Mineralogy: An Introduction to Minerals, Rocks, and Mineral Deposits [1st ed.] 9783662573143, 9783662573167

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Mineralogy: An Introduction to Minerals, Rocks, and Mineral Deposits [1st ed.]
 9783662573143, 9783662573167

Table of contents :
Front Matter ....Pages I-XI
Front Matter ....Pages 1-1
Crystals (Martin Okrusch, Hartwig E. Frimmel)....Pages 3-30
Minerals (Martin Okrusch, Hartwig E. Frimmel)....Pages 31-53
Rocks (Martin Okrusch, Hartwig E. Frimmel)....Pages 55-67
Front Matter ....Pages 69-69
Elements (Martin Okrusch, Hartwig E. Frimmel)....Pages 71-86
Sulfides, Arsenides and Complex Sulfides (Sulfosalts) (Martin Okrusch, Hartwig E. Frimmel)....Pages 87-104
Halides (Martin Okrusch, Hartwig E. Frimmel)....Pages 105-109
Oxides and Hydroxides (Martin Okrusch, Hartwig E. Frimmel)....Pages 111-126
Carbonates, Nitrates and Borates (Martin Okrusch, Hartwig E. Frimmel)....Pages 127-140
Sulfates, Chromates, Molybdates, Wolframates (Martin Okrusch, Hartwig E. Frimmel)....Pages 141-149
Phosphates, Arsenates, Vanadates (Martin Okrusch, Hartwig E. Frimmel)....Pages 151-156
Silicates (Martin Okrusch, Hartwig E. Frimmel)....Pages 157-239
Fluid Inclusions in Minerals (Martin Okrusch, Hartwig E. Frimmel)....Pages 241-246
Front Matter ....Pages 247-248
Igneous Rocks (Martin Okrusch, Hartwig E. Frimmel)....Pages 249-274
Volcanism (Martin Okrusch, Hartwig E. Frimmel)....Pages 275-289
Plutonism (Martin Okrusch, Hartwig E. Frimmel)....Pages 291-297
Magma and Lava (Martin Okrusch, Hartwig E. Frimmel)....Pages 299-305
Formation and Evolution of Magmas (Martin Okrusch, Hartwig E. Frimmel)....Pages 307-315
Experiments on Magmatic Model Systems (Martin Okrusch, Hartwig E. Frimmel)....Pages 317-340
The Origin of Basalt (Martin Okrusch, Hartwig E. Frimmel)....Pages 341-346
The Origin of Granite (Martin Okrusch, Hartwig E. Frimmel)....Pages 347-357
Orthomagmatic Mineral Deposits (Martin Okrusch, Hartwig E. Frimmel)....Pages 359-369
Pegmatites (Martin Okrusch, Hartwig E. Frimmel)....Pages 371-377
Hydrothermal Mineral Deposits (Martin Okrusch, Hartwig E. Frimmel)....Pages 379-404
Weathering and Mineral Formation in Soils (Martin Okrusch, Hartwig E. Frimmel)....Pages 405-415
Sediments and Sedimentary Rocks (Martin Okrusch, Hartwig E. Frimmel)....Pages 417-452
Metamorphic Rocks (Martin Okrusch, Hartwig E. Frimmel)....Pages 453-497
Phase Relations and Mineral Reactions in Metamorphic Rocks (Martin Okrusch, Hartwig E. Frimmel)....Pages 499-528
Metamorphic Facies and Facies Series (Martin Okrusch, Hartwig E. Frimmel)....Pages 529-546
Front Matter ....Pages 547-548
Earth’s Interior (Martin Okrusch, Hartwig E. Frimmel)....Pages 549-571
Lunar Rocks and the Moon’s Interior (Martin Okrusch, Hartwig E. Frimmel)....Pages 573-582
Meteorites (Martin Okrusch, Hartwig E. Frimmel)....Pages 583-603
The Planets, Their Satellites and Smaller Planetary Bodies (Martin Okrusch, Hartwig E. Frimmel)....Pages 605-633
Introduction to Geochemistry (Martin Okrusch, Hartwig E. Frimmel)....Pages 635-665
The Genesis of Our Solar System (Martin Okrusch, Hartwig E. Frimmel)....Pages 667-676
Back Matter ....Pages 677-719

Citation preview

Martin Okrusch Hartwig E. Frimmel

Mineralogy An Introduction to Minerals, Rocks, and Mineral Deposits

Springer Textbooks in Earth Sciences, Geography and Environment

The Springer Textbooks series publishes a broad portfolio of textbooks on Earth Sciences, Geography and Environmental Science. Springer textbooks provide comprehensive introductions as well as in-depth knowledge for advanced studies. A clear, reader-friendly layout and features such as end-of-chapter summaries, work examples, exercises, and glossaries help the reader to access the subject. Springer textbooks are essential for students, researchers and applied scientists. More information about this series at 7 http://www.springer.com/series/15201

Martin Okrusch Hartwig E. Frimmel

Mineralogy An Introduction to Minerals, Rocks, and Mineral Deposits

Martin Okrusch Institute of Geography and Geology University of Würzburg Würzburg, Bavaria, Germany

Hartwig E. Frimmel Institute of Geography and Geology University of Würzburg Würzburg, Bavaria, Germany Department of Geological Sciences University of Cape Town Rondebosch, South Africa

ISSN  2510-1307 ISSN  2510-1315  (electronic) Springer Textbooks in Earth Sciences, Geography and Environment ISBN 978-3-662-57314-3 ISBN 978-3-662-57316-7  (eBook) https://doi.org/10.1007/978-3-662-57316-7 Library of Congress Control Number: 2018945875 © Springer-Verlag GmbH Germany, part of Springer Nature 2020 This work is subject to copyright. All rights are reserved by the Publisher, whether the whole or part of the material is concerned, specifically the rights of translation, reprinting, reuse of illustrations, recitation, broadcasting, reproduction on microfilms or in any other physical way, and transmission or information storage and retrieval, electronic adaptation, computer software, or by similar or dissimilar methodology now known or hereafter developed. The use of general descriptive names, registered names, trademarks, service marks, etc. in this publication does not imply, even in the absence of a specific statement, that such names are exempt from the relevant protective laws and regulations and therefore free for general use. The publisher, the authors and the editors are safe to assume that the advice and information in this book are believed to be true and accurate at the date of publication. Neither the publisher nor the authors or the editors give a warranty, expressed or implied, with respect to the material contained herein or for any errors or omissions that may have been made. The publisher remains neutral with regard to jurisdictional claims in published maps and institutional affiliations. Cover photograph: Azurite crystals, partly replaced by malachite, Tsumeb mine, Namibia; width of view ca. 3 cm; collection H.E. Frimmel, photo: K.-P. Kelber. This Springer imprint is published by the registered company Springer-Verlag GmbH, DE part of Springer Nature The registered company address is: Heidelberger Platz 3, 14197 Berlin, Germany

V

Preface Thirty-five years ago Siegfried Matthes, chair of mineralogy at the University of Würzburg at that time, wrote a German book on Mineralogy, designed as introduction to the fields of mineralogy, petrology and mineral deposits. That book, mainly aimed at students of mineralogy and geological sciences in general, quickly became one of the most popular textbooks of its kind amongst German-speaking students. Encouraged by its success, author and publisher (Springer) repeatedly improved the contents in a series of new revised editions. After Matthes’ death in 1999, one of us, Martin Okrusch, who had succeeded Matthes as chair of mineralogy, took up the challenge to continue this tradition and produced a series of new editions of the book—the ninth edition was published in 2014. Over the past years he continuously updated the content and added new chapters, in some cases with the help from other experts, to ensure that the book remained and still remains the German geoscience students’ top choice. One of the reasons for the high popularity of the book has been its comprehensiveness. Many subjects are dealt with to an extent that goes beyond what is typically expected from students at exams and thus the book naturally survives the student days of its owner, when he or she finds it a valuable reference during ­post-graduate employment, be it in academia, as consultant or in industry. In the meantime the baton of the chair of mineralogy had been passed on to the next generation, that is, the junior author of this book, and the invitation was expressed to team up and continue with the by now long tradition of this highly successful textbook. At that stage the idea was born to write the next edition for an English-speaking readership. This resulted on the one hand from the wish to make available this sought-after textbook also to students and interested professionals outside of the German-speaking countries and, on the other hand, from the realisation that even for German-speaking students, it can only be beneficial if they become familiar, already at undergraduate level, with the English professional terminology since English has become the global language in geoscientific communication. Although we tried to refer to places and/or people from all over the world when illustrating certain processes or concepts, the attentive reader might notice some bias towards German examples in

the text. This is admittedly an artefact from taking the German version of the textbook as a blueprint but also brings to the fore the high significance of German speaking scientists in the historic development of mineralogy and petrology as well as the fact that many type localities of rocks, minerals or deposits happen to be located in Germany. Invariably there are numerous books on the topics of crystallography, mineralogy, petrology, geochemistry, mineral deposits etc. What makes this book different is that much of these topics is treated under a single cover. This reflects not only a trend in geoscience training at universities from individual semester courses on these topics to combined courses that deal with all of the above but also the interconnectivity between these traditional fields of geosciences. For example, with the growing public awareness of the need for discovering new mineral resources to maintain the high standards of living we have become used to in the industrialised nations and to enable those in the emerging countries to achieve such standards in the future, an improved knowledge of mineral deposits is required. This cannot be achieved without a sound understanding of mineralogy, petrology and geochemistry. The book is subdivided into four parts. First, some basic crystallographic, mineralogical and petrographic terms and concepts are explained, all of which is necessary to characterise geomaterials, such as minerals, rocks and ores. The second part deals with systematic mineralogy and thus lays the foundation for the third and largest part, the description of the various rock types and mineral deposits. Emphasis is placed on genetic concepts in order to ignite an interest in, and understanding of, processes that shaped and shape our surroundings on this planet as well as those that form rocks of economic value. The final part concerns the chemical make-up not only of planet Earth but also of extraterrestrial material, such as meteorites or lunar rocks, as well as that of the other planets. This leads to a concluding review of our current thinking about the origin of our solar system. The reader is expected to bring along some understanding of physical geology as well as basic knowledge of college-level inorganic ­chemistry. For those who want to go beyond the level of a textbook or look for more detailed background

VI

Preface

information, a bibliography is given at the end of each chapter. These bibliographies are by far not exhaustive but can only be snapshots of the current professional literature. Not only the references cited in the text are listed therein but also references that are particularly useful for follow-up reading on the various topics dealt with. The project of writing such a book cannot be achieved without the help, in various kinds, from numerous people. Several colleagues had contributed to the German version of the textbook and their contributions have been incorporated into this book as well. Major updates and revisions were provided by Hans Ulrich Bambauer (Münster/Ostbevern), Gerd Geyer (Würzburg), Reiner Klemd (Erlangen), Herbert Kroll (Münster), Karl Mannheim (Würzburg), and Dieter Stöffler (Berlin). For constructive criticism and important hints as well as providing new photographs and diagrams fort the English edition thanks are due, in addition to the above, to Christian Hager (Engelschoff), François Holtz (Hannover), Dorothée Kleinschrot (Würzburg), Michael Kleber (Mutmannsreuth), Nikola Koglin (Hannover), Herbert Kroll (Münster), Rob Lavisnky (Dallas, USA), Heike Lehner (Heidelberg), Joachim Lorenz (Karlstein am Main), Karl Mannheim (Würzburg), Vesna Marchig (Hannover), Pete Mouginis-Mark (Honolulu), Andrea Murphy (Adelaide, Australia), Martin Pfleghaar (Heidelberg), Cornelia Schmitt-Riegraf (Münster), Ulrich Schüßler (Würzburg), Hans Adolf Seck† (Köln), Volker von Seckendorff (Würzburg),

