deposits have isotopic compositions of organic C which appear unaffected by metamorphism. It is also reported that oxygen isotope variation could be a guide to certain epigenetic ore deposits. The study cases quoted above were selected from a plethora of isotope investigations to serve as an indication of the enormous possibilities and prospects of isotope geochemistry application in metallogeny (see also Applications of trace elements and isotopes to environmental biogeochemistry and mineral resources evaluation, 1978; Eds.: Hurst, Davis and Augustithis).

Chapter 40

Mass-Replacement of Rocks by Ores and Palaeo-Karst-Type Deposits

In contradistinction to the textural patterns presented in Part I (Chapters 5 and 6), mass-replacement of rocks by ore is most significant and widespread and occurs under several geoenvironmental conditions. Following the inductive approach, study cases will be presented, showing replacement of different rock types by ore forming solutions or fluids. The staff of Reeves Macdonald (1961) supports that the lead-zinc ore bodies of Reeves Macdonald Mine (northwest of Nelway, British Columbia, Canada) are sulphide replacement deposits in the Reeves limestone member of the Laib formation (lower Cambrian). Furthermore, the ores lie on the footwall of a major thrust fault of regional extent. The ore minerals are pyrite, sphalerite and galena; the gangue is essentially dolomite. In contrast to the paragenesis described, Takeuchi et al. (1960) report a complex bedded-type of deposit due to metasomatic replacement of a calcareous bed. They state that minerals found in this deposit are blende, chalcopyrite, pyrite, galena, wollastonite, bustamite, garnet, diopside, rhodonite, epidote, quartz, calcite and rhodochrosite. In addition, Heath (1961) reports that the deposit at Broken Hill (Northern Zambia) consists mainly of galena and sphalerite, associated with cerussite, willemite, hemimorphite, descloizite and vanadinite which occur replacing dolomite country rock. Concerning carbonate rock replacement, Doe (1962), writing in "Distribution and composition of sulphide minerals at Balmat, New York", reports that sphalerite and locally, pyrrhotine and galena, replaced carbonate minerals in siliceous magnesian marbles of the Precambrian Grenville Series. Elaborate geochemical studies by him indicate the "the uniform isotopic composition of lead from galena in contrast to its non-uniform composition in the marble is interpreted as indicating that the lead in the ore was not derived from the surrounding marble." He also suggests that Co and Ni in the pyrite in the wallrock are significantly higher than in pyrite in the ore. Furthermore, the ratio of the concentrations of minor elements between sphaleritepyrite pairs varies widely, suggesting that exchange during the formation of the ores was slow and incomplete. A complex case of replacement is reported by Lopez et al. (1944): the copper deposits of the Aroa district, Estado Yaracuy, northern Venezuela, resulted from the

hydrothermal replacement of schistose graphitic and sericitic limestones interbedded with strongly folded calcareous, graphitic and micaceous schists. It is believed that the mineralization followed the low-grade regional metamorphism of the sediments and a period of intense local silicification and introduction of sericite, zoisite, epidote, tremolite, actinolite, diopside, chlorite, talc, calcite and dolomite. They also state that the deposition of quartz, calcite and lesser amounts of clinozoisite, zoisite, sericite and microcline accompanied the successive introduction of pyrite, galena, sphalerite, chalcopyrite and minor bornite. It is suggested that the mineralization was related to granitic intrusion, although no acid igneous rocks are exposed in the area. As Reynolds (1965) reports, Salsigne is a producing gold mine situated on a Hercynian fold belt in Southern France. The basic structure is an overturned syncline cut on the south by a thrust fault of unknown throw. The mineralization is developed on a series of normal faults that cut these major structures at approximately right angles, large scale replacement of limestone and phyllitic quartzite occurred. According to Reynolds, three phases of mineralization are apparent: (1) arsenopyrite and quartz, (2) pyrite, (3) pyrrhotite, chalcopyrite, gold, silver, bismuth minerals and quartz; the distribution of these phases is not always identical. Considering the possibility of replacement type deposits, Lewis and Narväez (1955), describing the PbZn-Ag deposits in the Cajatambo province, Department of Lima, regard the deposits as mostly being fracture fillings in the vicinity of intrusions, but they reported that some mantos have been formed through the replacement of limestones. Concerning the replacement of carbonates, an interesting contribution was made by Osman and Piestrzynski (1989), in "Mechanism of sulphide mineralization through successive metasomatic replacement stages of zoned host dolomite in Cracow-Silesian Zn-Pb deposits (Mississippi Valley type), Pomorzany Mine, Poland". As they report "replacement is controlled by two factors, i. e., microporosity and composition". The epigenetic host dolomite rhombs are zoned into three main zones; a) an outer rhombohedral rim which is hard and Fe-rich with remarkable Zn content; b) an inner zone which is porous and Zn-rich but usually low in Fe con195

tent, and c) a core of the dolomite rhombs which is Fefree but may contain some Zn. According to them, "zones rich in Zn were the first to be replaced by sphalerite; the same holds true for zones with higher microporosity showing a high degree of crystallographic disorder of cations ... The greater the zinc content in the host dolomite, the more reactive it is to replacement by sphalerite." In contrast to the typical replacement type of deposits, Leleu (1969) supports that the Pb-Zn mineralization at Laurium (Greece) includes an anomalous veintype deposit. From a study of the rhythmic morphology and mineralogy of these veins, and from thermodynamic arguments, it is concluded by Leleu that they represent pseudo-veins, formed by descending aqueous solutions identical with the solutions responsible for the karstic mineralization. The subject of replacement deposits would not be complete without special reference to skams (see Chapter 25). As a corollary to Mg metasomatism related to a basic front as the case might be with some skarns, Geijer (1963), in his contribution "On the association of magnesium and sulphide ores in metasomatic mineralization" reports the following: "Two types of sulphide mineralization are reviewed and compared: (1) the common type of replacement deposit in limestone accompanied by large-scale dolomitization and (2) the magnesia metasomatism in the Svencofennian of central Sweden and adjacent parts of Finland (Falun-Orijärvi type). The nature of the solutions active in both cases was essentially the same, involving enormous quantities of magnesium, although the general geological environment of the two types is radically different." Besides calcareous rocks, sandstones may be replaced by ore minerals. Attia (1956) reports that the Zenime district of the Sinaia Peninsula pyrolusite, psilomelane, wad, haematite and goethite occur as replacement deposits of Carboniferous sandstones and limestones. In support of the replacement of sandstones, Banas et al. (1982) report that replacement of terrigenous quartz and other minerals by sulphides was observed in the grey sandstone of Lower Permian, southwest Poland. According to them, "the process is connected with diagenetic mobilization of metals in the copper-bearing shale and their migration into the underlying porous sandstones". Furthermore, "the role of clay-organic microlayers acting as semipermeable membranes seems to be valid, controlling the conditions of the microenvironment and the extraction of metals." Concha et al. (1952) report that the Santa Barbara Mine in Huancavelica, Peru, was probably productive to a depth of about 300 m. Ore bodies in sandstone resulted largely from the replacement of the cement by the mineralizing solutions. Cinnabar, the major ore mineral, was associated with pyrite and arsenopyrite, smaller amounts of galena, sphalerite and stibnite, minor native mercury and possibly metacinnabar; ar196

senopyrite was replaced by realgar and orpiment. According to them, gangue minerals included sparse quartz, calcite, barite and hydrocarbons. In 1943, Lopez and Brineman reported that the host rocks of the San Jacinto mercury deposits, near Corora, Venezuela, are Lower Eocene sandstones, black shales and thin limestones. These suffered severe post-Miocene folding and faulting, and the ore bodies are closely related to major faultzone, forming small, irregular fillings, and replacement bodies in sandstones, often beneath impermeable shales in the vicinity of fractures. Faulting continued after the mineralization and truncated many of the ore bodies. Lopez and Brineman stated that cinnabar, the only Hg mineral, was associated with pyrite, quartz, sericite and a hydrocarbon (probably idrialite). Markham (1961) also reports that the Peelwood PbZn-Cu-Ag-Au ores, New South Wales, Australia, form massive replacement bodies parallel to the foliation of the host rocks and frequently along the contacts of metasedimentary quartz-muscovite schist and metavolcanics or smaller cross-cutting quartz-sulphide veins. According to Markham, the ores are very finegrained and banded; sphalerite and pyrite are associated with galena, arsenopyrite, chalcopyrite, tetrahedrite (an antimonian end member) and minor pyrrhotite, marcasite, gold and bornite. Furthermore, Watson (1959) supported that the banded sulphide deposits of Minadamar Mine, Stirling, southeastern Cape Breton Island, probably originated by replacement of relic lamination in the schist. The host rock is a laminated siltstone of the Bourinot Group (Middel Cambrian) transformed by low-grade metamorphism into siliceous sericite schist. The metallic minerals are pyrite, blende, chalcopyrite, galena and tennatite. The banded structure in which layers rich in sulphides alternate with those rich in gangue minerals (as mentioned), probably originated by replacement of relic lamination in the schist. It is suggested that sulphide replacement started in the coarser, quartz-rich portions of the laminae. Watson maintains that chalcopyrite was among the last of the ore minerals to be deposited and where chalcopyrite-rich layers alternate with pyrite and blende-rich layers, it is presumed that chalcopyrite may have replaced the highlyschistose sericite-rich layers more readily than it did the sulphide-rich layers. Another example put forward by Fleming (1961) described the Murray deposit, Restingouche County, New Brunswick, Canada, as a replacement body in chlorite schist along the axis of a reverse-type drag fold. The ore minerals consist of pyrite, sphalerite, galena and chalcopyrite. In contrast to the replacement types presented so far, Sorem and Cameron (1960), in "Manganese oxides and associated minerals of Nsuta manganese deposits, Ghana, West Africa" support that the principal ore minerals are 'Nsuta Mn02\ cryptomelane, pyrolusite and lithiophorite; goethite is present in some ores, and

muscovite, quartz, garnet, zircon and amphibole are also present. According to X-ray data, 'Nsuta Mn02' is unlike pyrolusite, ramsdellite or γ-Μη0 2 ; one type is similar to p-Mn0 2 . Oxide ore was formed by replacement of rocks rich in quartz, muscovite and garnet by filling of fractures and other openings, and by leaching of impurities from low-grade ore. Besides quartz being replaced by manganese-rich ores, a complex case of quartz replacement is reported by Stemprock (1960). According to him, "the Sn-W mineralization in the Cinovec district (Czechoslovakia) is associated with a system of flat veins developed inside a granite dome and proceeding for a short distance into the neighbouring quartz porphyry. The ore forming process was preceded by albitization of the granite, flat fissures were formed and filled with quartz, and subsequently greissens were formed and extensive intravenous metamorphism took place during which quartz veins were replaced by zinnwaldite, topaz, cassiterite and wolframite. Later a potassium feldspar metasomatism occurred and finally sulphides were deposited". Not only sedimentary rocks suffer replacement, but also volcanics are subject to replacement processes. Geological and mineralogical studies of two sulphide ore bodies at Kokkinopezula (Cyprus) by Kattamis (1961) show that they are mesothermal substitutions of lava. Kattamis maintains that geothermal and X-ray studies indicate formations at low temperature and pressure. The deposits are characterized by the associations pyrite-chalcopyrite and pyrite-chalcopyritesphalerite. Geoffroy and Koulomzine (1960) described the sulphide deposit in Barraute area, northern Quebec, as a complex replacement of silicified volcanic rocks by sulphides with gold and silver. Despite extensive replacement of the volcanics and sediments of the Rio Tin to, Williams (1962), in "Further reflections on the origin of the porphyries and the ores of Rio Tin to, Spain" maintains that on the basis of the consistent association of the pyrite bodies with the pyroclastic deposits that the mineralization may be exhalative-sedimentary in origin; however, extensive replacement of the volcanics and sediment is also accepted by him. In contrast to William's interpretation of Rio Tinto (exhalative sedimentary), Kinkel (1966), in "Massive pyritic deposits related to volcanism, and possible methods of emplacements" supports the "observations on modern and recent massive pyrite and pyrite-sulphur deposits associated with volcanic centres in Japan and Taiwan, that they form by replacement of lavas and pyroclastics affected by volcanic emanations at or near surface, on land, or in crater lakes". Kinkel's view is that older, underformed massive pyrite deposits in Cyprus and deposits within deformed rocks, e. g., at Rio Tinto, Spain, and in Shasta County, California, are thought to have been formed by similar processes operating under submarine conditions. According to

Kinkel, such deposits are commonly peneconcordant and stratabound but not syngenetic in origin. From the plethora of literature available on mass-replacement by ore, Hall (1959), in "Geochemical study of Pb-Ag-Zn ore from the Darwin Mine, Inyo County, California" reports that ore bodies occur as pipe-like replacements in rocks consisting mainly of wollastonite, grossular-andradite garnet, idiocrase and diopside. The ore consists of galena, sphalerite, pyrite, pyrrhotite and chalcopyrite with small amounts of andorite, matildite, scheelite, tetrahedrite and a bismuth-seleniumbearing galena with high-silver lead. Considering the en masse replacement of rocks by ore, the processes could involve dissolution-assimilation of elements comprising the rock and precipitation of the ores either in voids or fronts of dissolutions might be immediately followed by deposition-precipitation. The microscopic scale of a comparable process is extensively treated in the case of replacement textures (see Chapter 5). As already mentioned, Leleu (1969) explained the mineralization at Laurium 13 , Attika, Greece, as karstmineralization in contrast to previous interpretations which considered the metallogeny as due to replacement. Also as a corollary to void fillings, karst mineralization attains a particular significance in reconsidering many calcareous replacement type deposits, in accordance to the hypothesis of karst-type mineralization. Maria Boni and Amstutz (1982), in "The PermoTriassic Palaeokarst ores of southwest Sardinia (Iglesiente Sulcis): an attempt at a reconstruction of Palaeokarst conditions", emphasized that "karst-type ore deposits are a major new discovery in many countries made by many mining geologists during the past decade". Furthermore, they maintain that karstification took place at various intervals, producing superimposed karst systems and local sulphide-barite concentrations. In conjunction with the hypothesis of karst-mineralization, Lacerda and Bernard (1984), in "Existence de min6ralisations plumbo-zinciferes syngdnetiques du substratum Cambrien du District des Malines (Gard, France)", maintain that "most of the ores extracted from Les Malines mining district come from Pb-Zn deposits apparently formed by the infilling of paleokarst cavities". They also support that these cavities derive from Cambrian dolomites solution by meteoric water circulations below the unconformity separating the Triassic and/or Mesozoic strata from the Hercynian basement. A very sound criticism concerning the karst-mineralization hypothesis versus thermal mineralization is presented by Schulz (1982), in "Karst or thermal mineralization interpreted in the light of sedimentary ore 13

The mineral paragenesis of Laurium is most adequately treated in the paper "New observations on ore formation and weathering of the Kamariza deposit, Laurion, SE Attica (Greece)", by Meixner and Paar (1982). 197

fabrics". According to him, "in the wake of recently obtained data on Zn-Pb-Fe-Ba-F-enrichment or erosion surfaces and paleo-karsts in carbonate rocks, there is a

198

tendency towards unjustified generalization of ideas beyond immediate straightforward interpretations".

Chapter 41

Hypogene, Supergene and Oxidation Mineralizations

(a) General The terms hypogene and supergene (hypergene) mineralization have been widely used, the first referring to "primary" mineralization at some depth, often due to ascending solutions from a source, and the second (supergene) to near surface or upper part due mainly to descending solutions caused by alteration or leaching of hypogene or other primary mineralizations. Due to the fact that the hypogene mineralization is mainly the result of ascending solutions - the supply may be in stages or intermittent; zonation and stages of mineralization are often easily recognized. As a result of multistage mineralization, extensive replacement phenomena are common in the case of hypogene mineralization. In contradistinction, supergene mineralization is due to solutions and mobilizations which are the result of alteration of primary mineralization either hypogene or metamorphogenic and the descending solutions usually result in the enrichment of the effected hypogene mineralization by solutions (elements) supplied by the supergene mineralization. Often the formation of the supergene mineralization is entirely at the expense of hypogene mineralization or parts of it. Also common are those cases where the surface "zone" is relatively depleted in some metallic elements and where enrichment can take place within the supergene mineralization itself. Since mixed hypogene and supergene mobilizations can occur and since the same mineral species can be either of hypogene or supergene derivation strict criteria and characteristic mineral species recognition is difficult to define or to state with any certainty. As mentioned, Edwards (1960, see Chapters 2 and 5), emphasized the importance of replacement processes in the hypogene mineralization itself which may be due to multistage supply of solutions. However, it should be stressed that most impressive replacement patterns are shown when supergene solution affects hypogene mineralization. In fact, supergene mineralization exhibits replacement textural patterns in most cases. Since hypogene, supergene and oxidation mineralization are of great significance, it is not surprising that the literature available on these topics is enormous. Therefore, study cases will be presented under the sub-

headings: hypogene, hypogene-supergene and supergene mineralizations.

(b) Hypogene

Mineralization

As mentioned, stages and zonation are often recognizable in hypogene mineralization. Markham and Lawrence (1962), in "Primary ore minerals of the Consols Lode, Broken Hill, New South Wales" describe the transgressive silver-bearing vein of the Australian Broken Hill Consols Mine, which lies about 600 yards to the east of the main conformable Broken Hill lode system. Argentian tetrahedrite, dyscrasite and acanthite were the major ore minerals. Furthermore, according to Markham and Lawrence, the general order of hypogene deposition was arsenides-sulphides, native elements, and intermetallic compounds silver sulphates; the paragenesis is comparable to the St. Andreasberg, Harz, and the later AgSb-rich stage in the main Broken Hill lode. The recognition of an order of deposition is compatible with a sequence of supply or even stages of supply. However, a more distinct case of hypogene zoning is described by Buryak (1967), in "Hypogene zoning in the gold provinces of Siberia". Zoning in the ore bodies consists of gold-sulphide veins in deeper horizons compared to gold-quartz veins at higher levels. In contradistinction to zoning, Apel'tsin and Savel'yev (1960) describe leucocratic dykes changing from aplitic quartz porphyries near the main massif to greisen rock and to veins essentially of tourmaline and quartz and of quartz away from the massif. According to them, the ore minerals refer to two hypogene phases; those of the first include wolframite, gold, bismuthinite, scheelite, löllingite and rammelsbergite, and those of the second bismuthinite, tetradymite, bismuth and microscopic gold. Furthermore, Gorovoy (1968), in "Paragenesis of hypogene minerals in a mercury deposit in Donets basin" supports that the process of mineralization took place in five stages: metacrystals of arsenopyrite, pyrite and quartz; pyrite and quartz; carbonates, quartz, stibnite and cinnabar; dickite, pyrite and marcasite. The presence of a single mercury-antimony stage, several stages with pyrite, and the appearance of hypo199

gene marcasite are characteristic for the mineralization in this deposit. Bindeman (1963), in "Types of hypogene mineralization in Shilka-Arbayar region and the Novoberezovskaya depression (eastern Transbaikalia)" describes complex gold-tourmaline and gold-antimony associations in the hypogene ores of the ShilkaArbagar districts and the Novoberezovskaya basin, eastern Transbaikalia. Hypogene environmental controls are considered by Moore et al. (1966) for the largely coextensive areas of locally high Bi-Cu-Mo concentrations which define a central zone; Pb-Zn areas of greater lateral extent partly overlap the central zone; locally high As-Sb concentrations characterize an outer zone with Β extending beyond the limits of known significant mineralization. Ag is erratically distributed. In addition to the silver in veins of hypogene manganese oxides, Hewett (1968) suggests that Ag-bearing sulphides may exist below barite-manganese deposits in Nova Scotia, again pointing to hypogene "mineralization differentiation".

(c) Hypogene-Supergene Mineralization (Direct Relationship and Derivation of Supergene from Hypogene Mineralization) The coexistence of hypogene and supergene is not only common but also indicates their relationship and the derivation of supergene mineralization from hypogene or other primary mineralizations. Study cases presenting the relationship of hypogene and supergene mineralization are presented as follows, selected from the vast plethora of available literature. Beall (1962), in "Tsumeb enters a new stage of development" reports that the ore body in South West Africa (Namibia) is a breccia pipe with peripheral veins of massive ore accompanied by disseminated mineralization in the core. The hypogene minerals include galena, sphalerite, tennantite, chalcopyrite, bornite, pyrite, germanite, renierite, enargite, neodigenite and chalcopyrite. Furthermore, Beall recognized as important supergene minerals: native copper, silver, cuprite, azurite, malachite, cerussite, smithsonite, anglesite, duftite, chrysocolla and olivenite. Hall and MacKevett (1958) in the Darwin quadrangle, Inyo County, California, also recognized as main hypogene minerals in the lead-silver-zinc deposits argentiferous galena, sphalerite, pyrite, phyrrotite, chalcopyrite, tetrahedrite, scheelite, andorite, franckeite, stannite, enargite-famatinite (?), matildite, bismuth (?) and a probable lead-bismuth-selemium sulphosalt. According to them, supergene minerals include limonite, hemimorphite, cerussite, anglesite, hydrozincite, plumbojarosite, pyromorphite, smithsonite and locally also aurichalcite, azurite, brochantite, caledonite, chrysocolla, linarite and malachite. 200

Considering certain aspects of the geochemistry of hypogene and supergene mineralizations, of significance are studies by Gavelin et al. (1960) on sulphur isotope fractionation in sulphide mineralization. According to them, "although hypogene sulphides deposited at lower temperatures exhibit greater variation in isotopic composition than those in the same deposit formed at higher temperatures, isotope analyses on sulphides and sulphates from many localities indicate that no correlation is established between temperature stages in hypogene mineralization and isotopic composition even where metal zoning can be demonstrated". Furthermore, significant divergencies in isotope ratios in coexisting sulphides and sulphates suggest that some fractionation may take place during hypogene mineralization. No fractionation is apparent in supergene oxidation of sulphides to sulphate. Vincienne and Mozaffari (1966) report "that hypogene mineralization in chlorite schists and dolomites has produced arsenopyrite, löllingite, siderite, pyrrhotite, pyrite, wolframite, scheelite, rutile, sphalerite, chalcopyrite, quartz, galena, cosalite, bismuthinite, polybasite and native gold." Furthermore, they maintain that supergene minerals include calcite, covellite, scorodite, arsenopyrite, chalcopyrite and pyrrhotite. It should be noted though that arsenopyrite, chalcopyrite and pyrrhotite are reported both as hypogene and supergene in the Salsigne, Aude, France. A special case of primary mineralization followed by supergene is reported by Marinelli (1959). A bismuth-rich paragenesis is reported in the Falcacci zone of the Rio Marina deposit, eastern Elba. The deposition of pyrite, antimony, bismuthinite, sphalerite and galena has taken place at relatively high temperatures, followed by supergene formation of bismoclite, bismuthinite, cerussite, anglesite and smithsonite. Another case of supergene mineralization as derivative of primary mineralization is reported by Shnaider and Shnaider (1966), in "The occurrence of realgar in sulphide ores of Novo-Zolofushinskii and Kamyshinskii (Rudney Altai) deposits, USSR". According to them, the realgar in these deposits occurs with arsenopyrite, tennantite and rare löllingite, enargite and proustite and is one of the latest minerals. It was probably produced by the supergene leaching and redeposition of As from the As containing minerals. The realgar was found in fractures cutting pyrite-chalcopyrite and polymetallic ores. The study cases presented above support derivation of supergene mineralization from hypogene or other primary mineralization by dissolution, leaching and reprecipitation.

(d) Special Cases of Supergene

Mineralization

As already pointed out, supergene mineralization is often a derivative of hypogene mineralization due to dissolution, leaching and, in general, element mobili-

zation processes. Within this general framework, an attempt will be made to present some study cases further elucidating the processes. Bezrodnykh (1967), in "Conditions of formation of associations of some copper and silver minerals in the supergene zone of the Udokanst (USSR) copper deposit" states the following: "Calculated Eh-pH and P 0 2 - P C 0 2 diagrams based on thermochemical data from the literature indicate that native silver should be associated with native copper, cuprite, tenorite, malachite, azurite and chalcocite in the supergene zone of most ore deposits". Bezrodnykh suggests that Ag is separated from Cu in alkaline media but not in acid, and silver carbonate occurs only rarely in highly oxidizing and C0 2 -rich environments. In addition, studies on idaite, Cu5FeS6, by Krause (1965) show that specimens from old mine dumps from Konnerud, Norway, are shown under polished section to contain idaite as fine lamellae within bornite. According to Krause, the idaite is interpreted as a very early supergene alteration product of bornite. Considering the mineralogical and geochemical changes in the supergene zone, Chitayeva (1965) states that in the oxidized parts of the Fe-Cu sulphide ore bodies of the southern Urals, Se and Te are concentrated in supergene sulphides, quartz-native sulphur "sand" and (Se only) in gossan limonites, but depleted in the sulphate zones relative to the unoxidized ores. In cases, the Se and Te form haloes around some deposits in sulphate and limonite impregnated wallrocks, in other instances they are apparently removed. Additional studies by Kolkovsky (1966) support that two types of vertical zoning are discerned corresponding to the degree of oxidation and leaching (type I) and to enrichment in the principal supergene minerals (type II). According to Kolkovsky, the first type has an upper sub-zone of intense oxidation and leaching coinciding with the zone of aeration, and a lower sub-zone of limited oxidation. The second type has five levels: hydrogoethite-beudantite-plumboj arosite, hydrogoethite-pyromorphite, anglesite-covelline, hydrogoethite-cerussite, and limonite-hydrogoethite. Supergene gold occurs only in the last level. The work of Kolkovsky shows the intimate relationship between supergene minerals and the oxidation zone (discussed further on p. 202). Another case of geochemical-mineralogical studies concerning the formation of supergene minerals related to the oxidation zone is reported by Warden (1970). He maintains that at the Forari Mn deposit, Efate Island, New Hybrides, the ore occurs as stratiform and lensoid concentrations of Mn0 2 with todorokite as the commonest recognizable mineral, in marine volcanic sediments and in contact with overlying soil and limestone. The distribution of the principal oxides (Mn, Fe, Si, Al) and some trace elements (Ba, Sr, Mo, P, Co, Ni) in several profiles show a distinctive covariance pattern. Fe is concentrated at or near the surface and Mn at a lower level; Mo and Ba are at a maximum in the ore

horizon where there is a decrease in Si0 2 and A1 2 0 3 . According to Warden, these features suggest supergene enrichment in deep-weathering topical soil due to precipitation of Mn and other metals from groundwater circulating in the zone of oxidation. The Mn was probably leached from submarine pyroclastic sediments of andesitic composition and transported in solution or partly in suspension. Further studies on oxide-type supergene mineralization by Edel'shteyn (1965) support that in some deposits the nickel is concentrated at the bottom of the nontronite and nontronized serpentine zones and in others it is the ochre developed through decomposition of nontronite and other silicates. He also states that the oxidation of sulphides during supergene alteration of nontronite residium may produce a mobile N i S O ^ H j O type compound of Ni which migrates to lower horizons. The study cases quoted above clearly show the significance of supergene mineralization and the relationship between supergene mineralization and oxidation. In order to understand the weathering of ore minerals in addition to leaching and supergene mineralization, it is necessary to consider study cases of the oxidation zone, in spite of those already quoted, where supergene mineralization and oxidation have been treated together.

(e) Oxidation Zone

Mineralization

Oxidation is a process that can extend to considerable depth. Taylor (1958) reports that oxidation persists to a depth of at least 1150 ft below surface in the case of Broken Hill lead-zinc ore bodies in Zambia. The oxidation process often results in the weathering of minerals and in the formation of more stable alterationoxidation secondarily formed minerals (see also Chapter 15). Extensive studies have been carried out on sulphide ores and from the available literature characteristic cases will be presented. Timofeeva (1965), in "Some minerals in the oxidation zone of the Mosrif deposit" reports that primary minerals in the Mosrif deposit (USSR) are magnetite, löllingite, arsenopyrite, chalcopyrite, pyrite and occasional native bismuth, bismuthinite, tetradymite and wittichenite. According to Timofeeva, the oxidation zone is thin but very rich mineralogically, including goethite, scorodite, chalcanthite, azurite, malachite, halotrichite, copiapite, pickeringite, arsenosiderite, and in addition very rare Cu, Zn and Bi arsenates (olivenite, adamite, atelestite, chalcophyllite, tyrolite and mixite). Timofeeva further reports that the formation of this large group of rare arsenates was by direct precipitation from acid solutions oversaturated with Cu, Zn and Bi (olivenite, chalcanthite, tyrolite and mixite) and by replacement of primary ores and newly formed sulphates and arsenates in weakly acid and neutral me201

dium (adamite, atelestite, chalcophyllite). The formation of such a variety of "secondary" minerals in the oxidation zone of Mosrif deposits indicates extensive element remobilization and geochemically complex geoenvironmental conditions. Similarly, Morävek (1958), considering the mineralogy of the complex gold-bearing ore veins in Jilovd, central Bohemia, reports that the veins contain gold, bismuth, pyrite, phyrrhotine, chalcopyrite, galena, sphalerite, arsenopyrite, bismuthinite, covelline, cosalite, tetrahedrite, tetradymite, tellurobismuthite, coloradoite, meneghinite, scheelite. Furthermore, the secondary minerals in the oxidation zone are limonite, haematite, gypsum, malachite, azurite and yellowish bismuth ochres. It should be pointed out that in the case of the Jilov6 deposit, both the gold-bearing veins paragenesis and minerals formed in the oxidation zone indicate a complex geochemical system comprised of the vein minerals and of oxidation zone secondary mineral formation. An additional case of oxidation zone mineralization is described by Deb (1968), in "Mineralogical study of the gossans and oxidized ores from Roam-Rakha area, Bihar, India". The gossans are essentially composed of goethite and hydrohaematite with minor magnetite and martite. The oxidized ore minerals are iron hydroxides, chalcocite, covellite, native copper, tenorite, malachite, marcasite, pyrite and violarite. It should be noted that the oxidation zone mineralization is typical of copper sulphides "primary" mineralization. Gaspar (1967), considering the main mineralization of the Cerro do Algard deposit in Beja, states that the main minerals are pyrite, arsenopyrite and löllingite; chalcopyrite, enargite, stannite and galena are also present. The mine operates mainly in the oxidation and cementation zone where the secondary minerals limonite, cuprite, covellite, chalcocite, malachite, azurite, scorodite and bornite occur. It should be noted that bornite is a mineral present in the oxidation zone. Another case of oxidation zone mineralization as derivative of primary mineralization is reported by Eliseev et al. (1961), in "Ultrabasic and basic intrusion of Pechenga' Petsamo". They maintain that the principal primary minerals are pentlandite, pyrrhotite, chalcopyrite, pyrite, ilmenite and magnetite. Furthermore, they state that the zone of oxidation includes covellite, violarite, malachite, azurite and, among others, ironmagnesium retgersite. Boyle (1960), discussing the occurrence and geochemistry of native silver in the lead-zinc-silver lodes of the Keno Hill/Galena Hill area, Yukon, Canada, reports that native silver is of secondary origin, derived by the oxidation of freibergite and argentiferous galena; such deeply oxidized lodes are uncommon in Canada. According to him, the lenses of ice in the permafrost zone which extend to depths of 75 to 200 ft. are unusual. Splendid leaf silver occurs mainly in clear ice lenses in the oxidized zones. In the oxidized parts of the lodes anglesites, bindheimite, scorodite, melan202

terite, acanthite and Ag-bearing beudantite and plumbojarosite are also included. Most complex element mobilization processes must have taken place for the formation of the oxidation zone secondary minerals from the reported primary paragenesis and their geoenvironment. Threadgold (1958), considering the antimony-gold mineralization at Steel's Creek, near Yarra Glen, Victoria, Australia, reports that narrow quartz veins altered quartz-feldspar porphyry contain gold, stibnite, berthierite and other sulphides. Furthermore, the oxidation of the antimony ores yielded stibiconite, valentinite, senarmontite and tripuhyite. Comparing the primary mineralization to the oxidation zone mineralization, complex element processes must have taken place. In addition to the comparison of primary and oxidation mineralization, Kulikova (1965) supports that the effect of enclosing rocks on the behaviour of rare elements in the oxidized zones of the lead-zinc deposits of eastern Transbaikalia are of significance. Kulikova states that concentrations of trace elements (Tl, In, Cd, Ga) remain high in the oxidized zones of sulphide ore bodies where the host rocks are carbonates but ar r lower where the host rocks are aluminosilicates. This is in support of the significance of the geoenvironmental factors for element mobilization in the oxidation zone. Concerning minerals in the oxidation zone or environment, Nickel (1977), in "Mineralogy of the 'Green leader' gold ore at Kaligoorlie, Western Australia" supports that gold occurs as native gold and as calaverite and petzite; the native gold has been derived, at least in part, by the oxidation of calaverite. Machairas (1967) also supports that "the distribution of Ag in exogenous gold which crystallized in the oxidation zones as nuggets, shows that the recrystallization of the gold occurred in several stages". Furthermore, the solution and recrystallization of gold are related to the oxidation of sulphides. In their study on the oxidation of sulphides, Szolnoki and Bongär (1964) report that the oxidation of pyrite, chalcopyrite and sphalerite is more rapid for samples inoculated with bacteria (Thiobacillus ferrooxidans) than for sterile samples, but according to them, it is still a slow process. Also, Ivanov (1962) supports that Thiobacillus ferrooxidans is effective in the oxidation of pyrite and covellite to form sulphate solutions whereas Thiobacillus ferrooxidans is merely effective in the oxidation of sulphur. Ivanov also maintains that the combination of the two organisms produces faster oxidation rates than Thiobacillus ferrooxidans alone. Borchert (1960), in "Genesis of marine sedimentary iron ores" also points out that trivalent iron is practically insoluble in the presence of oxygen and that the mobilization and precipitation of iron must have been within the oceans. The existence of a C0 2 zone at moderate depths in the ocean, when Fe is dissolved to be later precipitated in oxygenated shallow waters, is emphasized.

On the subject of iron ores, Mukherhjee (1966), in "Minerography of the iron ores around Khetri, Jhunjhunu district, Rajasthan" reports that the ore contains magnetite, maghemite, martite, goethite and lepidocrocite which show replacement in the same order, with magnetite being replaced by the above minerals during different stages of oxidation. Il'vitskiy and Romanesko (1964) in their study on ore mineralization at granite-serpentine contacts recorded metamorphic ultrabasic rocks near the Sur magnetic anomalies of the Dnieper region containing magnetite-nickeliferous pyrite-millerite concentrations resulting from the effects of hydrothermal solutions of young granitoids whose dykes cut the serpentinites. According to them, the oxidation of the magnetite, sulphides and serpentinites produced secondary minerals such as martite, goethite-hydrogoethite, jefferisite

and ferri-halloysite which contain up to 6.06% NiO and 0.23% CoO. The example quoted supports that element mobilization resulted in the formation of ferrihalloysite containing Ni and Co which are most probably derivatives of the nickeliferous pyrite or other sulphides. In cases, an element of great geochemical mobility, such as U+6, may undergo repeated mobilization and finally form exogenic uranium deposits only under warm acid conditions. Krasnikov and Sharkov (1962) support such mobilizations of uranium in their contribution "Spatial and genetic relations between the exogenic and metamorphic uranium deposits and the acid zones of the geological past". It should also be noted that U+6 is more mobile than U+4 (see Augustithis, 1964, 1974).

203

Chapter 42

Some Aspects of Manganese Mineral Formation - Transformation - Alteration Oxidation and in General MnMobilization/Remobilization

When considering manganese minerals, it should be kept in mind that Mn is a contributing element in the composition of many ore minerals and as already pointed out, belongs to the mantle derived element group: Cr, Μη, Fe, Co, Ni, Cu, Zn, Cd, Ag, Au, Hg (see Chapter 1 and Part III). Aspects of the role of manganese in metallogeny should be emphasized when discussing the empirical "laws" of element segregation. Manganese mineral formation will also be discussed in respect to the superimposed paragenesis concept (see Chapter 51) and the mobilization of Mn by leaching processes from intercontinental-rift basalts has been mentioned. Since Mn and manganese minerals occur in a wide spectrum of rock and ore types, some are discussed in the pertinent Chapters (e. g., manganese nodule formation, Chapter 13), an approach will be made to consider the 'plexus of processes' involving hypogene hydrothermal solutions, supergene mobilization, alteration-oxidation, synsedimentary formation of manganese ores as well as aspects of recrystallization and metamorphism. The term 'plexus of processes' is preferred over 'cycle of processes' since in the same deposit metamorphic, colloform and alteration manganese minerals may coexist. In other cases, hydrothermal replacement minerals may be associated with alteration-oxidation minerals. Study cases indicating some of the complexities of Mn mineral parageneses will be presented in support of the concept of 'plexus of processes'. Roy (1961), in "Mineralogy, textures and paragenesis of the manganese ores of Gumgaon-Ramdongri mine area, Nagapur district, Maharashtra, India" states that the paragenesis of manganese ore minerals is classified broadly into metamorphic and colloidal types. Braunite-free 3(Mn,Fe) 2 0 3 MnSi0 3 , hausmannite-free Mn 3 0 4 , jacobsite (Mn,Fe) 3 0 4 , vredenburgite (Mn +2 ,Fe +2 )(Mn +3 ,Fe +3 ) 2 0 4 - hematite (Fe 2 0 3 ) assemblages are regarded as of metamorphic origin, while the cryptomelane (XR 2 Mn 6 4 0 16 , where X = Kj, Ba, Pb; R = Mn, Fe, Cu, Zn) - manganite (MnO(OH)) - pyrolusite (Mn0 2 ) assemblages are considered to have been formed by colloidal influx on the metamorphic ores. In contradistinction, Roy (1959), in "Mineralogy and texture of manganese ore bodies of Dongari, Buzburg, 204

Bhandara district, Bombay State, India, with a note to their genesis" supports that the manganese ore have two principal modes of occurrence (1) massive ore bodies associated with gondites containing braunite (3(Mn,Fe) 2 0 3 MnSi0 3 ), manganite (MnO(OH)), jacobsite ((Mn,Fe) 3 0 4 ) and vredenburgite ((Mn +2 ,Fe +2 )(Mn +3 ,Fe +3 ) 2 0 4 ) with minor secondary pyrolusite (Mn0 2 ) and cryptomelane (XR 2 Mn 6 4 0 16 , where X = Kj, Ba, Pb; R = Mn, Fe, Cu, Zn) and (2) cavernous, botryoidal, kidney-shaped, bounded, principally manganese dioxide ores containing cryptomelane, coronadite (MnPbMn 6 4 0 14 ), pyrolusite and manganite. Furthermore, in the mixed ore of the two groups jacobsite is enveloped in minerals of colloidal origin (i. e., pyrolusite and cryptomelane) with little or no replacement. Roy also supports that the paragenetic sequence indicates that the massive gonditic ore was formed initially by metamorphism of early manganiferous sedimentary rocks. Elevation of temperature resulted in recrystallization and growth of coarse braunite and jacobsite and the destruction of previous pyrolusite and manganite. The secondary (cavernous) ore type is much later and is the result of the influx of colloidal gels into porous and folded mica-schists. The above cases rather support that the colloform (secondary ores) were derived most probably by the remobilization due to solutions of the metamorphogenic ores; thus, they support the sequence: sedimentary manganese metamorphosed —» mobilized by solutions -»colloform ores. An attempt to distinguish between pyrolusite (Mn0 2 ), ex-manganite (MnO(OH)) from primary polianite (Mn0 2 ) (pyrolusite) by crystallographic and optical studies was carried out by Gaudefroy (1960). Overall, it is difficult to specify which 'manganese mineral' is "primary" sensu stricto, since, as already pointed out, it is a 'plexus of processes'. Transitions of hydrothermal manganese oxides to rhodochrosite (MnC0 3 ), rhodonite (CaMn 4 (Si 5 0 15 )) and albandite (MnS) are reported by Hewitt (1963), supporting in these cases the derivation of oxides and hydroxides of manganese to vein hypogene manganese minerals. Another case indicating a "derivation" of vein manganese minerals partly from metamorphics is reported

by Perseil (1968). She maintains that "in the manganese deposits situated between Peyresounde and Vieille Aure, Hautes Pyr6n6es, the Mn is associated in part with a microquartzite horizon and in part with quartz and carbonate veins. The manganese oxides in the microquartzite have a microrhythmic structure and include ramsdellite (Mn0 2 ), groutite (HMn0 2 ), birnessite (Na 4 Mn 1 4 0 2 7 -9H 2 0), manganite (MnO(OH)) and todorokite [(Mn +2 CaBa) 2 Mn +3 ,Mn u +4 0 3 3-8H 2 0]; in the veins they include polianite, cryptomelane, albandite (MnS) and hiibnerite (MnW0 4 ). In this particular case, it is possible that bimessite, manganite and todorokite reported in association with the microquartzite are alteration derivatives themselves. Kabesh (1961), considering the "Manganese ore deposits of the Sudan" supports that the main ore minerals as pyrolusite (MnO), manganite, psilomelane (BaMn 2 Mn 4 0 1 6 (0H) 4 ), braunite and wad. Primary Mn minerals are reported as manganiferous garnet, rhodonite (CaMn 4 (Si 5 0 15 ) and rhodochrosite (MnC0 3 ) and perhaps the reported main manganese mineral braunite is, in a sense, also primary. The recognition of primary and secondary Mn minerals should perhaps be considered in any case as arbitrary since hydrothermal manganese minerals might be derivatives of earlier sedimentogenics or metamorphics. Considering further the plexus of manganese mineral transformations, alterations, oxidations and, in general, mobility, Crittenden et al. (1961) report that manganese carbonate ores were formed by hypogene hydrothermal replacement of impure dolomite, where preore faults that served as channels cut the carbonate rocks. According to them, the fissures locally have veins in which rhodochrosite is associated with base metal sulphides. The hypogene veins include the following Mn minerals: pyrolusite, 'wad' psilomelane (BaMn 2 Mn 4 0 1 6 (0H) 4 ), rhodochrosite manganoan calcite, and manganoan dolomite. Pyrolusite, 'wad' and the psilomelane are probably derivatives of previous manganese minerals. The ore deposits near the surface are oxidized and enriched by the action of surface water. Brooks (1962), in his study of the Mary Valley manganese deposits of Queensland, reports that the manganese deposits are confined to shale, quartzite, andesite and jasper and that the country rocks are slightly metamorphosed but highly deformed. Important manganese cryptomelane (XR 2 Mn 6 4 0 16 , where X = K^, Ba, Pb; R = Mn, Fe, Cu, Zn), braunite (3(Mn,Fe) 2 0 3 MnSi0 3 ) and pyrolusite (Mn0 2 ), white hausmannite (Μη 3 0 4 ), γ-Μη0 2 , rhodonite (CaMn 4 (Si 5 0 15 ) and piemontite (manganoan epidote) were also identified. Considering the mentioned manganese minerals reported by Brooks, braunite, hausmannite, rhodonite and piemontite are most likely metamorphogenic, while cryptomelane, pyrolusite and γ-Μη0 2 are most probably derivatives of metamorphogenic manganese minerals. Furthermore, supergene enrichment of manganese oxides was controlled by

fractures, faults and lenses, commonly by fold structures. The plexus of manganese mineral formation/remobilization can be outlined as sedimentary manganese —> transformed to metamorphogenic manganese assemblage —> by 'hydrothermal' solution to derivative manganese minerals and finally supergene oxides enrichment has taken place. Okada and Kitamura (1965) report that secondary manganese ores are formed after rhodochrosite (MnC0 3 ) from Inakuraishi and Ohe mines, Hokkaido, Japan. According to them, the secondary minerals formed are: manganite (MnO(OH)), groutite (HMn0 2 ), pyrolusite (Mn0 2 ), yokosukaite (γ-Μη0 2 ), bimessite (Na 4 Mn 14 0 27 -9H 2 0), cryptomelane (XR 2 Mn 6 4 0 1 6 , where X = Kj, Ba, Pb; R = Mn, Fe, Cu, Zn), todorokite [(Mn +2 CaBa) 2 Mn +3 ,Mn 11 +4 0 33 -8H 2 0], lithiophorite ((Li, Al)Mn0 2 (0H) 2 ) and hetaerolite (ZnMn 2 0 4 ). Considering the composition of rhodochrosite and of the secondary manganese minerals reported by Okada and Kitamura, it is obvious that in addition to Mn many elements were mobilized/remobilized for their formation. Nambu and Tanida (1961), in "Progressive alteration of manganese dioxide observed at the Toyoguchi mine, Iwate Prefecture, Japan" report that the Toyoguchi deposit is one of the banded manganese deposits in Palaeozoic black banded chert. Namdu and Tanida conclude that rhodochrosite ore has been altered to pyrolusite through Mn0 2 gel, bimessite and ramsdellite (Mn0 2 ), and then to cryptomelane by potassium absorption. However, also in this case, a wide range of elements in addition to Mn have been involved in the alteration of the rhodochrosite to the recorded secondary manganese minerals. Concerning the alteration of primary manganese minerals, Prinz (1961) reports that sphalerite, silver and galena accompany rhodochrosite in primary replacement deposits. The oxidation of rhodochrosite has produced: todorokite ((Mn + 2 CaBa) 2 Mn + 3 ,Mn n + 4 0 3 3 -8H 2 0), cryptomelane (XR 2 Mn 6 4 0 16 , where X = IC,, Ba, Pb; R = Mn, Fe, Cu, Zn), γ-Μη0 2 , pyrolusite (Mn0 2 ), chalcophanite (ΖηΜη 3 0 7 ·3Η 2 0), hetaerolite (Zn,Mn 2 0 4 ) and rare manganite (MnO(OH)). In this case rhodochrosite may also have been the main supplier of Mn for the formation of the reported oxidation manganese minerals, however, many other elements have been mobilized or remobilized (i. e., Zn, Pb, i y for the formation of the reported oxidation zone manganese minerals. It should be noted that according to Prinz, the depth of the oxidation ranges from 100 to more than 850 ft. Furthermore, Roy (1960), in "Mineralogy and texture of the manganese ores of Kodur, Srikakulam district, Andhra Pradesh, India" reports an additional case of primary Mn minerals with an influx of colloidal solutions resulting in a manganese mineral assemblage of cryptomelane, coronadite (MnPbMn 6 4 0 14 ) and γ205

Mn0 2 . According to Roy, pyrolusite and cryptomelane are also present as supergene minerals. As primary manganese minerals are reported braunite, jacobsite with hausmannite as ex-solutions in jacobsite and minor pyrolusite, associated with garnet and sillimanite paragneisses and schists, and garnetiferous quartzites, intercalated calc-silicate gneisses and marbles. The primary Mn minerals reported are most probably initial sedimentary manganese ore metamorphosed. The influx of hydrothermal to colloidal solutions might be derivatives by dissolution of the metamorphic manganese and the supergene manganese might be due to mobilization/remobilization both of metamorphic and colloidal manganese minerals. The manganese ores from Kudor, India, clearly furnish a pattern of element transportation, transformation, mobilization and remobilization characteristic of the 'plexus of processes' of Mn mineral formation and mobilization. Since many of the 'primary' manganese mineral parageneses reported are actually metamorphic and derivatives of sedimentary manganese ores, it is important to consider some study cases of sedimentary and metamorphic manganese mineral formation and their relationship to the 'plexus of processes' (?cycle of processes). Byramjee and Meindre (1956) report a syngenetic sedimentary manganese occurrence at Guettara, southwest Algeria, which is associated with a pyroclastic intercalation in Precambrian rhyolites. According to them, although the ore is essentially braunite (3(Mn,Fe) 2 0 3 MnSi0 3 ), psilomelane (BaMn 2 Mn 4 0 16 (0H) 4 ), and pyrolusite (Mn0 2 ), other minerals are found: rhodonite (CaMn 4 (Si 5 0 15 )), bustamite (Ca3Mn 3 [Si 3 0 9 ] 2 ), spessartite (Mn 3 Al 2 [Si0 4 ] 3 with Fe), winchite (Si0 2 58, A1 2 0 3 3.6, Ti0 2 0.1, FeO 0.6, MnO 2.7, MgO 20.4, CaO 5.2, N a ^ 6.2, IC,0 1.4, F^O 1.5 = 99.7) and the arsenates titasite and possibly adelite. Considering these manganese minerals, psilomelane and pyrolusite might be derivatives of the metamorphic minerals representatives of which are ?rhodonite, spessartite, winchite. These manganese minerals might be metamorphic derivatives of syngenetic sedimentary minerals initially formed under syngenetic sedimentary conditions. The Guettara manganese occurrence reported indicates another example of the 'plexus of processes' responsible for the discussed manganese mineral occurrence. In contradistinction to the metamorphosed, initially sedimentary manganese deposits, the stratiform and lensoid Mn0 2 concentrations at the Forari Mn deposit, Efate island, New Hebrides, with todorokite occur in marine volcanic sediments which are in contact with overlying soil and limestone. According to Warden (1970), the distribution of the principal oxides (Mn, Si, Al), and some trace elements (Ba, Sr, Mo, P, Co, Ni), in several profiles show a dis206

tinctive covariance pattern. Fe is concentrated at or near the surface and Mn at a lower level; Mo and Ba are at a maximum in the ore horizon level, where there is a decrease in Si0 2 and A1 2 0 3 . As Warden supports, these features suggest supergene enrichment in deepweathering tropical soils due to precipitation of Mn and other metals from groundwater circulating in the zone of oxidation. It is suggested by Warden that the Mn was probably leached from submarine pyroclastic sediments of andesitic composition and transported in solution or partly in suspension. The 'plexus of processes' in the Forari manganese deposits suggests a volcanic sedimentary formation with main mineral todorokite and supergene Mn enrichment due to deep-weathering conditions of tropical soil (element leaching conditions). Discussing the 'plexus of processes' responsible for the formation, transformation and alteration of manganese minerals, metamorphic minerals (a part of the cycle of manganese mobilization processes) have been mentioned. However, it is necessary to elaborate on the processes responsible for the formation of manganese sedimentary or alteration minerals to metamorphic manganese assemblages. Roy (1964), considering the stability relations of manganese minerals as determined experimentally and the trend of mineral transformations in natural manganese ores subjected to regional metamorphism, support the following: "With the onset of metamorphism, the higher oxides of manganese with Mn 4+ dominant in the sedimentary deposits transform to Mn 2 0 3 and combine with available silica to form braunites; the stability range of braunite (3(Mn,Fe) 2 0 3 MnSi0 3 ) is very wide, but bixbyite ((Mn,Fe) 2 0 3 ) occurs, according to Roy, only in medium-grade metamorphism at about 500° C and that jacobsite ((Mn,Fe) 3 0 4 ) and vredenburgite ((Mn +2 ,Fe +2 )(Mn +3 ,Fe +3 ) 2 0 4 ) at higher grades of metamorphism". Roy's extrapolation of experimental results to natural metamorphic conditions should be subjected to general critics of metamorphism, namely metamorphic mineral crystalloblastic growths and may show a wide range of conditions of formation under metamorphism. As a corollary to the experimental results of Roy, Chattopadhyay (1967) supports in his description of manganese ores from Orissa, that they contain braunite, pyrolusite, cryptomelane, jacobsite, hausmannite, vredenburgite, hollandite (composition: cryptomelane-hollandite-coronadite), manganite (MnO(OH)), bixbyite ((Mn,Fe) 2 0 3 ) and that the paragenetic sequence indicates medium- to high-grade metamorphism. Additional studies on the transformation of Mn minerals by Roy and Purkait (1965) support that "the paragenesis and textural relations of the manganese oxide minerals from Guwari Wadhoma, Madhya Pradesh, India, mainly consisting of braunite (3(Mn,Fe) 2 0 3 MnSi0 3 ), bixbyite ((Mn,Fe) 2 0 3 ), hausmannite (Mn 3 0 4 ), jacobsite ((Mn,Fe) 3 Q 4 ),

hollandite (Ba(Mn +2 ,Mn +4 ,Fe +3 ) g 0 16 ), manganite (MnO(OH)), pyrolusite (Mn0 2 ), and cryptomelane (XR 2 Mn 6 4 0 16 , where X = IC,, Ba, Pb; R = Mn, Fe, Cu, Zn) are correlated with formation and transformation of phases in the system Fe 2 0 3 - Mn 2 0 3 (- air)".

All these study cases from the plethora of available international literature showing manganese mineral formation, transformation, alteration oxidation are in accordance with the stability of manganese mineral phases and the great mobilization and remobilization capacity of the element Mn.

207

Chapter 43

The Significance of Leaching and Diffusion Processes in Ore Formation

Augustithis (1983) edited an international volume on Leaching and diffusion in rocks and their weathering products, where the major fields were treated by the following researchers: Theory and Application of Adsorption and Ion Exchange Reaction Kinetics to in Situ Leaching of Ores by D. R. Cole; Kinetic Processes and Nature of Product Separated from Sulfate Solutions of Aluminium Ion at Various Hydrothermal Conditions by R. Vracar et al.; Kinetics of Dissolution and Structure of Aluminium Hydroxide Polymers by J. D. Heim; Kinetics and Mechanism of the Acid Kaolin Leaching Process by C. Z. Zivkovic et al.; Leaching in Rocks: Some Physical Principles by L. Stegena; Experimental Investigation of Water Diffusion in Silicate Rock Glasses at Elevated P-T Conditions by L. Pesty; Theoretical Evaluation of Diffusion-Controlled Oxygen Isotopic Exchange Between Silicates and Fluids at Elevated Temperatures by D. R. Cole; Origin by Metasomatic Diffusion of the Calc-Silicate Rocks at the Scheelite Rich Area of Morille-Salamanca. An Example of Metamorphic Differentiation in Upper Proterozoic Rocks from Iberian Peninsula by J. Saavedra and E. Pellitero; On leaching and Diffusion-Rings in Charnockites, Basalt, Trachytic Tuffite, Bauxite, Alunite and Andesite by S. S. Augustithis and A. Vgenopoulos; Pyritic Coal Spoils: Their Chemistry and Water Interactions by V. P. Evangelou; Leaching and Diffusion Studies in Manganese Oxides by M. A. Malati; Alteration in the Surface Composition of Some Silicate Minerals after Hydrothermal Treatment, Studies by Sims by J.-M. Beusen and R. Gijbels; Pianation Surfaces, Epeirogenesis, Water Tables by P. A. Hill; Chemical Changes of Granite During its Weathering by L. Minarik et al.; Detection of Neoformed Adularia by Rb-Sr Age Determinations of Granitic Rocks in Ohio by G. Faure and F. C. Barbis; Weathering and Leaching Patterns Associated with Some Gibbsitic Soils in Southern Africa by P. L. C. Grubb; Petrography and Geochemistry on the Weathering Crust of Cherty Iron-Formation in Wuyang Iron Mine Area, Henan Province and Quaternary Laterization in Hainan Island, both by W. Shousong et al.; Significance of Leaching Studies in Relation to the Development of Bauxite Profiles of India by K. S. Balasubramaniam; Leaching-Diffusion-Adsorption Processes in Rocks, Clays and Laterites from Kerala, 208

India by A. M. Nair and Narayanaswayn; About the Favourable Conditions for Bauxite Deferrification and the Problem of the White Bauxites in Greece by D. A. Kiskyras; The Role of Chemical Equilibria in the Leaching of Metal Ions from Soil Components by W. F. Pickering; Plant Nutrient Leaching in Tropical Soils by J. A. Zusevics and J. J. LaBrecque; Experimental Study of the Phosphate Sorption and Immobilization in Soil Components by P. Skrivan and J. Fafejtova; Mathematical Simulation of the Dispersion of Nitrates in the Fluvial Zone of the River Danube with Respect ofOutwash of Fertilizers from Soil by S. Gazda and K. Kovarik; Principles and Practice of Modeling Nitrate and Water Movement in Agricultural Soils by I. G. Burns and D. J. Greenwood. In addition to the extensive studies mentioned above, a number of other study cases pertaining to leaching and formation of ores are presented as follows. Peck (1967), in "Mass transport in porous rocks" outlines the physical processes and diffusion coefficients of movement of materials within impermeable and permeable porous rocks, considered with reference to the leaching and deposition of ores. It is found by Bj0rlykke (1960) that Norwegian copper-bearing pyrite deposits show a considerable higher content of Cu near the surface than in deeper parts. In the majority of cases it is assumed that such a zone existed prior to the period of denudation during the ice age, and that the Cu enrichment, now near the surface, was due to leaching of Cu from this zone and redeposited at lower levels. Davidson (1966), in "A possible mode of origin of stratabound copper ores" presents the hypothesis that stratabound copper ores originated, where chemical and structural conditions were favourable, from ascending brines which had leached the metal from primary sulfides of magmatic origin. The brines were derived from evaporite horizons from which they descended under gravity and after the reaction with the primary sulfides reascended, reaching regions of higher geothermal gradient. The hypothesis is further supported by a wide ranging examination of the geographical association of this type of Cu deposits with evaporite series and is applied especially to the Kupferschiefer of Germany and Poland.

Additional patterns of element leaching are also discussed. Vasilyev (1965), in "Tetrahedrite as a source of secondary cinnabar in Gorney Altai" supports that secondary cinnabar occurs in an oxidized ore body in the Dzhylkydal area (Gorney Altai) and the lack of unaltered cinnabar or its relics and the lack of native mercury, the most common alteration product of cinnabar, indicate that the secondary cinnabar (HgS) did not form from reprecipitation of dissolved cinnabar. Furthermore, as Vasilyev supports, tetrahedrite (Cu3SbS3; where Cu can be replaced but not completely by Ag, Zn, Fe, Hg) is the source of the secondarily formed cinnabar. Some of the vugs found have a tetrahedral outline, and locally contain relics of tetrahedrite. X-ray powder patterns and semi-quantitative spectrographic analysis of the tetrahedrite relics show that one sample contains mercury and silver and another resembles schwartzite (all the constituent elements of the tetrahedrite occur as trace elements in the cinnabar). It can thus be suggested that cinnabar was formed by Hg leached from tetrahedrite. In addition, Shnaider and Shnaider (1966), in "The occurrence of realgar in sulfide ores of NovoZolotushinskii and Kamyshinskii (Rudney Altai) deposit" report that the realgar in these deposits occurs with arsenopyrite, tennantite and rare löllingite, enargite and proustite and is the later formed mineral. According to them, realgar (AsS) was produced by supergene leaching and redeposition of As and As containing minerals. In the Kamyshinskii deposits, realgar was found in the near surface part of the ore zone where it formed fine crusts and elongated plates in fractures cutting pyrite-chalcopyrite and polymetallic ores. A special case of leaching in sulfides is reported by Poplavko (1967), in "On the mode of occurrence of rhenium in the sulfide ores of Dzhezkazgan". Poplavko carried out a variety of leaching experiments on the sulfide minerals dzhezkazganite, tailings and synthetic ReS 2 and Re 2 S 7 . The results indicate that solution of Re is enhanced by oxidation conditions and is rather variable in the Dzhezkazgan ores. Furthermore, it is concluded that most rhenium in the Dzhezkazgan sulfide ores occurs as micro-inclusions of rhenium minerals similar in their capacity for leaching and oxidation to Re 2 S 7 and dzhezkazganite. An interesting case of preferential leaching of Ag is reported by Lawrence (1958), in "The mineralogy and genetic significance of a Consuls-type vein in the main lode horizon, Broken Hill, NSW". As he reports, a silver-bearing siderite vein and a galena vein cut across the Main Lode horizon in the South Mine at Broken Hill. It is suggested by Lawrence, that the siderite vein and the mineralogically very similar Consuls Lode are of secondary hydrothermal origin and were formed as a result of preferential leaching of Ag from the Main Lode ore horizon. He furthermore proposed a tentative scheme involving the breakdown of primary argentian tetrahedrite to Ag sulfosalts and hence to silver-allargentum-dyscrassite assemblages. It is contended that

the Ag-rich minerals were deposited from metahydrothermal fluids. Uranium, and particularly U +6 , is geochemically very soluble and mobile and as already discussed (Chapter 15 and 16), leaching of U from primary uranium minerals containing U may involve complex mineralogical geochemical processes. Krasnikov and Sharkov (1962), in "Spatial and genetic relation between the exogenetic and metamorphic uranium deposits and the arid zones of the geological past" report that "large exogenetic uranium deposits can form only under arid conditions where: (1) intensive chemical weathering is accompanied by leaching of uranium from a thick permeable weathered mantle; (2) the uranium is highly mobile in the supergene zone, when water is available, so insuring accumulation of uranium from extensive areas and large volumes of weathered material; and (3) conditions are favourable for the precipitation of uranium from natural solutions on land and in adjacent epicontinental seas". They also point out that when plotted on palaeoclimatic maps, known exogenetic U deposits fit well into the arid zones of the past. Davidson (1965), in "The origin of banket ore bodies" suggests that the metals of Au + U deposits were leached from acid-intermediate volcanic material and that U of Au-free deposits was leached from granitic debris. In elaborating his hypothesis, Davidson states that "the occurrence of uraniferous and auriferous conglomerates close to major unconformities within or at the base of deep confined basins of Proterozoic molasse-type sediments is compatible with the view that the metals have been leached from overlying geological formations". The hypothesis involves post-leaching migration of the heavy solutions to the deepest permeable horizons, their reascent when these levels became hot in the course of orogenesis, and deposition of the metals in cooler higher regions where relatively open channels were available as in conglomerates. Davidson applied this hypothesis to interpret the deposits of the Dominion Reef, Witwatersrand, Blind River, Serra de Jacobina, Westmoreland (Australia), Krivoy Rog and South Karelia. Leaching of Fe is a common process and a brief consideration is included in Chapter 16. Similarly, manganese leaching is of great genetical significance and many of the manganese occurrences adjacent to the Red-Sea could perhaps be interpreted due to Mn leaching from intercontinental basalt buried in the Great Rift (see Tooms, 1976; Augustithis, 1982; and also Chapter 51). As a corollary to the leaching studies presented so far, additional study cases of Fe leaching are put forward as follows. Brandt (1964), in "The iron ore deposits of the Mount Goldsworthy area, Port Hedland district, Western Australia" reports that the iron ore deposits are associated with banded iron formations belonging to the (Middle Archaean) Warrawoona formation. The 209

two main types of the ore deposits where leaching processes played an important role are recognized by Brandt: (1) residual replacement deposits involving preferential leaching of gangue minerals from the protore beds and residual concentration of the ferric oxides; and (2) derived deposits in which concentration of iron oxides has occurred by leaching of iron-bearing detritus outside the protore beds. A most interesting case of Fe deposits involving Feleaching from basalts (comparable to Mn leaching from intercontinental basalts buried in the Great Rift, which resulted in Mn vein deposits in the adjacent Red Sea regions) is reported by Rosier (1962). According to him, the Upper Devonian layered iron ores of LahnDill type in the 'Schleizer Trog1 in east Thuringia are not exhalative (by transportation of Fe and Si as chlorides) but by hydrothermal solutions formed by heating of wet sediments during Upper Devonian volcanism. Furthermore, Rosier suggests that hot fluids decomposed basalts and tuffs and removed Fe, Si, Mg, Al, etc. It followed flocculation of iron silicate gel at the sea bottom and later oxidation of the silicate to haematite. It is suggested that the different types of ore were formed by successive C0 2 - and HjO-bearing hydrothermal solutions and their alteration effects. A case comparable to iron deposit formation involving leaching is reported by Treiber (1966), according to whom the Leuta-Vlahita iron deposits in the MioPliocene sediments are the result of post-volcanic processes. Limonite occurred during the early eruptive phase in the central Harghita area, but due to pH, iron hydroxide was leached by vadose water and siderite

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was later deposited by (vadose) metasomatism. As mentioned, the involvement of leaching processes is believed to be significant for the formation of this iron deposit. Somehow comparable to the formation of manganese minerals and deposits formed adjacent to the Red Sea Rift are cases of copper mineralization described by Ilani et al. (1987), in "Copper mineralization in sedimentary cover associated with tectonic elements and volcanism in Israel", where they maintain that fault zones may have provided conduits along which solutions in places possibly briny, could have risen and leached subsurface mineralized volcanic or sedimentary bodies. Also significant in this case were leaching, brines and faulting (as was the case with the manganese mineralization adjacent to the Red Sea Rift). Bonavia et al. (1986), in "Geology and ore enrichment factors at the Radiore mine, Quebec", considering "ore enrichment factors (wt% metal in ore/wt% metal in source rock) calculated from analyses of ore and source rocks, assuming that seawater-derived brines leached ore minerals from underlying rocks and precipitated them at the point of brine discharge onto the seafactor, that Cd, Zn, Cu, Au and Ag are most highly enriched, followed by Bi, Pb, Sn, As and Co. Mo, W, V, Cr, Mn and Ni are not enriched at Radiore." Concluding, it should be mentioned that most of the supergene processes involve element leaching and that leaching and diffusion processes are of significance in many weathering processes and deposits (see Augustithis, 1980, 1981, 1982; and Augustithis and Vgenopoulos, 1983).

Chapter 44

Redistribution - Mobilization Remobilization

The subject of element recycling, mobilization and remobilization is treated under several topics in the present volume, either in conjunction with syngenetic ore deposits or, in the case of remobilization, as affecting the element and mineral distribution (Part III, Chapter 62). Nevertheless, in the present Chapter study cases will be introduced showing that extensive mobilization of elements is possible in ore which can be due to diagenetic, metamorphic, tectonic and, in general, palingenetic processes. Shterenberg et al. (1964), in "The role of diagenesis in the formation of manganese ores" present analyses of concretions and the enclosing rock in Transuralian manganese carbonate ores of undoubted diagenetic origin indicating that manganese possesses great mobility during diagenesis. Therefore, diagenesis plays an important role for the formation of the ore deposits. According to them, Co is also concentrated during diagenesis. Fe concentration occurs in one group of deposits. Ni, P, Cr, V and Cu are not concentrated during diagenesis. In contradistinction, as Nixon (1963) reports, stratiform lead and copper sulfide mineralization of Mississippi Valley type occurs within a sedimentary sequence of lower Cambrian age. Mineralization occurs over a stratigraphic thickness of 400 ft. within algal dolomites, sandy dolomites, sandstones and shale. Sulfide minerals have been deposited contemporaneously with the sediments, though with subsequent recrystallization and mobilization. King (1966) reports epi-syngenetic mineralization in the English Midlands. The concepts of plutonism and neptunism are examined by King and applied to mineralogical environments. Furthermore, many mineralized bodies in the English Midlands, hitherto considered to be of hydrothermal origin, are now considered pseudo-neptunian or neo-neptunian 'dykes' filled with deposits of epigenetic mineral matter derived from a remobilized earlier syngenetic deposit. Von Rahden (1965), as mentioned, discussing the apparent fineness values of gold from two Witwatersrand gold mines, supports that variations in the apparent fineness values of gold, 1000Au/(Au+Ag) in crushed ore from two Witwatersrand conglomerates (Ventersdorp Contact Reef and Main Reef) show that

the samples with a high gold content have a high apparent fineness whereas low apparent fineness is more typical of low-grade ores. Von Rahden maintains that the answer to the variations in apparent fineness observed, lies in the redistribution of Ag relative to Au, during the 'solution' stage of the modified placer theory. Also Murray (1961), considering some aspects of the geology of Mount Isa, supports that previous theories on the mode of ore emplacements, involving at least four periods of deposition, are too complex, and according to him, the balance of evidence suggests that the mineralization is essentially syngenetic, and suffered later restricted mobilization during folding and fracturing. In contrast to sedimentogenic ore bodies, mobilization cases are presented where mobilization is attributed to metamorphism. George (1969), in "Sulfide vein formation during metamorphism of the Nairne pyrite deposit (South Australia)" supports that pyrite-rich axial plane and pyrrhotite-rich tension gash veins have originated by remobilization of sulfides during metamorphism. The mechanism of formation is believed to be that of diffusion of a dispersed phase along free energy gradients produced by rock fractures. Differences in mineralogy between these two vein types may be due to differences in their structural orientation and/or to the relative ages of the veins. Cox (1967), considering the regional environment of the Jacobina auriferous conglomerate in Brazil, reports that gold-bearing conglomerates are interbedded with pelitic schists and mafic and ultramafic rocks in the Precambrian Jacobina series of northeast Brazil. It is believed by Cox that the gold had a placer origin but that it was partially redistributed during a metamorphism that transformed the rocks into the amphibolite facies. Furthermore, Hiigi et al. (1967) report that an uranium-bearing zone occurs near the village Isdrables in the Rhone-Valley and that the ore occurs in chloritesericite-albite schists which grade into phyllites and quartzites. The most abundant U mineral is pitchblende and is often accompanied by abundant pyrite and small amounts of typical Cu-As-Pb sulfide mineralization. Minerals in the latter association include some uncommon minerals in the Alps, such as enargite, lantite, 211

idaite, covellite, cobaltite, molybdenite, gold and rutile. It is tentatively suggested by Hügi et al. that the original ore formation took place in clastic Permian deposits, and that remobilization occurred in the Tertiary Alpine metamorphism. Oelsner (1960) suggests that the criteria for mobilized primary Pb-Zn ore deposits are considered to include a relatively low iron content of the deposit overall, persistence of primary textures, and high vanadium contents particularly of the galena. Thermal energy is considered by him to represent the solvent leading to a re-solution of the precipitated material (without the addition of solvent chemicals). Tectonic heat (generated by friction) may give a temperature rise of up to about 100° C, which, according to Oelsner, is nearly twice that necessary for mobilization. In addition, Kornilov (1963) in his contribution on the redeposition of disseminated sulfides during the formation of copper-nickel sulfide deposits, reports that rootless sulfide veins are concentrated in shear zones near the contact of ultramafic rocks which contain disseminated ore minerals similar to the veins. Komilov supports that the ore bodies were formed after the ultramafics had completely consolidated and after the biotitization and talcification associated with the shearing. Another interesting interpretation concerning ore mobilization is reported by Dunnet and Moore (1970), in "Inhomogenous strain and the remobilization of ore and minerals". They include that the response of a rock sequence to a differential stress system includes a change of shape or strain of the rock mass and the migration of pore fluids. If the fluids carry ores in solution, it can lead to their selective remobilization and redeposition. The fluid will normally migrate from zones of high strain to those of low strain. However, the stress drop associated with late brittle structures may produce hydraulic fractures in zones of high strain

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which would become channel-ways for migrating fluids. As stated by Dunnet and Moore, examples of structural localization of Pb-Zn ores in southwest Sardinia, suggest that in different localities both high and low strain environments are locally enriched in remobilized ores. In addition, Tweto (1960), in "Scheelite in the Precambrian gneiss of Colorado" reports that scheelite is disseminated chiefly in calcsilicate gneiss but occurs in amphibolite as well. The deposits appear to vary in age and exhibit no consistent genetic associations. Therefore the tungsten in the region may have been repeatedly redistributed and recrystallized. In Chapter 1, in addition to pitchblende remobilization, cassiterite was also reported to be palingenic. furthermore, radiogenic lead is similarly considered to be remobilized. It should be emphasized that mineral or element remobilization is in accordance to the concept of element recycling. The fact that there is a great age gap (difference) between a metallic element generation (time of its generation or ? creation) and the last mobilization or remobilization of the particular element(s) to form or participate in the formation of a particular ore deposit, supports that most upper crust deposits consist of elements that have been subjected to multiple recycling processes. A deposit, the geological age of it, and geochronological determinations indicate that it is, for example, 200 m. y. old, might actually contain or consist of elements that have an age equal to that of the earth (with the exception of radiogenic elements) and that since then, they have been subjected to multiple recycling processes before they participated in the formation of the particular deposit, which might, in turn, be a temporary resting position before the element(s) are remobilized again in accordance with the unfolding of the geological spiral (see Chapter 1).

Chapter 45

Zonal Distribution of Elements and Minerals

As already mentioned in the case of metallogeny related to a 'granite intrusion' (see Edwards, 1960), the recognition of the pneumatolytic-pegmatitic phase, followed by hypothermal, mesothermal, hydrothermal, epithermal, telethermal has been widely accepted to be due to emanations and fluids-solutions from an intrusive granitic melt, after the orthomagmatic crystallization phase. Also, as pointed out, Augustithis (1990) considers perigranitic mineralization to represent a metamorphic-metasomatic differentiation of which the granite formation is a part of the processes in the order: granitization —> skarnification —> perigranitic zonal metallogeny. The pegmatites and aplites are seen as exudation products within the "field" (Bereich) of metamorphism-ultrametamorphism-granitization. The following extract is quoted from Augustithis (1990), Atlas of metamorphic-metasomatic textures and processes. "Considering the relationship between skarns and "hydrothermal perigranitic metallogeny" in many of the aureoles of perigranitic origin and the contention of Daly (1917) that the ore mineral parageneses are in reality metasomatic mobilizations, then we have once more an example of bulk composition changes due to perigranitic metamorphism-metasomatism. Furthermore, taking into consideration the three related processes namely granitization, skarnification and perigranitic metallogeny, it seems that we have a metamorphic-metasomatic differentiation where bulk compositions have been greatly modified." Consideration of Study Cases on Zonal Mineralization (Metallogeny) Based on Orthodox Views Yershov and Popova (1967), in "Primary zoning of copper-nickel sulfide ores at the outer contact of the Talnakh deposit", support that "outward from the gabbro-dolerite intrusive at the Talnakh deposit the following ores occur: (1) pyrrhotite with lesser pentlandite and chalcopyrite, (2) pentlandite-chalcopyrite with pyrrhotite, (3) pentlandite-chalcopyrite without pyrrhotite, (4) pentlandite-chalcopyrite with millerite, and (5) millerite-bornite-chalcopyrite with pyrite". According to Yershov and Popova, the zoning is attributed to differentiation by crystallization of a single ore forming liquid, emanating from the intrusion under the in-

fluence of isotherms concentric with the contact of the intrusion. Chasovitin and Pozdnyak (1965) report that ore occurs in quartz veins, altered granophyre dykes and brecciated sandstone and shale adjacent to the Vodorazdel'nyy granitoid (Chukotka). Mineral zones away from the west side of the stock are: (1) hypothermal tourmaline-quartz veins in the outer part of the hornfels with wolframite, scheelite, pyrrhotite, arsenopyrite, magnetite and subordinate gold and copper and zinc sulfides, (2) sulfide-quartz veins with arsenopyrite, chalcopyrite, pyrite, galena and gold, (3) carbonate-quartz veins with magnetite, hematite, arsenopyrite, chalcopyrite, sphalerite, boulangerite, galena and gold, (4) albite-carbonate-quartz veins with chalcopyrite, sphalerite, galena, gold, jamesonite, (5) collomorphic quartz veins with pyrite, chalcopyrite, galena, gold, silver tellurides, antimonite and cinnabar, and (6) sporadic quartz-carbonate veins with galena, gold and pyrite. The east side of the stock chiastolite hornfels contains stockwork-type quartz stringers with cassiterite, arsenopyrite and pyrite. Furthermore, Chasovitin (1964) reports in "Postmagmatic mineral zones of the Pyrkanay granite stock in northeast Kolyma (USSR)", that the six zones to the east are: (1) barren pegmatite veins, (2) hornfels with arsenopyrite, pyrite, pyrrhotite, galena and gold in its outer part, (3) gold, scheelite and sulfide-bearing quartz and epidote-quartz veins, (4) gold and sulfide in carbonate-quartz veins, (5) quartz veins with goethite, and (6) antimony in quartz veins. Considering the role of solutions, Korzhinsky (1959), in "The advancing wave of acidic components in ascending solutions and hydrothermal acid-base differentiation" supports that "in rocks altered by the flow of ascending solutions, vertical zoning should give (from bottom to top): a zone of increasing acidity, essentially the zones of autometasomatic alteration of magmatic rocks; a zone of maximum leaching; and a zone of fading acidity of the advancing wave, giving a predominance of precipitation over leaching in telethermal ore veins". In contrast to concepts of zonal differentiation, Pisa (1966), in "Minerogenesis of the Pb-Zn deposit at Bohutin near Pribram" reports that "after the intrusion of the Bohutin quartz diorite and the respective pegma213

tite, aplite, and quartz-molybdenite veins, gold-bearing quartz veins and later the youngest polymetallic ore veins were formed". Furthermore, the classical Pribram polyascendent Ag-Pb-Zn-Sb mineralization, whose most typical representative are the Klement veins at Bohutin, was formed in the course of four development stages of the deposit (polymetallic-galena-sphalerite and sulphantimonite, antimonite, and carbonate stages). Within these stages twelve periods of introduction can be discerned. Monoascendent and polyascendent zoning has been recognized and described by Pisa in the case of the complex mineralization occurring at Bohutin. Smirnov (I960), in "Types of hypogene zonality of hydrothermal ore bodies" supports that zonality in the distribution of ore forming mineral complexes in the outskirts of hydrothermal ore bodies may be due to successive elimination from the magma of ore-bearing solutions of different composition, or to a change in geological and physico-chemicals conditions of the ore-bearing solutions along their routes. Buryak (1967), in "Hypogene zoning in the old gold province of Siberia" supports that zoning in the ore bodies of the Lena goldfield consists of gold-sulfide veins in deeper horizons compared with gold-quartz veins at higher level. An additional type of zoning is described by Markham (1963), in his contribution "An interpretation of the Mount Lyell (Tasmania) copper ore pangenesis". Generally the paragenetic sequence is pyrite —> pyrite + chalcopyrite or bornite bornite + chalcocite -» bornite + digenite. According to Markham, this corresponds to a progressive increase in the ratios Cu/Fe and metal/sulfur and is in accordance to the mineral zoning in the district. In contradistinction to vertical zoning, Petrulian et al. (1965) describe lateral and vertical zoning in the case of copper mineralization at Deva. The mineralization is a stock-work type in a hypabyssal andesite which is slightly hydrothermally altered. The Cu minerals are mainly chalcopyrite and bornite with sparse chalcocite, digenite and covellite, some pyrite, molybdenite, gold, clausthalite, tetrahedrite, magnetite, haematite, ilmenite and gangue of quartz and calcite. In contrast to the cases of mineral zones due to solution derived from a magmatic source considered so far, a number of diverse cases of zoning will also be presented. Buryak (1965), in "Relationship between the mineralization of Precambrian strata and regional metamorphic zoning in the Vitim-Paton uplands" states that

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mineralization appears related to metamorphic grade so that (1) rocks of the sericite-chlorite subfacies of the greenschist facies contain quartz-carbonate veins, much pyrite and rich gold ore, (2) rocks of the biotitechlorite subfacies of the greenschist facies contain quartz and quartz-carbonate veins, pyrite and some pyrrhotite and some gold, (3) rocks of the epidote-amphibolite facies contain quartz veins, pyrrhotite, and a little gold, and (4) rocks of the amphibolite facies contain pegmatites, pyrrhotite and chalcopyrite and no gold. According to Buryak, a genetic relationship exists between ore genesis and metamorphism and granite emplacement. An additional case of zonal distribution not directly related to solutions derived from a magmatic source is described by Shawe (1966), in "Zonal distribution of elements in some uranium-vanadium roll and tabular ore bodies on the Colorado Plateau". According to him, U, V and Se are zonally distributed. The concentric elements were precipitated at the interface between two solutions with different properties, now represented by the ore boundary. A special case of zoning is reported by Iwao et al. (1954) within the wallrock alteration of the Kosaka mine, Akita Prefecture, Japan. They state that the limited mineral assemblages of the altered rocks associated with these epithermal Cu-Pb-Zn deposits in Tertiary dacite or rhyolite indicate formation at 350-400° C at an estimated depth of 2-3 km. They support that in this restricted pressure-temperature range, pH is considered the principal factor determining zoning within the altered rocks. In contrast to the hypotheses proposed by the presented studies, the calculated relative stabilities of covalent-bonded complexes is supported by Barnes (1962) to correspond more closely to the position of metals in zoned deposits than do the physical and chemical properties of the elements and minerals. It is further suggested by Barnes that ore transport is probably effected by this mechanism. Thus, the zonal distribution of minerals in a metallogenic field (Bereich) depends on the generation of a 'stimulus' (and its progressive diminishing outwards), source of materials and mechanism of transportation of elements or compounds. Since in this case too, element distribution-segregation is also concerned, the factors discussed in Part III (see Chapters 60, 61 and 62) are also applicable in the case of zonal distribution of elements and minerals, as outlined in the study cases quoted on ore mineral zoning.

Chapter 46

Source and Recipient Geoenvironments of Mineralization

(a) General For the discussion of the concept of metallogenic geoenvironment, two basic aspects should be briefly discussed: (i) the source geoenvironment and (ii) the recipient geoenvironment (see also Chapter 1).

(b) The Source Geoenvironment Concerning possible source geoenvironments, Goldschmidt (1922) indicated the existence of a sulfide-oxide level in the deeper interior of the earth, formed by primary liquation in the primordial evolution of the earth and augmented ever since by a rain of sulfide drops and oxides from cooling silicate magmas at higher levels. Goldschmidt (1954) further supposed "that there must be an intermediate level in the interior of the earth, below the silicate mantle, where sulfides and heavy oxide minerals increasingly accumulate according to their density", and for this level, of yet unknown depth and thickness, the name sulfide-oxide level is retained, as in any case, it must contain the main mass of sulfides of the globe and also a very great amount of oxides of the iron family. Whether Goldschmidt's sulfide-oxide level, which is supposed to be at a depths of more than 800 km (thickness of crystalloplastic mantle) could be the source environment of mineralization in the upper crust, gives rise to scepticism since even the deepest possible fractures or faults could not transect 800 km of crystalloplastic mantle in which forsteritic olivine most probably predominates. Nevertheless, in Chapter 18 study cases were quoted which support a relationship between metallogeny and deep crustal tectonics. Another source geoenvironment which could be of significance, is the mantle. The platinoid group of elements, as well as the B-group of elements (i. e., Cr, Μη, Fe, Co, Ni, Cu, Zn, Cd, Ag, Au), represented in many cases of mineralization (see Chapter 1 and Part III) could be derivatives of mantle. As mantle- projects, particularly those launched in the decade of the 70s support, mantle involvement in the upper crust is more significant than hitherto thought. Furthermore, mantle diapirism and lower crust/upper mantle involvement in crustal mobile zones renders mantle a significant metallogenic source environment.

When considering the element distribution-segregation (element concentration to form ore deposits - see Chapters 1 and 49), the group of elements Se, Te, Bi, Sb, As, Pb, Sn, (Tl, In), has been recognized as a distinct group, elements that show interrelations in accordance to the empirical laws of the periodic system. It has also been suggested that these elements are most likely lower crust derivatives and as isotope studies show, some can also be mantle derivatives, e. g., Pb. Researching the possible source geoenvironments for this group of elements, a crust derivation is possible. However, no concrete evidence can be provided that these elements are restricted to, or even abundant in, any primordial layer or even related to certain rock types. It is thus more difficult to speak of a specific geoenvironment than tentatively to suggest lower crust. Perhaps one of the current and future tasks of isotope geochemistry is to istermine the geoenvironment of derivation of each element involved in a paragenetic association. This is an enormous task but with modern instrumentation, it is within present possibilities. In contradistinction, the group of elements Zr, Hf (REE), Ti, Nb, Ta, Mo, W, U and ?Sn, are considered as elements more abundant in the granitic-pegmatitic crust geoenvironment and some of them could be characterized as granitophile elements. Perhaps the source and recipient geoenvironment of these elements is the realm of granitization including gneisses and the pegmatitic-aplitic types which are interpreted as exudation of gneisses-granites (see Drescher-Kaden, 1948, 1969, 1974; and Augustithis, 1962, 1973, 1993). Admittedly, the present theoretical source environments are only suggested tentatively, and whether a specific element is confined to a particular source geoenvironment is also arbitrary. Another factor complicating matters is that perhaps there are no vestiges left of the initial earth's crust. The earth's upper crust, as known, is a product of multiple geological cycles operated in the past, or more correctly, the product of an unfolding spiral (see Augustithis, 1973, and Chapter 1). Mobilization and remobilization processes and the recycling of elements, which nowadays is accepted as a geochemically plausible hypothesis further complicate the theory. In this connection, it should be emphasized that mobilization and remobilization include all processes 215

which involve material mobilization including volcanism, metasomatism, fluids and solution phases. Furthermore, when studying the recycling of elements, it should be emphasized (also mentioned in Chapter 62) that "a deposit, the geological age of it, and geochronical determinations indicate that it is for example, 200 m. y. old, might actually contain or consist of elements that have an age equal to that of the earth (with the exception of the radiogenic elements) which since then have been subjected to multiple recycling processes before they participated in the formation of the particular deposit, which might in turn be a temporary resting position before the element(s) are again remobilized in accordance with the unfolding of the geological spiral". Thus, elements in the process of recycling might change the geoenvironment, since their resting position might, among other factors, also depend on the agents of transportation and the stability of their compounds. So in addition to the source geoenvironment and to a temporary resting position in which an element may be found, the recipient geoenvironment (which could be a temporary resting position), should also be considered.

(c) Study Cases Taking into consideration the above mentioned plexus of processes, study cases on the environment (geoenvironment) of the generation and the resting position of metallic elements forming ore deposits, will briefly be presented. White (1968), in "Environments of generation of some base-metal ore deposits" presents four critical aspects in the generation of an ore deposit involving a hydrous phase. The four aspects are: (1) source of the ore constituents, (2) dissolution of the ore constituents in the hydrous phase, (3) migration of the metal-bearing fluid, and (4) selective precipitation of the ore constituents in favourable environments. Armands and Landergren (1960), in "Geochemical prospecting for uranium in northern Sweden; the enrichment of uranium in peat" report that a peat deposit 4.5 km northwest of Masugnsbyn, Norbotten County, Sweden, has a remarkably high average content of 600 ppm uranium. This is related to the occurrence of radioactive springs within the area of the peat. It is suggested that the peat has served as a "collector" for the uranium, REE, etc. According to them, the enrichment in uranium is due to long distance of Mg- and Ca-bicarbonates emanating from nearby dolomite whereby the bicarbonate waters serve as carriers of uranium. It should be pointed out that in this case, the dolomite might itself represent a 'temporary' resting position for the U and the REE and that the peat "collector" represents the recipient environment. Considering the significance of peat as a "collector" for U, it can be regarded as the recipient environment in which U and REE occur. Comparable is the pres216

ence of U in sandstones overlaying granitic-gneissic massifs (usually eroded granitic-gneissic complexes) which acted as the geoenvironment of the U and REE mineral formation, which by subsequent alterationweathering provided the solutions which impregnated the overlying sandstones and which often deposited complex secondary radioactive minerals (see "Formation of ore deposits", Vienna, 1974); the overlying sandstones acted as recipient geoenvironment. A special case where mineral paragenesis (ore mineral associations) depends on the geological environment is presented by Ishibashi (1960), in "Au - Ag tellurides from the Dat6 mine Hokkaido, Japan". According to him, the combination of Au, Bi and Te varies according to differences of geological environment. Under volcanic hydrothermal conditions, Au and Te combination is common while under plutonic hydrothermal conditions, Te combines with Bi rather than Au. When studying the prevailing geoenvironmental conditions of the past, it is often emphasized that contrary to the concept of 'uniformitarianism', conditions of diverse sedimentation have prevailed. Rutten (1969), in "Sedimentary ores of the Early and Middle Precambrian and the history of atmospheric oxygen" suggests that the evidence provided by the gold-pyrite "placers" of the Witwatersrand and by banded iron formations support an anoxygenic atmosphere in Precambrian time. This was modified by the evolution of photosynthesizing organisms and the enrichment of atmospheric oxygen has continued ever since. Taylor (1910-68), in "Sedimentary ores of iron and manganese and their origin" raises the question whether differences exist between Precambrian and later deposits. He supports that the present day formation of ferruginous and manganiferous sediments of all known occurrences indicate that chamosite lies within 10° of the equator. However, it might be difficult to extrapolate on this information, i. e., the environment of formation of Precambrian deposits. In this respect, Appel and LaBerge (1987) edited an international volume entitled Precambrian Iron Formations (Theophrastus Pubs., Athens, Greece). The following is stated (which indicates the complexity of interpreting the geoenvironmental conditions of some Precambrian deposits' formation): "Precambrian ironformation is an extremely variable rock type that was deposited over a long span of early Earth history. Because the rock is variable, and was deposited in a wide variety of environments, all authors do not agree on a common definition of iron-formation. Some authors restrict the definition of chert-bearing rocks; some include pyritic slates; others include sulfide-rich tuff or graywacke. Similarly, authors of papers on iron-formation have very different backgrounds and orientations; some are sedimentologists, others are stratigraphers; some are geochemists, other are microbiologists, while others again are economic geologists. This diversity leads to a broad range of ideas concerning the geology

of iron formations and contribute to a greater understanding of the nature and origin of this fascinating rock known as iron-formation." Considering the source and recipient geoenvironments, isotope geochemistry may provide evidence

that elements comprising a deposit may have a derivation incompatible with accepted interpretations, e. g., studies by Vinogradov and Grinenko (1966) support that 30-50% of sulfur in Sudbury-type deposits is derived from CaS0 4 (see Chapter 27).

217

Part III

On the Distribution of Elements and Ore Parageneses. The Empirical Laws of Element Segregation-Concentration in Ores.

Part HI, Dedicated to the Memory of Dimitri Ivanovitch Mendelejeff and Lothar Meyer

Chapter 47

The Empirical "Laws" of Element Segregation/Crystallochemistry/Isotope Chemistry Versus Genesis of Ores - State of the Art

Considering the textural patterns of ore minerals (Part I), it was often stated that these patterns represent chaotic relationships, in the sense that dissolution of a mineral might precede replacement by another mineral whereby a chaotic relationship of the two may be possible, that is no rules or laws can interpret this relationship. However, what seems most significant is that the textural patterns exhibited are understandable by the application of causation principles. In contrast to the author's concept of chaotic relationships, such as dissolution and replacement, researchers of ore genesis and especially ore microscopists often referred to ore genesis as the state of the art, as, for example, Ramdohr (1937) in a review paper did. Another reason for this description of ore genesis was the fact that the entire gamut of geological processes may be involved in ore genesis, therefore rendering the establishment of "rules" and "laws" almost impossible, and so scientists drew parallels between the scientific understanding of natural processes - what the genesis of ores actually is - and the art of interpreting ore genesis. It should be emphasized that ore microscopy was developing as a special approach for studying ore minerals, and, among others, the following contributions should be quoted: "Microscopical determination of the opaque minerals", (Murdoch, 1916); "Geology of the ore deposits in Kennecott, Alaska", (Bateman and McLaughlin, 1920); "Die mikroskopische Untersuchung undurchsichtiger Mineralien und Erze im auffallenden Licht und ihre Bedeutung für Mineralogie und Lagerstättenkunde", (Schneiderhöhn, 1920); "Mineral Deposits", (Lindgren, 1924); "Minerography and ore deposition", (van der Veen, 1925); "Etude microscopique de quelques minerals m6talliques", (Orcel and Plaza, 1928); "Lehrbuch der Erzmikroskopie", (Schneiderhöhn and Ramdohr, 1931); "Opaque oxides and sulphides in common igneous rocks", (Newhouse, 1936); "Economic mineral deposits", (Bateman, 1942); "Tables for microscopic identification of ore-minerals", (Uytenbogaardt and Westerveld, 1951); "Erzmikroskopisches Praktikum", (Schneiderhöhn, 1952); "On the opaque mineral constituents... Norra Storfjället, Västerbottoen, Sweden", (Uytenbogaardt, 1953); "Die Erzlagerstätten, Kurzvorlesungen", (Schneiderhöhn, 1958); "Texturi i Strukturi

Rud.", (Betechtin et al., 1958); "Textures of the ore minerals and their significance", (Edwards, 1960); "Atlas der wichtigsten Mineralparagenesen im Mikroskopischen Bild", (Öelsner, 1961); "Die Bildkartei der Erzmikroskopie", (Eds., Maucher and Rehwald, 1961); "Sulphide minerals - crystal chemistry parageneses and systematics", (Kostov and Minceva-Stefanova, 1981); "Ore microscopy and ore petrography", (Craig and Vaughan, 1981); "Atlas of ore minerals", (Picot and Johan, 1982); "Atlas of opaque and ore minerals", (Ixer, 1991). However, the period of 1920-1937 was the time in which the work of Ramdohr and Schneiderhöhn attained great significance, reaching its peak in 1960 with Ramdohr's "Erzmineralien und ihre Verwachsungen", (1960 - and its subsequent editions). During the same period of 1923-1933, Goldschmidt and his Oslo/Göttingen school were unfolding the Verteilungsgesetze der Elemente. Despite the fact that both these developments can be described as running parallel to one another, and that both contributed enormously to the increase of our knowledge, somehow each one respected the other trend and possible confrontation of principles was avoided. In time, geochemistry evolved and embraced almost all fields of mineralogical and penological research. The great revolution in ore microscopy though took place with the introduction of microprobe, which in fact is a spot-microanalysis inevitably bringing more geochemistry to ore microscopy. The mutual interdependence of ore microscopy-microanalysis-geochemistry is of decisive significance in the progress of both ore genesis and geochemistry. Even in 1990, and despite these important developments, the genesis of ores is still claimed to be "a state of the art"; while in the meantime, geochemistry with the "phase rule" and the application of the principles of the "laws of the distribution of elements" of Goldschmidt, and the analytical techniques which have developed, have given new perspectives to mineralogy and petrology in general. If ore mineralogy though remains a state of the art and does not embrace the progress made in petrology and geochemistry, and particularly isotope geochemistry, it may turn into a museum exhibit. However, the developments of the late 80s show ore mineral studies are more and more 219

dependent on geochemistry and especially isotope geochemistry and microprobe analysis. Ore microscopists have, of course, learned (by trial and error) that "laws" and "mles" for general acceptance are difficult to find, and thus they preferred the "study case" approach (which is also followed by the author, in Part II). However, an attempt should be made, using the inductive approach, to come to some conclusions where unifying principles could evolve. Augustithis (1964), in "Geochemical and ore microscopic studies of hydrothermal and pegmatitic primary uranium parageneses", tried by applying the interrelationships of elements in accordance to the empirical "laws" of the periodic system, to explain the common segregation of rare elements in mineral associations (parageneses) in conjunction with Goldschmidt's "laws" of element distribution. Since 1964, the author has attempted to interpret other mineral paragenetic associations by referring to these interrelationships (see Augustithis, 1967, 1979; and Augustithis and Vgenopoulos, 1982). Furthermore, considering the segregation of metallic elements to form ore concentrations, the author wants to reiterate that the following are of great significance: (i) The relationship (or interrelationship) of the elements according the empirical "laws" of the periodic system, which in turn ultimately results in the common segregation of metallic elements in ore concentrations (associations or parageneses) which were originally distributed in minute quantities in the crust or mantle.

220

(ii) The mechanism of transportation of these elements and whether they have been transported or mobilized in the state of compounds, in melts, fluids, gases, solutions (hydrothermal or low temperature solutions). The generation of these mechanisms of transportation in diverse geoenvironmental conditions (e. g., mantle diapirism, lower crust/upper mantle obduction, magmatism, volcanism, recycling of elements in the geological cycle or in the unfolding of the geological spiral) should also be taken into consideration. (iii) The crystallization of minerals depends on the physico-chemical conditions, i. e., temperature, pressure and crystallochemical factors (as shown most extensively in the work of Goldschmidt (1922-33; 1938; 1954). The relationship between crystal structure and chemical composition can be summarized as depending on: (1) The various proportions of the various kinds of atoms (or ions) in the chemical formula, (2) The relative size (radii) of the various kinds of particles, i. e., atoms (or ions) in a crystal. (iv) The geoenvironmental condition of element mobilization and remobilization. Each of the above mentioned factors is in itself an enormous field of knowledge or research potential (and each one presents a specialized field) which is beyond the scope of the present volume. An attempt will nevertheless be made to introduce examples of such studies.

Chapter 48

Segregation of Elements in Accordance with Their Interrelationships to Form Mineral Associations-Parageneses

(a) General Principles - General of the Periodic System

Conception

Considering the relationships between chemical properties and one fundamental constant, the atomic weight, Meyer (1868) and Mendelejeff (1869) 13 , found that the chemical and partly the physical properties of elements are a periodic function of this constant. The concept of periodicity is expressed by Remy as follows: "In the natural system of elements four types of periods fundamentally exist. The first of these contains only two elements, hydrogen and helium, the following two contain 8 each, and the two which follow after 18 each; but the period which follows after contains 32 elements, namely all elements from cesium to radon included. This particularly extensive period includes as a special sub-division the lanthanides. With the elements following the radon a further extensive period begins. The sequence of the elements is interrupted, however, before it reaches its end. At one time uranium was the heaviest element known but now still heavier elements referred to as the transuranides are known". Table II (from Remy's Lehrbuch der anorganischen Chemie, Bd. 1, p. 10), illustrates this classification. Moreover, in Bohr's theory of atomic structure the periodicity of the elements is actually a function of the atomic structure and the theory also provides a basis for the increasing lengths of the periods. The grouping of the elements into periods can be expressed as 2:8:18:32. Thus, the progressive increase in the number of elements per group follows a relatively simple law, so that the periodicity is a function of the repetition of homologies in the structure of the atoms. Besides the division of elements into periods, main groups and subgroups are recognized. The elements of the main groups are considered as those with ordinal numbers either 1 to 3 units greater than, or 4 to 1 units less than the ordinal number of a rare gas element. The subgroups contain the remaining elements excluding 13 As is supported in van Spronsen's "The periodic system of chemical elements", (1969), in the section entitled 'Priority conflict between Mendelejeff and Meyer' (p. 342), Meyer has acknowledged that Mendelejeff preceded him in recognizing the periodicity of elements.

14 which follow lanthanium and which are consequently classified as lanthanides. The elements of the subgroups of the periodic system, as well as the elements of the main groups, follow the laws of periodicity. Elements belonging to the same subgroup show close similarity. Every subgroup is also connected with a main group, and together they form a family. If elements show similarity in their behaviour, or if in their properties their compounds correspond, and even if the elements do not belong to the same main or subgroup, it can be said that analogies exist between the elements or their compounds and consequently these elements are analogous. In contradistinction, elements belonging to the same main or subgroup of the periodic system are referred to as homologous.

(b) The Empirical "Laws" of the Periodic System The following empirical laws and discussions of their causes are again based on the work of H. Remy. (i) The similarity between the main and the subgroup elements of the same family increases strongly from families I to IV and decreases strongly from IV to VIII. In trying to understand this empirical law, the very recognition of the main and the subgroups depends on the conception of the ordinal number and on the rare gases ("inert gases"), the recognition of which depends in turn on their valency. Consequently, valency is the primary factor on which the law depends. The relationship expressed in this law can be understood by considering Table III (also from Remy's Lehrbuch der anorganischen Chemie, Bd. 1, listing the elements on the basis of the "inert gases"). If families I to III are considered, the valency of the elements of the main and the subgroups is positive and uniform within each separate family, although in proceeding from families I to III the value of the valency increases. Consequently, the increase in similarity of the main and subgroup elements is in accordance with this increase in valency. Comparing the valencies of the main and subgroups of families IV to VII, however, it can be seen that whereas the valencies of the subgroups are 221

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positive and increasing, those of the main groups can be both positive and negative, with the positive increasing and the negative decreasing from family IV to VII. In general, an explanation should be sought for the increasing and decreasing similarity of the main and subgroups. (ii) Between the main and the subgroup belonging to the same family there is a marked similarity between the second element of the main group and the first of the subgroup. This law is subject to an increase from families I to IV and a decrease from IV to VII (the interrelationship between the VIII main and subgroup elemnets is almost non-existent). This tendency should be seen from the aspect of valency since it depends on the relationship between the elements of the main and subgroups. (iii) The similarities and affinities between the elements are illustrated by two additional empirical laws, which also depend on the valency and on the atomic structure of the elements: (1) Among elements belonging to the same subgroup there are always two that are particularly similar to each other, and for those subgroups which are made up of transition elements the following is applicable: from the elements (within the same subgroup) listed vertically within the same column, the second element reading downwards shows a greater similarity with the third than with the first. The similarity is greater than that which is observed between the third and the fourth element. (2) The similarity in behaviour of elements listed along the same horizontal row increases from subgroups III to VIII but decreases from subgroups I to III.

(c) Special Cases of Element (Interrelationship)

Segregation

Regarding the similarity and relationship of elements, special cases of strong interrelationships between homologous elements must be considered. Those ho-

224

mologous elements which are transition elements, and which in their chemical behaviour and in the properties of their compounds are closely similar, possess atomic radii differing little from one another. It must be mentioned, however, that besides analogies in the structure of the outer layers and small differences in the atomic radii, there are other factors which are active and which determine special cases of similarity. The greatest similarity in behaviour and in properties of their compounds occurs between the two elements zirconium and hafnium. In this case, however, the atomic radii differ more from each other than in the case of the two elements niobium and tantalum, where on the other hand, the chemical similarity is undoubtedly less. Moreover, whether this tendency for association of elements results from the atomic weight of these homologous elements being comparable or whether from other special causes, can be computed from a simple mathematical consideration. In the subgroup of family IV the difference between the atomic weights of Hf (178.6) and Zr (91.2) is 87.4. This small calculation apparently suggests that the greater tendency for association between Hf and Zr compared to that of Hf and Ti is due to the smaller difference in atomic weights. Considerations of the atomic weight relations of the elements, however, show that the attribution of tendency for association to resemblance between atomic weights would not be valid genuinely. This can be seen from the following calculations: the difference between the atomic weights of Zr (91.2) and Ti (47.9) is 43.3 and this difference is nevertheless smaller than that between Hf (178.6) and Zr (91.2), which is 87.4. The similarity between Hf and Zr is one of the greatest known. According to Remy, the interrelationship is as follows: "In nature Hf always occurs with Zr, to which it is so similar that the chemical separation of Hf from Zr is difficult". It can thus be repeated and emphasized that in general the main cause for the empirical laws of element affinities and relations is the atomic structure of the elements.

Chapter 49

Common (Joint) Segregation of Elements

(a) The Interrelationship of the PGE (and Other Related Elements) and Their Common Segregation in Forming PGM in Diverse Geoenvironmental Conditions It is interesting to note that the Pt group elements and the formation of the Pt group minerals can be present in a wide range of distribution and genetic possibilities, e. g., mantle, magmatic rocks, pneumatolytic-hydrothermal occurrences and finally as lateritic paragenetic associations. Considering most of the known Pt group minerals from a wide range of parageneses (Table IV), it can be seen that the mineral building elements can be grouped into two independent groups. Group A (most probably ?lower crust/crust derivatives, since in this particular case they form PGM in association with PGE) consisting of the elements As, Sb, Bi, Sn, Te and Pb, Se14 and group B, consisting of Fe, Ni, Co, Jlu, Rh, Pd, Os, Ir, Pt (subgroup elements of the VIII family of the periodic system) and Cu and Au15. The elements As, Sb, Bi are related to one another as elements of the main group, homologue elements of the V family of the periodic system. Also the elements Sn, Sb, Te are related as main group elements, next to each other. Therefore, it is to be seen that all the elements of the A group are interrelated in accordance with the empirical laws of the periodic system (see Augustithis, 1964). Considering the element geochemistry and particularly the abundance of the elements comprising groups Β and A, which build up most of the Pt group minerals (see Table IV), it is clear that the joint segregation of very rare elements in the earth's outer part cannot be accidental but is governed by "laws" which enable their joint segregation in Pt minerals and parageneses. These "laws" are proposed as the empirical "laws" of the periodic system which suggest the relationships of elements between themselves. Another interesting case of PGM association is reported by Watkinson et al. (1991), in "Platinum group

mineralization in gabbroic rocks, Two Duck Lake intrusion, Coldwell Complex, Canada". According to them, PGM occur in coarse-grained to pegmatitic fades of the Two Duck Lake intrusion. They report that the deposit comprises Cu-rich and Pd-rich sulphides and, in general, the following ore minerals: pyrrhotite (FeS), chalcopyrite (CuFeS2), cubanite (CuFe2S2), mackinawite ((Fe,Co,Ni)S?), pentlandite ((FeNi)S), Ag-pentlandite, magnetite (Fe2Fe304), ilmenite (FeTiOj), kotulskite (PdTe), lead-rich kotulskite, zvyaginstevite (Pc^Asj), merteite II (PdgSb3), palladoarsenide (Pd^As), majakite (PdNiAs), hollingsworthite (RhAsS), and an unnamed PdsAs2. Considering the elements comprising the minerals mentioned, the following two groups are tentatively recognized: group A: Pb As, Sn, Te, and group B: Fe, (Ti), Ni, Co, Pt, Pd, Rh, Cu and Ag. It should be noted that group A and group Β of the Two Duck Lake intrusion are geochemically comparable to group A and B. In this case too, the elements comprising each group are interrelated in accordance with the empirical "laws" of the periodic system as mentioned. Comparably, the consideration of PGM of ophiolitic mantle and cumulates by Ohnenstetter et al. (1991), in "Platinum-group minerals in ophiolitic mantle and cumulates of Albania: preliminary results" support that the metallic elements characteristic of these associations are interrelated in accordance with the empirical "laws" of the periodic system. In particular they report, that mantle chromites show high (Ru + Os + Ir)/(Pt + Pd) compatible with the presence of rutheriridosmine, osmiridium, iridian ruthenium, leurite, erlichmanite and Ir-bearing sulphoarsenides. In contradistinction, PGE mineralization in cumulates has the highest PGE tenors [(Pt + Pd) up to 9 g/t] with correspondingly low (Ru + Os + Ir)/(Pt + Pd). Considering the elements reported, in addition to the PGE which belong to the subgroup of the VIII family of the periodic system, Cu is also reported (Fe, Ni, Co are also subgroup elements of the VIII family) which again is interrelated to these elements in accordance with the empirical "laws" of the periodic system.

14

Pb is related to Bi as next to it and Se is related to Te as homologues. Considering Pb isotope studies (see page 193) - support both crust and mantle derivation. 15 Cu and Au are related as subgroup elements of the I family and Hg to Au as adjacent elements. 225

Table IVa (based on the glossary of platinum-group minerals by Cabri, 1976) Mineral

General Composition

Mineral

General Composition

Arsenopalladinite Atheneite

Pds(As,Sb)2 (Pd,Hg,Au,Cu)3(As,Sb) Pd, Hg, As major (Pd, Pt)3Sn Pd > Pt (Pt,Pd)(Te,Bi)2 Pt > Pd; Te > Bi (Pd,Pt,Ni)(Te,Bi)2 Pd1+,(As,Pb)2 As= Pb; χ < 0.2 (Pd,Pt,Ni)15(Sb,Bi)05Te200 Pd > Pt, Ni; Sb > Bi (Pt,Pd,Ni)S Pt > Pd > Ni (Pt,Pd,Ni)S Pt > Pd, Ni (Cu,Pt)2AsS2 Cu > Pt (Os,Ir,Rh,Ru,Pd)S2 Os > other Pt group (Pd,Pt)(Bi,Te)2 Pd > Pt; Bi > Te Pt(Sb,AsBi)2 Sb > As, Bi PdjAs (Ni,Pd)2(Sb,Bi)Te Ni > Pd; Sb > Bi (Rh,Ru,Pt,Pd,Co,Ni)AsS Rh > other Pt group PtCuAs (Pt,Pd,Ni)(Bi,Te,Sb,Sn)2 Pt > Pd, Ni; Bi > Te, Sb, Sn (Ir,Ru,Rh,Pt,Pd,Os,Ni,Co)AsS Ir > other metals (Ir,Ru,Os,Rh,Pt,Cu)(As,S)2 Ir > other metals Ir, Os, Ru, Pt, Pd, Rh, Fe, Ni Ir major Os, Ir, Ru, Pt, Pd, Rh, Fe, Cu, Ni Os, Ir major ~(Pt,Ir,Os,Ru,Rh)3(Fe,Ni,Cu,Sb) Pt > other Pt group; Fe > other metals (Pd,Cu)5(Sb,As)2 Sb = As (Pd,Ni)(Te,Bi,Sb,Pb) Pd > Ni; Te > Bi, Sb, Pb (Ru,Ir,Os)S2 Ru > Ir, Os (Cu,Pt,Ir,Pd,Fe,Ni)S2 Cu, Pt, Ir major (Pd,Pt,Ni)(Te,Bi,Sb)2 Pd > Pt, Ni; Te > Bi, Sb

Mertieite

~(Pd,Cu)5(Sb,As)2 Pd > Cu; Sb = As = Group I Sb > As = Group II (Pd,Pt,Ni)(Bi,Sb)Te Pd > Pt, Ni; Bi > Sb (Pt,Pd,Ni)(Te,Bi,Sb)2 Pt > Pd, Ni; Te > Bi, Sb (Pt,Pd)(Sn,Bi,Te) Pt > Pd; Sn > Bi, Te (Pd,Cu)7Ses Pd > Cu (Os,Ru,Ni,Ir,Pd,Pt,Rh)AsS Os > other metals Ir, Os, Ru, Pt, Pd, Rh, Fe, Cu, Ni Ir, Os major Os, Ir, Ru, Pt, Pd, Rh, Fe, Cu, Ni Os major Pd, Pb, Rh, Pt, Os, Ir Pd > other metals (Pd,Pt,Au,Cu)2(As,Sb,Te) Pd > Pt, Au, Cu; As > Sb, Te

Atokite Biteplapalladite Biteplatinite Borishanskiite Borovskite Braggite Cooperite Daomanite Erlichmanite Froodite Geversite Guanglinite Hexatestibiopanickelite Hollingworthite Hongshiite Insizwaite Irarsite Iridarsenite Iridium Iridosmine Isoferroplatinum

Isomertieite Kotulskite

Malanite Merenskyite

226

Michenerite Moncheite Niggliite Ooosterboschite Osarsite Osmiridium Osmium Palladium Palladoarsenide Palladobismutharsenide Paolovite Platiniridium Platinum Plumbopalladinite Polarite Potarite Rhodium Rustenburgite Ruthenarsenite Rutheniridosmine Ruthenium Ruthenosmiridium Sobolevskite Sperrylite Stannopalladinite

Pd

2Aso.8Bio.2

(Pd,Pt)2Sn Pd > Pt Ir, Pt, Os, Ru, Fe, Cu, Ni Ir, Pt > other metals Pt, Pd, Ir etc. Pt > other metals (Pd,Ag)3(Pb,Bi,Sn,Cu,Sb)2 Pd > Ag; Pb > other elements Pd(Bi,Pb) Bi> Pb PdHg Rh, Pt Rh > Pt (Pt,Pd)3Sn Pt > Pd (Ru,Ni,Rh,Ir,Pd,Os)As Ru, Os, Ir, Pt, Pd, Rh, Fe, Cu, Ni Ru, Os, Ir major Ru, Ir, Rh, Pt, Os, Pd, Fe Ru major Ir, Os, Ru, Pt, Pd, Rh, Fe, Ni Ir, Os, Ru major Pd 107 Bi minor Pt, Pb, Sb, Te reported (Pt,Rh,Ir)(As,Sb,S)2 Pt > Rh, Ir; As > Sb, S (Pd,Cu)3Sn2 Pd > Cu

(continued on next page)

Mineral

General Composition

Stibiopalladinite

(Pd,Cu)5+x(Sb,As,Sn)2.x Pd » Cu; Sb » As, Sn Pdg(As,Sb,Te,Sn,Bi)3 As > Sb, Te, Sn, Bi Pt(Sb,Bi) Sb > Bi (Pd,Ni)(Sb,Bi,Te,As) Pd > Ni; Sb > Bi, Te, As (Pd,Ag,Pb,Bi)4+x(Te,Se) Pd, Ag > Pb, Bi; Te > Se Pd2Hg(Te,Bi,Sb)3 Te > Bi, Sb (Pd,Ni)(Sb,Bi)Te Pd > Ni; Sb > Bi Pt, Fe, Ir, Cu, Ni, Sb Pt > Fe > Ir, Cu, Ni, Sb (Pt,Ir)2(Fe,Cu,Ni,Sb)2 Pt > Ir; Fe > Cu > Ni, Sb (Pd,Pt)3(As,Sb,Te) Pd > Pt; As:(Sb + Te) =1:1 (Pd,Ni,Pt)S Pd > Ni, Pt (Ir,Cu,Rh,Pb,Os,Pt,Fe)S Ir, Cu, Rh major

Stillwaterite Stumpflite Sudburyite Telargpalite Temagamite Testibiopalladite Tetraferroplatinum Tulameenite Vincentite Vysotskite Xingzhongite Yixunite Zvyagintsevite

(Pd,Pt,Au)3(Pb,Sn) Pd > Pt, Au; Pb > Sn

Table IVb: Precious metals in the upper lithosphere (compiled from Goldschmidt, Mason and Wedepohl) Β group 0.4 Ru = Rh = 0.4 4 Pd = Os = 0.4 Ir = 0.4 2 Pt = Fe = 50,000 Ni = 80 Co = 23 Cu = 70 0.001 Au =

A group ppb ppb ppb ppb ppb ppb ppm ppm ppm ppm ppm

As = 5 Sb = 21 Bi = 0.2 Sn = 40 Te = 0.002 Pb = 16

ppm ppm ppm ppm ppm ppm

(b) The Common Segregation ofPGE in Forming PGM Under Lateritic Conditions (Pt Nugget Formation) Due to the alteration of the ultrabasics and birbirite of the Yubdo complex, a lateritic cover has been formed in which Pt nuggets occur. The absence of native Pt in

the form and size of the nuggets in the ultrabasics and of birbirite and, in addition, the angular protuberances of the nuggets in the laterite suggest that they have grown by element agglutination/(accretion) as a result of the lateritization of the ultrabasics and birbirite. Additional evidence and geochemical studies of these nuggets by Ottemann and Augustithis (1967), Augustithis (1967) support this mode of origin. X-ray studies by Wolbeck (In: Augustithis, 1965) proved the presence of sperrylite in the Yubdo dunite which may be one of the carriers of the platinoid element. Decomposition of the primary minerals and subsequent agglutination of the platinoid elements has most probably taken place as a result of lateritization. Chromite grains originally present in the Yubdo dunite have partly resisted alteration (due to birbiritization or lateritization) and have acted as nuclei around which "ferroplatin" and other platinoid minerals have formed. Often a limonitic coating covers the nuggets which has also grown within the laterite (Fig. 723). Ore microscopic and microprobe studies by Ottemann and Augustithis showed that the main mass of the nuggets consists of "ferroplatin", composed of Pt, Fe, Au16 and little Co. In addition to osmiridium, other platinoid minerals were determined (by electron scanning) and new minerals were discovered, e. g., roseite, Os, Ir(S). The following is a tentative list of the mineral phases as determined so far, which constitute the nuggets: (i) Chromite nuclei of the nuggets (partly altered remnants of the dunite), essentially: FeCr 2 0 4 . (ii) Limonite, 2Fe03-3H20 (as coatings of the nuggets; some Mn may be present). (iii) "Ferroplatin"; Pt, Fe, Au and little Co. (iv) Osmiridium; Os, Ir. (v) Roseite; Os, Ir(S). (vi) Mineral (a)17, consisting of Ni, S, Pd, Rh and Fe. (vii) Mineral (b), consisting of Ni, Co, Fe and a little Pd. (viii) Mineral (c), consisting of Rh, Pd and Pt. (ix) Mineral (d), consisting of Ru, some Rh and Pd.

16 Concerning the relationship of Au with PGE and in general with basic and ultrabasic rocks, Hu Lunchi and Qi Changmuo (1985) support that gold is more abundant with basic and ultrabasic rocks than with acid. As a corollary, Panagos et al. (1982) found gold by neutron activation analysis inhomogeneously distributed in Greek chromites. Furthermore, Boyd et al. (1991) in their study of Norwegian ophiolites documented Pt-Pd-Au mineralization in ultramafic cumulates in the Leka and Lyngen ophiolites. The author suggested that the Au in the Yubdo Pt nuggets might be derivative of weathered Au quartz veins common in Wollaga, W. Ethiopia. However, the possibility should not be excluded that the gold in the Yubdo Pt nuggets might be derivative from the ultrabasics of the Yubdo complex. 17

Minerals (a), (b), (c) and (d) have been determined in the nuggets by microprobe analysis (electron scanning).

227

A geochemical consideration of the elements constituting the mineral which form the nuggets show that these elements are interrelated and show affinities in accordance with the empirical laws of the periodic system (Augustithis, 1967). On the basis of the mineral composition forming the nuggets, the following metallic mineral building elements are found: Pt, Ir, Os, Pd, Rh, Ru, Fe, Co, Ni (i. e., elements in the subgroup of the VIII family of the periodic system) and, in addition, Mn, Cr and Au. Considering the periodic table (Remy, 1959) where the elements are grouped on the basis of "inert" gases 0Gruppierung der Elemente um die Edelgase), it can be seen that the elements Pt, Ir and Os are interrelated they are next to each other and belong to the same subgroup of the VIII family. The same is true for the elements Ru, Rh, Pd and Fe, Co, Ni. All these elements are interrelated because of an increase of the horizontal and vertical affinities of the elements in the VIII subgroup. The elements Mn and Cr are related since they are next to each other and are similarly related to the group Fe, Co, Ni. Also, Au is related to Pt, being adjacent in the periodic table. Considering the relative abundance of the mineral building elements in the earth's crust present in the nuggets (Table V), it can be seen that very rare elements have segregated jointly to form these nuggets. Table V: Abundance of the metallic mineral building elements of the nuggets, based on Goldschmidt (1954) Elements

Abundance[ppm]

Fe Co Ni Ru Rh Pd Os Ir Pt Mn Cr18 Au

50000 40 100 0.001 0.010 0.001 0.005 1000 200 0.001

It is proposed that the joint segregation/(accretion) of these rare elements, which constitute the mineral building elements of the nuggets, depends on their affinities and chemical relations as outlined in the empirical "laws" of the periodic system and in compliance to the factors mentioned in Chapters 48 and 49.

18

Cr is represented by the chromite grains (relics of the dunite) which have acted as nuclei for the agglutination of the other elements of the nuggets; the chemical affinities of Cr with these elements might have played a role. 228

(c) The Segregation (Distribution) of Fe, Ni, Co, Cu and Zn - The Common Segregation and Mobilization (Remobilization) of Fe, Ni, Co, Cu and Zn in Diverse Geoenvironments As it was the case with Cr and Pt minerals, Fe, Ni, Co, Cu, Zn minerals may exhibit the same patterns of distribution, namely: The elements Ni, Co, Cu, Zn may occur as trace elements in the mineral components of the mantle. This might be referred to as the primary distribution of these elements. In contrast, the high temperature sulphide cumulus, particularly in igneous rocks, could again represent a primary distribution as mobilized derivatives from the mantle. In opposition, the possibility that in orogenic ultrabasic bodies a progressive recrystallization of the sulphides takes place with serpentinization, could be interpreted as to serpentinization resulting in a remobilization of the elements Fe, Ni, Co, Cu, Zn from primary distribution in mantle (the orogenic ultrabasics represent mantle diapirs, see Chapter 4) to sulphides within the serpentinized ultrabasics. The "re-equilibration" of sulphide ores in ultrabasics under low temperatures (Kullerud et al., 1959) could be understood as a case of remobilization of these elements (Fe, Ni, Co, Cu, Zn). The particular remobilization of Ni under laterization is discussed by Augustithis (1962) and Agiorgitis (1972). The element geochemistry of the group Fe, Ni, Co, Cu, Zn can be understood on the basis of their relationship in accordance with the empirical "laws" of the periodic system. The elements Fe, Ni, Co are related as subgroup elements of the VIII family, the element Cu is related to Ni as next to it and similarly, Zn with Cu. The joint segregation of the elements in the mantle (primary distribution) and in the magmatic (lava and sills sulphides) as well as a re-equilibration phase of Fe-Ni-Cu-Co-Zn sulphides in ultrabasics is explicable by their interrelationship (see above). In this connection it should be pointed out that the intercontinental ore deposits are recently considered to represent leached out elements from basaltic flows (or from the crust) due to the percolating brines (sea water). They are thus considered to represent a redistribution of elements. In contradistinction the Ni deposits which are associated with Archaean greenstones, again represent a redistribution of elements originally present in the mantle (particularly in the mantle forsterites), the mechanism, however, being remelting of material.

(d) The Common Segregation ofCu, Zn, Ag, Cd and Pb, Sb and, as Traces, As, Bi, Sn and Ge in (Mainly) Hydrothermal Sulphide Paragenesis Microscopic and X-ray diffraction studies by Augustithis and Vgenopoulos (1982) determined a complex hydrothermal paragenesis associated with mylonitized muscovite-gneiss in Ragada, Rhodope, northern Greece. The main gangue mineral is dolomite which as veinlets invades the mylonitized gneiss. The main ore minerals determined ore microscopically, by X-ray diffraction and by microprobe scanning are sphalerite, wurtzite, galena, chalcopyrite, tetrahedrite (fahlore) and associated with sphalerite/wurtzite and hawleyite. Ore microscopically other sulphides are paragenetically associated with the sphalerite/wurtzite. Sphalerite is associated with galena and tetrahedrite. In contradistinction to the idiomorphic pyrite, sphaeroids of pyrite (perhaps gel-pyrite) are associated with tetrahedrite, galena and chalcopyrite. Tetrahedrite, often associated with chalcopyrite, is a common and abundant mineral phase in the sulphides paragenesis of Ragada. Quantitative X-ray fluorescence spectroanalyses show that isolated sphalerite/wurtzite crystal aggregates contain 1.8% maximum of cadmium. As a corollary, semi-quantitative X-ray diffraction studies show 2% maximum of hawleyite in sphalerite/wurtzite. The differentiation of wurtzite from sphalerite was only possible by X-ray analysis. The association of hawleyite (usually greenockite is present) with sphalerite/wurtzite can be understood by considering the element geochemistry of Cd and Zn

(they are related as subgroup elements of the II family of the periodic table). Furthermore, the common occurrence in this paragenesis of the elements Cu, Ag, Zn and Cd is again explicable on the basis of the relationships of these elements in accordance with the periodic table. Cu and Zn are related as next to each other subgroup elements of the I and II family, and similarly Ag and Cd (horizontal relationship of the subgroup elements). In this connection, the presence of Ag is understandable due to its interrelationship with both Cu and Cd. As the composition of the main sulphides (i. e., chalcopyrite, pyrite, sphalerite, tetrahedrite and hawleyite) of the Ragada paragenesis shows, Cu, Zn, Ag and Cd are the main metallic elements of this paragenesis, in addition to Fe (which is also related to Cu in accordance with the empirical "laws" of the periodic table). In contradistinction to the group Cu, Zn, Ag and Cd, there are also Pb, Sb and as traces, As, Bi, Sn and Ge present in the paragenesis of Ragada. Pb is present in galena and Sb in the tetrahedrite, the other elements occur as traces in the tetrahedrite. The elements As, Sb and Bi are related as main elements of the V family of the periodic system. The elements Ge, Sn and Pb are similarly interrelated as main elements of the IV family of the periodic system. Furthermore, Pb, Sb, As, Bi, Sn and Ge are interrelated in accordance with the empirical "laws" of the periodic system. As the element geochemistry of the metallic elements of the Ragada paragenesis shows, two independent groups of elements are present, the members of each group being interrelated in accordance with the empirical "laws" of the periodic system.

229

Chapter 50

(a) U-Parageneses

Hydrothermal and Pegmatitic Element Segregation to Form U-Parageneses

(General)

When considering element segregations to form mineral associations-paragenesis the hydrothermal and pegmatitic U-parageneses are chosen, because in any case, there are not only most complex mineral associations present but also a wide range of minerals. It should be emphasized that the U-hydrothermal parageneses not only represent typical vein hydrothermal deposits but that they also represent the most complex cases possible. Similarly the U-pegmatitic paragenesis (chosen as an example to illustrate the common segregation of rare elements) also exemplifies typical and most complex pegmatitic parageneses. Furthermore, a geochemical comparison of these two cases illustrates that diverse elements are segregated under hydrothermal and pegmatitic conditions, while others are common in both.

(a) Hydrothermal U02

Parageneses

General Considerations Regarding the Conception of Paragenesis Broadly-speaking, the term paragenesis is used to signify genetic associations of minerals (see Chapters 1 and 48). It should be kept in mind, however, that a mineral association can be the result either of a single period of crystallization or of more. (Examples illustrating parageneses of multiple periods of crystallization are quoted in Augustithis, 1964). Tishkin et al. (1958), describing types of uranium parageneses from the USSR, use the following terms: (i) Period of mineralization - a long period of time during which a complex of asynchronous hydrothermal formations connected with a definite tectonomagmatic cycle was produced. (ii) Stage of mineralization - a space of time within the period of mineralization during which hydrothermal formations were produced (veins, nests, zones of dissemination, metasomatic bodies, etc.). Hydrothermal formations formed during the same stage of mineralization have more or less similar compositions, and are chronologically separated from other stages by tectonic movements. 230

(iii) Association - a group of minerals in the same place, irrespective of their genesis or age. (iv) Paragenetic association - a group of minerals, as a rule formed together and at the same time, deposited from only one portion of single solution. (v) Generation - asynchronous separation of a single mineral, deposited during the same stage of mineralization. In contradistinction to the above terms, however, other authors have introduced other terms describing the periodic formation of minerals in other uranium parageneses. It should also be pointed out that the Russian term stage of mineralization corresponds to the following terms used by the respective authors: (i) Formation - term used by Leutwein to describe the mineralization stage of the Erzgebirge uranium paragenetic associations. (ii) Type - term used by Kidd and Haycock to describe the mineralization stages of the Great Bear Lake deposits. These authors have also introduced the term stage which corresponds rather to the Russian term period of mineralization. (iii) Phase - term used by Derriks and Oosterbosch to describe the mineralization stages of the Shinkolobwe, Swambo and Kalongwe deposits. In the present work, however, the term paragenesis is not used to describe every type of mineral association in rocks but cases of metallic minerals in hydrothermal and pegmatitic occurrences. Consequently, the paragenetic association of minerals (provided that the minerals are genetically interrelated) actually represents a geochemical interrelationship between elements segregated at one or more stages of mineralization. Irrespective of whether the minerals belong to one or more of these stages, two groups of elements forming the metallic minerals of hydrothermal pitchblende parageneses have been recognized. In the ideal case, recognized by considering the metallic elements present in a number of "pitchblende occurrences" of world-wide distribution (see Appendix A - Geochemical and ore microscopic studies of hydrothermal and pegmatitic uranium parageneses, Augustithis, 1964), the following elements are included in each group: Group A (mainly crust-derivative elements) = U, (Th), Pb, Bi, As, Sb, Se, (Sn and Mo); Group Β (mainly mantle-derivative elements) = Fe, Co, Ni, Cu, Zn, Au and Ag.

Nevertheless, the crystallization of minerals in a paragenesis depends on the physico-chemical conditions, i. e., temperature, pressure and crystallochemical factors (as most extensively shown in the work of Goldschmidt, 1954). The relationship between crystal structure and chemical composition can be summarized depending on: (i) The relative proportions of the various kinds of atoms (or ions) in a chemical formula. (ii) The relative sizes of the various kinds of particles, i. e., atoms (or ions) in a crystal. Geochemical Consideration of Pitchblende Paragenesis A number of well-known pitchblende occurrences have been ore-microscopically studied. On the basis of the minerals identified, a list of the elements present has been prepared, particularly in order to show the main metallic elements. The elements listed occur in sufficiently large quantities to contribute to or form ore minerals. The elements present as traces only have been omitted. A study of listed elements (see Augustithis, 1964) shows that they clearly belong to two groups: group A, consisting of U, Th, Pb, Bi, As, Sb, Se, Sn and (Mo) the three last ones merit special consideration - and group B, consisting of Fe, (Μη), Ni, Co, Cu, Ag, Au andZn. It can be seen from group A that As, Sb and Bi are interrelated as main elements of family V of the periodic system. In contrast Se, Sn and (Mo) are only in cases typical for hydrothermal pitchblende paragenesis. The elements of group Β also show certain relationships in accordance with the periodic system. Fe, Co and Ni are elements of the subgroup of family VIII. Cu is related to Ni as the element next to it; Zn is similarly related to Cu. The elements Cu, Ag and Au are related as homologues. Variations in the number and the identities of the elements present have been observed in different pitchblende pangeneses. Also, the minerals present can vary from one occurrence to the other. A selected number of hydrothermal pitchblende occurrences illustrates these points. Samples of pitchblende in association with other ore minerals from Joachimsthal (Jächymov, Bohemia) (Zwittermühl, Segen Gottes) show the following: Pitchblende U 0 2 Galena PbS Chalcopyrite CuFeS2 Smaltite (Ni, Co)As3 Pyrargyrite Ag3SbS3 Niccolite NiAs Bismuthinite Bi 2 S 3 Marcasite FeS2 Arsenopyrite FeAsS Sphalerite ZnS Polybasite Ag 16 Sb 2 S n The mineral building elements of this particular pitchblende paragenesis can be grouped into two units, group A: U, (Th), Pb, Bi, Sb and As; and group B: Fe,

Cu, Co, Ni, Zn and Ag. Comparing the elements present in this case with ideal units, it can be seen that almost all the elements in these units are also present here, except for Se, Sn and Mo. Comparable pitchblende samples from Schneeberg, Saxony, contain the following ore minerals: Pitchblende U0 2 Niccolite NiAs Smaltite (Ni, Co)As3 Pararammelsbergite NiAs2 Rammelsbergite NiAs2 Argentiferous minerals Safflorite CoAs2 Pyrite FeS 2 Tetrahedrite Cu3SbS3 Lepidocrocite F e ^ - H j O Emplectite CuBiS2 Bi minerals Maucherite Ni4As3 Cosalite Pb 2 Bi 2 S 5 On the basis of the mineral building elements the following groups are recognized: group A: U, (Th), Pb, Bi, Sb and As; and group B: Fe, Ni, Co, Cu and Ag. If the mineral building elements of the above mentioned occurrences are compared, it is obvious that despite differences in the minerals present, the two occurrences contain almost the same elements. Mo, Se and Sn as Mineral Building Elements in Hydrothermal Pitchblende Paragenesis I. Mo in pitchblende paragenesis The presence of Mo minerals in pitchblende paragenesis is somewhat problematic. Overall, their presence is not characteristic, although occurrences with molybdenite and gel pitchblende are known to exist. Ore microscopic studies by Ramdohr (1960) and studies on the metallogenesis of "Kupferschiefer von Mansfeld" by Schüller (1958) showed that molybdenite is present in the hydrothermal pitchblende paragenesis of Rucken Mansfeld, Germany. The following minerals were identified: Pitchblende U0 2 Native Bi Niccolite NiAs Molybdenite MoS2 Maucherite Ni4As3 Bornite Cu5FeS4 Rammelsbergite NiAs2 Chalcocite Cu2S Safflorite CoAs, Haematite Fe 2 0 3 Barite BaSO, Cobaltite (Co,Fe)AsS Chalcocite with Ag Cu 2 S Galena PbS Sphalerite ZnS Betechtinite Pb2(Cu,Fe)S15 The mineral building elements can be divided into group A, consisting of U, (?Th), Bi, As, Pb and Mo; and group B, containing Fe, Ν, Cu, Co, Zn and ?Ag. The presence of Mo, Ag and Ba can be attributed to lateral segregation. Moreover, the geochemical interrelationships of uranium and molybdenum have been studied by a number of workers who tentatively suggest that the paragenetic association is probably governed by chemical relations. The following are certain selected cases of paragenetic associations, i. e., geochemical and mineralogical interrelationships of pitchblende with Mo-containing minerals: 231

(i) According to Tishkin et al. (1958), pitchblende and molybdenum minerals occur paragenetically within Russia (former Soviet Union). They quote one instance which shows that quartz, pitchblende, molybdenite, galenite and carbonates are paragenetically associated and belong to the same stage of mineralization. (ii) Kopchenova and Skvortsova (1958), also described the paramagnetic association of pitchblende with molybdenite. According to these authors, the two minerals often form original structures of joint isolation, although the properties of the molybdenite are noticeably abnormal. It forms fine or cryptoscally aggregates which macroscopically have the form of a black film or crust with metallic lustre. (Anisotropy and no double-reflection are characteristic for this variety.) Moreover, its colloform appearance and its association with gel pitchblende suggests a simultaneous precipitation by coagulation. The authors also suggest that this association is due to recrystallization from an originally amorphous gel pitchblende and jordisite (a powdery colloidal molybdenum sulphide). Another case of pitchblende associated with molybdenum minerals is quoted by Kerr et al. (1952). Jordisite and isemannite are associated with oxidized uranium minerals and uraninite. Furthermore, the geochemical interrelationships of uranium and molybdenum have been studied by a number of researchers who tentatively suggest that the paragenetic association is probably governed by chemical relations. Concerning the distribution of Mo and its paragenetic associations, see also Chapters 56 and 57. (iii) Another case of uranium-molybdenum association is advanced by Nekrasova (1958, quoting Davidson), according to whom concentration of molybdenum, thorium, rare earths and other elements occur together in some coals of the United States. Nevertheless, such a paragenetic association should not be seen as more than tentatively, owing to the problematic complexity of the presence of uranium in coals. (iv) A further indication of the geochemical paragenetic association of U and Mo is the presence of the latter in small quantities in pitchblende, as chemical analyses from Johanngeorgenstadt show, i. e., V 2 0 J ( MO03, W 0 3 = 0.75%. (v) Additional cases of geochemical interrelation of uranium with Mo are indicated by the formation of minerals containing both elements as products of the primary pitchblende-molybdenite association after its subjection to variable chemical milieu of mineral weathering. The following cases are quoted from different workers: (i) Due to the disintegration of uraninite (pitchblende) and molybdenite colloform associations, amorphous substances known as "blacks" exist with a wide range of uranium-molybdenum. These substances are the intermediate products of the disintegration of 232

primary ore deposits. In the oxidation zone they are unstable and can easily be replaced by other minerals, e. g., uranium micas and wulfenite. Similarly, an uranium containing powellite is formed during the oxidation of uranium-molybdenum blacks in the hypergene zones of molybdenite-pitchblende deposits. (ii) Rudnitskaya (1958) showed that a calcium-uranium molybdate (Ca(U0 2 )3(M00 4 ) 3 (0H) 2 -8H20) exists in the lower part of the oxidation zone of hydrothermal U-Mo veinlet deposits. It is developed in those places where uranium-bearing chalcedony veins, almost devoid of sulphides except for molybdenite, occur. Rekharskii (1959), discussing the metasomatic processes and geological conditions of the association of Mo and U, points out the interrelationship of these elements in accordance with the periodic table. "The problems of the interrelation of Mo and U minerals are more interesting because the elements of these minerals belong to the same subgroup of the periodic table and therefore possess many similar characteristics." He also points out that tungsten similarly belongs to the same subgroup, and forms compounds that occur together with Mo and U and in addition is paragenetically associated with them. Moreover, X-ray fluorescence spectroanalysis of uraninites often show tungsten present in small quantities. Molybdenite and pitchblende occur as fine-grained aggregates or colloform masses. In some cases, molybdenite passes gradually into pitchblende through a "mixture" of molybdenite-pitchblende. Their association probably occurs from crystallization of simultaneously coagulated gels. Furthermore, the geological interrelations of uranium and molybdenum have been studied by a number of workers who tentatively suggest that the paragenetic association is probably governed by chemical relations. The association of Mo and Rh is discussed in the geochemical work of Rekharskii and their genetic association is pointed out by Schüller and Myers et al. (1960), who by simple water extraction and concentration technique prior to spectrographic analyses determined the approximate quantities of Rh associated with Mo. According to Rekharskii20, the paragenetic association of Mo and Rh results from the properties of their salts.

20

It is well-known that in Mendelejeffs table (1869) the basicity of the elements decreases from the first to the last group (excluding the noble gases) and increases within each subgroup with atomic weight of the element. In many cases, therefore, the elements lying along the diagonals may form salts soluble under similar conditions. An example of this is the existence of rhenium, potassium and sodium thiosalts similar to those of molybdenum which, upon decomposition and removal of a part of the sulfur, yield rhenium sulfide.

II. Se in pitchblende paragenesis In addition to Mo and Sn, which cannot be regarded as typical elements of hydrothermal pitchblende paragenesis with certainty, Se has also been recorded, e. g., Se (and Te 21 ) occur in the Shinkolobwe paragenesis. It is difficult to be sure whether Se is a typical element. Often the association of Se minerals with pitchblende is attributed to later Se mobilization and to lateral segregation. Since U is extremely "mobile" and palingenic structural associations often occur, the structural relationship between colloform pitchblende and Se minerals is difficult to interpret. Ore microscopic observations indicate the relation of colloform pitchblende to Se minerals. Fig. 713 shows a colloform pitchblende, the central areas of which have been dissolved and subsequently occupied by clausthalite. The later age of the clausthalite is often indicated by its occupying inter-shell layers and cracks in the colloform U0 2 structures (Robinson, 1955). In addition to the association of pitchblende with Se minerals (clausthalite, klockmannite, umangite), association with native Se can exist. Fig. 925 shows pitchblende (half-moon in shape) with native Se and clausthalite. However, where this type of association occurs, is unclear and it is difficult to draw conclusions regarding the age relation. Studies on polished sections from Pinky Fault, Beaverlodge Lake, Saskatchewan, Canada, show the following mineralogical compositions for this pitchblende paragenesis: Pitchblende UO., Bornite Cu5FeS4 Clausthalite PbSe Galena PbS Klockmannite CuSe Berzelianite Cu2Se Native Se Crookesite (Cu,Tl,Ag)2Se CovelliteCuS Limonite Fe 2 0 3 -H 2 0 Umangite Cu3Se2(Cu-Ag) Bornite Cu5FeS4 Se-Linneite R2(Fe,Ni,Co,Cu) Blue digenite Cu9S5 R23(Co,Ni)S4 (selenium is also present) The mineral building elements can be divided into group A: U, (?Th), Pb, (?T1) and Se; and group B: Fe, Ni, Co, Cu and Ag. III. Sn in pitchblende paragenesis Ore microscopic studies of specimens from Grampound Road, St. Stephens, Cornwall, UK, show that the following minerals occur: Pitchblende U0 2 Galena PbS Pyrite FeS2 Cassiterite SnO, Chalcopyrite CuFeS2 Siderite FeC0 3 Chalcocite Cu2S Bornite CusFeS4 Covellite CuS Tetrahedrite Cu3SbS3 The presence of cassiterite is of special interest. The mineral building elements can be divided into group A: U, (Th), Pb, Sb and Sn; and group B: Fe and Cu.

Similar studies of specimens from St. Just, Cornwall, UK, show that cassiterite is again present. The following minerals have been observed: Pitchblende U0 2 Galena PbS Chalcopyrite CuFeS2 Cassiterite Sn0 2 Pyrite FeS2 Bismuthinite Bi2S3 Sphalerite ZnS The mineral building elements can be divided into group A: U, Pb, Bi and Sn; and group B: Fe, Cu and Zn. The Grampound Road cassiterite occurs partly as veinlets cutting through pitchblende, illustrating a palingenic origin (Fig. 809)22. Uraninite and U-Containing geneses

Pegmatitic

Mineral

Pan-

In contrast to the hydrothermal pitchblende paragenesis where, as already mentioned, the mineral building elements can be divided into two groups, group A: U, Th, Pb, Bi, As, Sb, (Sn, Se, Mo), and group B: Fe, Ni, Co, Cu, Zn, Ag and Au. The pegmatitic uranium paragenesis is mainly characterized by the presence of the following elements: U, Th, Y, Zr, Hf, Nb, Ta, Ti and rare earths. [The elements Sn, Mo and W also occur in granites and pegmatites and belong to the "granitophile" group of elements, see Chapters 56 and 57.] The abundance of Th, Y, Zr, Hf, Nb, Ta, Ti and rare earths in the pegmatitic uranium paragenesis are illustrated by the following factors: (i) The mineral building elements of pegmatitic uranium parageneses, some of them are mineralogically characterized (or often characterized) by the presence of zircon, monazite, Ta - Nb, and rare earth minerals. (ii) The composition of some of the uranium primary minerals, present in pegmatitic occurrences, examples of which are: Aeschynite The type formula is AB 2 0 6 priorite series (A = Ce, Y, Ca, Fe2, Th and minor Mg, Mn, Pb, U, Sn, Zr; Β = Ti, Nb, minor Ta) Betafite (U,Ca)(Nb,Ta,Ti) 3 0 9 nH20(?) Blomstradite Synonymous with priorite Davidite General formula possibly: AB 3 (0,0H) 7 (A = Fe2, rare earths, U, Ca, Na, Zr, Th; Β = Ti4, Fe3, U, (V3, Cr3)) Delorenzite (Y,U,Fe2)(Ti,Sn)3Og(?) Euxenite A complex oxide of the type polycrase series AB 2 0 6 (A = Y, Ce, U, Th, minor Ca, Mg, Μη, Fe2; Β = Ti, Nb, Ta, minor Fe3) Fergusonite The type formula for the series is formanite series AB0 4 (A = Y, Er, minor Ce, U4,

22

21

Se and Te often occur together.

Another interesting case is the presence of Sn0 2 with uraninite of pegmatitic origin (Augustithis, 1964). 233

Hjelmite

Ishikawaite Monazite Polymignite

Pyrochlore microlite series Samarskite yttrotantalite series Thorianite Yttrocrasite

Th, Zr, Fe 2 and Ca; Β = Nb (fergusonite) and Ta (formanite)) Complex oxide of Ta, subordinate Nb, also containing Y, Fe 2 , U 4 , Ca, Μη, and Sn (U 4 , Fe 2 , Y, etc.)(Nb, T a ) 0 4 (Ce,La)P0 4 The type formula is A B 0 4 (A = Ca, Fe 2 , Y, Ce, Zr, minor Th; Β = Nb, Ti, Ta, Fe 3 ) NaCaNb 2 0 6 F - pyrochlore (Na,Ca) 2 (Ta,Nb) 2 0 6 (0,0H,F) microlite The probable formula is A B 2 0 6 (A = Y group, some Ce group, U, Th, Fe 2 , Ca; Β = Nb, Ta, Fe 3 , minor Ti, Sn, W) (Th, U 4 ) 0 2 (Y,Th,U,Ca) 2 Ti 4 0 1 1

The mineral building elements of these minerals are U, Th, Y, Zr, Hf, Ta, Nb and Ti and the rare earths. Fe is also present, and often Mg, Mn and Ca; to a lesser extent P, Sn and W. (iii) A study of selected occurrences of davidite, the [U, Nb, Y] mineral from Harrar, Ethiopia, and pegmatitic uraninite (Augustithis, 1964) illustrates that the mineral building elements are similarly U, Th, Y, Zr, Hf, Nb, Ta and Ti and the rare earths. (iv) Comparative geochemical studies of pitchblende of hydrothermal origin and of uraninite of pegmatitic origin illustrate that Th, Y, Zr, Hf, rare earths, Ti possess the greater importance in the pegmatitic uraninite (Augustithis, 1964). (v) Geochemical studies of malacon (zircon) by Kluth and Ottemann (in: Augustithis, 1964) show uranium as an accessory in pegmatitic occurrences (see Augustithis, 1964 - Section A.3(b), pp. 16-17). Nevertheless, it should be emphasized that the elements present in the pegmatitic uranium paragenesis show certain interrelationships in accordance with the periodic system. The elements Th, (Hf), Zr and Ti are homologous (all being elements of the subgroup IV); the elements Nb and Y are interrelated as being elements next to Zr in the periodic table and the rare earths as next to Hf (Hf and Zr possess special bonds of interrelationship). These interrelations strongly suggest that elements present in pegmatitic uranium paragenesis show certain relationships and that they differ from the elements forming hydrothermal pitchblende paragenesis. Whereas in the case of hydrothermal pitchblende paragenesis the elements of the A and Β groups are most common, in the pegmatitic paragenesis of uranium the special group (C), U, Th, Y, Zr (Hf), Nb, Ta and Ti and the rare earths predominates (although certain exceptions and special cases are considered, see Augustithis, 1964). 234

Chemical and Geochemical Considerations of the Metallic Elements Present in Primary Uranium Parageneses The relative distribution of elements in the earth's crust (Tables Via and VIb - based on Goldschmidt, 1954) shows that some are present in minute quantities. Nevertheless, segregation of these elements often occurs and forms special concentrations, ore deposits are examples. Regarding the abundance of uranium, and also certain tendencies in the abundance, the following relationships are summarized by Goldschmidt in his

Verteilungsgesetze der Elemente.

(i) Von Hevesy et al. (1930) determined 4 gr uranium per ton in eruptive rocks. (ii) In chondritic meteorites he estimated 0.12 gr Ra per ton. (iii) In iron meteorites, Paneth and coworkers (1930) determined 0.03 gr Ra per ton, corresponding to 0.1 gr uranium per ton. In accordance with these estimates the distribution quotient between silicates and nickel-iron is 4:1, so that uranium is definitely a lithophile element. The elements present in uranium paragenesis show that relatively rare elements are segregated, with uranium forming paragenetic associations. The elements of groups A and Β of hydrothermal paragenesis, already described, show the following abundance values according to Goldschmidt (1954): Table Via Group A (parts per million) u Th Pb As Bi Sb Mo Se Te Sn

4 11.5 16 5 0.2 1 2.3 0.09 0.0018 40

Group Β (parts per million) Fe Ni Co Cu Zn Ag Au

50000 100 40 70 80 0.02 0.001

Similarly for the special group (group C) of elements characteristic of pegmatitic uranium paragenesis, the corresponding values are: Table VIb Group C (parts per million) u Th Y Zr Hf Nb Ti Ta Ce (Fe

4 11.5 28.1 220 4.5 20 4400 2.1 41.6 50000)

Consequently it is more than coincidence that the rare elements occur together in mineral paragenetic associations of worldwide distribution. Between these elements, it appears to exist a chemical-geochemical affinity and some kind of interrelationship. It can be best understood by considering the periodic system and its empirical laws which determine the broad lines of the affinities of all elements. In this respect it should be emphasized that the periodic system is a scheme of element classification as proposed by Ihde (1969). A Geochemical Comparison of Hydtothermal and Pegmatitic Uranium Parageneses Comparing the hydrothermal paragenesis consisting of groups A (= U, (Th), Pb, Bi, As, Sb, (Se, Sn, [Mo?])) and Β (= Fe, (Μη), Co, Ni, Ag, Zn, Cu, Au) with the pegmatitic uranium paragenesis which mainly consists of the group C (= U, Th, Zr, Hf, Y, Nb, Ta, Ti and REE) certain geochemical tendencies become apparent. As already mentioned, Bi, Sb, As are homologous as elements of the main group of family V and are therefore interrelated. Also Mo, W and U are homologous and belong to the subgroup of family VI. It is therefore not surprising that in hydrothermal parageneses23 they occur together. The elements comprising group C show interrelationships. Th, Hf, Zr, Ti are homologous and belong to the subgroup of family IV. Y and Nb are interrelated to Zr as elements next to it in the periodic table, and are also related to Hf. Ta is related both to Nb (homologous) and to Hf as next to it. The rare earths are related to Hf as elements next to it. Thus, it can be suggested that the elements which comprise group C and which are often the main mineral building ele23

Mo, W and U also occur in pegmatitic parageneses.

ments in pegmatitic uranium parageneses are interrelated in accordance with the empirical "laws" of the periodic system. Moreover, on the basis of the geochemical studies (Augustithis, 1964), it has been determined by both chemical and X-ray fluorescence spectroanalysis that Th, Zr (?Hf), Ti are also present in hydrothermal uranium parageneses in rather restricted quantities, whereas the elements Ta and Nb are restricted to pegmatitic uranium parageneses only. Thus, comparison of hydrothermal and pegmatitic uranium parageneses shows that Bi, Sb, As are characteristic elements of the hydrothermal paragenesis whereas Nb and Ta of the pegmatitic. The geochemical incompatibility can best be explained by the following empirical law: the similarity between the main and the subgroup elements of the same family increases strongly from family I to IV and decreases strongly from IV to VIII. The elements Bi, Sb, As belong to the main group and Ta, Nb to the subgroup of family V, in which the similarity between the two groups decreases. The present geochemical studies show that Th is the key element for understanding the uranium paragenesis. According to Dwight (1960), there is also a geochemical interrelation between U and Th. Besides their interrelationship in accordance with the periodic system, U0 2 and Th0 2 form an isomorphous series. Th is the key element for group C (Th, Zr, Hf, Ti, Y, Nb, Ta and the rare earths), and investigations showed that when thorium is in excess, the other elements of the group are also present in an uranium paragenesis. Geochemical studies by Mineev (1959) show that within the pegmatitic uranium minerals characterized by group C, the chemical similarity of the elements within the group also determines the predominance of certain elements over others. Thus, the similarity of the crystallochemical and geochemical properties of uranium and yttrium lead to the conclusion that uranium predominates over thorium in all yttrium complex minerals. Despite this conclusion, which exemplifies an interrelation between elements based on chemical similarities independent of those specified by the periodic system (i. e., the relation between U and Th is regarded as stronger than that between U and Y), overall, the presence of elements characteristic of primary uranium minerals depends on their similarities and interrelations - which as pointed out, depend in turn on similarities as expressed by the periodic system. Conclusions Regarding the Interrelationship Between Metallic Elements in Uranium Parageneses A number of observed cases suggest that, in nature, elements strongly homologous or analogous tend to occur together. This tendency derives from similarities and analogies, and indeed the repetition of analogies, in the atomic structure. Consequently, for element segregation/distribution there is an internal cause, i. e., the 235

segregation/distribution depends on the structure of the atoms themselves- a factor fundamentally important for the understanding of the joint segregation of the less abundant elements in the earth's crust into concentration of deposits, a factor that was not taken into consideration in Goldschmidt's Verteilungsgesetze der Elemente. Whereas the periodic system (table) considers the relationship between chemical properties and one fundamental constant, i. e., atomic weight, Goldschmidt considers as the principle factor for regulating the entrance of the atoms and ions and the distribution of rare elements in the crystalline phases of igneous and metamorphic rocks the size of the atoms and ions and not the atomic weight. The two concepts however are not contradictory or incompatible, they are complimentary. From the elements present in primary uranium parageneses (see Augustithis, 1964), and also from the quoted affinities and interrelationships of all elements in accordance with the empirical laws of the periodic system, the following conclusions can be drawn: (i) In hydrothermal primary uranium parageneses mineral building elements can be divided into group A = U, (Th), Pb, As, Sb, Bi, Se, Mo, Sn; and group Β = Fe, Ni, Co, Cu, Zn, Ag, Au. (1) In group A only the elements As, Bi and Sb are interrelated as homologues, so that the recognition and acceptance of this group as a whole is no more than tentative, i. e., no strict interrelationships in accordance with the periodic system exist. (2) The elements of group Β show a stronger interrelationship in accordance with the periodic system. Nevertheless, their affinities are not as great as theoretically possible since other elements not present in the group show stronger bonds of affinity with elements which are present in the group.

(ii) More than one unit of element assemblage can occur in paragenesis, i. e., the A and Β groups in the case of uranium hydrothermal paragenesis. In such cases no affinity or interrelationship exists between the groups24. (iii) In primary pegmatitic uranium paragenesis as a whole another group of elements is more common, consisting of U, Th, Zr, Y, Nb, Ta, Ti (Sn, Mo, W) and the rare earths. These elements show interrelationships in accordance with the periodic system. Despite variations and transitions between the composition of hydrothermal and pegmatitic uraninites (Augustithis, 1964), and also despite gradations and transitions between hydrothermal and pegmatitic parageneses as described by Robinson (1955), the pegmatitic uranium paragenesis is characterized by elements other than the hydrothermal. (iv) The elements present in the above mentioned groups (A, Β and the special group C) do not strictly follow the rule that elements with the strongest affinities and relationships are always paragenetically associated, i. e., occur as mineral-building elements in the paragenesis. Elements with greater affinities (as stated by the empirical "laws" of the periodic system) than those of elements included in a group may be absent. (v) The empirical "laws" of element affinities and interrelationships should not be considered as rules governing the segregation/distribution of elements forming parageneses. They just serve as an explanation of relationships existing among elements of the three groups recognized. Nevertheless the fact that uranium parageneses of worldwide distribution show similarities in the elements comprising their metallic minerals can be understood to some extent from the point of view of the chemical and geochemical affinities of these elements based on the periodic system.

24

On this occasion, it should be pointed out that group A elements are mainly crust derivatives and the elements of group Β are mantle derivatives (see Chapters 1, 39, 48, 49); thus, a multi-source derivation is suggested.

236

Chapter 51

Superimposed Paragenesis (Element Segregation/Distribution Processes)

In contradistinction to the U pegmatitic paragenesis already described (see page 232 and Augustithis, 1964), a complex of parageneses in the apogranite of Abu Dabbab, Eastern Desert, Egypt, has been described by Augustithis (1982). In addition to typical elements of the U paragenesis (Nb, Y, Ta, Zr), also elements belonging to groups A and Β (Μη, Fe, Cu, Zn, Ni, Cd, (Sb), Ag, (Sn)) of the hydrothermal U paragenesis (see page 229) are present. Whereas the elements Nb, Y, Ta, Zr, (Sn) belong to the granitization formation of the apogranite, the elements Mn, Fe, Cu, Zn, Ni, Cd, Ag, (Sb), (Te) and (Sn) belong to a much later introduction "by hydrothermal" solutions derived by leaching from buried basalts from the adjacent intercontinental rift formation of the Red Sea. In this connection it should be mentioned that of the "hydrothermally" introduced elements, the elements Mn, Fe, Cu, Zn, Ni, Cd, Ag correspond to group Β of the hydrothermal paragenesis (see page 230) and as mentioned, are derivatives by leaching from basalts, which in turn might be remobilized from mantle or the protolyte layer. Xu Keqin et al. (1982) described granitoids from southeastern China where two distinct types of metallogeny are characteristic, namely granitoids with Fe, Cr, Ni, Co, Zn, Cu and (Pb), and granitoids with Sn, W, Mo, Nb, Ta and U. The case of the Abu Dabbab apogranites is interesting since there are both types present in the same apogranitic bodies. As mentioned, the elements Nb, Y, Ta, Zr belong to the pegmatitic-granitic type of paragenesis related to the apogranite formation, whereas Mn, Fe, Cu, Zn, Ni, Cd, Sb, Ag and Te are superimposed derivative from the leaching of basalt buried in the Red Sea. Interesting is the presence of mineral phases present in the Abu Dabbab apogranite which is the result of reaction of the two parageneses. Extracts from a paper by Augustithis (1982), entitled "Textural and mineralogical studies of the apogranites of Abu Dabbab, Eastern Desert, Egypt" are quoted as a synopsis, presenting the mineralogy and geochemistry (parageneses) of these apogranites and the problems involved. "The apogranitic bodies in the Abu Dabbab area of the Eastern Desert, Egypt, are concordant and discordant with the gneisses and contain small concentrations

of Nb, Y, Ta, Zr as well as Cu, Zn, (Sn), Ni, Cd, (Sb), Ag and (Te) 25 Metamorphic-metasomatic interlocking quartz and feldspar intergrowths, as well as idioblastic quartz and poikiloblastic orthoclase support the transformist's explanation for the origin of these apogranites. The Nb, Y, Ta and Zr present are a part of the typical "pegmatitic group of elements" in contrast to the Mn, Fe, Cu, Zn, Ni, Cd, Ag, (Sn), (Sb) and (Te) which are characteristic for the "superimposed" intercontinental rift mobilizations (deposits) typical of the Red Sea Rift." Often associated with pegmatitic and apogranitic bodies Nb, Y and Ta concentrations have been reported by Strunz (1961), Augustithis (1964), Augustithis et al. (1974), Odikadze (1967), and Kalinin and Goldin (1967). Considering the concentration of Nb and Ta in pegmatitic paragenesis in addition to Nb and Ta, the elements U, Th, Zr, Y, Ti and the REE commonly are geochemically associated. Augustithis (1964), studying the geochemical concept of paragenesis, regards the group of elements U, Th, Zr, Y, Ta, Nb and Ti and the REE as characteristic for the pegmatitic paragenesis. The joint segregation and association of these elements are furthermore attributed to their interrelationship in accordance with the empirical "laws" of the periodic system. These interrelationships are extensively discussed in a monograph by the author (1964), "Geochemical and ore microscopic studies of hydrothermal and pegmatitic primary uranium parageneses". There it is stated: "A number of observed cases suggest that, in nature, elements strongly homologous or analogous tend to occur together. This tendency derives from similarities and analogies, and indeed the repetition of analogies, in the atomic structure. Consequently, for element segregation/common distribution there is an internal cause, i. e., the segregation/distribution depends on the structure of the atoms themselves - a factor fundamentally important for the understanding of the joint seg-

The elements Sn, Sb, Te actually belong to a special group of interrelated elements consisting of Sn, As, Sb, Se, Te and In (see page 241) and are most likely derivatives of the crust or the lower crust (see also Chapter 39). 237

regation of the less abundant elements in the earth's crust into concentration of deposits." Furthermore, considering that the Ta and Nb concentration in the apogranites of the Man'Khambo granitegranodiorite massif (Southern Pechora region, Ural Mountains) is attributed by Kalinin and Goldin (1967) to the presence of interspersed furgusonite (type formula AB0 4 , A = Y, Er, minor Ce, U4, Th, Zr, Fe 2 and Ca; Β = Nb (furgusonite) and Ta (formanite)) and samarskite (AB 2 0 6 , A = Y group, some Ce group, U, Th, Fe 2 , Ca; B= Nb, Ta, Fe 3 , minor Ti, Sn, W), it is evident that an analogy and correspondence exists, geochemically, between the pegmatitic euxenite (AB 2 0 6 , A = Y, Ce, U, Th, minor Ca, Mg, Μη, Fe2; Β = Ti, Nb, Ta and minor Fe 3 ) paragenesis of the graphic quartz/K-feldspar containing pegmatites of Harrar, Ethiopia (Augustithis, 1964, Augustithis et al., 1974), and the apogranitic parageneses in which, as mentioned, fergusonite and samarskite occur. Preliminary trace element studies (semi-quantitative X-ray fluorescence spectroanalysis) by Vgenopoulos (in: Augustithis, 1982) and previous geochemical investigations by Russian researchers show that the Abu Dabbab apogranites besides the elements Ta, Nb, Zr, Y, also contain in trace quantities Zn, Cu, Ni, Cd, Ag, Sb, Sn and Te, elements characteristic of the hydrothermal paragenesis (see Augustithis, 1964), which are jointly segregated according their interrelationships in agreement with the empirical laws of the periodic system (see Augustithis and Vgenopoulos, 1981). The coexistence of these two different groups of elements in the apogranites of Abu Dabbab renders more complex the understanding of their metallogenesis. The presence of the group Nb, Y, Ta, Ti, Zr, (HO, U and Th can be understood by the availability of these elements in the initial sedimentogenic-metamorphic granitization environment (responsible for the genesis of the apogranites) and on their relationship according the empirical laws of the periodic system. In addition, ore microscopic and microprobe studies show interspersed small ore mineral grains within the apogranite. Spot-analysis of these grains by E. Mposkos (in: Augustithis, 1982) showed Ta, Nb and (Mn).

As can be seen in Fig. 926, a leaching halo" surrounds these metamictic grains, to similar dispersion (diffusion) haloes Augustithis (1964) and Augustithis et al.

238

"dispersion comparable studied by (1974) and

owing to the differential leaching of elements and their migration in the pegmatite, Nb and U have been determined. In contrast, the group of elements Mn, (Fe), Zn, Cu, Ni, Cd, Sn, Sb, Ag and Te in the Abu Dabbab apogranites is understandable as elements representing a "superimposed" paragenesis (as a geochemical rather than mineralogical concept, sulphides are also ore microscopically observed) which is characteristic of the intercontinental rift type of deposits (very abundant in the adjacent areas of the Red Sea Rift, see Tooms, 1976) and is derived by leaching of metallic elements from buried intercontinental rift basalts. The special significance of Mn both as late impregnations in the Abu Dabbab apogranites and as independent manganese deposits in the region of the Eastern Desert support the hypothesis of "superimposed paragenesis, geochemical concept" for the elements Mn, (Fe) Zn, Cu, Ni, Cd, Sn, Sb, Ag and Te in addition to the true apomagmatic paragenesis represented by the elements Nb, Y, Ta, Ti, Zr, (Hf), U and Th and their corresponding mineralization. As mentioned, in addition to the true apogranitic paragenesis represented by Ta-Nb minerals, zircons and perhaps by cassiterite and the "superimposed paragenesis" Mn (Fe), Zn, Cu, Ni, Cd, Sn, Sb, Ag and Te which is represented mainly by ramsdellite (Mn0 2 ), some sulphides (?) and perhaps other minerals of the group; X-ray diffraction analyses by Vgenopoulos determined the presence of nickmanite (MnSn(OH)3), spiroffite (MnTe 3 0 8 ) and manganotantalite (MnTaO e ). It is possible that some of these minerals and particularly the manganotantalite represent a "reactive mineral phase" of the tantalium of the true apogranitic paragenesis and of the "superimposed paragenesis". As a corollary to this "reactive phase" (manganotantalite) is the presence of some Mn (determined by microprobe spot-analysis) in the Ta-Nb interspersed grains of the apogranitic paragenesis. In conclusion, it should be pointed out that regarding the elements W and particularly tin, doubts are expressed whether they belong to the element group derived by leaching of intercontinental rift basalts, since, as well-known, Sn and W are granitophile elements. However, the fact that nickmanite also occurs in the Abu Dabbab apogranites, makes the problem even more complex. The concept of superimposed and reactive mineral paragenesis is further considered in Chapter 25.