Denis Smith (Adelaide, Australia), Wilhelm Stürmer (Erlangen), Ekkehart Tillmanns (Wien), Thomas Will (Würzburg) and Armin Zeh (Karlsruhe). If not otherwise stated, line drawings as well as photographs of samples and thin sections were made by Klaus-Peter Kelber (Würzburg), which is highly appreciated. We also thank Stefan Höhn (Würzburg) for designing the English versions of the line drawings. Our special thanks go to Armin Stasch (Bayreuth) for careful editorial editing and compiling the subject and geographical indices. The endeavour of writing this book would not have happened without the enthusiastic encouragement and support by Annett Büttner from Springer Nature, the production of the book not possible without the excellent work by the entire Springer Nature team, especially J. Viju Falgon and Karthik Raj Selvaraj for the technical editing and lay-out. We would also like to extend our thanks to our wives, Irene Okrusch and Elisabeth Nachtnebel, respectively, for their patience with us while preparing the manuscript. We hope that this book will prove useful and helpful to both students of, and lecturers in, geosciences but also to professionals who need a reliable reference for basic concepts they might have learned at some stage at university as well as to interested mineral collectors. Martin Okrusch Hartwig Frimmel

Würzburg, Germany January 2019

VII

Contents Part I  Introduction and Basic Concepts 1 1.1 1.2 1.3 1.4 1.5

Crystals. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3 Crystal Morphology. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4 Crystal Structures. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 Crystal Chemistry. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 12 Physical Properties of Crystals. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 Optical Crystallography. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29

2 2.1 2.2 2.3 2.4 2.5

Minerals. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 31

3 3.1 3.2 3.3 3.4 3.5 3.6

Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 Mineralogical Composition of Rocks. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 Relationships Between Lithogeochemistry and Mineralogy. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 Rock Fabric . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 Field Relationships. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59 Principal Rock-Forming Processes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 Mineral Deposits. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 64 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67

Definition of the Term Mineral. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32 Identification and Classification of Minerals. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 34 Mode of Occurrence. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35 Rock-Forming and Economic Minerals. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 37 Biomineralisation and Medical Mineralogy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 42 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52

Part II  Systematic Mineralogy 4 4.1 4.2 4.3

Elements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 71 Metals. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 72 Metalloids (Semi-metals). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 78 Non-metals . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 78 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 85

5 5.1 5.2 5.3 5.4 5.5

Sulfides, Arsenides and Complex Sulfides (Sulfosalts). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87

6

Halides . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 105 References. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 109

7 7.1 7.2 7.3

Oxides and Hydroxides . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 111

Metal Sulfides with M:S > 1:1 (Generally 2:1) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 88 Metal Sulfides and Arsenides with M:S = 1:1. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 90 Metal Sulfides, Sulfarsenides and Arsenides with M:S = 1:2. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 96 Arsenic Sulfides. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 102 Complex Metal Sulfides (Sulfosalts). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 102 References. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104

M2O Compounds. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 M3O4 Compounds. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 112 M2O3 Compounds. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 116

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Contents

7.4 7.5

MO2 Compounds. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 120 Hydroxides. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 124 References. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 126

8 8.1 8.2 8.3 8.4 8.5 8.6

Carbonates, Nitrates and Borates. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 127 Calcite Group, 3¯2/m . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 128 Aragonite Group, 2/m2/m2/m. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 132 Dolomite Group. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 135 Azurite-Malachite Group. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 136 Nitrates. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 137 Borates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 138 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 140

9 9.1 9.2 9.3

Sulfates, Chromates, Molybdates, Wolframates. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 141

10

Phosphates, Arsenates, Vanadates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 151 Suggestions for Further Reading . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 155

11 11.1 11.2 11.3 11.4 11.5 11.6

Silicates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 157 Orthosilicates (Nesosilicates). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 160 Disilicates (Sorosilicates). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 169 Ring Silicates (Cyclosilicates) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 172 Chain Silicates (Inosilicates) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 177 Sheet Silicates (Phyllosilicates) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 189 Framework Silicates (Tectosilicates). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 201 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 237

12

Fluid Inclusions in Minerals. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 241

Sulfates. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 142 Chromates. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 147 Molybdates and Wolframates. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 147 Suggestions for Further Reading . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 149

References. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 246

Part III  Petrology and Metallogenesis 13 13.1 13.2

Igneous Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 249 Classification of Igneous Rocks. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 250 Petrography of Igneous Rocks. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 257 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 273

14 14.1 14.2 14.3 14.4 14.5

Volcanism. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 275

15 15.1 15.2 15.3

Plutonism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 291 Volcanic Roots and Magma Chambers. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 292 Shapes of Plutonic and Subvolcanic Intrusive Bodies. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 293 Internal Structure and Emplacement of Intrusive Bodies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 294 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 297

Effusive Volcanism: Lava Flows . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 277 Extrusive Volcanism. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 280 Explosive Volcanism. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 281 Mixed Volcanic Activity: Stratovolcanoes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 286 Volcanic Exhalations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 286 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 289

IX Contents

16 16.1 16.2 16.3 16.4 16.5

Magma and Lava . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 299

17 17.1 17.2 17.3 17.4 17.5

Formation and Evolution of Magmas. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 307 Magma Series. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 308 Primary and Parental Melts. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 308 Magma Mixing. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 310 Magmatic Differentiation. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 311 Assimilation. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 314 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 314

18 18.1 18.2 18.3 18.4 18.5

Experiments on Magmatic Model Systems. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 317 The Gibbs’ Phase Rule. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 318 Experiments in Binary and Ternary Systems. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 319 Bowen’s Reaction Series. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 334 The Basalt Tetrahedron of Yoder and Tilley (1962). . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 337 Equilibrium Melting and Fractionated Melting. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 338 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 339

19 19.1 19.2

The Origin of Basalt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 341

20 20.1 20.2

The Origin of Granite. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 347

21 21.1 21.2 21.3 21.4

Orthomagmatic Mineral Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 359 Introduction. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 360 Mineralisation Due to Fractional Crystallisation. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 360 Mineralisation Due to Liquid Immiscibility. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 364 Carbonatite- and Alkaline-Magmatic Rock-Hosted Deposits. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 367 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 368

22 22.1 22.2 22.3 22.4

Pegmatites. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 371

23 23.1 23.2 23.3 23.4 23.5 23.6 23.7

Hydrothermal Mineral Deposits. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 379 Basic Principles . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 380 Hydrothermal Impregnation Deposits. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 382 Hydrothermal Replacement Deposits. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 386 Hydrothermal Vein-Type Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 388 Volcanogenic-Sedimentary Ore Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 395 Non-magmatic Stratabound Hydrothermal Deposits. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 400 Unconformity-Related Uranium Deposits. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 401 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 402

Chemical Composition and Structure of Magma . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 300 Volcanic Gases. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 300 Temperatures of Magmas. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 301 Viscosity of Magmas and Lavas. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 303 Solubility of Volatiles in Magma . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 304 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 305

Basalt Types and Plate Tectonics. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 342 Formation of Basaltic Melts by Partial Melting of Peridotite in Earth’s Upper Mantle . . . . . . . . . . . . . . . 343 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 346

Petrogenetic Classification of Granitoids Based on Their Chemical Composition . . . . . . . . . . . . . . . . . . . 348 Experiments on the Petrogenesis of Granite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 349 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 357

Theoretical Considerations. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 372 Field Relations, Petrography and Petrogenesis of Pegmatites. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 373 Pegmatites as Sources of Economic Minerals. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 375 Geochemical Classification of Granitic Pegmatites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 376 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 377

X

Contents

24 24.1 24.2 24.3 24.4 24.5 24.6

Weathering and Mineral Formation in Soils . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 405 Mechanical Weathering . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 406 Chemical Weathering. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 406 Subaerial Weathering and Climate Zones. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 409 On the Definition of the Term Soil . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 409 Weathering of Silicate Rocks and Related Deposits. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 410 Weathering of Sulfidic Ore Bodies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 412 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 415

25 25.1 25.2 25.3 25.4 25.5 25.6 25.7

Sediments and Sedimentary Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 417 Basic Principles . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 419 Clastic Sediments and Sedimentary Rocks. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 420 Chemical and Biochemical Sediments and Sedimentary Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 435 Iron- and Manganese-rich Sediments and Sedimentary Rocks. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 441 Siliceous Sediments and Sedimentary Rocks. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 445 Sedimentary Phosphate Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 446 Evaporites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 447 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 451

26 26.1 26.2 26.3 26.4 26.5 26.6

Metamorphic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 453

27 27.1 27.2 27.3 27.4 27.5

Phase Relations and Mineral Reactions in Metamorphic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . 499

28 28.1 28.2 28.3

Metamorphic Facies and Facies Series. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 529 Principles of Metamorphic Facies. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 530 Metamorphic Facies Series . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 531 Mineralogical Characteristics of Individual Metamorphic Facies. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 532 References (see also Chapters 26 and 27) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 545

Basic Principles . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 454 Metamorphism as a Geological Process. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 459 Nomenclature of Regional and Contact Metamorphic Rocks. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 473 Structure and Texture of Metamorphic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 479 Formation of Migmatites by Anatexis. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 487 Metasomatism. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 491 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 495

Mineral Equilibria in Metamorphic Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 500 Metamorphic Mineral Reactions. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 504 Geothermometry and Geobarometry. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 517 Pressure-Temperature Evolution of Metamorphic Complexes. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 520 Graphical Presentation of Metamorphic Mineral Assemblages. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 522 References (see also Chapters 26 and 28) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 527

Part IV  Our Planetary System 29 29.1 29.2 29.3 29.4

Earth’s Interior. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 549

30 30.1 30.2 30.3

Lunar Rocks and the Moon’s Interior . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 573 The Lunar Crust. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 574 Moon’s Internal Layering. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 578 Geological History of the Moon. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 580 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 582

Seismic Evidence of Earth’s Overall Structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 550 The Crust . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 553 The Mantle. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 557 The Core. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 568 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 569

XI Contents

31 31.1 31.2 31.3 31.4 31.5

Meteorites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 583 Fall Phenomena. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 584 Frequency of Falls and Finds. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 587 Classification of Meteorites Derived from the Asteroid Belt. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 588 Planetary Meteorites. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 599 Tektites. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 601 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 601

32 32.1 32.2 32.3 32.4 32.5

The Planets, Their Satellites and Smaller Planetary Bodies. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 605 The Terrestrial Planets. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 608 Asteroids . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 618 The Giant Planets and Their Satellites. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 621 The Trans-Neptune Objects (TNO) in the Kuiper Belt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 630 The Dwarf Planet Pluto and Its Moon Charon: A Double Planet. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 630 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 631

33 33.1 33.2 33.3 33.4 33.5 33.6

Introduction to Geochemistry. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 635 Geochemical Classification of the Elements. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 636 Chemical Composition of the Bulk Earth. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 638 Chemical Composition of the Earth’s Crust. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 640 Trace-Element Partitioning and Magmatic Processes. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 642 Isotope Geochemistry. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 650 The Formation of the Chemical Elements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 663 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 664

34 34.1 34.2 34.3 34.4

The Genesis of Our Solar System. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 667 Earlier Theories . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 668 Formation of Stars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 669 Composition of the Solar Nebula . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 670 Formation of Planets. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 672 References and Suggestions for Further Reading. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 675

Supplementary Information. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 677 Appendix. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 678 A.1 Important Ionic Radii and the Coordination of Cations With O2−. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 678 A.2 Calculation of Mineral Formulae. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 678 Copyright permissions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 681 Geographical Index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 683 Subject Index. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 691

1

Introduction and Basic Concepts Mineralogy Literally, “Mineralogy” means science of minerals. The term “mineral” was coined in the late Middle Ages; it is derived from the Medieval Latin word mina = shaft (minare = to mine). The antique people, e.g., the Greeks and Romans, only used the word “stone”. Special attention was dedicated to precious stones, distinguished by their lustre, colour and hardness, which were highly estimated by the old civilised nations already in pre-Greek times. The treatise De lapidibus (“On stones”) by Theophrastos (371– 287 bc), partly based on lost texts of Aristotle (384–322 bc), provided a wealth of observations and sound considerations on minerals and rocks, and on their practical use. In times of the Roman Empire, Pliny the Elder (ad 23/24–79) wrote his Naturalis historia that summarised the contemporary knowledge about minerals and rocks. Minerals are chemically uniform, natural constituents of the Earth and other planetary bodies, like the Moon, meteorites and the terrestrial planets of our and other solar systems. With few exceptions, minerals are inorganic, solid and crystalline (. Fig. 1.1). According to this extremely general definition that will be explained, in more detail, in 7 Chap. 2, most minerals are also crystals (7 Chap. 1). They are the most important constituents of rocks (7 Chap. 3).