Chapter 52

Ti, V, Cr - Their Interrelationships and Antipathies

When considering the common segregation of elements in ore concentrations in addition to the interrelationship of the elements in accordance with the empirical "laws" of the periodic system, as mentioned, the behaviour and proportions of the compounds in which the element transportation/mobilization/concentration takes place, is just as important. Thus, although the elements Ti, V and Cr are interrelated as subgroup elements next to each other, nevertheless certain geochemical antipathies have been recorded (see Augustithis, 1979). Whereas Cr is more abundant in ultrabasics where the MgO/FeO ratio shows a predominance in MgO, the presence of Ti is more marked in ultrabasics in which FeO reaches a certain value of the MgO/FeO ratio. The presence and prevalence of Cr in MgO-rich ultrabasics is understood on the basis of the strong interrelationship of Cr with Ni and Fe in accordance with the empirical "laws" of the periodic system. Furthermore, as Goldschmidt stated in his 'law' of element distribution (1922-1933), Mg can be substituted in the forsterite (olivine) lattice by Ni (since the size of the Mg +2 = 0.78 Ä and that of Ni+2 = 0.78 Ä). In addition, taking into consideration that the elements Cr, Fe and Ni are interrelated in accordance with the empirical "laws" of the periodic system (as subgroup elements in a horizontal line) and the prevalence of Cr depends on the ratio of MgO/FeO (when MgO is prevalent), this is an example of crystallization of minerals depending on the three factors initially mentioned: (i) the interrelationship in accordance with the empirical "laws" of the periodic system, (ii) the ratios of their oxides, and (iii) the relative size of their atoms. In addition to the tendency mentioned for antipathy between Cr and Ti, Willemse (1969) reported an antipathy between Ti and V, considering the geochemistry of the magnetite paragenesis in ultrabasics. According to Willemse, it is suggested that a decrease in oxygen fugacity in magma determined the increase of the titanium content upward in the layer sequence of banded ultrabasics. This case of antipathy between the elements interrelated in accordance with the periodic system, also depends on the behaviour of compounds in which the elements are transported and proved to be decisive in determining the concentration

of elements in certain geoenvironments as it was the case of seams of iron and iron-plugs in ultrabasic complexes. As a corollary to the aforementioned factors which influence the concentration and segregation of the Ti, V and Cr elements, extracts from the monograph "Atlas of the textural patterns of basic and ultrabasic rocks and their significance", Augustithis (1979) are quoted: "Associated with basic and ultrabasic bodies are certainly more or less characteristic ore mineral parageneses and ore deposits. Of particular importance is the presence of minerals of the spinel group. The paragenetic association of spinel groups (of a certain range of the spinel composition) with types of the ultrabasic rocks is of petrogenetic importance". Augustithis (1978) studied the mineralogy and textures of "olivine bombs" in basalts. These bombs are interpreted as mantle pieces. Spinels of the range spinel-chromite-magnetite are in graphic intergrowth with olivines and pyroxenes. The graphic-shaped spinels represent skeleton crystals incompletely developed; however, transition and cases of more idiomorphic developed spinel have been observed. The mineralogical and petrographical studies of the mantle pieces clearly show that whereas the spinel range (Al spinels) predominates, nevertheless, chromite or chromite spinels are also important mineral constituents of the mantle, see Chapter 4. As a corollary to the importance and significance of chrome in the mantle is the abundance of chrome as a "trace element" in the different mineral constituents of mantle pieces (olivine bombs in basalt). As pointed out in Chapter 4, the tectonic (diapiric) mobilization 26 of mantle, rich in olivine (forsterite) and chrome-spinels could provide a satisfactory explanation of dunitic ultrabasic complexes such as Vourinos and Troodos. Of particular importance for the mantle diapirism is a forsterite-rich portion of the mantle (in the sense that forsterite prevails over the pyroxenes in a mantle portion, dunitic in composition). The relative abundance of forsterite renders greater plastic mobility to the mantle part as it is demonstrated by the deformation 26

or, in accordance to its alternative hypothesis, based on plate tectonics, oceanic crust/upper mantle obduction. 239

appearance and petrofabric analysis of dunites of such complexes (see Chapter 4). Another important factor for the understanding of the paragenetic association of chromite-dunite, i. e., the association of chrome spinels with forsterite-rich portions of the mantle, could be broadly understood by considering the geochemical interrelations of Mg in the olivines and Ni-Cr 27 (i. e., size of atomic radii Mg +2 and Ni +2 = 0.78 Ä). It is therefore probable that the paragenetic association of the different types of spinels with the ultrabasics depend on MgO/FeO ratios among other factors. Considering the paragenetic concept: spineltype/ultrabasic type, it is clear that distinguishing of the spinel paragenesis is possible, namely that chromespinels and chromites are prevailing in the forsteriterich dunites and magnetite-titanospinels prevailing in ultrabasic and basic rocks where the FeO is above a certain marginal value for the MgO/FeO ratio. Furthermore, these rather empirically proposed trends of preferential association of Cr (and to some extent Ni) in primary distribution with ultrabasics in which the MgO is predominating; and magnetite spinel, Al- and Ti-spinels in a wide range of ultrabasics in which the FeO reached a certain value range can provide a tentative explanation for the preference of Ti in pyroxenites and gabbros rather than dunites. As the geochemical comparison of trace element distribution (Augustithis, 1978) shows, Ti is virtually limited in abundance in mantle fragment (olivine bomb) whereas relatively high values are noticeable in the basalt enclosing the olivine bomb. The distribution of Ti in basic and ultrabasic rocks is also of interest, especially in the attempt to explain the virtual absence of titano-magnetites, Ti spinels in dunites, which in contrast, are very abundant in the gabbroic rocks (see Chapter 4). Considering the distribution of Ti in basalts, the following should be noted: (i) The range of titania content in terrestrial basalts varies within the following limits: 1.9-6.43% Ti0 2 . (ii) The terrestrial basalts are relatively poorer than the equivalent lunar basalts. (iii) There is a difference in the titania content between inter-oceanic and circum-oceanic basalt, i. e„ Ti values decrease with supposed depth of basalt derivation. The above geochemical consideration which synoptically reviews the Ti in basics and ultrabasics shows that Ti compared with Cr and Ni is more abundant in those basics and ultrabasics where FeO is above a certain value range and in which A1203 is also more abundant. In contradistinction, Cr and Ni tend to be more abundant in ultrabasics with a high ratio of MgO. In this connection, the following quotation from the work of Kern (1968) is interesting: 27

Ni and Cr are interrelated in accordance with the periodic system

240

"The FeO and the MgO contents (in chromites) are strictly antipathetic as well as the Fe 2 0 3 and A1203 resp. the Cr 2 0 3 and (A1 2 0 3 + Fe 2 0 3 ) values. The end members of the spinel group show similar reactions. Ferrite-chromite decreases on account of increasing values of picrochromite. The spinel and magnetite diagrams are contrary in the same way." Nevertheless, in dunitic complexes and intrusions, chromites may be present which show the presence of alumino-spinel and rutile ex-solutions. Fig. 295 shows the presence of rutile ex-solution lamellae oriented parallel to the [111] of the chromite. In addition to the ex-solution lamellae rounded rutile and twinned rutiles may be grain crystallization, included in the chromite from Rodiani, Greece. The presence of Ti and Cr together in the ultrabasic rocks of Rodiani, Greece, (Augustithis, 1960) is of geochemical significance. The Ti and chrome are related as subgroup elements of the periodic system which belong to the same horizontal line (Sc, Ti, V, Cr, Μη, Fe, Co, Ni) and as has been pointed out, despite Ti not being abundant in the dunites (nor in the dunitic mantle) and Ti containing spinels, are virt'ially absent from the dunitic types. In this connection, it is perhaps interesting to note the presence of Al spinels as oriented ex-solution lamellae in the same Rodiani chromite. Fig. 296 shows the coexistence of Al spinel ex-solution lamellae and rutile lamellae both following the [111] of the host chromite (Mg-rich chrome-spinel picrochromite). Similarly, Fig. 297 shows abundant Al spinels in the Rodiani chromite grains. Considering the statement by Kern (1968), it is clear that despite the antipathetic tendency of Cr 2 0 3 and (A1203 + Fe 2 0 3 ) values, Al spinel ex-solutions may be present in chromites. Rarely, chromites or actually idiomorphic chromites of dunitic "pipe" intrusions, show margins of magnetite (see Fig. 577). So far the magnetite associated with chromite is mostly free of ilmenite ex-solutions. It is interesting to note that the preferential paragenetic association of magnetite with ultrabasics "gabbroic in composition" is in accordance with the tendency of magnetite to disappear in olivine-pyroxenites and dunites. As already explained, this could depend on the MgO/FeO ratio in an ultrabasic (see Chapter 4). In this connection, it should be attempted to understand the preferential association of titaniferous minerals with the magnetite paragenesis in ultrabasics rather than with the chromite (see Chapter 20). Considering the geochemistry of the magnetite paragenesis in ultrabasics, another antipathy should be considered, namely that of Ti and V. Willemse (1969) in his study of the plug-like bodies and seams of vanadiferous magnetite in the Bushveld Complex states the following: "Many seams have a sharp basal contact against anorthosite, and a transitional hanging wall composed

of massive ore grading upwards into magnetite-anorthite and hyperite. A thin, highly puckered veneer of olivine and an intergrowth of pyroxene and plagioclase are developed at the sharp basal contact. The intergrowth could be due to diffusion of iron into the plagioclase. Irregularities at the contact consist of veinlike offshoots of magnetite into the foot wall and even brecciation of the foot wall is encountered. The grain size of the ore in the Main Seam is variable. The closely packed nature of the grains is best explained by enlargement after concentration. The magnetite started to crystallize earlier or simultaneously with the plagioclase and the seams could have formed in the same manner as any monomineralic rock in a layered sequence. In some of the country rocks magnetite continued to crystallize interstitially to the plagioclase. In these examples hydroxyl minerals such as biotite and amphibole are generally present. Titanium minerals in the ore consist mainly of (1) small ilmenite granules that probably formed earlier or simultaneously with the magnetite, (2) rather rare, broad lamellae of ilmenite which exsolved from magnetite at an early stage, (3) ulvite, which is ubiquitously present and forms the characteristic cloth-pattern and (4) a wide range of intergrowths of dispersed ilmenite ("proto-ilmenite") with magnetite, maghemite and martite. The intergrowths are considered to have formed from the ulvite cloth-pattern by surface or near surface oxidation. The cloth pattern of ulvite seems to have had a profound effect on the formation of maghemite and mar-

tite - it promoted maghemitization and retarded martitization. An antipathetic relationship prevails between V 2 0 5 and the T i 0 2 content of the ore. The lowermost seam contains about 2% V 2 O s and 14% T i 0 2 , the uppermost one about 0.3% V 2 O s and 18-20% T i 0 2 . As the titianium is mostly contained in ulvite it is considered that a decrease in oxygen fugacity in the magma determined the increase in the titanium content upward in the layered sequence. The V 2 0 5 content of the magnetite plugs seems to correspond approximately to that of the seams in their vicinity." The understanding of the antipathy of Ti to V as mentioned by Willemse, and also the mentioned "antipathy" of Ti to Cr are even more difficult since the three elements Ti - V - Cr are interrelated among each other according the empirical "laws" of the periodic system (Remy, 1961). "Ähnlichkeiten im Verhalten der im Periodensytem nebeneinanderstehenden Elemente nehmen von der III zur VIII Nebengruppe, von dort über die I bis zur III Nebengruppe wieder ab." Their "antipathy" should rather be understood by considering their paragenesis (i. e., paragenetic association of the minerals in which they tend to occur) than from the point of view of element to element interrelationship. Concerning the textural patterns of vanadiferous magnetites with ilmentite intergrowths, the reader is referred to Figs. 907 and 908 and Chapter 20.

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Chapter 53

The Te, Se, Bi, Au, Ag Element Segregation/Distribution (in Paragenetic Associations)

In addition to the hydrothermal uranium paragenesis (see page 230) where two groups of elements have been recognized, namely A = U (Th), Pb, Bi, As, Sb, (Se, Mo, Sn), and Β = Fe, Co, Ni, Cu, Au, Ag, similarly in the Te, Se, Bi, Au, Ag hydrothermal paragenesis - taking into consideration the metallic elements on the basis of the composition of the reported minerals - two comparable groups of elements are recognized which somehow correspond to the groups found in the case of hydrothermal uranium paragenesis. The two groups are: group A = Se, Te, Bi, Sb, As, Pb, Sn, (Tl, In), (?Mo), group Β = Fe, Co, Ni, Cu, Cd, Au, Ag. The recognition of these two groups is based on the composition of the reported minerals, considering a number of study cases of the above paragenesis 28 . As the discussion of the study cases will show, not all the elements of group A and Β are represented in 28

Considering the ideal cases of group A = Pb, Sb, As, Bi, Se, Te, Sn (Tl, In), ?Mo and group Β = Fe, (Μη), Ni, Co, Cu, Au, Ag, it is often found that in a paragenetic mineral association elements from both groups may be represented (and often build a mineral). Furthermore, as could be seen, group Β is rather mantle derivative, while group A can include elements from lower crust or mantle (e. g., Pb). The relationship of elements of group A and Β and the special case of the distribution of silver in connection with the formation of massive sulfides is considered based on the paper by Amcoff (1984): "Distribution of silver in massive sulfide ores". According to him "(1) The affinity between silver and galena-rich ores is only pronounced if antimony and/or bismuth are also present in significant amounts, and (2) the degree of correlation between silver and lead is partly a function of the Ag/Sb + Bi ratio in the ore and increasing with the decrease in ratio at least up to Ag/Sb + Bi = 1. Observations of correlated Cu/Ag ratios in complex ores indicate that some silver was solved in chalcopyrite at the time of deposition. This is often masked by the much higher silver contents associated with galena. The Bi/Sb ratio is expected to increase downwards stratigraphically in galena-rich ores, due to a larger solubility of silver-bismuth in galena as compared with silver-antimony, at a realistic deposition temperature (200-300°C)." Thus, in addition to the derivation (and distribution) of elements in accordance with the empirical "laws" of the periodic system, other factors should also be taken into consideration.

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selected examples of the minerals comprising the reported elements. Each study case shows that the minerals comprising the paragenesis have some elements from group A and B. The ideal case of group A and Β is derived by taking into consideration a number of study cases of the Te, Se, Bi, Au, Ag parageneses. More specifically, Ishibashi (1960) in his contribution "Au, Ag-tellurides from Dat6 mine, Hokkaido, Japan" reports the presence of krennerite (AuTe 2 ), hessite (Ag2Te) and petzite (Ag 3 AuTe 2 ) as Au-Ag-telluride minerals in association with sulphides (blende, galena, chalcopyrite and pyrite). Taking into consideration the metallic elements of this paragenesis, group A includes: Te, Pb and Bi (as reported) and group B: Zn, Fe, Cu, Au, Ag. In addition, Atanasov and Eskenazi (1964) in their study "Rare elements in galena from Madjarovo" determined Bi, Ag, Te and Sb in galena from the leadzinc mine of Madjarovo, eastern Rhodopes, as inclusions of tellurbismuth (Bi 2 Te 3 ), pyrargyrite (As3Sb[As]S3), stephanite (Ag 5 SbS 4 ) and polybasite (Ag 16 Sb 2 S u ). According to them, Bi, Ag and Te also take part in isomorphous substitution, whereas Cd and Tl are considered to occur only as isomorphous admixtures. Furthermore as is reported, in depth the content of Bi, Te and Tl increases while that of As, Sb and Au decreases. On the basis of the composition of the reported minerals, group A is represented by Bi, Te, Sb, As, (Tl), Pb, and group Β by Zn (Fe) Cd, Ag, Au. Morävek (1956), in "Bismuth minerals from the Slorir vein in the area of the mine Pepr near Jilov6", reports that besides metallic Bi and bismuthinite (Bi2S3) and tetradymite (Bi 2 Te 2 S) reported earlier, cosalite (Pb2Bi2S5), kobellite (Pb 6 FeBi 4 Sb 2 S 16 ) and tellurbismuth (Bi2Te3) occur, thus the paragenesis is characterized mainly by group A elements: Bi, Te, Pb, Sb. Another interesting, yet not typical case is reported by Meituv (1962) in his study of the rare elements in the lead-zinc deposit of the Klichkinskic region (eastern Transbaikaliya). According to him, cadmium is enriched in sphalerite (0.1 to 0.8%) and in galena, boulangerite (Pb s Sb 4 S u ) and chalcopyrite (0.001 to 0.005%). Later sphalerites contain more Cd than early formed sphalerites. Indium is concentrated in early

sphalerites (0.005 to 0.1%) along with tin, but is independent from the iron content. Furthermore, when the sulphide ores contain much alumino-silicate gangue, T1 and Ga are strongly dispersed; these elements are concentrated in sulphides in deposits localized in carbonate rocks and calcareous schists where aluminosilicate gangue is uncommon. Se and Te are found in pyrite but the highest concentration occurs with galena. On the basis of the reported geochemistry and mineralogy, group A is represented by: Pb, Sb, Se, Te, Sn (Tl), and group Β by Zn and Fe. In contradistinction to the cases presented so far in the presently reported paragenesis, Bi, Au, Ag are not represented in either group A or B. Considering the element segregation of Te, Se, Bi, Au, Ag paragenesis further, Shcherbina and Zar'yan (1964), in "Paragenesis of silver and gold tellurides as solid phases in the system Ag-Au-Te", report the following: "Parageneses of gold and silver tellurides from some of the sulphide deposits of Armenia are described and discussed in terms of the experimentally determined portions of the system Ag-Au-Te. The assemblages hessite-petzite-gold and hessite-petzitekrennerite were formed by deposition from hydrothermal solutions during the galena-sphalerite stage of mineralization and by later replacements. (Native silver and tellurium formed by the decomposition of gold-silver tellurides which occur in some assemblages)". According to the composition of the reported minerals, group A is represented by Te and Pb, and group Β by Zn, Au, Ag. An example of just how complex mineralogically and geochemically the Te, Se, Bi, Au, Ag paragenesis can be, is shown in the contribution of Markova (1967), "Occurrence of volynskite in a gold deposit of Central Asia". According to her, volynskite (AgBiTe 2 ) occurs with pyrite, tetrahedrite (Cu 3 SbS 3 ), galena, sphalerite and small amounts of Bi tellurides in a gangue of quartz, carbonates, sericite and rare baryte. The volynskite was the latest of the Bi tellurides to be formed and occurred in the sequence tetradymite (Bi 2 Te 2 S) - » joseite (Bi 3 TeS) —> petzite (Ag 3 AuTe 2 ) tellurobismuthite (Bi 2 Te 3 ) —> altaite (PbTe) sylvanite (AuAgTe 4 ) calaverite (AuTe 2 ) —> hessite (Ag 2 Te). Considering the composition of the reported minerals and their transitions, group A is represented by Bi, Pb, Sb, Te, and group Β by Fe, Cu, Au, Ag. The elements comprising group A are Se, Te, Bi, As, Pb, Sn, (Tl, In) (?Mo) and it is tentatively suggested that they are interrelated as follows (in accordance with the empirical "laws" of the periodic system):

Se and Te are related as homologous main elements of the VI family; Se is related to As as main elements of the V and VI family (next to each other); Te is related to Sb as main elements of the V and VI family (next to each other). As, Sb, Bi are homologous main elements of the V family; Sn is related to Sb as main elements (next to each other) of the IV and V family; Pb is related to Bi as main element (next to each other) of the IV and V family; Tl and In as main elements (homologues) of the III family and Tl and In as main elements of the III family next to Pb and Sn (main elements) of the IV family, respectively. The metallic elements comprising group Β = Fe, (Μη), Co, Ni, Cu, Cd, Au, Ag, are interrelated as follows: Fe, Co, Ni are subgroup elements of the VIII family; Mn is related to Fe as subgroup elements horizontally, Cu similarly to (Ni, Co, Fe); Cu and Zn as subgroup elements horizontally next to each other. The elements Cu, Ag, Au as vertically related subgroup elements and similarly Zn and Cd. As it was the case with the elements comprising the elements of the U paragenesis (A, B, and special group C), the "interrelated" elements do not strictly follow the rule that elements with the strongest affinities and relationships are paragenetically associated, i. e., occur as mineral building elements in the paragenesis. Elements with greater affinities (as stated by the empirical "laws" of the periodic system) than those of elements included in a group may be absent. The empirical "laws" of element affinities and interrelationships should not be considered as rules governing the segregation/distribution of elements forming parageneses but only as an explanation of relationships existing among elements of the groups recognized. Nevertheless the fact that Te, Se, Bi, Au, Ag parageneses of worldwide distribution show similarities in the elements comprising their metallic minerals, can be understood to some extent from the point of view of the chemical and geochemical affinities of these elements. Furthermore, even if the outlined affinities and interrelationships (according to the empirical "laws" of the periodic system) are not the strongest possible - often they are only weak - and despite all these reservations, some of these elements are interrelated among themselves, and that might be sufficient cause for their common segregation, since very scarce elements in the crust (or mantle) are segregated together in typical associations which in turn result in mineral assemblages and paragenetic associations of the upper crust.

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Chapter 54

Realgar, Orpiment - CinnabarMetacinnabar Parageneses

Considering the Te, Se, Bi, Au, Ag element segregation (paragenetic associations) and in addition to the cases discussed and the recognition of the two groups of elements, group A = Se, Te, Bi, Sb, As, Pb, Sn, (Tl, In), (?Mo), and group Β = Fe, Co, Ni, Cu, Cd, Au, Ag, certain paragenetic associations will be discussed which, despite their belonging to the Te, Se, Bi, Au, Ag-type of element segregation, exhibit a prevalence for the element As and to a lesser degree for Bi in group A, while some of the elements formerly prevalent (in the previous study cases) may be absent. In addition, in the elements of group B, the element Au appears to be prevalent in some cases, and the element Hg is also abundantly present. In this connection, it should be pointed out that Au and Hg are interrelated as subgroup elements horizontally and belong to the I and II family. Thus, Hg should be included in group B, so Β = Fe, Co, Ni, Cu, Cd, Au, Hg and Ag. It should also be emphasized when looking at these study cases, that As, Bi, Au and Hg are often strongly represented as elements segregated in these paragenetic associations. Boyer and Picot (1963) state that maldonite (BiAu2) occurs associated with bismuthinite (Bi2S3) and native gold in mineralized fault in low-grade schist about 50 km north of Carcassonne. According to them, other minerals included are pyrite (FeS2), pyrrhotite (FeS), chalcopyrite (Cu,FeS2) and arsenopyrite (FeAsS - with Co) and less frequently stannite (Cu2FeSnS4), polybasite (Ag 1 6 Sb 2 S n ) and guanajuatite (Bi2Se3); also present are wolframite ((FeMn)W0 4 ) and rutile (Ti0 2 ). Considering the metallic elements which comprise this mineral assemblage, it should be noted that the A group elements represented are As, Bi, Sn, Sb and the Β group elements represented in this mineral association are Fe (Ti), Mn, Cu, Ag, Au. The presence of W is difficult to explain (see page 246). Another comparable element segregation case is described by Kon'kova (1964), in "New minerals from one part of the Altyn-Topkan ore-field". Kon'kova claims that löllingite (FeAs2 with Co, Ni), realgar (AsS), orpiment (As 2 S 3 ), clausthalite (PbSe), dyscrasite (Ag3Sb) are reported from Altyn-Topkan ore field. In the löllingite, Co, Pb, Sb, Bi, Cu, Mn and Ag are shown in the spectrographical analysis. Furthermore, clausthalite together with dyscrasite and S occur as fine drop-shaped inclusions in galena (PbS). On the ba244

sis of the composition of the reported minerals, elements of the A group represented in this mineral association are As, Pb, Sb, Se and of the Β group Fe, Co, Ni, Ag, Cu. Tollon and Orliac (1966) support that "solutions from the auriferous veins which follow N-S faults of Salsigne (15 km north of Carcassonne) have mineralized the sandstone walls up to distances of 50 km". According to Tollon and Orliac, in the tough sandstones with splintery fracture, the ore and gangue minerals occur as fracture fillings normal to the bedding, and in softer bedded sandstones as impregnations along biotite-rich layers. The paragenesis is as follows: the earliest ore mineral is arsenopyrite (FeAsS with Co), Au, bismuthinite (Bi2S3) and maldonite (BiAu 2 ). Considering the composition of the reported minerals in this paragenetic association As and Bi are representatives of the A group and Fe, Co and Au of the Β group. Furthermore, Cheraitsyn and Apostolov (1966) report that the ore occurrences on the northern slope of the Greater Caucasus, in the pre-Jurassic basement of the epi-Hercynian platform mainly consist of arsenopyrite, native As, antimonite (Sb 2 S 3 ) and cinnabar (HgS). Less common according to them are sphalerite (ZnS), chalcopyrite (CuFeS2), realgar (AsS), orpiment (As2S3), galena (PbS) and magnetite (Fe 3 0 4 ). In this case too, considering the composition of the reported minerals, As, Sb and Pb are representatives of the A group, and Fe, Co, Zn and Hg of the Β group. ft should be note that the presence of cinnabar (HgS) is geochemically significant since Hg is a strongly represented element of the Β group (see page 243). According to Kulcsär (1970), Au and cinnabar occur at the eastern margin of the Tokaj range in Hungary. The gold occurs in thin plates, up to 0.3 mm across, in montmorillonite clay dykes in K-metasomatized rhyolite tuff in the Rudabanya-Berg, in which a mineralization of Pb, Zn, Ag and As is also found. Considering the reported mainly metallic elements present, Pb and As are representative of group A, and Zn, Ag, Au, Hg of group B. The association of native Au and cinnabar is perhaps best understood by considering the strong interrelationship of Au and Hg as subgroup elements next to each other. Furthermore, Maleyev (1967), in "Types of mercury mineralization and their relation to volcanism" reports

that in Transcarpathia mercury occurs as (1) disseminations in quartzite of cinnabar (HgS) and metacinnabar (HgS) associated with melnikovite (7FeS0 4 -7H20), wehlite (Bi 3 Te 2 ), native Bi, quartz and calcite deposited during orogenic volcanism, and (2) concentrations in altered clay rocks along fault zones and composed of cinnabar and metacinnabar associ-

ated with marcasite, quartz, chalcedony, siderite, dolomite, artinite, calcite and inorganic bitumen deposited during post orogenic volcanism from a subcrustal magma. According to the composition the reported minerals Bi and Te are representatives of the A group, and Fe and Hg of the Β group.

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Chapter 55

A Special Case of Non-Ferrous Metal Mineralization in Evaporites

Very different from the cases discussed so far is the non-ferrous metal mineralization in the evaporite deposit of Myrthengraben, Austria, presented by Tufar (1982) in his contribution "A new type of sulphosalt mineralization in the Myrthengraben gypsum deposit, Semmering, Lower Austria". According to him, "a Camian slate and dolomite series of the Semmering Mesozoic contains a gypsum-anhydrite deposit characterized by an extremely complex, iron-poor, non-ferrous metal mineralization (i. e., Cu, As, Pb, Zn, Fe, + Sb, Se, U, etc.). Sulphides and sulphosalts are the primary ore minerals". Furthermore, Tufar reports that major mineral constituents are enargite (Cu3AsS4), tennantite (Cu 3 AsS 3 2S), wurtzite (ZnS with minor Fe, Μη and Cd), galena (PbS), pyrite (FeS2), jordanite (Pb 4 As 2 S 7 ), luzonite (Cu2(As,Sb)S4), seligmannite (CuPbAsSj), stibnite (Sb2S3), a Cu-Zn-Sn-S mineral and a new AsS mineral. Also, an uranyl-carbonate, andersonite, occurs in the supergene zone. Tufar em-

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phasizes that "based on paragenesis and observed ore mineral intergrowths, the iron-poor, complex, non-ferrous metal mineralization in the evaporite deposit of Myrthengraben presents a totally new type of sulphosalt occurrence". On the basis of the recorded minerals and their composition, and of the elements mentioned as present, the following groups of elements are tentatively recognized (as was the case with other parageneses considered so far): group A = Pb, As, Sb, Sn (Se); and group Β = Fe, Μη, Cu, Zn, (Cd). It should also be pointed out that in this admittedly totally new paragenesis the elements of group Β (Fe, Μη, Cu, Zn, (Cd)) are most likely mantle derivatives, whereas the elements of group A (Pb, As, Sb, Sn (Se)) are perhaps derivatives of the lower crust. Element recycling and the role of brines have possibly played a significant role in this case, too (see Chapters 31 and 39).

Chapter 56

The Segregation (Distribution) of Sn, Mo and W to Form Concentrations or Ore Deposits

When studying the segregation of the Sn, Mo, W to form concentrations or ore deposits, the interrelationship of these elements with other elements, in which each one is interrelated in accordance with the empirical "laws" of the periodic system is of importance. Also the extent to which these elements are interrelated among each other is very significant. Perhaps a better understanding of the interrelationships of a wide range of elements that are interrelated with the elements Sn, Mo, W is easier by considering some theoretically possible interrelations of the "key" element Sn. (i) The element Sn is related to Sb as main elements next to each other of the IV and V family of the periodic system. Sn and Pb are also related as main element homologues. There is a weak interrelationship between Sn and Bi, since Bi is homologous to Sb. Thus, the "key" element Sn is interrelated on the basis of the empirical "laws" of the periodic system with Pb, Sb and Bi (and also with As, In and Ga, see page 247). (ii) Furthermore, the "key" element Sn is interrelated with Zr (and Hf) as main element (Sn) and subgroup elements (Zr, Hf) of the same IV family. In turn, the elements Zr, Nb and Mo are interrelated as subgroup elements (horizontally); thus Sn is interrelated with the elements Zr (Hf), Nb and Mo. It should be noted that Mo and W are strongly interrelated as homologous subgroup elements of the VI family. As a consequence, the element Sn is interrelated with elements Zr (Hf), Nb, Mo, W. Furthermore, Nb and Ta are vertically interrelated as subgroup elements of the V family. Thus, Sn is interrelated with Zr (Hf), Nb, Ta, Mo and W. In this connection it should be noted that Zr (Hf), Nb, Ta 2 9 , Mo and W are elements characteristic of the granitic-pegmatitic geoenvironment. In contrast to the above mentioned groups of elements with which Sn is related, namely Pb, Sb and Bi, and as mentioned under (ii) Zr (Hf), Nb, Ta, Mo and W, there is another group of elements, while not inter29

As a corollary to the interrelationship of Sn with Nb and Ta, Hosking (1982) reports that "intensely red-pale coloured cassiterites containing Ta and possibly Nb in the lattice are restricted to the west Southeast Asian Tin Belt, while those that are brown-pale coloured pleochroic contain also Nb/Ta or W in their lattice".

related with Sn, that nevertheless is often found in paragenetic (geochemical) association not only with Sn, but also with Mo and W. This group of elements is Fe (Μη) Co, Ni, Cu, Zn (Ag, Au) which is characteristic or derivative of basic rocks or mantle. An attempt will be made to understand the common segregation of metallic elements with each of the elements Sn, Mo and W or with all or some of them (e. g., Mo, W). Radkewitsche and Poljakowa (1960) presented a most interesting element segregation concerning Sn. Complex Pb-Sn sulphide mineralizations of Boliviantype are described from Sinancha, Sikhote-Alin and from Smirnovsk, Transbaikalia, both in the USSR. According to them, in Sinancha the ores occur in quartz porphyries and in Smirnovsk they are in carbonate rocks. The ore minerals are galena (PbS), stannite (Cu2FeSnS4), franckeite (Pb 2 Sn 3 Sb 2 S u ) and tealite (PbSnS2) with cylindrite (Pb 3 Sn 4 Sb 2 S 14 ) only at Smirnovsk. On the basis of compositions of the reported mineral paragenesis, the Sn is related to Pb and Sb as it is supported by (i). Thus, the elements comprising this paragenesis can be considered as the two groups mentioned. As a corollary to the above mentioned hypothesis of groups of interrelated elements, occurring simultaneously in mineral associations is the study case presented by Lawrence and Golding (1969) of cubaniterich sulphide ores from the McAlpine copper mine, which in addition to cubanite (CuFe 2 S 2 ) contain cassiterite (SnO), sphalerite (ZnS), chalcopyrite (CuFeS2), magnetite (Fe 3 0 4 ), galena (PbS) and pyrite (FeS2) together with a number of ore minerals containing Bi, Cu, Sn and Ni. What seems to be of significance is that the ore metals are believed to be of hybrid origin derived partly from a granite and partly from an ultramafic magmatic source. This interpretation is furthermore supported by considering the metallic elements characteristic of this mineral association, more in detail: the group Sn, Pb, Bi are, as pointed out, granitophile elements (see (i)) whereas the elements Fe, Ni, Cu, Zn are interrelated elements of basic mantle derivation (see (iii) page 229). In addition, a rather more complex case is presented by Dudykina (1959) on the paragenetic association of trace elements in cassiterites of various genetic types of tin ore. According to Dudykina, the study of trace 247

elements in 271 cassiterites of different genesis showed that every genetic type is accompanied by a deposit assemblage of trace elements. The cassiterites of sulphide-cassiterite formation, for example, are accompanied by W, Pb, Sb, Ag, In, As and Ga. Cassiterites from pegmatites are accompanied by Zr, Nb, Ta. Considering the trace element associations as presented by Dudykina, the cassiterites (from pegmatites) characterized by the trace elements Zr, Nb, Ta, clearly represent case no. (ii) of the present interpretation of element segregation (page 246). Furthermore, the group of trace elements W, Pb, Sb, Ag, In, As and Ga characteristic of the sulphide-cassiterites can be subdivided (according to the above mentioned interpretation of element segregation) as follows: Sn, Pb, Sb, In, Ga, As as elements interrelated according the empirical "laws" of the periodic system and as mainly of crust derivation. W is interrelated to Sn (as stated in (ii) on page 246); and Ag is interrelated to the group of elements Fe, (Μη), Co, Ni, Cu, Zn, (Ag, Au) which, as pointed out, are mantle derivatives. When studying the paragenetic associations of trace elements in cassiterites as presented by Dudykina, and on the basis of the present hypothesis of element segregation, the elements can be grouped as follows: (i) Sn, Pb, In, Ga, As (ii) Zr, Nb, Ta, W (iii) Ag (which represents the group Fe, (Μη), Co, Cu, Ni, Zn (Ag, Au). On the basis of the interpretations presented concerning the Sn, Mo, W segregation and concentrations, the significant role of the empirical "laws" of the periodic system has been emphasized. An attempt will be made to present study cases concerning various mineralizations of these elements (Sn, Mo, W) on the basis of the orthodox views, i. e., based on a granitic magma intrusion. Most of the available study cases are based on the granitic magma hypothesis. However, the present cases could equally be interpreted on the granitization hypothesis as presented in Chapter 22. Since in these study cases the mineral associations (parageneses) are of importance, more emphasis will be put on geochemical aspects of mineralization than on the controversy of granites-pegmatites. Furthermore, as the empirical "laws" of element segregation support, considering the study cases of the elements Sn, Mo and W, interpretations based on the empirical "laws" of the periodic system will be used to interpret mineral associations as far as possible. Stumpfl (1963) in his geochemical study of some Bushveld tin deposits suggested a critical review of genetic significance of certain trace elements in cassiterite. According to him, Nb and Sc contents in cassiterites depend on the temperature of formation. Older, high-temperature cassiterites in lode deposits average 40 ppm Nb and 130 ppm Sc. Younger cassiterites in replacement ore bodies have 15 ppm Nb and 33 ppm Sc. Furthermore, Ti, Pb and Ag cannot be 248

used for genetic conclusions. Regarding the results of Stumpfl, two significant points should be emphasized: (i) The trace elements Nb, Sc present in the cassiterites are related to Sn in accordance with the empirical "laws" of the periodic system. (Nb is related to Zr and Zr to Sn - as main and subgroup elements of the IV family). Similarly Ti is related to Zr as subgroup elements of the IV family (vertical relations); and Sc to Ti as subgroup elements horizontally. In addition, Sb is related to Sn as main elements next to each other. In contrast, Ag is mantle derived and belongs to the group Fe, Co, Ni, Cu, (Ag), Zn, as pointed out earlier. (ii) The second significant point presented by Stumpfl - higher Nb and Sc contents are present in higher temperature cassiterites - is explainable on the basis of Goldschmidt's 'laws' of distribution of elements and depends on the state of the cassiterite lattice to incorporate various sizes of atoms. Comparable studies by Nikulin (1967) on the distribution of indium, niobium and scandium in the cassiterites of the Khingan deposit, eastern Siberia, support changes in the concentration of In, Nb, Sc with depth of cassiterites. (As pointed out, In is interrelated with Sn as main elements of the III and IV family, respectively). Thus, also in this case, the trace elements present in cassiterite are interrelated with it. Their incorporation into the lattice of Sn0 2 is to be understood in accordance with Goldschmidt's 'laws' of element distribution. Concerning the relationship of Sn to Nb, de Kun (1958) reported of occurrences of cassiterite of Nord Lugulu, Kiva, Zaire, with ratios of Sn:Nb of the order of 15:1, which again is in accordance with the periodic system and with the laws of element distribution by Goldschmidt. In contradistinction to the cases described, Ramdohr (1960) described cassiterite with thin ex-solution particles of tapiolite and others of tapiolite and columbite. Also in this case, Nb and Ta are interrelated with Sn by the empirical "laws" of the periodic system (see page 246). Furthermore, Fe belongs to the group Fe, Ni, Co, Cu, Zn, Ag and Au as pointed out. Further consideration of the segregation and concentration of the Sn, W, Mo elements to form deposits, shows that besides the cases of Sn, the association of Sn and W is of particular significance. As was discussed, Sn is interrelated with W in accordance with the empirical "laws" of the periodic system. However, this relationship is not particularly strong, which may perhaps explain certain cases of Sn and W differentiation within the same mineralization. Furthermore, Gouanvic and Babkine (1985), considering the mineralization of tungsten in Monteneme (northwestern Galice, Spain) report that the tungsten-Sn mineralization forms a Stockwerk enclosed in monzonites. Three patterns of mineralized veins, successively emplaced, show evolusive paragenesis coherent with decrease in crystallization temperature. Also many Bi, Pb, Ag-SbSe phases are recognizable and the veins are related to

an apogranitic cupola. Both the W-Sn mineralization and the Bi-Pb-Ag-Sb-Se phases considered geochemically show elements that can be grouped together according their relationships in accordance with the empirical "laws" of the periodic system. The following interrelated groups are tentatively recognized: (1) Sn, Sb, Pb, (Se); (2) Sn, W; (3) Ag (Fe, Cu). The chemical interrelationship of the elements comprising each group has already been discussed (see pages 229, 246). Sainsbury and Hamilton (1967) support that the principle lode-tin deposits of the world are closely related to biotite-muscovite granite or to subvolcanic dacite or rhyolitic rocks along major orogenic belts. They also maintain that in Precambrian rocks these belts are broader than in younger rocks, and most of the deposits are in pegmatite with spodumene, tantalite, columbite and other heavy minerals. In post-Cambrian rocks contact metamorphic pneumatolytic-hydrothermal, subvolcanic or tin-silver "fumarolic" and disseminated deposits are recognized. Also according to Sainsbury and Hamilton, the principle deposits are of the pneumatolytic-hydrothermal and tin-silver types, and are localized along zones of deep faulting within the broader granitic belts. In addition, within each district, the Sn deposits have been formed in the higher temperature or deeper parts of the lodes and grade upwards or outwards into Cu and W deposits. Considering the main metallic elements associated with Sn mineralization as supported by Sainsbury and Hamilton, there appears to be a tendency for Sn mineralization to be of relatively higher temperature than Cu and W mineralization and a "tendency" for "dissociation" of Sn and W, despite the often reported common occurrences of Sn and W. In contrast, Edwards and Lyon (1957), studying the mineralization at Aberfoyle Tin Mine, Rossarden, Tasmania, report that within the productive zone there is a marked downward increase in the wolframite content of the veins at the expense of cassiterite. In this case too, a "tendency" for differentiation "disassociation" of Sn and W is perhaps recognizable. Adam (1960) supports that in the Klappa Kampit Hill area "cassiterite is never found in association with wolframite, although veins containing these minerals singly may occur in close proximity". This tendency for Sn and W to "disassociate" might well be due, in this case, to their weak interrelationship as the empirical "laws" of the periodic system suggest. Despite these reported cases where a tendency for differentiation or "disassociation" of Sn and W is found, cases of mineralization are reported where Sn and W occur together in paragenetic associations. According to Zen Chin-Fung and Yang Bo-Lin (1965) in the Nanling region of China, a high-temperature W-Sn vein deposit occurs in metamorphic rocks (of Caledonian age) and is believed to be related to granites of the Eanshanian orogeny. In addition, Stemprock (1967) supports that the "relations between granite contacts and tin-tungsten

mineralization in the Erzgebirge (Kru§n6 Hory), Czechoslovakia, result from the structural control of solidified granite margins on the passage of post-magmatic solutions that originated at depth in the earth's crust". Furthermore, according to Williams (1958), in "Tintungsten mineralization at Moina, Tasmania", "tintungsten ores of the Murphy and Shepherd Mine, Moina, occur in a series of vertical quartz veins within the contact aureole of a Devonian granite". In addition, Williams states that the vein minerals include wolframite, cassiterite, pyrite, marcasite, magnetite, haematite, pyrrhotite, arsenopyrite, molybdenite, chalcopyrite, sphalerite, bismuthinite, bismuth, galena and scheelite. Considering the elements present Pb, Sn, As, Bi are interrelated as supported on page 230. The elements Fe, Zn, Cu are related as indicated on page 230 and are most likely derivatives of basic rocks or mantle. Often there is a common segregation or concentration of the elements Sn, W and Mo as already pointed out. Whereas Sn and W or Mo are weakly interrelated in accordance with the empirical "laws" of the periodic system, the elements W and Mo are strongly related as subgroup elements (vertically related) of the VI family of the periodic system. As a consequence of these interrelationship Sn, W and Mo occur together, and cases of W and Mo segregated/(mobilized) together are common. According to Pavlov (1965), in "New data on the patterns of distribution of endogenic mineralization in the Soviet Maritime province", tin, tungsten and molybdenum mineralization is generally concentrated in the granitoid bodies localized by tectonic sutures (see Chapter 18). Yet another case of Sn, W and Mo mineralization is reported by Stemprock (1964) in his consideration of the tin-tungsten ore paragenesis in the Κηιδηέ Hory Mountains. Observations support that quartz-micafeldspar assemblages occurring in the tin and tungsten veins are not real pegmatites but feldspathized portions of the earlier veins. According to Stemprock, the feldspathization occurred after greisenization as an independent process. During feldspathization replacement of quartz mainly by K-feldspar (adularia) took place. This young feldspathization was observed recently in the district with tin, tungsten and molybdenum metallization. In this case too, the elements Sn, W and Mo representative of this type of mineralization are interrelated in accordance with the empirical "laws" of the periodic system, as pointed out on page 246. In contradistinction to the common segregation of Sn, W and Mo as already pointed out, cases are described where segregation of elements occurs in which W and Mo are represented and where Sn is missing. According to Sharp (1958), in "Mineralization in the intrusive rocks in Little Cottonwood Canyon, Utah", scheelite and molybdenite occur in zones of jointing 249

and fracturing which form a concentric pattern around an intensely fractured central zone, (the more intensely fractured zones are mainly within a leuco-quartz monzonite). In contrast to the rather simple scheelite-molybdenite mineral paragenetic association as pointed out by Sharp, a more complex paragenesis is reported from Chhendapathar, Bankura, West Bengal by Chakravarty (1958). It consists of wolframite ((Fe,Mn)W04, scheelite (CaW0 4 ), molybdenite (MoS2), bismuthinite (Bi2S3), magnetite (Fe 3 0 4 ), ilmenite (TeTi03), pyrite (FeS2), arsenopyrite (FeAsS, with Co), pyrrhotite (FeS), and chalcopyrite (CuFeS2). Considering the

250

composition of the reported minerals by Chakravarty, the following groups of elements interrelated in accordance with the empirical "laws" of the periodic system are recognized: W and Mo (subgroup elements of the VI family, vertically interrelated), As, Bi (as elements of the V family) and Cu, Co, Fe as subgroup elements (horizontally related). Ti is related with Zr (and Zr, Nb, Mo are interrelated horizontally as sub-group elements, see page 246). In this case too, the metallic elements comprising the paragenetic mineral associations are interrelated in accordance with relationships based on the empirical "laws" of the periodic system.

Chapter 57

(a) Sn-Associated Pegmatites

Special Cases of Element Segregation/Distribution

Element-Segregation

in

Sn is interrelated to Zr as a main group element and subgroup element of the IV family of the periodic system as already pointed out. Furthermore, Nb is related to Zr as subgroup elements next to each other and Nb and Ta as subgroup elements of the V family, vertically related. Independently to these relationships, Fe is related to Mn (horizontally interrelated subgroup element). These elements characterize the Β group: Mn, Fe, Ni, Co, Cu, Zn, Ag, Au (- mantle derivatives, see page 230). Kazmitcheff (1959-60), in his study "Observations sur la pegmatite de Tshonka (Nord-Lugulu, Kivu)" reports that a pegmatite occurs as replacement veins and pockets in a dyke of fine-grained granitoid rock which is mineralized with cassiterite (Sn0 2 ) and columbotantalite (Niobite (Fe,Mn)Nb 2 0 6 tantalite (Fe,Mn)Ta20 6 ). Considering the metallic elements segregated in this pegmatite, as mentioned, Sn is related to Nb and Ta, and Fe is interrelated to Mn (B-group of elements, mantle derivatives). Sn is related to Nb and Ta, characteristic elements of pegmatitic-granitic geoenvironment (see page 246). Bernard (1954-55) described zoned pegmatitic veins from Manono. According to him, "two flat veins break off from the north of the dyke of pegmatite at ManonoKatanga. They are zoned, with three main phases of deposition. Spodumene is described (Li carrier). The heavy minerals reported are: tantalocolumbite [Niobite (Fe,Mn)Nb 2 0 6 , tantalite (Fe.M^TajOg], cassiterite (Sn0 2 ), thoreaulite [Sn(Ta,Nb) 2 0 7 ], löllingite (FeAs2), pyrite (FeS2), and galena (PbS)." Considering the composition of the reported minerals, Sn, Nb, Ta are interrelated as pointed out; Sn and Pb are related as main elements of the IV family of the periodic system. As is related to Nb as main and subgroup elements of the V family; and the elements Fe and Mn (horizontally interrelated subgroup elements, and as belonging to mantle-derived Β group elements: Mn, Fe, Co, Ni, Cu, Zn, Cd, Ag and Au). Thus, the elements segregated in the pegmatite at Manona (Katanga) are Sn, Nb, Ta (typical pegmatitic elements), Pb, As (as crust elements, see page 193) and Mn, Fe most likely mantle derivatives (B group elements).

Another study case from north Lugula, Zaire, by de Kun (1954-55) describes tin-bearing pegmatites with cassiterite (Sn0 2 ), columbite [niobite (Fe, Mn)Nb 2 0 6 tantalite (Fe.M^TajO,;], tourmaline (B-bearing); ixiolite (Sn-tantalite) and tantaliferous cassiterite (Ta Sn0 2 ). As the composition of the observed minerals indicates, the interrelated elements Sn, Nb, Ta represent typical pegmatitic-granitic elements and Mn, Fe ore elements of the Β group (see page 230, most probably mantle derivatives). Considering the reported cases of tin-bearing pegmatites in addition to the typical pegmatitic-granitic metallic elements (Sn, Nb, Ta) also mantle in derivation elements such as Fe, Mn are jointly segregated in these pegmatites, indicative of element recycling and multisource derivation.

(b) Be-Associated Elements (Segregation) in Pegmatites with Transition to Vein Deposits According to Stoll (1965), the metalliferous deposits of the Sierras Pampeanas, Argentina, are Precambrian and Tertiary in age. In particular, the Precambrian deposits include pegmatites with exploitable beryl, feldspar, mica, lithium minerals, columbite-tantalite and minerals of the U, W, Mo, Sn and quartz deposits of cassiterite (Sn0 2 ), Au, Pb, Ag, Zn, Cu and V. Genetically, these deposits are related to granitic intrusions and probably form a great zonal series ranging from pegmatitic and pneumatolytic stage deposits to those of the mesothermal and hydrothermal stages. Considering the composition of the ore minerals reported (including beryl [Al 2 Be 3 (Si 6 0 lg )]) and that Li is also present, the elements comprising the pegmatiticpneumatolytic and hydrothermal stages would be: Sn related to the pegmatitic-granitic elements, V, Nb, Ta, Mo, W and U (related horizontally Zr to Nb and Mo), vertically V, Nb, Ta subgroup elements of the V family and the elements Mo, W and U vertically related subgroup elements of the VI family of the periodic system. Furthermore, the elements Sn, Pb, Sb and Bi are interrelated as pointed out on page 246 (as mainly crust elements). In addition, the elements Mn, Fe, Cu, Ag and Au are interrelated as pointed out on page 230 and 251

belong to the Β group most probably as derivatives of mantle. Thus, typical interrelated pegmatitic-granitic elements might be in mineralogical association of pegmatitic or hydrothermal stage, but also typical mantle elements (i. e., Β group elements: Mn, Fe, Cu, Ag and Au) might be in mineralogical associations with both pegmatitic and vein deposits. However, the group Pb, Sb and Bi is usually in hydrothermal mineral associations. On the basis of the present study case, it can be concluded that pegmatites might contain ore minerals composed of both groups, i. e., of the group V, Nb, Ta, Mo, W and U (which is characterized in accordance with their interrelationships according to the empirical "laws" as pegmatitic-granitic group of elements) and Fe, Μη of the mantle derived group (B group: Mn, Fe, Co, Ni, Cu, Zn, Cd, Ag, Au). In this connection Li, Be, Β and (Al) could be considered as elements interrelated in accordance with the empirical "laws" of the periodic system and most probably "pneumatolytically" mobilized (or in solutions). The transition from pegmatitic-pneumatolytic and finally to hydrothermal phase can be understood in the way that each phase may contain elements from the following groups which have been recognized according to relationships on the basis of the empirical "laws" of the periodic system: (i) (Sn), V, Nb, Ta, Mo and U; (ii) Sn, Pb, Sb; (iii) Mn, Fe, Cu, Ag, Au; (iv) Li, Be, B? Al. It should be noted that these elements comprise the main ore minerals of the Sierras Pampeanas mineralization (Precambrian). Thus, the mobilization of elements from these groups is of great importance for the understanding of the transition-stages from pegmatitic-pneumatolytic to hydrothermal. The geochemical concept of paragenesis based on the interrelationships of the elements is of significance for the understanding of the common segregation (of rare elements), however, the crystallization of minerals will be influenced by factors (i) and (ii) as pointed out in Chapter 47. A complex pegmatitic paragenesis containing Be is described by Lawrence and Markham (1963), in "The petrology and mineralogy of the pegmatite complex at Bismuth, Torrington, NSW." According to them, the pegmatitic complex consists of an elongated intrusion of a granitoid quartz-topaz (topaz: A12(F2/Si04)) rock (silexite) together with a series of pegmatites of various composition. The principle pegmatite consists of orthoclase, biotite, quartz and beryl with concentric zoning passing outwards into fine-grained biotite-beryl rock containing a number of ore minerals: arsenides of Co, Fe, Ni, wolframite ((Fe, Mn)W0 4 ), bismuth, bismuthinite (Bi2S3), molybdenite (MoS2), joseite (Bi 4 (TeS) 3 ), rutile (Ti0 2 ), cassiterite (Sn0 2 ), uraninite (U0 2 ) and monazite (Ce[P0 2 ]), and beryl 252

(Al 2 Be 3 (Si 6 O u )) as mentioned. Furthermore, according to Lawrence and Markham, small pegmatite veins issuing from the main body contain high temperature tetrahedrite (Cu2SbS2), chalcopyrite (CuFeS 2 ) and sphalerite (ZnS). Considering the composition of the ore minerals reported (including beryl and monazite) the following groups of interrelated elements can tentatively be recognized: (i) Sn, As, Sb, Bi; (ii) Mn, Fe, Co, Ni, Cu, Zn; (iii) Mo, W, U (subgroup element of the VI family vertically interrelated) and interrelated to Hf and Ce. Ti is also tentatively put into this group (see page 247). (iv) Be (related to Li and B); F is also considered to be a "pneumatolytic" element. However, in the paragenesis reported by Lawrence and Markham, only Be and F are present. Also in this complex pegmatitic paragenesis, it is noted that group (i) (crust elements) and group (iii) (Mo, W, U, Ce, Ti) of the granitic-pegmatitic geoenvironment are represented (see page 232). Furthermore, the mantle derived elements (ii) are present as well in association with minerals of this pegmatite complex. As mentioned, Be is also present. In contradistinction to the complex pegmatite paragenesis from Sierras Pampeanas as described by Stoll, a different type of Be mineralization associated with greisen formation, is reported by Hawley (1969) from Lake George (or Badger Flats) beryllium area, Park and Jefferson Counties, Colorado. According to Hawley, the area on the south-west side of Front Range is underlain mainly by the Precambrian Idaho Springs formation, a granodiorite and various granites. The main mineral resource is Be ore, which has been mined (principally at the Boomer Mine) from veins, pipes, pods and complex irregular bodies encased in greisen. Furthermore, the greisen and associated veins also contain local concentrations of W, Mo, Pb, Zn, Ag and fluorite. Also in this case, there is a transition from greisen Be deposits to vein deposits. Considering the composition of the main ore minerals, the following groups of interrelated elements are recognized: (i) Pb belonging to the group: Sn, As, Bi (mainly lower crust/crust elements); (ii) Fe, Zn, Ag as Β group elements (Fe/Μη, Co, Ni, Cu, Zn, Cd, Ag, Au, see page 230); (iii) W and Mo (vertically interrelated subgroup elements of the VI family of the periodic system). (iv) Be and F (belong to the "pneumatolytic" phase and related to greisenization). As can be seen by comparing the related groups of elements in both Be parageneses, namely that of the pegmatitic complex, Torrington, New South Wales, Australia, described by Lawrence and Markham, and that of the Boomer Mine, Colorado, USA, described by Hawley, comparable element segregation processes have taken place where interrelated element groups from crust, granitic-pegmatitic geoenvironments and

mantle-derived elements form paragenetically related mineral associations. In contrast to the cases of Be mineralization discussed so far, Gies (1970) on the basis of analytical work and considerations, supports that the cobalt ores from Siegerland (Germany) consist of danaäite with 5.2 to 9.62% Co and glaucodote with 20.2 to 30.8% Co, also Co-arsenopyrite (Fe(Co)AsS) is reported. What is very interesting though, is that the following trace elements Ag, Be, Bi, In, Li, Mn, Mo, Ni, Pb, Sb and V were determined in 25 analyzed samples. Furthermore, considering that the cobalt ores of Siegerland are hydrothermal, the segregation of elements occurs in hydrothermal veins despite the diverse derivation for the cobalt ores reported by Gies, the following groups of interrelated elements are tentatively recognized: (i) As, Sb, Pb, Bi, In (interrelated elements, see page 246); (ii) Mn, (Fe), Cu, Ni (interrelated elements, Β group, most likely mantle derivatives, see page 229); (iii) Mo, V (interrelated elements of the graniticpegmatitic geoenvironment, see pages 232, 256); (iv) Be, Li (interrelated elements, most probably "pneumatolytically" segregated?). It seems noteworthy though, that in this case too, the reported groups of segregated elements are comparable to the other cases of Be mineralizations that have been studied.

(c)W- Vein Mineralization Element Segregation

and Associated

In contradistinction to the cases of W segregation in pegmatitic mineralization, complex element segregation can also take place in hydrothermal W-vein mineralizations. Bellindo and Manrique (1954), in "Geologia de los yacimientos de tungsten de Mundo Nuevo y La Victoria", report that "the tungsten deposits of the Mundo Nuevo (Tamboras) area, Cachicadän district, La Libertad department, northern Peru, are mesodermal or leptothermal fissure fillings cutting Upper Jurassic shales and Lower Cretaceous quartzites and sandstones". According to them, the emplacement of intermediate and acidic minor intrusions, probably representing the extremities of an extensive granitic batholith, closely preceded and was intimately related to the mineralization. Furthermore, deposition of quartz, pyrite, wolframite ((Fe,Mn)W0 4 ) preceded and followed that of sphalerite (ZnS), galena (PbS), tetrahedrite (Cu 2 SbS 2 ) and chalcopyrite (CuFeS2). Less common minerals recorded from the mineralized area include hübnerite (MnW0 4 ), ?ferberite (FeW0 4 ), realgar (AsS), orpiment (As2S2), arsenopyrite (FeAsS), pyrrhotite (FeS), marcasite (FeS2), polybasite (Ag 1 6 Sb 2 S n ), freibergite (Cu-Ag-Sb Fahlerz), bournonite (CuPbSbS3), boulangerite (Pb 5 Sb 4 S n ),

chalcocite (Cu2S), bornite (Cu 5 FeS 4 ), covellite (CuS), Au, Ag, Sb and barite. Considering the comparison of the minerals forming this complex mineralization, the following groups of related elements (in accordance with the empirical "laws" of the periodic system) are identified: (i) Pb, Sb, As (representing ?lower crust/crust elements and interrelated as described on page 229); (ii) Fe, Mn, Cu, Zn, Ag, Au (representing the Β group and horizontally interrelated as subgroup elements; most probably mantle derivatives); (iii) W (representing elements characterizing the granitic-pegmatitic geoenvironment). As is indicated by the elements comprising the minerals of the W vein mineralization, elements of crust derivation and elements of mantle derivation are jointly segregated and form mineral associations characteristic of this mesothermal-leptothermal pangenesis. In addition, Kantor (1965), in "Tungsten in the West Carpathian metallogenetic province" reports that the area has generally been considered to contain little W, but finds of scheelite (CaW0 4 ) and wolframite ((Fe,Mn)W0 4 ) from ores of Au, Ag, Pb-Zn-Cu and stibnite (Sb2S2) at several localities are reported. Considering the composition of the minerals comprising this mineralization, the following groups of related elements are tentatively recognized: (and interrelated as described;) (i) Sb, Pb (representative of ?lower crust/crust elements, see page 229); (ii) Fe (Mn), Cu, Zn, Ag, Au (B group elements; mantle derivatives, see page 230); (iii) W (representative of the pegmatitic-granitic geoenvironment elements, see page 232). In this case too, elements of diverse derivation are also jointly segregated to form mineral associations of these vein mineralizations.

(d) Mo-Vein Mineralization Element Segregation

and

Associated

As in the case of the W vein mineralization, Mo vein mineralization and element segregation are comparable and commensurable. Moore et al. (1966), in "Distribution of selected metals in the Stockton district, Utah" report that largely coextensive areas of locally high Bi-Cu-Mo concentration define a central zone; Pb-Zn areas of greater lateral extent partly overlap the central zone; locally high As-Sb concentrations characterize an outer zone with Β extending beyond the limits of known significant mineralization. Ag is erratically distributed. Hypogene geoenvironmental controls are indicated. From the elements reported, the following groups of related elements are tentatively distinguished: (i) As, Sb, Pb, Bi (representative of ?lower crust/crust elements, see page 229); 253

(ii) Cu, Zn, Ag (representatives of the Β group elements, most probably mantle derivatives); (iii) Mo (representative of the pegmatitic-granitic geoenvironment elements, see page 232); (iv) Β (representative element of the "pneumatolytically" mobilized elements). Furthermore, Miller and Elliott (1969) in their paper "Metalliferous deposits near Granite Mountain, eastern Seward Peninsula, Alaska" report deposits of Pb, Zn and Ag. Deposits of Mo, Bi and Ag are also recorded from a previously reported U, Cu, Pb and Zn deposit northeast of Granite Mountain. Both groups of deposits are spatially related to felsic plutonic rocks.

254

The following groups of related elements are identified: (i) Pb, Bi (representatives of the crust element group, see page 229); (ii) (Fe) Cu, Zn, Ag (representatives of the Β group elements, most probably mantle derivatives); (iii) Mo, U (representatives of the pegmatitic-granitic geoenvironment elements, see page 232). Also in this case, elements of diverse derivation are jointly segregated to form mineral associations of these vein mineralizations.

Chapter 58

Element Segregation/Distribution in the Manganese Parageneses

The theoretical scheme of element groups interrelated in accordance with the empirical "laws" of the periodic system, namely: group A = As, Sb, Bi, Pb, Sn, Se, Te; group Β = Fe (Μη), Co, Ni, Cu, Zn, Cd, Ag, Au, Hg (Cr, V, (Ti?)); group C = Zr (Hf), Ti, Nb, Ta, Mo, W, U; should be considered as an expression of the relationships of the elements forming characteristic manganese mineral parageneses. As it has been supported before, the theoretical scheme presents elements that might be represented in a manganese paragenesis. Considering that the element group A is thought to be crust derived, group B, in contrast, mantle derived and furthermore group C related to the granitic-pegmatitic geoenvironment, it is perhaps possible to see the elements comprising the Mn paragenesis as derivatives from those sources. In an attempt to interpret the segregation (?distribution) of elements comprising various manganese parageneses, study cases will be considered. According to Bourgignon and Toussaint (1954), the haematites of the Ardennes leave a black streak on porcelain when they are particularly rich in Mn. According to them, samples from Bihain gave the following results: Sample 1 = Fe: 39.6; Mn: 17.64; Si0 2 : 9.4; CaO: 0.5; P: 0.5 Sample 2 = Fe: 42.3; Mn: 16.84; Si0 2 : 5.8; CaO: 0.3; P: 0.3. It is interesting that the analyzed haematites show a strong presence of Mn. Considering that Mn and Fe are interrelated in accordance with the empirical "laws" of the periodic system as subgroup elements next to each other (horizontally; subgroup elements of the VII and VIII family), their common segregation (distribution) as mantle derivatives 30 is understandable. Interesting indirect support of the interrelationship between Fe and Mn is presented by Taylor (19101968) in his contribution "Sedimentary ores of iron and manganese and their origin". The distribution of Fe and Mn ore in time and space is briefly reviewed in his 30

It should though be admitted that in general the derivation of Fe from mantle is arbitrary since it may equally well be a derivative of the initial earth crust and recycled due to the unfolding of the geological spiral.

paper and consideration is given to the question whether significant differences exist between Precambrian and later deposits. He maintains that present day formation of ferruginous and manganiferous sediments show that all the known occurrences of chamosite? lie within 10° of the equator. Supporting the derivation of Mn and Fe from basic-ultrabasic source, Fyfe and Reed (1959) report small pockets of high-grade, but non-economic manganese ore, mainly braunite [3(Mn,Fe) 2 ) 3 MnSi0 3 ] occurring with basaltic lava and jasper intercalated with Cretaceous sedimentary rocks. Furthermore, as a corollary to the derivation of Mn from basic rocks, mainly buried intercontinental rift basalts, are the Red Sea region manganese deposits which are believed to be due to leaching of Mn from basalts (Tooms, 1976; Augustithis, 1982). In addition to the occurrence of Mn and Fe representing the mantle derived Β group (Mn, Fe, Co, Ni, Cu, Zn, Cd, Ag, Au, Hg (Cr)), additional study cases will be considered supporting the interrelationship of Mn with elements of the Β group (mantle derivatives). Hewitt (1968), in "Silver in veins of hypogene manganese oxides" supports that in southwestern USA and eastern Canada hypogene manganese oxides containing appreciable amounts of base metals occur. Furthermore, Hewitt reports Ag-bearing sulphides below a barite-manganese deposit in Nova Scotia. Dzhumaylo et al. (1966) describe manganese deposits occurring at the north end of the Urup chalcopyrite (CuFeS2) deposit (northern Caucasus) in mottled red tuffaceous silt stones of rhyolitic composition. According to them, manganese occurs in piemontite (Mn-epidote), bixbyite [(Mn,Fe) 2 0 3 ], sitaparite (bixbyite with Ca partly replacing Mn 3 ) and rhodochrosite (MnC0 3 ). Ferric iron and manganese contents increase together. According to Dzhumaylo et al. the occurrence is considered to be exhalative sedimentary. In this case, Mn, Fe and Cu of the mantle derived Β group are represented. Prinz (1967) reports that "the replacement deposits rich in Mn are irregularly distributed in favourable host beds along with Ag and Zn bearing quartz-veins", and in this case the reported elements Mn, Fe, Ag, Zn are representatives of group Β (mantle derivative). Comparable is the element segregation (distribution) of primary metamorphic Mn ore (a gondite) and of manganese veins occurring at Mount Branduten, 255

Botnedal, southern Norway. According to Westerveld (1961), a vein is massive and consists mainly of coarsely crystalline rhodonite (CaMn4[Si5)15]) and spessartite (Mn 3 Al[Si0 4 ] 3 ) with lenses of Mn oxides. The main Mn ore mineral is braunite [3(Mn,Fe) 2 0 3 ] and it is associated with some hausmannite (Mn 3 0 4 ) and jacobsite (MnFe0 4 ). Other subordinate minerals are fluorite (CaF2), rhodochrosite (MnC0 3 ), pyrite (FeS2), haematite (Fe0 3 ), and native Cu. Considering the composition of the main ore minerals, Mn, Fe and Cu are representatives of the Β group (mantle derivatives). Crittenden et al. (1961), in "Manganese deposits in the Drum Mountain, Juab and Millard counties, Utah", report that manganese carbonate ore was formed by hypogene hydrothermal replacement of impure dolomite (Cambrian) where pre-ore faults that served as channels for mineralizing solutions cut the carbonate rocks. According to them, the ore bodies are tubular and elongated parallel to the intersection of the host bed and controlling fault. Fissures locally have veins in which rhodochrosite is associated with base metal (Cu, Zn) sulphides. Manganese minerals include pyrolusite (ß-Mn0 2 ), 'wad' psilomelane ((Ba,H 2 O)Mn s O 10 ), rhodochrosite, manganoan calcite, and manganoan dolomite. In addition to Mn and Fe, base metals mainly belonging to group Β elements and derivatives of basic rocks or mantle are also reported. In this case too, the main metallic elements are interrelated in accordance to the empirical "laws" of the periodic system (i. e., Fe, Mn, Cu, Zn...) In contrast to the cases discussed so far, where the main metallic elements paragenetically associated with Mn were the elements of the Β group (Fe, Co, Ni, Cu, Cd, Zn, Ag, Au) mainly basic rocks or mantle derivatives, cases are reported where representative elements of the A group (Sn, As, Pb, Sb, Bi, Se, Te - most probably ?lower crust/crust derivatives) may be represented in the metallic elements forming the mineral pangenesis of manganese deposits. Kamp (1953) reports that in both the Pb-Ag-Zn deposits of Mashcan, department of Junin, Peru, and the Pb deposit of Mashcan, department of Pasco, Peru, vein filling and mantos of pre-mineralization of Mn oxides are associated with, and replaced by high-grade concentrations of oxidized Pb minerals. According to Kamp, there is no evidence that the latter have resulted through replacement of galena, and it is suggested that the manganese oxides have caused the oxidation of the hypogene ore solution. Considering the reported metallic elements, in addition to Mn, Fe, Zn and Ag (representatives of the mantle derivative group), Pb 31 as a representative of the A group is also present in paragenetic mineral association with the manganese deposits. 31

As mentioned in Chapter 39, isotope analysis could provide indicative information concerning whether Pb is mantle or crust derivative.

256

In addition, Boström (1964) in "Some aspects of the analysis of mineral-forming conditions" considers stability conditions of native lead and arsenides in some manganese ores. He also considers the reaction conditions between PbS - Mn 3 0 4 - CaC0 3 - PbC0 3 - BaC0 3 - MnC0 3 . It is further reported that minerals like native lead, pyrochroite (Mn(OH) 2 ) and hydrocerussite (Pb 3 [0HC0 3 ] 2 ) can be formed by redox processes of this complex system. Here again, when considering the main metallic elements present in this theoretical manganese paragenesis (which is compared to the native lead-pyrochroite occurrences of Längbau and Franklin), it can be seen that in addition to Mn, Fe (and Cu, Zn) belonging to the mantle derivative group B, also As and Pb as representatives of group A are present in association with minerals of the manganese mineral paragenesis. In contrast to the cases discussed, Augustithis (1982; see Chapter 51), considering the mineralization and geochemistry of the Abu Dabbab apogranite, Eastern Desert, Egypt, reports that granitophile elements (i. e., elements of the granitic pegmatitic geoenvironment Nb, Y, Ta, Zr and REE, together with Sn and ?Te, represent a granitophile Precambrian mineralization associated with the apogranite formation) and that in contrast Mn, Fe, Cu, Zn, Ni, Cd, Ag are derivatives from intercontinental rift basalts buried in the Great Rift. Augustithis considers the manganese mineral paragenesis as superimposed on the apogranite (granitic pegmatitic). From the composition of the reported minerals of the Abu Dabbab apogranite the following element groups can be recognized: A = Sn, Te (belonging to A group and are ?lower crust/crust derivatives); Β = Mn, Fe, Cu, Zn, Ni, Cd, Ag (mantle derivatives, in this particular case, by leaching of buried intercontinental rift basalts geoenvironment), and C = Y, Ta, Nb, Zr and REE (derivatives of the granitic-pegmatitic geoenvironment). When the wide gamut of geochemical and mineralogical possibilities for the formation of mineral p a n geneses is further studied, two additional study cases will be quoted, involving V and Ti in manganese mineral pangeneses. Ljunggren (1958) reports that at Bölet, Undenäs, in southern Sweden, a manganese mineralization has taken place, localized in the brecciated zones in granitic bedrock. Most of the manganese ore of Bölet consists of almost pure manganite [MnO(OH)] but a partial replacement of manganite by pyrolusite (ß-Mn0 2 ) is not unusual. The ore also contains calcite, barite, fluorite and in small portions vanadinite (Pb 5 [Cl/V0 4 ) 3 ]) and rhodochrosite (MnC0 3 ). Considering the main metallic elements of the main ore minerals in addition to Mn, (Fe) representing the mantle derivative Β group, also Pb representative of the A group is present in addition to V.

The derivation of V is dubious since it can be considered to be interrelated to the mantle group Ti, V, Cr, Μη (horizontally interrelated subgroup elements) or related vertically to V, Nb, Ta (as elements belonging to the granitic-pegmatitic geoenvironment, see page 232). Furthermore, Petruk (1963) reports the occurrence of manganite [MnO(OH)] and pyrolusite (ß-Mn02) as pods of manganese in a breccia along a shear zone at the contact between Cretaceous shales and sills and an intrusive basalt on Queen Charlotte Islands. Closely associated minerals, according to Petruk, are hausmannite (Mn 3 0 4 ), maghemite (y-Fe203), ilmenite (FeTi03) and haematite (Fe 2 0 3 ). Besides the main metallic elements on the basis of the reported minerals in addition to Fe and Mn (representative of the mantle or basic rock derivatives),

Ti is also present. As it was the case with V, in this case also with Ti occurring in the manganese paragenesis, the element interrelationships are dubious. Since, though, the manganese metallogeny is related to a basaltic intrusion, the possible theoretical relationship Ti, V, Cr, Μη, Fe (horizontally interrelated subgroup elements) derivatives from basic-ultrabasic rocks or mantle, is perhaps of greater significance. Also the chemical interrelationship of Ti and Fe is supportive of this interpretation. In addition to the consideration of the metallic elements building manganese parageneses, the mobilization of elements should be seen in conjunction with the transformation and alteration-oxidation of manganese minerals as reported in Chapter 42.

257

Chapter 59

Trace Elements in Sulfides (Compatible with a Joint Segregation of Elements in Accordance with the Empirical "Laws" of Element Interrelations)

The significance of trace elements in ore minerals, especially when solving petrogenetic problems is a vast topic with a plethora of contributions (e. g., The significance of trace elements in solving petrogenetic problems (1983), ed. S. S. Augustithis, Theophrastus Pubs.)· In addition to trace elements, isotope geochemistry has also contributed enormously to the international literature on this subject. Despite the fact that the geochemistry of trace elements in ore minerals and parageneses is a topic of great significance for the genesis of deposits, only a few selected study cases will be quoted in conjunction with the common concentration (segregation of elements/element distribution) to build paragenetic mineral associations. The selected examples will refer to the common sulphide minerals, pyrite, sphalerite and galena. The possibility that similar geochemical interrelationships prevail in the case of other ore minerals is open to further research. Badalov (1965), in "The role of major components in the geochemistry of minor and rare elements of ore deposits" reports that "determination of average contents of Au, Ag, Se and Te in disseminated Cu, Mo deposits of the Almalyk district, USSR, shows that the great bulk of these elements resides in abundant but rather low-grade minerals like pyrite (while certain other minerals, which concentrate the minor elements to a higher degree, are not sufficiently abundant to be dominant carriers". Considering the trace elements present in the pyrite from Almalyk, it can be suggested that pyrite contains Fe, Cu, Au, Ag - elements representative of the mantle derived Β group: Fe (Μη), Co, Ni, Cu, Zn, Cd, Au, Ag, Hg (see page 230) - and Se and Te as representatives of the lower crust group Sn, Sb, As, Pb, Bi, Te, Se, In, Ga, Ge (see page 230). Furthermore, Wright (1965), in "Syngenetic pyrite associated with a Precambrian iron ore deposit" gives an account of the geology, mineralogy and geochemistry of pyrite zones lying above the steep Rock Lake iron ore deposit in northwest Ontario, Canada. The zones contain about 50% pyrite which occurs as elongated lenses and according to Wright, approx. 90% of this pyrite consists of fragments with colloform and cryptocrystalline textures. A trace element study again 258

by Wright, reveals a marked paucity in the number and quantity of trace elements present in the zones. Combined with a low Co content, low Co/Ni ratios, presence of As in the pyrite, and negligible Se suggests a sedimentary rather than a hydrothermal origin. However, considering the trace elements in the pyrite, additionally to Fe, Co, Ni are elements representative of the Β group as well as As and Se of the lower crust/crust group (see page 230). In contrast, Yamaoka Kazuo (1958) in his contribution "Spectrographic studies on the trace elements in pyrite", suggests that petrogenetic elements, probably derived from the country rocks, are more or less uniformly distributed as trace elements in pyrite. According to Yamaoka, pyrite from the cupriferous pyrite deposits of Besshi-type generally contains Co in uniform distribution. Furthermore, trace elements suggest the character of the mineralization of some metallogenic provinces; e. g., Sn and Sb in pyrite from the Sn province on the periphery of the Sobo-Katamuki mountain block. In addition, Yamaoka supports that some trace elements are controlled by the temperature of the mineralizing fluid; Sb and Te are frequently recognized in pyrite from low temperature ore deposits. However, in this case too, the trace elements of the pyrite follow the general pattern of joint segregation of elements. Fe, Co and Cu are representative of the mantle derived Β group and the elements Sn, Sb and Te are comparably representatives of the lower crust elements (see page 230). Comparable considerations of the trace elements reported from study cases on the trace elements in sphalerite also support a relationship between trace elements, and their interrelationship is in accordance with the empirical "laws" of the periodic system. Studies by Takimoto and Minato (1961) on the relationship between temperature-type of deposits and distribution of minor constituents in sulphide minerals showed that the ratios of S/Se, Zn/Mn, Zn/Cd and Zn/Fe investigated, were in relation to the type of ore deposit and particularly to the temperature-type. According to them, it was found that Se concentrates in high temperature deposits. Mn concentrates in low temperature deposits at the Chichibu mine (Saitama Prefecture, Japan), while Cd shows a reverse relation-

ship. Fe shows concentration in high temperature deposits at Kamioka mine but no variation at Chichibu mine; Fe increases as Mn increases, while Cd decreases as Fe and Mn increase. However, consideration of the elements in these sphalerites shows that Zn, Fe, Cd and Mn are representatives of the mantle derived Β group, and Se is representative of lower crust/crust group (see page 230). Studies of minor elements in sphalerite by Muta (1960) show that the main minor elements detected in sphalerite from 120 mines all over Japan, are Ag, As, Au, Bi, Cd, Co, Cu, Fe, Ga, Ge, Sb, Sn, Mn, In, Ti, V, Si, Al, Mg and Ca. In contradistinction, sphalerites from xenothermal deposits show high concentrations of the minor elements Fe, Mn, In, Sn, Sb, Bi and As, Co and low concentrations of Ge, Ga, Ag and Hg. Considering the reported minor trace elements by Muta, the following groups of interrelated elements are recognized: Ag, Au, Cd, Co, Cu, Fe and Mn representative elements of the mantle derived Β group; Fe, Mn, Co, Ni, Cu, Zn, Cd, Ag, Au, Hg and As, Bi, Ga, Ge, Sb, Sn and In representative of the lower crust/crust group (Sn, As, Bi, Pb, Sb, Se, Te, Ga, Ge, In). In addition, the elements Ti, V reported by Muta are horizontally interrelated in accordance with the empirical "laws" of the periodic system. The elements Si, Al, Mg, Ca may be incorporated in the sphalerites from the petrogenetic minerals of their geoenvironment. Further studies on sphalerite by Noväk and Kvaccek (1964) support that the sphalerite from the Turkank zone in Kutnä Hora, Czechoslovakia, has an average crystallochemical formula (calculated) of (Zn 0 7 7 7 Fe 0 2 1 Mn 0 0 0 6 In 0 0 0 0 7 S 1 0 0 ). The microchemistry shows Mn, Cd > Sn, Cu, In, Pb, As > Ag, Sb, Bi > Hg, Ga, Co, Ni > Ge. Concerning the reported elements in these sphalerites, Zn, Fe, Cd, Mn, Cu, Ag, Hg, Co, Ni are representatives of the mantle derived Β group. In, Sn, Pb, As, Sb, Bi, Ga and Ge are representatives of the lower crust/crust elements. Additional investigation by Takahashi et al. (1961) of sphalerite from the Oppu mine, Aomori Prefecture, Japan, showed high contents of In, Sn, Bi and Co, and low contents of Ge and Ga. Considering the elements reported, Zn and Co represent elements of the mantle derived Β group, In Sn, Bi, Ge and Ga belong to the lower crust group (see pages 230, 247). Similar minor and trace element studies on galenas show comparable gecfchemical patterns. According to Oftedahl's (1959) geochemical studies of Norwegian galenas, the galenas of the Oslo district appear to be on average higher in Te than those from other areas. The Te values are comparable to TI, Bi and Sb. As Oftedahl states, Te entry supports high temperatures. Furthermore the element association Te, Bi, Pb, TI is characteristic of a sulphide environment. Considering the elements reported in these galenas (PbS), no representatives of the mantle derived group is found. In contradistinction, the elements reported

(Te, Bi, Pb, TI) are all representatives of the group Sn, As, SB, Pb, Bi, Te, Se, In, Ga, Ge and TI. Extensive geochemcial studies by Takahashi and Ito (1961) on galenas from Inner Northeast Japan and Southwest Hokkaido, Japan, metallogenic provinces show that Sb, Ag and Bi are most abundant and that As, Sn, TI, Cd, Se and Te are often present. Considering the elements reported in these galenas, Zn, Cd and Ag are representative elements of the mantle derived Β group, and in contradistinction, the additional elements detected in these galenas Sb, Bi, As, Sn, TI, Se and Te represent the lower crust/crust group (see pages 230, 247). As a corollary to the trace element studies on the galenas, Marshall and Joensuu (1961) support that "comparison of crystal habit and quantitative spectrographic determinations of Ag, Bi, Cu and Sb in galenas of the Upper Mississippi Valley district, the Picher field (Oklahoma, Kansas) and southwestern United States, reveals an enrichment of Sb in cubic as compared with octahedral galenas within a given district". Furthermore, they state that temperature appears to be a dominant factor in control of both habit and trace element content. It is thus clear that in addition to the interrelationship of the trace elements among themselves and with major elements of the minerals in which they occur, the crystal habit (crystal lattice) also controls the incorporation of elements in a lattice as Goldschmidt has pointed out in the relationship between composition and crystal structure. However, it should be emphasized that also in this case of the galenas studied by Marshall and Joensuu, the reported elements Zn, Ag and Cu are representatives of the mantle derived Β group and Bi and Sb belong to the lower crust/crust group. The recognition of the groups Fe (Mn), Ni, Co, Cu, Zn, Cd, Ag, Au, Hg (mantle derivatives) and Sn, As, Sb, Pb, Bi, Te, Se, Ga, Ge, In, TI (lower crust/crust) which are represented either as main, minor or trace elements of the common sulphide minerals: pyrite, sphalerite and galena should be considered as a corollary to the common (joint) segregation of elements in accordance to their relationships as suggested by the empirical "laws" of the periodic system to form paragenetic mineral associations. In pyrite and sphalerite Fe and Zn are main elements and as a consequence of this fact, the group of mantle derived elements are abundant as trace elements in addition to elements of the lower crust/crust group derivatives. In contradistinction, in the case of galena where Pb is a predominant element, the elements of the group Sn, As, Sb, Pb, Bi, Te, Se, Ga, Ge, In, TI are more abundant as trace elements than the mantle derived group. However, the presence of Ag as a trace element in galena should not be ignored and its chemical relationship with Pb should also be stressed.

259

Chapter 60

Study Cases of Agents of Metal Transportation

As mentioned in Chapter 47, the segregation/distribution of metallic elements, in addition to their relationship in accordance with the empirical "laws" of the periodic system, also depends on the agents of metal transportation. Thus, from the plethora of experimental studies concerning the agents of metal transportation (i. e., the possibility to be in solution or any chemical combination and be transported), a few selected study cases will be presented, not to exhaust the subject but to present the pattern of the processes. Schröcke (1963), in "Einige Gleichgewichte pneumatolytischer Paragenesen" supports that "calculations based on reactions important in pneumatolytic deposition of tin ores, show that Ti, Si and Al cannot be transported in the magmatic gas phase as TiCl4, SiCl4 and A1C13, because of their unfavourable equilibrium constants and small partial pressure". According to Schröcke though, the partial pressure reaches an amount sufficient for the deposition of ores. Furthermore, the deposition of cassiterite is considered for near surface and deep intrusion bodies. In contrast, Beruskov and Kurilfhikova (1966), give the results of an experimental study, of the influence of chlorine, carbon dioxide, boron, silicic acid and fluorine (according to them, the chief components of the hydrothermal solutions in sulphide-cassiterite deposition) on tin solubility in sodic and potassic solutions under conditions corresponding to the hydrothermal process. For the formation of deposits of quartz-cassiterite type as well as of the sulphide-cassiterite type, the most probable form of tin transportation by hydrothermal solutions is the hydroxyfluorstannite complex of tin of the [SnF x (OH) 6 J type. Investigating the transportation of W, Sushchevskaya and Ivanova in "Composition of mineralizing solutions of some wolframite deposits of eastern Transbaykaliya (based on a study of gas-liquid inclusions)" support that the quartz-wolframite ore bodies crystallized at temperatures Cu 2 Te + FeS, are used to explain the presence of certain tellurides and selenides in the ores of the Noril'sk, Kola and Sudbury deposits. Considering further the elements mentioned in the above thermodynamic investigation, it should be mentioned that the elements involved in these reactions can be divided into two groups of interrelated elements in accordance with the empirical "laws" of the periodic system (see Chapters 47 and 59). (i) Sn, Sb, Bi, Te, Se (lower crust/crust elements); (ii) Fe, Co, Ni, Cu, Ag (mantle derived elements). In their consideration of the gamut of metallic element transportation, Lambet and Nikolayev (1962), in "The forms of occurrence of uranium in the waters of the Sea of Azov and some estuaries and rivers of the Azov-Black Sea basin", report that selective filtration and analysis indicate that uranium is transported in natural water in the form of the complex [U0 2 (C0 3 ) 3 ] 4 " or some other ion rather than in colloidal suspension. In contrast, Smellie and Laurikko (1984) reported in their paper "Skuppesavon, Northern Sweden: an uranium mineralization associated with alkali metasomatism", that "uranium, which has been transported mostly as uranyl carbonate complexes, has precipitated as uraninite which occurs as impregnations and only very rarely forms fracture infillings". From the experimental investigations and study cases quoted above, it can be suggested that in the common segregation (concentration)/distribution of metallic elements, in addition to the interrelationship of elements in accordance with the empirical "laws" of the periodic system, the agents of metal transportation are also of fundamental significance. In addition, it should be pointed out that in cases interrelated elements tend to be jointly transported by agents of transportation, or are subjected to similar transportation processes, since interrelated elements tend to possess comparable properties.