Contents Chapter 1   Crystals – 3 Chapter 2   Minerals – 31 Chapter 3   Rocks – 55

Part I

3

Crystals 1.1  Crystal Morphology – 4 1.1.1  Symmetry Operations and Symmetry Elements – 5 1.1.2  Crystal Systems and Classes – 6 1.1.3  The Law of Rational Indices – 8

1.2  Crystal Structures – 9 1.2.1  Bravais Lattices – 9 1.2.2  Space Groups – 9 1.2.3  Determination of Crystal Structures by X-Ray Diffraction – 11

1.3  Crystal Chemistry – 12 1.3.1  Basic Concepts – 12 1.3.2  Types of Chemical Bonds – 12 1.3.3  Some Important Terms of Crystal Chemistry – 15

1.4  Physical Properties of Crystals – 16 1.4.1  Hardness and Cohesion – 16 1.4.2  Thermal Conductivity – 17 1.4.3  Electrical Properties – 18 1.4.4  Magnetic Properties – 19

1.5  Optical Crystallography – 20 1.5.1  Basic Concepts – 20 1.5.2  Basic Principles of Microscopy in Transmitted Light – 22 1.5.3  Basic Principles of Microscopy in Reflected Light – 28

References and Suggestions for Further Reading – 29

© Springer-Verlag GmbH Germany, part of Springer Nature 2020 M. Okrusch, H. E. Frimmel, Mineralogy, Springer Textbooks in Earth Sciences, Geography and Environment, https://doi.org/10.1007/978-3-662-57316-7_1

1

4

1

Chapter 1 · Crystals

Introduction Crystals (from Greek κρὐσταλλος = ice, applied to rock crystal; . Fig. 1.1) are solid, chemically and structurally homogeneous bodies with a three-dimensional periodic arrangement of their chemical constituents (atoms, ions, molecules). The term crystal is not restricted to the realm of minerals but also includes all synthetic crystalline substances. Not only minerals but nearly all inorganic and many organic solids are crystalline. Man-made, synthetic crystals have a profound influence on our everyday life, from sugar to aspirin, from vibrating quartz in watches to microchips in computers, from laser crystals to catalysts in cars. Crystal structure refers to the internal make-up of crystals in which atoms, ions and/or molecules are periodically arranged to form three-dimensional space lattices. This means that, in a given direction, the smallest possible chemical building block is

repeated by equal distances, the so-called translation distances of lattice points (. Fig. 1.2). Each crystalline substance, including each mineral in the crystalline state, is distinguished by its characteristic, geometrically defined, crystal structure. The three-dimensional regular repetition of basic building blocks results in crystals being homogeneous, both physically and chemically. Homogeneity, in this regard, can be defined more specifically by looking at vectorial, i.e., direction-dependent, physical properties like hardness, cohesion, thermal and electrical conductivity, refraction and birefringence. A body is homogeneous if it displays the same behaviour along a given direction. Nevertheless, crystals are anisotropic, that is, their vectorial properties are different along different directions. Examples will be given in 7 Sect. 1.4. By contrast, in isotropic material, e.g., in glass, the vectorial properties are the same along different directions (. Fig. 1.3).

. Fig. 1.1  Group of quartz crystals, Arkansas, USA, approximately original size

1.1  Crystal Morphology

When crystals can grow unhindered, they develop even crystal faces (. Fig. 1.1) in a geometric arrangement that is dictated by their crystal lattice. Formation of crystal faces is a result of growth rate of crystal faces being anisotropic. A rapidly growing crystal face will become progressively smaller during crystal growth to eventually form an edge or a corner, while slow growth leads to the formation of relatively larger crystal faces (. Fig. 1.4). The combination of all faces of a crystal is designated as crystal form, whereas the size proportions of individual faces define the crystal habit. Crystals of the same form may display different habits, e.g., planar, isometric, columnar, spicular. Both crystal form and crystal habit are dependent on the physico-chemical conditions under which crystal growth takes place, especially temperature, pressure and the chemical composition of the melt or solution, in which the crystal is growing. Because of mutual hindrance during their growth, most crystals are not able to develop their ideal form or at least individual

crystal faces. Therefore, rock-forming minerals rarely show well-developed crystal faces (. Fig. 2.3). As early as 1669 the Danish physician and scientist Niels Stensen (Latinised Nicolaus Steno, 1638–1686) recognised a fundamental law that governs the geometric relationship between crystal faces, the so-called Steno’s Law of Constancy of Angles: In all crystals of the same substance, the angles between corresponding faces, measured at the same temperature and pressure, have a constant value.

This law applies even to highly distorted crystals (. Fig. 1.5). A variety of characteristic symmetry operations can be applied to crystals with regard to both their exterior appearance (i.e., crystal forms) and their spatial distribution of vectorial physical properties. These reflect the symmetrical arrangement of the chemical constituents in the crystal structure.

5

1.1 · Crystal Morphology

. Fig. 1.2  Three-dimensional crystal lattice of triclinic symmetry. The translation vectors a, b, c are of different length and are oriented at oblique angles to each other. Other translation vectors are, e.g., the face diagonals of the faces ab, bc, ac, or the space diagonals of the faces abc. The unit face is indicated in blue (modified after Kleber et al. 2010; courtesy J. Bohm and H. Klimm, Berlin, Germany)

. Fig. 1.4  Relative growth rate of different crystal faces of alunite, KAl(SO4)2·12H2O, starting from a polished sphere. The faces of the crystal forms {110}, {221} and {112} grow rapidly and thus become progressively smaller with on-going crystal growth, while the cubic {100} and octahedral {111} faces grow more slowly and become more and more dominant during crystal growth until only the form {111} remains; (after Spangenberg 1935, in Kleber et al. 2010; courtesy J. Bohm and H. Klimm, Berlin, Germany)

Definition Symmetry is the regular repetition of a motif, e.g., a structural element like a lattice point or a crystal face.

1.1.1  Symmetry Operations and Symmetry

Elements

. Fig. 1.3  Schematic diagrams illustrating the difference between anisotropic and isotropic material: a in anisotropic material physical properties are of identical magnitude along parallel directions but differ along all other directions; b in isotropic material the magnitude of physical properties is identical in all directions

. Fig. 1.5  Two- and threedimensional drawings of quartz crystals illustrating the effects of distortion on the absolute size of a given crystal face and the constancy of angles between corresponding faces; a–d projection along the c-axis; e–h parallel projections (from Ramdohr and Strunz 1978)

When describing the morphology of crystals, the following symmetry operations can be applied. The respective elements of symmetry are denoted by simple symbols (. Fig. 1.6): 5 rotation around a onefold, twofold, threefold, fourfold or sixfold axis of rotation (1, 2, 3, 4, 6)

1

6

Chapter 1 · Crystals

The coding of the 32 crystal classes follows the internationally accepted Hermann-Mauguin system (Hermann 1935; Hahn 2002) that is based on the combination of symbols for the symmetry elements. If the combination of two elements results in a third one, this can be omitted leading to a simplified notation. In the following, the HermannMauguin system will be shortly explained and exemplified by minerals, important ones being indicated by a star (*). For minerals otherwise not mentioned in the text, the chemical formula is given. In addition, we add the traditional names coined by Paul von Groth (1843–1927) after the respective determinant general crystal form. This is defined as a set of crystal faces, situated in a general, not a special position and related to each other by one or more symmetry operations.

1

For more details, the reader is referred to textbooks on crystallography (e.g., Buerger 1971, 1977; McKie and McKie 1990; Bloss 1994). . Fig. 1.6  Simple crystal forms generated by various symmetry ¯ pinacoid operations applied to a given face: a centre of symmetry 1: (two parallel faces); b mirror plane m: dome (Greek: δώμα = roof); c–f 2-, 3-, 4-, and 6-fold axes of rotation: c sphenoid (Greek: wedge), d trigonal pyramid, e tetragonal pyramid, f hexagonal pyramid

(= means equal, ≠ unequal)

5 reflection on a mirror plane (m)   5 inversion, i.e., reflection on a symmetry centre 1¯ 5 rotary inversion (rotoinversion), i.e., combination of rotation and inversion on a twofold, threefold, fourfold or sixfold rotary inversion axis (2¯  = m, 3¯, 4¯ , 6¯ ), whereby the two symmetry operations are applied simultaneously

Note that this definition applies to almost all cases but exceptions do exist. Even a triclinic crystal may have axes with equal unit lengths, e.g., a = b, or the same angles between axes, e.g., α = γ. The deciding factor for a crystal system is the combination of the symmetry elements present. For instance the dices with a = b = c, α = β = γ do not have a cubic but a triclinic symmetry, because—owing to the different numbers of points on each face—they possess no symmetry element at all. The same may apply to the other non-cubic crystal systems.

1.1.2  Crystal Systems and Classes

The combination of the above symmetry elements leads to a total of 32 crystal classes, as recognised as early as 1830 by Johann Friedrich Christian Hessel (1796–1872). These have been useful in describing the external shape of crystals. Within a crystal structure, however, additional symmetry operations are possible, that is, translation of individual lattice points. The combination of translation with rotation and reflection causes glide reflections and screw axes. Combining all of these symmetry operations gives rise to as many as 230 space groups, independently derived in 1891 by Arthur Moritz Schoenflies (1853–1928) and Efgraf Stepanowich Fedorov (1853–1919). For the mathematical description of crystal forms and crystal structures, especially the position of lattice points, point lines and lattice planes, different coordinate systems are used that are fitted to the respective symmetries. Altogether, seven such coordinate systems are distinguished, also known as the 7 crystal systems. They differ in the aspect ratios of their principal axes a, b, c, and the angles between theses axes α (between b and c), β (between a and c) and γ (between a and b). In the trigonal, tetragonal and hexagonal crystal systems, the units along the axes a and b have the same length and, therefore, are denoted a1, a2 and a3. An analogous notation is used for the cubic system in which the units along all axes have the same length.

z Triclinic Crystal System generally a ≠ b ≠ c, α ≠ β ≠ γ

5 1, trcl.-pedial: asymmetric. Examples: aramayoite, Ag(Sb,Bi)S2, lazurite-4A (7 Sect. 11.6.3). ¯ , trcl.-pinacoidal: symmetry centre. Examples are most 5 *1 important minerals like plagioclase, microcline (7 Sect. 11.6.2), kyanite (. Fig. 1.20, 7 Sect. 11.1).

z Monoclinic Crystal System generally a ≠ b ≠ c, α = γ = 90°, β > 90° 5 2, mcl.-sphenoidal: one twofold rotation axis

parallel to b (||b). Examples: clinotobermorite, Ca5(Si3O8OH)2·2H2O, cane sugar, C12H22O11, an economically important synthetic product. 5 m, mcl.-domatic: one mirror plane perpendicular to b (⊥b). Example: the zeolite mineral scolecite, CaAl2Si3O10·3H2O. 5 *2/m, mcl.-prismatic: one twofold rotation axis ||b and a mirror plane ⊥b. Examples: many important minerals like sanidine and orthoclase (7 Sect. 11.6.2), clinopyroxenes (7 Sect. 11.4.1), clinoamphiboles (7 Sect. 11.4.3), micas (7 Sect. 11.5.2), titanite (7 Sect. 11.1), and gypsum (7 Sect. 9.1).

z Orthorhombic Crystal System generally a ≠ b ≠ c,  α = β = γ = 90° 5 mm2, o.rh.-pyramidal: one twofold rotation axis

parallel to two mirror planes that are ⊥ to each other and intersect in c. Examples: hemimorphite, Zn4Si2O7(OH)2·H2O, bournonite, PbCuSbS3, enargite (7 Sect. 5.2).

1.1 · Crystal Morphology

5 222, o.rh.-disphenoidal: three twofold rotation axes ||a, b, c that are ⊥ to each other. Example: epsomite, MgSO4·7H2O. 5 *2/m2/m2/m (mmm), o.rh.-dipyramidal: three mirror planes, ⊥ to each other, intersect to form three twofold rotation axes ||a, b, and c, respectively. Examples are manifold among important minerals like olivine, andalusite, sillimanite, topaz (7 Sect. 11.1), orthopyroxenes (7 Sect. 11.4.1), orthoamphiboles (7 Sect. 11.4.3), anhydrite, baryte (7 Sect. 9.1), and aragonite (7 Sect. 8.2). z Tetragonal Crystal System generally a1 = a2 ≠ c, α = β = γ = 90°

The principal axes c is a fourfold rotation axis or a rotary inversion axis; ⊥ to c are the two secondary rotation axes a1 and a2. 5 4, tetr.-pyramidal: one fourfold rotation axis c. The only mineral example is the borate mineral pinnoite, MgB2O(OH)6. 5 4¯ , tetr.-disphenoidal: one fourfold rotary inversion axis in c. Example: the meteorite mineral schreibersite (Fe,Ni)3P. 5 4/m, tetr.-dipyramidal: one fourfold rotation axis in c, with a mirror plane ⊥c. Examples: scheelite (7 Sect. 9.3), leucite (low) (7 Sect. 11.6.3), scapolite group (7 Sect. 11.6.5). 5 4 mm, ditetr.-pyramidal: The fourfold rotation axis along c is combined with two mirror planes ⊥ to a1 and a2; this configuration leads to additional mirror planes ⊥ to the bisectors between a1 and a2. Example: diaboleite, Pb2Cu(OH)4Cl2. ¯ , tetr.-skalenohedral: In the principal fourfold 5 *42m inversion rotational axis c two mirror planes intersect both of which are ⊥ to each other; the bisectors between them are the twofold secondary rotation axes a1 and a2. Examples are important minerals like chalcopyrite (7 Sect. 5.2), stannite Cu2FeSnS4, melilite (7 Sect. 11.2). 5 422, tetr.-trapezohedral: ⊥ to the principal fourfold rotation axis along c, 2 + 2 twofold secondary rotation axes are present, two of them ||a1 and a2, the other two || to their bisectors. Example: (low) α-cristobalite, SiO2 (7 Sect. 11.6.1). 5 *4/m2/m2/m (4/mmm): ditetr.-dipyramidal: A mirror plane is ⊥ to the principal fourfold rotation axis c; 2 + 2 additional mirror planes intersect in c, two of them ⊥ to the secondary axes a1 and a2, two others ⊥ to the bisectors between them. This configuration leads to 2 + 2 twofold secondary axes of rotation, two of them ||a1 and a2, the other two || to their bisectors. Examples: several important minerals like rutile, TiO2 (7 Sect. 7.4), anatase, TiO2, stishovite, SiO2 (7 Sect. 11.6.1), cassiterite, SnO2 (7 Sect. 7.4), zircon (7 Sect. 11.1), vesuvianite (7 Sect. 11.2). z Trigonal Crystal System generally a1 = a2 = a3 ≠ c, α = β = 90°, γ = 120°

The principal axis c is a threefold rotation axis or a rotary inversion axis; three twofold secondary rotation axes a1, a2, a3 are ⊥ to c. The trigonal and the hexagonal

7

crystal systems are closely related to each other. In addition to this so-called hexagonal setting a so-called rhombohedral setting can also be chosen in the trigonal system with a1 = a2 = a3, α = β = γ. 5 3, trig.-pyramidal: one threefold rotation axis c. Example: carlinite Tl2S. 5 *3¯ , (trig.-)rhombohedral: one threefold inversion rotational axis c. Examples: ilmenite (7 Sect. 7.3), dolomite (7 Sect. 8.3), phenakite, Be2[SiO4], dioptase (7 Sect. 11.3). 5 *3m, ditrig.-pyramidal: three mirror planes, positioned ⊥ to a1, a2, a3, intersect in the principal threefold rotation axis c. Examples: millerite, NiS, proustite and pyrargyrite (7 Sect. 5.5), tourmaline (7 Sect. 11.3). 5 *3¯ 2/m (3¯ m), ditrig.-scalenohedral: three mirror planes, positioned ⊥ to a1, a2, a3, intersect in the principal threefold inversion rotational axis along c; this configuration leads to three twofold secondary rotation axes ||a1, a2 and a3. Examples: many important minerals like calcite (7 Sect. 8.1), corundum, haematite (7 Sect. 7.3), brucite, Mg(OH)2, native (elemental) bismuth, native antimony and native arsenic (7 Sect. 4.2). 5 *32, trig.-trapezohedral: three twofold secondary rotation axes a1, a2, a3, positioned ⊥ to the principal threefold rotation axis c. Examples: of paramount importance is the mineral (low) α-quartz (7 Sect. 11.6.1); other examples are cinnabarite (7 Sect. 5.2), native selenium Se and native tellurium Te. z Hexagonal Crystal System generally a1 = a2 = a3 ≠ c, α = β = 90°, γ = 120°

The principal axis c is a sixfold rotation axis or a rotary inversion axis; similarly as with the trigonal crystal system, three twofold secondary rotation axes a1, a2, a3 are ⊥ to c. 5 *6, hex.-pyramidal: one sixfold rotation axis c. Examples: nepheline (7 Sect. 11.6.3), cancrinite (7 Sect. 11.6.4). 5 6¯ (= 3/m), trig.-dipyramidal: one sixfold inversion rotational axis c, identical to the combination of a threefold rotation axis with a mirror plane ⊥ to it. Example: laurelite, Pb7F12Cl2. 5 *6/m, hex.-dipyramidal: a mirror plane ⊥ to the sixfold rotation axis c. The most important example is apatite (7 Chap. 10), other examples are pyromorphite, mimetite and vanadinite (7 Chap. 10). 5 6 mm, dihex.-pyramidal: In the principal sixfold rotation axis c, 3 + 3 mirror planes intersect all of which are positioned ⊥ to a1, a2, a3 and ⊥ to the bisectors between them. Examples: wurtzite (7 Sect. 5.2), greenockite, CdS, zincite, ZnO. 5 6¯ m2, ditrig.-dipyramidal: Three vertical mirror planes, situated ⊥ to a1, a2, a3, intersect in the principal sixfold rotary inversion axis along c; this configuration leads to three twofold secondary rotation axes, which are situated in the mirror planes and form the bisectors between a1, a2 and a3. Examples: bastnäsite, (Ce,La,Y)CO3F, benitoite, BaTi[Si3O9]. 5 *622, hex.-trapezohedral: 3 + 3 twofold secondary rotation axes ⊥ to the principal sixfold rotation axis c and

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Chapter 1 · Crystals

||a1, a2, a3 and the bisectors between them. Examples: (high) β-quartz (7 Sect. 11.6.1), kaliophilite, K[AlSiO4]. 5 *6/m2/m2/m (6/mmm), dihex.-dipyramidal: A mirror plane is situated ⊥ to the principal sixfold rotation axis along c; 3 + 3 additional mirror planes intersect in c, three of them ⊥ to the secondary rotation axes a1, a2, a3, three others ⊥ to the bisectors between them. This configuration leads to 3 + 3 twofold secondary rotation axes, three of them ||a1, a2, a3, the other three || to their bisectors. Examples include beryl (7 Sect. 11.3), molybdenite-2H (7 Sect. 5.3) and graphite-2H (7 Sect. 4.3). z Cubic Crystal System generally a1 = a2 = a3, α1 = α2 = α3 = 90°

A common characteristic of all five cubic crystal classes are threefold rotation or inversion rotational axes, oriented || to the space diagonals of the cube (SD) and listed in the second place of the Hermann-Mauguin notation, whereas the fourfold or twofold rotation or inversion rotational axes ||a1, a2, a3 as well as the mirror planes (m) ⊥ to them are listed in the first place. The third place take the twofold rotation axes || to the face diagonal of the cube (FD) and the mirror planes ⊥ to them. 5 23, cubic-tetartoidal (don’t confuse with the trigonal crystal class 32!): three twofold rotation axes ||a1, a2, a3 and four threefold rotation axes ||SD. Examples: ullmanite, NiSbS, gersdorffite, NiAsS, langbeinite, K2Mg2(SO4)3. 5 *2/m3¯ (m3¯ ), cubic-disdodecahedral: The combination of three mirror planes ⊥ to a1, a2, a3 and four threefold inversion rotational axes ||SD results in three twofold rotation axes ||a1, a2, a3. Examples: important ore minerals like pyrite, skutterudite (7 Sect. 5.3), sperrylite, PtAs2. 5 *4¯ 3m, cubic-hexakistetrahedral: three fourfold inversion rotational axes ||a1, a2, a3, four threefold rotation axes ||SD, six mirror planes ⊥FD. Examples: sphalerite (7 Sect. 5.2), tetrahedrite and tennantite (7 Sect. 5.5), β-boracite, β-Mg3B7O13Cl, sodalite (7 Sect. 11.6.3). 5 432, cubic-gyroidal: three fourfold rotation axes ||a1, a2, a3, four threefold rotation axes ||SD, six twofold rotation axes ||FD. Example: petzite, Ag3AuTe2. 5 *4/m3¯ 2/m (m3m), cubic-hexakisoctahedral: three mirror planes ⊥a1, a2, a3, four threefold inversion rotational axes ||SD, six mirror planes ⊥FD. This configuration leads to three fourfold rotation axes ||a1, a2, a3 and six twofold rotation axes ||FD. Examples include many native metals, such as copper, silver, gold and the platinum-group metals (7 Sect. 4.1), and other minerals, such as diamond (7 Sect. 4.3), argentite (7 Sect. 5.1), galena (7 Sect. 5.2), halite, fluorite (7 Chap. 6), periclase, MgO, uraninite (7 Sect. 7.4), spinel, magnetite, chromite (7 Sect. 7.2), and the garnet group (7 Sect. 11.1). 1.1.3  The Law of Rational Indices

The possibly most important law in crystallography is the Law of Rational Indices, first recognised by Abbé René-Just Haüy (1743–1822), the founder of crystallography, and it

serves as basis for quantitatively describing the position of crystal faces on a crystal polyhedron as well as that of lattice planes within a crystal structure. As already alluded to in the definition of the seven crystal classes, any three straight lines that are not in the same plane and parallel to actual or possible edges of a crystal can be chosen as reference axes, the crystallographic axes. For convenience these are chosen to be parallel to prominent axis of symmetry when possible. In a next step, a unit plane is chosen to define the units of measurement to be used when measuring along each of the crystallographic axes. This plane, which can be any plane that is parallel to a crystal face but not parallel to any of the crystallographic axes, is known as the parametral plane because it defines the units or parameters for the crystal. The intercepts made on the crystallographic axes by any other face can be expressed as a/h, b/k, c/l where h, k, l are simple rational numbers or zero. The Law of Rational Indices states that the indices of any face thus defined are always rational. Since William Hallowes Miller (1801–1880) popularised this notation, it became known as the Millerian Symbol or Miller Indices of a crystal face or lattice plane. The index of a particular face (or lattice plane) is then determined by dividing the intercepts made by the face (or lattice plane) by the standard parametral intercepts a, b, c and then multiplying the result, if necessary, to clear of fractions. Applying this principle to the parametral plane itself results in a/a, b/b, c/c, i.e., 111, which is the Millerian Symbol of the parametral plane in any crystal, irrespective of its symmetry (. Fig. 1.7). The Miller Indices, generally hkl, are enclosed in brackets when describing a crystal face, whereas the unbracketed symbol is used to denote a set of parallel planes within the crystal structure. A crystal face that only cuts the a-axis and is parallel to b and c (and thus intersects them at infinity), for example, has a Millerian Symbol of (100), resulting from the reciprocal of 1∞∞. By analogy, faces that only intersect the b- or c-axis are described by the Miller indices (010) and (001), respectively (. Fig. 1.7). Faces or planes parallel to c, that cut the a- and b-axes at the same axes intercepts (which are related to the crystallographic axial ratio defined by the unit plane), are expressed by the Millerian Symbol (110), faces that cut a at a single unit and b at a double unit have Miller Indices (210) as this is the reciprocal of 12∞. Negative axis intercepts are indicated by a bar (–) on top of the number. For the general indication of faces or planes, the indices (hkl), (hk0), (h0l) and (0kl) are used. In the trigonal and hexagonal crystal systems, the quadrinomial Bravais-Miller indices (hkil) are applied, where i = –(h + k), e.g., (101¯0) for the hexagonal prism face of quartz. Miller indices not related to a single crystal face but to the whole crystal form, i.e., to a set of equivalent crystal faces, are written between braces, e.g., {101¯0} for the hexagonal prism as observed on quartz crystals. Every crystalline mineral and synthetic crystalline material has a characteristic size and shape of its unit cell, the smallest building block for its crystal lattice. It is defined by the lattice constants or lattice parameters, which comprise the intercepts at the three crystallographic axes a, b, c, and the angles at which the three crystallographic axes intersect each other (as dictated by the crystal system to which