261

Chapter 61

Goldschmidt's 'Laws of Element Distribution' and the Empirical "Laws" of Interrelated Element Segregation (Metallic Element Concentration)

The third factor (see Chapter 47) controlling the segregation/concentration of metallic elements in paragenetic mineral association, is the concept of element distribution of Goldschmidt (1938; 1954). The crystallization of minerals depends on the physic-chemical conditions, i. e., temperature, pressure and crystallochemical factors (as most extensively presented in the work of Goldschmidt). The relationship between crystal structure and chemical composition (provided factors (i) and (ii) are taken into consideration - see Chapters 47 and 60), can be summarized as depending on: (i) the various proportions of various kinds of atoms (or ions) in the chemical formula. (ii) the relative size (radii) of the various kinds of particles, i. e., atoms (or ions) in a crystal. It is beyond the scope of the present volume to elaborate on these principles which are most extensively discussed in "Geochemistry" by Goldschmidt (1954). An approach, though, will be made to present extracts from Rosebaud (republished in 1992) and from Goldschmidt's "Geochemistry", in order to clarify what Goldschmidt's 'laws of element distribution' comprise in contrast to 'element segregation/concentration to form mineral paragenetic associations' as explained in Chapter 47-59 concerning the "empirical laws of element segregation/concentration". According to Rosbaud, "One of the problems in the history of the earth that Goldschmidt approached, is the partition of the chemical elements during the geochemical evolution between gas and coexisting liquid phases, the subsequent crystallization of these liquid phases, of molten iron, iron sulphides and fused silicates, and the distribution of the chemical constituents on these phases. According to their tendency to enter one or the other of these phases, the elements could be classified into siderophile (metal melt), chalcophile (sulphide melt) and lithophile groups (silicate melt). A fourth group, the atmophile group, includes those elements which in pregeological times differentiated into primordial atmosphere. It is of great interest that the distribution of the chemical elements among the three phases (iron, sulphide, silicate), is closely related to their atomic volume and the structures of their electronic shells. 262

The main controlling factors during these stages of differentiation are the chemical affinities of the various elements towards oxygen and sulphur and their latent heat of vaporization as compared with the affinity of the most common terrestrial heavy metal, iron. Goldschmidt took as a measure of these affinities the free energy of oxidation per gram atom of oxygen of the lowest oxides of the electropositive elements and the corresponding data for their sulphur compounds. Elements which are extremely rare in the earth's crust the lithosphere - are gold, the elements of the platinum group, and also nickel, cobalt and germanium. They can be expected in the iron phase, the siderosphere. On the other hand, the alkali and alkali-earth elements, silicon, aluminum, titanium, elements with a higher free energy of oxidation than iron, will, in the primordial differentiation have concentrated in the outer silicate crust of the earth". Furthermore, Rosbaud adds, "Goldschmidt attacked the problem of finding the general laws and principles of geochemistry 'from the viewpoint of atomic physics and atomic chemistry and to find out the relationships between the geochemical distribution of the various elements and the measurable properties of their atoms and ions'. He soon discovered that the principle factor regulating the entrance of the atoms and ions and the distribution of the rarer elements in the crystalline phases of igneous and metamorphic rocks was the size of the atoms and ions and not their weight. During the gradual crystallization of liquid solutions, those atoms or ions of rare metals are caught in the already existing three-dimensional lattice which, because of their size, fit into this lattice. Those atoms and ions which are either too small or too big to be caught remain in the liquid. This idea led to the discovery of the fundamental relationship between crystal structure and chemical constitution". Concerning the evolution of thinking which led to the "laws of element distribution" and the relationship between crystal structure and chemical constitution, the following extracts from Goldschmidt's "Geochemistry" seem to be of significance: "The task of crystal chemistry is to find systematic relationships between chemical composition and physical properties of the crystalline substances, and in

particular to find how crystal structure, i. e., the arrangement of atoms or ions in crystals, depend on chemical composition. From earlier times, it has been clear to the student of crystallography and the chemist that systematic relationships exist between chemical composition and crystal structure. The first step in the study of such relations was the observation by R. Hauy, in the last quarter of the eighteenth century that to any homogenous substance (or chemical individual) there belongs a certain geometrical complex of crystal faces which must be determined by the inner structure (or molecular arrangement) of that substance. The next step was the discovery by E. Mitscherlich in 1819, that substances which are analogous in chemical composition have in many cases a similarity of crystalline form also; the first examples studied were the two salts KHjPC^ and KJ^AsO^ In 1822 Mitscherlich discovered another fundamental phenomenon of crystal chemistry, viz. polymorphism, the faculty of certain substances to form different species of crystals, with different structure and different properties, under different conditions of temperature or pressure." Commenting on the relationship between composition and crystal structure, Goldschmidt continues: "In 1821 Mitscherlich, after having discovered isomorphism, formulated the result of his observations in the following manner: 'Eine gleiche Anzahl Atome, wenn sie in gleicher Weise verbunden sind, bringen gleiche Krystallformen hervor, und die Krystallform beruht nicht auf der Natur der Atome, sondern auf deren Anzahl und Verbindungsweise.' (The same number of atoms, combined in the same manner, produces the same crystalline forms, and the form of a crystal is not dependent upon the nature of the constituent atoms but upon their number and manner of binding.) In this way, Mitscherlich thought he had contradicted the idea that every chemical individual has its own characteristic crystalline form and structure. Today we know that the results of Hauy and Mitscherlich do not contradict each other, but are complimentary. Quite generally, the law of Hauy is valid: each chemically homogeneous substance has a crystal form (and a structural arrangement) of its own such that any chemical substitution causes an alteration of structural arrangement or at least of structural dimensions, and thereby of the crystalline form, or at

least of molecular volume. But if one limits the chemical substitution to a mutual replacement of atoms or atom groups (such as P 0 4 and As0 4 ) the properties of which are not very different, then the crystal structures (and the crystalline form) are not very much affected by the substitution, and the alteration of the external form of the crystals are not in many cases very conspicuous." Thus, after commenting the possibility of element substitution in the crystal structure, Goldschmidt points out that "it is most important, therefore, to find those properties of atoms, atom groups, ions and composite ions which are 'essential' for the crystalline structure of their compounds". Goldschmidt considers "ionic radii" as an essential property for the crystalline structure and he emphasizes that "one of the next problems is to find the 'size' of the ions. Generally the study of crystal structures of ionic crystals only furnishes data for the total distance between adjacent anions and cations. How this distance, for instance between chlorine and sodium in sodium chloride, is to be divided between the part belonging to chlorine and the part belonging to sodium is a separate problem". In order to solve this problem, he studies a series of compounds, "those with the chemical formula type AX2, such as fluorides of divalent metals ZnF 2 and SrF2, and the oxides of tetravalent metals, Ge0 2 , RU02, Ce0 2 , Th0 2 . It was clearly found by measuring the interatomic distances in these two series of crystals that the occurrence of the different types of structures is closely related to the size of the individual ions. Table VII gives the interatomic distances and radius ratios. Table VII reveals that a certain limiting quotients of ionic radii one type of crystal structure may be followed by another structure, each of them being conditioned by a certain range of radius ratio. Both for the dioxides and for the difluorides the limiting ratio between the fluorite structure and the rutile structure is about 0.7. If the proportion between the cation radius and the anion radius is in excess of 0.7 there will be 8 anions around each cation, whereas if the ratio is less than 0.7 but in excess of 0.4 we shall find 6 anions coordinated around each cation. If, however, the ratio is between 0.4 and 0.22 there is only space for 4 anions in contact with each cation: thus, the various Si0 2 structures are stable in this special range".

263

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Fig. 7. Chromite with a magnetite margin free of ilmenite ex-solutions. Ch = chromite. Μ = magnetite. S = platinoid mineral (often Ni, Co sulfides occur in this form). Dunitic pipe. Yubdo, W. Ethiopia. Polished section (with oil-immersion, one nicol).

W ^ I S S I β·

I C h

0.0125

Fig. 8. Sperrylite in intergrowth with magnetite (probably representing replacement of sperrylite by the magnetite). S = sperrylite. Μ = magnetite. Yubdo, W. Ethiopia. Oil-immersion with one nicol. χ 800.

Fig. 9. Intergranular martitized magnetite with sperrylite. S = sperrylite. Μ = magnetite. Yubdo, W. Ethiopia. Oil-immersion with one nicol. χ 800.

279

Fig. 10. Dunite, olivine predominant (with magnetite), crushed and invaded by mobilized antigorite "veinlets" which extend into the dunite following its cracks, ol = olivine, m = magnetite. an = antigorite. Arrow "a" shows antigorite invading the dunite along cracks. Serpentinized dunite. Xerolivado, Vourinos, N. Greece. Thin section, without crossed nicols.

0.1mm

Fig. 11. Rounded granular chromite with a margin of blastically grown uvarovite and intergranular chalcedony, c = chalcedony, ch = chromite. m-u = uvarovite marginal to the granular chromite. "Chromite-rich band". Bushveld, Igneous Complex, Bushveld, Transvaal, South Africa. Thin section, without crossed nicols.

Fig. 12. Chromite grains partly rounded and surrounded by uvarovite crystalloblastesis. Ch = chromite. Uv = uvarovite occupying the spaces between the chromite. "Chromite-rich band". Bushveld, Transvaal, South Africa. Thin section, without crossed nicols.

280

Fig. 13. Crystalloblastic magnetite enclosing pyroxenes and plagioclases. However, there is a phase-b pyroxene following the intergranular between plagioclase and magnetite. PI = plagioclase enclosed by the blastic magnetite. ΡΥ = pyroxene enclosed by the blastic magnetite. py-ig = pyroxene (phase-b) following the intergranular between pyroxene and magnetite (black). Average ferrogabbro (plagioclase-augiteolivine-magnetite "cumulate"). House area, Skaergaard, Greenland. Thin section, with crossed nicols.

Fig. 14. Plagioclase olivinefels enclosed by a late blastogenic magnetite phase which occupies the interleptonic spaces of the intergranular and corrodes olivine and pyroxene crystal grains. Ol = olivine. PI = plagioclase. m = mica, magnetite (black). Arrow "a" shows corroded outline of the plagioclase. Perpendicular feldspar rock (a peculiar eucrite with feldspars roughly perpendicular to margin of intrusion). 55 m from W. margin of Mellemo, Skaergaard, Greenland. Thin section, with crossed nicols.

Fig. 15 Plagioclase corroded and rounded by a late phase magnetite (black). A late phase pyroxene crystallization occupies the interleptonic spaces between the magnetite and the included plagioclase. PI = plagioclase. py = pyroxene (diopside), average ferrogabbro (plagioclase-augiteolivine-magnetite "cumulate"). House area, Skaergaard, Greenland. Thin section, with crossed nicols.

281

Fig. 16 Crystalloblastic pyroxene between corroded plagioclase and crystalloblastic magnetite. PI = plagioclase. Py = pyroxene. Μ = magnetite crystalloblast, averagt ferrogabbro (plagioclase-augiteolivine-magnetite "cumulate"). House area, Skaergaard, Greenland. Thin section, with crossed nicols.

Fig. 17 Pyroxene invaded by magnetite with extension of the ore mineral following the interleptonic spaces of the pyroxene cleavage, m = magnetite. py = pyroxene, arrow "a" shows the magnetite following the interleptonic cleavage spaces of the pyroxene. "Purple band" ferrogabbro (iron wollastonite-plagioclase-oli vinemagnetite-apatite "cumulate"). Skaergaard, Greenland. Thin section, with crossed nicols.

Fig. 18 Pyroxene including mica invaded and replaced by later iron oxides clearly indicating a later "infiltration" of the iron minerals in the pyroxene, py = pyroxene, ο = iron oxides, m = mica. Skaergaard, Greenland. Thin section with one nicol.

282

Fig. 19 Magnetite with pyroxene and a late generation of diopside occupying the interleptonic spaces between the magnetite and the pyroxene. Magnetite (black), py = pyroxene (augite). n-p = new generation of pyroxene between magnetite and pyroxene. Melanocratic average rock (plagioclase-pyroxene-ore "cumulate"). Foot of Boubouca Ridge at sea level. Skaergaard, Greenland. Thin section, with crossed nicols.

Fig. 20 Rounded magnetite (black) surrounded by blastic biotite; also crystalloblastic apatite is present, ma = magnetite, bi = biotite. ap = crystalloblastic apatite, ρ = ore pigments. Theralite. Alceria, Serra de Monchique, South Portugal. Thin section, with crossed nicols.

Fig. 21 Symplectic intergrowth structure of magnetite with hypersthene. The magnetite (black) transgresses from symplectic to more compact magnetite, m = magnetite, hy = hypersthene. Norite. Hitteroe, Norway. Without crossed nicols.

283

Fig. 22 Martitized magnetite with extensions attaining myrmekitic form in the pyroxenes. Yubdo, W. Ethiopia. Oil-immersion with one nicol.

χ 800.

Fig. 23 Spinel surrounded by an olivine-pyroxene crushed zone with marginal transitions to opaque iron oxides, s = spinel. i-o = iron oxides as opaque margins of the spinel, due to the thermal influence by the basaltic groundmass. c-z = crushed zone of pyroxene and olivine surrounding the spinel, bg = basaltic groundmass. Olivine bombs (mantle fragments) in basalt. Jato, Lekempti, W. Ethiopia. Thin section, without crossed nicols.

0.1mm

Fig. 24 Chromospinel xenocrysts in the groundmass of olivine basalt. The spinel xenocryst has a magnetite margin. Both the magnetite margin and the chrome-spinel show "myrmekitic intergrowths" as a result of the synantetic reaction of chromespinel-magnetite with the basaltic groundmass which has corroded and infiltrated into the metallic minerals, ch = chrome-spinel, m = magnetite margin, mi = myrmekitic intergrowth of basaltic melts and chrome-spinel magnetite. g = basaltic groundmass. Mantle fragments in basalt. Jato, Lekempti, W. Ethiopia. Polished section, oil-immersion with one nicol.

284

Fig. 25 Chromospinel with a margin of magnetite in contact with olivine crystal grain, components of olivine bomb in basalt. The basaltic melts show myrmekitic intergrowths with the magnetite and the chrome-spinel, g = olivine of "olivine bomb" and basaltic melt infiltrations, ma = magnetite margin, my = myrmekitic intergrowths of basaltic melts with chromospinel and magnetite, ch = chromospinel. Olivine bomb in basalt. Jato, Lekempti, W. Ethiopia. Oil-immersion with one nicol.

Fig. 26 Basaltic titanomagnetite (actually magnetite with an increased Ti content), t = titanomagnetite. The titanomagnetite shows alteration to maghemite. Olivine basalt. Debra-Sina, Ethiopia. Oil-immersion with one nicol.

Fig. 27 Zoned augite with broad polysynthetic lamellae with magnetite idiomorphic crystal grains, partly enclosed in the phenocryst and partly in the groundmass. A = twinned and zoned augite phenocrysts. Μ = idiomorphic phenocrysts (magnetite). G = basaltic groundmass. Augite-rich basalt. Selale Mountain region, Ethiopian Plateau. Thin section, with crossed nicols.

285

Fig. 28 Idiomorphic apatite surrounded by an aggregate of fine magnetites in basaltic groundmass. m = magnetite, a = apatite. Minderberg near Linz on the Rhine River, Germany. Thin section, with crossed nicols.

Fig. 29 Apatite sub-phenocrystalline in size, surrounded and corroded by magnetite. A symplectic pattern of apatite and the intracrystalline diffusion melts is also illustrated, a = apatite, m = magnetite. Basalt columns. Minderberg near Linz on the Rhine River, Germany. Thin section, with crossed nicols.

Fig. 30 Ilmenite with a "rounded" outline in kimberlitic groundmass. il = ilmenite. G = kimberlitic groundmass. Kimberlite. M'sipashi Plateau of Kundelungu, Zaire. Oil-immersion with one nicol.

Fig. 31 X-ray fluorescence spectroanalysis of rounded ilmenite from the tuffitic kimberlite of M'sipashi Plateau. Kundelungu, Zaire. Diagram by Professor A. Vgenopoulos.

Fig. 32 Magnetite replacing pyrrhotite marginally and along cracks of the pyrrhotite. m = magnetite ρ = pyrrhotite. Alexo Mine, Dundonald Township, Ontario, Canada. Oil-immersion with one nicol.

Fig. 33 Magnetite replacing pyrrhotite along fracture lines. Myrmekitic magnetite/pyrrhotite symplectite is also formed, m = magnetite ρ = pyrrhotite. s = magnetite/pyrrhotite myrmekitic symplectites. Silicate also present. Alexo Mine, Dundonald Township, Ontario, Canada. Oil-immersion with one nicol.

287

Fig. 34 Magnetite and titanite replacing ilmenite (with haematite oriented lamellae), m = magnetite, il = ilmenite. t = titanite. Otanmäki, Finland. Oil-immersion with one nicol.

Fig. 35 Magnetite replacing ilmenite (with haematite oriented lamellae) marginally and along cracks. m = magnetite, il = ilmenite. mc = magnetite extension replacing ilmenite along a crack. Otanmäki, Finland. Oil-immersion with one nicol.

mm

Fig. 36 Ilmenite replacing magnetite (with oriented spinel lamellae). Also spinel follows the contact magnetite/ilmenite. Chalcopyrite "invades" ilmenite and magnetite along cracks. il = ilmenite (atoll replacement of magnetite), m = magnetite, s = spinel, c = chalcopyrite. Grube Concordia, near O'okiep, South Africa. Oil-immersion with one nicol.

0.1 mm

288

Fig. 37 Titanomagnetite replaced by ilmenite. s = oriented spinel in the magnetite, si = spinel following the contact ilmenite-magnetite. c = chalcopyrite. il = ilmenite. Grube Concordia, near O'okiep, South Africa. Oil-immersion with one nicol.

Fig. 38 Crystalloblastic pyrite replacing magnetite. The relics of magnetite are apparently "oriented" within the later pyrite. In cases maghemite is associated with magnetite, ρ = pyrite. m = magnetite, mh = maghemite. Ryllshyttan, near Garpenberg, Sweden. Oil-immersion with one nicol.

Fig. 39 Tectonogranular magnetite replaced by intergranular pyrite. m = tectonogranular magnetite, ρ = pyrite. Dognacska, Banat, Croatia. Oil-immersion with one nicol.

289

Fig. 40 Tectonogranular magnetite replaced intergranularly by pyrite. Also large pyrite marginally replacing the tectonogranular magnetite is shown. m = magnetite (tectonogranular). ρ = pyrite intergranular replacing the tectonogranular magnetite. p-1 = large pyrite replacing the tectonogranular magnetite. Dognacska, Banat, Croatia. Oil-immersion with one nicol.

Fig. 41 Tectonogranular magnetite replaced along fractures by pyrite often occupying the intergranular spaces of the magnetite. Dognacska, Banat, Croatia. Oil-immersion with one nicol.

Fig. 42 Atoll replacement of magnetite by pyrite. m = magnetite, ρ = pyrite. Otanmäki, Finland. Oil-immersion with one nicol.

290

Fig. 43 Magnetite replaced by pyrite. m = magnetite, ρ = pyrite. s = spaces within the magnetite (silicates?) partly occupied by pyrite. Feldbach, Binnental, Switzerland. Oil-immersion with one nicol.

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ml·1 Fig. 44 Magnetite with maghemite replaced by pyrite; (maghemite also replaces magnetite), m = magnetite, mh = maghemite. s = symplectic pyrite. ma = marcasite. ρ = pyrite. Outukumpu, Finland. Oil-immersion with one nicol.

Fig. 45 Magnetite with ilmenite replaced by pyrite. Relic structures of the titanomagnetite (ilmenite) are present in the pyrite. m = magnetite, ρ = pyrite. r = relic structure of the titanomagnetite in the pyrite. Otanmäki, Finland. Oil-immersion with one nicol.

291

Fig. 46 Magnetite with oriented lamellae of ilmenite replaced by pyrite in which relics of the magnetite and ilmenite are preserved, m = magnetite. il = oriented lamellae of ilmenite. ρ = pyrite. r = relics of ilmenite in pyrite. Otanmäki, Finland. Oil-immersion with one nicol.

Fig. 47 Pyrite replacing titanomagnetite. The pyrite partly follows interlamellar spaces of the magnetite (between ilmenite lamellae), m = magnetite, il = ilmenite. ρ = pyrite replacing magnetite between ilmenite lamellae. Otanmäki, Finland. Oil-immersion with one nicol.

Fig. 48 Crystalloblastic pyrite replacing enargite, relics of which are preserved in the pyrite. e = enargite. re = relics of enargite in pyrite. ρ = pyrite crystalloblast. Otjosondu, Sandfeld, S. W. Africa. Oil-immersion with one nicol.

292

Fig. 49 Crystalloblastic pyrite replacing enargite. The pattern could be interpreted as enargite veinlets invading the crystalloblastic pyrite. However, relics of enargite are interspersed in the pyrite. e = enargite. re = relics of enargite in the pyrite. ve = apparent veinlet of enargite "extending into the pyrite". ρ = pyrite crystalloblast. Otjosondu, Sandfeld, S. W. Africa. Oil-immersion with one nicol.

Fig. 50 Pyrite replacing berthierite. ρ = pyrite. b = berthierite. Freiberg, Saxony, Germany. Oil-immersion with one nicol.

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Fig. 51 Atoll and marginal replacement of berthierite by pyrite. ρ = pyrite. b = berthierite. Freiberg, Saxony, Germany.

«HAift 0

0.05

293

Fig. 52 Berthierite marginally replaced by pyrite attaining cubic forms (idiomorphism). b = berthierite. ρ = pyrite. Freiberg, Saxony, Germany.

Fig. 53 Zoned cassiterite replaced by pyrite. c = zoned cassiterite. ρ = pyrite, arrow shows pyrite extending into the zoned cassiterite. Shizhu Yuan, near Chenzhou, Hunan Province, China. Oil-immersion with one nicol.

Fig. 54 Skeletal pyrargyrite (actually pyrargyrite occupying the spaces between developed gangue crystals) is replaced by Wasserkies. g = gangue crystals. Py - pyrargyrite. wp = Wasserkies pyrite. r = relics of pyrargyrite in the Wasserkies pyrite. Freiberg, Saxony, Germany. Oil-immersion with one nicol.

294

Fig. 55 Initial magnetite with ilmenite lamellae (titanomagnetite) replaced by chalcopyrite and pyrrhotite. Relics of the ilmenite lamellae are preserved in the chalcopyrite. c = chalcopyrite. pyr = pyrrhotite. il = ilmenite. Frood Mine, Sudbury, Canada. Oil-immersion with one nicol.

Fig. 56 Initial titanomagnetite replaced by pyrrhotite. Relics of ilmenite (initially associated with the magnetite) are preserved in the pyrrhotite. Frood Mine, Sudbury, Canada. Oil-immersion with one nicol.

Fig. 57 Magnetite marginally replaced by silicate and enargite. Haematite is also present, m = magnetite, g = gangue (silicate), e = enargite. h = haematite. San Cristobal Mine, Cerro de Pasco, Peru. Oil-immersion with one nicol.

295

Fig. 58 Magnetite replaced by adjacent chalcopyrite which sends veinform extensions into the magnetite, m = magnetite, c = chalcopyrite. Peko Mine, N. Territory, Australia. Oil-immersion with one nicol.

Fig. 59 Magnetite surrounded by chalcopyrite. Atoll-type replacement of the magnetite by chalcopyrite has taken place, m = magnetite, c = chalcopyrite. Peko Mine, N. Territory, Australia. Oil-immersion with one nicol.

Fig. 60 Magnetite replaced by boulangerite. m = magnetite replaced by boulangerite. b = boulangerite. Naica near Parrel, Mexico. Oil-immersion with one nicol.

Fig. 61 Rutile in intergrowth with silicate (black) marginal to chalcopyrite. The rutile/silicate intergrowth is invaded and replaced by bornite. c = chalcopyrite. r = rutile. b = bornite. Spectacel Mine near O'okiep, Namibia, S.W. Africa. Oil-immersion with one nicol.

Fig. 62 Rutile in intergrowth with silicate. The rutile/silicate intergrowth is invaded and replaced by bornite. r = rutile. b = bornite, silicate (black). Spectacel Mine near O'okiep, Namibia, S.W. Africa. Oil-immersion with one nicol.

Fig. 63 Bornite replacing pyrite. b = bornite. ρ = pyrite. Tilva Mica, Bor, Serbia. Oil-immersion with one nicol.

297

Fig. 64 Chalcopyrite replacing bornite which in turn is replaced by covellite and a network of fine anastomosing covellite veinlets extend into the bornite. c = chalcopyrite. b = bornite. cv = covellite. cl = chalcocite. Glava (West) near Arvika, Sweden. Oil-immersion with one nicol.

Fig. 65 Pyrite with idiomorphic tendency surrounded by copper minerals. Bornite replaced by chalcopyrite and both replaced by later formed chalcocite. ρ = pyrite. c = chalcopyrite. cl = chalcocite. b = bornite. Matchless Mine Farm, Friedenau, Namibia, S. W. Africa. Oil-immersion with one nicol.

,1mm

Fig. 66 Stromeyerite replacing bornite. s = stromeyerite. b = bornite. Aurora, Nevada, USA. Oil-immersion with one nicol.

298

Fig. 67 Chalcopyrite replacing enargite. e = enargite. c = chalcopyrite. S. Joäo Deserto, Aljustral, S. Portugal. Oil-immersion with one nicol.

Fig. 68 Tennantite marginally replaced by banded colloform chalcocite (also chalcocite veinlets transverse the tennantite). The colloform chalcocite is often followed by cuprite, t = tennantite. c = chalcocite. cu = cuprite vc = veinlets of chalcocite replacing tennatite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 69 Tennantite rhythmicallymarginally and in veinform replaced by chalcocite. Cuprite is also present marginal to the chalcocite. t = tennantite. c = chalcocite. cu = cuprite. vc = veinform replacement of tennatite by chalcocite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

299

Fig. 70 Pentlandite surrounded by chalcopyrite. Also chalcopyrite replaces pentlandite along fractures, ρ = pentlandite. c = chalcopyrite. c-1 = chalcopyrite lamellae replacing pentlandite. Scattered sulfides in norite. Falconbridge, Ontario, Canada. Oil-immersion with one nicol.

blend

Fig. 71 Pyrite and sphalerite (black) replaced by chalcopyrite. ρ = pyrite. c = chalcopyrite. s = sphalerite (black due to photocontrast developing), cs = chalcopyrite extending into the sphalerite. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

Fig. 72 Sphalerite marginally replaced by chalcopyrite and extensions of it attaining ex-solutionlike forms in the blende, s = sphalerite. cm = marginal chalcopyrite. e = chalcopyrite extensions of the marginal chalcopyrite in the blende attaining ex-solution-like forms of chalcopyrite intergrown with sphalerite. Supija Stujena, Montenegro. Oil-immersion with one nicol.

300

Fig. 73 Sphalerite transversed and replaced by chalcopyrite veinlets. Also pyrite replaces the sphalerite, s = sphalerite. c = veinlets of chalcopyrite and myrmekitic-like replacement of the sphalerite by the chalcopyrite. ρ = pyrite replacing sphalerite. Teufelsgrund, Münstertal, Black Forest, Germany. Oil-immersion with one nicol.

Fig. 74 Chalcocite replacing marginally and in veinform chalcopyrite. c = chalcocite. ch = chalcopyrite. ν = veinform replacement of chalcopyrite by chalcocite. North Broken Hill, Australia. Oil-immersion with one nicol.

Fig. 75 Pyrite (idiomorphic) surrounded with chalcocite which has veinform extensions into the adjacent chalcopyrite. ρ = pyrite. c = chalcocite. ch = chalcopyrite. Matchless Mine. Friedenau, Namibia, S. W. Africa. Oil-immersion with one nicol.

301

Fig. 76 Chalcopyrite marginally replaced by neodigenite which marginally transgresses into brown iron. Also veinlets of neodigenite replace the chalcopyrite. ch = chalcopyrite. η = neodigenite. b = brown iron. Gorob, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 77 Rhythmical interbanding of brown iron and neodigenite. η = neodigenite. b = brown iron. Gorob, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 78 Neodigenite and brown iron exhibiting colloform interbanding. b = brown iron, η = neodigenite. Gorob, Namibia, S. W. Africa. Oil-immersion with one nicol.

302

Fig. 79 Enargite replaced by haematite (atoll-type replacement), g = gangue mineral, e = enargite. h = haematite. San Cristobal Mine, Cerro de Pasco, Peru. Oil-immersion with one nicol.

Fig. 80 Chalcopyrite marginally replaced by covellite. ch = chalcopyrite. co = covellite. Naica, near Parrel, Mexico. Oil-immersion with one nicol.

Fig. 81 Chalcocite replaced by covellite. c = chalcocite. co = covellite. Bisbee, Arizona, USA. Oil-immersion with one nicol.

303

Fig. 82 Neodigenite replaced by covellite. The covellite replacement also extends into the neodigenite. η = neodigenite. co = covellite. ce = covellite extensions into the neodigenite. Bisbee, Arizona, USA. Oil-immersion with one nicol.

Fig. 83 Galena replaced by covellite. Tennantite and ?sphalerite are also shown, g = galena, co = covellite. t = tennantite. s = ?sphalerite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 84 Boulangerite marginally and in its central part replaced by covellite. b = boulangerite. co = covellite. Naica, near Parrel, Mexico. Oil-immersion with one nicol.

304

Fig. 85 Bornite replaced by covellite (also a fine network of covellite extends into the bornite). In the covellite, neocrystallization of chalcocite and chalcopyrite is also exhibited. be = bornite with a fine network of covellite. c = chalcocite neocrystallization in the covellite. ch = chalcopyrite neocrystallization in the covellite. co = covellite. Glava (West) near Arvika, Sweden. Oil-immersion with one nicol.

Fig. 86 Bornite replaced by fine veinlets of covellite. Also massive covellite replaces bornite. In the covellite, neocrystallizations of chalcopyrite have taken place along fractures of the initial bornite filled with covellite. b-c = bornite with a fine network of covellite. co = covellite. ch = chalcopyrite neocrystallizations. Glava (West) near Arvika, Sweden. Oil-immersion with one nicol.

Fig. 87 Nucleus of older colloform pyrite surrounded by chalcocite with marginal neocrystallization of pyrite. cp = colloform pyrite (older?), c = chalcocite. ρ = neocrystallizations of pyrite. Reichenberg Mine near Dens, Kreis Rothenburg, Germany. Oil-immersion with one nicol.

305

Fig. 88 Cuprite replaced by colloform malachite, c = cuprite, m = malachite. Viscachani, Vetas, Coro-Coro, Bolivia. Oil-immersion with one nicol.

Fig. 89 Cuprite marginally replaced by malachite, c = cuprite, m = malachite. Viscachani, Vetas, Coro-Coro, Bolivia. Oil-immersion with one nicol.

Fig. 90 Cuprite and idiomorphic cuprite transversed and replaced by malachite veinlets. c = cuprite, m = malachite. Livadi, Seriphos, Greece. Oil-immersion with one nicol.

306

Fig. 91 Cuprite replaced by native copper. cu = cuprite. nc = native copper. Peko, Northern Territory, Australia. Oil-immersion with one nicol.

Fig. 92 Tetrahedrite marginally replacing niccolite and sending extensions into it. t = tetrahedrite. η = niccolite. te = extensions of tetrahedrite in niccolite. Talmessi Mine, Anarek, Iran. Oil-immersion with one nicol.

Fig. 93 Tetrahedrite replacing niccolite in the marginal part of the niccolite. t = tetrahedrite. η = niccolite. Talmessi Mine, Anarek, Iran. Oil-immersion with one nicol.

307

Fig. 94 Germanite enclosed, corroded and replaced by tetrahedrite. g = germanite. t = tetrahedrite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 95 Renierite surrounded and replaced by germanite in tennantite. Galena, also present, replaces both the germanite and the tennantite. r = renierite. ge = germanite. t = tennantite. g = galena. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 96 Sphalerite partly surrounded and corroded by renierite which in turn is "invaded" and replaced by germanite. Both renierite and germanite transverse sphalerite in veinform type of replacement. Tennantite is also present, s = sphalerite, t = tennantite. r = renierite. ge = germanite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

308

Fig. 97 Sphalerite corroded and replaced by tennantite. Relics of sphalerite are present in the tennantite. s = sphalerite. t = tennantite, (tetrahedrite). r = relics of sphalerite in the tennantite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 98 Pyrite replaced by sphalerite, extending from the margins of the pyrite and following also its cracked pattern, ρ = pyrite. s = sphalerite. Vihanti, Finland. Oil-immersion with one nicol.

Fig. 99 Chalcopyrite replacing pyrite along a cleavage fracture system and also extending along cracks of the pyrite. c = chalcopyrite. ρ = pyrite. Sparneck, Fichtelgebirge, Germany. Oil-immersion with one nicol.

309

Fig. 100 Chalcopyrite replacing pyrite (in cases veinlets of chalcopyrite transverse the pyrite). ρ = pyrite. c = chalcopyrite. Tilva Mica, Bor, Serbia. Oil-immersion with one nicol.

Fig. 101 Pyrite replaced by veinlets of enargite. ρ = pyrite, arrows show corroded margins of the pyrite by enargite. e = enargite. S. Joäo Deserto, Aljustral, S. Portugal. Oil-immersion with one nicol.

Fig. 102 Pyrite corroded and replaced by millerite. ρ = pyrite. m = millerite, often following cracks of the pyrite. Temagami, Ontario, Canada. Oil-immersion with one nicol.

310

Fig. 103 Pyrite invaded and replaced by millerite. m = millerite. ρ = pyrite. Temagami, Ontario, Canada. Oil-immersion with one nicol.

Fig. 104 Pyrite invaded and replaced by millerite. m = millerite. ρ = pyrite. Temagami, Ontario, Canada. Oil-immersion with one nicol.

Fig. 105 Pyrite invaded and replaced by millerite. m = millerite. ρ = pyrite. Temagami, Ontario, Canada. Oil-immersion with one nicol.

311

Fig. 106 Pyrite idioblast replaced (atoll-type replacement) by neodigenite. ρ = pyrite. η = neodigenite. Bisbee, Arizona, USA. Oil-immersion with one nicol.

Fig. 107 Pyrite replaced by neodigenite. ρ = pyrite (idioblasts). η = neodigenite surrounding, invading and replacing the pyrite. Tilva Mica, Bor, Serbia. Oil-immersion with one nicol.

Fig. 108 Pyrite marginally replaced by chalcocite, resulting in indentation margins of the latter, c = chalcocite. ρ = pyrite. Debarau, Erythrea. Oil-immersion with one nicol.

312

Fig. 109 Chalcocite surrounding, corroding and replacing the pyrite. Also veinlets of chalcocite transverse and replace the pyrite. ρ = pyrite. c = chalcocite. Morenci, Arizona, USA. Oil-immersion with one nicol.

Fig. 110 Fractured pyrite invaded and replaced by chalcocite. ρ = pyrite. c = chalcocite. North Broken Hill, Australia. Oil-immersion with one nicol.

Fig. I l l Pyrite surrounded, corroded and replaced by chalcocite. The chalcocite replaces pyrite along cleavage and fracture spaces, ρ = pyrite. c = chalcocite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

313

Fig. 112 Pyrite surrounded, corroded and replaced by chalcocite. The chalcocite replaces pyrite along cleavage and fracture spaces, ρ = pyrite. c = chalcocite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 113 Pyrite surrounded, corroded and replaced by chalcocite. The chalcocite replaces pyrite along cleavage and fracture spaces, ρ = pyrite. c = chalcocite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 114 Pyrite replaced by prismatic molybdenite, ρ = pyrite. m = molybdenite, arrows show molybdenite extending into pyrite. Grant Mine, Columbia, USA. Oil-immersion with one nicol.

314

Fig. 115 Pyrite replaced by molybdenite with pyrite relics left in the molybdenite. ρ = pyrite. m = molybdenite. rp = relics of pyrite left in the molybdenite. Grant Mine, Columbia, USA. Oil-immersion with one nicol.

Fig. 116 Pyrite marginally replaced by realgar, ρ = pyrite. r = realgar. Skopje region. Oil-immersion with one nicol.

Fig. 117 Fractured pyrite partly replaced by realgar, ρ = pyrite. r = realgar. Skopje region. Oil-immersion with one nicol.

315

Fig. 118 Pyrite and enargite replaced by chalcocite (banded colloform). e = enargite. ρ = pyrite. c = chalcocite. Cusihuiriachic, Mexico. Oil-immersion with one nicol.

Fig. 119 Sphalerite marginally replaced by marcasite. s = sphalerite, m = marcasite. Joachimstal, Bohemia, Czech Republic. Oil-immersion with one nicol.

Fig. 120 Marcasite extensively replacing sphalerite. Relics of sphalerite preserved in marcasite. s = sphalerite, m = marcasite. Joachimstal, Bohemia, Czech Republic. Oil-immersion with one nicol.

316

Fig. 121 Radiating and star safflorite replacing sphalerite. Atoll replacement of sphalerite by safflorite. sf = safflorite. s-s = star safflorite. s = sphalerite. Joachimstal, Bohemia, Czech Republic. Oil-immersion with one nicol.

Fig. 122 Sphalerite replaced by bornite and chalcocite. Also idiomorphic (idioblastic) pyrite is shown. s = sphalerite, b = bornite. c = chalcocite. ρ = pyrite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 123 Sphalerite replaced by bornite and chalcocite. Also idiomorphic (idioblastic) pyrite is shown, s = sphalerite, b = bornite. c = chalcocite. ρ = pyrite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

317

Fig. 124 Idiomorphic (idioblastic) pyrite with marginal chalcocite extending into and replacing sphalerite, s = sphalerite, c = chalcocite. ρ = pyrite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 125 Pyrite marginally replaced by galena. Sphalerite also shown, ρ = pyrite. g = galena. s = sphalerite. Arrow shows corroded pyrite outline. Turkey, Asia Minor. Oil-immersion with one nicol.

Fig. 126 Pyrite enclosed and marginally replaced by galena. The relics of pyrite in the galena assume myrmekitic-like symplectic forms, ρ = pyrite. g = galena, s = sphalerite. Turkey, Asia Minor. Oil-immersion with one nicol.

A 318

Fig. 127 Pyrite interzonally replaced by galena, ρ = pyrite. g = galena. Joachimstal, Bohemia, Czech Republic. Oil-immersion with one nicol.

Fig. 128 Pyrite marginally and interzonally replaced by galena, ρ = pyrite. g = galena. Seriphos Island, Greece. Oil-immersion with one nicol.

Fig. 129 Pyrite replaced by galena along a fracture system, ρ = pyrite. g = galena. f = fracture system of pyrite replaced by galena. Supija Stujena, Montenegro. Oil-immersion with one nicol.

319

Fig. 130 Pyrrhotite replaced (along crystallographic directions) by galena, ρ = pyrrhotite. g = galena. Boliden, Sweden. Oil-immersion with one nicol.

Fig. 131 Arsenopyrite replaced by galena. The replacement is partly controlled by the crystalline outline of the arsenopyrite. a = arsenopyrite. g = galena. Mourros Gold Mine, Campo Jalte, N. Portugal. Oil-immersion with one nicol.

Fig. 132 Arsenopyrite corroded and replaced by galena, a = arsenopyrite. g = galena, arrow shows galena replacing arsenopyrite along? fractures. Mourros Gold Mine, Campo Jalte, N. Portugal. Oil-immersion with one nicol.

320

Fig. 133 Galena replacing chalcopyrite. Rounded relics of chalcopyrite are shown in the galena, c = chalcopyrite. g = galena. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

Fig. 134 Cubanite replaced along crystallographic directions by galena, c = cubanite. g = galena. Magnet Heights, Lydenburg District, Transvaal, South Africa. Oil-immersion with one nicol.

Fig. 135 Galena marginally replacing sphalerite. The galena replacement attains a myrmekitic-graphic-like intergrowth in the sphalerite, g = galena, s = sphalerite, gn = gangue. Weisser Hirsch, Goslar, Harz, Germany. Oil-immersion with one nicol.

321

Fig. 136 Sphalerite replaced by galena. Tennantite is also shown, s = sphalerite, g = galena, t = tennantite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 137 Sphalerite corroded and replaced by galena, s = sphalerite. g = galena, arrow shows extension of galena into and replacing the sphalerite. Weisser Hirsch Vein, Goslar, Harz, Germany. Oil-immersion with one nicol.

Fig. 138 Sphalerite replaced by galena, in cases veinform or attaining myrmekitic-like intergrowth patterns, s = sphalerite, g = galena, ρ = pyrite. Weisser Hirsch Vein, Goslar, Harz, Germany. Oil-immersion with one nicol.

Fig. 139 Sphalerite replaced by galena attaining myrmekitic-like intergrowth patterns. Weisser Hirsch Vein, Goslar, Harz, Germany. Oil-immersion with one nicol.

Fig. 140 Galena and silicate replacing sphalerite. Often the galena forms a fringe (margin) between the silicate and the sphalerite, s = sphalerite, g = galena. si = silicate (gangue), arrow shows galena margin between the silicate and the sphalerite. Kvcatha, Yugoslavia. Oil-immersion with one nicol.

Fig. 141 Sphalerite replaced by gangue (silicate) and galena which often forms a margin between the silicate and the sphalerite, s = sphalerite, g = galena, si = silicate, ρ = pyrite. Kvcatha, Yugoslavia. Oil-immersion with one nicol.

323

Fig. 142 Galena replacing tennantite and sphalerite, g = galena, t = tennantite. s = sphalerite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 143 Galena replacing tennantite. g = galena, t = tennantite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 144 Galena replacing neodigenite (atoll-type of replacement), g = galena, η = neodigenite. Bisbee, Arizona, USA. Oil-immersion with one nicol.

324

Fig. 145 Galena marginally replacing neodigenite and with extensions of galena replacing neodigenite in which relics of the latter are included in the galena, g = galena, η = neodigenite. Bisbee, Arizona, USA. Oil-immersion with one nicol.

Fig. 146 Jamesonite replacing pyrite. j = jamesonite. ρ = pyrite, arrows show jamesonite replacing (corroding and extending into) pyrite. s = sphalerite. Veta Clavo Bravo, Huari-Huari, Bolivia. Oil-immersion with one nicol.

Fig. 147 Jamesonite replacing pyrite. j = jamesonite. ρ = pyrite, arrows show jamesonite replacing (corroding and extending into) pyrite. Veta Clavo Bravo, Huari-Huari, Bolivia. Oil-immersion with one nicol.

325

Fig. 148 Atoll- and zonal-type replacement of sphalerite by jamesonite. s = sphalerite, j = jamesonite. Poopo, Bolivia. Oil-immersion with one nicol.

Fig. 149 Jamesonite replacing sphalerite. As a result of the replacement, jamesonite crystals are formed in the sphalerite, j = jamesonite. s = sphalerite. Poopo, Bolivia. Oil-immersion with one nicol.

Fig. 150 Colloform sphalerite replaced by gratonite. s = sphalerite, g = gratonite. Wiesloch, Baden, Germany. Oil-immersion with one nicol.

326

Fig. 151 Anglesite replacing skeletal pyrite (pyrite between well-developed crystals of gangue). g = gangue. ρ = pyrite. a = anglesite. Freiberg, Saxony, Germany. Oil-immersion with one nicol.

Fig. 152 Anglesite replacing rammelsbergite and its margin of safflorite. The anglesite transverses and replaces both the safflorite and the rammelsbergite extending from the outside inwards, a = anglesite. s = safflorite. r = rammelsbergite. Joachimstal, Bohemia, Czech Republic. Oil-immersion with one nicol.

Fig. 153 Oriented carbonates in the niccolite partly replaced by anglesite. c = carbonates (partly solved), η = niccolite. a = anglesite. Joachimstal, Bohemia, Czech Republic. Oil-immersion with one nicol.

327

Fig. 154 Oriented initial carbonates in niccolite to a great extent replaced by anglesite. Also relics of the initial carbonate are shown, η = niccolite. a = anglesite. c = relics of the carbonate initially oriented in the niccolite. Cobalt City, Ontario, Canada. Oil-immersion with one nicol.

Fig. 155 Ullmannite replaced by anglesite which from the margins extend inwards (following intracrystalline penetrability directions), u = ullmannite. a = anglesite. Montenarba, Argentina. Oil-immersion with one nicol.

Fig. 156 Ullmannite replaced by anglesite following the cleavage pattern of the host, a = anglesite. u = ullmannite. Montenarba, Argentina. Oil-immersion with one nicol.

mm$m 328

Fig. 157 Sphalerite replaced by argentinite; pyrite is also shown, s = sphalerite, a = argentite. ρ = pyrite. Joachimstal, Bohemia, Czech Republic. Oil-immersion with one nicol.

Fig. 158 Pyrargyrite replacing sphalerite, s = sphalerite, py = pyrargyrite. g = galena, ρ = pyrite. Zacatecas, Mexico. Oil-immersion with one nicol.

Fig. 159 Pyrargyrite replacing sphalerite and also pyrite. Atoll-type replacement of pyrite by pyrargyrite. s = sphalerite, py = pyrargyrite. ρ = pyrite. Zacatecas, Mexico. Oil-immersion with one nicol.

329

Fig. 160 Cobaltite with a safflorite margin is transversed and replaced by argentite. co = cobaltite. s = safflorite. a = argentite. Joachimstal, Bohemia, Czech Republic. Oil-immersion with one nicol.

Fig. 161 Cobaltite crystals initially associated with (included in) sphalerite. Due to the replacement of sphalerite by argentite, some are now in the argentite. s = sphalerite, co = cobaltite. a = argentite. Marcasite (m) is present and also safflorite - and partly starsafflorite (sf) can be seen. Joachimstal, Bohemia, Czech Republic. Oil-immersion with one nicol.

Fig. 162 Cobaltite is corroded and replaced by argentite. co = cobaltite. r = relics of cobaltite in argentite. a = argentite. Joachimstal, Bohemia, Czech Republic. Oil-immersion with one nicol.

330

Fig. 163 Zoned cobaltite "invaded" and replaced by proustite. co = cobaltite. pr = proustite. Cobalt City, Ontario, Canada. Oil-immersion with one nicol.

Fig. 164 Wire-silver replacing argentite. s = silver, a = argentite. Guanajuato, Mexico. Oil-immersion with one nicol.

Fig. 165 Argentite marginally replaced by silver, a = argentite. s = silver. Freiberg, Saxony, Germany. Oil-immersion with one nicol.

331

Fig. 166 Proustite invaded and replaced by pyrargyrite. Pyrite is also present. pr = proustite. py = pyrargyrite. ρ = pyrite. Cobalt City, Ontario, Canada. Oil-immersion with one nicol.

Fig. 167 Gold tellurite and silver, both replacing argentite. gt = gold tellurite, s = silver, a = argentite. Guanajuato, Mexico. Oil-immersion with one nicol.

Fig. 168 Argentite associated with gold tellurite. The gold tellurite is replaced by gold and the argentite also by a network system of gold veinlets. gt = gold tellurite, g = gold, a = argentite. Guanajuato, Mexico. Oil-immersion with one nicol.

332

Fig. 169 Argentite replaced by gold and silver, a = argentite. g = gold, s = silver. Guanajuato, Mexico. Oil-immersion with one nicol.

Fig. 170 Argentite marginally replaced by gold, a = argentite. g = gold. Nancy Mine, Boulder County, Colorado, USA. Oil-immersion with one nicol.

Fig. 171 Argentite marginally replaced by petzite. Also petzite replacing argentite extends from the margin into the argentite (see arrow), ρ = petzite. a = argentite. Nancy Mine, Boulder County, Colorado, USA. Oil-immersion with one nicol.

333

Fig. 172 Sylvanite replacing petzite. s = sylvanite. ρ = petzite. Nancy Mine, Boulder County, Colorado, USA. Oil-immersion with one nicol.

Fig. 173 Cobaltite partly replaced by niccolite. co = cobaltite. η = niccolite. r = relic of cobaltite in niccolite. Kiruna, N. Sweden. Oil-immersion with one nicol.

Fig. 174 Cobaltite corroded, surrounded and replaced by niccolite. co = cobaltite. η = niccolite. Kiruna, N. Sweden. Oil-immersion with one nicol.

334

Fig. 175 Zoned cobaltite crystals are partly replaced by niccolite which also extends between the cobaltite crystals, η = niccolite. co = cobaltite. Kiruna, N. Sweden. Oil-immersion with one nicol.

0.1mm

Fig. 176 Cassiterite zonally replaced by stannite. c = cassiterite. s = stannite. Shizhu Yuan near Chenzhou, Hunan Province, China. Oil-immersion with one nicol.

Fig. 177 Cassiterite "invaded" and replaced by stannite and haematite, c = cassiterite. s = stannite. h = haematite. Altenberg, Saxony, Germany. Oil-immersion with one nicol.

335

Fig. 178 Stannite and haematite replacing cassiterite. The stannite attains veinform replacement patterns in the cassiterite. s = stannite. h = haematite, c = cassiterite. Altenberg, Saxony, Germany. Oil-immersion with one nicol.

m

jjL -

• r —

•

/ w W

•-

^ • jr^MMHfl

_

Λ

ΪΜ MM 4

ι

m

Fig. 179 Elongated "prismatic" stannite crystals "invading" and replacing cassiterite. s = stannite. c = cassiterite. Carrock Mine, Grainsgill, Caldbeck Fells, England. Oil-immersion with one nicol.

j,

-

lifo!·

Fig. 180 Stannite replaced by cassiterite extending from the margins inwards, s = stannite. c = cassiterite. Carrock Mine, Grainsgill, Caldbeck Fells, England. Oil-immersion with one nicol.

336

Fig. 181 Cassiterite replacing stannite. In cases the replacement of stannite by cassiterite "attains" oriented lamellae-pattern (see arrow), c = cassiterite. s = stannite. Carrock Mine, Grainsgill, Caldbeck Fells, England. Oil-immersion with one nicol.

Fig. 182 Wolframite replaced and "invaded" by cassiterite. Prismatic cassiterite develops in the wolframite, c = cassiterite. w = wolframite. Colquiri, Bolivia. Oil-immersion with one nicol.

Fig. 183 Pyrite replaced by lamellae of teallite. s = stannite. ρ = pyrite. t = teallite. Monserrat Mine, Bolivia. Oil-immersion with one nicol.

337

Fig. 184 Crystalloblastic pyrite replacing stannite. Lamellar teallite develops "bending" over pyrite and stannite. ρ = pyrite. s = stannite. t = teallite. Monserrat Mine, Bolivia. Oil-immersion with one nicol.

Fig. 185 Stannite replaced by lamellae of teallite. s = stannite. t = teallite. Monserrat Mine, Bolivia. Oil-immersion with one nicol.

Fig. 186 Pyrite replacing stannite. ρ = pyrite. s = stannite. Monserrat Mine, Bolivia. Oil-immersion with one nicol.

0.1mm

338

Fig. 187 Pyrite replacing cassiterite. Arrow shows pyrite replacing the cassiterite along crystal-cleavage directions. Stannite is also present, c = cassiterite. ρ = pyrite. s = stannite. Monserrat Mine, Bolivia. Oil-immersion with one nicol.

Fig. 188 Idiomorphic cassiterite replaced along cracks by the surrounding sphalerite, c = cassiterite. s = sphalerite. Chang Poy, Dachang, S. E. China. Oil-immersion with one nicol.

Fig. 189 Cassiterite corroded and replaced by sphalerite, c = cassiterite. s = sphalerite, ρ = ?pyrite. Chang Poy, Dachang, S. E. China. Oil-immersion with one nicol.

339

Fig. 190 Pyrite replacing cassiterite and sphalerite, ρ = pyrite. c = cassiterite. s = sphalerite. Chang Poy, Dachang, S. E. China. Oil-immersion with one nicol.

Fig. 191 "Atoll"-type of replacement. Bi replacing haematite. b = bismuth. h = haematite. S. Joäo Deserto, Aljustral, S. Portugal. Oil-immersion with one nicol.

Fig. 192 Bismuth including relics of haematite and replacing the haematite lamellae along penetrability directions interlamellar spaces (see arrow), b = bismuth, h = haematite. Tarvis Township, Algoma District, Ontario, Canada. Oil-immersion with one nicol.

340

Fig. 193 Tetrahedrite replaced by Bi. Tetrahedrite is also replaced by wittichenite. b = Bi. t = tetrahedrite. w = wittichenite. Daniel Mine near Wittichen, Germany. Oil-immersion with one nicol.

Fig. 194 Bi replacing wittichenite and emplectite. Tetrahedrite is also present, b = Bi. w = wittichenite. t = tetrahedrite. e = emplectite?. Daniel Mine near Wittichen, Germany. Oil-immersion with one nicol.

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Fig. 195 Bi replaced and surrounded by chloanthite. Also chloanthite extends from the margins of the Bi inwards. c = chloanthite margins surrounding Bi and extending in the central part (see arrow), b = Bi. Luisito Mine near Pozoblanco, Cordoba, Spain. Oil-immersion with one nicol.

341

Fig. 196 Marginal chloanthite and ramified extension in the Bi. b = Bi. cm = marginal chloanthite. er = ramified (extending and branching of) chloanthite in the Bi. Loma Desola, Pizarra (Espuela Mine), Spain. Oil-immersion with one nicol.

Fig. 197 Native Bi replaced by ring structures of chloanthite. b = Bi. rc = ring structure of chloanthite. Loma Desola, Pizarra (Espuela Mine), Spain. Oil-immersion with one nicol.

Fig. 198 Chloanthite marginal to native Bi with Bi bursting through the chloanthitic margin and by mobilization extending into the gangue. b = native Bi. c = chloanthitic margin. be = bismuth extension (mobilization) into the gangue. Loma Desola, Pizarra (Espuela Mine), Spain. Oil-immersion with one nicol.

342

Fig. 199 "Atoll"-type of replacement. Bi replacing niccolite and with veinform extension into it. b = Bi. η = niccolite. Loma Desola, Pizarra (Espuela Mine), Spain. Oil-immersion with one nicol.

Fig. 200 Bi replacing niccolite and predominantly a banded type of replacement of niccolite by Bi is shown, b = Bi. η = niccolite. Loma Desola, Pizarra (Espuela Mine), Spain. Oil-immersion with one nicol.

Fig. 201 Niccolite with a chloanthite margin. A Bi overgrowth attaining crystalline outlines has inward veinform extensions which transverse the chloanthite margin and ramify in the niccolite, eventually attaining an atoll-type replacement of niccolite by Bi. bo = Bi overgrowth. be = Bi transversing the chloanthite margin. bn = Bi in the niccolite. η = niccolite. c = chloanthite. Loma Desola, Pizarra (Espuela Mine), Spain. Oil-immersion with one nicol.

343

Fig. 202 Niccolite with a chloanthite margin has an overgrowth of Bi, an extension (vein) of which transverses the chloanthite and ramifies in the niccolite resulting in a banded replacement of niccolite by Bi. bo = Bi overgrowth, be = Bi extension (veinlet) transversing both the chloanthite margin and the niccolite. bb = banded replacements of the niccolite by Bi. η = niccolite. c = chloanthite. Loma Desola, Pizarra (Espuela Mine), Spain. Oil-immersion with one nicol.

Fig. 203 Bi replacing (cutting across) zoned cobaltite and safflorite. b = Bi. c = cobaltite. s = safflorite. Cobalt City, Ontario, Canada. Oil-immersion with one nicol.

Fig. 204 Relics of cobaltite in the Bi. b = Bi. c = cobaltite. r = relics of the cobaltite in the Bi. Cobalt City, Ontario, Canada. Oil-immersion with one nicol.

344

Fig. 205 Bi with a margin of chloanthite is replaced by lamellar sphalerite. Also sphalerite is present outside the chloanthite. b = Bi. c = chloanthite. si = sphalerite lamellae replacing Bi. s = sphalerite. Luisito Mine near Pozoblanco, Cordoba, Spain. Oil-immersion with one nicol.

Fig. 206 Bi corroded and replaced by sphalerite; also with lamellar extensions of sphalerite in the bismuth, b = Bi. s = sphalerite. si = lamellar sphalerite replacements in Bi. Luisito Mine near Pozoblanco, Cordoba, Spain. Oil-immersion with one nicol.

r

·.- · < · • · -f

Fig. 207 Bi corroded and replaced by sphalerite; also with lamellar extensions of sphalerite in the bismuth, b = Bi. s = sphalerite. si = lamellar sphalerite replacements inBi. Luisito Mine near Pozoblanco, Cordoba, Spain. Oil-immersion with one nicol.

345

Fig. 208 Mutual replacement and remobilization pattern of sphalerite replacing Bi and Bi sending extensions into the sphalerite. Bi is partly marginally surrounded by chloanthite. b = Bi. s = sphalerite. be = Bi extension in sphalerite, se = sphalerite replacing Bi and extending into it. c = chloanthite. Luisito Mine near Pozoblanco, Cordoba, Spain. Oil-immersion with one nicol.

Fig. 209 A pattern of branching veinlets of millerite replacing bismuthinite. b = bismuthinite. m = millerite. Wissen near Sieg, Westphalia, Germany. Oil-immersion with one nicol.

Fig. 210 Millerite marginal and as veinlet extensions replacing bismuthinite. m = millerite. b = bismuthinite. Wissen near Sieg, Westphalia, Germany. Oil-immersion with one nicol.

346

Fig. 211 Hauchecornite replaced by millerite. Relics of hauchecornite are left in the millerite. h = hauchecornite. m = millerite. r = relics of hauchecornite left in the millerite. Wissen near Sieg, Westphalia, Germany. Oil-immersion with one nicol.

Fig. 212 Trogtalite invaded and replaced by clausthalite. t = trogtalite. c = clausthalite. Trogtal near Lautental, Harz, Germany. Oil-immersion with one nicol.

Fig. 213 Trogtalite with a margin of hastite invaded and replaced by clausthalite. t = trogtalite. h = hastite. c = clausthalite. cl = chloanthite. Trogtal near Lautental, Harz, Germany. Oil-immersion with one nicol.

347

Fig. 214 Clausthalite with a margin of hastite. Such patterns could suggest mutual replacements, c = clausthalite. h = hastite. Trogtal near Lautental, Harz, Germany. Oil-immersion with one nicol.

005 mm

Fig. 215 Complex patterns of mutual replacement of hastite/clausthalite. Often the hastite forms margins. Also in cases an atoll-type replacement of hastite by clausthalite is possible, h = hastite. c = clausthalite. Trogtal near Lautental, Harz, Germany. Oil-immersion with one nicol.

Fig. 216 Pitchblende interzonally invades and replaces smaltite. ρ = pitchblende, s = smaltite. Eldorado Mine, Great Bear Lake, Canada. Oil-immersion with one nicol. χ 1000.

348

Fig. 217 Smaltite crystals interzonally replaced by pitchblende, s = smaltite. ρ = pitchblende. Eldorado Mine, Great Bear Lake, Canada. Oil-immersion with one nicol.

χ 1000.

Fig. 218 Native silver with marginal U0 2 , apparently later than the Ag replacing the silver, s = silver, ρ = pitchblende. Eldorado Mine, Great Bear Lake, Canada. Oil-immersion with one nicol.

Fig. 219 Tetrahedrite replacing pitchblende along cracks (initial gelstructures crushed), t = tetrahedrite. ρ = pitchblende, s = smaltite. Wittichen, Black Forest, Germany. Oil-immersion with one nicol.

349

Fig. 220 Marcasite replacing pitchblende along shrinkage cracks, ρ = pitchblende, m = marcasite. Wölsendorf, Bavaria, Germany. Oil-immersion with one nicol.

Fig. 221 Original gel pitchblende, deformed and invaded along cracks by later Bi. b = Bi. ρ = pitchblende. Sophia Mine, Halde, Wittichen, Baden, Germany. Oil-immersion with one nicol.

Fig. 222 Deformed initial gel pitchblende, invaded and replaced by Bi with a relic of pitchblende left in the central part of the Bi. b = Bi. ρ = deformed pitchblende, r = relic of pitchblende in the Bi. Sophia Mine, Halde, Wittichen, Baden, Germany. Oil-immersion with one nicol.

350

Fig. 223 Original gel pitchblende, deformed and invaded along cracks by later Bi. Gel spheroids of the pitchblende are enclosed in the Bi and a case is exhibited of an intact gel spheroid enclosed in the later Bi. b = Bi. ρ = pitchblende. s = intact spheroid surrounded by the later Bi. Sophia Mine, Halde, Wittichen, Baden, Germany. Oil-immersion with one nicol.

Fig. 224 Cell wood structure replaced by uraninite and pyrite. u = uraninite. ρ = pyrite. Happy Jack, White Canyon, Utah, USA. Oil-immersion with one nicol.

Fig. 225 Native B i replaced marginally by brown iron, b = Bi. bi = brown iron. Espuela San Miguel, Villanueva de Cordoba, Spain. Oil-immersion with one nicol.

351

Fig. 226 Initial pyrrhotite dissolved and replaced by lepidocrocite. A marginal gel lepidocrocite partly includes initial limonitic-oolitic relics. I = lepidocrocite replacing initial pyrrhotite. 1-g = gel lepidocrocite enclosing relics of initial limonitic oolites. Larymna, Lokris, Greece. Oil-immersion with one nicol.

Fig. 227 Haematite containing martite-magnetite. Another mineral grain outline is also shown, almost completely replaced by the haematite and probably representing former davidite as a relic of this replacement, m = magnetite, h = haematite. hd = initial davidite replaced by haematite. Wipper Aminga, Southern Australia. Oil-immersion with one nicol.

ι",· < « * ψ A. j ;'s.!.."ill f l - / J / . " \ ** -J» . » w »

V

V

Fig. 228 Initial magnetite with oriented lamellae of ilmenite selectively replaced by pyrite. As a result, the ilmenite lamellae are left as relics in the pyrite following their initial orientation pattern in the titanomagnetite. ρ = pyrite. i = ilmenite. Frood Mine, Sudbury, Canada. Oil-immersion with one nicol. (Figs. 228 and 229 show iron-oxides substitution)

Fig. 229 Titanomagnetite (magnetite with ex-solutions of ilmenite). The magnetite is solved out and the oriented ilmenite lamellae are left as relics. m = initial magnetite solved out. il = ilmenite lamellae left as relics. Jay, Cap Colony. Oil-immersion with one nicol.

Fig. 230 A non-deformation pattern of initial lamellar haematite replaced by induced reduction to magnetite, m = magnetite. Mavro Punti, Seriphos, Greece. Oil-immersion with one nicol.

Fig. 231 Initial lievrite changed by "induced reduction" to magnetite. Mavro Punti, Seriphos, Greece. Oil-immersion with one nicol.

353

Fig. 232 Initial colloform radiating hydro-oxides prismatic crystals changed finally to magnetite, m = magnetite, ρ = pyrite. Peko Mine, Northern Territory, Australia. Oil-immersion with one nicol.

Fig. 233 Initial carbonate exhibiting gel patterns is replaced by magnetite in which the initial gel patterns are preserved, m = magnetite, ρ = pyrite. g = initial gel structure. Peko Mine, Northern Territory, Australia. Oil-immersion with one nicol.

Fig. 234 Initial zoned carbonate replaced by magnetite (the zonal structure of the carbonate is preserved). ζ = initial zonal structure of the carbonate now replaced by magnetite, m = magnetite. Peko Mine, Northern Territory, Australia. Oil-immersion with one nicol.

354

Fig. 235 Galena replaced by a later ? crystalloblastic growth of magnetite partly attaining idiomorphism. g = galena, m = magnetite. Trepca, Serbia. Oil-immersion with one nicol.

Fig. 236 Magnetite successively replaced by sphalerite and chalcopyrite. m = magnetite, s = sphalerite, c = chalcopyrite. Naica, near Parrel, Mexico. Oil-immersion with one nicol.

Fig. 237 Multiple successive replacement of pyrite by galena and chalcopyrite. ρ = pyrite. g = galena, c = chalcopyrite. Rammelsberg, Harz, Germany. Oil-immersion with one nicol.

355

Fig. 238 Multiple replacement of pyrite by galena and sphalerite. The pyrite is also marginally replaced by galena, ρ = pyrite. g = galena, s = sphalerite. Websky, Turkey, Asia Minor. Oil-immersion with one nicol.

Fig. 239 Multiple replacement of arsenopyrite by galena and gold, a = arsenopyrite. g = galena, au = gold. Mourros Gold Mine, Campo Jalte, N. Portugal. Oil-immersion with one nicol.

Fig. 240 Successive replacement of quartz by bornite and tetrahedrite. q = quartz, b = bornite. t = tetrahedrite. San Marcos Mine, Conception del Oro. Oil-immersion with one nicol.

356

Fig. 241 Chalcocite successively replaced by malachite and native copper. ch = chalcocite. m = malachite, c = native copper. Bisbee, Arizona, USA. Oil-immersion with one nicol.

•ν

%

\ '

-

Fig. 242 Multiple successive replacement of pyrite by bornite, chalcopyre and chalcocite (replacing both the bornite and the chalcopyrite). ρ = pyrite. b = bornite (also included in the chalcopyrite). ch = chalcopyrite. c = chalcocite. Matchless Mine, Friedenau, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 243 Successive replacement of pyrite by chalcopyrite which is transversed by chalcocite. ρ = pyrite. ch = chalcopyrite. c = chalcocite. Matchless Mine, Friedenau, Namibia, S. W. Africa. Oil-immersion with one nicol.

357

Fig. 244 Bornite in replacement intergrowth with later chalcopyrite. Both the bornite and the chalcopyrite are replaced by veinlets of chalcocite. Pyrite is also present, b = bornite. ch = chalcopyrite. c = chalcocite. ρ = pyrite. Matchless Mine, Friedenau, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 245 Tetrahedrite replaced by pyrite and both are replaced and transected by later chalcocite. t = tetrahedrite. ρ = pyrite. c = chalcocite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 246 Enargite multiple replacement by chalcopyrite, covellite and haematite, e = enargite. ch = chalcopyrite (as well as fine chalcopyrite infiltrations), co = covellite. h = haematite. Feldbach Mine, Binnental, Switzerland. Oil-immersion with one nicol.

358

Fig. 247 Multiple replacements. Bornite replaced by chalcopyrite and pyrite by atoll-type chalcopyrite replacement. Both bornite and chalcopyrite are replaced by veinlets of chalcocite. b = bornite. ch = chalcopyrite. ρ = pyrite. c = chalcocite. Matchless Mine, Friedenau, Namibia, S . W . Africa. Oil-immersion with one nicol.

Fig. 248 Magnetite associated with silicate surrounded and replaced by chalcopyrite which in turn is rounded and replaced by blastic pyrite. s = silicate, m = magnetite, c = chalcopyrite. ρ = blastic pyrite. Dognacska, Banat, Croatia. Oil-immersion with one nicol.

0.1mm

Fig. 249 Bi replaced by sphalerite and chloanthite. b = Bi. s = sphalerite, c = chloanthite. Luisito Mine near Pozoblanco, Cordoba, Spain. Oil-immersion with one nicol.

359

Fig. 250 Sphalerite replaced by bismuthinite which in turn is replaced by Bi. s = sphalerite, b = Bi. bi = bismuthinite. Shizhu Yuan near Chenzhou, Hunan Province, China. Oil-immersion with one nicol.

Fig. 251 Sphalerite replaced by crystalloblastic pyrite and transected by a network of silicates and galena veinlets. Often the galena is marginal to the silicate veinlets. s = sphalerite, ρ = crystalloblastic pyrite. g = galena, black veinlets transecting the sphalerite consist of silicate. Kvcatha, Yugoslavia. Oil-immersion with one nicol.

Fig. 252 Older galena (cross shaped structures due to corrosion) enclosed in pitchblende. Also a later (younger generation) of galena forms a margin on the pitchblende. g = interspersed galena in pitchblende (for details see Fig. 253). ρ = pitchblende, gl = later galena margin. Weberbruck, Johannesschacht, Wölsendorf, Bavaria, Germany. Oil-immersion with one nicol.

360

Ψ*

Fig. 253 Star-shaped galena ?relics in pitchblende representing the first generation of galena (see Fig. 252). g = first generation of galena (probably representing relics of replacement), ρ = pitchblende. Weberbruck, Johannesschacht, Wölsendorf, Bavaria, Germany. Oil-immersion with one nicol.

Fig. 254 Successive formation of smaltite core with pitchblende interzonal (followed again by smaltite crystallization). A successive formation of smaltite pitchblende is suggested on the basis of the texture (see also Fig. 255). ρ = pitchblende, s = smaltite. Wittichen, Black Forest, Germany. Oil-immersion with one nicol.

Fig. 255 Successive formation of smaltite core with pitchblende interzonal (followed again by smaltite crystallization as shown in Fig. 254). A successive formation of smaltite pitchblende is suggested on the basis of the texture (see also Fig. 254). ρ = pitchblende, s = smaltite. Wittichen, Black Forest, Germany. Oil-immersion with one nicol.

361

Fig. 256 Gold enclosed (rounded) by pitchblende which is in turn surrounded and transversed by gangue and also gold is in the gangue. ρ = pitchblende, g = gold. Shinkolobwe, Katanga, Zaire. Oil-immersion with one nicol.

Fig. 257 Gold overgrowth on pitchblende and fine particles of it together with silicate extending into cracks of the uraninite. g = gold u = uranium Shinkolobwe, Katanga, Zaire. Oil-immersion with one nicol.

Fig. 258 The pattern could perhaps be interpreted as due to successive formations: pitchblende partly surrounded by an intergrowth of bornite with bismuthinite (or emplectite). In turn the outer gel layer of pitchblende surrounds and encloses part of the bornite-bismuthinite (or emplectite) symplectite. ρ = spheroidal pitchblende, corroded by a symplectite consisting of bornite (b) and bismuthinite (bi). Arrow shows outer gel layer surrounding and enclosing part of the symplectite. Schniedeberg, Kowary, Poland. Oil-immersion with one nicol.

362

Fig. 259 Gold and chalcopyrite "veinlets" invading and replacing pyrite. g = gold veinlet (including chalcopyrite). c = chalcopyrite "veinlet". ρ = pyrite. Quebec, Canada. Oil-immersion with one nicol.

Fig. 260 Gold and chalcopyrite replacing pyrite. g = gold. c = chalcopyrite. ρ = pyrite. Quebec, Canada. Oil-immersion with one nicol.

Fig. 261 Gold and chalcopyrite "veinlets" replacing pyrite. Also granular gold due to replacement is shown. g = gold veinlet. gg = granular gold, c = chalcopyrite. ρ = pyrite. Quebec, Canada. Oil-immersion with one nicol.

363

Fig. 262 Gold replacing chalcopyrite (and pyrite) which in turn is replacing pyrite. ρ = pyrite. c = chalcopyrite. g = gold. Quebec, Canada. Oil-immersion with one nicol.

Fig. 263 Successive crystallizations (formations). Smaltite, pitchblende and tetrahedrite invading and replacing the deformed (initial gelspheroids) of the pitchblende, s = smaltite. ρ = pitchblende, py = pyrite. t = tetrahedrite. Wittichen, Black Forest, Germany. Oil-immersion with one nicol.

Fig. 264 Chromite grains included and corroded by garnet crystalloblast which is surrounded and in cases invaded by chalcopyrite. c = chromite. g = garnet, ch = chalcopyrite. Qutukumpu, Finland. Oil-immersion with one nicol.

364

Fig. 265 Chromite grains included and corroded by garnet crystalloblast which is surrounded and in cases invaded by chalcopyrite. c = chromite. g = garnet. ch = chalcopyrite, arrow shows chalcopyrite corroding and marginal to an included chromite grain in the garnet crystalloblast. Outukumpu, Finland. Oil-immersion with one nicol.

Fig. 2 6 6 Garnets corroded and invaded by chalcopyrite which surrounds the garnet and sends extensions into the garnet, g = garnet, ch = chalcopyrite. Outukumpu, Finland. Oil-immersion with one nicol.

0 05 mm

Fig. 267 Complex crystallization patterns ? successive crystallizations. Braunite surrounded and enclosed by jacobsite which in turn is surrounded by marginal braunite. b, = braunite enclosed in jacobsite. j = jacobsite. b 2 = braunite margin of jacobsite. Possible crystallization sequence ? b, jacobsite b 2 . Otjosondu, Sandfeld, Namibia, S. W. Africa. Oil-immersion with one nicol.

365

Fig. 268 ?Successive crystallizations. Braunite (b,) surrounded by jacobsite (jj) and in turn surrounded by braunite (b2). Jacobsite (j 2 ) partly surrounds braunite (b2). Otjosondu, Sandfeld, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 269 Braunite in places invaded by jacobsite; the braunite is also shown to be transected by a veinlet of jacobsite. b = braunite. j = jacobsite. vj = veinlet of jacobsite. Otjosondu, Sandfeld, Namibia, S. W. Africa. Oil-immersion with one nicol.

A

Fig. 270 Ilmenite veinlet transversing magnetite with extensions of ilmenite attaining lamellar oriented bodies within the magnetite. m = magnetite. vl = veinlet of ilmenite. e = extension of ilmenite attaining oriented lamellar form. dm = martitization. Sphinx near Bon Accord, near Pretoria, Transvaal, South Africa. Oil-immersion with one nicol.

>

m \

366

Fig. 271 Oriented lamellae of ilmenite in magnetite extending beyond their intergrowth with the magnetite and occurring together with silicate or pyrite. m = magnetite. il-e = ilmenite extending beyond its intergrowth with the magnetite, il = ilmenite. Otanmäki, Finland. Oil-immersion with one nicol.

Fig. 272 Oriented lamellae of ilmenite in magnetite extending beyond their intergrowth with the magnetite and occurring together with silicate or pyrite. m = magnetite. il-e = ilmenite extending beyond its intergrowth with the magnetite, il = ilmenite. ρ = pyrite. Otanmäki, Finland. Oil-immersion with one nicol.

Fig. 273 Oriented lamellar ilmenite as extension of ilmenite replacing magnetite, m = magnetite. il = ilmenite (replacing magnetite), arrows show extension of ilmenite attaining lamellar forms (?pseudo exsolutions). Otanmäki, Finland. Oil-immersion with one nicol.

367

Fig. 274 Ilmenite lamellae oriented parallel to the octahedral magnetite face. However, as arrows "a" show, the ilmenite is not delimited by a straight boundary. It appears that ilmenite bodies diffuse into the magnetite, most probably suggesting an ilmenite mobilization along the weakness planes [111] of the host magnetite, m = magnetite. il = ilmenite (following the [111] of the magnetite). Otanmäki, Finland. Oil-immersion with one nicol.

Fig. 275 Magnetite with lamellar body of ilmenite. Spinel bodies occur either parallel oriented (parallel to the [100] of the magnetite) or occur as intergranular-interleptonic space fillings between the ilmenite and the magnetite, m = magnetite. il = ilmenite lamellae oriented parallel to the [111] of the magnetite (arrow "a"). o-s = oriented spinel bodies most probably parallel to the [100] of the magnetite (arrow "b"). s = spinel occupying the interleptonic space between the ilmenite lamella and the magnetite (arrow "c"). Lac de la Blanche Saguenay, Quebec, Canada. Oil-immersion with one nicol.

Fig. 276 Ilmenite in contact with magnetite. Ilmenite lamellae are parallel oriented to the [111] of the host magnetite. Spinel occurs as lamellar bodies following the [111] of the host magnetite and also as intergranular bodies between the ilmenite and the magnetite, m = magnetite, il = ilmenite. 1-il = lamellar ilmenite following the [111] of the magnetite, i-s = intergranular spinel between magnetite and ilmenite (arrow "a" shows intergranular spinel with extensions of it as lamellar spinel into the adjacent magnetite). Otanmäki, Finland. Oil-immersion with one nicol.

368

m

ul

(

#

^

Fig. 277 Magnetite with spinel and tablecloth structure, m = magnetite. ul = ulvite (tablecloth structure), s = spinel. Lac de la Blanche Saguenay, Quebec, Canada. Oil-immersion with one nicol.

wsmm

Fig. 278 Spinel oriented parallel to the [100] cleavage of the magnetite. Ulvospinel is also shown, m = magnetite, s = spinel, u = ulvospinel. Lac de la Blanche Saguenay, Quebec, Canada. Oil-immersion with one nicol.

Fig. 279 Magnetite with lines (white in color) and with lens-shaped ilmenite bodies following the white lines. m = magnetite, il = ilmenite. s = spinel. w = white lines in the magnetite. Bushveld Norite, Phoenix, Pretoria, Transvaal, South Africa. Oil-immersion with one nicol.

369

Fig. 280 Ilmenite lamellae following the [111] of the magnetite (marginal to the ilmenite dots of spinel), m = magnetite. il = ilmenite, dots marginal to ilmenite spinel, s = spinel. Concordia Mine near O'okiep, South Africa. Oil-immersion with one nicol.

Fig. 281 Spinel replacing magnetite along the [100] and ilmenite along the [111] (face) direction, m = magnetite, il = ilmenite. s = spinel, c = chalcopyrite. Concordia Mine near O'okiep, South Africa. Oil-immersion with one nicol.

i i *

>•

f

Τ" *

&

ν

\

X i

*. ^

1

>4 % * J

_

t

k ·;·i'il Fig. 282 Symplectic spinel replacing magnetite along the [100] face. Ilmenite also follows the [111] face of the magnetite, m = magnetite, il = ilmenite. s = spinel, c = chalcopyrite. Concordia Mine near O'okiep, South Africa. Oil-immersion with one nicol.

370

Fig. 283 Symplectic (skeletal in form) myrmekitic-like spinel bodies following the [100] face of the magnetite. Ilmenite also follows the [111] of the magnetite, m = magnetite, s = symplectic spinel bodies, il = ilmenite. Concordia Mine near O'okiep, South Africa. Oil-immersion with one nicol.

Fig. 284 Symplectic (skeletal in form) myrmekitic-like spinel bodies following the [100] face of the magnetite. Ilmenite also follows the [ 111] of the magnetite, m = magnetite, s = symplectic spinel bodies, il = ilmenite. Concordia Mine near O'okiep, South Africa. Oil-immersion with one nicol.

Fig. 285 Bands of spinel replacing both ilmenite lamella and the magnetite, m = magnetite, il = ilmenite. s = spinel (bands following both the ilmenite lamella and the magnetite), arrow shows spinel following the [100] of the magnetite. Concordia Mine near O'okiep, South Africa. Oil-immersion with one nicol.

371

Fig. 286 Spinel replacing magnetite. Chalcopyrite partly replacing the spinel bodies (lamellae), m = magnetite, s = spinel. c = chalcopyrite replacing the spinel, il = ilmenite lamellae, ch = chalcopyrite. Concordia Mine near O'okiep, South Africa. Oil-immersion with one nicol.

\s \sv

Fig. 287 Spinel veinlets with extensions following the [100] face of the magnetite. sv = spinel veinlets. s = spinel following the [100] of the magnetite. m = magnetite. The same pattern occurs in many localities such as vanadiferous magnetite and Lac de la Blanche Saguenay Compte, Quebec Province, Canada. Oil-immersion with one nicol.

m

Fig. 288 Veinform spinels with oriented extension and sheathing ilmenite lamella in magnetite, m = magnetite, s = spinel. e = spinel extension, il = ilmenite. The same pattern occurs in many localities such as vanadiferous magnetite and Lac de la Blanche Saguenay Compte, Quebec Province, Canada. Oil-immersion with one nicol.

372

, './ 11\

I.

ν

«Ι 1

s '

m

\ si

Fig. 289 Spinel following interleptonic spaces [100] and very fine spinel bodies also present in magnetite. s = spinel, arrows show fine spinels, m = magnetite. The same pattern occurs in many localities such as vanadiferous magnetite and Lac de la Blanche Saguenay Compte, Quebec Province, Canada. Oil-immersion with one nicol.

f .

Fig. 290 Magnetite with a fracture system, some occupied by spinel and ilmenite. m = magnetite, s = spinel. il = ilmenite, arrow "a" shows fractures not occupied by spinel or ilmenite, arrow "b" shows fine fractures occupied by ilmenite. Mooihoek, Lydenburg District, Transvaal, South Africa. Oil-immersion with one nicol.

Fig. 291 Magnetite with a fracture system, some occupied by spinel and ilmenite. m = magnetite. s = spinel. il = ilmenite. Mooihoek, Lydenburg District, Transvaal, South Africa. Oil-immersion with one nicol.

m

373

Fig. 292 Magnetite with a fracture system, some occupied by spinel and ilmenite. m = magnetite, s = spinel. il = ilmenite, arrow "a" shows fractures not occupied by spinel or ilmenite. Mooihoek, Lydenburg District, Transvaal, South Africa. Oil-immersion with one nicol.

Fig. 293 Oriented ilmenite exsolutions (lamellae in blastic magnetite in leuchtenbergite). m = blastic magnetite, il = ilmenite. s = spinel. Hadabudussa, Gari Boro, Adola, S. Ethiopia. Oil-immersion with one nicol.

r—S ν\

m

i

0.025 mm

Fig. 294 Blastic magnetite with ilmenite lamellae. Along the margins of the magnetite with the ilmenite are bodies of spinel (dark), m = magnetite, il = ilmenite. s = spinel. Hadabudussa, Gari Boro, Adola, S. Ethiopia. Oil-immersion with one nicol.

Fig. 295 Granular chromite with oriented ex-solutions of rutile (the rutile lamellae follow the octahedral face of the chromite). r = rutile. ch = chromite. Rodiani, N. Greece. Oil-immersion with one nicol.

374

Fig. 296 Chromite with rutile and spinel lamellae oriented parallel to the octahedral face of the chromite. ch = chromite. sr = serpentine, s = spinel, r = rutile lamellae. Rodiani, N. Greece. Oil-immersion with one nicol.

Fig. 297 Granular chromite with fine rutile grains and with rutile and spinel "lamellae" oriented parallel to the octahedral face of the chromite. ch = chromite. r-g = rutile grains. Rodiani, N. Greece. Oil-immersion with one nicol.

Fig. 298 Sphalerite with replacement bodies of chalcopyrite originally interpreted as ex-solutions of chalcopyrite in sphalerite and also . chalcopyrite replacing marginally the blende. Extensions of the marginal chalcopyrite attain "ex-solution" body forms in the sphalerite. Also relics of sphalerite are maintained in the marginal chalcopyrite. s = sphalerite. c = chalcopyrite replacement bodies in the blende. c-m = marginal chalcopyrite. Arrow "a" shows extensions of chalcopyrite in the blende, arrow "b" shows relics of replacement of the sphalerite in the marginal chalcopyrite. Supija Stujena, Montenegro. Oil-immersion with one nicol.

375

Fig. 299 Marginal chalcopyrite with replacement extensions into the sphalerite attaining ex-solution-like forms. s = sphalerite. c = chalcopyrite extending into the sphalerite. ch = chalcopyrite. Supija Stujena, Montenegro. Oil-immersion with one nicol.

Fig. 300 Chalcopyrite replacement bodies with extensions following fracture lines in the sphalerite, s = sphalerite, c = chalcopyrite. Arrow shows extension of chalcopyrite following a fracture line of the sphalerite. Supija Stujena, Montenegro. Oil-immersion with one nicol.

Fig. 301 A fracture in sphalerite followed by a veinlet of chalcopyrite. Also replacement bodies of chalcopyrite simulating ex-solution bodies are present in the blende, ν = veinlet of chalcopyrite in the sphalerite, s = sphalerite. c = chalcopyrite replacement bodies in the blende. Supija Stujena, Montenegro. Oil-immersion with one nicol.

376

Fig. 302 Chalcopyrite replacing sphalerite along a pattern of cracks; also replacement chalcopyrite bodies occur (which were previously believed to be ex-solutions), s = sphalerite, c = chalcopyrite. Supija Stujena, Montenegro. Oil-immersion with one nicol.

Fig. 303 Sphalerite with oriented replacement bodies of chalcopyrite. At one point the chalcopyrite follows a fracture of the blende, s = sphalerite, c = chalcopyrite. Arrow "V" shows chalcopyrite following a fracture (interleptonic space) of the sphalerite in an interrupted veinlet-form. Arrows show the oriented direction of the chalcopyrite in the sphalerite. Supija Stujena, Montenegro. Oil-immersion with one nicol.

Fig. 304 Chalcopyrite bodies related to cracks of the sphalerite, indicating replacement of the sphalerite mainly along cracks. s = sphalerite. c = cracks in the sphalerite. ch = chalcopyrite related to cracks. ρ = pyrite. Trepca, Serbia. Oil-immersion with one nicol.

377

Fig. 305 Veinlets of chalcopyrite along cracks of the sphalerite with extensions attaining "fine ex-solutionlike" bodies, s = sphalerite. cc = cracks followed by chalcopyrite. c = extensions of chalcopyrite attaining fine "ex-solution-like" bodies in the sphalerite. Mühlenbach Mine near Ahrenberg, Germany. Oil-immersion with one nicol.

Fig. 306 Sphalerite replaced by chalcopyrite along the crystallographic face, s = sphalerite, arrows "c" show fine chalcopyrite replacing the sphalerite. Mühlenbach Mine near Ahrenberg, Germany. Oil-immersion with one nicol.

Fig. 307 Chalcopyrite bodies in sphalerite exhibiting ring patterns, s = sphalerite, c = chalcopyrite. r = ring patterns of chalcopyrite replacing the sphalerite. Supija Stujena, Montenegro. Oil-immersion with one nicol.

Fig. 308 Sphalerite with chalcopyrite interspersed and marginal to the sphalerite. cm = marginal chalcopyrite. s = sphalerite. ic = interspersed chalcopyrite in sphalerite. Shizhu Yuan near Chenzhou, Hunan Province, China. Oil-immersion with one nicol.

Fig. 309 Sphalerite with replacement bodies of chalcopyrite which simulate ex-solution bodies and with a thin margin of chalcopyrite bodies, s = sphalerite, c = chalcopyrite. Arrow shows marginal bodies of chalcopyrite supporting a replacement texture for the chalcopyrite. Shizhu Yuan near Chenzhou, Hunan Province, China. Oil-immersion with one nicol.

Fig. 310 Galena following cracks in sphalerite (veinlets of galena) are replaced by chalcopyrite bodies (which were previously believed to be ex-solutions), s = sphalerite. g = galena veinlets (following a crack pattern in the sphalerite), c = chalcopyrite bodies replacing both the galena veinlets and the sphalerite. Supija Stujena, Montenegro. Oil-immersion with one nicol.

379

Fig. 311 Galena following cracks in sphalerite (veinlets of galena) are replaced by chalcopyrite bodies (which were previously believed to be ex-solutions), s = sphalerite. g = galena veinlets (following a crack pattern in the sphalerite), c = chalcopyrite bodies replacing both the galena veinlets and the sphalerite. Supija Stujena, Montenegro. Oil-immersion with one nicol.

Fig. 312 Galena following cracks in sphalerite (veinlets of galena) are replaced by chalcopyrite bodies (which were previously believed to be ex-solutions), s = sphalerite. g = galena veinlets (following a crack pattern in the sphalerite), c = chalcopyrite bodies replacing both the galena veinlets and the sphalerite. Supija Stujena, Montenegro. Oil-immersion with one nicol.

Fig. 313 Detail of Fig. 310, showing galena veinlet transected and replaced by chalcopyrite "ex-solution" bodies, c = chalcopyrite. g = galena, s = sphalerite.

380

Fig. 314 Galena replacing sphalerite and in turn replaced by chalcopyrite. These chalcopyrite bodies in the sphalerite were previously considered to be ex-solutions, s = sphalerite, g = galena, c = chalcopyrite. Supija Stujena, Montenegro. Oil-immersion with one nicol.

Fig. 315 Galena and chalcopyrite replacing sphalerite. Both the galena and chalcopyrite bodies attain forms which were previously interpreted as ex-solution bodies, s = sphalerite, g = galena, c = chalcopyrite. Harcit, Turkey, Asia Minor. Oil-immersion with one nicol.

Fig. 316 All transitions are shown from sphalerite replaced by chalcopyrite to relics of sphalerite (attaining ex-solution-like forms) in chalcopyrite. s = sphalerite replaced by chalcopyrite (in which case the chalcopyrite attains ex-solution-like forms), c = chalcopyrite with sphalerite relics, sr = sphalerite relics. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

381

Fig. 317 All transitions are shown from sphalerite replaced by chalcopyrite to relics of sphalerite (attaining ex-solution-like forms) in chalcopyrite. s = sphalerite replaced by chalcopyrite (in which case the chalcopyrite attains ex-solution-like forms), c = chalcopyrite with sphalerite relics, sr = sphalerite relics. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

Fig. 318 All transitions are shown from sphalerite replaced by chalcopyrite to relics of sphalerite (attaining ex-solution-like forms) in chalcopyrite. s = sphalerite replaced by chalcopyrite (in which case the chalcopyrite attains ex-solution-like forms), c = chalcopyrite with sphalerite relics, sr = sphalerite relics. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

Fig. 319 All transitions are shown from sphalerite replaced by chalcopyrite to relics of sphalerite (attaining ex-solution-like forms) in chalcopyrite. c = chalcopyrite with sphalerite relics, sr = sphalerite relics. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

382

Fig. 320 Sphalerite with zonal chalcopyrite and chalcopyrite attaining ex-solution-like forms. Chalcopyrite and pyrite are also present. s = sphalerite. zc = zonal chalcopyrite. c = chalcopyrite attaining "exsolution" forms. ch = chalcopyrite. ρ = pyrite. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

Fig. 321 Sphalerite with fine interzonal chalcopyrite (zonal distribution of chalcopyrite previously believed to be "ex-solutions" in the sphalerite), s = sphalerite. cz = interzonal chalcopyrite. ρ = pyrite. c = chalcopyrite (often attaining "exsolution"-like forms). Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

Fig. 322 Sphalerite with fine interzonal chalcopyrite also chalcopyrite attaining "ex-solution" forms. s = sphalerite. cz = interzonal chalcopyrite. c = chalcopyrite. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

383

Fig. 323 Sphalerite with interzonal and "ex-solution"-like forms of chalcopyrite. s = sphalerite. cz = interzonal chalcopyrite. c = chalcopyrite. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

Fig. 324 Sphalerite with interzonal and "ex-solution"-like forms of chalcopyrite. s = sphalerite. cz = interzonal chalcopyrite. c = chalcopyrite. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

Fig. 325 Sphalerite core free of chalcopyrite bodies, with "zonal" parts with fine interspersed chalcopyrite. c = core of sphalerite. zs = zonal sphalerite with fine chalcopyrite bodies. Supija Stujena, Montenegro. Oil-immersion with one nicol.

384

Fig. 326 Zonal replacement of sphalerite by stannite attaining "exsolution-like" forms, s = sphalerite, st = stannite. m = marcasite. Pali Mine, Dachang, China. Oil-immersion with one nicol.

Fig. 327 Zonal replacement of sphalerite by stannite attaining "exsolution-like" forms, s = sphalerite, st = stannite. m = marcasite. Pali Mine, Dachang, China. Oil-immersion with one nicol.

Fig. 328 Zonal replacement of sphalerite by stannite attaing "exsolution-like" forms, s = sphalerite, st = stannite. m = marcasite. Pali Mine, Dachang, China. Oil-immersion with one nicol.

385

Fig. 329 Galena and chalcopyrite as replacement bodies in sphalerite simulating ex-solution forms, g = galena, s = sphalerite, c = chalcopyrite. Wiesloch, Baden, Germany. Oil-immersion with one nicol.

Fig. 330 Galena replacing sphalerite and with extensions attaining "exsolution"-like shapes (bodies), s = sphalerite, g = galena. e = extensions of galena. Harcit, Turkey, Asia Minor. Oil-immersion with one nicol.

Fig. 331 Sphalerite with Staub (dust) fine interspersed bodies of pyrite. s = sphalerite. ρ = fine interspersed pyrite bodies in sphalerite. Websky, Turkey, Asia Minor. Oil-immersion with one nicol.

386

Fig. 332 Sphalerite with Staub (dust) fine interspersed bodies of pyrite. s = sphalerite. ρ = fine interspersed pyrite bodies in sphalerite. Websky, Turkey,Asia Minor. Oil-immersion with one nicol.

Fig. 333 Banded gel sphalerite with interzonal (inter-gel-zonal) pyrite. s = sphalerite, ρ = pyrite. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

Fig. 334 Pyrite replacing sphalerite and attaining pseudo "ex-solution" forms (see arrow "a"), ρ = pyrite. s = sphalerite. Chang Poy, Dachang, China. Oil-immersion with one nicol.

0.1 mm

387

Fig. 335 Arsenopyrite following cracks in the sphalerite and attaining "ex-solution" shapes in the host, a = arsenopyrite. s = sphalerite. San Martin Mine, Sombrerete, Zacatecas, Mexico. Oil-immersion with one nicol.

Fig. 336 Sphalerite (stars) X-shaped bodies metasomatically replacing chalcopyrite. c = chalcopyrite. s = sphalerite (X-shaped bodies), arrows show the trend followed by the sphalerite bodies in the chalcopyrite. Carris Mine, Serra Genez, N. Portugal. Oil-immersion with one nicol.

§ 9

• c

388

Fig. 337 In contradistinction to late metasomatic X-shaped sphalerite bodies in the chalcopyrite (see Fig. 336), the present photomicrograph shows sphalerite relics left in the chalcopyrite after the corrosion and replacement of sphalerite by chalcopyrite. s = sphalerite, c = chalcopyrite. g = galena. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

κ c

Fig. 338 Schematic diagram showing cross-shaped relics of sphalerite due to the replacement of sphalerite by pyrite. In the central part of the cross relic of sphalerite in the pyrite, chalcopyrite initially associated with the sphalerite is also preserved. The cross-shaped relic might indicate that it is related to the [100] face of the pyrite ? crystalloblast. ρ = pyrite. s = sphalerite, c = chalcopyrite.

Fig. 339 Pyrrhotite replaced by pentlandite along fractures or extending from fractures of the pyrrhotite. Py = pyrrhotite. ρ = pentlandite flame-like bodies starting from a crack in the pyrrhotite and extending inwards, pg = granular pentlandite having extensions in the crack in the pyrrhotite. Alexo Mine, Dundonald Township, Ontario, Canada. Oil-immersion with one nicol.

Fig. 340 Pentlandite metasomatically replacing pyrrhotite starting from the margin and extending inwards in the pyrrhotite. py = pyrrhotite. ρ = pentlandite replacing pyrrhotite marginally. fp = metasomatic flame pentlandite in the pyrrhotite. Alexo Mine, Dundonald Township, Ontario, Canada. Oil-immersion with one nicol.

389

Fig. 341 Pyrrhotite with pentlandite replacement along cracks in the central part of the pyrrhotite. Pentlandite also occurs marginally to the pyrrhotite. py = pyrrhotite. ρ = pentlandite. Alexo Mine, Dundonald Township, Ontario, Canada. Oil-immersion with one nicol.

Fig. 342 Pentlandite marginally and in flame form replacing pyrrhotite. py = pyrrhotite. ρ = pentlandite. fp = flame form pentlandite replacing pyrrhotite along a fracture. Alexo Mine, Dundonald Township, Ontario, Canada. Oil-immersion with one nicol.

Fig. 343 Zinkite with fine haematite pseudo-ex-solutions. ζ = zinkite. h = haematite pseudo-ex-solutions. Franklin-Fournace, New Jersey, USA. Oil-immersion with one nicol.

390

Fig. 344 Zinkite veinlets extending along fractures and replacing franklinite. ζ = zinkite. f = franklinite. Franklin-Fournace, New Jersey, USA. Oil-immersion with one nicol.

Fig. 345 Zinkite surrounding franklinite and with zinkite (often developed small crystals) interspersed in the franklinite. ζ = zinkite. f = franklinite. zc = small zinkite crystals interspersed particularly in the central part of the franklinite. Franklin-Fournace, New Jersey, USA. Oil-immersion with one nicol.

imjj-'W* Pfm Fig. 346 Zinkite (often exhibiting crystalline outlines) following directions or irregularly interspersed in franklinite. ζ = zinkite. f = franklinite. Franklin Fournace, New Jersey, USA. Oil-immersion with one nicol.

391

Fig. 347 Zinkite (often exhibiting crystalline outlines) following directions or irregularly interspersed in franklinite. ζ = zinkite. f = franklinite. Franklin-Fournace, New Jersey, USA. Oil-immersion with one nicol.

r*

-

'•»• : : * Λ-i -v ' ν Λ * k» 4

ν tj;

*

ί J

I*

* * - "»Ν * .

{

.4,1 Jb>4···

Λ;.,*

' ·,

-

«

fei \

: ί

> .

Fig. 348 Small zinkite crystals following crystallographic directions (oriented) within the franklinite host, ζ = zinkite. f = franklinite. Franklin-Fournace, New Jersey, USA. Oil-immersion with one nicol.

η

i ΛÄί Α? * · ί •··>·Λ } ',/· • : ;ί · ν. Ä i »*

·.

ι

Τ . , -«y

' '.J' % · ·,-»· · r

Fig. 349 Ilmenite with haematite replacement bodies. In cases the haematite follows boundaries of ilmenite grains, see arrows. Spinel replacing partly the haematite which in turn replaces ilmenite. il = ilmenite. h = haematite. s = spinel following cracks in the ilmenite. sr = spinel replacing the haematite. Otanmäki, Finland. Oil-immersion with one nicol.

392

Fig. 350 Spinel and haematite (with fine ilmenite ex-solutions) partly following a crack pattern or boundaries of ilmenite with fine haematite. il = ilmenite with fine haematite, s = spinel following a crack pattern in the ilmenite (or ilmenite boundaries), h = haematite (with fine ilmenite) following ilmenite boundaries (see arrows). Otanmäki, Finland. Oil-immersion with one nicol.

Fig. 351 Spinel following cracks in ilmenite. Also haematite lamellae partly following boundaries (see arrow) of ilmenite grains (of the main ilmenite mass), il = ilmenite. s = spinel. sr = spinels most probably replacing haematite lamellae. Otanmäki, Finland. Oil-immersion with one nicol.

Fig. 352 Ilmenite with haematite replacement bodies initiating from a curving crack in the ilmenite. Similarly spinel replaces the ilmenite. il = ilmenite. h = haematite, arrows show curving cracks, s = spinel. Otanmäki, Finland. Oil-immersion with one nicol.

393

Fig. 353 Ilmenite with fine haematite bodies. However, along a crack of the ilmenite the latter is free from haematite bodies, il = ilmenite. il-h = ilmenite with fine haematite bodies. Otanmäki, Finland. Oil-immersion with one nicol.

Fig. 354 Initial ilmenite with exsolutions of haematite partly enclosed by anthophyllite (with internal reflection). Hadabudussa, Gari-Boro, S. Ethiopia. Oil-immersion with one nicol.

Fig. 355 Corundum crystalloblast enclosing haematite with ilmenite "ex-solutions". The haematites enclosed by the blastic corundum show corroded and rounded outlines the rounding of the haematite grains took place after the formation of the ilmenite. c = corundum. a = haematite with lamellar ilmenite bodies enclosed and corroded by the corundum crystalloblast. Emery, Naxos, Greece. Oil-immersion with one nicol.

394

Fig. 356 Elongated haematite ilmenite crystal is partly altered to ilmenite and rutile. h = haematite, i = ilmenite associated with haematite. il = ilmenite replacing haematite, r = rutile. Emery, Naxos, Greece. Oil-immersion with one nicol. χ 750.

Fig. 357 Haematite crystals (h) with lamellae of rutile parallel to the second pyramid position (hh 2h 1) in contact with martitized magnetite. The rutile lamellae are also to be found in the magnetite which has replaced the haematite (by reduction), h = haematite, r = rutile. ma = magnetite, m = martite. Emery, Naxos, Greece. Oil-immersion with one nicol. χ 1500.

Fig. 358 Davidite with haematite lamellae parallel oriented, most probably representing a replacement of davidite by haematite, d = davidite. h = haematite with ilmenite. In the haematite fine bodies of ilmenite are present. Harrar, Harrarge, Ethiopia Oil-immersion with one nicol.

395

Fig. 359 Davidite associated with ilmenite bodies indicating complex myrmekitic replacement by haematite. Rutile grains are also present, d = davidite. il = ilmenite. h = complex myrmekitic haematite replacing ilmenite. r = rutile. da = alteration margins of davidite. Whip Lode, Radium Hill, Australia. Oil-immersion with one nicol.

Fig. 360 Davidite with granular aggregate of haematite with lamellar ilmenite as scattered inclusions partly replaced by the davidite in which they are included, d = davidite. h = haematite with ilmenite bodies. Olary, Southern Australia. Oil-immersion with one nicol.

Fig. 361 A gigantic rutile in davidite. The boundary rutile/davidite is indicating mutual replacements, r = rutile. d = davidite, arrows show mutual replacements. Olary, Southern Australia. Oil-immersion with one nicol.

396

Fig. 362 Homogenous davidite showing small rutile grains along curved lines caused by tectonic effect, d = davidite. r = rutile, arrows show curved tectonic line. Mount Victoria Mine, Southern Australia. Oil-immersion with one nicol.

Fig. 363 Columbite including uraninite infiltrations and replacements, c = columbite. ur = uraninite. Hagendorf, Bavaria, Germany. Oil-immersion with one nicol.

Fig. 364 Olivine with magnetite fine lamellae and with magnetite partly following cracks of the olivine, ol = olivine, m = magnetite. cm = magnetite following crack of the olivine. Pretoria, South Africa. Oil-immersion with one nicol.

397

Fig. 365 Oriented lamellar sphalerite in Bi with chloanthite margin, s = sphalerite lamellae, b = Bi. c = chloanthite. Luisito Mine near Pozoblanco, Cordoba, Spain. Oil-immersion with one nicol.

Fig. 366 Cuprite replaced by native copper. c = cuprite. cu = native copper. Cerro de Pasco, Peru. Oil-immersion with one nicol.

Fig. 367 Silver formed by leaching out of Sb and S from pyrargyrite. The irregular-in-shape silver grains are following an approximate zone in the host pyrargyrite. Py = pyrargyrite. s = silver. Beaver Mine, Cobalt City, Ontario, Canada. Oil-immersion with one nicol.

398

Fig. 368 Replacement symplectic intergrowth of siegenite replacing uraninite, resulting in crystal face oriented intergrowth. s = siegenite. u = uraninite. Shinkolobwe, Katanga, Zaire. Oil-immersion with one nicol.

Fig. 369 A Maltese cross-shaped siegenite skeleton replacement structure in uraninite. The siegenite replaces the uraninite but has not attained idiomorphic shape. It is an arrested state of development, s = siegenite. u = uraninite. Shinkolobwe, Katanga, Zaire. Oil-immersion with one nicol.

Fig. 370 Graphic-like replacement intergrowth of siegenite with uraninite. s = siegenite (skeletal crystal of siegenite). u = uraninite. Shinkolobwe, Katanga, Zaire. Oil-immersion with one nicol.

399

Fig. 371 Siegenite-uraninite symplectite. Siegenite replacing uraninite and resulting in face to face crystal orientation, s = siegenite "veinlets". se = siegenite extension in face to face crystal orientation alternating with uraninite. u = uraninite. Shinkolobwe, Katanga, Zaire. Oil-immersion with one nicol.

Fig. 372 Siegenite-uraninite symplectite. s = siegenite in face to face orientation with uraninite. u = uraninite. sv = veinlet of siegenite as extension of the main siegenite into the uraninite. Shinkolobwe, Katanga, Zaire. Oil-immersion with one nicol.

Fig. 373 Siegenite veinlet partly transecting the uraninite and partly replacing it. s = siegenite. u = uraninite, arrow shows a relic of uraninite in the siegenite. Similarly relics of uraninite are to be seen within the siegenite veinlet. Shinkolobwe, Katanga, Zaire. Oil-immersion with one nicol.

400

Fig. 374 Symplectic replacement intergrowth of uraninite and columbite. The uraninite partly exhibits crystalline outlines, u = uraninite. c = columbite. Hagendorf, Bavaria, Germany. Oil-immersion with one nicol.

w m m m m m m g

c

>

,

V

Fig. 375 Symplectic intergrowth of uraninite replacing columbite. Often the shape of the uraninite in the columbite is controlled by cracks of the latter (see arrow), u = uraninite. c = columbite. Hagendorf, Bavaria, Germany. Oil-immersion with one nicol.

Fig. 376 Symplectic uraninite replacing columbite. u = uraninite. c = columbite. Hagendorf, Bavaria, Germany. Oil-immersion with one nicol.

401

Fig. 377 Jacobsite in symplectic intergrowth replacing braunite (in cases the jacobsite is epitactic on the braunite). j = jacobsite. b = braunite. Otjosondu, Sandfeld, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 378 Titanite in symplectic intergrowth magnetite. The titanite replaces the magnetite marginally and extends inwards, m = magnetite, t = titanite. Londao, Coimbra, Portugal. Oil-immersion with one nicol.

Fig. 379 Oriented lamellae of ilmenite in pyrrhotite (probably replaced initial magnetite). The pyrrhotite is partly changed to marcasite. il = ilmenite. ρ = pyrrhotite. m = marcasite. Frood Mine, Sudbury, Ontario, Canada. Oil-immersion with one nicol.

402

Fig. 380 Pyrrhotite partly altered to marcasite in which part an ilmenite lamella is preserved. In contradistinction oriented ilmenite lamellae occur in the adjacent silicate. Most probably the ilmenite lamellae were in intergrowth with an initial magnetite which has been replaced successively by pyrrhotite and this by marcasite and finally silicate. The ilmenite lamellae (which were perhaps oriented in the initial magnetite) have been maintained despite the complete and extensive replacement of the original magnetite by pyrrhotite, marcasite and finally by the silicate, ρ = pyrrhotite. m = marcasite. s = silicate, il = ilmenite. Frood Mine, Sudbury, Ontario, Canada. Oil-immersion with one nicol.

Fig. 381 Graphic-like sphalerite (skeletal crystal) in intergrowth with silicate. s = sphalerite. si = silicate. Zacatecas, Mexico. Oil-immersion with one nicol.

Fig. 382 Graphic-like sphalerite (skeletal crystal) in intergrowth with silicate. s = sphalerite. si = silicate. Zacatecas, Mexico. Oil-immersion with one nicol.

403

Fig. 383 Sphalerite symplectically intergrown with silicates (quartz). Also galena is symplectically grown (replacing) sphalerite, s = sphalerite, si = silicate, g = galena, ρ = pyrite. Zacatecas, Mexico. Oil-immersion with one nicol.

Fig. 384 Gudmundite crystalloblastic growths in chalcopyrite and tetrahedrite. Also graphic-in-form sphalerite in intergrowth with chalcopyrite is shown, t = tetrahedrite. c = chalcopyrite. g = gudmundite crystalloblast. s = sphalerite micrographic-in-shape, as indicated by arrow. Seinajoki, Finland. Oil-immersion with one nicol.

Fig. 385 Skeletal crystals of galena and idioblastic pyrite replacing sphalerite, g = galena, s = sphalerite, ρ = pyrite. Roseberg Mine. Oil-immersion with one nicol.

404

Fig. 386 Chalcopyrite and pyrrhotite skeletal crystals (in graphic intergrowths with silicates - black), c = chalcopyrite. ρ = pyrrhotite. Armerzbalde, Insizwa, East Griqualand, South Africa. Oil-immersion with one nicol.

Fig. 387 Symplectic bornite with silicate (graphic intergrowths) and with chalcopyrite replacing the bornite. b = bornite. c = chalcopyrite. Roan Antelope, Zambia. Oil-immersion with one nicol.

Fig. 388 Galena and chalcopyrite in graphic intergrowth with silicates, c = chalcopyrite. g = galena. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

405

Fig. 389 Skeletal crystals of Bi in intergrowth with silicates. Marginally the Bi is replaced by brown iron, b = Bi. br = brown iron, s = silicate. Espuela San Miguel near Villanueva, Cordoba, Spain. Oil-immersion with one nicol.

Fig. 390 Native Bi in symplectic intergrowth with silicates, b = Bi. s = silicates. Daniel Mine near Wittichen, Germany. Oil-immersion with one nicol.

Fig. 391 Gold in symplectic intergrowth with silicate, g = gold s = silicate. Nancy Mine, Boulder County, Colorado, USA. Oil-immersion with one nicol.

406

Fig. 392 Sylvanite in graphic-like symplectic intergrowth with silicates, sy = sylvanite. s = silicates. Nancy Mine, Boulder County, Colorado, USA. Oil-immersion with one nicol.

Fig. 393 Symplectic (graphic-like) intergrowth of petzite with silicates. Sylvanite is also symplectic. ρ = petzite. s = sylvanite. si = silicate. Nancy Mine, Boulder County, Colorado, USA. Oil-immersion with one nicol.

Fig. 394 Symplectic petzite (graphiclike intergrowths) with silicates, ρ = petzite. q = silicate (quartz). Nancy Mine, Boulder County, Colorado, USA. Oil-immersion with one nicol.

407

Fig. 395 Skeletal clausthalite (graphic-like intergrowths) in symplectic intergrowth with calcite. cl = clausthalite. c = calcite. Tilkerode, Harz, Germany. Oil-immersion with one nicol.

Fig. 396 Graphic-in-shape (skeletal crystals) of clausthalite in intergrowth with calcite. Tilkerode, Harz, Germany. Oil-immersion with one nicol.

Fig. 397 Uraninite in graphic-like intergrowth with microcline. All transitions from graphic-like to idiomorphic crystalline form of the uraninite can be followed, u = uraninite. m = microcline. r = radioactive halo. Ruggles Mine, New Hampshire, USA. Oil-immersion with one nicol.

408

Fig. 398 Uraninite in graphic-like intergrowth with microcline. All transitions from graphic-like to idiomorphic crystalline form of the uraninite can be followed, u = uraninite. m = microcline. r = radioactive halo. Ruggles Mine, New Hampshire, USA. Oil-immersion with one nicol.

Fig. 399 A uraninite crystal which shows well-developed crystalline faces and simultaneously exhibits symplectic intergrowth with the microcline in which it is included, u = crystalline uraninite. m = microcline. Ruggles Mine, New Hampshire, USA. Oil-immersion with one nicol.

Fig. 400 A uraninite crystal which shows well-developed crystalline faces and simultaneously exhibits symplectic intergrowth with the microcline in which it is included, u = crystalline uraninite. m = microcline. Ruggles Mine, New Hampshire, USA. Oil-immersion with one nicol.

409

Fig. 401 The last phase of development of the series of patterns exhibited in Figs. 397-401, illustrating an idiomorphic crystal (cubic) of uraninite in the microcline. u = uraninite. r = radioactive halo, m = microcline. Ruggles Mine, New Hampshire, USA. Oil-immersion with one nicol.

Fig. 402 Skeletal pyrite (graphic-like intergrowths) crystalloblast replacing sphalerite, ρ = pyrite. s = sphalerite, g = galena. Roseberg Mine. Oil-immersion with one nicol.

Fig. 403 Idiomorphic pyrite crystals as the last phase of development of skeletal pyrite to fully idiomorphic forms. Roseberg Mine. Oil-immersion with one nicol.

410

Fig. 404 Skeletal (graphic-like intergrowths) of pyrite and chalcopyrite with silicates, ρ = skeletal pyrite. c = chalcopyrite (graphic-like intergrowth with the silicate), s = sphalerite. Roseberg Mine. Oil-immersion with one nicol.

Fig. 405 Skeletal pyrite (graphic-like forms) with pyrargyrite in which relics of prousite are present, py = pyrargyrite. pr = prousite. ρ = pyrite. Freiberg, Saxony, Germany. Oil-immersion with one nicol.

Fig. 406 Prismatic silicates in intergrowth with telluride and native gold. s = silicate, t = telluride. g = native gold. Lake Shore Mine, Kirkland Lake, Ontario, Canada. Oil-immersion with one nicol.

411

Fig. 407 Skeletal gold-silicate intergrowths. Siegenite is also present, g = gold, u = uraninite. se = siegenite. s = ?silicates. Shinkolobwe, Katanga, Zaire. Oil-immersion with one nicol.

Fig. 408 Rutile in davidite with haematite, symplectic (myrmekiticlike intergrowth) in the rutile. d = davidite. r = rutile. h = haematite (myrmekitic intergrowth with rutile). Whip Lode, Radium Hill, Australia. Oil-immersion with one nicol.

Fig. 409 Davidite including rutile grains. A case is exhibited of rutile replacement by haematite attaining a symplectic myrmekitic-like intergrowth of haematite and rutile. (Actually the rutile is replaced by a network of haematite.) d = davidite. r = rutile. rh = haematite in symplectic intergrowth with large rutile grain (a veinlet of haematite replaces the rutile). Whip Lode, Radium Hill, Australia. Oil-immersion with one nicol.

412

Fig. 410 Wiikite replaced by rutile (symplectic rutile/wiikite intergrowth). Also idioblastic rutile is shown. rs = symplectic rutile. br = blastic rutile. w = wiikite. Lake Ladoga, Impilahti, Finland. Oil-immersion with one nicol.

Fig. 411 Different symplectic patterns due to replacement of neodigenite by covellite whereby in cases a lamellar-type of replacement of neodigenite by covellite has taken place (see particularly Fig. 413). η = neodigenite. c = covellite. Bisbee, Arizona, USA. Oil-immersion with one nicol.

Fig. 412 Different symplectic patterns due to replacement of neodigenite by covellite whereby in cases a lamellar-type of replacement of neodigenite by covellite has taken place (see particularly Fig. 413). η = neodigenite. c = covellite. Bisbee, Arizona, USA. Oil-immersion with one nicol.

413

Fig. 413 Different symplectic patterns due to replacement of neodigenite by covellite whereby in cases a lamellar-type of replacement of neodigenite by covellite has taken place. η = neodigenite. c = covellite. Bisbee, Arizona, USA. Oil-immersion with one nicol.

Fig. 414 Drop-like intergrowths of sphalerite in intergrowth with gangue. Sphalerite spheroids are also exhibited. s = "drop"-like sphalerite in intergrowth with gangue. ss = sphalerite spheroids, g = galena. Mechernich, Eifel, Germany. Oil-immersion with one nicol.

Fig. 415 Skeletal sphalerite in intergrowth with gangue. Also spheroidal sphalerite in case with a central part of galena. Such spheroidal structures could represent (despite their relationship to volcanism) colloform structures from solutions, g = galena, s = sphalerite. Mechernich, Eifel, Germany. Oil-immersion with one nicol.

414

Fig. 416 Ring-like sphalerite and feather-like sphalerite structures in intergrowth with gangue. Despite their relationship to volcanic activity, the patterns could represent colloform structures, subsequently crystallized, rs = ring structure of sphalerite, fs = feather-like structure of sphalerite. Mechernich, Eifel, Germany. Oil-immersion with one nicol.

Fig. 417 Idiomorphic initial titanomagnetite, in the central part replaced by heat haematite and the ilmenite is changed to pseudobrookite. m = magnetite, h = heat haematite, ρ = pseudobrookite often exhibiting a distribution different to the initial oriented ilmenite lamellae in the magnetite. ma = heat martitization. Katzenbuckel, Germany. (Sanidinnephelinit). Oil-immersion with one nicol.

Fig. 418 Magnetite with ilmenite lamellae "pseudomorphosed" replaced by haematite and the ilmenite changed to pseudobrookite. Also heat martitization is shown, m = magnetite. h = haematite replacing the magnetite, ρ = pseudobrookite. hm = heat martitization. The haematite replacement of the magnetite attains graphic-like symplectic form (see arrows). Katzenbuckel, Germany. (Sanidinnephelinit). Oil-immersion with one nicol.

415

Fig. 419 Initial titanomagnetite pseudomorphosed. Haematite replaces the magnetite exhibiting symplectic forms. m = magnetite, h = haematite, ρ = pseudobrookite (?pseudomorphism of initial ilmenite lamellae in the magnetite). Katzenbuckel, Germany. (Sanidinnephelinit). Oil-immersion with one nicol.

Fig. 420 Initial titanomagnetite changed due to thermal effects in which case the pseudobrookite follows the initial lamellar distribution of the ilmenite in the magnetite. Magnetite is also replaced by heat haematite due to increased thermal influence, m = magnetite, ρ = pseudobrookite h = heat haematite. Katzenbuckel, Germany. (Sanidinnephelinit). Oil-immersion with one nicol.

Fig. 421 Magnetite and symplectic magnetite (partly changed into maghemite) intergranular to calcite (marble). m = magnetite. ms = symplectic magnetite. ma = marble. mh = maghemite. Halara, Seriphos, Greece. Oil-immersion with one nicol.

Fig. 422 Magnetite protruding into the marble and attaining myrmekitic symplectic form. The symplectic magnetite has partly changed to maghemite. m = magnetite. ms = magnetite symplectic (myrmekitic in form). mh = maghemite due to change of the symplectic magnetite. ma = marble. Halara, Seriphos, Greece. Oil-immersion with one nicol.

Fig. 423 Bornite in symplectic myrmekitic-like intergrowth (with bismuthinite or emplectite). The symplectite partly replaces pitchblende. However, as arrow shows, the outer layers of the pitchblende also replace and surround the symplectite. b = bornite. bi = bismuthinite (?emplectite). ρ = pitchblende. Schmiedeberg, Kowary, Poland. Oil-immersion with one nicol.

Fig. 424 Symplectite of bornite with bismuthinite or emplectite. Pitchblende is abundant in the section, b = bornite. bi = bismuthinite (or emplectite). ρ = pitchblende. Schmiedeberg, Kowary, Poland. Oil-immersion with one nicol.

417

Fig. 425 Sphalerite with pyrrhotite in symplectic intergrowth (the pyrrhotite replaces the sphalerite), b = blende (sphalerite), ρ = pyrrhotite. South Bay, Meldon Mine, Victoria, Australia. Oil-immersion with one nicol.

Fig. 426 Sphalerite and pyrrhotite symplectite with gangue. s = sphalerite, ρ = pyrrhotite. South Bay, Meldon Mine, Victoria, Australia. Oil-immersion with one nicol.

Fig. 427 Graphic/myrmekitic galena replacing sphalerite, g = galena, s = sphalerite. Tilkedore, Harz, Germany. Oil-immersion with one nicol.

418

Fig. 428 Graphic/myrmekitic galena replacing sphalerite, g = galena, s = sphalerite. Tilkedore, Harz, Germany. Oil-immersion with one nicol.

Fig. 429 Graphic/myrmekitic galena replacing sphalerite. Sub-spheroidal sphalerite relics are preserved in the galena, g = galena, s = sphalerite. sb = sub-spheroidal sphalerite. Tilkedore, Harz, Germany. Oil-immersion with one nicol.

Fig. 430 Graphic-like pyrite and galena (in symplectic intergrowth) replacing germanite. ρ = pyrite. ge = germanite. g = galena. Tsumeb, Otavi Bergland District, Namibia, S. W. Africa. Oil-immersion with one nicol.

419

Fig. 431 Germanite replaced by galena and pyrite; particularly galena shows patterns comparable to those of "unmixing", ge = germanite. g = galena, ρ = pyrite. Tsumeb, Otavi Bergland District, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 432 Germanite replaced by galena and both transected by a later veinlet of tennantite. ge = germanite. g = galena, t = tennantite. Tsumeb, Otavi Bergland District, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 433 Galena and graphic-like galena replacing sphalerite. Pyrite is also shown. g = galena. gg = graphic galena. s = sphalerite. ρ = pyrite. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

420

Fig. 434 A veinlet of galena (actually galena in symplectic intergrowth with sphalerite) extends into the sphalerite, g = galena, s = sphalerite, ρ = pyrite. Weisser Hirsch Vein, Goslar, Germany. Oil-immersion with one nicol.

Fig. 435 Galena marginally replacing sphalerite. Also graphic/myrmekitic intergrowths of galena (due to replacement) are in the sphalerite, g = galena, s = sphalerite. gg = graphic/myrmekitic-like galena in the sphalerite. Goppenstein, Wallis, Switzerland. Oil-immersion with one nicol.

Fig. 436 Both galena and chalcopyrite (myrmekitic/graphic) bodies replacing sphalerite, g = galena, s = sphalerite, c = chalcopyrite. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

421

Fig. 437 Chalcopyrite in symplectic intergrowth with sphalerite exhibiting all transitions from marginal chalcopyrite to chalcopyrite attaining myrmekitic-graphic forms in the sphalerite, ch = chalcopyrite. s = sphalerite. c = myrmekitic-like symplectite of chalcopyrite in replacement intergrowth with the sphalerite. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

Fig. 438 Galena and chalcopyrite veinlets transversing sphalerite and extensions of the veinlet attaining symplectic forms, s = sphalerite. vc = chalcopyrite of the veinlet. vg = galena of the veinlet, arrow shows extensions of the chalcopyrite veinlet attaining myrmekitic/graphiclike form, g = galena, c = chalcopyrite. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

Fig. 439 Chalcopyrite replacing bornite and in cases the replacement results in symplectitic intergrowth pattern of chalcopyrite and bornite. c = symplectite of bornite replaced by chalcopyrite. ch = chalcopyrite. b = bornite. Roan Antilope, Zambia. Oil-immersion with one nicol.

422

Fig. 440 Bornite relics in chalcopyrite. The relics attain a myrmekitic/graphic form in intergrowth with the replacing chalcopyrite. c = chalcopyrite. b = bornite. bs = bornite in intergrowth with silicates (which might also represent a replacement of the bornite by the gangue). However, this pattern can only be understood in conjunction with Fig. 439. Roan Antilope, Zambia. Oil-immersion with one nicol.

Fig. 441 Chalcopyrite replacing enargite. e = enargite. c = chalcopyrite extending into the enargite and resulting in an enarite chalcopyrite symplectite. S. Joäo Deserto, Aljustral, S. Portugal. Oil-immersion with one nicol.

Fig. 442 Chalcopyrite in form of myrmekite and lamellae replacing enargite. Also fine chalcopyrite lamellae are present in enargite. e = enargite. c = chalcopyrite myrmekitic in form replacing bornite. 1 = lamellae of chalcopyrite replacing enargite. Tarvis Township, Algoma District, Ontario, Canada. Oil-immersion with one nicol.

423

Fig. 443 Enargite replaced by chalcopyrite attaining myrmekitic-like forms, e = enargite. c = chalcopyrite. Tarvis Township, Algoma District, Ontario, Canada. Oil-immersion with one nicol.

Fig. 444 Bornite/chalcocite symplectite; the chalcocite replaces the bornite. b = bornite. c = chalcocite. Kupferberg, Fichtelgebirge, Germany. Oil-immersion with one nicol.

Fig. 445 Chalcocite with extensions attaining myrmekitic-like forms (intergrown with the bornite) in the adjacent bornite. b = bornite. c = chalcocite. Kupferberg, Fichtelgebirge, Germany. Oil-immersion with one nicol.

424

Fig. 446 Chalcocite replacing bornite. Bornite relics in the chalcocite result in myrmekitic-like intergrowths of bornite and chalcocite. Also extensions of chalcocite in the bornite is shown by arrow "a", b = bornite. c = chalcocite, arrows show probable extensions of the bornite into the chalcopyrite. Kupferberg, Fichtelgebirge, Germany. Oil-immersion with one nicol.

Fig. 447 Bomite replaced by marginal chalcocite with an extension of it in the bomite attaining myrmekitic/symplectic forms with the bornite (see also Fig. 448). b = bornite. c = chalcocite, arrow shows extension of chalcocite in the bornite. m = myrmekitic symplectite of the chalcocite extension with the bornite. s = sphalerite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 448 Bornite replaced by marginal chalcocite with an extension of it in the bornite attaining myrmekitic/symplectic forms with the bomite. b = bomite. c = chalcocite, arrow shows extension of chalcocite in bomite. m = myrmekitic symplectite of the chalcocite extension with the bomite. s = sphalerite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

425

Fig. 449 Chalcocite replacing bornite. Relics of bomite in the chalcocite result in bomite/chalcocite myrmekitic/symplectic intergrowths. Also extensions (see arrow) of chalcocite into the bomite attain myrmekitic-like symplectic forms, b = bornite. c = chalcocite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 450 Chalcocite replacing bornite. Relics of bornite in the chalcocite result in bornite/chalcocite myrmekitic/symplectic intergrowths. Also extensions (see arrows) of chalcocite into the bornite attain myrmekitic-like symplectic forms, b = bomite. c = chalcocite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 451 Bornite replaced by chalcocite and resulting in a myrmekitic-like bomite/chalcocite symplectite. As a result of the replacement of bomite by chalcocite, atoll-type structures are shown, b = bomite. c = chalcocite. ρ = pyrite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

426

Fig. 452 Bornite replaced by chalcocite and the replacement pattern attains symplectic myrmekitic form, b = bornite. c = chalcocite. ρ = pyrite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 453 Bornite replaced by chalcocite. A part of the bornite replaced by chalcocite exhibits myrmekitic-like intergrowth of bornite and chalcocite (which is due to replacement of bornite by chalcocite). b = bornite. c = chalcocite. m = myrmekitic symplectic intergrowth of bornite and chalcocite due to the replacement of bornite by chalcocite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 454 Symplectic myrmekitic intergrowth of bornite and chalcocite in which case the chalcocite has replaced the bornite. c = chalcocite. b = bornite. ρ = idioblastic pyrite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

427

Fig. 455 An advanced phase of replacement of bornite by chalcocite resulting in a symplectic myrmekitic pattern of bomite and chalcocite. c = chalcocite. b = bornite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 456 Pyrrhotite replaced by galena where the latter simulates exsolution forms. ΡΥ = pyrrhotite. g = galena. Boliden, Sweden. Oil-immersion with one nicol.

Fig. 457 Corollary to Fig. 456, showing galena replacing pyrrhotite. ρ = pyrrhotite. g = galena. Boliden, Sweden. Oil-immersion with one nicol.

428

Fig. 4 5 8 Tetrahedrite and wittichenite in symplectic intergrowth with Bi. t = tetrahedrite. w = wittichenite. b = Bi. Daniel Mine near Wittichen, Black Forest, Germany. Oil-immersion with one nicol.

Fig. 459 Tetrahedrite and wittichenite in symplectic intergrowth (replacement) by B i . t = tetrahedrite. w = wittichenite. b = Bi. Daniel Mine near Wittichen, Black Forest, Germany. Oil-immersion with one nicol.

Fig. 4 6 0 B i in symplectic intergrowth (replacement) with wittichenite. b = Bi. w = wittichenite. Daniel Mine near Wittichen, Black Forest, Germany. Oil-immersion with one nicol.

429

Fig. 461 Tetrahedrite replaced by wittichenite (arrow shows tetrahedrite symplectite relics in the wittichenite). Bi in symplectic intergrowth with (replacing) wittichenite. t = tetrahedrite. w = wittichenite. b = Bi. Daniel Mine near Wittichen, Black Forest, Germany. Oil-immersion with one nicol.

Fig. 462 Chloanthite in symplectic intergrowth with Bi. c = chloanthite. b = Bi. Luisito Mine near Pozoblanco, Cordoba, Spain. Oil-immersion with one nicol.

Fig. 463 Bornite in symplectic intergrowth with chalcopyrite. Myrmekitic-like relics of bornite due to its replacement by chalcopyrite. b = bornite. c = chalcopyrite. ch = chalcocite veinlets replacing chalcopyrite. (occasionally pyrite is present). Matchless Mine, Friedenau Farm, Windhoek, South Africa. Oil-immersion with one nicol.

430

Fig. 464 Bornite in symplectic intergrowth with chalcopyrite. Myrmekitic-like relics of bornite due to its replacement by chalcopyrite. b = bornite. c = chalcopyrite. ch = chalcocite veinlets replacing chalcopyrite. (occasionally ρ = pyrite is present). Matchless Mine, Friedenau Farm, Windhoek, South Africa. Oil-immersion with one nicol.

Fig. 465 Bornite symplectically intergrown with chalcopyrite, both transversed by later chalcocite veinlets. An idiomorphic pyrite is also shown, ρ = pyrite. b = bornite. c = chalcopyrite. ch = chalcocite. Matchless Mine, Friedenau Farm, Windhoek, South Africa. Oil-immersion with one nicol.

Fig. 466 Pyrite in symplectic intergrowth (replacement) with magnetite which in turn is replaced by haematite, m = magnetite, ρ = pyrite. h = haematite. Outukumpu, Finland. Oil-immersion with one nicol.

431

Fig. 467 Pyrite in symplectic intergrowth (replacement) with magnetite which in turn is replaced by haematite, m = magnetite, ρ = pyrite. h = haematite. Outukumpu, Finland. Oil-immersion with one nicol.

Fig. 468 Haematite in symplectic myrmekitic intergrowth with pyrite (actually pyrite replacing the initial magnetite), h = haematite, ρ = pyrite. ma = marcasite. Outukumpu, Finland. Oil-immersion with one nicol.

Fig. 469 Clausthalite replaced by later covellite. cl = clausthalite. co = covellite. Pinky Fault, Beaverlodge Lake, Saskatchewan, Canada. Oil-immersion with one nicol.

mm

432

Fig. 470 Clausthalite in symplectic (myrmekitic-like) intergrowth with covellite. The covellite replaces the clausthalite. cl = clausthalite. co = covellite. Pinky Fault, Beaverlodge Lake, Saskatchewan, Canada. Oil-immersion with one nicol.

/

Fig. 471 Myrmekitic-like symplectite of sphalerite and bravoite. s= sphalerite, b = bravoite. Mechernich, Eifel, Germany. Oil-immersion with one nicol.

Fig. 472 Myrmekitic-like symplectite of sphalerite and bravoite. s= sphalerite, b = bravoite. Mechernich, Eifel, Germany. Oil-immersion with one nicol.

433

Fig. 473 Myrmekitic-like symplectite of sphalerite and bravoite. s= sphalerite, b = bravoite. Mechernich, Eifel, Germany. Oil-immersion with one nicol.

Fig. 474 Myrmekitic-like symplectite of sphalerite and bravoite. s= sphalerite, b = bravoite. Mechernich, Eifel, Germany. Oil-immersion with one nicol.

Fig. 475 Zerfall texture. Breakdown texture of original mix-crystal into arsen (black) and white stibarsen. Initial mix-crystal probably preserved, m = ?initial mix-crystal, black = arsen. white = stibarsen. Allemont, Dauphine, France. Oil-immersion with one nicol.

434

Fig. 476 Zerfall texture. Breakdown texture of original mix-crystal into arsen (black) and white stibarsen. Initial mix-crystal probably preserved, m = ?initial mix-crystal, black = arsen. white = stibarsen. Allemont, Dauphine, France. Oil-immersion with one nicol.

Fig. 477 Zerfall texture. Breakdown texture of original mix-crystal into arsen (black) and white stibarsen. Initial mix-crystal probably preserved, m = ?initial mix-crystal, black = arsen. white = stibarsen. Allemont, Dauphine, France. Oil-immersion with one nicol.

Fig. 478 Breakdown symplectite texture. Allemontite Zerfall (breakdown) texture of initial mixcrystal into As (black) and stibarsen (white). Allemont, Dauphine, France. Oil-immersion with one nicol.

0.1 mm

435

Fig. 479 Breakdown symplectite texture. Allemontite Zerfall (breakdown) texture of initial mixcrystal into As (black) and stibarsen (white). Allemont, Dauphine, France. Oil-immersion with one nicol.

Fig. 480 Breakdown symplectite texture. Allemontite Zerfall (breakdown) texture of initial mixcrystal into As (black) and stibarsen (white). Allemont, Dauphine, France. Oil-immersion with one nicol.

Fig. 481 Zonal texture of allemontite ΠΙ from Pzribram, showing rhythmic distribution of AsSb-phase (black) and As-phase (white). Etched with K.S. In: Textures of the Ore Minerals and their Significance, by A. B. Edwards, 1960. χ 25.

436

Fig. 482 Cuprite with marginal malachite which replaces the cuprite and results in a myrmekitic-like symplectite of malachite and cuprite, m = malachite, c = cuprite. Viscachani, Vetas, Coro-Coro, Bolivia. Oil-immersion with one nicol.

Fig. 483 Chalcocite with cracks occupied by worm-like in appearance later cuprite replacing the chalcocite. The cuprite attains myrmekitic-like patterns along the cracks of the chalcocite. Also a cleavage pattern of chalcocite is exhibited, ch = chalcocite. c = cuprite. Batebawana, Algoma, Ontario, Canada. Oil-immersion with one nicol.

A-tfff, ;

Fig. 484 Complex relic pattern of magnetite (attaining symplectic form) in blastic pyrite, in cases exhibiting crystalline outlines, m = magnetite, ρ = pyrite. Otanmäki, Finland. Oil-immersion with one nicol.

437

Fig. 485 Cassiterite marginally replacing stannite. Myrmekitoidshaped relics of stannite are left in the marginal cassiterite which has replaced the stannite. s = stannite. c = cassiterite. m = myrmekitoid relics of stannite. Carrock Mine, Grainsgill, Caldbeck Fells, England. Oil-immersion with one nicol.

Fig. 4 8 6 Chromospinel with an alteration margin (due to the influence of basaltic melts) of magnetite. Symplectic silicates (myrmekite-like bodies) occur both in the magnetite margin as well as in the chromospinel in contact with the basalt (see Fig. 487). c = chromospinel. m = magnetite. my = myrmekitic-like silicate, b = basalt. Jato, Lekempti, W. Ethiopia. Oil-immersion with one nicol.

Fig. 487 Chromospinel with a magnetite margin in which symplectic silicates (myrmekitic-like bodies) occur. c = chromospinel. m = magnetite margin, my = myrmekitic-like intergrowth of silicate with the magnetite, b = basalt. Jato, Lekempti, W. Ethiopia. Oil-immersion with one nicol.

438

Fig. 488 Chromite with myrmekitic Serpentine bodies invading the chromite. c = chromite. s = serpentine, arrow shows extension of the serpentine into the chromite. Rodiani, N. Greece. Oil-immersion with one nicol.

Fig. 489 Brown serpentine forms a margin outside the chromite (b-s) with extensions of it following cracks of the adjacent chromite and attaining a myrmekitic intergrowth with the chromite. b-s = brown serpentine, ch = chromite. m-s = myrmekitic serpentine. Rodiani, N. Greece. Oil-immersion with one nicol.

Fig. 490 Chromite corroded and invaded by serpentine (attaining myrmekitic forms, chromiteserpentine symplectite). c = chromite. s = serpentine surrounding the chromite. m = myrmekitic-like serpentine in intergrowth with the chromite. Rodiani, N. Greece. Oil-immersion with one nicol.

439

Fig. 491 Chromite with serpentine in myrmekitic intergrowth (myrmekitic/symplectite). The myrmekitic serpentine follows oriented directions within the chromite host (see arrows), c = chromite. s = serpentine, arrows show serpentine invading the chromite and attaining myrmekitic forms. Rodiani, N. Greece. Oil-immersion with one nicol.

Fig. 492 Chromite with fracture and serpentine extending into the chromite and attaining myrmekitic-like character (intergrowth). c = chromite. cr = cracks of chromite invaded by serpentine, s = serpentine. Rodiani, N. Greece. Oil-immersion with one nicol.

0025

Fig. 493 Chromite in symplectic intergrowth with serpentine (attaining myrmekitic-like intergrowth). The serpentine extends from the margin of the chromite and is mainly restricted in the decoloration margin, c = chromite. s = serpentine outside the chromite with extension into the chromite. d = decoloration margin of the chromite. g = granular chromite. mc = massive chromite. m = myrmekitic serpentine in the massive chromite as extensions of the intergranular serpentine. Rodiani, N. Greece. Oil-immersion with one nicol.

440

Fig. 494 Cataclastic chromite with Serpentine filling the cataclastic cracks and acting as a Wiederverkittungs material. Also the serpentine extends into the chromite attaining myrmekitic symplectic forms. ch = chromite. w = Wiederverkittungs serpentine, m = myrmekitic serpentine, arrow "a" shows the extension of Wiederverkittungs serpentine as myrmekitic serpentine and arrow "b" shows external serpentine extending into the cataclastic cracks of chromite. Rodiani, N. Greece. Oil-immersion with one nicol.

Fig. 495 Sphalerite marginally invaded and replaced by silicates which result in a symplectic myrmekitic intergrowth. s = sphalerite. si = silicate in symplectic myrmekitic intergrowth with the sphalerite. Roan Antelope, Zambia. Oil-immersion with one nicol.

Fig. 496 Chalcocite in symplectic intergrowth with silicate (quartz). The chalcocite also corrodes and replaces the pyrite. cs = chalcocite in symplectic intergrowth with quartz, ρ = pyrite corroded and replaced by chalcocite. Debarau, Tigre, Erythrea. Oil-immersion with one nicol.

441

Fig. 497 Pyrite with a margin of neodigenite partly corroding the pyrite and with quartz extending into and replacing the pyrite. A symplectic quartz/pyrite intergrowth texture resulted due to infiltration and replacement of the pyrite by the silicate, ρ = pyrite. s = silicate (quartz) infiltration and replacement bodies, η = neodigenite. Tilva Mica, Bor, Serbia. Oil-immersion with one nicol.

Fig. 498 Quartz replacing pyrite along crystal penetrability directions (cleavage and cracks), q = quartz, ρ = pyrite. Ondonock, Wolaga, W. Ethiopia. Oil-immersion with one nicol.

Fig. 499 Pyrrhotite invaded and replaced by (quartz) silicate, resulting in myrmekitic symplectic "margins" and in the quartz following crystal penetrability directions of the pyrrhotite. ρ = pyrrhotite. m = myrmekitic quartz/pyrrhotite symplectite. q = quartz infiltrating into the pyrrhotite along crystallographic directions of the pyrrhotite. a = albandite. Kaiserstuhl, Germany. Oil-immersion with one nicol.

442

Fig. 500 Zoned cobaltite with gangue (quartz) invading the cobaltite interzonally. c = cobaltite. q = quartz invading the cobaltite interzonally. ρ = pitchblende. Joachimstal, Bohemia, Czech Republic. Oil-immersion with one nicol.

Fig. 501 Tourmaline needles transversing pyrite. ρ = pyrite. t = tourmaline needles. Pailaviri, Potosi, Bolivia. Oil-immersion with one nicol.

Fig. 502 Mica flakes oriented within the sphalerite (?representing mica schist relics) replaced by the sphalerite. Also chalcopyrite replaces the sphalerite, m = mica flakes, s = sphalerite, c = chalcopyrite. ps = pyrite spheroids (due to bacterial action). Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

443

Fig. 503 Symplectic intergrowth of idiomorphic hornblende and galena. In the section quartz, mica, andalusite or cordierite are also observed in intergrowth with galena, h = hornblende, g = galena. Boliden, N. Sweden. Oil-immersion with one nicol.

Fig. 504 Antimonite intergranular to silicates. s = silicates. a = antimonite. Bell Mine, Zimbabwe. Oil-immersion with one nicol.

Fig. 505 Antimonite intergranular to silicates. s = silicates. a = antimonite. Bell Mine, Zimbabwe. Oil-immersion with one nicol.

Fig. 506 Galena enclosing or partly enclosing quartz (idiomorphic). q = quartz, g = galena, b = bournonite. San Marcos Mine, Concepcion del Oro, Mexico. Oil-immersion with one nicol.

Fig. 507 Galena extending between the intergranular of idiomorphic quartzes, g = galena, q = quartz. San Marcos Mine, Concepcion del Oro, Mexico. Oil-immersion with one nicol.

Fig. 508 Galena and bournonite enclosing or partly enclosing quartz, b = bournonite. g = galena, q = quartz. San Marcos Mine, Concepcion del Oro, Mexico. Oil-immersion with one nicol.

445

Fig. 509 Quartz partly corroded and enclosed by gold, q = quartz. g = gold, arrow shows corroded margins of quartz. Nancy Mine, Boulder County, Colorado, USA. Oil-immersion with one nicol.

Fig. 5 1 0 Sylvanite enclosing quartz and marginally partly enclosed quartz, s = sylvanite. q = quartz. mq = marginal quartz partly enclosed by sylvanite. a = argentite. Nancy Mine, Boulder County, Colorado, USA. Oil-immersion with one nicol.

Fig. 511 Gangue (silicate) surrounded, partly corroded and invaded by sphalerite. Due to gangue replacement by the invading sphalerite, myrmekitic-like sphalerite/gangue symplectite results, s = sphalerite, g = gangue. m = sphalerite/silicate symplectic (myrmekitic in form), in cases in the outer zone of the silicate. Weisser Hirsch vein, Rammelsberg, Harz, Germany. Oil-immersion with one nicol.

446

Fig. 512 Gangue (silicate) surrounded, partly corroded and invaded by sphalerite. Due to gangue replacement by the invading sphalerite, myrmekitic-like sphalerite/gangue symplectite results, s = sphalerite, g = gangue. m = sphalerite/silicate symplectic (myrmekitic in form), in cases in the outer zone of the silicate. Weisser Hirsch vein, Rammelsberg, Harz, Germany. Oil-immersion with one nicol.

Fig. 513 Idiomorphic quartzes often exhibiting replacement by the later sphalerite which encloses them, q = quartz, s = sphalerite. Weisser Hirsch vein, Goslar, Harz, Germany. Oil-immersion with one nicol.

Fig. 514 Tetrahedrite enclosing idiomorphic quartz, t = tetrahedrite. q = quartz, s = sphalerite, ρ = pyrite. San Marcos Mine, Concepcion del Oro, Mexico. Oil-immersion with one nicol.

447

Fig. 515 Skeletal pyrite (occupying intercrystalline spaces) and enclosing gangue is replaced by argentite, which forms a net in the pyrite. ρ = pyrite. a = argentite? (Pyrargyrite is also present) e = pyrite extending between gangue crystals, g = gangue. Freiberg, Saxony, Germany. Oil-immersion with one nicol.

Fig. 516 Garnet replaced by sphalerite, g = garnet. s = sphalerite (arrows show sphalerite replacing the garnet), ρ = pyrite. Avalos, Conception del Oro, Mexico. Oil-immersion with one nicol.

Fig. 517 Garnet transversed and replaced by veinlet of sphalerite and partly pyrite. s = sphalerite surrounding the garnet, sv = veinlets of sphalerite, ρ = pyrite, associated with sphalerite, g = garnet. Avalos, Conception del Oro, Mexico. Oil-immersion with one nicol.

448

Fig. 518 Garnet crystalloblast enclosing corroded chromite crystals and chalcopyrite. Chalcopyrite also surrounds the garnet crystalloblast which also exhibits crystalline outlines, c = chromite. ch = chalcopyrite. g = garnet. Outukumpu, Finland. Oil-immersion with one nicol.

Fig. 519 Tetrahedrite with a margin of chalcopyrite transversed by elongated gudmundite crystalloblasts. t = tetrahedrite. c = chalcopyrite. g = gudmundite (elongated) crystalloblasts. Seinajuki, Finland. Oil-immersion with one nicol.

Fig. 520 Tetrahedrite with elongated gudmundite crystalloblasts transversing the tetrahedrite and with marginal chalcopyrite. g = gudmundite. t = tetrahedrite. c = chalcopyrite. Seinajuki, Finland. Oil-immersion with one nicol.

449

Fig. 521 Zoned smaltite transversed by safflorite crystalloblasts. s = smaltite. sa = safflorite crystalloblasts transversing the zoned smaltite. Morocco. Oil-immersion with one nicol.

Fig. 522 Star-shaped safflorite crystalloblasts with loellingite. s = safflorite. 1 = loellingite. South Terras, St. Austell, Cornwall, S. W. England. Oil-immersion with one nicol.

Fig. 523 Rammelsbergite with crystalloblastic safflorite stars, partly or fully developed, r = rammelsbergite. s = safflorite stars. Richelsdorfer Gebirge, Grube, Munich, Germany. Oil-immersion with one nicol.

450

Fig. 524 Rammelsbergite with crystalloblastic safflorite stars, partly or fully developed, r = rammelsbergite. s = safflorite stars. Richelsdorfer Gebirge, Grube, Munich, Germany. Oil-immersion with one nicol.

Fig. 525 Wolframite idioblast in a mass of marcasite. w = wolframite crystalloblast. m = marcasite. Carroch Mine, Grainsgill, Goldbeck Falls, England. Oil-immersion with one nicol.

Fig. 526 Wolframite idioblast in a mass of lamellar marcasite. w = wolframite, m = marcasite (lamellar). Carroch Mine, Grainsgill, Goldbeck Falls, England. Oil-immersion with one nicol.

451

Fig. 527 Wolframite idioblasts associated with scheelite in a mass of marcasite. w = wolframite, s = scheelite. m = marcasite. Carroch Mine, Grainsgill, Goldbeck Falls, England. Oil-immersion with one nicol.

Fig. 528 Arsenopyrite crystalloblasts (some exhibiting idiomorphism) enclosing, corroding and replacing sphalerite. a = arsenopyrite crystalloblasts. s = sphalerite. Ramsbeck, Westphalia, Germany. Oil-immersion with one nicol.

Fig. 529 Idioblastic arsenopyrite transversing a "band" of sphalerite, a = arsenopyrite crystalloblast. s = "band" of sphalerite transversed by arsenopyrite crystalloblasts. Ramsbeck, Westphalia, Germany. Oil-immersion with one nicol.

452

Fig. 530 Radiating arsenopyrite crystalloblasts enclosing and replacing sphalerite. a = radiating arsenopyrite crystalloblasts. s = sphalerite enclosed in the arsenopyrite. Ramsbeck, Westphalia, Germany. Oil-immersion with one nicol.

Fig. 531 Radiating arsenopyrite crystalloblasts enclosing and replacing sphalerite. a = radiating arsenopyrite crystalloblasts. s = sphalerite enclosed in the arsenopyrite. Ramsbeck, Westphalia, Germany. Oil-immersion with one nicol.

Fig. 532 Feather-like pattern of arsenopyrite crystalloblasts. a = arsenopyrite. g = gangue. Ramsbeck, Westphalia, Germany. Oil-immersion with one nicol.

453

Fig. 533 Haematite idioblast associated with native selenium and other selenium minerals, h = haematite, s = native selenium. Pinky Fault, Beaverlodge Lake, Saskatchewan, Canada. Oil-immersion with one nicol.

0025

Fig. 534 Clausthalite with crystalloblastic haematite, c = clausthalite. h = crystalloblastic haematite. Tilkerode, Harz, Germany. Oil-immersion with one nicol.

Fig. 535 Interpenetrating haematite crystalloblasts enclosing clausthalite. h = haematite crystalloblasts. c = clausthalite enclosed in haematite crystalloblasts. cl = clausthalite in which the haematite crystalloblasts have grown. Tilkerode, Harz, Germany. Oil-immersion with one nicol.

454

Fig. 5 3 6 Haematite crystalloblast transversing the contact clausthalite/calcite. h = haematite crystalloblast. ca = calcite. c = clausthalite. Tilkerode, Harz, Germany. Oil-immersion with one nicol.

Fig. 537 Prismatic haematite (velonoblasts) in enargite. The haematite crystalloblast includes relics of enargite. h = haematite velonoblast (needleshaped). e = enargite inclusions in the velonoblast. en = enargite. Tarvis Township, Algoma District, Ontario, Canada. Oil-immersion with one nicol.

Fig. 538 Haematite crystalloblast with prismatic protuberances of haematite velonoblasts. he = haematitic crystalloblast. h = haematite velonoblast. Tarvis Township, Algoma District, Ontario, Canada. Oil-immersion with one nicol.

455

Fig. 539 Ilmenite crystalloblast enclosing gangue. il = ilmenite crystalloblast. g = gangue. Murgang near Casaccia, Lukmanier, Switzerland. Oil-immersion with one nicol.

Fig. 540 Niccolite with pyrite crystalloblasts. η = niccolite. ρ = pyrite crystalloblast. Chang Poy, Dachang, China. Oil-immersion with one nicol.

Fig. 541 Initial colloform pyrite (now crystalline) transecting marcasite. ρ = pyrite. m = marcasite. Arrows show initial colloform structure of pyrite. Afrikander Mine, Dominion Reef, South Africa. Oil-immersion with one nicol.

456

Fig. 542 Initial colloform iron oxides changed due to induced reduction into magnetite. However, the initial colloform pattern of the iron hydroxides is preserved, replaced by later pyrite. ρ = pyrite. m = magnetite. Peko Mine, N. Territory, Australia. Oil-immersion with one nicol.

Fig. 543 Blastic pyrite enclosing relics of sphalerite and with pyrite extensions in the sphalerite enclosing sphalerite. ρ = pyrite crystalloblast. rs = relics of sphalerite, s = sphalerite. Ruhrgebiet, Germany. Oil-immersion with one nicol.

Fig. 544 Blastic pyrite enclosing fragments of sphalerite (often corroded) and silicates (crystalline outline often exhibited), ρ = pyrite. s = sphalerite (corroded). si = silicates exhibiting crystalline outline. Ruhrgebiet, Germany. Oil-immersion with one nicol.

Fig. 545 Granular pyrrhotite with pyrite crystalloblast including pyrrhotite relics, py = pyrrhotite (granular), ρ = pyrite crystalloblast. Talvey Metal Mine, Township 157, Algoma District, Ontario, Canada. Oil-immersion with one nicol.

Fig. 546 Pyrite crystalloblast enclosing chalcocite. ρ = pyrite crystalloblast. c = chalcocite. Tilva Mica, Bor, Serbia. Oil-immersion with one nicol.

Fig. 547 Blastic pyrite bending chalcopyrite. ρ = pyrite crystalloblast. c = chalcopyrite. g = galena. Bayerland, Waldsassen, Germany. Oil-immersion with one nicol.

458

Fig. 548 Bornite with chalcocite transversed by later crystalloblastic pyrite (often idioblast). b = bornite. c = chalcocite. py = pyrite. Tsumeb, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 549 Wedge-shaped pyrite crystalloblast transversing pattern of bornite replaced by stromeyerite. ρ = pyrite. b = bornite. s = stromeyerite. Magma Mine, Superior, Arizona, USA. Oil-immersion with one nicol.

Fig. 550 Pyrite crystalloblast transecting bornite replaced by stromeyerite. Also, galena is enclosed in the pyrite crystalloblast. ρ = pyrite crystalloblast. b = bornite. s = stromeyerite. g = galena. Magma Mine, Superior, Arizona, USA. Oil-immersion with one nicol.

459

Fig. 551 Idioblastic pyrite enclosing relics of corroded and replaced magnetite. ρ = pyrite idioblast. m = magnetite relics in pyrite. Peko Mine, N. Territory, Australia. Oil-immersion with one nicol.

Fig. 552 Blastic pyrite enclosing relics of corroded and replaced sphalerite. ρ = pyrite crystalloblast. r = relics of sphalerite, s = sphalerite. Chang Poy, Dachang, China. Oil-immersion with one nicol.

Fig. 553 Pyrite idioblast including zonally corroded relics of sphalerite, ρ = pyrite. s = sphalerite. Chang Poy, Dachang, China. Oil-immersion with one nicol.

0.1 mm

460

Fig. 554 Marginal area of pyrite crystalloblast including relics (corroded sphalerite in cases associated with silicates), ρ = pyrite. r = relics of sphalerite. Chang Poy, Dachang, China. Oil-immersion with one nicol.

Fig. 555 Pyrite crystalloblast with relics of sphalerite irregularly distributed within the pyrite crystal, ρ = pyrite. r = relics of sphalerite. Chang Poy, Dachang, China. Oil-immersion with one nicol.

Fig. 556 Idioblastic pyrite including relics of sphalerite and also pigment size relics of sphalerite are restricted in a marginal part of the pyrite. ρ = pyrite. r = relics of sphalerite (relics left due to the replacement of sphalerite by the blastic pyrite). pi = pigment relics of sphalerite in a marginal area of the pyrite crystalloblast. Chang Poy, Dachang, China. Oil-immersion with one nicol.

0.1mm

461

H -sv***

Ρ

Fig. 557 Zonal pattern of galena as indicated by cleavage (shown on polishing), g = galena, c = cleavage. Laurium, Greece. Oil-immersion with one nicol.

^

Fig. 558 Zonal pattern of galena as indicated by cleavage (shown on polishing). Pyrite is also present, g = galena, c = cleavage, ρ = pyrite. Laurium, Greece. Oil-immersion with one nicol.

Fig. 559 Zoned cobaltite. Cobalt City, Ontario, Canada. Oil-immersion with one nicol.

462

Fig. 560 Zoned and twinned smaltite. t = twin plane. Bou Azzer, Morocco, Africa. Oil-immersion with one nicol.

Fig. 561 Zoned smaltite. Bou Azzer, Morocco, Africa. Oil-immersion with one nicol.

Fig. 562 Zoned safflorite, most probably derived from gels, s = zoned safflorite. Eldorado Mine, Great Bear Lake, Saskatchewan, Canada. Oil-immersion with one nicol.

463

Fig. 563 Gel-derived zoned safflorite pattern. s = safflorite. Eldorado Mine, Great Bear Lake, Saskatchewan, Canada. Oil-immersion with one nicol.

Fig. 564 Zonal safflorite interbanded with gangue, most probably derived from colloidal solutions. Cobalt City, Ontario, Canada. Oil-immersion with one nicol.

Fig. 565 Zonal safflorite interbanded with gangue, most probably derived from colloidal solutions. Cobalt City, Ontario, Canada. Oil-immersion with one nicol.

464

Fig. 566 Zoned malacon. Whereas the outer zoning is due to growth, some of the internal colloform-like zones might be due to water intake of the malacon (metamictic) due to the change of zircon to malacon. ζ = normal zoning, cz = colloform zoning. Faraday Mine, Bancroft, Ontario, Canada. Oil-immersion with one nicol.

Fig. 567a Malacon exhibiting zonal growth and change in form and habitus. ζ = zoned malacon. m = malacon forming a marginal growth of magnetites, ma = magnetite. Faraday Mine, Bancroft, Ontario, Canada. Oil-immersion with one nicol.

Fig. 567b Bravoite exhibiting zoning and changing of habitus. Photograph from RamdohrEhrenberg. Mechernich, Eifel, Germany, χ 200.

465

Fig. 568 Zoned tetrahedrite, the central crystal exhibits zonal growth, whereas the outer zones of growth exhibit differences in habitus as compared with the central zoned crystal. Djebel Ouenza, Constantine, Algeria. Oil-immersion with one nicol.

Fig. 569 Zoned cuprite crystals with interzonal malachite in cases exhibiting sphaeroidal forms in the cuprite, c = cuprite. iz = interzonal malachite, s = sphaeroidal malachite. Livadi, Seriphos, Greece. Oil-immersion with one nicol.

Fig. 570 Cuprite crystals with interzonal malachite gel-structures, c = cuprite, m = malachite. Livadi, Seriphos, Greece. Oil-immersion with one nicol.

466

Fig. 571 Cuprite crystal enclosed by a margin of malachite which in turn is enclosed by a larger cuprite, c = cuprite crystal, m = malachite margin, e = cuprite crystal enclosing the idiomorphic cuprite with the malachite margin. In cases the malachite exhibits colloform structure (see arrows). Cerro de Pasco, Peru. Oil-immersion with one nicol.

Fig. 572 Zoned cassiterite idioblast. ζ = zonal growth of cassiterite. Shizhu Yuan near Chenzhou, Hunan Province, China. Oil-immersion with one nicol.

X

Fig. 573 Idiomorphic zoned cassiterite with variations in the pleochroism of the zones. ζ = zoned cassiterite. Shizhu Yuan near Chenzhou, Hunan Province, China. Oil-immersion with one nicol.

467

Fig. 574 Zoned cassiterite exhibiting alternating change in orientation of the zoned cassiterite. ζ = zones of cassiterite. Shizhu Yuan near Chenzhou, Hunan Province, China. Oil-immersion with one nicol.

Fig. 575 Pyrite exhibiting zonal structure with changes in form and habitus. Peko Mine, N. Territory, Australia. Oil-immersion with one nicol.

Fig. 576 Metacinnabar (elongated prismatic crystals - derived from gels) exhibiting zonal growth, m = metacinnabar. z= zonal growth of metacinnabar. Mt. Diablo, Contra Costa County, California, USA. Oil-immersion with one nicol.

468

Fig. 577 Chromite idiomorphic with a margin of epitactic magnetite, c = chromite. m = magnetite (martitized). Yubdo, W. Ethiopia. Oil-immersion with one nicol. χ 800.

Fig. 578 Corroded and rounded chromite with an overgrowth of epitactic blastic magnetite attaining idiomorphism in lateric iron, nickel and chromite occurrences and partly transecting iron pisolite-oolite, c = chromite. m = magnetite, ρ = iron pisolite. Kozani, Greece. Oil-immersion with one nicol.

Fig. 579 Multiple epitaxis. Magnetite epitactic on chromite and sulfide (pyrite?) on magnetite, c = chromite. m = magnetite, ρ = sulfide (pyrite). Transvaal, South Africa. Oil-immersion with one nicol.

469

Fig. 580 Epitaxis of uraninite on columbite, according to Strunz. Hagendorf, Bavaria, Germany.

Fig. 581 Epitactic uraninite on columbite. u = uraninite. c = columbite. Hagendorf, Bavaria, Germany. Oil-immersion with one nicol.

Fig. 582 Uraninite epitactic on columbite. A radioactive halo surrounds the uraninite (sulfide also occurs within the halo), u = uraninite. c = columbite. r = radioactive halo, ρ = sulfide. Hagendorf, Bavaria, Germany. Oil-immersion with one nicol.

470

Fig. 583 Columbite enclosing uraninite and partly engulfing marginally a larger epitactic uraninite which exhibits developed crystal faces and is surrounded by a radioactive halo margined by sulfides, c = columbite. u = uraninite. r = radioactive halo, s = sulfide margin to the radioactive halo. Hagendorf, Bavaria, Germany. Oil-immersion with one nicol.

Fig. 584 Well-formed crystal of uraninite epitactic on columbite in which case a great part of the crystalline uraninite is partly engulfed by the columbite. Pyrite develops on the radioactive dispersion halo of the uraninite in the feldspars of the pegmatite. u = crystalline uraninite. c = columbite. ρ = pyrite. Hagendorf, Bavaria, Germany. Oil-immersion with one nicol.

Fig. 585 Columbite enclosing uraninite exhibiting partly developed crystal faces and having another uraninite crystal epitactic on the columbite. u = uraninite. eu = epitactic uraninite. c = columbite. h = radioactive halo. Hagendorf, Bavaria, Germany. Oil-immersion with one nicol.

471

Fig. 586 Epitactic uraninite on columbite with extensions of the uraninite in the columbite. u = uraninite. c = columbite. e = extensions of the uraninite in the columbite. h = radioactive halo surrounding the uraninite. Hagendorf, Bavaria, Germany. Oil-immersion with one nicol.

Fig. 587 Epitactic ilmenite on magnetite. (Ilmenite can also be seen as extension or veinform in the magnetite, see Fig. 292). m = magnetite, il = ilmenite. Mooihoek, Lydenburg District, Transvaal, South Africa. Oil-immersion with one nicol.

Fig. 588 Titanomagnetite (with ilmenite and spinel) with epitactic spinel which also extends into the magnetite. m = titanomagnetite. es = epitactic spinel. e = extension of spinel in the magnetite. c = chalcopyrite. Concordia Mine near O'okiep, South Africa. Oil-immersion with one nicol.

472

Fig. 589 Titanomagnetite with epitactic spinel and with spinel following oriented directions within the titanomagnetite. The epitactic spinel is intergrown with chalcopyrite. m = magnetite, es = epitactic spinel, s = spinel, often oriented within the chalcopyrite. il = oriented lamellae of ilmenite in magnetite. c = chalcopyrite. Concordia Mine near O'okiep, South Africa. Oil-immersion with one nicol.

Fig. 590 Titanomagnetite with spinels and with marginal ilmenite ?epitactic which in turn has a margin of epitactic spinel, m = titanomagnetite. il = ilmenite ?epitactic. sr = spinel following the reaction margin titanomagnetite with epitactic ilmenite. es = epitactic spinel, c = chalcopyrite. Concordia Mine near O'okiep, South Africa. Oil-immersion with one nicol.

Fig. 591 Jacobsite epitactic on braunite. The epitactic jacobsite exhibits a developed crystal face, b = braunite. j = jacobsite epitactic, arrow shows developed crystal face of jacobsite. Otjosondu, Sandfeld, Namibia, S . W . Africa. Oil-immersion with one nicol.

473

Fig. 592 Braunite with epitactic jacobsite in cases including a relic of the braunite in the epitactic jacobsite. The epitactic jacobsite extends also into the braunite. b = braunite. j = jacobsite, arrow shows relic of braunite in epitactic jacobsite. Otjosondu, Sandfeld, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 593 Jacobsite epitactic on braunite. j = epitactic margins of jacobsite often with extensions into the braunite. b = braunite. Otjosondu, Sandfeld, Namibia, S.W. Africa. Oil-immersion with one nicol.

Fig. 594 Jacobsite in intergrowth with the silicates and as a margin epitactically grown on the braunite. Also extensions of the jacobsite in the braunite are to be seen, b = braunite. j = jacobsite, arrow shows epitactic margin of jacobsite on braunite. Otjosondu, Sandfeld, Namibia, S. W. Africa. Oil-immersion with one nicol.

474

Fig. 595 Pyrrhotite epitactic on sphalerite. In cases the sphalerite is corroded and partly replaced by the epitactic pyrrhotite. ρ = pyrrhotite epitactic on sphalerite, s = sphalerite. sc = sphalerite corroded and included in epitactic pyrrhotite. Otjosondu, Sandfeld, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 596 Epitactic pyrrhotite on sphalerite. The epitactic pyrrhotite has corroded and replaced the sphalerite marginally; pyrrhotite is also included in the sphalerite, ρ = pyrrhotite. s = sphalerite. Otjosondu, Sandfeld, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 597 Arsenopyrite replacing and being epitactic on sphalerite, a = arsenopyrite (epitactic). s = sphalerite. San Martin Mine, Sombrerete, Zacatecas, Mexico. Oil-immersion with one nicol.

475

Fig. 598 Arsenopyrite replacing and being epitactic on sphalerite, a = arsenopyrite (epitactic). s = sphalerite. San Martin Mine, Sombrerete, Zacatecas, Mexico. Oil-immersion with one nicol.

Fig. 599 Pyrite replacing and being epitactic on magnetite, m = magnetite. ρ = pyrite, arrow shows epitactic pyrite. s = silicates. Seriphos Island, Greece. Oil-immersion with one nicol.

Fig. 600 Pyrite epitactic with welldeveloped crystal outlines corroding and replacing sphalerite, ρ = pyrite. s = sphalerite. Chang Poy, Dachang, China. Oil-immersion with one nicol.

476

Fig. 601 Epitactic bournonite on tetrahedrite (the bournonite also replaces the tetrahedrite). t = tetrahedrite. be = bournonite (epitactic). b = bournonite replacing the tetrahedrite. q = quartz idiomorphic is included in the tetrahedrite/bournonite replacement symplectite. San Marcos Mine, Conception del Oro, Mexico. Oil-immersion with one nicol.

Fig. 602 Bornite with well-developed crystalline faces epitactic on a symplectite of rutile and silicate. The bornite is also symplectic with the rutile. b = epitactic bornite, arrow shows symplectic bornite with rutile. s = silicate symplectic with rutile. r = rutile. Spectacel Mine near O'okiep, Namibia, S. W. Africa. Oil-immersion with one nicol.

0 05

Fig. 603 Pyrrhotite with an epitactic margin of iron albandite. a = albandite. ρ = pyrrhotite. Kaiserstuhl, Germany. Oil-immersion with one nicol.

477

Fig. 604 Pyrrhotite enclosing albandite. ρ = pyrrhotite. a = albandite. Kaiserstuhl, Germany. Oil-immersion with one nicol.

-.

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•1S. -S Ι & ' ^ ί έ λ Fig. 605 Pyrite surrounding gersdorffite. The pyrite extends into the sphalerite and colloform patterns of pyrite are exhibited. ρ = pyrite. g = gersdorffite. pc = colloform pyrite. s = sphalerite. Fastrada Vein, Ramsbeck, Westphalia, Germany. Oil-immersion with one nicol.

Fig. 606 Chromite including olivine centrally; zonal distribution of corroded olivine relics is also shown, ο = olivine (centrally), ζ = zonal distribution of olivines, c = chromite. Transvaal, South Africa. Oil-immersion with one nicol.

478

Fig. 607 Zonal distribution of corroded olivine relics in idiomorphic chromite. r = corroded silicate relics zonally distributed in chromite. c = idiomorphic chromite. Transvaal, South Africa. Oil-immersion with one nicol.

Fig. 608 Chromite engulfing olivine marginally (see arrow) and also including corroded olivine relics, c = chromite. r = corroded olivine relics. Transvaal, South Africa. Oil-immersion with one nicol.

Fig. 609 Rutile (partly rounded) included in chromite which in turn exhibits spinel and rutile ("exsolutions"), r = rutile. c = chromite. s = spinel ("ex-solutions"), re = rutile ("ex-solutions"). Rodiani, Greece. Oil-immersion with one nicol.

0.025

479

Fig. 610 Platin (Fe-Au-Pt) nugget including osmiridium crystals indicating a zonal distribution within the nugget. Also corroded chromite is included in the nugget. A Mn-Fe oxides margin of the nugget is also exhibited. m = Mn-Fe oxides margin, ρ = ferroplatin. ο = osmiridium. c = corroded chromite. Yubdo, W. Ethiopia. Oil-immersion with one nicol.

Fig. 611 Platin PGM marginally included in a ferroplatin nugget. (For detailed description of the Pt group minerals see Figs. 735-739).

Fig. 612 Uraninite partly exhibiting well-developed crystalline faces included in columbite. u = uraninite. c = columbite. Hagendorf, Bavaria, Germany. Oil-immersion with one nicol.

480

Fig. 613 Uraninite with some welldeveloped faces included in columbite. u = uraninite. c = columbite. Hagendorf, Bavaria, Germany. Oil-immersion with one nicol.

Fig. 614 Uraninite partly exhibiting developed crystalline outline included in columbite. Also epitactic uraninite on columbite is shown, u = uraninite (partly exhibiting welldeveloped crystal faces) included in the columbite. c = columbite. e = epitactic uraninite on columbite exhibiting well-developed crystal faces, and also a radioactive halo (h). Hagendorf, Bavaria, Germany. Oil-immersion with one nicol.

Fig. 615 Hawleyite exhibiting cubic outline and strong lemon-yellow internal reflection, marginal to sphalerite/wurtzite. h = hawleyite. s = sphalerite/wurtzite. c = carbonate gangue. Ragada, Rhodope, N. Greece. Oil-immersion with one nicol.

481

Fig. 616 Hawleyite with strong lemon-yellow internal reflection, included in sphalerite/wurtzite. s = wurtzite. h = hawleyite. Ragada, Rhodope, N. Greece. Oil-immersion with one nicol.

Fig. 617 Zoned cobaltite partly enclosed and "surrounded" by safflorite. c = cobaltite zone, s = safflorite. South Terras Mine, St. Austell, Cornwall, England. Oil-immersion with one nicol.

Fig. 618 Colloform pitchblende enclosed in coffinite (well-developed crystals exhibited), ρ = pitchblende, c = coffinite. Pribram, Czech Republic. Oil-immersion with one nicol.