9

1.2 · Crystal Structures

. Fig. 1.7 Miller indices for some important crystal faces in an orthorhombic system of coordinates. Out of the general face positions (hkl), a unit face (111) is selected

the material in question belongs). By convention, the intercept on the b-axis is set as equal to one, and the axial ratio is then a:1:c. In the past, this most fundamental characteristic of a given crystalline substance could be calculated from the combination of Miller indices of a series of crystal faces. Today, the axial ratios are obtained directly from X-ray diffractometric studies of the crystal structure (see 7 Sect. 1.2.3). Using the orthorhombic topaz as an example, a crystallographic axial ratio of a:b:c = 0.528:1:0.955 is obtained from its crystal faces. Structure analysis by X-ray diffraction yielded the following lattice constants for the topaz structure: a0 = 4.65  Å, b0 = 8.80  Å, c0 = 8.40  Å (1 Å = 10−8 cm); a0:b0:c0 = 4.65:8.80:8.40 = 0.528:1:0.955. A zone describes a specific direction that is common to a set of crystal faces or lattice planes (h1k1l1), (h2k2l2), (h3k3l3). Faces that belong to one zone are tautozonal. The Miller indices for a zone [uvw] are placed between square brackets. For indexing crystal zones, the zone equation hu + kv + lw = 0 applies. For instance, the faces of the ¯ ) all belong to the zone ¯ ) and (010 tetragonal prism (100), (010), (100 ¯ ) [001] that is orientated parallel to the c-axis. The faces (100) and (100 also belong to the zone [010] that is parallel to b, whereas the faces ¯ ) belong to the same zone [100] parallel to a. (010) and (010

For a more in-depth explanation of the indication procedures, the reader is referred to the textbook of Klein (1989) who presents useful exercises. 1.2  Crystal Structures 1.2.1  Bravais Lattices

It has been established above that the external crystal form is intrinsically related to the internal crystal structure. The fundamental difference between the crystal form and crystal

structure lies in the fact that translation features as additional symmetry operation in crystal structure. In crystal morphology, translation is not visible, because the translation distances are on the order of magnitude of Ångströms (1 Å = 10−8 cm). As was shown in 1842 by Moritz Ludwig Frankenheim (1801–1869) and in 1850 by August Bravais (1811–1863), 14 different translation groups can be deduced, called Bravais lattices (. Fig. 1.8). These may be primitive (P), body-centred (I), side-centred (A, B or C), face-centred (F) or rhombohedral (R). The Bravais lattices belong to six crystal families that conform to the morphological crystal systems: 5 triclinic (“three times inclined” or anorthic), abbreviated a 5 monoclinic (“one times inclined”), abbreviated m 5 (ortho-)rhombic, abbreviated o 5 tetragonal, abbreviated t 5 hexagonal, abbreviated h 5 cubic, abbreviated c Note that both the hexagonal and trigonal crystal systems belong to the hexagonal crystal family. In the respective crystal system, each Bravais lattice has the highest possible symmetry. The elementary cells of the 14 Bravais lattices are shown in . Fig. 1.8 and listed in . Table 1.1. 1.2.2  Space Groups

Combining twofold, threefold, fourfold and sixfold rotation axes 2, 3, 4 and 6, with different extents of translation that is directed along the translational vector τ, results in twofold, threefold, fourfold and sixfold screw axes 21, 31, 32, 41, 42, 43, 61, 62, 63, 64, 65, respectively. As an example, a sixfold

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Chapter 1 · Crystals

. Fig. 1.8  The 14 translation groups of crystals (Bravais lattices) and their symmetries (modified after Ramdohr and Strunz 1978)

. Table 1.1  Elementary cells of the 14 Bravais lattices aP

Triclinic primitive lattice

a, b, c and α, β, γ optional In most cases a ≠ b ≠ c  and α ≠ β ≠ γ  

mP

Monoclinic primitive lattice

a, b, c optional; α = γ = 90°, β optional

mC

Monoclinic base face-centered lattice

In most cases a ≠ b ≠ c  and β ≠ 90°

oP

Rhombic primitive lattice

a, b, c optional; α = β =  γ = 90°

oI

Rhombic space-centered lattice

In most cases a ≠ b ≠ c

oC

Rhombic base face-centered lattice

In most cases a ≠ b ≠ c

oF

Rhombic face-centered lattice

In most cases a ≠ b ≠ c

tP

Tetragonal primitive lattice

a = b, c optional; α = β =  γ = 90°

tI

Tetragonal space-centered lattice

In most cases c ≠ a, b; (a ≡ a1, b ≡ a2)

hP

Hexagonal primitive lattice

a = b; c optional; α = β = 90°, γ = 120°

hR

Hexagonal rhombohedral lattice

In most cases c ≠ a, b; (a = a1, b = a2), if a rhombohedral cell in the trigonal system is described in a hexagonal unit cell

cP

Cubic primitive lattice

a = b = c; α = β = γ = 90°; (a ≡ a1, b ≡ a2, c ≡ a3)

cI

Cubic space-centered lattice

a = b = c; α = β = γ = 90°; (a ≡ a1, b ≡ a2, c ≡ a3)

cF

Cubic face-centered lattice

a = b = c; α = β = γ = 90°; (a ≡ a1, b ≡ a2, c ≡ a3)

screw axis 61 is shown in . Fig. 1.9a, where the lattice points 1, 2, 3, … are arranged like in a spiral staircase. After each rotation at an angle of ε = 60°, the extent of translation equals τ/6; on translation along the c-axis this is 1/6c0. In . Fig. 1.9b, the screw axes 31 and 32 are shown, where a lattice point is translated by 1/3c0 and 2/3c0 after each rotation at an angle σ = 120°. It is readily recognised that both screw axes are antimer (mirror-inverted) to each other: a lefthanded screw is formed by 31, a right-handed one by 32. A prominent example for an antimer crystal is α-quartz. Note,

however, that left-handed quartz has the screw axis 32, righthanded quartz 31! This—apparently opposite—denotion was initially derived from the external crystal form and is based on the position of the trigonal trapezohedral faces {516¯ 1} (. Fig. 11.47b, c), at a time when determination of crystal structure by X-ray diffraction was not possible yet. Combining a mirror plane with translation at an extent of a half lattice constant a0/2, b0/2 or c0/2 parallel to the a, b or c axes, results in glide planes a, b or c, respectively (. Fig. 1.10). The symbol n describes a glide plane with

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1.2 · Crystal Structures

the symmetry elements, e.g., P1¯ in the crystal class 1¯, P2/m, P21/m, C2/m, P2/c, P21/c, C2/c in the crystal class 2/m etc. Examples are P312 and P322 for trigonal (low) α-quartz, P6222 and P6422 for hexagonal (high) β-quartz (. Fig. 11.42), C2/m for monoclinic sanidine and P1¯ for the triclinic plagioclases (. Fig. 11.65c, d). 1.2.3  Determination of Crystal Structures

by X-Ray Diffraction

. Fig. 1.9  Modes of operation of screw axes. a Sixfold screw axis 61 with a rotation angle ε = 60° and translation parallel to the c-axis with 1/6c0. b Threefold screw axes with the rotation angle ε = 120° and translation parallel the c-axis with 1/3c0 (31) = left-handed screw or 2/3c0 (32) = right-handed screw, respectively. Both screw axes are antimer (mirror-inverted) to the mirror plane m (after Borchardt-Ott 2008)

. Fig. 1.10  Mode of operation a of a mirror plane m = (010) and b a glide plane c = (010) with the gliding component ½c0 (after BorchardtOtt 2008)

diagonal gliding components, i.e., (a0 + b0)/2, (a0 + c0)/2 or (b0 + c0)/2; glide planes denoted with symbol d involve the gliding components (a0 + b0)/4, (a0 + c0)/4 or (b0 + c0)/4. Using mathematical methods, Federov and Schoenflies demonstrated that the combination of the 14 translation groups with all possible symmetry operations like rotation, mirror reflection, inversion, rotary inversion (rotoinversion), screw operation and glide operation leads to 230 possibilities, known as the 230 space groups. A space group describes all the possible symmetry operations in a crystal structure or a group of symmetry operations including the translation of lattice points.

The space-group symbols after Hermann-Mauguin comprise the symbols for the Bravais lattices P, I, C, R and

We have seen that in crystal structures, atoms, ions or groups of molecules are arranged to form lattice planes that are repeated periodically by specific, constant distances (. Fig. 1.2). The extents of translation along the crystallographic axes a, b and c are called lattice constants a0, b0, c0. The angles between these axes are denoted α, β, γ. The distances between the lattice planes and the lattice constants are typically in the range of several Ångströms to tens of Ångströms (1 Å = 10−8 cm). This had remained unknown until Max von Laue (1879–1960), assuming that the wavelengths of X-rays are of the same order of magnitude, proved by means of X-ray diffraction experiments that crystals act as diffraction gratings when bombarded by X-rays. With the help from Walter Friedrich and Paul Knipping, the experiments were performed on a sphalerite crystal (Friedrich et al. 1912). It did not take long until it was realised that all crystalline matter behaves in this way. Thus von Laue managed in a single experiment to prove the periodicity of the internal structure of crystals and the electromagnetic wave nature of X-rays. In this context, it should be recalled that the wavelengths of the visible light were conventionally determined using optical diffraction gratings. When radiating a crystal with X-rays, the primary X-ray is diffracted at the lattice planes in various directions and interference takes place. The diffracted waves hit a photographic plate or film and produce density spots forming an interference pattern. On a so-called Laue diagram thus generated, the symmetry of the crystal is clearly visible, provided the crystal was properly orientated, that is, it was radiated parallel to a rotation axis (. Fig. 1.11). These experiments laid the foundation for determining: 5 the symmetry 5 the space group, and 5 the fine structure of crystals The relations between the wavelength of the X-rays λ, the distance of lattice planes d and the diffraction angle θ were formulated by William H. Bragg (1862–1942) and his son William L. Bragg (1890–1971) in a simple equation, known as Bragg Equation or Bragg’s Law:

n = 2d sin θ

(1.1)

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Chapter 1 · Crystals

1

. Fig. 1.11  The first Laue diagrams that were produced in 1912 by Friedrich, Knipping and Laue on sphalerite: a parallel to the fourfold rotoinversion axis, b parallel to the threefold rotation axis of the cubic crystal (crystal class 4/3m). The distribution of interference patterns clearly demonstrates the respective symmetries (after Friedrich et al. 1912)

In both cases, the interference pattern can be documented as diffraction spots or lines on a film. Alternatively, the X-ray interferences are recorded by a Geiger-Müller counter and appear as intensity peaks on a so-called powder diffractogram (. Fig. 1.13b). The different X-ray diffraction methods are used to determine the crystal structure of minerals or other crystalline substances. For instance, the periodical distribution of electron density can be determined by mathematical Fourier or direct methods and, thus, the position of the lattice points of atoms, ions or groups of molecules can be established (. Fig. 1.14a, b). For practical purposes in mineralogy and material sciences, X-ray powder methods, especially X-ray powder diffractometry, are of paramount importance. They can be used for rapid and simple identification of minerals and other crystalline substances and to determine their lattice constants, even in very fine-grained samples. Moreover, X-ray powder methods are suitable to identify the constituents of rocks or mixtures of technical products. 1.3  Crystal Chemistry 1.3.1  Basic Concepts

In crystal structures, the individual components like atoms, ions and groups of molecules interact with each other. The kind and strength of the interaction forces are fundamentally related to the kind of bonding, which in turn determines the physical properties of the crystals. As a first approximation, atoms and ions can be regarded as ideal spheres that are arranged to form closest packings of spheres. Following Victor Moritz Goldschmidt (1888–1947) and Fritz Laves (1906–1978), the following three principles of order are applicable to these packings. . Fig. 1.12  Single-crystal diagram (precession method) of beryl, parallel to the sixfold rotation axis of the hexagonal crystal (crystal class 6/m2/m2/m) (from Buerger 1977)

where n is an integer variable, the order of interference. This important basic equation tells us that 5 at a constant diffraction angle θ, the wavelength of X-rays λ is variable (“white X-radiation”), yielding a Laue diagram (. Fig. 1.11), 5 at constant wavelength of X-rays λ (monochromatic X-radiation) the diffraction angle θ is variable. Basically, the variation of θ can be achieved by two different methods: 5 During bombardement with X-rays, the crystal is rotated, yielding a rotating crystal or a precession diagram (. Fig. 1.12), 5 or a great number of tiny crystals of random orientation are radiated, simultaneously satisfying the Bragg Equation [1.1] and yielding a powder or Debye-Scherrer diagram (. Fig. 1.13a).