482

Fig. 619 Magnetite with a margin of malacon and with a developed epitactic malacon. m = magnetite. ma = margin of malacon. e = epitactic developed malacon. r = rutile. Faraday Mine, Bancroft, Ontario, Canada. Oil-immersion with one nicol.

Fig. 620 Magnetite with a margin of malacon (actually enclosing the magnetite). Also zoned malacon. m = magnetite, ζ = malacon margin, zm = zoned malacon. Faraday Mine, Bancroft, Ontario, Canada. Oil-immersion with one nicol.

Fig. 621 Zoned malacon enclosing magnetite. The zoned malacon has extensions following cracks of the magnetite which surrounds the zoned malacon. zm = zoned malacon. m = magnetite enclosed in malacon. e = extensions of malacon in the outer magnetite. mo = outer magnetite surrounding the malacon. Faraday Mine, Bancroft, Ontario, Canada. Oil-immersion with one nicol.

483

Fig. 622 Zoned malacon in uraninite. Cracks radiate from the malacon in the uraninite. ζ = malacon. u = uraninite. m = magnetite. Faraday Mine, Bancroft, Ontario, Canada. Oil-immersion with one nicol.

Fig. 623 Uraninite enclosing gold. Gold is also included in the gangue. u = uraninite. g = gold. Shinkolobwe, Katanga, Zaire. Oil-immersion with one nicol.

Fig. 624 "Kidneyforms" of cryptomelane (K(Mn 2+ , Mn 4+ ) 8 0 16 ). Low temperature hydrothermal formations. Kisenge Kamata, Zaire. Oil-immersion with one nicol. About natural size.

484

Fig. 625 Small sphaeroids of colloform pitchblende in sandstone, ρ = pitchblende, q = quartz of the sandstone. Val Rendena, Trentino, Italy. Oil-immersion with one nicol.

Fig. 626 Gel pitchblende with a gel layer lighter in colour, perhaps due to variations in composition. 1 = lighter layer of pitchblende, d = dark outer zone due to ?alteration. r = radiogenic lead. Ransom Property, Montreal River, east of Lake Superior, Canada. Oil-immersion with one nicol.

Fig. 627 Fine gel layers of pitchblende of complex appearance, grading to broad smooth running outer zones. Look-out Claim, Marshall Pass, Saguachelo, Colorado, USA. Oil-immersion with one nicol.

485

Fig. 628 Complex colloform structures of pitchblende which result from rhythmical precipitation of layers under coagulation conditions of uo2. Look-out Claim, Marshall Pass, Saguachelo, Colorado, USA. Oil-immersion with one nicol.

Fig. 629 Broad layers of colloform pitchblende with finer, more complex gel structures between broad zones of pitchblende. Look-out Claim, Marshall Pass, Saguachelo, Colorado, USA. Oil-immersion with one nicol.

0025

Fig. 630 Gel pitchblende with colloform rhythmical zoning and with cracks and margins occupied by younger pitchblende (lighter in color) probably due to solution and redeposition. Look-out Claim, Marshall Pass, Saguachelo, Colorado, USA. Oil-immersion with one nicol.

486

Fig. 631 A fragment of colloform pitchblende is surrounded by younger pitchblende similarly colloform. An "unconformity" in the directions of the rhythmical zones of coagulation marks the two generations of precipitation. Look-out Claim, Marshall Pass, Saguachelo, Colorado, USA. Oil-immersion with one nicol.

Fig. 632 Colloform pitchblende with central parts and layers removed by solution (differential solution of pitchblende layers). Dark-grey gangue. Gryssgruvan, Dalarna, Sweden. Oil-immersion with one nicol.

Fig. 633 Pitchblende with fine rhythmical structures, indicating a difference in reflectivity o f the various shell layers, is dependent on porosity. The structures represent relics of former gel structure (the pitchblende now being crystalline). The white flecks (filling small cracks) are radiogenic lead. Joachimsthal, Bohemia, Czech Republic. Oil-immersion with one nicol.

487

Fig. 634 A rhythmical structure due to alternating zones of gangue and U0 2 . This structure probably represents early conditions of simultaneous precipitation of U 0 2 and gangue material. Johanngeorgenstadt, Saxony, Germany. Oil-immersion with one nicol. χ 1000.

Fig. 635 Uraninite with colloform relic structure. Johanngeorgenstadt, Saxony, Germany. Oil-immersion with one nicol.

I ύχ&ζΧΑ '{tit·· 4*'·. '.'λ λ ·* '«'£',

488

Fig. 636 Colloform relic patterns in manganite. cr = colloform relic patterns, m = manganite. Ilfeld, Harz, Germany. Oil-immersion with one nicol.

Fig. 637 Colloform relic patterns in manganite. Ilfeld, Harz, Germany. Oil-immersion with one nicol.

V

iL' ·. ν ' *

>

''

ι

" . * · Μ

Fig. 638 Niccolite with gel-relic structure preserved, η = niccolite. r = relic colloform texture in the niccolite. Azegour, S. Morocco. Oil-immersion with one nicol.

Fig. 639 Gel marcasite and sphalerite. m = marcasite. cm = crystalline marcasite. s = sphalerite. Limni, Cyprus. Oil-immersion with one nicol.

489

Fig. 640 Colloform sphalerite with marcasite. s = sphalerite, m = marcasite. Limni, Cyprus. Oil-immersion with one nicol.

Fig. 641 Banded gel-structure of colloform sphalerite and colloform gratonite. s = colloform sphalerite, g = colloform gratonite. Wiesloch, Baden, Germany. Oil-immersion with one nicol.

0 05 mm

Fig. 642 Interbanded colloform anglesite and galena, a = anglesite. g = galena. Lemitar Mountains, Soccoro Co., New Mexico, USA. Oil-immersion with one nicol.

490

Fig. 643 Crystalline cuprite with interzonal colloform malachite, c = cuprite. m = interzonal malachite. Livadi, Seriphos, Greece. Oil-immersion with one nicol.

Fig. 644 Colloform hutchinsonite including gratonite crystals. Wiesloch, Königszug, Baden, Germany. Oil-immersion with one nicol.

Fig. 645 Colloform hutchinsonite including gratonite crystals, h = hutchinsonite. g = gratonite. Wiesloch, Königszug, Baden, Germany. Oil-immersion with crossed nicols.

491

Fig. 646 Colloform plattnerite with galena. ρ = plattnerite. g - galena. Mapimi, Durago, Mexico. Oil-immersion with one nicol.

Fig. 647 Colloform structures of plattnerite and galena, ρ = plattnerite. Mapimi, Durago, Mexico. Oil-immersion with one nicol.

Fig. 648 Rhythmically banded colloform tenorite with chrysocolla and malachite. t = tenorite rhythmically banded with chrysocolla. Torreon, Cohukla, Mexico. Oil-immersion with one nicol.

492

Fig. 649 Rhythmically banded colloform tenorite with chrysocolla and malachite. t = tenorite rhythmically banded with chrysocolla. c = chrysocolla. Torreon, Cohukla, Mexico. Oil-immersion with one nicol.

Fig. 650 Rhythmically banded colloform tenorite with chrysocolla and malachite. t = tenorite rhythmically banded with chrysocolla. c = chrysocolla. Torreon, Cohukla, Mexico. Oil-immersion with one nicol.

Fig. 651 Rhythmically banded colloform tenorite with chrysocolla and malachite. t = tenorite rhythmically banded with chrysocolla. c = chrysocolla. Torreon, Cohukla, Mexico. Oil-immersion with one nicol.

493

Fig. 652 Colloform tenorite (with chrysocolla and malachite). Fine outer rhythmical band of tenorite showing complex indentation pattern (see arrows). Torreon, Cohukla, Mexico. Oil-immersion with one nicol.

Fig. 653 Lievrite and brown-iron colloform interbanded structure. 1 = lievrite. b = brown iron. Mavro Pounti, Seriphos, Greece. Oil-immersion with one nicol.

Fig. 654 Colloform structure of interbanded haematite and lievrite. 1 = lievrite. h = haematite. r = haematite/lievrite replacement patterns also shown. Mavro Pounti, Seriphos, Greece. Oil-immersion with one nicol.

494

Fig. 655 Gel sphaeroids of polianite with dark bands of an unidentified manganese mineral, ρ = polianite. d = dark unidentified manganese mineral. Goa, India. Oil-immersion with one nicol.

Fig. 656 Polianite with a band consisting of fractured dark manganese minerals cemented by polianite. The bands appear to be part of the colloform banding of the manganese minerals, ρ = polianite. d = dark manganese mineral, fractured and held together by polianite forming an inter-gel band within the polianite. Goa, India. Oil-immersion with one nicol.

Fig. 657 Gel-sphaeroid bands of polianite interbanded with a dark unidentified manganese mineral. The polianite sphaeroid is partly intergrown and surrounded by cryptomelane. ρ = polianite. cy = cryptomelane. Goa, India. Oil-immersion with one nicol.

495

Fig. 658 Polianite initial gelsphaeroid exhibiting atollreplacement by psilomelane and surrounded by psilomelane in which scattered dark patches of an unidentified manganese mineral occur. ρ = polianite. cy = cryptomelane. ip = atoll-type replacement of polianite by cryptomelane. Goa, India. Oil-immersion with one nicol.

Fig. 659 Polianite with a gel-band of psilomelane scattered. Dark unidentified manganese minerals are also present, ρ = polianite. cy = cryptomelane. Goa, India. Oil-immersion with one nicol.

Fig. 660 Colloform pyrite (with chalcopyrite colloform bands) and with uraninite also of colloform derivation, ρ = colloform pyrite. c = chalcopyrite. u = uraninite. Grampound Road Mine, St. Stephens, Cornwall, S. W. England. Oil-immersion with one nicol.

496

Fig. 661 Colloform limonite as a replacement structure of ilmenite with oriented haematite bodies. 1 = interbanded limonite colloform replacement structure, r = ilmenite with haematite bodies. Harrar, Ethiopia. Oil-immersion with one nicol.

Fig. 662 Colloform limonite as a replacement structure of ilmenite with intergranular replacement haematite bodies of the ilmenite. 1 = interbanded limonite colloform replacement structure, il = ilmenite. h = haematite. Harrar, Ethiopia. Oil-immersion with one nicol.

Fig. 663 Clastic chrome spinel surrounded by colloform structure of radiating bands of needle iron interbanded with fine colloform bands of haematite, c = chrome spinel, η = needle iron. h = haematite colloform bands. Larymna (lateritic nickel-iron deposit), Lokris, Greece. Oil-immersion with one nicol.

497

Fig. 664 Colloform structure of interbanded needle iron and fine haematite bands, surrounding and assimilating a partly resorbed aggregate of initial oolitic structures, ο = partly resorbed oolites, η = needle iron, h = haematite. Larymna (lateritic nickel-iron deposit), Lokris, Greece. Oil-immersion with one nicol.

Fig. 665 Colloform structure of interbanded needle iron and fine haematite bands, surrounding and assimilating a partly resorbed aggregate of initial oolitic structures, ο = partly resorbed oolites, η = needle iron. Larymna (lateritic nickel-iron deposit), Lokris, Greece. Oil-immersion with one nicol.

Fig. 666 Gel pitchblende (colloform) together with cubes of uraninite (white = Co-Ni minerals). In addition the subsequently formed sulfide margins of the small uraninite cubes also consist of Co-Ni. ρ = colloform pitchblende sphaeroids. uc = uraninite crystals. Schneeberg, Saxony, Germany. Oil-immersion with one nicol. χ 700.

498

Fig. 667 A general view of Fig. 668, which shows that in addition to the marginal aggregate of uraninite crystals surrounding replaced Co-Ni minerals. Gel pitchblende marginal to Co-Ni minerals is also exhibited, c = Co-Ni minerals, ρ = pitchblende marginal and replacing Co-Ni minerals, cu = crystals of uraninite forming a margin surrounding Co-Ni minerals. Schneeberg, Saxony, Germany. Oil-immersion with one nicol.

Fig. 668 Uraninite crystals forming an aggregate surrounding a Co-Ni mineral. c = Co-Ni mineral, uc = uraninite crystals. Schneeberg, Saxony, Germany. Oil-immersion with one nicol. χ 700.

Fig. 669 Colloform pyrite surrounded by a margin of pyrite crystals, c = colloform pyrite. m = margin of pyrite crystals, ρ = pyrite. Dillich near Friedendorf, Germany. Oil-immersion with one nicol.

499

Fig. 670 Psilomelane (gel-sphaeroid) and gel braunite (with fine needle crystallites), ρ = psilomelane. b = braunite, arrows show crystallites. Pforzheim, Germany. Oil-immersion with one nicol.

Fig. 671 Colloform braunite with elongated needleform crystallites, η = needleform crystallites, b = colloform banding of braunite. Pforzheim, Germany. Oil-immersion with one nicol.

Fig. 672 Dendritic-like pattern of psilomelane. ρ = psilomelane. Ilmenau, Thüringen, Germany. Oil-immersion with one nicol.

500

Fig. 673 Dendritic-like pattern of psilomelane derivative from gels, ρ = psilomelane. Ilmenau, Thüringen, Germany. Oil-immersion with one nicol.

Fig. 674 Dendritic-like pattern of psilomelane derivative from gels, ρ = psilomelane. Ilmenau, Thüringen, Germany. Oil-immersion with one nicol.

Fig. 675 Dendritic pattern of silver (mineral) associated with chalcocite. s = silver, c = chalcocite. Ilmenau, Thüringen, Germany. Oil-immersion with one nicol.

501

Fig. 676 A dendritic pattern of silver, most probably a derivative from gels, s = silver. Kupferberg, Germany. Oil-immersion with one nicol.

Fig. 677 A dendritic pattern of silver, most probably a derivative from gels, s = silver. Kupferberg, Germany. Oil-immersion with one nicol.

Fig. 678 A dendritic brown-iron pattern (gel derived dendritic texture), partly surrounding chalcocite. b = brown iron, c = chalcocite. Batebawana, Algoma, Ontario, Canada. Oil-immersion with one nicol.

502

Fig. 679 A dendritic brown-iron pattern (gel derived dendritic texture), partly surrounding chalcocite. b = brown iron, c = chalcocite. Batebawana, Algoma, Ontario, Canada. Oil-immersion with one nicol.

Fig. 680 Gel (colloform) derived lepidocrocite with prismatic haematite crystallites. 1 = lepidocrocite. h = haematite. Bieber, Hessen, Germany. Oil-immersion with one nicol.

Fig. 681 Lepidocrocite gel (colloform) and crystalline mass, lg = lepidocrocite gel. cl = crystalline lepidocrocite. Bieber, Hessen, Germany. Oil-immersion with one nicol.

503

Fig. 682 Lepidocrocite crystallites derived from gel. lc = lepidocrocite crystallites. Bieber, Hessen, Germany. Oil-immersion with one nicol.

Fig. 683 Gel (colloform) and crystallites of lepidocrocite, the latter derived from gel. lg = gel lepidocrocite. cl = crystallites of lepidocrocite. h-g = haematite-gel. Bieber, Hessen, Germany. Oil-immersion with one nicol.

Fig. 684 Gel (colloform) and crystalline (crystallites) lepidocrocite. gl = gel lepidocrocite. cl = crystalline lepidocrocite (crystallites). Bieber, Hessen, Germany. Oil-immersion with one nicol.

504

Fig. 685 Gel (colloform) and crystalline (crystallites) lepidocrocites. gl = gel (colloform) lepidocrocite. cl = crystalline lepidocrocites. Bieber, Hessen, Germany. Oil-immersion with one nicol.

Fig. 686 Gel (colloform) and crystalline lepidocrocite. gl = gel (colloform) lepidocrocite. cl = crystalline lepidocrocite. h = haematite. Bieber, Hessen, Germany. Oil-immersion with one nicol.

Fig. 687 Elongated (radiating and in cases zoned) metacinnabar crystals; crystallization from gels. Mt. Diablo, Contra Costa County, California, USA. Oil-immersion with one nicol.

505

Fig. 688 Interpenetrating metacinnabar crystals, derivatives from gel. Mt. Diablo, Contra Costa County, California, USA. Oil-immersion with one nicol.

Fig. 689 Zoned and interpenetrating metacinnabar crystals, crystallization from gels. Mt. Diablo, Contra Costa County, California, USA. Oil-immersion with one nicol.

Fig. 690 Elongated prismatic metacinnabar crystals, interpenetrating crystallization from gels. Mt. Diablo, Contra Costa County, California, USA. Oil-immersion with one nicol.

506

Fig. 691 Elongated, in cases radiating and occasionally interpenetrating hausmannite and ?cryptomelane prismatic or needle-shaped crystals, h = hausmannite. c = cryptomelane. Ilmenau, Thüringen, Germany. Oil-immersion with one nicol.

Fig. 692 Gel-derived texture of prismatic brown iron and malachite, b = brown iron, m = malachite. Bisbee, Arizona, USA. Oil-immersion with one nicol.

Fig. 693 Fine (trichite) crystallites of cassiterite derived from gels. Colloform cassiterite pattern, c = cassiterite (trichites). Shizhu Yuan near Chenzhou, Hunan Province, China. Oil-immersion with one nicol.

507

Fig. 694 Fine (trichite) crystallites of cassiterite derived from gels. Colloform cassiterite pattern, c = cassiterite (trichites). Shizhu Yuan near Chenzhou, Hunan Province, China. Oil-immersion with one nicol.

Fig. 695 Radiating prismatic cassiterite (?derivative from gel) and stannite replaced by pyrite. c = cassiterite. s = stannite. ρ = pyrite. Mina Montserrat, Bolivia. Oil-immersion with one nicol.

Fig. 696 Colloform marcasite including and replacing sphalerite, m = marcasite. s = sphalerite, arrow shows atollreplacement of sphalerite by marcasite. Peko Mine, N. Territory, Australia. Oil-immersion with one nicol.

508

Fig. 697 Relic gel pattern of galenobismuthinite. g = galenobismuthinite. r = relic gel structure. Peko Mine, N. Territory, Australia. Oil-immersion with one nicol.

Fig. 698 Niccolite associated with gangue (?carbonates) forming the central part of a colloform "sphaeroid", which is followed by a broad initial gel layer of rammelsbergite or pararammelsbergite margined by a thin rim of safflorite. c = carbonates, η = niccolite. r = rammelsbergite. s = thin margin of safflorite. Cobalt City, Ontario, Canada. Oil-immersion with one nicol.

Fig. 699 Niccolite associated with gangue (?carbonates) forming the central part of a colloform "sphaeroid", which is followed by a broad initial gel layer of rammelsbergite or pararammelsbergite margined by a thin rim of safflorite. c = carbonates, η = niccolite. r = rammelsbergite. Cobalt City, Ontario, Canada. Oil-immersion with one nicol.

509

Fig. 700 Zoned malacon with internal "diffusion-colloform patterns", m = magnetite, ma = zoned malacon. d = diffusion colloform patterns, u = uraninite. Faraday Mine, Bancroft, Ontario, Canada. Oil-immersion with one nicol.

Fig. 701 Zoned malacon with internal "diffusion-colloform patterns". m = magnetite. ma = zoned malacon. d = diffusion colloform patterns. Faraday Mine, Bancroft, Ontario, Canada. Oil-immersion with one nicol.

Fig. 702 Bird's eye structure due to alteration of pyrrhotite which results in marcasite alteration sphaeroids. ρ = pyrrhotite. m = marcasite. Bisbee, Arizona, USA. Oil-immersion with one nicol.

510

Fig. 703 Rhythmical alteration effects in davidite, assuming colloform patterns. Spring Hill, Southern Australia. Oil-immersion with one nicol.

Fig. 704 A polygonal pattern of synaeresis cracks in uraninite occupied by coffinite. u = uraninite. c = coffinite (occupying synaeresis cracks). La Crouzille, France. Oil-immersion with one nicol.

Fig. 705 Circular synaeresis cracks in uraninite partly occupied by coffinite. u = uraninite. c = circular synaeresis cracks occupied by coffinite. La Crouzille, France. Oil-immersion with one nicol.

511

Fig. 706 Synaeresis cracks of uraninite partly occupied by coffinite and galena (?radiogenic galena), u = uraninite. c = coffinite. g = galena. La Crouzille, France. Oil-immersion with one nicol.

Fig. 707 Colloform pitchblende with radiating synaeresis cracks and with galena grains present in it. ρ = colloform pitchblende, s = synaeresis cracks, g = galena grains. Look-out Claim, Marshall Pass, Saguachelo, Colorado, USA. Oil-immersion with one nicol.

Fig. 708 Colloform pitchblende with synaeresis cracks, ud = darker uraninite. u = uraninite. s = synaeresis cracks, r = radiogenic galena (lead) veinform partly occupying a synaeresis crack. Look-out Claim, Marshall Pass, Saguachelo, Colorado, USA. Oil-immersion with one nicol.

512

Fig. 709 Colloform pitchblende exhibiting curved and polygonal synaeresis cracks, u = pitchblende (uraninite). ud = dark pitchblende? ρ = polygonal synaeresis cracks, s = curved synaeresis cracks, r = radiogenic lead. Look-out Claim, Marshall Pass, Saguachelo, Colorado, USA. Oil-immersion with one nicol.

Fig. 710 Small half-moon shaped pitchblende structures often surrounding (or partly surrounding) small smaltite crystal grains; (also non-typical compact safflorite is identified by P. Ramdohr). The fine white "threads" in the pitchblende are more recent arsenopyrite. s = smaltite. ρ = pitchblende, a = arsenopyrite. Wittichen, Black Forest, Germany. Oil-immersion with one nicol. χ 1000.

Fig. 711 Gel pitchblende with intershell material consisting of clausthalite and blue-remaining covellite. u = uraninite. (p = pitchblende.) cl = clausthalite. Pinky Fault, Beaverlodge Lake, Saskatchewan, Canada. Oil-immersion with one nicol.

513

Fig. 712 Gel pitchblende with intershell material consisting of clausthalite and blue-remaining covellite. u = uraninite. cl = clausthalite. co = covellite. Pinky Fault, Beaverlodge Lake, Saskatchewan, Canada. Oil-immersion with one nicol.

Fig. 713 Gel pitchblende with intershell spaces (or layers) occupied by later clausthalite (white) and blueremaining covellite in a radiating form in the central parts of the gel pitchblende. Clausthalite occupying the central part of a gel sphaeroid is of particular interest and indicates solution of the pitchblende, ρ = pitchblende, co = covellite. cl = clausthalite. Pinky Fault, Beaverlodge Lake, Saskatchewan, Canada. Oil-immersion with one nicol.

Fig. 714 Synaeresis cracks of pitchblende occupied by strongly pleochroic Bi 2 S 3 . Galena also is present. ρ = pitchblende, b = bismuthinite. g = galena. Katherina Flacher, Weisser Hirsch, Schneeberg, Saxony, Germany. Oil-immersion with one nicol. χ 700.

514

Fig. 715 Synaeresis cracks in pitchblende occupied by "veinlets" of tetrahedrite margined by galena. Johann Evangel. Gang (Vein), Daniel, Joachimstal, Bohemia. Oil-immersion with one nicol.

Fig. 716 Synaeresis cracks in pitchblende occupied by "veinlets" of tetrahedrite margined by galena, ρ = pitchblende, t = tetrahedrite. g = galena. Johann Evangel. Gang (Vein), Daniel, Joachimstal, Bohemia. Oil-immersion with one nicol.

Fig. 717 Manganese nodule from the Pacific Ocean near Hawaii, USA. About natural size.

515

Fig. 718 Section of manganese nodule showing the nucleus and the metallic outer shell, η = nucleus. m-s = metallic outer shell. Manganese nodule from the Pacific floor, near Hawaii, USA. About natural size.

Fig. 719 Altered angular fragment of rock (basaltic piece) surrounded by fine bands of manganese oxides (ramsdellite Mn0 2 ) alternating with fine transparent mineral bands. The banding and forms developed indicate colloform structures, b = altered basaltic fragment, c-b = colloform bands of manganese minerals. Detail of manganese nodule. Pacific Ocean, near Hawaii, USA. Oil-immersion with one nicol.

Fig. 720 Banded "colloform manganese" oxides (hydro-oxides) with alternating transparent mineral bands. Also fine interspersed metallic grains in matrix. m-o = manganese colloform bands, t-m = transparent mineral bands alternating with manganese "metallic" colloform bands. i-m = interspersed metallic granules (manganese oxides) in matrix. Manganese nodule. Pacific Ocean, near Hawaii, USA. Oil-immersion with one nicol.

516

Fig. 721 Colloform structure simulating "cauliflower structure" often shown in stromatolites. Detail of manganese nodule. Pacific Ocean, near Hawaii, USA. Oil-immersion with one nicol.

Fig. 722 A nugget of "ferroplatin" with a limonitic coating and with angular protuberances grown due to accretion in the lateritic eluvial cover of altered ultrabasics. Lateritic cover of Yubdo ultrabasic ring intrusion. Yubdo, W. Ethiopia.

Fig. 723 A polished section view of a platinum "ferroplatin" nugget consisting essentially of "ferroplatin" and including corroded chromite grains. A limonitic coating of the ferroplatin is also indicated. Osmiridium is included in the "ferroplatin". ρ = "ferroplatin". c = chromite grains (arrow "a" shows corrosion prior to inclusion in the "ferroplatin"). m = limonitic coating, ο = osmiridium. Eluvial lateritic cover of the Yubdo dunite. Yubdo, W. Ethiopia. Oil-immersion with one nicol.

517

Fig. 724 Detail of Fig. 723. The corrosion of the chromite grain (c) included in the ferroplatin (ρ) is indicated by arrow, ο = osmiridium. m = limonite (and manganese oxide). Yubdo, W. Ethiopia. Oil-immersion with one nicol.

Fig. 725 Scattered chromite grains in cases with a Fe-Mn oxides margin, held together by ferroplatin. c = chromite. m = Fe-Mn oxides margin, ρ = platinum. Yubdo, W. Ethiopia. Oil-immersion with one nicol.

Fig. 726 Chromite grain with a partial coating of limonite enclosed in "ferroplatin". c = chromite. ρ = "ferroplatin". m = Fe-Mn oxides. Yubdo, W. Ethiopia. Polished section, without crossed nicols.

518

Fig. 727 Ferroplatin with a margin of Fe-Mn oxides in which scattered grains of ferroplatin are also included. Later ferroplatin mobilization marginal to the Fe-Mn oxides margin is also shown (see arrow), ρ = ferroplatin. m = Fe-Mn oxides margin, p, = later mobilized ferroplatin marginal to the Fe-Mn oxides margin. Yubdo, W. Ethiopia. Oil-immersion with one nicol.

Fig. 728 Chromite grains held together by ferroplatin which has a Fe-Mn oxides margin shown to extend into the ferroplatin. c = chromite. ρ = ferroplatin. m = Fe-Mn oxides margin, arrow "m" shows Fe-Mn oxides margin extending along cracks of the ferroplatin. Yubdo, W. Ethiopia. Oil-immersion with one nicol.

Fig. 729 Chromite grains held together by ferroplatin which has a Fe-Mn oxides margin shown to extend into the ferroplatin. ρ = ferroplatin. m = Fe-Mn oxides margin, arrow "m" shows Fe-Mn oxides margin extending along cracks of the ferroplatin. Yubdo, W. Ethiopia. Oil-immersion with one nicol.

519

Fig. 730 Detail of Fig. 728. An altered chromite grain adjacent to the ferroplatin which in turn is surrounded by a Fe-Mn margin sending a protuberance into the ferroplatin. m = Fe-Mn oxides margin, ρ = ferroplatin. c = chromite. arrow "m" = protuberance of the FeMn oxides margin into the ferroplatin. cc = corroded chromite grain. Yubdo, W. Ethiopia. Oil-immersion with one nicol.

Fig. 731 Osmiridium marginal to the ferroplatin nugget which is surrounded by a margin of Fe-Mn oxides. As arrow "m" shows, the FeMn oxides margin extends and partly engulfs and replaces the osmiridium grains in the ferroplatin. ο = osmiridium. ρ = ferroplatin. m = Fe-Mn oxides margin. Yubdo, W. Ethiopia. Oil-immersion with one nicol.

Fig. 732 Ferroplatin with a margin of Fe-Mn oxides which is mobilized along a crack in the ferroplatin. At the same time ferroplatin is mobilized and forms a thin margin on the Fe-Mn oxides (margin of the Pt nugget), ρ = ferroplatin. m = Fe-Mn oxides margin, arrow shows Fe-Mn oxides margin mobilized along a crack of the ferroplatin. arrow "p" = mobilized ferroplatin marginal to the Fe-Mn oxides margin. Yubdo, W. Ethiopia. Oil-immersion with one nicol.

520

Fig. 733 Fibriolitic silicates. Chromite with epitactic magnetite (martitized) is also included in the ferroplatin nugget, c = chromite. m = epitactic magnetite (martitized). s = fibriolitic asbestos included in the nugget. ρ = ferroplatin nugget. Yubdo, W. Ethiopia. Oil-immersion with one nicol.

Fig. 734 Chromite grain with epitactic magnetite (martitized) included in ferroplatin. The ferroplatin replaces the epitactic magnetite as is shown by the corroded outlines of the contact magnetiteferroplatin (see also arrow), c = chromite. m = martitized magnetite, ρ = ferroplatin. Yubdo, W. Ethiopia. Oil-immersion with one nicol.

Fig. 735 A general view of ferroplatin nugget with a marginal area (broad zone) in which other platinoid minerals are included (see also Figs. 737 - 739). ρ = ferroplatin, marginally including platinoid minerals (e. g., osmiridium). Yubdo, W. Ethiopia. Oil-immersion with one nicol.

0.1mm

521

Fig. 736 A general view of ferroplatin nugget with a marginal area (broad zone) in which other platinoid minerals are included (see also Figs. 737 - 739). ρ = ferroplatin, marginally including platinoid minerals (e. g., osmiridium). oi = osmiridium. r = roseite. Yubdo, W. Ethiopia. Oil-immersion with one nicol.

^

Ü

1

Fig. 737 "Roseite" (r) in intergrowth with silicate (black) in a background mass of "ferroplatin" (fp). Also other minerals containing Ni, Fe and platinoid elements are recognized, marked as (a) and (b). Yubdo, W. Ethiopia. Oil-immersion with one nicol.

fa b

or «Ρ

Fig. 738 Detail of Fig. 736. Minerals (a), (b) and (c) in a background mass of "ferroplatin" (fp). Yubdo, W. Ethiopia. Oil-immersion without crossed nicols.

522

Fig. 739 Detail of Fig. 736. Osmiridium (oi) partly surrounded by "roseite" (r) in "ferroplatin" (fp). The dark "veinlets" in the "roseite" are silicates. Yubdo, W. Ethiopia. Oil-immersion without crossed nicols.

Fig. 740 Chromite grain partly engulfed by ferroplatin in which fibriolitic silicates are included and partly replaced by later roseite. c = chromite. fp = ferroplatin. s = elongated fibriolitic silicates partly replaced by roseite. r = roseite. Yubdo, W. Ethiopia. Oil-immersion with one nicol.

Fig. 741 Ferroplatin including chromite and fibriolitic silicates (antigorite, asbestos), replaced by roseite. c = chromite. fp = ferroplatin. s = fibriolitic silicates. r = roseite, arrow shows replacement of the fibriolitic silicates by roseite. Yubdo, W. Ethiopia. Oil-immersion with one nicol.

0.1mm

523

Fig. 742 A complex textural pattern exhibiting all transitions from slightly leached to completely deferrified bauxite parts or layers of the oolites. "A": Iron-rich bauxite in which oolitic textures begin to be distinguishable (due to deferrification). "B": leached part of the bauxite with relics of non-deferrified oolites (arrow "a"). All transitions and irregularities are indicated concerning the deferrification of the oolitic layer. Arrow "b": oolite with iron-rich central part and deferrified outer layer. Arrow "c": iron-rich outer oolitic layer and a deferrified central part. The selectivity of leaching as far as the oolitic structure is concerned, does not follow the rules of layering. Parnass Bauxite, Greece. Thin section, with one nicol.

Fig. 743 Detail of Fig. 742. Bauxite part, deferrified in which relatively iron-rich oolites are left as relics (see A). Also, oolitic structure composed of a bauxite nucleus (deferrified) and surrounded by iron-depleted and ironrich oolitic layers is shown in B. This oolitic structure is probably a multigeneration oolite. Parnass Bauxite, Greece. Thin section, with one nicol.

Fig. 744 Oolites and matrix with limited iron-leaching (grey-green in color hand specimen) in contact with a pocket (patch) of bauxite showing a more advanced phase of leaching out of iron. Within this pocket relics of oolites exist exhibiting a variable degree of iron-leaching, ο = oolites with rel. limited iron-leaching (deformation, elongation of oolites is probably due to tectonic effects), ρ = pocket of iron-leached bauxite, r-o = relics of oolites, o-l = oolites with bands or central part exhibiting ironleaching. Arrow "a" = oolite elongated due to deformation with its central part leached of iron. Diasporic Bauxite, Helicon, Greece. Without crossed nicol.

524

Fig. 745 Ooids exhibiting selective Fe-leaching from some of their layers. The leaching patterns simulate microdiffusion rings. i-m = iron-rich bauxite matrix, i-r = iron-rich layer of the bauxite ooid. d-i = iron-depleted layer of the ooids. Diasporic-boehmitic Bauxite, Parnass, Greece. Thin section.

Fig. 746 Oolitic structure consisting of iron-rich bauxite bands (shells) and bands with iron leached out. i-b = iron-rich bauxite. 1 = bauxite with iron leached out. b-d = bauxite with iron relatively depleted. Bauxite, Parnass, Greece. Oil-immersion with one nicol.

Fig. 747 An advanced phase of bauxite homogenization. Relics of initial iron-rich oolitic bauxitic structures. The central part of the initial iron-rich bauxite is indicated by arrow "a". Also a diffused zone with iron leached out is indicated, d-b = diffused band (zone) of the oolite with iron depletion. b-1 = bauxite matrix. Bauxite, Parnass, Greece. Thin section without crossed nicols.

525

Fig. 748 A small oolite (a) in a later formed larger oolite (b). Parnass, Greece. Thin section, with one nicol.

Fig. 749 Oolitic structures enclosed in a larger oolite. Comparable textures have also been observed in oolitic limestones. The examination of bauxite with reflected microscopy, in addition to revealing textural details, also gives an approximate indication of the haematite (iron) content of the examined textures. 1, 2, 3, 4 = oolitic structures enclosed in a larger oolite (l-o). Bauxite, Parnass, Greece. Oil-immersion without crossed nicols.

Fig. 750 A general microscopic view of the pisolite sphaeroid. The fine circular white spots represent initial oolites equally interspersed in the sphaeroid and in the matrix (background mass). Arrow "a" shows iron-rich outer "diffusion"-ring of another sphaeroidal structure. Bauxite (gibbsitic). Arkalik, Kazakhstan. Thin section, χ 5.

526

Fig. 751 A general microscopic view of a sphaeroidal structure (pisolite) showing that it consists of relatively iron-rich and depleted rings (zones). The circular white spots interspersed equally in the iron-rich and depleted rings as well as in the matrix outside the sphaeroids represent an initial oolitic phase. Arrow "a" shows the iron-rich rings and "b" the relatively iron-depleted zones. Arrow "c" shows an oolitic structure. Bauxite (gibbsitic). Arkalik, Kazakhstan.

Fig. 752 Pisolitic sphaeroids, gibbsitic in composition, in a predominantly gibbsitic matrix (with kaolinite segregations). The pisolitic sphaeroid indicates fragmentation of its external layer. Also a gibbsitic fragment is present in the matrix indicating partial deferrification. ο = oolite. p-g = gibbsitic pisolites (sphaeroid). m = gibbsitic matrix, k-s = kaolinitic segregation within gibbsitic matrix. Arrow "a" shows fragmentation of the external layer of the pisolite. Gibbsitic Bauxite, Gargoti, Kolhapur, Maharashtra District, India.

Fig. 753 Pisolites, gibbsitic in composition, in gibbsitic matrix. Kaolinite segregations in the gibbsitic pisolites are also shown, ρ = gibbsitic pisolites, k = segregations of kaolinite. g-m = gibbsitic matrix. Gibbsitic Bauxite, Gargoti, Kolhapur, Maharashtra District, India. Oil-immersion with one nicol.

527

Fig. 754 A most complex textural pattern consisting of diffusion sphaeroids and of rounded apatite structures. r-i = rings rich in iron (diffusion sphaeroids). il = rings leached in iron, r-ap = rounded apatite. The rhythmical banding exhibited by the diffusion sphaeroids could be attributed to the operation of colloidal solutions. Boehmitic Bauxite, Niksic, Montenegro.

Fig. 755 Limonitic oolitic structures either with a nucleus of clastic chromite grain (or silicate fragment) or free of nucleus. Some of the oolites show a banded structure, others appear massive. o-c = oolitic structures compact without apparent "layering". 0-b = oolitic structure with differentiation of a central part and an outer layer. ch-n = chromite nucleus of an oolitic structure. As arrow "a" indicates a reaction took place between the clastic chromite and the limonite, often the reaction follows cracks of the chromite. The average of several spot analyses by microprobe (Augustithis and Mposkos, 1980) show: [Fe in dark margin]/[Fe in unaltered nucleus] = 0.83; [Cr in dark margin]/[Cr in unaltered nucleus] = 1.02; s-n = silicate nucleus. 1-m = limonitic matrix. Oolitic/Pisolitic Ni-laterite. Larymna, Locris, Greece (see also Fig. 819).

Fig. 756 Clastic chromite grain as nucleus and forming the greatest part of an oolitic structure. The chromite grain shows a corroded outline and indentations. Also compact oolites without clastic nuclei are shown, in = indentations, ch-n = clastic chromite grain, ο = iron-rich oolitic layer surrounding the chromite nucleus, m = limonitic matrix. Oolitic/pisolitic Ni-laterite. Larymna, Locris, Greece. Oil-immersion with one nicol.

528

Fig. 757 Haematitic pisolite marginally showing a transition zone to hydro-haematite. Also chromite remnants in the margin (partly hydrohaematized pisolite) remain unaffected. Additional clastic chromite grains are present in the hydro-haematized matrix, h-p = haematite pisolite, hy-m = hydro-haematized matrix, m-hy = partly hydro-haematized margins of the haematite pisolite, c = clastic chromite grains either present in the matrix or in the pisolite. Oolitic/pisolitic Ni-laterite. Larymna, Locris, Greece. Oil-immersion with one nicol.

Fig. 758 Haematitic pisolite (metamorphosed initial limonitic pisolite) partly transversed by crystalloblastic magnetite which has a corroded clastic chromite as nucleus, h-p = haematite pisolite, ch-n = chromite nucleus, c-m = crystalloblastic magnetite, m = martitization of the magnetite crystalloblast. Oolitic/pisolitic Ni-laterite. Near Kozani, N. Greece. Oil-immersion with one nicol.

Fig. 759 Phosphorite with phosphate ooids exhibiting different stages of deferrification. One ooid exhibits a margin rich in iron (not deferrified), a zone intensely deferrified, and a central part exhibiting an intermediate stage of deferrification. ο = ooids completely deferrified. o-i = iron-rich ooids. m = matrix. n-o = ooids exhibiting an intermediate stage of deferrification. a = margin not deferrified. b = zone intensely deferrified. c = central part exhibiting an intermediate stage of deferrification. Phosphorite. Tennessee, USA. Thin section.

529

Fig. 760 Ellipsoidal-shaped iron "ooids" often with clastic quartz grains as nucleus. c-q = clastic quartz grains (free). n-g = clastic quartz as nuclei of the ellipsoidal ooids. c-c = CaCo 3 cement (matrix). i-o = iron oolites. Iron oolite. Salzgitter, Niedersachsen, Germany. Without crossed nicols.

Fig. 761 Sphaeroidal pyrite derivative of bacterial action included in sphalerite which contains scattered chalcopyrite. ρ = pyrite. s = sphalerite. Wiesloch, Baden, Germany. Oil-immersion with one nicol.

Fig. 762 Sphaeroidal sphalerite associated with chalcopyrite. s = sphaeroidal sphalerite, c = chalcopyrite. Aljustral, S. Portugal. Oil-immersion with one nicol.

530

Fig. 763 Bacterially formed pyrite sphaeroids, an aggregate which is surrounded by covellite. ρ = pyrite sphaeroids. c = covellite. Reichenberg Mine near Dens Kreis, Rothenburg, Germany. Oil-immersion with one nicol.

Fig. 764 Bacterially formed sphaeroids of pyrite associated with galena, probably also a product of bacterial action, ps = pyrite sphaeroid. g = galena, s = sphalerite. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

Fig. 765 An aggregate pyrite sphaeroid formed by bacterial action, ρ = pyrite sphaeroids. b = blende, s = silicates. Rammeisberg, Harz, Germany. Oil-immersion with one nicol.

531

Fig. 766 Layered sphaeroid consisting of a central sphaeroid of sphalerite surrounded by a later bravoite layer which in turn is surrounded by an outer layer of sphalerite. Galena is probably replacing the bravoite. s = sphalerite sphaeroid. b = bravoite. g = galena. si = sphalerite outer layer. Mechernich, Eifel, Germany. Oil-immersion with one nicol.

Fig. 767 A sphaeroidal structure consisting centrally of small sphalerite sphaeroids surrounded by bravoite which in turn is surrounded by an outer layer of sphalerite. Galena is also present associated with the bravoite. g = galena, b = bravoite. s = sphalerite. Mechernich, Eifel, Germany. Oil-immersion with one nicol.

Fig. 768 A sphaeroidal structure consisting of sphalerite replaced by galena, s = sphalerite, g = galena, b = bravoite. Mechernich, Eifel, Germany. Oil-immersion with one nicol.

532

Fig. 769 "Sphaeroidal structure" consisting of "crystallized" carbon (thucholite) and uraninite. t = thucholite "sphaeroid". u = uraninite (crystallized), r = rutiles. Baldwyn Township, Blind River, Ontario, Canada. Oil-immersion with one nicol.

Fig. 770 Deformed molybdenite exhibiting undulating extinction. Pine Creek Mine, Inyo County, California, USA. Oil-immersion with half-crossed nicols.

Fig. 771 Deformed and fractured molybdenite, exhibiting undulating extinction. Pine Creek Mine, Inyo County, California, USA. Oil-immersion with half-crossed nicols.

533

Fig. 772 Deformed molybdenite, exhibiting undulating extinction and fracturing. The mineral has passed the limits of crystal plasticity and fracturing occurs. Pine Creek Mine, Inyo County, California, USA. Oil-immersion with half-crossed nicols.

Fig. 773 Compact chromite, transversed by a micro-mylonitic zone. myl = mylonitized chromite consisting of fine fragmented chromite. c-ch = compact chromite. Rodiani, Greece. Oil-immersion with one nicol.

Fig. 774 Curved cataclasis in chromite. ch = chromite. Sr = serpentine. Rodiani, Greece. Oil-immersion with one nicol.

534

Fig. 775 Cataclastic effects on chromite, often related with fractures. There is chromite alteration. ch = chromite. a = alteration effects. ws = Wiederverkittungs serpentine. Rodiani, Greece. Oil-immersion with one nicol.

Fig. 776 Chromite cataclastically affected with "rebinding" serpentine holding together the fractured chromite grains. ch = chromite cataclastically affected, sr-r = re-binding serpentine holding together the fractured chromite grains, s = serpentinized dunite. Xerolivado, Vourinos, Northern Greece. Thin section, with crossed nicols.

Fig. 777 Fracture pattern of chromite showing the general orientation of the cracks. Rodiani, Greece. Oil-immersion with one nicol.

535

Fig. 778 Compact cataclastically affected chromite transgressing marginally into a zone of fractured granular-chromite with intergranular brown serpentine (recrystallized serpentine). ch = chromite cataclastically affected, c = cataclastic fractures of compact chromite. s = serpentine. g-c = granular chromite zone due to chromite fracturing. Rodiani, Greece. Oil-immersion with one nicol.

Fig. 779 Transitions of massive chromite in fine granular Schlieren type due to tectonic deformation, m = massive chromite. s = Schlieren type of chromite. Rodiani, Greece. Oil-immersion with one nicol.

Fig. 780 Angular fragments of chromite in a recrystallized calcite mass. ch = chromite. ca = recrystallized calcite (marble). Drepanon, Kozani, Greece. Oil-immersion with one nicol.

536

'Γ*·* * ·

Fig. 781 Tectonically shaped Schlieren type chromite banding in serpentinized dunite. c = chromite (Schlieren bands), s = serpentine. Chalkidiki, Greece. Oil-immersion with one nicol. About natural size.

TOOTH

Fig. 782 Fine chromite "bands" folded and ptygmatically folded in anorthosite beds. c = fine chromite "bands" overfolded. pf = ptygmatically folded chromite in anorthosite beds, a = anorthosite. Dwars River, Bushveld, Transvaal, South Africa.

Fig. 783 General view of the banded chromite-anorthosite occurrence at Dwars River, Bushveld, Transvaal, South Africa.

537

Fig. 7 8 4 Chromite-anorthosite banding. Thick anorthosite at the bottom bed is free of chromite. The overlying chromite band exhibits tectonically caused irregularities (see arrow "a"). an = anorthosite thick bottom "bed" free of chromite layers, ch-a = chromite/anorthosite layering. Dwars River, Bushveld, Transvaal, South Africa. Chromite layer about 10-20 cm thick.

Fig. 785 "Plastically deformed" anorthosite body and anorthosite lenses in chromite band, an = anorthosite. ch = chromite. p-a = "plastically deformed " anorthosite band. 1 = anorthosite lens. Anorthosite/chromite layers. Dwars River, Bushveld, Transvaal, South Africa. Chromite band about 5 0 cm thick.

Fig. 7 8 6 Undulating contact of the chromite-anorthosite banding rather indicating tectonic deformation, a = anorthosite. c = chromite, arrow "c" shows fine chromite band in anorthosite. Dwars River, Bushveld, South Africa.

538

Fig. 787 Anorthosite-chromite banding with rounded anorthosite "bodies" included in a chromite band, c = chromite bands, an = anorthosite bands, a = anorthosite rounded and included in the heavier chromite. Dwars River, Bushveld, Transvaal, South Africa. ϊΐ:·* •«til

Fig. 788 Chromite bodies in anorthosite bands, often with resorbed boundaries and extensions filling cracks of the anorthosite. Arrow "a" shows chromite filling cracks of the anorthosite (remobilized chromite in cracks of the anorthosite), arrow "b" shows resorbed chromite boundaries, ch = chromite. an = anorthosite. Dwars River, Bushveld, Transvaal, South Africa.

Fig. 789 Chromite "potato", actually a tectonically fragmented chromite piece with polished faces due to tectonic effects. Domokos, Greece. Hand specimen approximately natural size.

539

Fig. 790 Boudinage of chromite bodies in serpentine and altered serpentine, ch = chromite. s = serpentine. s-a = altered serpentine, arrows indicate chromite boudinage. Kursumia, Vourinos, Greece. Hand specimen, natural size (sample with courtesy of Eng. G. Kanellopoulos).

Fig. 791 Pyrite and sphalerite tectonically fractured (or rounded) with galena tectonoplastically mobilized in the fractured zone, ρ = pyrite. s = sphalerite, g = galena. Sadon, Caucasus, Russia. Oil-immersion with one nicol.

Fig. 792 Pyrite and sphalerite tectonically fractured (or rounded) with galena tectonoplastically mobilized in the fractured zone, ρ = pyrite. s = sphalerite, g = galena. Sadon, Caucasus, Russia. Oil-immersion with one nicol.

540

Fig. 793 Sphalerite fractured and with galena plastically mobilized into the fractured sphalerite, s = sphalerite, g = galena. fg = fractures of sphalerite occupied by galena. Sadon, Caucasus, Russia. Oil-immersion with one nicol.

Fig. 794 Cobaltite exhibiting a pattern of fractures due to tectonic deformation. Blackbird District, Idaho, Colorado, USA. Oil-immersion with one nicol.

Fig. 795 Fragmentation of uraninite due to tectonic deformation (fracturing), u = uraninite. Shinkolobwe, Katanga, Zaire. Oil-immersion with one nicol.

541

Fig. 796 A chain of pitchblende fragments in gangue between two large crystal grains of rammelsbergite. The fragmentation of the pitchblende is caused by tectonic effects. r = rammelsbergite. u = uraninite. dark grey mass = gangue. Wolfschacht, Eisleben, Harz, Germany. Oil-immersion with one nicol.

Fig. 797 Zoned zircon (malacon) in uraninite. The uraninite contains cracks which do not radiate equally in all directions from the malacon margin. Magnetite is also present, u = uraninite. ma = malacon. m = magnetite, f = fractures. Faraday Mine, Bancroft, Ontario, Canada. Oil-immersion with one nicol.

Fig. 798 Fractured magnetite with malacon marginal to the fractured magnetite and with extensions of malacon along the fractures, m = magnetite, ζ = malacon (zircon), zm = marginal malacon. Faraday Mine, Bancroft, Ontario, Canada. Oil-immersion with one nicol.

542

Fig. 799 Homogenous davidite without ex-solutions, showing small rutile grains along curved lines caused by tectonic effects. black = gangue minerals. grey = davidite. whitish = rutile. Mt. Victoria Mines, Southern Australia. Oil-immersion with one nicol.

Fig. 800 Fractured colloform banded pitchblende surrounded by later mobilized (or remobilized) colloform uraninite. f = fragment of banded colloform pitchblende. rp = remobilized pitchblende surrounding the fragment. Look-out Claim, Marshall Pass, Saguachelo, Colorado, USA. Oil-immersion with one nicol.

Fig. 801 Fractured colloform banded pitchblende surrounded by later mobilized (or remobilized) colloform uraninite. f = fragment of banded colloform pitchblende. rp = remobilized pitchblende surrounding the fragment. Look-out Claim, Marshall Pass, Saguachelo, Colorado, USA. Oil-immersion with one nicol.

543

Fig. 802 Fractured colloform banded pitchblende surrounded by later mobilized (or remobilized) colloform uraninite. f = fragment of banded colloform pitchblende. rp = remobilized pitchblende surrounding the fragment. Look-out Claim, Marshall Pass, Saguachelo, Colorado, USA. Oil-immersion with one nicol.

Fig. 803 Fractured colloform banded pitchblende surrounded by later colloform pitchblende (exhibiting most complex colloform banded patterns). f = fragment of banded colloform pitchblende. sp = second "generation" colloform banded pitchblende exhibiting complex pattern. Look-out Claim, Marshall Pass, Saguachelo, Colorado, USA. Oil-immersion with one nicol.

Fig. 804 Three generations of colloform pitchblende, gl = colloform banded pitchblende fragment. g2 = second generation banded colloform pitchblende on the fragment of generation 1. g3 = third generation of pitchblende following generation 1 and 2 as a margin. Look-out Claim, Marshall Pass, Saguachelo, Colorado, USA. Oil-immersion with one nicol.

544

Fig. 805 Banded colloform pitchblende fracture with a remobilized later pitchblende following cracks of the pitchblende fragment and also forming a margin, bp = banded colloform pitchblende, rp = remobilized second generation pitchblende forming margins and following cracks of the first generation. Look-out Claim, Marshall Pass, Saguachelo, Colorado, USA. Oil-immersion with one nicol.

Fig. 806 Fine remobilized pitchblende following penetrability direction and fractures of gangue minerals. ρ = pitchblende following penetrability directions, g = gangue minerals. La Crouzille, France. Oil-immersion with one nicol.

Fig. 807 Complex pattern of association of pitchblende (remobilized) with gangue minerals. It is uncertain whether these structures result from cataclastic effects or from mobilization and recrystallization. In this case recrystallization from solutions has certainly played a role, ρ = pitchblende, g = gangue. La Crouzille, France. Oil-immersion with one nicol.

545

Fig. 808 Fractured pitchblende transversed by veinlets of cassiterite (palingenic remobilization of cassiterite). ρ = pitchblende, c = cassiterite. Grampound Road Mine, St. Stephens, Cornwall, England. Oil-immersion with one nicol.

Fig. 809 Fractured pitchblende transversed by veinlets of cassiterite (palingenic remobilization of cassiterite). ρ = pitchblende, c = cassiterite. Grampound Road Mine, St. Stephens, Cornwall, England. Oil-immersion with one nicol.

Fig. 810 Bands of chromite indicating disruptions and discontinuities within, or interbanded with, anorthosite. c = chromite. a = anorthosite, arrows show interruptions of the chromite banding. Dwars River, Bushveld, South Africa.

546

Fig. 811 Fine bands of chromite in anorthosite. The fine bands of chromite show interruptions which considering Fig. 812, are not due to faulting. The chromite fine bands represent chromite mobilizations along "bedding" or penetrability directions of the anorthosite. c = chromite. a = anorthosite, arrow shows interruptions of chromite banding. Dwars River, Bushveld, South Africa.

Fig. 812 Two fine bands of chromite interbedded (or interbanded) with anorthosite. There is a "disruption" in the continuation of only one chromite band (arrow "a"), whereas the other chromite band is continuous (see arrow "b"). c = chromite. an = anorthosite. Dwars River, Bushveld, South Africa. First chromite band is approximately 5 cm thick.

Fig. 813 A crack of anorthosite filled with chromite. c = chromite band, a = anorthosite. ρ = pyroxene segregations in the anorthosite. Dwars River, Bushveld, South Africa. Approximately natural size.

547

Fig. 814 Detail of Fig. 813. Chromite crystals following a crack within the anorthosite. ch = chromite. pi = plagioclase. g = glass of the slide. Dwars River, Bushveld, Transvaal, South Africa. Oil-immersion with crossed nicols.

Fig. 815 "Sand" grain (pebble) of chromite (transported) with a decoloration margin and worm-like solved out spaces, c = chromite. d = decoloration margin with solved out "worm-like" spaces. Magnetic fraction of rutile sand from Australia. Oil-immersion with one nicol.

Fig. 816 Chromite with solved out solution features (see also Fig. 818) in birbirite. Yubdo, Ethiopia, χ 800.

548

Fig. 817 Chromite with decoloration margin with which myrmekitic-like serpentine is often associated, c = chromite. d = decolorized margin of chromite. m= myrmekitic-like bodies of serpentine (often associated with the decolorized margins). Rodiani, N. Greece. Oil-immersion with one nicol.

Fig. 818 Chromite with solved out "channels" and decolorations associated with the solved out spaces (see also Fig. 816). c = chromite. s = solved out channels, d = decolorations, b = birbirite. Yubdo, Ethiopia. Oil-immersion with one nicol. χ 800.

Fig. 819 Limonitic oolitic structures either with a nucleus of clastic chromite grain (or silicate fragment) or free of nucleus. Some of the oolites show a banded structure, others appear massive. In cases a reaction took place between the clastic chromite nucleus and the surrounding limonitic material. 1 = limonite of the oolite, d-m = darker limonitic margins also following cracks in the chromite. c = chromite nucleus of oolitic structure. Arrows indicate that a reaction took place between the clastic chromite and the limonite, often the reaction follows cracks of the chromite. The average of several spot analyses by microprobe (Augustithis and Mposkos, 1980) show: [Fe in dark margin]/[Fe in unaltered nucleus] = 0.83; [Cr in dark margin]/[Cr in unaltered nucleus] = 1.02; s-n = silicate nucleus. 1-m = limonitic matrix. Larymna, Locris, Greece. Oil-immersion with one nicol.

549

Fig. 820 Clastic chromite grain with haematite veinlets transversing and replacing the chromite. The chromite exhibits alteration (darker margins) in contact with the haematite, c = chromite. h = haematite. m = darker margins of the chromite. Larymna, Locris, Greece. Oil-immersion with one nicol.

Fig. 821 Alteration of chromite (Mg increase due to Fe [? and chrome] depletion) following a crack pattern, c = chromite. a = alteration following a crack in the chromite. Rodiani, N. Greece. Oil-immersion with one nicol.

Fig. 822 White patches alteration of the chromite, with dark lamellae. (The white patches in this case are also associated with cracks), c = chromite. a = alteration patches, m = manganese? lamellae associated with alteration patches. Rodiani, N. Greece. Oil-immersion with one nicol.

550

Fig. 823 Chromite almost completely altered (solved, replaced and surrounded) by haematite, c = chromite relics, vh = haematite veinlet transversing altered chromite. h = haematite. Larymna, Locris, Greece. Oil-immersion with one nicol.

Fig. 824 Chromite almost completely altered (solved, replaced and surrounded) by haematite, c = chromite relics, vh = haematite veinlet transversing altered chromite. h = haematite. Larymna, Locris, Greece. Oil-immersion with one nicol.

Fig. 825 Clastic chromite grain extensively replaced by haematite veinlets (ramified and replacing the chromite). c = chromite. hv = haematite veinlets in the chromite. h = haematite. Larymna, Locris, Greece. Oil-immersion with one nicol.

551

Fig. 826 Different chromospinels, partly solved and transversed by veinlets of haematite, c = chromospinels. h = haematites, hv = haematite veinlets. Larymna, Locris, Greece. Oil-immersion with one nicol.

Fig. 827 Chromite with an alteration margin due to talc "replacement" effects on the chromite. c = chromite. t = alteration margin of chromite due to talc influence. Rodiani, Greece. Oil-immersion with one nicol.

Fig. 828 Unaffected chromite in contact with alteration stage I, in which along cracks and margins, a more pinkish material has developed representing alteration stage Π. a = unaffected chromite. b = alteration stage I. c = pinkish material along cracks, alteration stage Π. d = ilmenites, rutiles, anatases and carbonates with sharp contacts in alteration stage I. Rodiani, Greece. Oil-immersion with one nicol.

552

Fig. 829 Unaffected chromite in contact with alteration stage I, in which along cracks and margins, a more pinkish material has developed representing alteration stage Π. a = unaffected chromite. b = alteration stage I. c = pinkish material along cracks, alteration stage Π. d = ilmenites, rutiles, anatases and carbonates with sharp contacts in alteration stage I. Rodiani, Greece. Oil-immersion with one nicol.

Fig. 830 Details of Fig. 828. Chromite alteration stage I (b) transgressing to pinkish material alteration stage II (c); associated with them are enclaves containing ilmenite, rutile, perovskite and carbonates (d). Rodiani, Greece. Oil-immersion with one nicol.

0.025

Fig. 831 Details of Fig. 828. Chromite alteration stage I (b) transgressing to pinkish material alteration stage Π (c); associated with them are enclaves containing ilmenite, rutile, perovskite and carbonates (d). Rodiani, Greece. Oil-immersion with one nicol.

553

Fig. 832 Ilmenites, rutiles, ?anatases and carbonates in contact with sharp margins to alteration stage I, intermediate stages are failing. The dark mass is talc. b = alteration stage I. il = ilmenites. r = rutiles/anatases. c = carbonates, strong reflexionpleochroism. Rodiani, Greece. Oil-immersion with one nicol.

Fig. 833 Ilmenites, rutiles, ?anatases and carbonates in contact with sharp margins to alteration stage I, intermediate stages are failing. The dark mass is talc. b = alteration stage I. il = ilmenites. r = rutiles/anatases. c = carbonates, strong reflexionpleochroism. Rodiani, Greece. Oil-immersion with one nicol.

Fig. 834 Magnetite with a crack showing leaching-oxidation effects, m = magnetite. 1 = leaching (oxidation) effects. Feldbach, Binnental, Wallis, Switzerland. Oil-immersion with one nicol.

554

Fig. 835 Magnetite with oriented maghemite lamellae initiating from a crack in the magnetite, m = magnetite, ma = maghemite. c = crack (partly occupied by pyrite). Platin Mooihoek Pipe, Lydenburg, Transvaal, South Africa. Oil-immersion with one nicol.

Fig. 836 Magnetite (clastic grain) martitized with more pronounced martitization marginally, m = magnetite. mm = more intensely martitized margin. Magnetic fraction of rutile sand from Australia. Oil-immersion with one nicol.

Fig. 837 Clastic magnetite grain rounded (due to attrition) and exhibiting martitization as well as haematite margins, m = magnetite. ma = martitization lamellae, arrow shows haematite margin of the magnetite. Magnetic fraction of rutile sand from Australia. Oil-immersion with one nicol.

555

Fig. 838 Clastic magnetite exhibiting an oriented pattern of lamellae and "flakes" of martite haematite, m = magnetite, ma = martite. Magnetic fraction of rutile sand from Australia. Oil-immersion with one nicol.

Fig. 839 Idiomorphic magnetite with oriented and marginal martitization. m = magnetite, ma = martite. mg = marginal martitization. Rio Fradas Mine, Arcona, Portugal. Oil-immersion with one nicol.

Fig. 840 Idiomorphic magnetite martitized and due to extensive martitization marginally a haematite replacement margin has been formed, m = magnetite, arrow shows martitization lamella, mg = martitized (replacement margin) of magnetite. Rio Fradas Mine, Aronca, Portugal. Oil-immersion with one nicol.

556

Fig. 841 Idiomorphic magnetite also martitized with maghemite. m = magnetite, ma = martite. mg = maghemite. Rio Fradas Mine, Aronca, Portugal. Oil-immersion with one nicol.

Fig. 842 Clastic magnetite grain exhibiting marginal and zonal martitization. m = magnetite, ma = marginal martitization. ζ = zonal martite. b = broad martitization lamellae. Larymna, Locris, Greece. Oil-immersion with one nicol.

Fig. 843 Ferberite due to alteration (weathering) partly altered to haematite (and iron hydro-oxides), f = ferberite. h = haematite. Huosier Mine, Nederland, Colorado, USA. Oil-immersion with one nicol.

557

Fig. 844 Bird's eye structure in pyrrhotite. b = bird's eye structure, ρ = pyrrhotite. Trepca, Serbia. Oil-immersion with one nicol.

Fig. 845 Intermediate product in the alteration of pyrrhotite to marcasite. i = intermediate product, ρ = pyrrhotite. Trepca, Serbia. Oil-immersion with one nicol.

Fig. 846 Pyrrhotite with intermediate product of alteration, ρ = pyrrhotite. py = pyrite. i = intermediate product. Trepca, Serbia. Oil-immersion with one nicol.

558

0.1mm

Fig. 847 Colloform alteration of pyrrhotite changing to marcasite. Also idioblasts of pyrite are marginally developed. c = colloform alteration of pyrrhotite to marcasite. ρ = pyrite crystalloblasts. Laurium, Greece. Oil-immersion with one nicol.

Fig. 848 Box-work structure of lepidocrocite and haematite (due to replacement of pyrrhotite). A newlyformed margin of needle-iron is also indicated. b = box-work structure, η = needle-iron margin. Larymna, Locris, Greece. Oil-immersion with one nicol.

Fig. 849 Sphaeroidal structure of marcasite exhibiting colloform banding due to alteration of pyrrhotite. ρ = pyrrhotite. m = marcasite. Murganz near Casaccia, Tessin, Switzerland. Oil-immersion with one nicol.

559

Fig. 850 Marginal alteration of pyrrhotite into marcasite which attains sphaeroidal forms, m = marcasite sphaeroids due to alteration of pyrrhotite. ρ = pyrrhotite. Bisbee, Arizona, USA. Oil-immersion with one nicol.

Fig. 851 Rhythmic alteration of chalcopyrite to brown iron and with colloform structures of needle-iron neocrystallization. b = brown iron (alteration of chalcopyrite). η = colloform bands of needle iron. Gorob, Namibia, S. W. Africa. Oil-immersion with one nicol.

Fig. 852 Hausmannite altered to psilomelane. h = hausmannite. ρ = psilomelane. Ilmenau, Thüringen, Germany. Oil-immersion with one nicol.

560

Fig. 853 Idiomorphic crystals of siderite altered to limonite (complete pseudomorphism). 1 = limonite. r = relics of siderite. i = iron hydro-oxides. Megalo Livadi, Seriphos, Greece. Oil-immersion with one nicol.

Fig. 854 Maghemite replaced by brown iron, m = maghemite. b = brown iron. Skarn, Seriphos, Greece. Oil-immersion with one nicol.

Fig. 855 Native Bi marginally replaced by brown iron resulting in a symplectic intergrowth of the two minerals. The brown iron is most probably an alteration derivative of other iron minerals in the paragenesis. The above pattern strictly speaking is a replacement pattern of native Bi by brown iron which is most probably a weathering derivative of ironcontaining minerals, bi = Bi. b = brown iron. Espuela San Miguel, near Villanueva de Cordoba, Spain. Oil-immersion with one nicol.

561

Fig. 856 Lievrite replaced by brown iron. 1 = lievrite. b = brown iron (often exhibiting colloform structure). Mavro Pounti, Seriphos, Greece. Oil-immersion with one nicol.

Fig. 857 Lievrite replaced by brown iron. 1 = lievrite. b = brown iron (often exhibiting colloform structure). Mavro Pounti, Seriphos, Greece. Oil-immersion with one nicol.

Fig. 858 Magnetite with a replacement (alteration) margin of maghemite. ma = magnetite, m = maghemite. Halara, Seriphos, Greece. Oil-immersion with one nicol.

562

Fig. 859 Magnetite with an alteration margin of maghemite which also replaces pyrite. m = magnetite, mg = maghemite. ρ = pyrite. Halara, Seriphos, Greece. Oil-immersion with one nicol.

Fig. 860 Sphalerite corroded due to weathering and exhibiting a rim of chalcopyrite. s = sphalerite. c = chalcopyrite rim. Avalos, Concepcion del Oro, Mexico. Oil-immersion with one nicol.

Fig. 861 Uraninite showing zones due to selective leaching of altered and unaltered pegmatitic uraninite. 1 = unaltered uraninite. 2 = black margin (altered uraninite). 3 = yellow margin of secondary uranium minerals. 4 = gangue minerals, feldspars. The white circles represent areas analyzed by x-ray fluorescence spectroscopy. Gordonia, South Africa. 2x natural size.

563

Fig. 862 Detail showing the contact of 1-2 as illustrated in Fig. 861. 1 = unaltered uraninite. 2 = black margin (altered uraninite). Gordonia, South Africa. Oil-immersion with one nicol.

0.05 mm

Fig. 863 Detail showing the contact of 3-4 as illustrated in Fig. 861. 3 = yellow margin of secondary uranium minerals. 4 = gangue minerals, feldspars. Gordonia, South Africa. Oil-immersion with one nicol.

Fig. 864 Crystalline uraninite with alteration into fourmarierite and compact yellow-ore. The white parts are the unaltered uraninite (due to hard reproduction of the photomicrograph); the grey are the altered. In this instance a transition from the unaltered to altered uraninite is illustrated, f = fourmarierite. y = compact yellow ore. u = uraninite. Shinkolobwe, Katanga, Zaire. Oil-immersion with one nicol.

564

Fig. 865 Colloform pitchblende showing differential resistance to solution. Often the central layers are dissolved and the outer remain, u = uraninite. d = dissolved uraninite. Wölsendorf, Bavaria, Germany. Oil-immersion with one nicol.

Fig. 866 Former gel pitchblende changed into a mass of secondary uranium minerals. The white parts of the photomicrograph represent relics of the gel pitchblende. (The pitchblende is whitish owing to the reproduction of the photograph.) Wölsendorf, Bavaria, Germany. Oil-immersion with one nicol.

Fig. 867 Uranium minerals, secondary in formation and due to the alteration of former gel pitchblende. The former gel structure is preserved even after the pitchblende has completely altered. Wölsendorf, Bavaria, Germany. Oil-immersion with one nicol.

565

Fig. 868 Euxenite with an alteration margin in graphic pegmatite. e = unaltered euxenite. am = alteration margin of euxenite. Harrar, Ethiopia. Oil-immersion with one nicol.

Fig. 869 Euxenite with an alteration margin in graphic pegmatite, e = unaltered euxenite. am = alteration margin of euxenite. Harrar, Ethiopia. Oil-immersion with crossed nicols.

Fig. 870 Euxenite with alteration following a fracture in the euxenite. e = unaltered euxenite. a = alteration following a crack in the euxenite. Harrar, Ethiopia. Oil-immersion with one nicol.

566

Fig. 871 Euxenite with alteration following a fracture in the euxenite. e = unaltered euxenite. a = alteration following a crack in the euxenite. Harrar, Ethiopia. Oil-immersion with crossed nicols.

Fig. 872 Ilmenite replaced or altered by sphene. il = ilmenite. s = sphene (scattered sulfides in norite). Falconbridge, Ontario, Canada. Oil-immersion with one nicol.

Fig. 873 An advanced stage of davidite alteration, no original davidite is preserved. The whole mass consists of fine rutiles derived from the disintegration. Apart from the fine rutile aggregates, relatively coarse grains representing ex-solution phases in the original davidite are present. Houghton, Southern Australia. Oil-immersion with one nicol.

567

Fig. 874 Rhythmical alteration effects in davidite. These structures represent the formation of minute secondary minerals as products of the alteration and disintegration, and are considered to originate from the effects of advancing solutions which are most probably "hydrothermal". Spring Hill, Southern Australia. Oil-immersion with one nicol.

Fig. 875 Detail of Fig. 873. Alteration of davidite into an aggregate of fine titanium oxides. Houghton, Southern Australia. Oil-immersion with one nicol.

Fig. 876 Homogenous davidite, without ex-solutions, and with cracks or probable cleavage pattern along which alteration of the davidite has taken place. The alteration minerals consist of minute rutiles, haematites and ilmenites. a = alteration following cracks of davidite. d = davidite. Mount Pleasant, Australia. Oil-immersion with one nicol.

568

Fig. 877 Blomstrandine with alteration. b = blomstrandine. s = alteration shadows (channels). Hitterö, Norway. Oil-immersion with one nicol.

Fig. 878 Limonite rings and crack fillings in a leached rhyolite. rh = rhyolite. lr = limonite rings. lc = limonite crack fillings. Entoto, near Addis Abeba, Ethiopia.

Fig. 879 Feldspar sericitized and invaded by limonite. f = feldspar, s = sericite. 1 = limonite. g = rhyolitic groundmass. Entoto, near Addis Abeba, Ethiopia. Thin section, with half-crossed nicols.

569

Fig. 880 Limonite globules due to intracrystalline penetration of solutions in feldspars of the Entoto rhyolite. f = feldspar. 1 = limonite globules intracrystalline infiltration into the feldspars, g = rhyolitic groundmass. Entoto, near Addis Abeba, Ethiopia. Without crossed nicols.

Fig. 881 Limonite globules due to intracrystalline penetration of solutions in feldspars of the Entoto rhyolite. f = feldspar. 1 = limonite globules intracrystalline infiltration into the feldspars. Entoto, near Addis Abeba, Ethiopia. Without crossed nicols.

Fig. 882 Serpentinization of dunite. The original granular structure of dunite is preserved with development of antigorite and iron-oxides, a = antigorite. i = iron oxides occupying spaces of original olivine grains due to the serpentinization of the initial olivine, ic = iron oxide concentration along cracks of the serpentinized dunite. Lizard Peninsula, Cornwall, England. With crossed nicols.

570

Fig. 883 A pattern of structures due to leaching out and subsequent concentration of iron in the reddish Adigrat sandstone (Triassic) near Adigrat, Tigre, Ethiopia. The pattern shown is about 1 m across.

Fig. 884 Worm-like structures of white (grey) bauxite in red bauxite. Analyses as follows. w= white bauxite, r = red bauxite. Niksic, Montenegro. TABLE....

Table to Fig. 884. Si0 2 A1 2 0 3 Fe 2 0 3 CaO

Ti0 2

Si0 2 A1 2 0 3

Fe 2 0 3 CaO

Ti0 2

H02

Red bauxite (r)

2.4 66.3

26.9 0.21

3.63

2.1

58.0

23.6 0.18

3.18

12.3

Worm-like structure (w)

4.3 86.01

0.9 0.35

3.63

3.6

72.3

0.75 0.26

3.04

16

Mineralogical Composition (according to abundance) Boehmite, Haematite, Anatase Boehmite, Anatase, Kaolin (Goethite)

571

Fig. 885 Leopard bauxite ore consisting of red and white (grey) bauxite. Analyses as follows. w= white (grey) bauxite, r = red bauxite. Niksic, Montenegro. TABLE....

Table to Fig.885. Si0 2 AI2O3 Fe 2 0 3 CaO

Ti0 2

Si0 2 AI2O3

Fe 2 0 3 CaO

Tio 2

HO2

Mineralogical Composition (according to abundance)

29.7

0.21

2.34

12.1

6.3

0.17

2.7

14.8

Boehmite, Haematite, Goethite, Kaolin, Anatase Boehmite, Haematite, (Goethite), Anatase, (Quartz)

Leopard-ore Red

6.1

57.1

33.8

0.25

2.67

5.36 50.15

Leopard-ore White

10.5

77.1

7.4

0.2

3.16

8.94 65.7

Chemical composition of red and white parts of the leopard bauxite ore. Leopard bauxite, Niksic, Montenegro.

Fig. 886 Leopard bauxite ore. Red bauxite relic pisolitic structures in a "matrix" of relatively Fe-leached bauxite. ρ = pisolitic relic structure, m = Fe-leached bauxite matrix. Helicon, Greece. About 2/3 natural size.

572

Fig. 887 Diffusion rings (sphaeroids) in bauxite, showing a pattern controlled by a system of initial cracks (see arrows). The sphaeroids consist of Fe-depleted "zones" (white rings), of Fe-rich "zones" (a), and of "zones" intermediate in composition (b). Klisoura (Dunionia), Helicon, Greece. Approx. natural size.

Fig. 888 Diffusion rings (sphaeroids) in bauxite, consisting of rings (shells) with variable composition. The Roman numbers I-VI show the different analyzed rings (zones) the composition of which is given at the end of the caption. A = sphaeroid in which Fe-depleted zones predominate. Β = sphaeroid in which Fe-rich zones predominate. Arrow "a" shows zones intermediate in composition. Similarly, zone Π is intermediate in composition between an Fe-depleted and an Fe-rich zone. Arrows "b" show an initial pattern of cracks which control the diffusion-ring pattern. Klisoura (Dunionia), Helicon, Greece. Approx. natural size. TABLE....

Table to Fig.888.

Si0 2 Zone I (Ring) Zone Π (Ring) ZoneΙΠ (Ring) Zone IV (Ring) Zone V (Ring) Zone VI (Ring)

Major elements (oxides) [wt.%] HO2 A1 2 0 3 Fe 2 0 3 CaO Ti0 2

Zn

Ni

Trace elements [ppm] Sr Zr Y Nb

Rb

10.1

69.9

5.8

0.1

3.2

11.7

64

21

280

17

32

110

77

8.1

61.5

13.7

0.1

3.1

13.6

76

26

385

19

29

91

68

4.9

57.2

17.0

0.2

3.3

15.3

82

16

320

11

47

82

46

7.2

71.2

5.1

0.3

3.2

13.4

76

17

240

16

28

88

63

6.1

54.4

18.6

0.4

3.5

14.9

62

18

240

13

27

86

59

8.5

67.9

8.8

0.1

3.0

11.4

48

21

375

16

36

94

54

573

Fig. 889 Oolites in bauxite, ο = oolites. 1 = shell of bauxite depleted due to leaching of iron. Parnass, Greece. Thin section, without crossed nicols.

Fig. 890 Bauxite with oolites (grey in color, hand specimen). Both the matrix and the oolites have been subjected to the process of ironleaching. In contrast to these oolites and matrix which were subjected to a relatively limited iron-leaching (o-m), patches of bauxite are indicated where iron is almost depleted (p-1). r-o = relic oolites in the patches of bauxite with the iron almost depleted. 1 = leached shells or bands of oolites, r = oolite-rich in iron in the patch depleted of iron relic. Helicon, Greece. Without crossed nicols.

Fig. 891 Diffusion-alteration rings of chalcopyrite. Enrichment of copper takes place in the chalcocite ring, c = chalcopyrite. ch = chalcocite. Brown iron surrounding chalcocite. Mindouli, Cameroon, Central Africa. Oil-immersion with one nicol.

c

574

Fig. 892 A rare case of chromite cleavage exhibited in a chrome-spinel clastic grain in the Larymna laterite. c = chromite. h = haematite, cl = cleavage of chromite. Clastic chromite grain. Larymna, Locris, Greece. Oil-immersion with one nicol.

Fig. 893 Silicate following albandite cleavage, a = albandite. s = silicate. Nagyag, Siebenbürgen. Oil-immersion with one nicol.

Fig. 894 Uraninite exhibiting cleavage pattern. Arrows show cleavage following the [100] face of the uraninite. Shinkolobwe, Katanga, Zaire. Oil-immersion with one nicol.

575

Fig. 895 Triangular cleavage pattern exhibited due to defects in polishing in galena. Arrows show triangular pattern representing the cleavage of [ill]. Teufelsgrund, Münstertal, Germany. Oil-immersion with one nicol.

Fig. 896 Cleavage pattern of galena (shown due to polishing, showing galena crystals differently oriented), g = galena, c = cleavage pattern, ρ = pyrite. Laurium, Greece. Oil-immersion with one nicol.

Fig. 897 Galena crystal orientation indicated by cleavage (zonal growth of galena indicated), g = galena, c = cleavage, ρ = pyrite. Laurium, Greece. Oil-immersion with one nicol.

576

Fig. 898 A ramified crack pattern exhibited by uraninite. Compact pitchblende with branching synaeresis cracks. Rayrock Mine, Marion River District, Great Slave Lake, Canada. Oil-immersion with one nicol.

Fig. 899 Pyrite exhibiting a ramified crack pattern, c = crack, ρ = pyrite. Kvcatha, Yugoslavia. Oil-immersion with one nicol.

Fig. 900 Crack pattern as revealed on polishing of tetrahedrite. Arrows show crack pattern. Cracks became apparent due to polishing effects (since the crystal surface has been exposed to the atmosphere for several decades, Anlaufen effects). Djebel Ouenza, Constantine, Algeria. Oil-immersion with one nicol.

577

Fig. 901 Ilmenite exhibiting crosstwinning. Mühlenbach, Germany. Oil-immersion with one nicol.

Fig. 902 Covellite exhibiting curved twinning due to anisotropy. Pinky Fault, Beaverlodge Lake, Saskatchewan, Canada. Oil-immersion with slightly crossed nicols.

Fig. 903 Etch pattern of pyrite. ρ = etch pattern of pyrite. Chang Poy, Dachang, China. Oil-immersion with one nicol.

578

Fig. 904 Zircon enclosed in biotite with a typical radioactive halo. Radioactive halo also follows U-Th oxides following a crack in the biotite. ζ = zircon. h = radioactive haloes. u = U-Th granules following a crack in the biotite. Harrar granite, Ethiopia. Thin section.

Fig. 905 A double radioactive halo in biotite produced by the emanation of radioactivity from U and Th included in the zircon which occurs in the biotite. h, = first radioactive halo. h 2 = second radioactive halo, ζ = zircon. Kalbach (two mica granite), Schenkenzell, Black Forest, Germany. With crossed nicols.

Fig. 906 Uraninite and columbite in contact with microcline of pegmatite. Strong radioactive alteration halo is exhibited at the contact of uraninite with the microcline. u = uraninite. c = columbite. r-h = radioactive halo in the microcline. Hagendorf, Bavaria, Germany. Oil-immersion with one nicol. (Ramdohr has also shown a radioactive halo on the columbite.)

579

Fig. 907 Ilmenite filling a crack pattern in vanadiferous magnetite. Spinels are also following a crack pattern in the magnetite, m = magnetite, il = ilmenite. s = spinel. sc = spinels following a crack pattern in the magnetite. Magnetite plug. Kennedy's Vale, Bushveld, South Africa. Oil-immersion with one nicol.

Fig. 908 Ilmenite lamella in the vanadiferous magnetite with spinel following the margin of ilmenite/magnetite and also in intergrowth with the ilmenite. In addition, spinel follows a crack pattern in the magnetite, m = magnetite, il = ilmenite lamella, s = spinel marginal to the ilmenite lamella. sc = spinels following a crack pattern in the magnetite. Magnetite plug. Kennedy's Vale, Bushveld, South Africa. Oil-immersion with one nicol.

Fig. 909 Chromite clastic grains, some rounded, others angular, some recrystallized, associated with clinoenstatite and ? plagioclase (x-ray diffraction also shows the presence of phyllosilicates) as components of an enclave in recrystallized chromite. cr = coarse-grained chromite, a part of the main band of Wintervelt chromite layer. cc = clastic chromite grains of the enclave. g = gangue of the enclave. Chromite. Wintervelt Mine near Burgersfort, Bushveld, South Africa.

580

Fig. 910 Chromite clastic grains, some rounded, others angular, some recrystallized, associated with clinoenstatite and ?plagioclase (x-ray diffraction also shows the presence of phyllosilicates) as components of an enclave in recrystallized chromite. cr = coarse-grained chromite, a part of the main band of Wintervelt chromite layer. cc = clastic chromite grains of the enclave. g = gangue of the enclave. Chromite. Wintervelt Mine near Burgersfort, Bushveld, South Africa.

Fig. 911 Coarse-grained chromite (recrystallized clastic chromite due to metamorphism). cr = coarse-grained chromite. Chromite. Wintervelt Mine near Burgersfort, Bushveld, South Africa.

Fig. 912 Angular and rounded clastic chromite with initial "matrix" (recrystallized). c = clastic chromite grain, ?clinoenstatites metamorphically formed (exhibiting an orientation of the elongated crystal direction - see arrows). Chromite. Wintervelt Mine near Burgersfort, Bushveld, South Africa.

581

Fig. 913 Angular and rounded clastic chromite with initial "matrix" (recrystallized). c = clastic chromite grain, ?clinoenstatites metamorphically formed (exhibiting an orientation of the elongated crystal direction - see arrows). Chromite. Wintervelt Mine near Burgersfort, Bushveld, South Africa.

Fig. 914 Rounded clastic chromite grains interspersed in the Merensky Reef silicates. Pyroxene pegmatoid. Merensky Reef, Dsiekop, Bushveld, South Africa.

Fig. 915 Rounded and angular clastic chromite interspersed in the intracumulate anorthite-rich plagioclase of the Merensky Reef pegmatite and also partly present in the pyroxene, c = chromite. f = plagioclase. ρ = pyroxene crystalloblasts. Pyroxene pegmatoid. Merensky Reef, Dsiekop, Bushveld, South Africa.

582

Fig. 916 Rounded chromite (clastic grains) interspersed in the pyroxene idioblasts and in the An-rich plagioclase. c = chromite. ρ = pyroxene, pi = plagioclase. Pyroxene pegmatoid. Merensky Reef, Dsiekop, Bushveld, South Africa. With crossed nicols.

Fig. 917 Idioblastic pyroxene and xenoblastic intracrystalline plagioclase (An-rich) between the pyroxene idioblasts. ρ = pyroxene idioblast. pi = plagioclase, xenoblastic in form (intracrystalline [intracumulate]). Pyroxene pegmatoid. Merensky Reef, Dsiekop, Bushveld, South Africa. With one nicol.