In crystal structures, the individual components strive for an order that achieves 5 the densest possible volume ratio (principle of closest packing), 5 the highest possible symmetry (principle of symmetry), and 5 the highest possible coordination, by which the highest possible number of components interact with each other (principle of interaction).

1.3.2  Types of Chemical Bonds

In principle, chemical bonding is based on the interaction between the valence electrons in the outermost valence shells of the atoms. Four principal types of bonds are distinguished. In most crystal structures, these do not occur in isolation but in various combinations, thus giving rise to the wide range in physical properties observed in crystalline materials.

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1.3 · Crystal Chemistry

. Fig. 1.13  a X-ray powder photograph of halite NaCl. b X-ray powder diffractogram of α-quartz SiO2

both atoms to achieve a noble-gas electron configuration. Thus ionic crystals consist of positively charged cations and negatively charged anions that attract each other by electrostatic forces and, in general, are of different size. Following the Coulomb Law

K=

. Fig. 1.14  Distribution of electron density in the crystal structures of a halite NaCl, projection on the (100) plane, and b diamond C, pro¯ jection on the (110) plane. In the halite structure, there is no overlap between the electron shells of the Na+ and Cl− ions (heteropolar bond). In contrast, in the diamond structure, a significant overlap is recorded between the electron shells (atomic bond) (courtesy of Armin Kirfel, University of Bonn, Germany)

z Ionic Bond (Heteropolar Bond)

Ionic bonds are formed by transfer of valence electrons between atoms. One or more electrons from one atom are removed and attached to another atom in order to enable

e1 e2 d2

(1.2)

the force of attraction or bond strength K is proportional to the respective ionic charges e1 and e2 and inversely proportional to the square of the distance d2. In ideal ionic crystals, e.g., in halite NaCl, the cations (Na+) and anions (Cl–) form ideal spheres, whose electron shells do not overlap. This fact is evident from the distribution of the electron density (. Fig. 1.14a). From the distances between the central points of the respective anions and cations, their ionic radii can be calculated (see appendix, Fig. A.1). The ions tend to approach a closest packing of spheres where, depending on their respective sizes, different coordinations are achieved. For instance, in the halite structure, each Na+ cation is surrounded by 6 Cl− anions and vice versa, i.e., both ions are in [6]-coordination (see . Fig. 6.2).

14

Chapter 1 · Crystals

leading to the polar covalent bond type. In most ore minerals that are opaque even in thin section, a certain amount of metallic bonding is present as well.

1

z Metallic Bond

. Fig. 1.15  The four sp3-hybrid orbitals in the diamond structure display tetrahedral arrangement (after Borchardt-Ott 2008)

z Atomic Bond (Covalent Bond)

Crystals that consist of not more than one chemical component cannot crystallise in ionic structures. Instead, they are made up of electronically neutral building blocks, atoms. These atoms are bonded by electronic interaction between positively charged nuclei of atoms and negatively charged electron shells that is by sharing of electrons. The paired valence electrons rotate in a common electron orbit, where the orbital plane is perpendicular to the connecting line between the nuclei of the atoms. For instance, in the diamond structure (. Fig. 4.11), the outermost electron shell of the carbon atom is occupied by 2s22p2 electrons, whereas in the excited state each of the electrons is situated in the 2s and the 2px, 2py, 2pz orbitals. This results in four new sp3-hybrid orbitals that are directed to the corners of a tetrahedron (. Fig. 1.15). Consequently, each C atom can connect with up to four other C atoms leading to a structure with tetrahedral or [4]-coordination. Although the atomic bond approaches a noble gas configuration as well, the pairing of the valence electrons results in an overlap of the outer electron shells (. Fig. 1.16) as documented by the distribution of electron density, e.g., in the diamond structure (. Fig. 1.14b). Thus the concept of the atoms representing isolated ideal spheres in the crystal structure is not appropriate and the structure is better described by the space-filling calotte model. Atomic bonds are generally very strong as exemplified by diamond, which is by far the hardest of all minerals (. Table 1.2). Most minerals consist of more than one chemical element. Consequently, pure atomic bonds are very rare, and a variable proportion of ionic bonding is often present

. Fig. 1.16  Transition between ionic and atomic bonding (after Fajans from Kleber et al. 1998; courtesy of J. Bohm and H. Klimm, Berlin, Germany)

In contrast to the ionic and atomic bonds, the valence electrons in metallic structures are not localised or assigned to specific protons. They form—metaphorically speaking—a negatively charged electron cloud that moves, with a certain sojourn probability, between the positively charged atomic nuclei (these are not ions but a combination of protons and neutrons) and shields them from each other. In metallic structures, the atomic nuclei approach sphere packing, each with 12 nearest neighbours, i.e., they are [12]-coordinated. Judging from the stacking sequences of the atomic layers, two different types can be distinguished, the cubic close packing with the sequence ABCABC … and the hexagonal close packing with ABABAB … (. Fig. 4.1a, b). The physical properties of most metals, metal alloys, metal sulfides and metal oxides are governed by a dominant, or at least considerable extent of, metallic bonding, especially the opaque behaviour in transmitted and the strong reflectivity in reflected light (7 Sect. 1.5.3). z Van der Waals Bond

Compared to the principal bond types so far described, van der Waals bonds are relatively weak. They are based on residual valences that are caused by an uneven distribution of the positive and negative charges in groups of atoms or molecules, although, in principle, these are electrostatically neutral. As the centres of gravity of the positive and negative charges do not coincide, an electrostatic attraction between these dipoles is possible. However, the mutual potential energy decreases with the distance to the 6th power, i.e., the bond strength K is proportional to 1/d6. An important example of van der Waals bonding is the sheet structure of graphite (. Fig. 4.11c), in which each C is covalently linked, in [3]-coordination, to three neighbouring C atoms. A fourth, much weaker bond is formed in wide distance to another C atom in the overlying or underlying sheet, respectively. Therefore, only weak van der Waals forces exist between the individual sheets, explaining the perfect flaky cleavage and extremely low hardness of graphite. Van der Waals bonds play also an important role in most sheet silicates, especially in pyrophyllite and talc (7 Sect. 11.5.1, . Table 1.2).

15

1.3 · Crystal Chemistry

1.3.3  Some Important Terms of Crystal

Chemistry

z Isotypy

(Greek ίσος = equal, τύπος = nature, character) Crystals that crystallise in the same structure belong to a common structure type. They are called isotypous or isostructural. Generally, isostructural crystals belong to the same space group, are described by the same chemical formula and show a similar arrangement of the coordination polyhedra. The size of the chemical constituents and their bond types are less important. For instance, the structures of halite, NaCl, with pure ionic bonding (. Fig. 6.2) and of galena, PbS, (. Fig. 5.3) with dominant metallic bonding are mutually isotypous. z Solid Solutions

The relationship between isostructural crystals becomes even closer if their chemical components can mutually replace each other. By such a diadoch behaviour, solid solutions are formed. These play an important role in many mineral groups.

Na[AlSi3O8]. In contrast to the cations Na+ and Ca2+, which have about the same size of 1.24 and 1.20 Å, respectively, K+ has a distinctly larger ionic radius of 1.59 Å and, therefore, its coexistence with Na+ in the feldspar structure is tolerated only at high temperatures. z Polymorphism

(Greek πόλυ = much, μορϕή = form) Definition Polymorphism is a property of many chemical substances to crystallise in more than one crystal structure, depending on the respective thermodynamic conditions of state.

Examples of polymorphic minerals are orthorhombic and monoclinic sulfur, S, graphite and diamond, C, calcite and aragonite, CaCO3, α-quartz, β-quartz, tridymite, christobalite, coesite and stishovite, SiO2. Polymorphic transformations can take place in different ways: 5 At transformations in first coordination, the coordination number of the chemical constituents is changed. For instance, during the transformation calcite,

Definition Crystals, in which one or more point position(s) are statistically occupied by two or more chemical components are called solid solutions.

For instance, a complete solid-solution series is formed by the metals silver, Ag, and gold, Au. Such an (Ag,Au) alloy can be enforced by mechanically pressing together pure crystals of silver and gold at temperatures that are elevated but far below the respective melting points of the two metals. Thereby a process of diffusion is induced that leads—via various transitional states—to a final state with statistical Ag–Au distribution (. Fig. 1.17). Many rock-forming minerals form solid-solution series, e.g., olivine, (Mg,Fe)2SiO4, with the pure end members forsterite, Mg2SiO4, and fayalite, Fe2SiO4. Diadoch substitution can take place only if the atoms or ions involved are of similar size. Differences in chemical valence can be adjusted, however, by coupled substitution (coupled valence equalisation). The most prominent example is the plagioclase solid-solution series with the end members albite, Na[AlSi3O8], and anorthite, Ca[Al2Si2O8], in which the valence adjustment is achieved by the coupled substitution Na+Si4+ ⇌ Ca2+Al3+ (7 Sect. 11.6.2). The formation of solid solutions is facilitated by elevated temperatures. When the temperature decreases, exsolution takes place. For instance, the alkali feldspars form a complete solid solution series at temperatures above ca. 500 °C, but exsolve upon cooling below this temperature to form the virtually pure end members microcline, K[AlSi3O8], and albite,

. Fig. 1.17  a Structure of a silver and a gold crystal, projected onto (001). b By mechanical compressing diffusion of the Ag and Au atoms takes place and an (Ag,Au) solid solution is formed, in which these metal atoms are statistically distributed (after Borchardt-Ott 2008)

1

16

1

Chapter 1 · Crystals

Ca[6]CO3 ⇌ aragonite, Ca[9]CO3, the coordination number of Ca increases, which requires the bonds between Ca2+ and (CO3)2− to become broken. In the same way, the coordination number of Si increases from [4] to [6] during the transformation coesite ⇌ stishovite. As shown in the pressure-temperature diagrams for CaCO3 (. Fig. 8.8) and SiO2 (. Fig. 11.44), both systems conform to the general rule that the coordination number increases with rising pressure but decreases with rising temperature. 5 At transformations in second coordination, the coordination numbers of the chemical constituents do not change. Whereas the arrangement of the adjacent atoms or ions is maintained, the arrangement of the next atom or ion but one is changed. This is exemplified by the SiO2 polymorphs α-quartz, β-quartz and tridymite, whose transformations may be either displacive or reconstructive (. Fig. 1.18): a. During the displacive reversible transformation of the low-temperature polymorph α-quartz to the high-temperature polymorph β-quartz at a temperature of 573 °C (at a pressure of 1 bar) the SiO4 tetrahedra are merely rotated. b. In contrast, the reconstructive transformation from the relatively lower temperature polymorph β-quartz to the even higher temperature phase tridymite, involves the total breaking of the bonds between the SiO4 tetrahedra and a complete reconstruction to form a structure composed of six-membered rings (see also 7 Sect. 11.6.1). There are two more possibilities to achieve transformation in second coordination: c. Transformations by order-disorder processes play an important role in some mineral groups, especially the feldspars (see 7 Sect. 11.6.2). d. Transformation by alteration of the bond type, as performed by the graphite ⇌ diamond transformation, described above. Moreover, the sheet structure of graphite conforms to different stacking sequences in which the hexagonal graphite-2H displays a twolayer structure, the trigonal graphite-3R a three-layer structural unit. This special type of one-dimensional polymorphism that also plays an important role in the sheet-silicate structures (7 Sect. 11.5) is known as polytypism. 1.4  Physical Properties of Crystals 1.4.1  Hardness and Cohesion

Many minerals and synthetic crystals display a marked anisotropy of hardness. Moreover, due to the anisotropic distribution of cohesion properties, they are characterised by pronounced cleavages parallel to one or more crystal faces. For instance, halite, NaCl, shows a perfect cleavage parallel to the cube faces {100}, whereas the

. Fig. 1.18  Transformation in second coordination for Si[4]O2 structures, projected onto (0001): a tridymite (space group P63/mmc), b β-quartz (P6222), c α-quartz (P322); displacive transformation: α-quartz ↔ β-quartz, reconstructive transformation: β-quartz ↔ tridymite (see text). ● Si, 〇 oxygen (from Borchardt-Ott 2008)

perfect cleavage of fluorite, CaF2, follows the octahedral faces {111} (. Fig. 1.19a, b). Moreover, the hardness of fluorite is distinctly higher on {100} than on {111}, whereas the respective faces of halite do not show marked differences in hardness (. Fig. 1.19a, b). A remarkable anisotropy of hardness is recorded for kyanite Al2O[SiO4], leading to the old name disthene (Greek δις = twice, double, σθένος = strength). Based on its crystal structure, kyanite has a Mohs’ hardness of 4–4½ parallel to its c-axis, but 6–7 perpendicular to it (. Fig. 1.20). The term hardness describes the resistance of a solid against mechanical force, e.g., scratching or indentation. The scratch test to distinguish minerals by their hardness has been already known in ancient times, as mentioned by Theophrast (371–287 bc) and Pliny the Elder (ad 23/24–79). In 1812 Friedrich Mohs (1773–1839) proposed a hardness scale that is based on the ability of minerals and artificial materials to scratch each other. He selected 10 reference minerals, starting with talc as the softest (Mohs’ hardness 1) and ending up with diamond as the hardest (hardness 10). Materials with hardness 1 and 2 can be scratched by the fingernail, those up to 5 by a pocket-knife, up to 6 by a file or a steel

1

17

1.4 · Physical Properties of Crystals

an imprint of definite depth. As shown in . Table 1.2, the microhardness of the Mohs’ reference minerals increases in a non-linear way and in unequal steps. 1.4.2  Thermal Conductivity

. Fig. 1.19  Two minerals, crystallized in the form of the cube show different distribution of hardness and direction of cleavage. a Halite, NaCl, cleavage along the cube faces {100}. b Fluorite, CaF2 cleavage along the octahedral faces {111} (after Kleber et al. 2010, courtesy J. Bohm and H. Klimm, Berlin)

. Fig. 1.20  Strong anisotropy of hardness displayed by kyanite (disthene) Al2O[SiO4], depending on the crystal structure. Kyanite can be scratched by a steel needle (hardness 6) parallel to its longer axis c, but not in transverse direction (after Ramdohr and Strunz 1978)

needle; materials with a hardness of 7 or more are able to scratch glass. The Mohs’ hardness scale remains highly popular because of its simplicity but it is only a relative scale. Quantitative determination of the hardness is achieved by microhardness tests. The microhardness indenter is a small pyramid of diamond that is placed at the front lens of the objective of a microscope. The microhardness or indentation hardness, given in kilobars (kbar), is proportional to the pressure exerted onto a mineral surface to achieve

In most crystals, the thermal conductivity also behaves in an anisotropic way, as already demonstrated by the classical experiments of Hureau de Sénarmont (1808–1862). He covered a crystal face with a layer of wax and, after cooling, pressed a hot spike into it. It acted as a pointshaped heat source, from which the wax started to melt in outward direction. After removing the hot spike, a linear melt protrusion had formed that indicated the position of the melting isotherm at the moment when the heat supply was interrupted. On faces of anisotropic crystals, the melt protrusion has an elliptical shape. Taking gypsum, CaSO4·2H2O, as an example, the longest axis of this melt ellipse is inclined against the c-axis of the crystal by 16°; in this direction the thermal conductivity is higher by 20 % than perpendicular to it (. Fig. 1.21). The thermal conductivity λ is defined by the amount of heat Q that moves, per unit time t, through a rod of length l and with a cross section A, with a temperature difference ΔT between its upper and lower ends. According to the definition of Q in Eq. (1.3) it is proportional to A, ΔT and t, but inversely proportional to l whereby the material constant λ is a material-dependent proportionality factor:

Q = A ·

T ·t l

(1.3)

. Table 1.2  Microhardness of the standard minerals of Mohs’ scale of hardness (after Broz et al. 2006) Mohs’ hardness

Mineral

Microhardness (in kilobars)

1

Talc

1.4 ± 0.3

2

Gypsum

6.1 ± 1.5

3

Calcite

14.9 ± 1.1

4

Fluorite

20.0 ± 1.0

5

Apatite

54.3 ± 3.3

6

Orthoclase

68.7 ± 6.6

7

Quartz

122 ± 6

8

Topaz

176 ± 10

9

Corundum

196 ± 5

10

Diamond

1150

. Fig. 1.21  Crystal of gypsum with a linear melt protrusion of wax (blue) on the (010) crystal plane. The blue ellipse represents an isotherm and demonstrates the anisotropy of thermal conductivity (after Kleber et al. 2010; courtesy J. Bohm and H. Klimm, Berlin)

18

1

Chapter 1 · Crystals

Crystals with predominantly atomic or ionic bonding are generally marked by low thermal conductivity, whereas metallic bonds favour higher conductivity. For instance, at 0 °C, the thermal conductivity of silver is 419 W m−1 K−1, in the case of quartz only 7.25 perpendicular to its c-axis and 13.2 W m−1 K−1 parallel to it. 1.4.3  Electrical Properties z Electrical Conductivity

The electrical conductivity of crystals varies within an extremely wide range. For a good metallic conductor like silver, it amounts to 6 × 1017 Ω−1 m−1 whereas, for quartz, an insulator, it is only 3 × 10−5 Ω−1 m−1 (⊥c), a difference of as much as 22 orders of magnitude. As in any other electron conductor two kinds of electric conduction can be distinguished in crystals: 1. Ion conduction by diffusion of ions across the crystal structure plays a role in ionic crystals, especially at elevated temperatures. 2. Electron conduction occurs predominantly in crystals of metals and metal alloys with dominantly metallic bonding. The complex interactions between the atomic nuclei and the electron clouds, best described in terms of quantum mechanics, can be illustrated, in rough simplification, by the so-called band theory (e.g., Kittel 1996; Kleber et al. 1998, 2010). Electrons of single, isolated atoms occupy atomic orbitals, which form, at a temperature of absolute zero, a discrete set of energy levels. In crystals at geological temperatures, the different atomic orbitals are closely combined to form continuous energy bands of finite width (allowed bands), which are interrupted by band gaps or forbidden bands. The highest, fully occupied band is the valence band (VB), the next higher one the conductivity band (CB). The total chemical potential of all the electrons of a crystal is called the characteristic energy level or Fermi level EF that, in plots of the band structure, is set at zero. Due to thermal excitation, the Fermi level is somewhat blurred at higher temperatures but still distinct. The following cases can be distinguished (. Fig. 1.22): 5 In electronic conductors, i.e., in metallic or semi-metallic crystal structures, the Fermi level is situated inside the conductivity band which, consequently, is half filled, half empty. If an electric field is applied, the electrons in the conductivity band can gain energy; when accelerated they fill up the adjacent next higher energy level. The movement of the electrons in a given direction yields an electric current. 5 In insulators, i.e., in crystals with ionic or atomic bonds, the valence band is fully occupied, whereas the conductivity band is empty. Between both bands, a forbidden zone of several electron volts exists. As all reachable energy levels are occupied, the electrons cannot gain kinetic energy from an electric field applied, and

. Fig. 1.22  The band theory to explain a metallic conductors, b semiconductors, c insolators. E energy of the electronic state in electron volts (eV); VB, valence band; CB, conduction band; EF, Fermi energy (after Kleber et al. 1998; courtesy J. Bohm and H. Klimm, Berlin)

electronic conduction does not happen. However, minor ionic conduction is possible under the influence of lattice defects. At elevated temperatures, especially close to the melting temperature, the electrical conductivity of insulators increases. 5 In semiconductors, the forbidden zone between the valence band and the conductivity band is relatively narrow, varying between 0.1 and 3 eV. Different processes, e.g., thermal excitation, can lift electrons to fill the conductivity band and facilitate a certain amount of electric conduction, the n-conduction. Moreover, “holes” that are left in the valence band due to the loss of electrons may cause an additional effect, the p-conduction. The different influences that control the electronic properties of semiconductors lead to a wealth of technical applications, virtually in all fields of electronic hardware. In some cases, ultrapure semiconductor crystals are used whereas others are specifically doped with crystal impurities. For instance, the relatively low intrinsic conductivity of pure silicon or pure germanium can be considerably enhanced by doping the Si- or Ge-structure with traces of pentads like P5+ or As5+. Atomic centres formed by doping with crystal impurities are called donators if they emit electrons to the conductivity band, and acceptors if gaining electrons from the valence band, thereby generating movable holes in the crystal structure. z Piezoelectricity

(Greek πιέζειν = to press, to tense) Owing to the piezoelectric effect, an unequal distribution of electrical charge is generated when mechanical forces, like pressure or tension, are directly applied to the crystal. Hereby, microscopic dipoles are formed within the unit cell and an electric potential or tension voltage arises. This is a reversible process: By attaching an alternating electric field, mechanical vibrations arise due to intermittent compression and dilatation. Piezoelectric effects are possible only if pressure and tension are applied

1.4 · Physical Properties of Crystals

19

. Fig. 1.23 Piezoelectric effect on quartz. a Position of a quartz plate cut out of a quartz crystal; b quartz plate with the polar axes a1, a2 and a3; c piezoelectric effect generated by pressure along a polar axis, here a1 (after Borchardt-Ott 2008)

parallel to a polar axis, that is, if the structure at one and the other end of that axis differs. This condition is fulfilled by the twofold axes a1, a2 and a3 of quartz (crystal class 32, . Fig. 1.23). Therefore, pure, untwinned quartz crystals can be used, in a variety of ways, such as crystal oscillators for quartz clocks and watches (7 Sect. 11.6.1). The production of such devices is, however, dependent on large-scale crystallisation of synthetic “quartz”, because in the vast majority of natural quartz crystals the twofold axes have lost their polar character due to twinning. Piezoelectric effects are also shown by the minerals tourmaline (crystal class ¯ ) 3m, 7 Sect. 11.3) and sphalerite ZnS (crystal class 43m as well as the crystals of levo-(L-) and dextro-(D-)tartaric acid C4H6O8 (crystal class 2). z Pyroelectricity