Fig. 918 Platinoid mineral following a fracture in the pyroxene pegmatoid of the Merensky Reef, Dsiekop, Bushveld, South Africa. Oil-immersion with one nicol.

0 05 mm

583

Fig. 919 A general diagram showing the conditions which prevail for the deposition of gold in the Witwatersrand system, South Africa, according to a diagram of the Chamber of Mines of South Africa.

Fig. 920 A rounded gold grain (small pebble) partly engulfed by later (metamorphically) mobilized pyrite. g = gold rounded grain, ρ = pyrite margin. Witwatersrand, South Africa. Oil-immersion with one nicol.

584

Fig. 921 Recrystallized hexagonally shaped gold with extension invading and replacing an adjacent pyrite. g = hexagonal in shape gold crystal, ρ = pyrite, arrow shows relics of the replaced pyrite in the gold. Witwatersrand, South Africa. Oil-immersion with one nicol.

Fig. 9 2 2 Gold in pyrite, related to gangue veinlet. g = gold, ρ = pyrite. ν = gangue veinlet. Gold-quartz veins, Ondonoc, W. Ethiopia. Oil-immersion with one nicol.

585

IX

Fig. 923 Photograph of the model of the gold-quartz-tourmaline veins of Ondonoc, based on a sketch map included in Usoni's book. The relation of the veins to the theoretical strain ellipsoid is also indicated. The two sets of veins correspond to the nondistortion planes of the theoretical strain ellipsoid. AB and CD = nondistortion planes of the theoretical strain ellipsoid. EF = acute disectrix. HG = obtuse angle bisectrix.

Ο SULPHUR

OCCURRENCES

LOCATED O N

THE

HELICOPTER

BASIS

OF

RECONNAISSANCE

BY S . S . A U G U S T I T H I S - 1 9 6 4

(Approximate compass traverse)

νιιι 'S V.

CHEBRIT VII

Ö

LEGEND

/ Ο

/ Distance Sulphur

in km

I

occurrence

(not associated a volcanic

with

cone)

\

586

Fig. 924 A sketch diagram based on helicopter reconnaissance by Augustithis, (1964), showing the volcanic line of the Dallol-Lake Julietta region of Dankalia, Great Rift Valley, Ethiopia.

Fig. 925 Pitchblende in cases colloform associated with native Se and clausthalite. Crystalloblastic haematite is also present, ρ = pitchblende, s = native Se. cl = clausthalite. h = crystalloblastic haematite. Pinky Fault, Beaverlodge, Saskatchewan, Canada. Oil-immersion with one nicol. χ 1000 (approx.).

Fig. 926 Small Ta, Nb minerals interspersed in the apogranite. m = metamictic grain, d = secondarily formed dispersion halo, due to alteration, ap = apogranite. a = altered margin of the Ta, Nb mineral. Apogranite. Abu Dabbab, Eastern Desert, Egypt. Oil-immersion without crossed nicols.

0.05 mm

|j§ \ : f i"

^^fc?=

-

J

587

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Author Index

Abdullaev, 119 Adam, 136, 248 Adegbuyi, 176, 178 Agard, 121, 135 Agiorgitis, 227, 269 Aksyuk, 131 Amcoff, 59, 60, 241 Amov, 193 Amstutz, 148, 197 Andreyev, 147 Andrushchuk, 141 Anger, 146, 192 Anon, 134, 146 Apel'tsin, 199 Apostolov, 243 Appel, 150, 151, 173,216 Armands, 216 Artus, 1 Ashley, 125 Atanasov, 241 Attia, 196 Augustithis, 1, 2, 3, 4, 5, 6, 8, 16, 18, 19, 20, 21, 23, 39, 44, 53, 55, 61, 64, 66, 69, 75, 76, 78, 80, 85, 86, 87, 88, 89, 90, 94, 95, 97, 98, 99, 101, 102, 103, 109, 111, 114, 115, 116, 119, 123, 125, 134, 139, 141, 143, 144, 150, 153, 155, 156, 167, 168, 169, 183, 184, 185, 186, 193, 194, 203, 208, 209, 210, 213, 219, 224, 226, 227, 228, 229, 230, 233, 234, 235, 236, 237, 238, 239, 254, 255, 266, 267, 268, 269, 270, 271 Aull, 191 Ayora, 138

Babkine, 247 Bächtiger, 146 Badalov, 257 Bai Jin, 175, 179 Bajwah, 170 Ball, 122 Banäs, 136, 196 Baneijee, 109 Bannister, 127 Barnes, 151, 168,214 Barrass, 87 Barrow, 120 Barton, 58 Basham, 122 Bateman, 218 Baumann, 134 Beales, 164

Beall, 200 Becke 55, 64 Behr, 185 Bell, 109 Beloussov, 7, 144 Benvenuti, 184 Beran, 148 Bernard, 141, 170, 197, 250 Bernhard, 170 Bertolani, 139 Beruskov, 259 Betechtin, 64, 218 Bezrodnykh, 201 Bezsmertnaya, 134 Biely, 151, 193 Binda, 91 Bindeman, 200 Birchenall, 58 Biste, 140 Bj0rlykke, 208 Blanchard, 161 Bogdanov, 149, 172 Bohr, 220, 264 Boludan, 134 Bonavia, 210 Bongär Ivanov, 202 Boni, 197 Bonin, 267 Borchert, 202 Borcos, 140, 181 Bose, 109 Boström, 255 Bourgignon, 254 Bowen, 17, 153, 155, 169 Bowens, 111 Bowles, 86 Boyd, 226 Boyer, 243 Boyle, 151, 168, 202 Brandt, 209, 210 Braterman, 176 Breithaupt, 8 Brett, 58 Brineman, 196 Brodin, 32, 58 Bromley, 108 Bronner, 176 Brooks, 205 Brotzen, 124 Buchanan, 192 Buddington, 17, 136

Buryak, 120, 199, 214 Buseck, 135 Bush, 163, 182 Butler, 121 Byramjee, 206

Cabri, 58, 86, 87 Calembert, 149 Callahan, 147 Callow, 32 Cameron, 112, 153, 155, 196 Campana, 109 Campbell, 69 Cann, 102 Cannon, 68, 193 Carstens, 112 Casquet, 130 Cerny, 270 Cevales, 141 Chaikovsky, 109 Chakraborty, 114 Chamber of Mines, 157 Chasovitin, 213 Chattopadhyay, 206 Chen, 148 Cheney, 10, 149, 150, 189 Cheng Yuqui, 177 Chernitsyn, 243 Chitayeva, 201 Chowdhury, 148 Chudinov, 110 Chumachenko, 147 Cioflica, 138 Cissarz, 109, 110 Clark, 58, 127 Cohen, 171 Coleman, 140 Collins, 55, 193 Concha, 196 Condon, 159 Cordwell, 161, 162 Cornelius, 151 Cortecci, 192 Cousins, 87, 88, 116 Cox, 211

Degens, 164 Derre, 153 Derriks, 229 Desborough, 112, 115, 153, 155 Di Colbertaldo, 127, 134, 141, 145 Diamond, 180 Dickson, 109 Dill, 145 Dimitrov, 110 Doe, 194, 195 Dokov, 137 Dorosh, 138 Dozy, 164, 165 Drescher-Kaden, 1, 2, 3, 6, 13, 53, 55, 64, 71, 80, 110, 111, 118, 125, 127, 139, 150, 153, 167, 168, 183, 185, 268 Drovenik, 270 Duda, 171 Dudek, 141 Dudykina, 246, 247 Dunham, 144, 163, 165 Dunnet, 152, 212 Duparc, 85 Durazzo, 47, 48, 58, 59 Dwight, 234 Dymkova, 32 Dzhumaylo, 254

Eales, 55 Edel'shteyn, 201 Edwards, 8, 9, 10, 11, 12, 13, 15, 16, 23, 39, 41, 42, 44, 46, 47,49, 58, 59, 1 9 9 , 2 1 3 , 2 1 8 , 2 4 8 Eisenlohr, 152 Ekwere, 122 Eliseev, 116, 202 Elliott, 126, 127, 253 Ellis, 260 Emerson, 112, 153, 155 Engin, 108 England, 91 Erdmannsdörffer, 64, 118, 125 Ericksen, 180 Ernst, 18 Eskenazi, 241 Evans, 108

Crittenden, 205, 255 Crockett, 107, 108

Dahanayake, 91 Daly, 64, 118, 183, 213 Darnley, 68 Dash, 109 Davidson, 151, 163, 208, 209, 231 Davis, 149, 194 De Carvalho, 124 de Kun, 247, 250 De Wiesse, 35 Deb, 202 Dechow, 146, 191 Deer, 20, 153, 155 618

Fabricius, 91 Favorskaya, 110 Fedink, 140 Fersmann, 167 Fettel, 109 Figueiredo, 60 Finlow-Bates, 102, 161 Firsov, 141 Fisher, 161 Fleming, 196 Rood, 47 Floyd, 102 Foissy, 139 Fominykh, 115 Fontbote, 148

Fonteilles, 135 Ford, 151 Forster, 117 Foster, 174, 176, 179 Fox, 189, 194 Fralick, 174, 177 Freas, 181, 182 Frenzel, 68, 171 Friedrich, 95, 152 Fruth, 192 Fuchs, 86, 92, 270 Fyfe, 254

Gabert, 135 Garson, 108 Gaspar, 202 Gaspar da Cruz, 140 Gastil, 149 Gaudefroy, 204 Gauthier, 179 Gavelin, 191, 200 Gavrilov, 267 Geffroy, 135 Gehör, 179 Geijer, 120, 150, 196 Genkin, 6, 89, 117, 269, 270 Geoffroy, 197 George, 211 German, 170 Gibbs, 174, 175 Giblin, 181 Gies, 252 Gilevich, 58 Gilligan, 174, 176, 179 Giuliani, 69, 259 Gjelsvik, 152 Goffinet, 121 Goldich, 173 Goldin, 236, 237 Golding, 246 Goldschmidt, 2, 3, 132, 150, 168, 188, 189, 190, 191, 215, 218, 218, 227, 233, 235, 238, 247, 258, 261, 262, 263, 264, 266, 268 Goni, 86, 88, 168, 269 Gorovoy, 199 Goryainov, 150 Gorzhersky, 134 Gottfried, 112 Gouanvic, 247 Graeser, 185, 186 Graham, 78 Graig, 67, 218 Grains-Smith, 176 Gregory, 193 Griffiths, 172 Grigorieff, 3 Grigorjew, 98 Grinenko, 10, 149, 150, 189, 191, 217 Grip, 112 Gross, 175 Groves, 182 Gruner, 58

Grushkin, 181 Gruszczyk, 145 Guha, 171 Guillou, 149 Gunn, 89

Hall, 161, 197, 200 Hallbauer, 158 Hamilton, 121, 122, 194, 248 Harder, 151 Hardy, 77 Harris, 87 Hauy, 262 Haw, 60 Hawley, 60, 127, 146, 251 Haycock, 229 Heath, 195 Hein, 185 Heller, 110, 139, 184, 185 Hernes, 193 Heversy, 264 Hewitt, 200, 204, 254 Heyl, 165 Hilmy, 126 Hirst, 108 Hisahide, 269 Hladikovä, 184 Hock, 95 Hodge, 86, 88 Hoekstra, 92 Holl, 108, 110 Holland, 172 Holmquist, 3 Honma, 269 Hoppe, 174 Hors field, 127 Hoshino, 136 Hosking, 246 Hu Lunchi, 226 Hügi, 211 Hurst, 194 Huttl, 126 Hutton, 4, 175, 184, 269 Hyndman, 175

Idman, 70 Idriceanu, 140 Ihde, 234, 272 Il'vitskiy, 203 Ilani, 153, 210 Imeokparia, 122 Ishibashi, 216, 241 Ito, 258 Ivanova, 259 Iwao, 214 Ixer, 218

Jankovic, 108, 145 Jaskolski, 122, 149 Jedwab, 92 Jemson, 8 Jensen, 146, 147, 191, 192 Joensuu, 258 Johan, 218 Johannson, 189, 194

Kabesh, 205 Kalinin, 134, 236, 237 Kalliokoski, 67 Kaminskaya, 194 Kamp, 255 Kaneda, 48 Kantor, 151, 193, 252 Katayama, 146 Kato, 91 Kattamis, 197 Kautsky, 150 Kazanskii, 110 Kazmitcheff, 250 Kennedy, 17 Kern, 239 Kerr, 231 Khazov, 136 Kidd, 229 Kilikova, 202 Killingworth, 68 King, 109, 145, 146, 211 Kinkel, 170, 197 Kinloch, 87, 88 Kitamura, 205 Kittl, 146 Klemm, 108 Klominsky, 149 Kloosterman, 120 Kluth, 233 Knight, 148 Koark, 119, 150, 167 Kobe, 148 Koen, 117, 158 Kolkovsky, 201 Kon'kova, 243 Konstantinov, 109 Kopchenova, 231 Koppel, 4, 268, 270 Korhinsky, 138 Kornilov, 212 Korobeinikov, 138 Koroleva, 137 Korzhinsky, 152, 213 Kostakis, 21 Köster van Groos, 113 Kostov, 218 Koulomzine, 197 Kranz, 92 Krasnikov, 203, 209 Kraume, 35, 142 Krause, 201 Krauskopf, 260 Krige, 158 620

Krüger, 153 Krumbein, 91 Kulcsär, 172, 243 Kullerud, 58, 150, 227 Kulp, 191 Kurilifhikova, 259 Kusnir, 140 Kutina, 7, 109 Kuznetsov, 138 Kvaccek, 258

Laajoki, 179 LaBerge, 173, 175, 178, 216 Lacerda, 197 Lacy, 126 Lafforgue, 135 Lambert, 144 Lambet, 260 Landergren, 113, 150, 216 Landwehr, 107 Lange, 10, 149, 150, 189 Lanier, 177 Larsen, 112 Lattanzi, 184, 271 Laurikko, 260 Lawrence, 30, 33, 121, 209, 246, 251 Lazär, 141 Lebedev, 172 Leblance, 95 Legraye, 121 Lehne, 148 Leleu, 196, 197 Leonard, 136 Lepp, 173, 176 Leroy, 151 Leube, 109 Leutwein, 229 Levicki, 80, 83 Lewis, 195 Liebenberg, 158 Lindgren, 218 Lister, 115 Ljunggren, 125, 192, 255 Lopez, 120, 195, 196 Löttemoser, 159, 160 Lowell, 130, 132, 135 Lukas, 141 Lummen, 132 Lundberg, 133 Luo Hui, 175, 179 Lyell, 6, 183 Lyon, 248

Macdonald, 179 Machairas, 88, 135, 202, 269 MacKevett, 200 MacLean, 58 Macnamara, 128 Majumder, 177 Maleyev, 243

Malyutin, 112 Mangan, 151, 172 Marinelli, 200 Markham, 121, 196, 199, 214, 251 Markov, 176 Markova, 242 Marmö, 111, 124, 146 Marshall, 258 Mason, 107, 108 Matheson, 109 Matveyenko, 120 Maucher, 110, 171, 192, 218 McConchie, 176, 178 McLaughlin, 218 Mehnert, 111 Meindre, 206 Meituv, 241 Meixner, 197 Mendelejeff, 220, 231 Mendelsohn, 109 Menyaylova, 65 Meyer, 220 Miller, 253 Minato, 257 Mincheva-Stefanova, 193, 218 Mineev, 234 Mitchell, 108, 109 Mitscherlich, 262 Mochnacka, 122, 137 Monseur, 134, 136, 139, 145 Moore, 152, 200, 212, 252 Moravek, 202, 241 Morimoto, 58 Morton, 153 Mozaffari, 200 Mposkos, 99 Mukheijee, 114, 203 Murdoch, 218 Murray, 11, 161,211 Murthy, 174, 176 Muta, 258 Myers, 231

Nambu, 205 Narväez, 195 Nash, 152 Neinavaie, 271 Neiva, 32 Nekrasova, 231 Nesbitt, 185 Newberry, 131 Newhouse, 58, 218 Nickel, 202 Niggli, 17,94, 111, 153 Nikolayev, 260 Nikulin, 247 Nishihara, 147 Nixon, 211 Novak, 32, 258 Nysten, 60

O'Driscoll, 107 O'Meara, 23, 162 Odikadze, 236 Oelsner, 212, 218 Oen, 133 Oftedahl, 119, 150, 168, 258 Ogilvie, 58 Öhlander, 121 Ohnenstetter, 224 Okada, 205 Olade, 153, 176, 178 Olson, 126 Omenetto, 141 Onasick, 165 Onikhimovsky, 267 Oosterbosch, 229 Orberger, 108 Orcel, 218 Orliac, 243 Orlov, 33 Ortlepp, 35 Osman, 195 Ostwald, 77, 78, 91 Ottemann, 4, 68, 78, 86, 87, 88, 89, 98, 101, 103, 116, 153, 168, 171,226, 233, 269, 270 Ovchinnikov, 111, 112

Paar, 197 Pagnacco, 127 Panagos, 98, 115, 116,226 Paneth, 233 Papunen, 70 Park, 134 Pashinkin, 260 Patterson, 116 Pattrick, 192 Paul, 174 Paunen, 91 Pavlov, 110, 248 Peck, 112, 113, 208 Pedersen, 96 Perezhogin, 120, 125 Perichaud, 170 Perrin, 125 Perseil, 205 Pertsev, 130, 131 Petersen, 69, 193 Petrascheck, 114, 171, 183, 184 Petrov, 147 Petrovskaya, 272 Petruk, 256 Petrulian, 214 Philpotts, 76 Picot, 218, 243 Pieruccini, 4, 267 Piestrzynski, 195 Piirainen, 133 Piispanen, 117 Pilot, 191 Pisa, 213 Piznyur, 180 Plaza, 218 621

Plimer, 159, 160, 259 Podlessky, 131, 135 Pohl, 164, 183 Polferov, 24 Polge, 146 Poljakowa, 246 Poplavko, 209 Popov, 137, 147 Popova, 213 Poty, 151 Pozdnyak, 213 Prinz, 205, 254 Procyshyn, 145 Prouvost, 58 Purkait, 206

Qi Cbangmuo, 226 Quensel, 59

Radkewitsch, 246 Raguin, 145 Rai, 174 Ramdohr, 4, 6, 58, 64, 69, 89, 106, 115, 116, 117, 144, 150, 158, 218, 230, 247, 268, 269, 270 Randall, 127 Read, 111 Reed, 126, 127, 254 Reeves, 128 Reeves Macdonald, 195 Regan, 109 Reh, 158 Rehwald, 218 Reimer, 173 Rekharskii, 231 Remy, 220, 227, 240 Rentzsch, 146, 191 Reynolds, 56, 118, 133, 195 Richards, 145, 159, 193 Rickard, 91, 189, 194 Ripley, 192 Robbins, 177 Robinson, 189, 193, 232, 235 Romanesko, 203 Romes de Lille, 8 Rose, 86 Rosebaud, 261, 264 Rosier, 210 Ross, 164 Roubault, 125 Roy, 204, 205, 206 Rudnitskaya, 231 Rugheim, 163 Russell, 161 Rutten, 216

Saager, 58, 142, 157, 158 Sainsbury, 121, 122, 248 Salvadori, 141 622

Samal, 58 Sander, 64 Sanklov, 259 Sassano, 145 Savel'yev, 199 Sawkins, 108, 165 Sawlowicz, 91 Schachner-Korn, 69, 96 Schieber, 175 Schmidt, 181 Schmidt-Eisenlohr, 58 Schneiderhöhn, 11, 218, 269, 270 Scholl, 4 Schouten, 12 Schröcke, 136, 259 Schroll, 192, 268, 270 Schuiling, 110 Schüller, 230, 231 Schulz, 142, 147, 197, 270 Schwartz, 41, 42, 45, 46, 49, 57, 58 Schweigart, 170 Seccombe, 192 Segnit, 159 Seguin, 142, 145 Sekine, 171 Serebryakov, 122 Severgin, 8 Shabynin, 133 Shadlun, 145 Shaposhnikov, 108, 109 Sharkov, 203, 209 Sharma, 114 Sharp, 121, 248, 249 Shatolov, 120 Shawe, 214 Shcheglov, 171 Shcherbakov, 120 Shcherbina, 242 Shchlegov, 142 Shegelski, 174, 177, 179 Shei Kwang-Hong, 134 Sherbakov, 125 Shergina, 194 Sheridan, 140 Shields, 194 Shnaider, 200, 209 Shterenberg, 211 Sillitoe, 58 Sill, 178 Simic, 170 Simonson, 177 Sims, 137 Sinkovec, 29 Skinner, 112, 113 Skvortsova, 231 Slatkine, 32 Smellie, 260 Smirnov, 109, 214 Smith, 112 Snyman, 92 Sobotka, 121, 126 Sofoulis, 120 Soler, 138 Solomon, 162, 182, 192

Sorby, 180 Sorem, 196 Spronsen, 220 Staatz, 172 Stanton, 23, 147, 159, 161, 191 Starke, 134 Staudinger, 78 Stemprock, 197, 248 Stille, 110 Stillwell, 159 Stoll, 251, 252 Stolyarov, 134, 181 Strunz, 50, 71, 72, 106, 236 Strydom, 179 Studemeister, 260 Stulchikov, 174, 179 Stumpfl, 87, 89, 102, 117, 141, 247, 269, 270 Styles, 89 Sugaki, 58, 181 Sushchevskaya, 259 Suslova, 24 Suzuki, 189 Swayne, 128 Szolnoki, 202

Takahashi, 258 Takeuchi, 195 Takimoto, 257 Tan, 141 Tanelli, 271 Tanida, 205 Tarkian, 117 Taylor, 35, 47, 48, 58, 59, 91, 140, 201, 216, 254 Tchernokolev, 180 Teh, 104 Thalhammer, 116 Thayer, 114 Threadgold, 127, 202 Thum, 271 Timofeeva, 201 Tishkin, 229, 231 Tollon, 243 Tooms, 4, 163, 187, 209, 237, 254 Tornos, 130 Touray, 181 Toussaint, 254 Trask, 128 Travis, 116 Treiber, 210 Trepka-Bloch, 170 Trurnit, 171 Tschanz, 35 Tu Guangzhi, 133 Tufar, 58, 245 Tupper, 192 Turneaure, 140 Tweto, 212

Udubasa, 8 Uytenbogaardt, 218

Vaasjoki, 189, 193 Vakhrushev, 138 Valsami, 102 Van der Veen, 58, 218 Van Rensburg, 115 van Spronsen, 272 Vanderwilt, 121 Varvill, 166 Vasilyev, 209 Vaughan, 218 Velinov, 171 Venerandi-Pirri, 148 Verhoeven, 58 Vernadsky, 3, 7, 98, 185, 268 Vervoort, 189, 194 Vezrer, 8 Vgenopoulos, 4, 75, 210, 219, 228, 237 Victor, 121 Vincienne, 200 Vinken, 135 Vinogradov, 10, 149, 150, 189, 191, 217 Vistelius, 155 Vlad, 138 Vokes, 64, 67 von Gehlen, 158 Von Hevesy, 233 Von Rahden, 158, 211

Wager, 20, 153, 155 Wagner, 269, 270 Wandke, 58 Wang Shufeng, 130, 135 Warden, 200, 206 Warrington, 152 Watanabe, 136 Watkinson, 224 Watson, 196 Watts, 114 Wauschkuhn, 271 Wederpohl, 2, 188 Wegener, 17 Wei Diu-Yin, 259 Welin, 92, 138 Westenberger, 270, 271 Westerveld, 218, 255 Westland, 86, 88 White, 127, 189, 194, 216 Whittle, 138 Whittles, 147, 192 Willemse, 112, 115, 238, 239, 240 William, 120 Williams, 136, 197, 248 Wilson, 126, 128, 175 Winchester, 102 Wirth, 174, 175 Wolbeck, 226 Wolff, 170 Wood all, 116 Worley, 32 Wretblad, 59 Wright, 257 Wrucke, 126 623

Wu Jiashan, 177 Wu Liren, 135

Xang Xinyi, 5 Xu Keqin, 1, 236, 267

Yajima, 181 Yamaoka Kazuo, 257 Yang Bo-Lin, 248 YangMinzhi, 138 Yershov, 213 Yesikov, 193 Yuan Kuirong, 5 Yudalevich, 137

624

Zainullin, 260 Zang Qiling, 130 Zar'yan, 242 Zaraisky, 130, 131 Zarevich, 133 Zartman, 194 Zen Chin-Fung, 248 Zhan Quisheng, 179 Zhao Yiming, 137 Zharikov, 129, 138 Zheng Ming Hua, 130, 132 Zhu Shangqing, 133 Ziebold, 58 Zimmermann, 148 Zucchetti, 142 Zuev, 33 Zuffardi, 141, 148, 168

Subject Index to the Text Part

Ab/An of plagioclase, 155 abnormal concentrations of Pb, Zn and Cu, 148 acanthite, 199, 202 accretion, 4, 86, 87, 88, 89, 167, 168 accretion hypothesis, 87 accretion of PGE, 89 accretion processes, 168 actinolite, 195 actinolitic tremolite, 177 adamite, 201 addition of negatively charged electrolyte, 78 adularia, 248 adularization, 171 advancing fronts of solutions, 30 advancing wave, 152 aeschynite-priorite series, 232 Africa's golden arc, 157 Ag, 127, 146, 172, 184, 186, 215 Ag in exogenous gold, 202 Ag-bearing beudantite, 202 Ag-bearing sulphides, 200 Ag-pentlandite, 224 Ag-Sb deposits, 194 Ag-Sb-rich stage, 199 Ag-telluride, 241 Ag/Au ratio, 142, 157 Ag/Pb ratio, 145 Ag/Sb + Bi ratio, 241 age relations, 193 agents of metal transportation, 259, 260, 272 agents of transportation, 268 agglutination, 89, 227 agglutination process, 86 agglutinations of colloidal dimensions, 78 aikinite, 33, 162 albandite, 72, 105, 204, 205 albite, 181 algal dolomite, 211 Algoma-type, 175 Algoma-type BIF, 177 Algoma-type iron formation, 174 all water in the earth's crust, 98 allemontite, 59 allemontite-Π, 59 allemontite-ΠΙ, 59 allotrioblastic, 64

almandine, 134, 177 alpitic type of chromite deposits, 156 altaite, 32, 242 alteration, 4 alteration of chromite by talc, 99 alteration zone, 3 alteration-oxidation, 204 altered chromite grains, 87 alteriles, 170 alumina-enriched paleosols, 173, 174 aluminum-trend, 70 aluminum/gallium, 264 alunitized, 172 amphibole, 178, 181 amphibolite, 151, 177, 178, 212 amphibolite facies, 55, 148 amphibolites containing up to 2% W, 151 amygdales, 91 anaerobic, 163, 166 analogous, 220 anatase, 99, 102 anatectic granites, 18 anatectic granitoid, 125 anatectic interpretations, 18 anatexis, 111, 120, 123, 125, 139, 143, 144, 189, 267 anchi-sediments, 3, 153, 154, 155, 267 andersonite, 245 andesite, 141, 151, 170, 174, 193, 194, 205, 214 andorite, 197, 200 anglesite, 31, 81, 151, 193, 200, 202 anglesite-covelline, 201 anhydrite, 172, 181, 190 ankerite, 127, 142, 174, 185 Anlaufen effects, 68, 105 anorogenic granites, 5 anorogenic magmatic granitic complex, 122 anorthosite, 64, 239 anorthosite/chromite, 155 anorthosite/chromite banding, 5 anorthosite/chromite layering, 154 anoxygenic atmosphere, 216 anthophyllite blasts, 114 anticline, 146 antigorite, 94, 103 antigoritic asbestos, 89 antimonide, 116 625

antimonite, 32, 36, 62, 149, 213, 214, 243 antimony, 110, 140, 141, 142, 189, 200, 213, 270 antimony-gold, 202 antipathetic relationship between V 2 O s and the Ti0 2 content of the ore, 240 antipathy, 5, 44, 115, 238 antipathy of Cr to Ti, 74 antipathy of Ti to Cr, 240 antipathy of Ti to V, 240 apatite, 121, 122, 181 apatite inclusions, 181 aplite, 121, 123 aplite formation as products of substitution, 167 aplitic, 139 aplitic veins, 3 apogranite, 1, 3, 4, 39, 255 apogranite mineralization, 186 apogranitic bodies, 186 apogranitic cupola, 248 apogranitic paragenesis, 237 apomagmatic paragenesis, 187 apophysis, 141 apparition of minerals, 8 Appelella Ferrifera, 177 aramayoite, 137 Archaean iron formation, 174 Archaean iron formations, 178 Archaean lithologies, 175 Archaean or Proterozoic age, 178 arenites, 177 argentiferous galena, 127, 140, 202 argentiferous stannite, 171 argentiferous tetrahedrite, 127, 171 argentite, 31, 32, 44, 62, 128 argyrodite, 171 arid conditions, 209 arsenic, 59, 60, 110, 186 arsenic sulphides, 186 arsenide, 116, 251 arsenides-sulphides, 199 arsenopyrite, 30, 34, 37, 47, 65, 68, 72, 84, 121, 122, 126, 127, 135, 137, 140, 141, 171, 179, 186, 195, 196, 200, 201, 202, 209, 213, 230, 243, 248, 249, 252 arsenosiderite, 201 As, 184 As, Sb and Bi related as main elements of the V family, 228 As-Sb, 200 asbecasite, 186 asphaltite, 92 assimilation, 119 assimilation and contamination processes, 169 assimilation in basalts, 169 association, 229 atelestite, 201, 202 Atlantis Π Deep, 151 atmophile groups, 261 atmospheric oxygen, 216 atoll replacement of magnetite by chalcopyrite, 25 atoll type, 13, 30, 33, 58 626

atoll-type magnetite replacement, 24 atoll-type replacement, 26, 29, 33, 37, 58 atoll-type replacement of berthierite, 25 atomic diameters, 44 atomic migration, 14 atomic radii, 2, 14 atomic weight, 220, 235 Au, 124, 146, 172, 184, 215, 243 Au mobilization, 269 Au tellurite, 31 Au-Ag tellurides, 216 Au-Ag-telluride minerals, 241 aureole, 118, 129, 131, 134, 135, 136, 153, 213 aureoles of porphyries, 150 aurichalcite, 200 auriferous conglomerates, 209, 211 auriferous ferroplatin, 88, 89 auriferous granitic magmas, 120 auriferous quartz, 86 auriferous veins, 126 auto-aggregation, 168 autocathartically pushed, 67 autocreative skarns, 129 autohydration of ferriferous silicates, 134 autoliths, 114 autometasomatic alteration, 152, 213 autometasomatism, 116, 122 autometasomatites, 138 autoserpentinization, 24 autostibnite, 60, 140

azurite, 116, 128, 140, 145, 200, 201, 202

n

B / 1 0 B ratios, 194

B-group elements: Mn, Fe, Cu, Ag, Au, 251 bacteria, 162, 172, 173, 177, 192 bacterial action, 10 bacterial fission, 177 bacteriogenic reduction, 192 bacteriogenic-syngenetic, 148, 192 banded anorthosite chromite layers, 97 banded iron, 18 banded iron formations, 173, 174 barite, 145, 148, 164, 179, 196, 230, 255 barite-manganese deposits, 200, 254 baryte, 127, 141, 142, 147, 151, 165, 166 barytized, 172 Basal Reef conglomerate, 157 basalt, 39, 107, 152, 160, 170, 172, 174, 188, 237, 238, 239, 256 basalt conduit, 21 basalt sills, 142 basaltic flow, 227 basaltic lava, 254 basaltic magma, 153 basaltic magma differentiation, 111, 139, 143, 155, 169 basaltic materials, 267 basaltic melts, 150 base-metal stratabound deposits, 184

basements of the continents, 267 basic and ultrabasic rocks, 87 basic front, 132, 133, 139, 196 basic front mobilization, 6, 56, 129, 132, 134 basic front release, 118, 136, 137 basic-ultrabasic magma differentiation, 155 bauxite, 50, 85, 90, 103 Be, 121, 135, 172, 251 Be-associated elements, 250 berndtite, 104 berthierite, 24, 127, 140, 162, 189 beryl, 121, 250, 251 beryllium, 172 berzehanite, 137 berzelianite, 232 betafite, 181, 232 betechtinite, 230 Bi, 34, 35, 50, 53, 58, 216, 241 Bi, Sb, As as characteristic elements of hydrothermal paragenesis, 234 Bi minerals, 138, 230 Bi telluride, 242 Bi-Cu-Mo, 200 Bi-Te-S system, 121 Bi/Sb ratio, 241 BIF (banded iron formation), 173 big bang hypothesis, 5 bimetasomatic calcareous skarn, 130 bimetasomatic skarns, 129 bimodel iron-enrichment, 177 bindheimite, 202 biogenetic origin, 162, 192 biogenic influence, 146 biogenic origin, 191 biogeochemistry, 194 biological activity, 172 biological precipitation, 177 biostasy, 92 biotite, 121, 181 biotite-muscovite granite, 121, 248 biotitization, 212 birbirites, 6, 85, 86, 87, 168, 226 birbiritization, 4, 87, 89, 226 birbirization process, 85, 86 bird's eye structure, 53, 100 birnessite, 205 bismoclite, 200 bismuth, 122, 137, 199, 200, 202, 248, 251 bismuth minerals, 195 bismuth ochres, 202 bismuth-bismuthinite, 121 bismuth-rich paragenesis, 200 bismuth-selenium-bearing galena, 197 bismuthide, 116 bismuthinite, 33, 34, 37, 38, 39, 56, 84, 116, 135, 140, 141, 199, 200, 201, 202, 230, 232, 241, 243, 248, 249, 251 bitumen, 244 bixbyite, 206, 254 black smokers, 172

blacks with a wide range of uranium-molybdenum, 231 Blastese, 64 blastesis, 65 blastic gudmundite, 53 blastic magnetite, 36, 44, 75 blastic pyrite, 37, 38, 53 blastic-endoblastic growth, 132 blende, 35, 122, 142, 149, 152, 195, 241 blomstradite, 232 bodies with maximum reduced specific surface, 78 boghead structures of coal, 92 borate, 135 bornite, 25, 27, 29, 35, 37, 38, 39, 53, 56, 57, 58, 59, 116, 117, 120, 126, 128, 135, 136, 145, 146, 148, 149, 194, 195, 196, 200, 201, 202, 213, 214, 230, 232, 252 bornite overgrown on rutile, 72 bornite-chalcopyrite, 58 bornite-digenite, 149 bornite/chalcocite microenvironment, 66 boron, 259 boron isotopes, 194 boron minerals, 138 botryoidal structures, 79, 142, 204 boulangerite, 25, 27, 140, 213, 241, 252 boundary state, 77 bournonite, 62, 140, 141, 149, 252 boutinage, 18, 20, 114 boutinage structures, 5 box wood structure, 35 box-work, 100 braunite, 40, 52, 53, 72, 82, 204, 205, 206, 254, 255 bravoite, 55,59, 69, 92, 116 breakdown symplectites, 13, 59, 60 breccia pipes, 58 brecciation, 141 Briefe von der Tiefe, 18 brines, 153, 163, 181, 185, 208, 210, 227, 245 brittle quartz, 95 brochantite, 200 Broken Hill controversy, 159 bronzite-chromite veinlets, 155 brown iron, 35, 81, 101, 103, 104 brown iron crystallites, 83 brown iron dentritic pattern, 82 brown iron precipitation, 26 Brownian movement, 78 bulk composition changes, 213 Bushveld Complex, 112, 115, 116, 141, 153, 154, 155, 189, 269 Bushveld tin deposits, 247 bustamite, 137, 195, 206

C and Ο isotopes, 194 Ca skarn, 130 Ca-skarn types, 129 cadmium, 140, 228, 241 cafarsite, 186 calaverite, 32, 151, 202, 242 627

calc-alkaline series, 175 calcic Fe skarn, 130 calcic magma, 190 calcic W skarn, 130 calcioskarn, 135 calcite, 53, 66, 179, 181, 195 calcite-fluorite-fluorapatite assemblage, 159 calcium skarn, 130 calcium-uranium molybdate, 231 calcsilicate gneiss, 212 caledonite, 200 CaMg(C0 3 ) 2 , 173 camouflage, 264 carbon dioxide, 259 carbonates, 31, 36, 179 carbonization, 24 carrollite, 139, 149 cassiterite, 4, 25, 36, 60, 69, 80, 97, 101, 120, 121, 122, 125, 136, 140, 141, 149, 181, 187, 197, 212, 213, 232, 237, 246, 247, 248, 250, 251, 259 cassiterite crystallites, 83 cassiterite sphaeroid, 83 cassiterite-polymetallic sulphide deposits, 271 cassiterite-sulphide, 182 cataclasis, 94 catalytic role, 177 cauliflower structures, 85 causation principles, 218 cavernous, 204 cavity filling, 65 Cd, 186, 202, 215 cell wood structure, 35 cerussite, 127, 149, 151, 193, 195, 200 chalcanthite, 201 chalcedony, 154, 172 chalcedony spherules, 177 chalcocite, 26, 28, 29, 35, 37, 38, 44, 56, 58, 59, 60, 61, 66, 82, 91, 103, 126, 127, 128, 136, 140, 141, 147, 148, 149, 201, 202, 214, 230, 232, 252 chalcocite along the microcracks of the pyrite, 28 chalcocite layer, 104 chalcogenides, 260 chalcophanite, 205 chalcophile, 261, 266 chalcophyllite, 201 chalcopyrite, 24, 27, 28, 29, 30, 35, 36, 37, 38, 39, 40, 43, 44, 45, 46, 47, 53, 54, 56, 57, 58, 59, 60, 62, 65, 72, 76, 91, 101, 103, 115, 116, 117, 120, 121, 122, 126, 127, 128, 134, 135, 136, 139, 140, 141, 142, 145, 146, 148, 149, 150, 151, 152, 162, 172, 179, 186, 189, 194, 195, 197, 200, 201, 202, 213, 214, 224, 228, 230, 232, 241, 243, 246, 248, 249, 250, 252, 254 chalcopyrite + chalcocite, 59 chalcopyrite + covellite, 59 chalcopyrite infiltrations, 26 chalcopyrite lamellae, 66 chalcopyrite replacing bornite, 25 chalcostibnite, 141 chalmersite, 189 628

chamosite, 216, 254 chaotic conditions, 272 chaotic relationship, 76, 218 chemical composition and crystal structure, 262 chemical precipitate, 176, 177 chemical sedimentation, 159 chemical substitution, 262 chemical thermodynamics, 260 chemogenic sediments, 174 chert, 176, 177 chert-bearing rocks, 173, 216 chimneys, 172 chloanthite, 33, 34, 37, 38, 58 chloanthite-smaltite, 141 chloride complexes, 163 chloride fluids, 131 chlorine, 259 chlorite, 127, 181, 195 chlorite-rich brines, 163 chlorite-rich brines from Sabkha sediments, 182 chrome metasomatism, 269 chromite, 40, 44, 55, 61, 65, 69, 70, 71, 74, 75, 85, 86, 87, 88, 89, 90, 94, 95, 98, 104, 105, 108, 109, 112, 114, 115, 116, 143, 153, 157, 158, 224, 226, 227, 238, 239, 266, 269 chromite enclave, 154 chromite interbanded with anorthosite, 95 chromite leopard ores, 95 chromite mantle bodies, 156 chromite micromylonitized, 94 chromite nucleus, 90 chromite occupying a fine fracture, 97 chromite potato bodies, 269 chromite sand layers, 155 chromite sands, 154 chromite sphaericules, 20 chromite xenocryst, 75, 99 chromite-anorthosite, 155 chromite-anorthosite interbanding, 156 chromite-bronzite, 114 chromite-layering, 153 chromites in birbirites, 99 chromospinel, 19, 21, 40, 44, 61 chrompicolite, 134 chrysocolla, 81, 200 cinnabar, 29, 65, 141, 172, 196, 199, 209, 213, 243, 244, 260 classification device, 272 clastic chromite, 82 clastic chromite grains, 99, 154 clausthalite, 34, 53, 59, 66, 84, 137, 214, 232, 243 clausthalite occupies inter-shell spaces, 84 cleavage, 105 cleavage of galena, 105 cleavage pattern, 62, 68 clinopyroxene, 130, 178 clinozoisite, 195 closed system, 134 cloth-pattern, 240 Co, 215 Co and Cd in pyrite, 170

Co and Ni in pyrite, 195 Co content of pyrite, 68 Co-arsenopyrite, 252 Co/Ni ratios, 147, 152, 257 coagulation, 78, 81, 82, 91, 231 coagulation of colloids, 77 coagulation of gels, 78 coal substance, 35 coalified algae, 92 coarse dispersed, 78 coarsely dispersed aggregates, 78 cobalt, 40, 82, 261 cobalt ores, 252 cobaltite, 31, 32, 37, 38, 75, 96, 141, 212, 230 coffinite, 35, 75, 84, 137 collector mechanism, 162 collision, 130, 143 collisional environment, 108 colloform, 4, 36, 39, 85, 101, 133, 147, 204, 231, 257 colloform cassiterite, 83 colloform malachite, 81 colloform marcasite, 83 colloform patterns, 77 colloform pitchblende, 39, 40, 75, 81, 105, 268 colloform secondary pattern, 102 colloform sphaeroid, 100 colloform structure, 27, 77, 79 colloform structures secondarily formed due to alteration, 83 colloid particles, 78 colloid to crystalline, 70 colloidal condition, 77, 78 colloidal dimensions, 77 colloidal discontinuities, 77 colloidal dispersed, 78 colloidal influx, 204 colloidal precipitates, 176 colloidal precipitation, 78 colloidal solution, 82 colloidal state, 77 colloidal sulphides, 170 colloidal suspension, 260 colloidal types, 204 colonies of sulfidized bacteria, 91 coloradoite, 32, 202 columbite, 50, 52, 53, 71, 72, 74, 75, 106, 121, 122, 247, 248 columbo-tantalite, 250 commencement of geological history, 267 common (joint) segregation of elements, 224, 226 common segregation and mobilization (remobilization) of Fe, Ni, Co, Cu, Zn, 227 common segregation of Cu, Zn, Ag, Cd, 228 common segregation of metallic elements, 219, 264, 265 comparative anatomy, i, 2, 15, 139, 154 comparative textural analysis, 153 complex gel patterns, 81 composition/crystal structure, 264 concentration-segregation of rare metals, 167 concordant syngenetic types, 151

condensation, 78, 79, 82 connate, 165 connate water, 163, 165, 182 constructive activity of solutions, 183 contact metamorphism, 56 contact metasomatic, 134 contact metasomatism, 131 continental drift, 110 continental drift theory, 17 continental margins, 175 continental rifting, 161 continental runoff, 175 continuous haematite-ilmenite series, 49 continuous isomorphic crystallization, 14 continuous isomorphic series, 15 continuous solid solution series, 41, 48, 49, 52 controversy, 143, 144, 173 controversy over platinum nuggets, 85 convection current hypothesis, 17 convergent boundary environment, 108 copiapite, 201 copper, 151, 159, 213 copper belt, 109 copper deposits, 172 copper mineralization, 161, 210 copper ore deposit, 134 copper ore paragenesis, 214 copper ores, 161 copper porphyry, 128 copper-bearing pyrite, 208 copper-molybdenum, 180 copper-nickel, 149 copper-nickel sulphide deposits, 212 core, 266 core replacement, 13 core-mantle mobilization, 107 coronadite, 204, 205 corpuscular dispersive state, 77 correlation coefficients, 175 corundum crystalloblasts, 50 cosalite, 136, 200, 202, 230, 241 covalent character, 188 covariance pattern, 201, 206 covelline, 146, 149, 202 covellite, 27, 35, 37, 38, 54, 55, 59, 91, 105, 116, 126, 128, 136, 140, 141, 145, 148, 200, 202, 212, 214, 232 covellite veinlets, 25 Cr, 215 Cr leaching, 100 crater, 197 crenulations, 161 critical zone, a layered series of pyroxenite, norite and anorthosite, 155 crookesite, 232 cross twinning, 105 cross-shaped sphalerite in pyrite, 47 crushed zone, 19 crustal mush, 108 crustal sources, 191 629

crustification, 9, 141 cryptocrystalline structures, 257 cryptocrystalline textures, 147 cryptomelane, 196, 204, 205, 206, 207 cryptoscally aggregates, 231 crystal fractionation, 269 crystal plasticity, 5, 7 crystal structure and chemical constitution, 261, 264 crystal-lattice, 169 crystal-lattice relationships, 72 crystalline structure, 262 crystalline uraninite, 102, 105 crystallite aggregates, 82 crystallites, 75 crystallization sequence, 10, 23, 40 crystallization sequence in mantle, 74 crystallization temperatures, 180 crystalloblast, 44, 154 crystalloblast-poikiloblast, 21 crystalloblastesis, 11, 64, 67 crystalloblastesis of pyrite, 66 crystalloblastic, 32 crystalloblastic corundum, 50 crystalloblastic garnet, 40 crystalloblastic growth, 65, 67 crystalloblastic pyrite, 29, 100 crystalloblastic sequence, 167 crystallochemical factors, 219 crystalloplastic, 18, 114, 189 crystalloplastic forsterite-rich olivines, 266 crystalloplastic medium, 5 crystals and gel pitchblende, 79 Cu, 120, 127, 146, 152, 172, 186, 215 Cu and W mineralization, 248 Cu deposits, 146 Cu-As ore, 186 Cu-As-Pb sulfide mineralization, 211 Cu-Au-Ag ore, 151 63 Cu/ 65 Cu ratios, 194 Cu-Ni ore of Pechenga, 116 Cu-Ni sulphide deposit, 194 Cu-Pb-W veins, 171 Cu-Pb-Zn-Ag sedimentary deposits, 148 Cu/Ag ratios, 241 Cu/Pb/Zn ratios, 170 Cu/Zn ratios, 170 cubanite, 30, 47, 115, 224, 246 cumulate chromite, 112, 155 cumulates, 226 cumuloporphyroblastic, 64 cupriferous sandstones, 147, 149 cuprite, 26, 27, 50, 60, 69, 81, 128, 200, 201 curved twin intergrowth "plane", 68 cycle of processes, 204 cyclic-metamorphosed banded iron, 176 cyclothemic pattern, 174 cylindrite, 246

630

dacite, 170, 214, 248 dacite porphyry, 128 danaäite, 252 dark alteration margins in clastic chromite grains, 99 datolite, 141 davidite, 35, 50, 54, 84, 96, 102, 232, 233 debasification, 189 debasification processes, 18 decoloration, 98, 99 decoloration margin, 5, 61, 98, 104 decoloration of chromite, 61 deductive, 12, 14, 15 deductive approach, 10 deep crustal ruptures, 107 deep-seated fracture zones, 109 deep-weathering tropical soil, 201 deferrification, 90 deformability of ore minerals, 94 deformation, 96, 185 deformation lamellae, 19 delorenzite, 232 dentritic, 65 dentritic pattern of lepidocrocite, 82 depositional environments, 173 depth factor, 131 derivation environments, 5 derivation of metals from a mantle source, 189 descloizite, 195 de volatilized mantle, 160 diagenesis, 91, 139. 148, 164, 165, 175, 184, 211 diagenetic, 176, 177, 178, 211 diagenetic iron silicate minerals, 175 diagenetic mobilization, 196 diamictite, 179 diapiric bodies, 20 diapiric mantle, 269 diapiric mobilizations, 18 diapiric ring complexes, 19 diapirism, 7, 108, 116 diapirs, 5, 19, 164 diastrophic phase, 114 diatoms, 91 dickite, 199 differences in colloidal solutions, 79 different layers of the earth, 266 different valency states, 188 differential leaching, 2, 4, 90, 267 differential zonality, 119 differentiated layer, 116 differentiated sills, 116 differentiation, 5, 6, 18, 20, 39, 118, 119, 121, 122, 126, 143, 152, 156, 169, 200, 213, 247, 248, 261 differentiation gravitative layering, 153 differentiation theory, 143 difformation, 77, 78 diffusion, 12,14, 15, 42, 43, 59, 131 diffusion banding, 142 diffusion of solution, 167 diffusion rate, 15, 60

diffusion replacement, 12 diffusion ring structure, 90, 103 diffusion sphaeroid, 90 digenite, 44, 149, 214, 232 dimension definition, 78 diopside, 135, 138, 195, 197 disequilibrium, 60 disordered state, 14 dispersion, 77, 78, 82 dissociation of Sn and W, 248 dissolution of chromospinel 99 dissolution of ore constituents, 216 dissolution of the central part, 80 dissolution-assimilation, 197 distal Sn skarns, 129 distal turbidites, 174 distribution behaviour of metallic elements, 272 divergent boundary environment, 108 diversification, 11 djurleite, 145 dolomite, 130, 179, 195 dolomite rhombus, 195, 196 dolomite-ankerite, 179 dolomitic marbles, 130 dolomitization, 134 double halo in biotite, 106 double radioactive haloes, 106 double salts, 134 dual interpretation, 145 duftite, 200 dunite, 19, 85 dunitic complexes, 18 dyscrasite, 157, 199, 243 dysoxic muds, 175, 178 dzhezkazganite, 209

earth's core, 5, 7 edenite, 178 electrically conducting mineral, 12 electron pair forces, 264 electronegative, 41 electronegative effect, 15 electronegativity, 188 electropositive, 41, 188 electrum, 126 element agglutination, 87, 226 element distribution, 261, 272 element interrelationships, 272 element leaching, 210 element migration, 27, 31, 37, 52, 57 element mobilization, 56 element mobilization in solid state, 125 element recycling, 211, 212, 250, 266 elements of the main groups, 220 elements present in primary U-paragenesis, 233 ellipsoid shaped chromite, 95 eluvial lateric cover, 87 emanation fluids, 132

emanations, 213 emery, 50 empirical laws, 1, 2, 8, 52 empirical laws of element interrelations, 86 empirical laws of element interrelationship according to the periodic system, 272 empirical laws of element segregation, 204 empirical laws of element segregation/concentration, 261, 272 empirical laws of the periodic system, 87, 115, 142, 168, 169, 186, 219, 220, 224, 227, 228, 234, 235, 242, 246, 247, 248, 249, 254, 257, 258, 259, 260, 266, 272 emplacement, 11 emplacement of granitoid plutons, 110 emplectite, 33, 230 emulsion, 42 emulsion ex-solution, 41 emulsion-like bodies of chalcopyrite, 46 emulsion-like type of distribution of fine chalcopyrite in blende, 46 emulsoid bodies, 46 en mass tectonic mobilization, 268 enargite, 25, 26, 27, 28, 37, 38, 56, 66, 128, 140, 194, 200, 202, 209, 211, 245 enargite replacement by pyrite, 24 enargite-famatinite, 200 enclave of clastic chromite grains, 153 Endoblastese, 64 endoblastesis, 143 endoblastic growth, 143 endogeneous activity, 110 endogenetic, 267 endogenic, 5 endogenic deposits, 109 endogenic mineralization, 110, 248 endogenic ore formation, 267 endogenic tin, 5 endogenous, 145 endoskarn, 130 epi-magmatic, 121, 168 epi-mesothermal, 127 epi-syngenetic, 211 epi-syngenetic mineralization, 145 epicontinental seas, 209 epidosites, 102 epidote, 195 epigenetic, 11, 124, 142, 144, 146, 148, 150, 151, 152, 153, 159, 161, 162, 163, 169, 193, 195, 211 epigenetic controversy, 157 epigenetic deposits, 151 epigenetic hydrothermal, 159 epigenetic manganese mineralization, 153 epigenetic metallogeny, 153 epigenetic mineralization, 152 epigenetic minerals, 145 epigenetic ore deposits, 194 epigenetic origin, 119 epigenetic Pb-Zn deposits, 153 epigenetic telethermal, 151 631

epigenetic U-bearing, 152 epigenetic unconformity, 145 epigenetic versus synsedimentary, 144 epirogenic, 110 epitactic, 18, 19 epitactic bournonite, 72 epitactic growth, 71 epitactic growth of magnetite, 71 epitactic growths and inclusions, 74 epitactic jacobsite, 72 epitactic magnetite, 20, 71, 85, 89 epitactic magnetite on chromite, 71 epitactic patterns, 72 epitactic relationships of minerals, 73 epitactic spinel, 43, 72 epitactic uraninite, 50, 75 epitaxis, 71 epitaxis of magnetite on chromite, 72 epithermal, 9, 12, 118, 139, 141, 142, 151, 159, 213 epithermal Cu-Pb-Zn deposits, 214 epithermal metallogeny, 152 epithermal solutions, 169 epithermal stage, 4 epithermal-subvolcanic, 141 epithermal-telethermal, 34 equilibrium conditions, 8 equilibrium constants, 136,259 erlichmanite, 224 erniggliite, 186 etch-figures, 105 eugeosyncline, 138 eutectic conditions, 53 eutectic crystallization, 15, 23, 41, 53, 56, 57, 60 eutectic intergrowth, 13 eutectic simultaneous, 54 eutectics, 15 eutectoid, 30, 54 euxenite-polycrase series, 232 euxinite, 102 evaporite, 165 evaporite horizons, 208 evaporite series, 163, 190 ex-solution, 9, 11, 12, 13, 14, 15, 16, 30, 41, 42, 44, 45, 47, 48, 49, 50, 52, 56, 57, 58, 59, 71, 74, 115, 247 ex-solution bodies, 8, 45, 50 ex-solution formation, 10 ex-solution forms, 26 ex-solution intergrowth, 9 ex-solutions of ilmenite, 42 ex-solve, 14, 23, 96 ex-solved mineral, 42 ex-solved phases, 10 ex-solved spinel, 44 exchange of cations, 58 exhalative, 160, 179, 210 exhalative S due to volcanic activity, 189 exhalative sedimentary, 120, 145, 150, 170, 172, 194, 197, 254 exhalative sedimentary ores, 119 632

exhalative sediments linked to volcanic exhalative sulphide deposits, 170 exhalative submarine, 173 exhalative type, 172 exhalative volcanic deposits, 170 exhaling hydrothermal fluids, 174 exhalites, 159, 170 exogenetic, 209, 267 exogenetic uranium, 209 exogenous, 145 exoskarn, 130, 148 expansion of earth-hypothesis, 7 expansion of the earth, 17, 143 experimental ex-solution textures, 59 experimental mineralogy, 41, 144 experimental petrofabrics, 6 extraterrestrial, 7 extraterrestrial bodies, 2, 7 exudation, 3, 5, 6, 133, 167, 185, 213 exudation phase, 118, 123 exudation products, 154

falkmanite, 27 fan-shaped deltas, 157 Fe, 146 Fe, Cu (Co) Zn, Be, 129 Fe, Co, Ni, Cu, Ag (mantle derived elements), 260 Fe and Cr mobilization, 85 Fe-Cu-Co sulphide, 149 Fe-Mn margin, 88 Fe-Mn oxide, 89 Fe-Ni-rich laterites, 90 Fe-rich microenvironment, 99 Fe-Zn-Mn oxide, 147 Fe/Mg ratios, 155 feather-like, 65 Fe/Cr ratios, 155 feldspathization, 159, 248 ferberite, 32, 100, 252 fergusonite-formanite series, 232 ferri-halloysite, 203 ferriferous ilmenite, 48 ferrihydrite, 175, 178 ferrimagnetic properties, 91 ferro-salite, 148 ferroan pargasite, 178 ferrohypersthene, 177 ferromagnesian carbonates, 65 ferroplatin, 86, 87, 88, 89, 226 ferroplatin nuggets, 87, 89, 168, 269 ferroplatinum, 116 ferrosalite, 177 ferruginous and manganiferous sediments, 216 ferruginous quartzite, 150, 151 fibriolitic asbestos antigorite, 89 fibriolitic crystallites, 83 field of FeS, 161 field of stability, 2

fillings of platinoid minerals, 154 fixed-continent concept, 17 flexural flow, 179 flocculation of iron silicate, 210 fluid exhalation, 160 fluid hydrocarbons, 93 fluid inclusion, 140, 151, 165, 180, 181, 182, 192, 259, 260 fluid inclusion cation ratios, 180 fluid introduction, 132 fluid mixing, 165 fluid-solution, 152, 213 fluids, 11, 119, 131, 136, 151, 152, 170, 195, 216, 260, 265, 268 fluids in metamorphic ores, 183 fluids in metamorphics, 183 fluids in metamorphism, 183 fluids in skarns, 132 fluorborite, 135 fluorides of divalent elements, 263 fluorine, 259 fluorite, 121, 127, 136, 141, 142, 145, 147, 164, 166, 172, 181,255 fluorite structure, 263 fluorite-barite accumulations, 145 fluorite-baryte zoning, 165 fluvial model of iron introduction, 175 fore-arc basin setting, 175 formation, 229 forsterite crystalloblastesis, 134 forsterite-rich olivines, 19 forsteritic olivine, 215 fossil placer, 114, 158 fossil rift valley structures, 109 fossilization, 35 fossils, 35 fourmarierite, 102 fractional crystallization, 9, 111, 112, 114, 153 fractional differentiation, 112 fractionation, 124, 200 fractured pyrite, 96 fragment of rock, 85 fragmentation, 96 fragments of the mantle, 21 framboid, 91 framboidal, 133 framboidal clusters of digenite, 91 framboidal pyrite, 91 franckeite, 140, 200, 246 franklinite, 48 free energy of formation, 9 freibergite, 141, 202, 252 friction, 139 friction generated heat as energy source for hydrothermal processes, 184 friction generating heat, 185 friction of various crust rocks, 169 friction zone, 21 front replacement, 30 fuchsite, 114, 269

fumarolic, 121 fumarolic sulphide, 192 fumarolic-solfatara stage, 170 fundamental source of metals, 267 fused silicates, 261 fusion chambers, 268 fyalite-rich molecule, 50

Ga, 202 gahnite, 159 galena, 27, 28, 30, 31, 34, 35, 36, 37, 38, 39, 45, 46, 47, 50, 53, 55, 56, 58, 59, 62, 66, 81, 83, 84, 92, 105, 120, 121, 122, 126, 127, 128, 135, 136, 140, 141, 142, 145, 146, 147, 149, 150, 151, 152, 161, 166, 172, 179, 186, 191, 192, 193, 194, 195, 196, 197, 200, 202, 205, 209, 212, 213, 228, 230, 232, 241, 242, 243, 245, 246, 248, 250, 252, 257, 258 galena as a substitute, 29 galena crystals, 193 galenite, 231 galenobismuthinite, 83, 135 gallium-bearing sphalerite, 140 gangue, 61, 83 gangue minerals, 37 gangue-galena veinlets, 38 garnet, 40, 62, 65, 130, 135, 178, 206 garnet-quartzite, 159 garnetiferous rock, 62 garnetiferous skarn, 138 garnetization, 159 gas and coexisting liquid phases, 261 gas-liquid inclusion, 181, 183, 259 gas-liquid inclusions in quartz, 132 gaseous emanations, 132 gaseous exhalations, 191 gases, 265 gathering together into paragenetic associations of elements rarely abundant in the earth's lithosphere, 272 Ge, Sn and Pb interrelated as main elements of the IV family, 228 Gefügekunde, 96 gel (colloform) structures, 78 gel coagulation, 85 gel gratonite, 80 gel minerals, 79 gel phase interbanded, 80 gel pitchblende, 79, 102, 105, 231 gel pitchblende and crystalline uraninite, 82 gel pitchblende sphaeroids, 80 gel psilomelane, 82 gel pyrite, 82 gel relic structures, 80 gel sphaeroids, 70, 78, 84 gel sphaeroids of lepidocrocite, 82 gel sphalerite, 80 gel state of silica, 177 gel structures, 77, 78 gel to sol reversibility, 82 633

gel-hyaloid structures, 78 gel-pyrite, 228 general state of matter, 77, 78 generation, 229 generations of ore minerals, 11 geochemical culminations, 110 geochemical differentiation, 270 geochemical index horizon, 110 geochemical interrelationship between segregated elements, 229 geochemical signature, 152 geochronology, 144 geodes with galena crystals, 146 geoenvironment, 3, 4, 5, 18, 60, 61, 118, 202, 238 geoenvironmental conditions, 20, 89, 195, 216, 219, 265 geological spiral, 2, 3, 6, 18, 20, 188, 189, 212, 216 geopetal textures, 148 geosynclinal, 7, 110 geosynclinal hypothesis, 143 geosynclinal sediments, 125 geosynclinal theory, 17 geosynclines, 109, 145, 149, 152 geotectonic events, 119 geotectonic evolution, 119 geotectonic model, 110 geotectonic setting, 108, 130 geotectonics, 7 geothermal gradient, 181 geothermometry, 12 geringeres Diffusionsvermögen, 79 germanite, 27, 28, 30, 56, 140, 200 germanium, 261 germanotype settling, 152 gersdorffite, 60, 72, 116, 140 Gibbs free energy, 260 gibbsitic, 90 gigantic boutinage structures, 155 glaucodote, 252 global settings, 108 global tectonic concept, 143 global tectonics, 7, 107, 109, 110, 143 global tectonics and metallogeny, 266 globular structures, 79 gneiss, 178 goethite, 101, 126, 127, 147, 196, 201, 202, 203 goethite-hydrogoethite, 203 gold, 31, 32, 37, 39, 58, 60, 62, 68, 69, 76, 86, 88, 117, 120, 121, 125, 126, 127, 128, 135, 139, 140, 141, 142, 145, 151, 152, 157, 158, 168, 172, 179, 195, 196, 197, 199, 201, 202, 211, 212, 213, 214, 226, 243, 261 gold accretion, 168 gold and silver tellurides, 242 gold coagulation, 269 gold deposits, 152, 157 gold mineralization, 126, 138, 174, 179 gold mineralization in iron formations, 179 gold nodules, 85 gold nuggets, 88, 168, 269 gold provinces of Siberia, 199 634

gold replacing an adjacent pyrite, 158 gold-antimony, 200 gold-arsenopyrite skarn, 138 gold-bearing, 214 gold-bearing conglomerates, 211 gold-bearing district, 121 gold-bearing ore veins, 202 gold-bearing reefs, 157 gold-boron mobilizations, 268 gold-pyrite placers, 216 gold-quartz lenses, 168 gold-quartz veins, 180 gold-quartz-tourmaline veins, 168, 185 gold-rich zones, 68 gold-skarn formation, 138 gold-tourmaline, 200 gold-tourmaline-pyrite, 269 gold-uranium placers, 158 gold/quartz, 53 Goldschmidt's law of element distribution, 272 gondite, 204 gossan, 193, 201, 202 grado-separation, 2, 4, 102, 267 grain boundaries, 16 granite controversy, 108, 137, 144 granite emplacement, 189 granite porphyry, 130, 135 granite-endogenic tin, 5 granites and 'granites', 123 granitic emplacement, 133 granitic intrusions, 250 granitic intrusive magma, 189 granitic magma, 132 granitic magma hypothesis, 118 granitic magma intrusion, 124, 247 granitization, 39, 56, 111, 118, 119, 124, 137, 139, 143, 144, 146, 153, 183, 186, 188, 189, 213, 247, 267, 268 granitization related to mineralization, 123 granitization sensu stricto, 111 granitization theory, 18 granitization-skarnification, 271 granitophile, 3, 4 granitophile element, 39, 119, 187, 246, 255 granitophile group of elements, 232 granoblastic, 64 granoblastic magnetite, 60 granoblastic textures, 67 granodiorite, 126, 168, 185, 269 granodiorite porphyry, 127 granodioritization, 126 granodioritization-granitization, 168 granolepidoblastic, 64 granonematoblastic, 64 granular chromite, 61 granules, 103 granulite facies, 177, 178 granulites, 177 graphic, 15, 19, 53, 54 graphic intergrowth, 13

graphic pegmatites, 167 graphic quartz/feldspar intergrowth, 150 graphic skeletal spinel, 20 graphic symplectites, 13 graphic-like intergrowth, 53 graphic-like spinel, 19 graphic-like symplectites, 23, 59 graphic-myrmekitic, 54 graphite, 179, 183 graphitized, 92 gratonite, 81 gravitational differentiation, 95 gravitational settling, 20 gravitative accumulation, 112, 155 gravitative crystal settling, 155 gravitative crystal settling/differentiation, 155 gravitative differentiation, 112 gravitative slumping, 161 gravity anomalies, 108 gravity differentiation, 114 graywacke, 173, 216 great reactivity, 55 greenalite, 178 greenockite, 228 greenschist, 152, 214 greenschist facies, 149, 270 greenschist grade metamorphism, 151 greenstone, 116, 146, 161, 168, 174, 227 greenstone belts, 152, 179, 194 greenstone schists, 179 greisen, 120, 132, 136, 197, 199, 251 greisenization, 248, 251 Grenzschicht Zustand, 77 grossular, 138 grossular-andradite garnet, 197 group A (mainly crust-derivative elements) = U, (Th), Pb, Bi, As, Sb, Se (Sn and Mo), 229, 230, 234, 235 group A = As, Sb, Bi, Pb, Sn Se, Te, 254, 255 group A = Pb, As, Sb, Sn (Se), 245 group A = Se, Te, Bi, Sb, As, Pb, Sn (Tl, In) (?Mo), 241, 242, 243 group A mainly crust derivatives, 235 group Β (mainly mantle-derivative elements) = Fe, Co, Ni, Cu, Zn, Au, Ag, 229, 230, 234, 235, 241, 242, 243 group Β = Fe, Co, Ni, Cu, Zn, Au, Hg, Ag, 243 group Β = Fe, (Μη), Co, Ni, Cu, Zn, Cd, Au, Ag, Hg, (Cr, V, (Ti?)), 254 group Β = Fe, Μη, Cu, Zn (Cd), 245 group Β = Μη, Fe, Co, Ni, Cu, Zn, Cd, Au, Ag, 251 group Β = Μη, Fe, Co, Ni, Cu, Zn, Cd, Au, Ag, Hg, 258 group Β mainly mantle derivatives, 235 group C = U, Th, Y, Zr, (HO, Nb, Ta, Ti and rare earth, 233, 234 group C = Zr, (Hf), Ti, Nb, Ta, Mo, W, U, 254 group V, Nb, Ta, Mo, W, U, 251 groutite, 205 growth zones, 193 grunerite, 178, 179 guanajuatite, 243

gudmundite, 53, 65, 121, 140 gypsum, 147, 202 gypsum-anhydrite deposit, 245

habit and trace element content, 258 habitus, 69, 70 haematite, 27, 33, 35, 36, 37, 38, 54, 55, 58, 59, 65, 66, 75, 81, 82, 90, 100, 105, 128, 136, 138, 147, 148, 157, 175, 178, 196, 202, 210, 214, 230, 254, 255, 256 haematite subidioblast, 66 halite, 181 halogen compounds, 150, 259 halotrichite, 201 hastingsite, 177 hastite, 34 hauchecornite, 34 hausmannite, 83, 101, 204, 205, 206, 255, 256 hawleyite, 75, 228 heavy isotope, 194 hematite, 213 hemimorphite, 195, 200 hercynite, 43, 44, 49, 72 herzenbergite, 104 hessite, 32, 241, 242 hessite-petzite-gold, 242 hessite-petzite-krennerite, 242 hetaerolite, 205 hexagonal outline, 158 Hg, 243 hiding of a rare element, 264 high grade iron ores, 178, 179 higher free energy of oxidation, 261 hjelmite, 232 hollandite, 206, 207 hollingsworthite, 224 homogeneity of lead isotope abundance data, 193 homogeneous substance, 262 homogenization, 14 homogenization of crystal inclusions, 180 homogenization of detrital gold, 157 homogenization of gas-liquid inclusions, 180 homogenization of inclusions, 181 homogenization of liquid-gas inclusions, 12 homogenization of sulfur isotopes, 270 homogenization rates, 14 homogenization temperatures of fluid inclusions, 181 homologous, 220 homologous elements, 223 hornblende, 62, 179 hornblende schists, 178 hornblendic rocks, 177 hot brines, 163, 165 hübnerite, 32, 121, 205, 252 hiibnerite/ferberite ratio, 134 hutchinsonite, 81 Huttonian geology, 267 hybridism, 119 hybridization, 188 635

hydro-oxides, 36 hydrocarbon, 196 hydrocerussite, 255 hydrogels, 25 hydrogoethite-beudantite-plumbojarosite, 201 hydrogoethite-cerussite, 201 hydrogoethite-pyromorphite, 201 hydrohaematite, 90, 202 hydrosphere, 191 hydrothermal, 9, 118, 121, 124, 133, 139, 146, 158, 159, 182, 192, 213, 250, 269 hydrothermal activity, 186 hydrothermal alteration hypothesis, 179 hydrothermal conditions, 124 hydrothermal emanations, 122 hydrothermal exudation, 168 hydrothermal in zones of interlayer brecciation, 145 hydrothermal mineralization, 139 hydrothermal origin, 159 hydrothermal paragenesis, 234, 237 hydrothermal perigranitic metallogeny, 213 hydrothermal polymetallic fluorite, 137 hydrothermal replacement, 144, 150, 159 hydrothermal solutions, 110, 259 hydrothermal uranium paragenesis, 241 hydrothermal volcanogenic sphaericules, 92 hydrothermal W-vein mineralization, 252 hydrothermal-magmatic, 166 hydrothermal-metasomatic, 152 hydrotherms, 172 hydroxyfluorstannite complex of tin, 259 hydrozincite, 200 hypabyssal intrusives, 172 hypabyssal rock, 127 hypergene, 231 hyperite, 240 hypersaline brines, 151, 163, 165 hypersaline fluorine-bearing fluids, 159 hypersthene, 21 hypidioblastic, 64 hypogene, 3, 11, 126, 127, 128, 199, 200, 204, 205, 252, 255 hypogene anhydrite, 134 hypogene hydrothermal replacement, 205 hypogene manganese, 200, 254 hypogene mineralization, 191, 199, 200 hypogene replacement, 12 hypogene sulphides, 191 hypogene zonality, 214 hypogene zoning, 199, 214 hypogene-supergene mineralization, 200 hypomesothermal stage, 140 hypothermal, 12, 118, 124, 140, 213 hypothermal-hydrothermal phase, 137 hypothermal-mesothermal, 139 hypothesis of granitic magma, 119

I-type, 123 idaite, 145, 171, 201, 212 636

idioblasic gudmundite, 65 idioblast, 44, 52, 53, 65 idioblastic, 64 idioblastic form, 54 idioblastic growth, 65 idioblastic pyrite, 66 idioblastic tendencies, 29 idiocrase, 135, 197 idiomorphic apatites, 21 idiomorphic chromite, 20 idiomorphic quartz, 62 idrialite, 196 ignimbritic tuffs, 150 ilmenite, 24, 35, 42, 43, 44, 50, 53, 55, 66, 71, 72, 75, 81, 100, 102, 115, 116, 120, 122, 127, 136, 141, 202, 214, 224, 240, 249, 256 ilmenite lamellae, 25, 42, 43 ilmenite margins epitactic on magnetite, 72 ilmenite replacing magnetite, 24 ilmenite-haematite granular pattern, 82 ilvaite, 135 immediate source, 173 immiscibility, 189 immiscibility of melts, 112, 189 immiscible sulfide liquid, 113 immiscible sulphide-rich liquids, 113 In, 202 in the solid state during metamorphism, 64 inclusion, 74 inclusion brines, 163, 182 incompatibilities, 16, 21 incorporation and substitution of metallic elements in the silicate lattice, 267 incorporation of elements in a crystal structure, 264 increase in solubility in intergranular, 268 indium, 241 induced reduction, 35, 36 inductive, 12 inductive approach, 219 infiltration replacement, 52 initial clastic chromite sand layer, 153 initial earth crust, 6, 266, 267 initial source, 173 innate flexibility, 272 inorganic nature, 188 inter-rift buried basaltic flows, 6 interatomic distances, 262 interbanded gel phases, 81 interbanded pitchblende, 81 interbanded pyrite, chalcopyrite, uraninite structures, 81 interbanding of colloform mineral phases, 81 interbanding of lievrite and haematite, 81 interbanding of neodigenite and brown iron, 26 intercontinental basalts, 209, 210 intercontinental rift basalts, 187, 204, 254, 255 intercontinental rift deposits, 109 intercontinental rift type, 187, 237 intercumulus, 154 interfacial free energy, 23

intergrowths of ilmenite-haematite spinels, 48 interleptonic spaces, 97 intermediate discontinuity, 77 intermediate phases, 35 intermediate product, 100 interrelationship of elements, 272 interrelationship of elements according the empirical laws of the periodic system, 265 interrelationship of metallic elements, 268 interrelationship of the PGE, 224 interstitial solid solutions, 41 interzonal chalcopyrite, 46 interzonal penetrability directions, 46 interzonal replacement, 34 intracontinental rift basin, 153 intracrystalline infiltration, 83 intrusive complexes, 19 inversion point, 12 ionic radii, 262 ionic radii and crystal structure, 264 ionic substitution, 264 ionic substitution in minerals, 264 ionized radii state, 49 Ir-bearing sulphoarsenides, 224 iridian ruthenium, 224 iridium, 87 iridosmine, 158 iron hat, 168 iron hydro-oxides, 66 iron sulphides, 261 iron-bacteria, 177 iron-bearing sediments, 149 iron-formations and paleosols, 173 iron-leached bauxites, 103 iron-trend, 70 irreversible rupture, 94 isemannite, 230 ishikawaite, 232 island-arc evolution, 109 isomorphism, 262 isomorphous, 44 isomorphous crystallization, 53 isostructural, 44 isotherms concentric with the contact, 213 isotope analysis, 191 isotope data, 189 isotope fractionation, 191 isotope geochemistry, 1, 9, 144, 184, 185, 192, 217, 218, 219, 257 isotope pattern ratios, 191 isotope ratios, 68, 191, 200 isotope studies, 189 isotopes, 191, 194 isotopic composition, 146, 176, 180, 200 isotopic composition of lead, 193 isotopic composition of sulphur, 189 itabirites, 151 ixiolite, 250

J-type anomalous Pb-deposit, 193 J-type of Pb, 193 jacobsite, 40, 52, 53, 72, 204, 206, 255 jamesonite, 30, 31, 140, 141, 213 jasper, 174, 205, 254 jasperoids, 146 jaspilites, 174 jefferisite, 203 johannsenite, 137 joint segregation of less abundant elements in the earth's crust into concentration of deposits, 235 joint segregation of metallic elements, 272 joint segregation of rare elements in paragenetic associations, 264 joint segregation/(accretion), 227 Joplin-type leads, 193 jordanite, 186, 245 jordisite, 230 jordisite-ilsemannite, 147 joseite, 242, 252 juvenile simatic, 108 juvenile solutions, 165

K-feldspar, 55 kämmererite, 114 kappa-pyrite, 133 karst-mineralization, 197 karstic bauxites, 90 karstic mineralization, 196 keratophyre alumino-silicate formations, 174 kesterite, 104 key element Sn, 246 kidney forms, 79 kidney-shaped, 204 kimberlites, 21 kimberlitic pipes, 266 kinetics, 4, 7 klockmannite, 232 kobellite, 241 komatitic, 174 Kontinentalverschiebung und Erzprovinzen, 110 kotoite, 135, 138 kotulskite, 224 krennerite, 241 Kristallisationsfreudigkeit, 69, 79, 80, 81 Kupferschiefer, 148, 163, 208 kyanite, 150 Lake Superior-type, 175 lamellae of ilmenite, 50 lamellar bodies of haematite, 50 lamellar dispersive system, 77 lamellar haematite, 48 lamellar marcasite, 23 Landsat lineaments, 107 lanthanides, 220 lantite, 211 large scale replacement, 167 637

late-kinematic granites, 125 later pitchblende generation, 79 lateral and vertical zoning, 214 lateral leaching, 4 lateral secretion, 3, 4, 38, 152, 167, 168 lateral secretion processes, 271 lateral segregation, 39, 154, 167, 168, 230, 232, 269 lateric conditions, 117 latente, 173 lateritic conditions, 226 lateritic cover, 86, 226, 270 lateritic paragenetic associations, 224 lateritic soil, 86 lateritic weathering model, 173 lateritization, 86, 87, 226, 269 laterization process, 86 lattice intergrowth, 13 lattice of pyrrhotite, pentlandite, chalcopyrite, 270 laumontite, 136 laws of element distribution, 218, 219, 238, 261, 264 laws of geochemical distribution of elements, 264 layered bodies, 18 layered complexes, 153, 155 layered lead-zinc deposits, 192 layered pegmatoid, 112 layered sill, 60 layered structure of chromite-bearing zones, 112 layered ultrametamorphic bodies, 20 layering of the earth's crust, 17 leaching, 4, 11, 29, 31, 33, 52, 54, 85, 98, 101, 102, 104, 109, 152, 163, 194, 197, 199, 201, 204, 208, 209, 210, 213, 255, 269 leaching and diffusion, 103, 208, 210 leaching experiments, 209 leaching of elements, 98, 103 leaching of iron-bearing detritus, 210 leaching of Sn, 104 leaching out of Fe, 26, 38, 90, 103 leaching out of Mg, 99 leaching out of W, 100 lead, 140, 151, 159 lead isotope studies, 193 lead isotopes, 151, 193 lead isotopes in synaeresis cracks, 268 lead mobilization in synaeresis cracks, 268 lead-bismuth-selenium sulphosalt, 200 lead-copper-silver, 193 lead-zinc deposit, 270 lead-zinc ore, 152, 161, 195, 270 lead-zinc-silver deposits, 144, 200, 202 lead/zinc ratios, 147 leaf silver, 202 Leasengang rings, 103 lenticular chromite, 18 leopard chromite, 20, 269 leopard ore, 18 leopard structures in bauxites, 103 lepidocrocite, 82, 100, 101, 203, 230 lepidocrocite crystallites, 82 638

leptothermal, 252 leuchtenbergite, 44, 66, 114 leuchtenbergite schist, 49 leuco-quartz monzonite, 121 leucocratic type of mineralization, 110 leurite, 224 lherzolitic, 19, 107 Li, Be, B?, Al, 251 liebigite, 137 lievrite, 35, 101 lievrite-haematite, 81 lievrite-hedenbergite, 134 lievrite-hedenbergite skarns, 129 light coloured bands, 79 limits of elasticity, 94 limits of plastic deformation, 94 limonite, 35, 99, 101, 200, 226, 232 limonite-hydrogoethite, 201 limonitic iron, 103 linarite, 200 lineaments, 107, 152 linear correlation analysis, 161 linear trend extending from mantle to upper crust, 194 linnaeite, 58, 139, 141, 157 liquid immiscibility, 133 liquid inclusion, 181 liquid magmatic segregation, 134 liquid separation on release of pressure, 156 liquid unmixing, 154 liquid unmixing by pressure release, 155 liquid-gas inclusions, 11 lithic crust, 20, 155 lithic era, 2, 18, 20 lithiophorite, 196, 205 lithium, 264 lithium minerals, 250 lithophile, 3, 5, 6, 266 lithophile elements: Y, Zr, (Hf), (lanthanides), Nb, Ta, Mo, W, U, (?Ti), 266 lithophile factor, 122 lithophile groups, 261 lithophile-granitophile, 119 lithophilic, 6, 39 lithosphere, 6 lognormal frequency distribution, 158 löllingite, 116, 136, 137, 140, 199, 200, 201, 209, 243, 250 low Cu/Zn ratios, 159 low S content for the upper lithosphere silicate rocks, 189 low Sc, Th, Hf, Ta and Cr values, 160 low-salinity C0 2 -rich fluids, 185 lower crust/crust group (Sn, As, Bi, Pb, Sb, Se, Te, Ga, Ge, In), 258 lower crust/upper mantle obduction, 143, 219, 265 ludwigite, 135, 138 lumina, 35 lunar samples, 23 lutite, 177 luzonite, 245

mackinawite, 115, 224 macro molecules, 78 maghemite, 24, 55, 56, 100, 101, 115, 203, 256 maghemitization, 100 magma -crust interactions, 267 magma chamber, 150 magma immiscibility, 188, 189 magma reservoirs, 107 magma/metallogeny, 112 magmagene-ore systems, 268 magmatic anomaly, 109 magmatic differentiation, 17, 21, 97, 112, 113, 114, 117, 153, 155, 189, 270 magmatic differentiation/fractional crystallization, 153 magmatic emanations, 122 magmatic gravitation-differentiation, 18, 95 magmatic hydrothermal origin, 191, 193 magmatic immiscibility, 113 magmatic origin, 193 magmatism, 265 magmato-endogene ore deposits, 110 magmatogene ore genesis, 111 magmatogenic derivations, 7 magmatogenic origin, 192 magnesian Sn(-W) skarn, 130 magnesioferrite sphaeroids, 91 magnesioskarn, 135 magnesite, 4, 179, 271 magnesium skarn, 130, 131 magnetic attraction, 91 magnetite, 20, 21, 23, 24, 35, 36, 37, 38, 42, 43, 44, 50, 53, 55, 56, 59, 60, 61, 64, 66, 71, 72, 75, 96, 99, 100, 101, 115, 116, 117, 119, 133, 134, 136, 138, 141, 149, 162, 174, 175, 176, 177, 178, 179, 201, 202, 203, 213, 214, 224, 243, 246, 248, 249, 266 magnetite band, 21 magnetite crystals, 177 magnetite epitactically overgrown, 90 magnetite octahedra, 66 magnetite octahedra metacrysts, 64 magnetite relics, 66 magnetite replaced by pyrite, 24 magnetite seams, 115 magnetite-forming fluids, 176 magnetite-graphite assemblage, 179 magnetite-nickeliferous pyrite-millerite, 203 magnetite-pyrite-chalcopyrite, 172 magnetite-pyrrhotite-pyrite, 134 magnetite-titanospinel, 114 magnetitic, 179 magnetitic margin, 21 majakite, 224 malachite, 27, 37, 38, 60, 69, 83, 116, 128, 140, 145, 200, 201, 202 malacon, 69, 72, 75, 83, 96, 233 malacon epitactic, 72 malacon included in uraninite, 75 malakton, 118, 132, 133, 137 malayaite, 136

maldonite, 243 malnikovite, 149 Maltese cross structure, 39 mammilary, 79 mammilary-botryoidal, 79 manganese, 4, 176, 187, 210, 216, 254 manganese deposits, 109, 254 manganese garnet, 159 manganese leaching, 209 manganese mineral deposits, 170 manganese minerals, 81, 204 manganese nodules, 85, 90 manganese ores, 211 manganese paragenesis, 39, 254 manganiferous garnet, 205 manganiferous sediments, 254 manganite, 80, 204, 205, 207, 205, 255, 256 manganoan calcite, 205, 255 manganoan dolomite, 205, 255 manganoan skarns, 137 manganotantalite, 187 mantle, 107, 108, 110, 114, 116, 153, 154, 155, 169, 189, 190, 204, 215, 219, 224, 227, 241, 242, 250, 258, 259, 267, 269 mantle acting as impermeable barrier, 266 mantle derivation, 246, 269 mantle derivative, 253, 254, 255, 266 mantle derived elements,: Fe (Μη), Ni, Co, Cu, Zn, Cd, Ag, 266 mantle derived Pb, 193 mantle diapir, 153, 156 mantle diapirism, 3, 143, 215, 219, 265 mantle environment, 5 mantle fragments, 266 mantle geoenvironment, 266 mantle gneiss domes, 146 mantle metasomatism, 159, 160 mantle obduction, 7, 18 mantle patterns, 19 mantle petrography, 18 mantle plume, 194 mantle sulphur, 189 mantle-group Ti, V, Cr, Μη, 256 mantle-like source, 194 mantle-mobilization, 107 mantle-Pb, 189 manto copper deposits, 147 manto ore (stratabound pyrite-chalcopyrite ore), 271 mantos, 195, 255 marble, 56, 94, 114, 167 marble replacement, 135 marcasite, 29, 31, 35, 36, 53, 65, 66, 100, 126, 127, 136, 140, 141, 142, 145, 147, 149, 162, 171, 199, 202, 230, 244, 248, 252 martite, 100, 101, 115, 202, 203 martitization, 35, 100 martitized, 89 mass-replacement, 195, 197 massive chromite, 266 639

massive gonditic ores, 204 massive ore, 204 massive pyrite deposits, 170, 197 massive replacement, 59 massive sulphide bodies, 170 massive sulphide deposits, 194 matildite, 197, 200 maucherite, 60, 116, 230 maximum possible increase of specific surface, 78 maximum possible reduction of specific surface, 78 mechanical injection, 114 mechanism of transportation, 219 mega-colloform structures, 134 meionite-rich scapolite, 148 melange type, 94 melanterite, 202 melnikovite, 244 melnikovite-pyrite, 145 melting by friction, 127 melts, 265 meneghinite, 140, 202 mercury, 110, 120, 142, 196, 209, 244 mercury mineralization, 243 mercury-antimony stage, 199 Merensky Reef, 112, 116, 153, 154, 270 merteite, 224 meso-epithermal, 141 mesocratic mineralization, 110 mesohypothermal stage, 140 mesothermal, 9, 12, 118, 124, 140, 141, 162, 181, 213, 252 mesothermal Au systems, 185 metabolic products of sulphur bacteria, 188 metacinnabar, 70, 83, 196, 244 metacinnabar sphaeroid, 83 metahydrothermal fluids, 209 metal zoning, 191, 200 metalliferous fluids, 268 metallogenic belts, 110 metallogenic processes, 107 metallogenic zones and belts, 268 metallogeny, 108, 109, 110, 114, 118, 120, 125, 126, 127, 143, 149, 163, 164, 169, 182, 183, 189, 194, 197 metallogeny of skarns, 129, 133 metallogeny related to granitization, 119 metallogeny/differentiation, 111 metallography, 88 metamictic euxenite, 102 metamorphic, 204 metamorphic effect, 129 metamorphic fluids, 152 metamorphic sedimentary, 159 metamorphic sedimentary origin, 159 metamorphic zoning, 120 metamorphic-metasomatic derivation, 155 metamorphic-metasomatic differentiation, 123, 213 metamorphic-metasomatic hypothesis, 155 metamorphics/granitization, 183 metamorphism, 178 metamorphism-granitization, 6 640

metamorphism-ultrametamorphism, 123 metamorphogenic, 183 metamorphogenic ore deposits, 183 metapelites, 178 metasedimentary-metamorphic, 21 metasomatic changes, 178 metasomatic differentiation, 133, 137, 188, 189, 268 metasomatic effect, 129 metasomatic filling, 45 metasomatic mobilization, 64, 100 metasomatic processes, 178 metasomatic replacement, 15, 118 metasomatic zonation, 131, 132 metasomatism, 67 metasomatism of calcareous wall rock, 134 metasomatism-replacement, 12, 23 metasomatized mantle, 160 metasome, 12, 13 meteoric, 192 meteoric fluids, 185 meteoric water, 185 Mg leaching, 85 Mg metasomatism, 196 Mg skarn, 130 Mg-skarn types, 129 MgO/FeO ratios, 114 miargyrite, 137, 171 mica, 62, 121 microbiota, 174, 177 microcline, 53, 106 microcolonies, 177 microcrystalline aggregates, 82 microdessication cracks, 177 microenvironment, 2 microfossils, 177 microgranite, 103 micrographic, 53, 54, 57 micrographic-like intergrowth, 57 microlaminated hematite, 174 micromylonitization, 94, 168 microprobe, 218 microprobe analysis, 219, 226 microprobe scanning, 228 microrhythmic structure, 205 microstructures, 177 microtectonic line in the davidite, 50 microtectonics, 96 migmatization, 139 migrated by diffusion, 168 migration of cations, 29 migration of elements, 29, 37, 38 migration of metal-bearing fluid, 216 migration of pore fluids, 212 millerite, 28, 34, 213 mimesis, 187 mineral assemblage (MAG), 8 mineral association (MAT), 8, 272 mineral forming solutions, 180 mineral paragenesis, 8, 145

mineral paragenetic associations, 267 mineral zones, 213 mineralization associated with acid magmatism, 108 mineralization of Mo, W, Sn and U, 121 mineralizing event coeval with deformation, 179 mineralizing fluids, 10, 11 minnesotaite, 175, 179 miogeosyncline, 138 mispickel, 134, 139 Mississippi Valley, 164 Mississippi Valley lead-zinc deposits, 180, 182 Mississippi Valley-type ore, 165, 195, 211 mixed fluids, 165 mixed liquids under great pressure separate, 154 mixing colloid solutions, 78 mixite, 201 mixture of molybdenite-pitchblende, 231 Mn, 215 Mn and phosphorite, 172 Mn in magnetite, 170 Mn-actinolite, 137 Mn-hedenbergite, 137 Mn-ilvaite, 137 γ-Μη0 2 , 196 Mn/Zn ratio, 164 Mo, 119, 121 Mo in pitchblende paragenesis, 230 Mo-Bi bearing quartz pipes, 121 Mo-mineralization, 135 Mo-vein mineralization, 252 mobilization, 4, 5, 6, 7, 20, 27, 33, 34, 36, 38, 49, 61, 84, 96, 103, 107, 109, 115, 116, 119, 124, 127, 133,134, 141, 146, 148, 149, 154, 155, 161, 167, 169, 171, 172, 179, 184, 185, 189, 193, 199, 202, 203, 204, 206, 207, 211, 216, 256, 265, 266, 268, 271 mobilization of boron (tourmaline), 268 mobilized, 20, 123 modified placer, 158 molecular disperse system, 77 molecular dispersed, 78 molten iron, 261 molybdenite, 28, 29, 94, 121, 128, 138, 141, 146, 147, 157, 181, 212, 214, 230, 248, 249, 251 molybdenite pipes, 121 molybdenite with Pb, Bi and Te-bearing minerals, 121 molybdenum, 110, 120, 142, 151, 271 molybdenum bearing granite, 124 molybdenum deposit, 121 molybdenum in brown coal, 147 molybdenum minerals, 231 monazite, 122, 125, 233, 251 monoascendent, 214 monotropic, 12 montmorillonite, 172 monzonites, 126, 247 Mount Isa controversy, 161 mud sulphur, 192 multigeneration deposits, 11 multigenetic concepts, 137

multimetallic mineralizations, 11 multiphase, multi-stage crystallizations, 180 multiphase inclusions, 180 multiple epitaxis, 71 multiple ex-solution phases, 42 multiple folding, 179 multiple magmatic intrusions, 141 multiple migrations, 149 multiple recycling, 212, 216 multiple replacement, 12, 29, 37, 38, 39, 45, 46, 47 multiple supplies, 10 multiple unmixing, 42 multisource derivation, 250 multisource origin, 152 multistage crystallization, 181 multistage mineralization, 199 multistage mobilization, 134 multistage processes, 192 muscovite, 121 mutual replacement, 34, 38 mutual replacement of atoms, 262 mutual solid solutions, 10, 46, 56 mylonization, 95 myrmekitic form, 57 myrmekitic intergrowth, 21 myrmekitic serpentine, 98 myrmekitic symplectic intergrowth, 56, 58 myrmekitic symplectic serpentine, 61 myrmekitic symplectic silicates, 62 myrmekitic symplectic spinel, 72 myrmekitic symplectite, 57, 58, 61 myrmekitic texture, 55 myrmekitic-like bodies, 57 myrmekitic-like forms, 15, 62 myrmekitic-like intergrowth, 23, 30 myrmekitic-like patterns of rutile, 25 myrmekitic-like pyrite, 24 myrmekitic-like symplectite, 54, 60 myrmekitoid pyrite, 59 myrmekitoid relic structure, 60

Na-Ca-K-chloride brine, 163, 182 Na/K ratios, 165, 182 Na/Li ratios, 182 nasturan, 137 native antimony, 189 native As, 243 native Bi, 33, 37, 38, 96, 101, 244 native bismuth, 121, 201 native copper, 37, 38, 50, 148, 200, 201, 202, 255 native elements, 199 native gold, 31, 37, 200 native lead, 255 native mercury, 209 native Se, 232 native silver, 31, 50, 82, 146, 171, 201, 202 naumannite, 137 Nb, 186 641

Nb and Sc contents in cassiterite, 247 Nb and Ta of pegmatitic paragenesis, 234 Nb-Sb type, 122 needle iron, 82 nematoblastic, 64 neo-crystallization, 26, 27 neo-crystallizations of chalcocite, 26 neo-neptunian "dykes", 145, 211 neodigenite, 26, 27, 28, 30, 54, 61, 62, 116, 200 Ni, 186, 215 Ni-Co arsenide/sulphide mineralization, 148 Ni-Cr lateric iron deposits, 82 Ni/Cu ratio, 112 niccolite, 27, 31, 32, 33, 36, 60, 66, 80, 83, 126, 230 nickel, 89, 116, 261 nickel ore, 112 nickeliferous, 116 nickmanite, 187, 237 niobian rutile, 125 niobite, 52, 250 noble gas isotopic geochemistry, 176 noble gases, 176 nodule (nuggets) formation, 168 nodule formation, 167 non-distortion plane, 185 non-ferous metal mineralization, 245 non-uniformitarianism, 18 nontronite, 201 noritic stage of evolution, 188 Nsuta manganese, 196, 197 nuggets, 4, 86, 87, 89, 202, 227, 269 nuggets of platinum, 90

obducted ocean flow, 109 obducted upper mantle, 5, 19 obduction, 17 ocean spreading, 17 ochre, 201 oil drops, 92 olivenite, 200, 201 olivine, 50, 69, 74, 76, 103, 108, 117, 150 olivine bombs, 18, 19, 21, 61, 107, 238 oolites, 82, 90, 103 oolitic ironstones, 175 oolitic pisolitic structures, 103 oolitic pisolitic textures, 85 opal, 141 open system, 134 ordered state, 14 ordinal number, 220 ore emplacement, 161 ore genesis as the state of art, 218 ore mineral gangue symplectites, 60 ore mineralization, 109 ore mobilization-remobilization, 268 ore-forming fluid, 132 organic material, 152 organic matter, 35, 92, 93, 178

642

organic plants, 91 organisms, 177, 188 organogenic (bacterial action), 91 organogenic sphaeroids, 91 oriented alteration patterns, 100 oriented symplectites, 52 orogenic, 110 orogenic belts, 248 orpiment, 29, 126, 186, 196, 243, 252 orthite-apatite enriched rock, 138 orthodox magmatic concepts, 111 orthomagmatic crystallization, 129, 213 orthomagmatic deposits, 12 orthomagmatic phase, 118, 167 orthopyroxenes, 178 orthothermal, 124 osmiridium, 86, 88, 89, 224, 226 osmiridium crystals, 74 osmium, 87 osmotic filtration, 163, 165 ottemannite, 104 overgrowth, 71, 74 overgrowths and epitactic growths, 74 overlapping crystallization, 9, 36, 40 owyheeite, 30 oxidation, 205 oxidation mineralization, 199 oxidation potential of magma, 112 oxidation zone, 202, 231 oxidation zone mineralization, 201 oxides of quadrivalent elements, 263 oxygen fugacity, 238 oxygenated atmosphere, 175

palaeo-atmospheric, 175, 176 palaeo-atmospheric composition, 176 palaeo-climatic maps, 209 palaeo-karst-type deposits, 195 palaeo-karsts, 198 palaeoslope, 178 palaeosoma, 154 palaeoweathering, 174 palingenesis of granitic melts, 125 palingenetic, 4, 18, 34 palingenetic processes, 211 palingenic, 212 palingenic origin, 232 palingenic remobilization, 97 palingenic veinlets of cassiterite, 268 palladium, 89 pallado-arsenide, 224 panidioblastic, 64 para-gabbros, 64 para-norites, 64 parablastic, 64 parablastic crystalloblastesis paragabbros, 20

pangenesis, 1, 2, 4, 6, 8, 9, 11, 39, 40, 50, 53, 78, 110, 118, 125, 140, 141, 149, 168, 169, 171, 185, 197, 216, 219, 224, 228, 229, 230, 243, 245, 252, 272 pangenesis of hypogene minerals, 199 paragenetic association, 1, 10, 36, 39, 50, 61, 95, 114, 116, 229, 239, 240, 242, 246 paragenetic association of Mo and Rh, 231 paragenetic formations, 39, 109 paragenetic interrelationship, 76 paragenetic mineral association, 241, 257 paranorites, 20 pararammelsbergite, 83, 230 parental basaltic magma, 18 parental geoenvironment, 5, 6 parental mantle geoenvironment, 6 partial dissolution, 9 passages in metallogeny, 107 Pb, 146, 152, 184 Pb isotopic and mineralogical studies, 133 Pb-As sulphosalts, 186 Pb-Bi-S system, 121 Pb-mantle derived, 194 206pb/204pb

ratios>

191)

1 9 2

Pb-Sn sulphide, 121 Pb-Zn (Ag) deposits, 137 Pb-Zn deposit, 166, 192, 197, 200, 213 Pb-Zn hydrothermal type, 137 Pb-Zn mineralization, 153, 196 Pb-Zn ores, 146, 161, 212 Pb-Zn skarn, 133, 137 Pb-Zn-Ag deposits, 159, 160, 195 Pb-Zn-Ag mineralization, 124, 126, 171 Pb-Zn-Ag ores, 150, 197 Pb-Zn-Ag veins, 184, 185 Pb-Zn-Cu deposits, 151 Pb-Zn-Cu ore, 146 peat as a collector, 216 pegmatite, 32, 92, 123 pegmatitic, 3, 12, 124, 139 pegmatitic crystallization phase, 118 pegmatitic paragenesis, 229 pegmatitic pneumatolytic, 250 pegmatitic U-paragenesis characterized by U, Th, Y, Zr, Hf, Nb, Ta, Ti, rare earth, 232 pegmatitic uraninite, 50 pegmatitic uranium paragenesis, 229, 232, 233, 234 pegmatitic-granitic geoenvironment, 253 pegmatoid bodies, 115 pelagic muds, 174 penetrability, 105 penetrability direction, 45, 105 pentlandite, 24, 26, 47, 48, 76, 115,116, 189, 213, 224, 270 pentlandite flame-shaped, 48 pentlandite flames, 48 pentlandite rims, 47 peridotite magma, 112 perigranitic, 2, 6 perigranitic (around the granite) S mobilization, 189 perigranitic (hydrothermal) vein deposits, 189

perigranitic metallogeny, 6, 119, 188, 189 perigranitic metamorphic-metasomatic mobilization, 118 perigranitic metamorphism, 133 perigranitic metasomatism, 123 perigranitic mineralization, 118, 123,213 perigranitic vein mineralization, 137 perigranitic zonal metallogeny, 213 period of mineralization, 229 periodic function, 220 periodic system, 220 periodic system as system of element classification, 272 periodic table, 14, 115, 272 periodical compressions and tensions, 267 periodicity as a function of the repetition of homologies in the structure of atoms, 220 perthites due to K-feldspar substitution, 167 perwoskite, 99,100 petrofabric analysis, 96, 239 petrofabrics, 21, 96 petroleum fluid, 163, 166 petzite, 31, 32, 202, 241, 242 petzite/quartz, 53 PGE, 266, 269 PGE and PGM, 86 PGM, 154 PGM in lateritic covers, 6 phase, 229 phase rule, 218 phase successions, 139 phenocrysts of alkali feldspar, 127 phlogopite, 129 phosphatic ooids, 90 phosphorites, 90 photo-oxidized by sunlight, 176 photo-precipitation, 176 photo-process, 176 photosynthesis, 177 photosynthesizing organisms, 216 photosynthetic, 175 phreatic, 7 phreatic cycle concept, 185 phreatic cycles, 3, 7, 98 phyllites, 184 pickeringite, 192, 200 picotite, 114 picrochromite, 239 piemontite, 205, 254 pigment-size sphalerite relic, 67 pillow lavas, 170 pink dolomite, 180 pipe-like bodies, 115 pisolite, 82, 90 pisolitic relics, 90 pisolitic textures, 149 pisolitic-oolitic lateritic, 75 pitchblende, 34, 35, 39, 56, 80, 84, 96, 97, 102, 137, 138, 151,211,212, 230, 231,232 pitchblende in sedimentary rocks, 147 pitchblende microsphaeroid, 79 643

pitchblende pangenesis, 78, 229, 230, 232 pitchblende sphaeroid, 34, 35, 102 placer origin, 211 plagioclase, 55 plant remains, 92 plastic deformation, 94, 95 plastic mobility of galena, 95 plastic tectonic mobilization, 95 plastically, 5 plastically mobilized galena, 95 plasticity, 4 plate junctions, 7 plate tectonic hypothesis, 7, 108, 109, 143 plate tectonic theory, 17, 110, 130 plate tectonics, 7, 107, 108, 144 plate tectonics junctions, 109 plateau basaltic flows, 143 platinoid element segregation to form Pt nuggets, 87 platinoid metals, 116 platinoid minerals, 226 platinum, 85, 87, 117, 261 platinum elements, 69 platinum group, 88 platinum group of elements, 5 platinum nuggets, 85, 86, 87, 153 platinum nuggets controversy, 144 plattnerite, 81 pleonaste, 19, 20 pleonaste-ilmenite lamellae, 60 plexus of processes, 204, 206, 216 plumbojarosite, 200, 202 pneumatolytic, 134, 139, 168, 269 pneumatolytic deposition, 136 pneumatolytic metamorphic bodies, 132 pneumatolytic phase, 119, 136, 251 pneumatolytic replacement, 136 pneumatolytic stage, 159 pneumatolytic-hydrothermal, 224, 248 pneumatolytic-hydrothermal deposits, 9, 122, 134 pneumatolytic-hydrothermal genesis, 139, 148 pneumatolytic-hydrothermal type, 139 pneumatolytic-katathermal transitional stage, 135 pneumatolytic-metasomatic phase, 118, 124 pneumatolytic-pegmatitic phase, 213 podiform chromite, 114 poikiloblastic pyrite, 66 polarization forces, 264 polianite, 204, 205 polyascendent Ag-Pb-Zn-Sb mineralization, 214 polybasite, 171, 200, 230, 241, 243, 252 polygenetic, 145, 158 polymetallic, 140, 168, 171 polymetallic character, 142 polymetallic deposits, 108, 109, 134, 137, 149 polymetallic hydrothermal ore, 181 polymetallic mineralization, 142 polymetallic mineralization with Zn, Cu, Sn and W, 130 polymetallic ore veins, 214 polymetallic ores, 209 644

polymetallic skarns, 130 polymetallic veins, 184 polymetamorphism, 11, 40 polymignite, 232 polymorphism, 39, 262 polyphase mobilization, 4 polysynthetic twinning, 105 pore fluids, 152 pore space, 69 porphyrite, 172 porphyrite lapis, 127 porphyritic schists, 184 porphyroblastic growth, 67, 270 porphyroblasts, 149 porphyry, 3, 108, 120, 126, 127, 139, 150, 184, 197, 202 porphyry ores, 7 powellite, 231 Precambrian iron formation, 172, 216 precursor magnetites, 177 preferential leaching, 209, 210 preferential replacement, 13 preferential substitution, 25 pressure conditions, 180 primary zoning, 213 primordial atmosphere, 191, 261 primordial layering, 266 principle lode-tin deposits of the world, 121, 248 prismatic silicates, 54 prograde, 11 prograde and retrograde scheelite, 135 prograde metamorphism, 134 Proterozoic, 178 Proterozoic atmosphere, 175 proto-ilmenite, 115, 240 protoliths, 102 protolyte, 267 protolytic, 5, 6 protolytic layer, 18, 21, 169 protolytic type, 177 protore, 176 prototectonic, 269 prototectonics, 18, 153, 155 protracted hydrothermal activity, 161 proustite, 31, 32, 54, 157, 200, 209 proximal Sn, Be-F-W skarns, 129 proximal turbidites, 174 proxy solid solution, 41 pseudo-eutectic intergrowth, 60 pseudo-eutectic textures, 13 pseudo-neptunian, 145, 211 pseudo-unmixing, 13 pseudo-veins, 196 pseudohydrothermal activity, 142, 157 pseudohydrothermal shaping, 158 pseudomorphs, 35, 101 pseudoporphyritic monzogranite, 135 psilomelane, 101, 170, 196, 205, 206 psilomelane sphaeroid, 82 Pt, Pd, Os and Rh enriched in skarn copper deposits, 138

Pt deposits, 155 Pt group minerals, 87, 224 Pt minerals, 112 Pt nugget formation, 226 Pt nuggets, 86, 87, 88, 143, 226 ptygmatic folding, 94, 95, 97 pulsating character, 140, 181 pulsations of temperature, 180 pyralspite, 178 pyrargyrite, 25, 31, 50, 51, 54, 62, 121, 171, 230, 242 pyrite, 24, 26, 27, 28, 29, 30, 31, 32, 35, 36, 37, 38, 42, 53, 54, 56, 59, 62, 66, 69, 72, 95, 96, 116, 119, 120, 121, 122, 126, 127, 128, 133, 134, 135, 136, 139, 140, 141, 142, 145, 146, 147, 148, 148, 150, 151, 161, 162, 166, 168, 170, 171, 172, 175, 179, 186, 189, 192, 195, 196,197, 199, 200, 202, 213, 230, 232, 242, 243, 245, 246, 248, 249, 250, 255, 257, 258 pyrite as a substitute, 24 pyrite bodies, 197 pyrite containing Fe, Cu, Au, Ag, 257 pyrite crystalloblastesis, 66 pyrite epitactic on sphalerite, 72 pyrite idioblast, 24 pyrite interzonal replacement of sphalerite, 47 pyrite polymetallic deposits, 145 pyrite sphaericule, 91 pyrite Staub, 47 pyrite subidioblast, 67 pyrite-chalcopyrite, 197, 209 pyrite-chalcopyrite-sphalerite, 197 pyrite-haematite-barite deposits, 184 pyrite-pyrrhotite, 121 pyrite-rich axial plane, 211 pyrite-sphalerite-(chalcopyrite), 67 pyrite-sphalerite-galena-tetrahedrite-albandite sequence, 141 pyritic conglomerates, 158 pyritic shale horizons, 175 pyritic slates, 173, 216 pyrochlore, 185 pyrochlore-microlite series, 232 pyrochroite, 255 pyroclastic deposits, 197 pyrolusite, 170, 196, 197, 204, 205, 206, 207, 255, 256 pyrometasomatic, 9, 56, 124, 139, 148 pyrometasomatic deposits, 12, 134, 135, 136 pyrometasomatic metallogeny, 129 pyrometasomatic mineralization, 137 pyrometasomatic origin, 136 pyrometasomatic-pneumatolytic, 118 pyromorphite, 200 pyrothermal conditions, 56 pyroxene, 178 pyroxene pegmatoid, 154 pyroxene-garnet skarn, 131 pyroxenite, 19, 20 pyroxferroite, 178 pyroxmangite, 137 pyrrhotine, 195, 202

pyrrhotite, 23, 24, 25, 30, 35, 47, 48, 50, 53, 56, 58, 60, 62, 66, 72, 76, 100, 115, 116, 120, 121, 122, 126, 127, 133, 135, 136, 140, 141, 142, 145, 151, 157, 161, 162, 171, 172, 179, 181, 189, 192, 195, 197, 200, 213, 214, 224, 243, 248, 249, 252 pyrrhotite disease, 67 pyrrhotite-chalcopyrite-iron sphalerite in galena, 121 pyrrhotite-marcasite-magnetite, 135 pyrrhotite-rich tension gash veins, 211

quartz, 37, 38, 39, 53, 62, 67, 94, 103, 121, 135, 178, 182, 195, 200 quartz being replaced by manganese-rich ores, 197 quartz diorite porphyry, 127 quartz epitaxis, 71 quartz monzonite, 135 quartz monzonite porphyry, 126, 127 quartz porphyry, 127, 197, 199 quartz porphyry monzonite, 126 quartz-magnetite-carbonates, 179 quartz-molybdenite, 214 quartz-specular haematite magnetite rock, 149 quartz-topaz, 251 quartzite, 178 quick-silver, 260

radially-fibroblastic, 64 radioactive cracks, 75 radioactive haloes, 105, 106 radioactive polymerization of fluid hydrocarbons, 92 radioaktive Sprengungen, 75 radiogenic, 4 radiogenic lead, 84, 212, 270 rain of sulfide drops, 215 rammelsbergite, 31, 65, 83, 96, 137, 199, 230 ramsdellite, 85, 187, 197, 205, 237 range of dispersion, 191 rare earth elements, 264 rare elements in pegmatites, 168 rates of unmixing, 42 ratio Cu/Fe, 214 ratio of sulfur, 11 ratios of S/Se, Zn/Mn, Zn/Cd, Zn/Fe, 257 reactive, 13 reactive mineral phase, 187, 237 reactive paragenesis, 3, 186 reactive phase, 187 reactiveness, i readiness to crystallize, 79 realgar, 29, 126, 141, 186, 196, 200, 209, 243, 252 recipient geoenvironment, 6, 61, 62, 169, 215, 216, 217 recrystallization, 148, 149, 159 recrystallization-mobilization, 184 recrystallized native gold grain, 158 recycling, 3, 4, 5, 6, 7, 107, 125, 194, 245 recycling of elements, 36, 216, 219 recycling of elements in the geological cycle, 265 645

red-pale coloured cassiterites containing Ta and possibly Nb, 246 redistribution, 211 redistribution of Ag relative to Au, 211 redistribution of elements, 227, 266, 268, 271 redistribution of Pt group elements, 269 reduced W-type skarn, 130 reduction of sulfur, 11 rehealed fracture, 45 relationship crystal structure-composition, 268 relative size of ions, 15 relic gel structures, 80 relics of colloform structure, 66 relics of ilmenite, 24 remobilization, 3, 4, 5, 6, 7, 8, 33, 36, 96, 97, 107, 114, 116, 119, 120, 123, 124, 125, 133, 139, 149, 151, 152, 184, 185, 186, 193, 202, 204, 206, 207, 211, 212, 219, 227, 265, 266, 268, 269, 270, 271 remobilization mainly of As in hydrothermal solution, 186 remobilization of cassiterite, 97 remobilized pitchblende, 96, 97 renierite, 27, 28, 200 reniform, 142 reorganization of the lattice, 27 repeated mineralization, 11 repetition of analogies in atomic structure, 234 replacement, 2, 9, 12, 13, 15, 23, 24, 26, 27, 29, 30, 31, 33, 34, 35, 36, 41, 42, 44, 46, 47, 48, 50, 52, 53, 57, 59, 64, 65, 67, 79, 89, 105, 126, 158, 161, 173, 180, 195, 199, 201, 250, 254, 255, 272 replacement-ex-solution replacement bodies, 196 replacement by copper minerals, 25 replacement deposit in limestone, 196 replacement minerals, 204 replacement of feldspar by graphic quartz, 71 replacement of lavas, 197 replacement of limestones, 195 replacement of magnetite by chalcopyrite, 25 replacement of silicates, 89 replacement of sphalerite by galena, 56 replacement of tennantite by chalcocite, 26 replacement of titanomagnetite, 24 replacement pattern, 11 replacement patterns of cassiterite, 32 replacement patterns of magnetite, 24 replacement patterns of pyrite, 28 replacement patterns of sphalerite, 29 replacement process, 10 replacement textures, 71 replacement-diffusion, 79 replacement-substitution, 15 replacements involving W minerals, 32 replacements of the volcanics, 197 residual liquid injection, 114 resorption of the oolitic phase, 90 resurgence, 11, 144 retgersite, 202 retrograde, 11 646

retrograde alteration, 132, 136 retrograde alteration of skarn, 135 retrograde changes, 60 reversibility of gel to sol, 79, 80 rhenium, 209 rhenium minerals, 209 rhenium sulfide, 231 rhexistasy, 92 rhodium, 89 rhodochrosite, 195, 204, 205, 254, 255 rhodonite, 137, 195, 204, 205, 206, 255 rhondigite formation, 99, 100 rhyolite, 23, 103, 170, 172, 206, 214, 243 rhyolitic, 248 rhythmical banding and interbanding of gels, 81 rhythmical replacement, 26 rhythmically interbanded lievrite, 81 rhythmite, 148 rift basin, 179 rift faults, 185 rift system, 109 ring and feather structures, 55 ring bornite, 57 ring complex, 4, 85 ring granite complexes ring intrusion, 18, 19 ring iron structures, 103 ring structure, 33, 92 Rio Tinto type of deposits, 119 Rogenpyrit, 91 role of brines, 163, 165 role of C0 2 -rich fluids, 259 role of fluids in metamorphics, 184 role of sulfur, 10, 25 roseite, 86, 89, 226 rounded ilmenites, 21 rudite, 177 ruthenium, 87 rutheriridosmine, 224 rutile, 44, 50, 54, 56, 72, 74, 100, 102, 115, 127, 200, 212, 243 rutile ex-solutions, 239 rutile lamellae, 50 rutile structure, 263 rutile-silicate symplectite, 25

S impregnations, 190 S leaching, 100 S-granites, 18 S-type, 123 S-type of granites, 122, 139 sabkha, 163 sabkha deposits, 163 sabkha sediments, 182 safflorite, 29, 31, 33, 65, 68, 70, 75, 83, 84, 157, 230 safflorite star-twins, 65 safflorite zoning, 69 salic magmatism, 110

saline brine solutions, 165 salt, 147 salt plane, 190 samarskite, 237 samarskite-yttrotantalite series, 232 Sammelkristallisation, 270 sandstones as recipient geoenvironment, 216 sanidine, 103 saw tooth structure, 35 Sb, 124, 127, 184, 186 Sb-As-Tl-Ba mineral assemblage, 145 Sb-Au mineralization, 124 Sb-content, 186 scandium, 264 scapolite, 134, 136 scapolite-bearing skams, 132 scapolitization, 134 Schallenblende, 80 schapbachite, 136 scheelite, 32, 33, 65, 121, 130, 132, 134, 135, 136, 139, 148, 151, 153, 197, 199, 200, 202, 212, 214, 248, 249, 252, 259, 271 scheelite remobilization, 270 scheelite-bearing quartz, 135 scheelite-bearing skams, 131 scbeelite-molybdenite mineral paragenetic association, 249 Schichtengitter, 94 Schlieren, 5, 95, 114, 153 Schlieren bodies, 156 Schlieren chromite, 18, 20, 94, 95, 269 Schouten's experiment, 12 schwartzite, 209 scorodite, 200, 201, 202 Se and Te concentrated in supergene sulphides, 202 Se in pitchblende paragenesis, 232 Se mobilization, 232 Se-linneite, 232 second generation ex-solutions, 49 secondary (cavernous) ore type, 204 secondary minute rutile grains, 84 secondary mobilization of C, 92 sediment participation in sulfides, 149 sedimentary, 123, 146 sedimentary origin, 157, 158 sedimentary process, 161 sedimentary submarine-hydrothermal formation, 142 sedimentary sulphur, 191 sedimentary syngenetic deposits, 151 sedimentary-diagenetic origin, 184 sedimentary-exhalative, 144, 150, 170 sedimentogenic, 110 sedimentogenic ore bodies, 211 sedimentogenic phase, 154 sedimentogenic relics, 154 sedimentogenic stratabound and vein deposits, 184 segregated together in typical associations, 242 segregation, 1, 2, 3, 4, 6, 9, 45, 141, 152, 167 segregation of rare elements, 229 segregation of Sn, Mo, W, 246

segregation of the platinoid elements, 86 segregation/concentration of metallic elements, 261 segregation/concentration of ores by solution/fluids, 268 segregation/distribution, 242 segregation/distribution depending on structure of atoms, 234 segregation/distribution of elements forming parageneses, 235 segregation/distribution of metallic elements, 259 Sekretionsdifferentiation, 153 selective precipitation, 216 selective replacement, 29 selective substitution, 161 selective zonal dissolution, 30 selenide, 116, 136 selenium, 138, 260 selenium minerals, 34, 65, 66 seligmannite, 245 senarmontite, 202 separation from metallic sulfide, 112 sequence of copper replacement, 26 sequence of crystallization, 264 sequence of fractionated crystallization, 264 sericite, 195 sericitization, 32, 159 serpentine, 61 serpentine myrmekites, 61 serpentine/chromite symplectites, 61 serpentinization, 98 sheet lattice, 94 shrinkage cracks, 84 shrinkage fractures, 32 shrinkage pattern of cracks, 105 sial, 108 sialic granitoid, 110 siderite, 101, 127, 140, 141, 174, 175, 178, 179, 200, 210, 232, 244 siderite spherules, 177 Siderocapsa, 177 siderophile, 261, 266 siegenite, 139 silexite, 251 silicate, 56, 60, 179 silicate infiltration, 105 silicates zonally incorporated in a chromite, 74 siliceous algae, 177 siliceous frustules, 175 silicic acid, 259 siliciclastic iron-formations, 174 sillimanite, 159 sillimanite paragenesis, 206 sills, 256 silver, 51, 128, 137, 145, 157, 158, 159, 195, 197, 200, 205, 209, 254 silver in massive sulphide ores, 241 silver lead, 197 silver telluride, 213 silver-all-argentum-dyscrassite assemblages, 209 silver-bearing siderite, 209 647

simatic, 110 simatic magmatism, 110 similarities as expressed by the periodic system, 234 similarity between Hf and Zr, 223 simple twinning simultaneous crystallization, 15 simultaneous deposition, 9 simultaneous precipitation, 58 single-fluid system, 165 Si0 2 , 173 sitaparite, 254 size of atoms and ions, 235, 265 size of metal atom, 14 skarn, 6, 9, 55, 56, 81, 118, 119, 129, 133, 134, 137, 138, 149, 150, 176, 181, 189, 196, 213 skarn bodies, 118, 132 skarn formation, 133 skarn zoning, 131 skarn-gold, 138 skarn-type bimetasomatic zoning, 131 skarnification, 119, 123, 132, 133, 134, 139, 189, 213 skaras derived from dolomites, 133 skarns formed in pelitic rocks, 132 skeletal Bi, 53 skeletal crystals, 52 skeletal growth of siegenite, 52 skeletal in form symplectites, 62 skeletal pyrite, 54 smaller capacity to diffusion, 79 smaltite, 34, 39, 65, 68, 75, 81, 84, 137, 230 smithsonite, 200 Sn, As, Sb, Bi, Pb, Se, Te - the so-called ?lower crust/crust elements, 266 Sn, Sb, Bi, Te, Se (lower crust/crust elements), 260 Sn, W, Mo mineralization, 248 Sn, 119, 184, 186, 187 Sn, W or Mo mineralization, 110 Sn and Sb in pyrite, 257 Sn deposits, 122 Sn in pitchblende paragenesis, 232 Sn mineralization, 248 Sn-associated element-segregation in pegmatites, 250 Sn-Nb mineralization, 122 Sn-W mineralization, 197 Sn-W type, 122 Sn-W veins, 153 Sn-W-Bi mineralization, 181 sodium metasomatism, 122 soft base/soft acid ligand complexes, 259 solid diffusion, 15, 42, 125 solid solution, 10, 12, 14, 15, 41, 42, 44, 45,48, 50, 53, 54, 171 solid solution hypothesis, 45 solid solution of pentlandite and pyrrhotite, 48 solid solutions of ilmenite-haematite, 48 solid solutions of PGE, 270 solid solutions/ex-solutions hypothesis, 50 solid state, 125 solid state process, 55 648

solubility of PGE in hydrothermal solution, 89 solute, 14 solution (dissolution), 12 solution path, 100, 103 solution-fluids, 268 solution-fluids mobilization, 139 solvent, 14 source geoenvironment, 215 source of ore constituents, 216 source rich in sulphate, 124 special cases of element segregation, 223 specialized granites, 122 sperrylite, 4, 19, 20, 85, 86, 87, 226, 266 spessartine, 137 spessartite, 206 sphaerical form, 77 sphaericules, 91, 92, 268 sphaeroidal forms, 78 sphaeroidal structure, 60 sphaeroids, 27, 55 sphaeroids consisting of radiating elongated crystals, 83 sphalerite, 28, 29, 30, 31, 32, 35, 37, 38, 44, 45, 46, 47, 53, 54, 55, 56, 57, 59, 61, 62, 65, 66, 67, 72, 83, 96, 101, 116, 120, 121, 126, 127, 128, 135, 136, 140, 141, 142, 145, 146, 147, 150, 151, 159, 161, 166, 171, 179, 181, 182, 186, 192, 194, 195, 196, 197, 200, 202, 205, 213, 228, 230, 232, 241, 242, 243, 246, 248, 251, 252, 257, 258 sphalerite nucleus, 92 sphalerite replacement, 47 sphalerite sphaericule, 91 sphalerite stars, 47 sphalerite-pyrite symplectites with garnet, 62 sphalerite-pyrrhotite, 139 sphalerite-wurtzite, 75, 171 spherical siderite, 178 spherules, 177 spilite-keratophyre formation, 145 spinel, 19, 44, 49, 55, 74, 115, 269 spinel bodies, 43 spinel formation, 24 spinel group, 238 spinel paragenesis, 114 spinel phases, 55 spinellids, 55 spiral, 6, 7 spiral changes, 36 spodumene, 248, 250 sprouting, 65 sprouting of a crystal, 64 stage of mineralization, 229 stages of deposition, 8 stannian garnet, 136 stannite, 32, 44, 46, 47, 60, 83, 104, 116, 140, 171, 200, 202, 243, 246 stannite/cassiterite symplectite, 60 stannoidite, 104 star-crust, 2 star-shaped bodies (Sternchen), 47 stars of sphalerite, 47

state of disorder, 4 4

successive ex-solutions, 4 3

state of the art, 2 1 8

successive metasomatic replacement stages, 195

stephanite, 241

successive mobilization, 36

stibarsen, 59, 6 0

successive replacement, 11, 15, 36, 37, 38, 39, 40, 42, 43, 5 6

stibiconite, 2 0 2

successive sequence, 37

stibnite, 60, 127, 140, 172, 194, 196, 199, 202, 245, 252, 259

successive waves, 11

stibnite mineralization, 189

successive waves of solutions, 11

stilpnomelane, 175

Sudbury deposits, 260, 2 7 0

stilpnomenale-iron, 179

Sudbury-type, 189, 2 1 6

stock-work type, 145

Sudbury-type ores, 10, 149

stock-works, 120

sulfide-graphite schists, 124

strain ellipsoid, 185

sulfidized bacteria, 91

stratabound, 133, 145, 148

sulphantimonides/sulpharsenides of copper, 140

stratabound copper ores, 163, 208

sulphantimonite, 214

stratabound copper sulphides, 173

sulphate ion, 188

stratabound deposits, 144, 148, 173, 187

sulphate ions in the sea, 191

stratabound Fe-Zn-Mn oxide-silicate mineralization, 147

sulphates, 188

stratabound hydrothermal-sedimentary, 145

sulphide melts, 112

stratabound iron-copper deposit, 170

sulphide metallogeny, 189

stratabound mineralization, 163

sulphide replacement, 196

stratabound scheelite, 150

sulphide replacement deposits, 195

stratabound scheelite mineralization, 173

sulphide-barite concentrations, 197

stratabound tourmaline-rich layers, 151

sulphide-graphite schists, 146

stratiform deposits, 151

sulphide-oxide level, 215

stratiform layers in dolomite, 186

sulphide-oxide level in the deeper interior of the earth, 188

stratiform massive sulphide ores, 170

sulphide-rich, 216

stratiform ore bodies, 146

sulphide-rich tuff, 173

stratiformed sheet o f igneous complexes, 17

sulphosalts, 188

stromatolites, 85

sulphur, 89, 141, 163, 166, 170, 189, 217

stromatolitic structures, 177

sulphur bacteria, 188, 192

stromeyerite, 26, 63, 66, 137, 146, 157

sulphur crystals, 192

structural corridors, 107

sulphur deposits, 189

study case approach, 219

sulphur equilibria, 192

subduction, 130, 143

sulphur in metallogeny, 188

subduction metamorphism, 143

sulphur in meteorites, 188, 189

subduction zone, 108

sulphur isotope composition, 192

subgroup of the VIII family o f the periodic system, 227

sulphur isotope studies, 10

subgroups of the periodic system, 220

sulphur isotopes, 124, 146, 147, 189, 191, 192, 2 0 0

subidioblastic arsenopyrite, 65

sulphur of the ores derived from country rocks, 149

submarine exhalative, 151, 153, 159, 160, 169, 170, 179

sulphurization, 10, 115, 149, 189

submarine exhalative gold, 151

super-saturation, 48

submarine fan successions, 175

supergene, 3, 11, 57, 120, 147, 199, 200, 201, 204, 205, 206,

submarine hot springs, 172

209, 2 1 0

submarine pyroclastic rocks, 152

supergene leaching, 200, 209

submarine volcanic, 169, 170

supergene mineralization, 199

submarine volcanism, 170, 184

supergene minerals, 145

submarine-hydrothermal, 35, 147

supergene oxidation of sulphides to sulphate, 191

submarine-hydrotbermal sedimentary, 146

supergene remobilization, 148

substances analogous in chemical composition have

supergene replacement, 12

similarity in crystalline form, 262

supergene zone, 25, 201, 245

substitution, 2

superheated H 2 0 , 150

substitution ex-solution, 4 2

superheating, 55

substitution solid solution, 41

superimposed, 3, 4, 6, 39, 255

substitution-diffusion, 15

superimposed hydrothermal paragenesis, 137, 141

substitution-replacement, 10

superimposed intercontinental rift mobilizations, 186

subvolcanic granites, 120

superimposed metallogeny, 129

successive crystallization, 11, 39 successive crystallization phases, 39 successive deposition, 9

superimposed mineral paragenesis, 115 superimposed paragenesis, 116, 136, 137, 186, 187, 2 0 4 superimposed paragenetic association, 116

649

superimposed phase, 116 superimposed sulphide phase, 116 supracrustal amphibolites, 178 surface zone, 199 suture zone, 108 sutures, 110 sylvanite, 32 sylvanite/quartz, 53 sylvine, 181 sylvinite, 62, 242 symbol population, 107 sympathetic relationship, 158 symplectic, 15, 16, 21, 52 symplectic form, 43, 46, 67 symplectic intergrowth, 26, 30, 47, 54, 71 symplectic pyrite, 47 symplectic replacement, 47 symplectic rutile, 54 symplectic sphalerite and pyrrhotite, 56 symplectic spinel, 24 symplectic-epitactic growth, 72 symplectically intergrown, 21 symplectite, 9, 11, 13, 39, 52, 53, 54, 55, 56, 57, 59, 61, 71 symplectite-like forms, 30 symplectites with silicates, 62 synaeresis, 4, 84, 105 synantectic intergrowth, 61 synantectic-symplectic intergrowth, 132 synantetic, 15, 16 synantetic reaction, 21, 55 syncline, 174 syndiagenetic, 142, 147, 192 syndiagenetic minerals, 145 syngenetic, 11, 119, 124, 145, 146, 148, 149, 152, 170, 186, 197, 211, 257 syngenetic deposit, 211 syngenetic hypothesis, 161 syngenetic ore deposits, 194 syngenetic origin, 146, 192 syngenetic Pb- and Fe-ores in dolomite, 186 syngenetic pyrite, 147 syngenetic-diagenetic fabrics, 270 syngenetic-epigenetic, 146 syngenetic-sedimentary, 123, 145, 147, 148, 152, 157, 159, 161, 163,206 syngenetic-sedimentary deposits, 147 syngenetic-sedimentary ores metamorphosed, 149 syngenetic-sedimentary origin, 148, 151 syninversion-orogenic granitoid intrusions, 138 synkinematic, 150 synmetamorphic barite, 184 synorogenic fluids, 185 synorogenic fluids and metallogeny, 185 synorogenic late Palaeozoic granite, 120 synorogenic ore deposition, 185 synorogenic stages, 109 synsedimentary, 142, 146, 147, 148, 149, 152, 153, 173, 179, 204 synsedimentary features in epigenetic ores, 145 650

syntectonic mineralization, 192 syntectonic recrystallization, 95 syntexis, 5 synthetic mineralogy, 12 szäjbelyite, 135, 138

Ta, 186 Ta-Nb minerals, 187 table-cloth structure, 43 tactite, 153 talc, 195 talcification, 212 tamatinite, 128 tantaliferous cassiterite, 250 tantalite, 52, 74, 121, 248, 250 tantalocolumbite, 250 tapiolite, 247 Te, Se, Bi, Au, Ag hydrothermal paragenesis, 241 Te, Se, Bi, Au, Ag paragenesis, 241, 242 Te, Se, Bi, Au Ag element segregation (paragenetic associations), 243 Te, 186 tealite, 32, 140, 246 tectonic belt, 175 tectonic brines, 185 tectonic differentiation, 5 tectonic effects, 94 tectonic heat (generated by friction), 212 tectonic mobilization, 4, 114 tectonically mobilized, 18, 156 tectonism-magmatism, 110 tectono-volcanic events, 169 tectonogenic features, 107 tectonogenic model, 110 tectonogranular, 95 tectonogranular chromite, 94 tectonogranular magnetite, 24 tectonomagmatic cycle, 229 tectonomagmatic granitic intrusion, 110 tectonomagmatic model, 107 tectonoplastic deformation, 95 tectonoplastic mobilization, 95, 97 tectonoplastic mobilization of chromite, 97 tectonothermal phenomena, 7 Tekoblastesc (tecoblastesis), 64 telemagmatic, 169 telescoped, 10 telescoping, 3 telethermal, 124, 139, 142, 152, 213 telethermal crystallization phase, 118 telethermal ore veins, 213 telethermal phase, 56 telethermal pitchblende paragenesis, 268 telethermal veins, 119 tellurbismuth, 241 telluride, 54, 116 tellurite, 32 tellurium, 138, 260

tellurium-bearing canfieldite, 136 tellurobismuthite, 151, 202, 242 temperature determination, 11, 12 temporary resting position, 212 tennantite, 27, 28, 30, 56, 151, 157, 186, 196, 200, 209, 245 tennantite-tetrahedrite, 186 Tennessee Valley -type of deposits, 163 tenorite, 81, 201, 202 tension faulting, 109 terrestrial basalts, 239 tetradymite, 151, 199, 201, 202, 241, 242 tetrahedral structure, 94 tetrahedrite, 27, 30, 33, 34, 38, 40, 53, 58, 62, 65, 69, 72, 84, 105, 137, 140, 141, 142, 146, 149, 194, 196, 197, 199, 200, 202, 209, 214, 228, 230, 232, 242, 251, 252 tetrahedrite-tennantite, 35 textural analysis, i, 2, 15, 111, 123, 153 thallium, 186 thermal metamorphic-metasomatic, 118 thermo-electric potential, 12 Thiobacillus ferrooxidans, 104, 202 thiosulphate complexes, 260 tholeiite, 114 tholeiitic, 174 tholeiitic basaltic magma, 143 tholeiitic series, 175 thoreaulite, 250 thorian brannerite, 125 thorianite, 232 thorium, 234 thrust fault, 195 thrust-orogenic evolution, 185 thucholite, 92, 93 thucholitic sphaeroids, 92 Thuchomyces lichenoides, 158 Ti, 256 Ti, V, Cr, Μη, Fe (horizontally related subgroup elements), 256 Ti-mineral intergrowth pattern, 49 tiemannite, 137 tin, 120, 125, 142, 151, 186, 187, 237, 242 tin deposits, 122, 148 tin granites, 130 tin mineralization, 122, 136 tin ore, 246 tin-bearing massif, 122 tin-bearing region, 125 tin-silver deposit, 140 tin-silver fumarolic, 248 tin-silver types, 122, 121, 248 tin-tungsten ores, 136 tin-tungsten province, 135 titaniferous haematite, 48 titanite, 24, 53, 75, 102 titanium-rich magnetite, 21 titanoferous magnetite, 112 titanomagnetite, 20, 21, 35, 43, 53, 72, 115 titanomagnetite replaced by chalcopyrite, 25 TI, 202

todorokite, 201, 205, 206 topaz, 121, 125, 136, 197 topazification, 125 topological, 45 topological enrichment, 27 topometasomatic, 49, 154 topometasomatic solutions, 269 topomorphic basis, 141 topomorphism, 141 tourmaline, 62, 121, 122, 125, 199, 250 tourmaline metasomatism, 168 tourmaline-quartz veins, 213 tourmalinite, 151 tourmalinization, 125 trace element, 146 trace elements in sulfides, 257 transition stages from pegmatitic-pneumatolytic to hydrothermal, 251 transuranides, 220 tremolite, 177, 195 tripuhyite, 202 trogtalite, 34 tungsten, 120, 142, 148, 150, 151, 231, 247, 259, 271 tungsten deposits, 252 tungsten-Sn mineralization, 247, 248 turbidite, 148, 178 twin planes, 105 twinning plane, 105 two-dimensional fibrillar dispersive system, 77 type, 229 typomorphism, 12 tyrolite, 201

U, V and Se zonally distributed, 214 U deposits, 152 U-containing pegmatitic mineral paragenesis, 232 U-hydrothermal paragenesis, 229 U-paragenesis, 229, 242 U-Th mineralization, 125 U-Th-Pb, 193 ullmannite, 31, 140 ultrametamorphic geoenvironment, 168 ultrametamorphic origin, 18 ultrametamorphic-metamorphic, 153 ultrametamorphic-sedimentogenic origin, 153 ultrametamorphism, 3, 20 ultrametamorphism of basic and ultrabasic rocks, 111 ultrametamorphism-granitization, 133 ultrametamorphosed, 6 ultrastructure of a microbial mat-generated phosphorite, 91 ultraviolet photochemical process, 176 ulvite, 43, 44, 115,240 ulvite cloth-pattern, 240 ulvospinel, 43, 115 umganite, 232 undulating extinction, 94 unfolding of the geological spiral, 219, 254, 265, 267 uniform 30% iron tenor, 173 651

uniform isotopic composition of lead, 195 uniformitarian concept, 155 uniformitarianism, 3, 20, 175, 216, 267 unifying principles, 219 unmixing, 9, 10, 12, 14, 15, 41, 42, 43, 45, 46, 48, 49, 50, 54, 133 unmixing hypothesis, 44 U 0 2 and T h 0 2 form isomorphous series, 234 upper lithosphere, 188, 189 upper mantle, 107, 108 upper mesothermal stage, 181 uraniferous skarn, 138 uraninite, 35, 39, 52, 53, 71, 72, 74, 75, 76, 84, 92, 96, 101, 102, 106, 122, 137, 147, 157, 158, 231, 232, 251 uraninite disseminations in skarn ores, 138 uraninite epitactic on columbite, 71 uraninite inclusion, 71 uranium, 112, 119, 193, 203, 209, 216, 220, 234, 260, 268 uranium in coals, 231 uranium is definitely a lithophile element, 233 uranium minerals, 231 uranium-vanadium roll, 214 uranothorite, 35 Uranpecherz, 158 uranyl carbonate complexes, 260 uvarovite, 20, 115, 116, 154

V and Ti in manganese mineral paragenesis, 255 vadose water, 210 valency, 14, 15, 41, 220, 223 valentinite, 202 valleriite, 115, 116, 136, 162 van der Waal's forces, 264 van't Hoff s law, 9 vanadiferous iron ore, 115 vanadiferous magnetite, 112, 115, 239 vanadinite, 195, 255 vanadium, 212 varlamoffite, 104 veinform character, 44 veinform jacobsite transecting braunite, 40 veinform siegenite veinform zinkite, 48 veinlet of silicate, 39 velonoblastic anthophyllite, 44 Vernadsky's phreatic cycles, 268 Verteilungsgesetz der Elemente, 2, 233 vertical zoning, 152, 201, 213, 214 vesuvianite, 136, 138 vesuvianite-hedenbergite skarns, 135 violarite, 116, 202 visco-elastic, 108, 155, 156 visco-plastic, 107 volatile activity, 150 volcanic emanation, 170 volcanic exhalations, 150, 152 volcanic gases, 150, 181

652

volcanic trend: basalts—»dolerites—»andesites—>rhyolites, 111 volcanic-sedimentary origin, 161 volcanic-sedimentary series, 145 volcanism, 110, 216 volcano-exhalative, 149, 170, 271 volcano-sedimentary, 150, 174, 206 volcano-sedimentary deposits, 184 volcano-sedimentary pile, 160 volcano-sedimentary rocks, 170, 184 volcano-sedimentary stratabound formations, 184 volcanogenic, 176 volynskite, 242 vonsenite, 136 vredenburgite, 204, 206 vugs, 209

W, Mo, Sn mobilization, 271 W, 119, 184, 187, 237 W skarn, 130 W-Bi and tourmaline mineralization, 181 W-Cu-Ag ores, 135 W-Mo-Cu deposit, 135 wad, 196, 205 wad psilomelane, 255 wallrock alteration, 128, 152, 159, 214 wallrock derived solutions, 180 wallrock metasomatic alteration, 170 wallrocks, 182, 201 Wasserkies pyrite, 25, 54, 62 wave replacement, 39 wave substitution, 15 weathered andesites, 173 weathering, 102 weathering and alteration of ore minerals, 98 weathering mantle, 5 weathering processes, 98 weathering products, 208 weathering zone, 11 weathering-alteration processes, 102 Wegsamkeit, 105 wehlite, 244 westerveldite, 116 white dolomitic rock, 185 Widmanstätten textures, 59 Wiederverkittung, 61 Wiederverkittungs serpentine, 94 wiikite, 54 willemite, 195 winchite, 206 wire silver, 31 wittichenite, 33, 58, 201 Witwatersrand controversy, 157 Witwatersrand system, 157, 159 W0 3 , 135 wolfram, 134, 139 wolframite, 32, 33, 65, 69, 121, 122, 126, 134, 136, 197, 199, 200, 213, 243, 248, 249, 251, 252, 271

wolframite deposit, 121 wolfiramite/ferbierite, 134 wollastonite, 195, 197 wollastonite skarns, 138 wood cassiterite, 83 working hypothesis, 272 worm-like relics, 58 wulfenite, 231 wurtzite, 142, 172, 228, 245

X-shaped bodies of sphalerite, 47 xenoblast, 154 xenoblastic, 64 xenoblastic pattern, 66 xenoblastic pyrite, 66 xenocryst, 5, 19, 21, 61 xenoliths, 6, 56, 118, 132, 133, 137, 155 xenothermal, 140 xenothermal deposits, 258

Y, 186

yellow ore, 102 yokosukaite, 205 Young granitic complexes yttrium, 234 yttrocrasite, 232

Zerfallstruktur, 60 zinc, 151, 159 zinc-lead ore, 145 zinc-lead-barite-fluorite deposits, 165 zincblende, 127 zinkite, 48

zinnwaldite, 197 zircon, 69, 75, 105, 106, 120, 122, 158, 181, 187, 237 zirconium/hafnium, 264 Zn, 127, 146, 172, 186 Zn and Ag in chalcopyrite, 170 Zn-Pb skarn, 130 Zn-Pb-Fe-Ba-F-enrichment, 198 Zn/Pb ratios, 170 zoisite, 195 zonal chalcopyrite, 46 zonal cobaltite, 62 zonal differentiation, 213 zonal distribution, 214 zonal growth, 68, 71 zonal growth of cobaltite, 68 zonal growth of galena, 68 zonal growth of smaltite, 68 zonal malacon, 75 zonal martitization, 100 zonal mineralization, 213 zonal replacement, 30, 47 zonal structure, 88 zonation, 199 zone of fading acidity, 213 zone of maximum leaching, 213 zone of oxidation, 201 zone replacement, 13 zoned carbonate, 36 zones of friction, 167 zones of rupture, 107 zoning, 133 zoning and epitactic growth, 74 zoning of pyrite, 69 Zr, 186 zvyaginstevite, 224 Zwischenprodukt, 100

653

Subject Index to the Illustrations

Accretion, 722 albandine, 893 albandite, 499, 603, 604 allemontite, 478, 479, 480. allemontite ΙΠ, 481 alteration following a crack, 870 alteration margin, 827, 868 altered chromite, 730. anatase, 828, 829, 832, 833 andalusite, 503 anglesite, 151, 152, 153, 154, 155, 156, 160, 515, 642 "Anlaufen effects", 900 anorthosite, 782, 784, 785, 786, 787, 788, 810, 811, 812, 813, 814 antigorite, 5, 10 - asbestos, 741 antimonite, 504, 505 apatite, 20, 28, 29 apogranite, 926 argentite, 157, 161, 162, 164, 167, 168, 169, 170, 171, 510 arsenopyrite, 130, 131, 132, 239, 335, 528, 597, 598, 600, 710 - crystalloblasts, 530, 531, 532 - epitactic on sphalerite, 597, 598 As-Sb-phase, 481 atoll replacement, 658 atoll-type of chalcopyrite replacement, 247 atoll-type replacement, 59, 79, 106, 144, 148, 159, 191, 199, 201

Bacterial action, 761, 764, 765 bacterially formed pyrite, 763 bacterially formed sphaeroids, 764 basalt, 24, 25, 486, 487 basaltic fragment, 719 - groundmass, 23, 24, 27 - pieces, 719 bauxite, 884, 886, 887, 890 berthierite, 50, 51, 52 Bi (bismuth), 191, 192, 193, 194, 195, 196, 197, 198, 199, 200, 201, 202, 203, 204, 205, 206, 207, 222, 223, 225, 249, 250, 365, 389, 390, 458, 460, 461, 462 Bi in symplectic intergrowth with silicate, 390 biotite, 904, 905 birbirite, 816, 818 bird's eye structure, 702, 844 bismuthinite, 209, 210, 250, 423, 424 blastic magnetite, 293

654

- pyrite, 543, 547, 552 - rutile, 410 blende, 72, 301, 425 blomstrandine, 877 boehmitic bauxite, 755 bornite, 61, 63, 65, 85, 86, 122, 123, 240, 242, 244, 247, 258, 387, 424, 439, 440, 444, 445, 446, 447, 448, 451, 452, 453, 454,455,463, 464, 465, 548, 549, 550, 602 bornite/chalcocite symplectite, 444 boudinage, 790 boulangerite, 60, 84 bournonite, 506, 508 box-work structure, 848 braunite, 267, 268, 377, 591, 592, 593, 594, 670, 671, 742, 743, 744, 746, 747, 749, 751 bravoite, 471, 472, 473, 474, 567(b), 766, 767 breakdown symplectite texture, 478, 479, 480 bronzite, 2 brown iron, 76, 77, 78, 225, 389, 653, 678, 679, 854, 856, 857, 891 brown serpentine, 490

Calcite, 395, 536, 780 carbonate exhibiting gel pattern, 233 carbonates, 153, 154, 689, 699, 828, 829, 832, 833 cassiterite, 53, 176, 177,178, 179, 180, 181, 182, 187, 188, 189, 190, 485, 572, 573, 574, 693, 694, 695, 808, 809 - idioblast, 572 cataclasis, 774, 775, 807 - cataclastic chromite, 494 cauliflower structure, 721 cell wood structure, 224 chalcedony, 11 chalcocite, 64, 65, 68, 69, 74, 75, 87, 108, 109, 110, 111, 112, 113, 118, 122, 123, 124, 241, 245, 247, 451, 452, 453, 454, 455, 483, 496, 546, 548, 675, 678, 679, 891 - "ring", 891 chalcopyrite, 55, 58, 59, 61, 64, 65, 67, 70, 71, 72, 73, 74, 75, 76, 80, 85, 99, 100, 201, 236, 237, 242, 243, 244, 246, 247, 248, 260, 262, 264, 265, 266, 298, 299, 300, 302, 303, 305, 306, 307, 308, 309, 313, 314, 316, 317, 318, 319, 320, 321, 322, 323, 324, 325, 337, 384, 386, 387, 388, 404, 436, 437, 438, 439, 440, 441, 442, 443, 452, 463, 464, 465, 518, 519, 520, 547, 588, 589, 590, 660, 761, 762, 851, 860, 891 - replacing sphalerite, 502 chloanthite, 95, 196, 197, 198, 201, 202, 205, 208, 259, 365, 462 chrome-spinel, 24, 663, 892

chromite, 5, 6, 7, 11, 12, 264, 265, 295, 297, 488, 489, 490, 492,493, 494, 518, 577, 579, 606, 607, 608, 609, 610, 723, 724, 725, 726, 728, 729, 730, 733, 734, 740, 741, 755, 756, 758, 773, 774, 775, 776, 777, 778, 779, 780, 781, 782, 784, 785, 786, 787, 788, 789, 790, 810, 811, 812, 813, 814, 815, 816, 818, 820, 822, 823, 824, 825, 827, 892, 909,910,911,916 - clastic grains, 909, 910 - "potatoes", 789 chromite-anorthosite, 783, 784, 786, 787 chromospinel, 25, 486, 487, 826 chrysocolla, 648, 649, 650, 651 circular synaeresis cracks, 705 clastic chromite, 663, 756, 819, 820, 912, 913, 914, 915 - magnetite, 837, 838, 842 clausthalite, 212, 213, 215, 395, 396, 469, 470, 534, 535, 536,711,712,713, 925 cleavage, 557, 558 - pattern, 895, 896 clinoenstatite, 912, 913 cobaltite, 160, 161, 162, 163, 173, 174, 203, 204, 249, 500, 559, 794 coffinite, 618, 704, 706 colloform, 800, 801, 802, 804, 805, 847, 925 - hutchinsonite, 645 - iron oxides, 542 - like zones, 566 - limonite, 661 - malachite, 88 - manganese, 720 - pitchblende, 618, 625, 629, 631, 632 - pyrite, 87, 605, 660 - sphalerite, 150, 640 - structure, 721 - tenorite, 652 - zoning, 566 columbite, 363, 374, 375, 376, 580, 581, 582, 583, 585, 586, 612, 613, 614, 617, 906 Co/Ni mineral, 668 cordierite, 503 corroded chromite, 610 corundum crystalloblast, 355 covellite, 64, 81, 82, 83, 84, 85, 86, 246, 469, 470, 711, 712, 713, 763, 902 crack pattern, 898, 899, 900, 907 cross-shaped structures, 252 cross twinning, 901 cryptomelane, 657, 658, 659, 691 crystalline uraninite, 864 crystallites, 683 crystalloblastic haematite, 534 - pyrite, 184, 251 cubanite, 134 cumulate, 15, 19 cuprite, 68, 69, 89, 90, 91, 365, 482,483, 569, 570, 571, 643 curved synaeresis, 709 - twinning, 902

Davidite, 358, 359, 360, 361, 362, 408, 409, 703, 799, 873, 874, 875, 876 - alteration, 873

decoloration, 815, 817, 818 - margin, 493 deferrification, 742, 752, 759 dentritic pattern of silver, 675, 676, 677 - brown iron, 678, 679 depleted rings, 751 diffusion rings, 750, 887, 888 - sphaeroids, 754 diopside, 15, 19 dispersion halo, 926 double radioactive halo, 905 "drop"-like intergrowths, 414 dunite, 5, 10, 776, 882

Emery, 355 emplectite, 194, 258, 423, 424 enargite, 48, 49, 57, 67, 101, 117, 246, 441, 442, 443, 537 epitactic, 578, 588, 619 - bornite, 602 - boumonite, 601 - development, 619 - ilmenite, 587 -jacobsite, 591, 592, 593 - magnetite, 577, 733, 734, 740 - pyrrhotite, 596 - spinel, 589, 590 - uraninite, 586, 614 epitaxis, 580 etch pattern, 903 euxenite, 868, 869, 870, 871 ex-solution, 609, 876 - body, 298 - -like shapes, 330 - -like forms, 316, 317, 318, 319, 322, 323, 324, 326, 327, 328 - of rutile, 295 extensions of chalcopyrite, 305

Feather-like pattern, 532 - sphalerite, 416 feldspar, 879 ferberite, 843 ferrogabbro, 15, 17 ferroplatin, 610, 611, 722, 723, 725, 726, 727, 728, 729, 730, 732, 733, 735, 736, 737, 738, 740, 741 - nugget, 731 fibriolitic silicates, 733, 740, 741 flame-form pentlandite, 342 flame pentlandite, 340 fourmarierite, 864 frame-like bodies, 339 franklinite, 344, 345, 346, 347, 348 Galena, 83, 95, 125, 126, 127, 128, 131, 132, 133, 134, 135, 136, 137, 138, 139, 140, 142, 143, 144, 145, 158, 235, 237, 238, 239, 251, 252, 310, 311, 312, 313, 329, 330, 337, 383, 385, 388, 414, 415, 429, 430, 431, 432, 433, 434, 435, 436, 456, 457, 503, 506, 507, 508, 547, 550, 557, 558, 642, 706, 715, 716, 764, 791, 792, 793, 896, 897 galenobismuthinite, 697

655

gangue, 135, 140, 141, 151, 198, 414, 415, 416, 427, 428, 440, 511, 512, 515, 539, 623, 634, 698, 699, 796, 799, 806, 807, 861, 863, 909, 910, 922 garnet crystalloblast, 264, 265, 518 garnets, 266,516,517,532 gel marcasite, 639 - pitchblende, 626, 630, 666, 711, 712, 713 - sphaeroids, 223, 655 - sphalerite, 333 - structures, 570 - structures of colloform sphalerite, 641 germanite, 94, 95, 96, 430, 431, 432 gersdorffite, 605 gibbsitic bauxite, 750, 752 - matrix, 753 - pisolite, 752 globules, 880, 881 gold, 132, 239, 256, 257, 259, 260, 261, 391, 406, 407, 509, 623, 919, 920, 921, 922 - crystal, 921 - in symplectic intergrowth with silicate, 391 - tellurite, 167, 168 granular gold, 261 - pyrrhotite, 545 graphic galena, 433 - in shape clausthalite, 396 - pegmatite, 868, 869 graphic-like uraninite, 397, 398 - galena, 433 - intergrowth, 402, 404 - pyrite, 430 - replacement, 370 - sphalerite, 381, 382 - symplectic haematite, 418 - symplectic intergrowth, 392 graphic/myrmekitic galena, 427, 428, 429 - -like galena, 435 gratonite, 150, 644, 645 gudmundite, 384, 519, 520

Habitus, 575 haematite, 34, 57, 177, 178, 191, 192, 227, 230, 246, 353, 354, 355, 356, 357, 360, 408, 409, 417, 418, 419, 420, 466, 467, 468, 533, 535, 537, 538, 654, 661, 662, 680, 683, 686, 749, 758, 820, 823, 824, 825, 826, 838, 840, 843, 848, 876, 892, 925 - lamellae, 358 - velonoblast, 538 haematitic pisolite, 757 - crystalloblast, 536, 538 hastite, 213, 214, 215, 343, 349, 350 hauchecornite, 211 hausmannite, 691, 852 hawleyite, 615, 616 heat haematite, 417, 418 - martitization, 418 hornblende, 503 hutchinsonite, 645 hydrohaematitized matrix, 757 hydrothermal formations (cryptomelane), 624 hypersthene, 21

656

Idioblastic, 124, 454 - arsenopyrite, 529 - pyrite, 551, 556 - pyroxene, 917 idioblasts, 107, 847, 917 idiomorphic chromite, 607 - hornblende, 503 -quartzes, 507,513,514 ilmenite, 7, 29, 31, 34, 35, 36, 37, 45, 46, 55, 228, 229, 270, 271, 272, 273, 274, 275, 276, 279, 280, 281, 282, 283, 284, 286, 288, 290, 291, 292, 293, 294, 349, 350, 351, 352, 353, 354, 355, 356, 358, 359, 360, 379, 380, 417, 418, 420, 539, 587, 589, 661, 662, 872, 876, 901, 907, 908 - crystalloblast, 538, 539 - ex-solutions, 293 - lamella, 285, 908 ilmenites, 828, 829, 830, 831, 832, 833 induced reduction, 231 intergranular serpentine, 493 interleptonic spaces, 17 interspersed pyrite bodies in sphalerite, 331, 332 interzonal chalcopyrite, 321 - malachite, 569 intracumulate, 915, 917 iron leached bauxite, 744 iron pisolite, 578

Jacobsite, 267, 268, 377, 592, 593, 594 - epitactic on braunite, 377, 591 jamesonite, 146, 147

Kaolinite, 752, 753 kidney forms, 624 kimberlite, 30, 31

Lamellae, 364, 822, 838 - replacing enargite, 442 laterite, 892 lateritic eluvial cover, 722 leaching patterns, 745 lenses, 785 leopard bauxite, 885, 886 lepidocrocite, 226, 680, 681, 682, 683, 684, 685, 686, 848 leuchtenbergite, 293 lievrite, 231, 653, 654, 856, 857 limonite, 661, 724, 726, 819, 853, 879, 880, 881 - rings, 878 limonitic coating, 722, 723 - matrix, 756 - -oolitic relics, 226, 755 loellingite, 522

Maghemite, 421, 835, 841, 854, 859 magnetite, 7, 8, 9, 10, 13, 14, 15, 16, 17, 19, 20, 21, 22, 25, 26, 29, 32, 33, 34, 35, 36, 38, 40, 41, 42, 43, 44, 45, 46, 47, 55, 57, 58, 59, 60, 227, 228, 229, 230, 231, 232, 233,

234, 235, 236, 270, 271, 272, 274, 275, 276, 277, 278, 279, 280, 281, 282, 283, 284, 285, 286, 287, 289, 290, 291, 293, 294, 364, 378, 417, 418, 419, 420, 421, 422, 466,467, 468, 484, 486, 487, 542, 551, 567(a), 579, 599, 619, 620, 621, 622, 700, 701, 758, 797, 798, 834, 835, 836, 837, 838, 842, 858, 859, 907, 908 - plug, 907 malachite, 88, 89, 90, 241, 482, 570, 571, 643, 692 malacon, 566, 567(a), 619, 620, 700, 701, 797, 798 Maltese cross-shaped siegenite, 369 manganese, 655, 656 - nodules, 717, 718, 721 - oxides, 724 manganite, 636, 637 mantle diapir, 3 marbles, 421 marcasite, 44, 120, 161, 220, 326, 327, 328, 379, 380, 468, 526, 527, 541, 639, 640, 702, 845, 849, 850 martite, 841, 843 - magnetite, 227 martitization, 270, 418, 758, 837, 842 martitized, 22, 836, 840 - magnetite, 357, 734 melanocratic, 19 metacinnabar, 576, 687, 688, 689 metamictic, 926 mica, 18 - flakes, 502 microcline, 397, 398, 399, 400, 401, 906 millerite, 102, 103, 104, 105, 209, 210, 211 micro-diffusion rings, 745 mobilization, 727 molybdenite, 115, 770, 771, 772 multiple epitaxis, 579 mylonitized, 773 myrmekite-like bornite/chalcocite symplectite, 449, 451 myrmekitic forms, 22, 24, 25, 73, 139, 442 - haematite, 359 - intergrowth, 491 myrmekitic-like, 283, 284, 492 - bodies, 486, 817 - forms, 443, 445 - intergrowth, 408, 409 - patterns, 483 - relics, 463, 464 - serpentine, 490 - symplectite, 438, 471, 472, 473, 474, 482 myrmekitic quartz/pyrrhotite, 499 - serpentine, 488, 489, 493, 494 - symplectic form, 421, 423, 447, 448 - symplectic intergrowths, 449, 450, 453, 454 - symplectite, 447, 448, 491 myrmekitic/graphic forms, 437 - galena, 436 myrmekitoid-shaped, 485

Native copper, 91, 126, 241, 366 - selenium, 533, 535 - silver, 218 needle form, 671 - iron, 663, 664, 665, 848, 851 neocrystallizations, 86, 87

neodigenite, 76, 77, 78, 82, 106, 107, 144, 145, 411, 412, 413, 497 niccoline, 154 niccolite, 92, 153, 173, 174, 175, 199, 200, 201, 202, 540, 638, 698, 699 norite, 70, 872 nugget, 610, 611,723, 733

Olivine, 606, 607 olivine-bomb, 1, 2, 23 olivine-pyroxene crushed zone, 23 olivinefels, 14 ooids, 745, 759, 760 oolites, 664, 665, 742, 743, 744, 747, 748, 749, 752, 756, 760, 890, oolitic limonite, 819 - structures, 746 oriented ilmenite, 380 osmiridium, 610, 723, 724, 731, 735, 736, 739

Pegmatite, 906, 915 pentlandite, 70, 339, 340, 341, 342 perovskite, 830, 831 petzite, 171, 172, 393 phosphate, 759 phosphorite, 759 phyllosilicates, 909, 910 pigment size relics, 556 pisolite, 752, 753, 758 - sphaeroid, 750 - -oolite, 578 pisolitic sphaeroids, 752 - structures, 886 pitchblende, 216, 217, 218, 220, 221, 222, 253, 254, 255, 256, 257, 258, 263, 423, 424, 500, 625, 626, 627, 628, 629, 630, 632, 633, 667, 707, 708, 709, 710, 715, 716, 796, 800, 801, 802, 803, 804, 805, 806, 807, 808, 809, 865, 866, 867, 925 p'agioclase, 13, 14, 15, 16, 17, 915 plastically deformed, 785 platin, 610, 611, 836 platinoid, 7, 737 - minerals, 735, 736, 918 platinum, 725 plattnerite, 647 polianite, 656, 657, 658, 659 - sphaeroid, 657 polygonal synaeresis, 709 proustite, 163, 166, 405 pseudo-ex-solutions, 273, 343 pseudobrookite, 417, 418, 419, 420 pseudomorphosed, 418, 419 psilomelane, 658, 659, 670, 672, 673, 674, 852 ptygmatically folded, 782 pyrargyrite, 54, 158, 159, 166, 367, 405 pyrite, 38, 39, 40, 41, 42, 43, 44, 45, 46, 47, 48, 49, 50, 51, 52, 53, 54, 63, 65, 71, 98, 99, 100, 102, 103, 104, 105, 106, 107, 108, 109, 110, 111, 112, 113, 114, 115, 116, 117, 122, 123, 124, 125, 126, 127, 128, 129, 146, 147, 157, 159, 166, 183, 184, 185, 186, 187, 189, 190, 224, 657

228, 260, 431, 501, 554, 763, 921,

232, 262, 434, 540, 555, 764, 922

233, 263, 451, 541, 556, 791,

237, 304, 452, 542, 558, 792,

238·, 334, 465, 544, 575, 847,

242, 338, 466, 545, 584, 859,

243, 383, 467, 546, 599, 896,

244, 385, 468, 548, 605, 897,

247, 402, 484, 549, 669, 899,

248, 405, 496, 550, 695, 903,

259, 430, 498, 553, 761, 920,

- crystalloblast, 546 - epitactic, 600 - sphaeroids, 502, 765 pyroxene, 13, 15, 16, 18, 19, 915, 916, 917 - idioblasts, 916 - pegmatoids, 918 pyrrhotite, 32, 55, 130, 226, 339, 340, 341, 342, 379, 380, 386, 425, 426, 456, 457, 499, 545, 596, 603, 702, 844, 845, 846, 847, 848, 849, 850 - symplectite, 33, 56 - symplectite with gangue, 426

Quartz, 240, 383, 394, 496, 506, 507, 508, 509, 510, 513, 760 quartz/pyrite symplectic intergrowth, 498, 500

Radioactive halo, 397, 398, 582, 586, 614, 904, 906 radiogenic galena, 706, 708 - lead, 709 ramsdellite, 719 rammelsbergite, 152, 523, 524, 689, 699, 796 realgar, 116, 117 rebinding serpentine, 776 renierite, 95, 96 rhyolite, 878, 879 rhythmic alteration, 851 rhythmical alteration, 874 rim, 860 ring-like sphalerite, 416 roseite, 735, 736, 739, 740, 741 rutile, 61, 62, 296, 297, 356, 357, 359, 361, 362, 408, 409, 602, 609, 799, 828, 829, 832, 833, 836, 876

Safflorite, 121, 152, 160, 161, 203, 522, 562, 563, 564, 565, 617 - crystalloblasts, 521 - stars, 523, 524 sandstone, 883 scheelite, 527 Schlieren chromite, 4, 779, 781 Se, 925 sericite, 879 serpentine, 488, 489, 490, 493, 774, 776, 778, 790 serpentinization, 882 shrinkage, 220 siegenite, 372, 373, 407 siegenite-uraninite symplectite, 371, 372 silicate, 57, 62, 140, 141, 251, 271, 272, 380, 381, 382, 383, 384, 387, 391, 404, 406, 407, 440, 495, 496, 497, 504, 505, 544, 594, 599, 607, 737, 765, 893 - fragment, 755 - symplectic with rutile, 602 658

silver, 164, 165, 167, 169, 218, 367, 675 skeletal clausthalite, 395 - crystal, 381, 382 - crystals of Bi, 389 - gold, 407 - pyrite, 151, 402, 403, 404, 515 - sphalerite, 415 smaltite, 216, 217, 219, 254, 255, 263, 521, 560, 710 sperrylite, 8, 9 sphaeroidal forms, 850 - pyrite, 761 - sphalerite, 762 - structures, 767 sphaeroids, 702, 750, 763, 766, 887 sphaeroid-structures, 751 sphalerite, 71, 72, 73, 83, 96, 97, 119, 120, 121, 122, 123, 124, 125, 135, 136, 137, 138, 139, 140, 141, 142, 148, 149, 150, 157, 158 159, 161, 205, 206, 207, 208, 236, 238, 249, 250, 251, 298, 299, 300, 302, 303, 305, 306, 307, 309, 310, 311, 312, 314, 316, 317, 318, 319, 320, 322, 323, 324, 325, 326, 327, 328, 329, 331, 332, 333, 335, 336, 337, 365, 381, 382, 383, 385, 414, 415, 425, 427, 428, 429, 433, 434, 435, 436, 437, 438, 471, 472, 473, 474, 495, 502, 511, 512, 513, 514, 517, 528, 530, 531, 543, 544, 552, 553, 554, 556, 596, 597, 598, 600, 625, 639, 696, 761, 766, 767, 791, 792, 793, 860 - micrographic-in-shape, 384 - sphaeroids, 414, 415, 766 sphalerite/gangue symplectite, 511, 512 sphene, 872 spinel, 1, 2, 3, 23, 36, 37, 275, 277, 278, 279, 280, 281, 283, 284, 286, 288, 290, 291, 294, 296, 349, 350, 351, 588, 590, 663, 907, 908 - veinlets, 287 strain ellipsoid, 923 stannite, 176, 177, 178, 180, 181, 183, 184, 185, 186, 187, 326, 327, 328, 485, 695 star safflorite, 121 - shaped, 253 - -shaped safflorite, 522 stibarsen, 475, 476, 477, 478, 479, 480 stromatolites, 721 stromeyerite, 66, 549, 550 sub-sphaeroidal sphalerite, 429 sulphides, 579, 872 sylvanite, 172, 392, 510 symplectic bomite, 387 - bornite with bismuthinite, 424 - forms, 126, 438, 484 - intergrowths, 21, 368, 419, 430, 434, 437, 439, 458, 459, 460, 461, 462, 463, 464, 466, 467, 493, 496, 503, 855 - magnetite, 421, 422 - myrmekitic form, 452 - myrmekitic intergrowth, 495 - myrmekitic pattern, 455 - patterns, 411, 412, 413 - petzite, 394 - petzite with silicates, 393 - quartz/pyrite, 497 - replacement, 374 - rutile, 410 - rutile/wiikite intergrowth, 410 - silicates, 486, 487 - spinel, 282

symplectically grown, 383 symplectite, 258 synaeresis, 704, 705, 707, 898 - cracks, 707, 714, 715, 716 synantetic reaction, 24

Table-cloth structure, 277 teallite, 183, 184, 185 tectonic deformation, 794 - effect, 362 tectonogranular, 39, 40, 41 telluride, 406 tennantite, 68, 69, 83, 85, 96, 97, 136, 143, 432 tenorite, 648, 649, 650, 651, 652 tetrahedrite, 92, 93, 94, 193, 194, 219, 240, 263, 384, 458, 459, 461, 514, 519, 520, 568, 715, 716, 900 theralite, 20 titanite, 378 titanium oxides, 875 titanomagnetite, 26, 37, 45, 47, 55, 56, 228, 229, 588, 589, 590 tourmaline needles, 501 trichyte crystallites, 693, 694 trogtalite, 212, 213

Ullmanite, 155, 156 ulvite, 277 ulvospinel, 278 unaffected chromite, 828, 829 undulating extinction, 772, 787 uraninite, 224, 257, 363, 368, 370, 371, 372, 373, 374, 375, 376, 397, 398, 399, 400, 401, 407, 580, 583, 584, 585, 586, 612, 613, 614, 622, 623, 635, 660, 666, 667, 668, 700, 701, 704, 705, 706, 711, 712, 796, 797, 861, 862, 863, 864, 865, 894, 898, 906 - epitactic on columbite, 582, 583, 584, 585 - in graphic-like intergrowth with microcline, 397, 398

uranium minerals secondarily formed, 866, 867 uvarovite, 11, 12

Vanadiferous magnetite, 289, 907, 908 veinlet of galena, 434 - of jacobsite, 269 - of magnetite, 270 - of sphalerite, 517 velonoblast, 537 volcanic line, 924

"Wasserkies", 54 "Wiederverkittungs"-serpentine, 494, 775 wiikite, 410 wire silver, 164 wittichenite, 193, 194, 458, 459, 460, 461 wolframite, 182, 525, 526, 527 wollastonite, 17 wurtzite, 616

X-shaped bodies, 336 - sphalerite bodies, 337 xenoblastic, 917 xenocrysts, 24

"Zerfall" texture, 475, 476, 477 zinkite, 343, 344, 345, 346, 347, 348 zircon, 797, 798, 904 zonal distribution, 606 zone cobaltite, 176 zoned malacon, 620, 621, 622 - smaltite, 561