(Greek πῦρ = fire) Similar to piezoelectricity, thermal treatment can also lead to electric charging, known as pyroelectricity, at the tip of a crystal if that tip represents the end of a polar axis. For instance, upon heating, tourmaline is charged positively at one end of the c-axis and negatively at the other; upon cooling the direction of charging is reversed. This effect results from the fact that tourmaline has a permanent electric dipole moment, the strength of which is temperature-dependent. As the polarity in the structure is a prerequisite of pyroelectricity, only certain crystal classes qualify for this specific property. These are crystals with no symmetry element (crystal class 1), with only one mirror plane (m), only one rotation axis (1, 2, 3, 4, 6) or one or two mirror planes parallel to a rotation axis (mm2, 3m, 4mm, 6mm). Crystals with higher symmetry, such as all cubic crystals or those with a centre of symmetry or of the type 222, 4¯ etc. cannot be pyroelectric. 1.4.4  Magnetic Properties

Each electron in an atom or an ion displays a magnetic moment that results from its spin and rotation. In the so-called transitional elements iron, Fe, and titanium, Ti, which are common constituents in many minerals, spin

plays the predominant role. The magnetic moments of two electrons with anti-parallel spin neutralise each other. Therefore, atoms or ions with paired electrons—summarised over all electrons—have no magnetic moment. These atoms as well as the crystals consisting of such atomic constituents are referred to as diamagnetic. In contrast, atoms and ions that contain one or more unpaired electrons, on average, display a magnetic moment; they are paramagnetic. The magnetic moment of an unpaired electron is a Bohr magneton μB = 0.9274  × 10−20 EMU (electromagnetic units). Most of the paramagnetic crystals possess only one unpaired electron and, consequently, have a magnetic moment of 1 μB. However, the transitional metals with the atomic numbers 21–30 show the tendency to fill the five places of the 3d level with a single electron and successively add further electrons until all these places are occupied by one electron (. Table 1.3). The occupation by only one of the five 3d-positions is known as high-spin condition, the most advanced double occupation as deep-spin condition. Thus Mn2+ and Fe3+ each display five unpaired 3d-electrons with parallel spin leading to a magnetic moment of 5 μB. In comparison, Fe0 and Fe2+ have six 3d-electrons, two of which are paired with anti-parallel spin and four of them unpaired with parallel spin, resulting in 4μB. In addition, Fe0 has two paired s electrons with anti-parallel spin. Ti3+ has one 3d-electron and, consequently 1 μB, whereas Ti4+ has no 3d-electrons and thus no magnetic moment (. Table 1.3). The distribution of paramagnetic atoms or ions in the crystal structure is critical for the magnetic behaviour of crystals: 5 In paramagnetic crystals, the paramagnetic atoms, their spins and their magnetic moments are statistically evenly distributed and thus neutralise each other. Therefore, the average magnetic moment μB is zero. 5 In ferromagnetic crystals, like α-iron, the exchange relationship between neighbouring Fe atoms leads to the magnetic moments in each of the crystal domains being parallel. This results in a high magnetic moment and, consequently, a high magnetic mass susceptibility. There is, however, also the possibility of two ferromagnetic substructures in which each of the unpaired electrons

1

20

1

Chapter 1 · Crystals

. Table 1.3  Electron configuration for iron and titanium Electron shell

K

L

M

N

Orbital

1s

2s

2p

3s

3p

3d

4s

Fe0

↑↓

↑↓

↑↓↑↓↑↓

↑↓

↑↓↑↓↑↓

↑↓↑↑↑↑

↑↓

Fe2+

↑↓

↑↓

↑↓↑↓↑↓

↑↓

↑↓↑↓↑↓

↑↓↑↑↑↑

Fe3+

↑↓

↑↓

↑↓↑↓↑↓

↑↓

↑↓↑↓↑↓

Ti0

↑↓

↑↓

↑↓↑↓↑↓

↑↓

↑↑↑↑↑

 ↑↓↑↓↑↓

Ti3+

↑↓

↑↓

↑↓↑↓↑↓

↑↓

↑↑

 ↑↓↑↓↑↓

Ti4+

↑↓

↑↓

↑↓↑↓↑↓

↑↓



 ↑↓↑↓↑↓

Magnetic moment

4μB 4μB 5μB

↑↓

2μB 1μB 0

↑↓ Paired electrons with anti-parallel spin; ↑ unpaired electrons

display anti-parallel spin. Two different possibilities can be distinguished: 5 In anti-ferromagnetic crystals, the magnetic moments exactly neutralise each other and, consequently, no magnetic moment results. 5 In ferrimagnetic crystals, like magnetite Fe2+Fe23+O4, the magnetic moments of the substructures do not exactly neutralise each other, because 5 the anti-parallel magnetic moments are slightly different, 5 the directions of spin are not exactly anti-parallel, or 5 there are structural defects or impurities in the structure. Therefore, if averaged over the whole structure, a magnetic moment results. Ferro- and ferrimagnetic crystals are characterised by a material-specific temperature at which they become paramagnetic. Above this so-called Curie point, the thermal vibrations in the crystal structure become so strong that the atomic magnets are no longer strictly parallel to each other. For magnetite, this occurs at a Curie temperature of 578 °C. Likewise, at the so-called Néel point, the anti-ferromagnetic crystals are transformed into paramagnetic crystal structures. For instance, ilmenite Fe2+Ti4+O3 is paramagnetic at room temperature but becomes anti-ferromagnetic at lower temperatures. For geomagnetic measurements and their interpretation, the Curie point is highly relevant (see 7 Sect. 7.2; Harrison and Feinberg 2009). 1.5  Optical Crystallography

The optical properties of crystals in polarised light are of paramount importance for the determination of rock-forming and economic minerals but also of crystalline substances in technical products. Therefore, a separate section of this book is devoted to this subject. Crystals that become transparent by grinding down to a thickness of about 25 µm (= 0.025 mm) are investigated in thin sections by using transmitted light, whereas opaque crystals, especially ore minerals, are studied in polished sections in reflected light. The light microscopic investigation of minerals, making use of

their optical properties, is the most common and important method of identifying rock types, establishing the crystallisation sequence as well as the mineral equilibrium assemblage and reaction textures, and thus provide most important clues for the reconstruction of rock-forming processes. In this textbook, only the basic concepts and methods of polarising microscopy will be described. For more details on optical mineralogy in transmitted light, the reader is referred to relevant textbooks, e.g., of Wahlstrom (1979) and Nesse (2004). For practical determination of rock-forming minerals, the tables of Tröger et al. (1979) are useful. Methods of ore microscopy have been described, e.g., by Craig and Vaughan (1981) while the classical text by one of the great ore petrologists, Paul Ramdohr (1976, 1980, 2013) remains indispensible for the practical determination of ore minerals. Moreover, the tables of Schouten (1962) provide a useful synoptical overview on the important ore minerals and their optical properties. 1.5.1  Basic Concepts

It is known that the nature of light and its interaction with matter can be described by two different theories, both of which complement each other: 5 The particle theory explains interaction of light with atoms, ions or molecules in terms of quantum physics. According to this theory light is composed of photons, that is, of small particles with a mass of close to zero, which move, like projectiles, from one point of matter, e.g., in a crystal structure, to another. 5 The wave theory considers light as energy of radiation that moves as electromagnetic waves from one point of matter to another. The optical properties that we observe under the microscope in transmitted or reflected light can be well described by the wave theory. Depending on the problem to be tackled, two alternative models may be applied: 5 The model of light as rays describes the propagation of light in space by geometric methods, their refraction and reflection as well as the path of rays in optical systems, especially under the microscope (geometric optics).

21

1.5 · Optical Crystallography

5 The wave model understands light in terms of a transverse wave that, by passing a crystal structure, is diffracted and polarised, leading to interference phenomena (physical optics). The visible light represents only a limited part of the continuous spectrum of electromagnetic radiation. It covers a range of wavelengths between 400 and 800 nm (1 nm = 10−7  cm), which includes the spectral lines, detected by Joseph von Fraunhofer (1787–1826) in the solar spectrum. In terms of the wave model, white light consists of a combination of an infinite number of light waves of different wavelength λ that oscillate with different amplitudes A. The intensity of light or brightness is proportional to A2. The relationship between the frequency of light f, i.e., the number of wave cycles per second (in Hz), the velocity of light ν and the wavelength λ is given by the simple equation

f=

ν 

(1.4)

According to Eq. (1.5) (see below), the refractive index of a medium is inversely proportional to the velocity of a propagating light wave whereas, with few exceptions, the frequency of a light wave remains constant irrespective of the medium it penetrates. As the velocity of light changes when entering a different medium of different density, the wavelength must change as well: in media with higher refractive indices, e.g., crystals or glass, ν and λ are smaller than in media with low refractive indices, e.g., in air (. Fig. 1.24). The following terms are useful for describing the transmission of light through different media (. Fig. 1.25a):

. Fig. 1.25  a Wave fronts are surfaces that connect equivalent points on adjacent waves; successive wave fronts are one wavelength apart. b In optically isotropic media, not only the wave normal but also the light ray is perpendicular to the wave front, c whereas in optically anisotropic media, this is no longer the case: wave normals and light rays are not parallel to each other (from Nesse 2004)

5 The wave front is a surface that connects equivalent points on adjacent waves. 5 The wave normal, a line perpendicular to the wave front, is the direction in which the light wave moves. 5 The light ray is the direction in which the light energy propagates. . Fig. 1.24  Passage of a light wave through two minerals, cassiterite SnO2 and rutile TiO2, with different indices of refraction of 2.0 and 2.9, respectively. On entering their crystal structures, light is slowed down to propagation velocities of 150,000 and 103,500 km−1, respectively. As the frequencies in air and the two minerals remain the same, the wavelengths in the crystals must be different and both shorter than in air (after Müller and Raith 1976)

In optically isotropic media like in glass, which is amorphous, or in cubic crystals, light moves with the same velocity in all directions. This implies that light ray and wave normal are parallel to each other (. Fig. 1.25b). In contrast, this is not the case in optically anisotropic media, i.e., in all

1

Chapter 1 · Crystals

22

1

non-cubic crystals, in which the velocity of light is different in different directions and, generally, light ray and wave normal are not parallel to each other (. Fig. 1.25c). 1.5.2  Basic Principles of Microscopy

in Transmitted Light

z Refraction and Birefringence

According to Snell’s Law (after Willebrord Snellius, 1580– 1626), a light ray entering from a medium with lower refractive index into a medium with higher refractive index is bent towards the axis of incidence and vice versa. This is due to the fact that wavelength and velocity of light changes when moving from one medium to another (. Fig. 1.24). The index of refraction is defined by the ratio of the light velocities in vacuum νV = 3.0  × 1010 cm s−1 = 300,00 0 km−1 and in a medium n with νn:

n=

νV νn

(1.5)

Consequently, a high refractive index always corresponds to a low velocity of light and vice versa. As νV is the highest possible light velocity, n must be always >1. Examples for refractive indices in cubic, i.e., optically isotropic, crystals are given in . Table 1.4. For diamond, the refractive index is strongly dependent on the wavelength of the light used: diamond has a high dispersion of refraction. In transmitted light, the relative differences in the refractive indices of two adjacent minerals, between a mineral and the mounting medium of a thin section, or between a mineral and an immersion oil in a grain mount is called the relief. It becomes visible, if the aperture of illumination is reduced by partly closing the iris diaphragm of the microscope. By this procedure, fissures or a delicate roughness, and thus positive or negative differences of relief, become visible. A useful aid is the Becke line method: By increasing the distance between the sample and the objective by lowering the mechanical stage of the microscope, a bright line moves into the medium with higher refractive index. Quantitative determinations are performed by the immersion method using grain mounts and immersion oils of different indices of refraction at . Table 1.4  Refractive indices of cubic, optically isotropic crystals Crystal

Formula

Refractive index n

Fluorite

CaF2

1.434

Halite

NaCl

1.544

Spinel

MgAl2O4

1.714

Almandine garnet

Fe2+3Al2[SiO4]3

1.830

Andradite garnet

Ca3Fe3+2[SiO4]3

1.887

Diamond for red light (λC = 656.3 nm)

C

2.410

Diamond for violet light (λF = 396.8 nm)

C

2.454

constant temperature; use of monochromatic light is advisable. More elaborate are the T, λ or λ-T variation methods, by which the temperature, the wavelength of light or both are gradually changed. In the past, one of these methods typically was employed to determine the chemical composition of solid solutions in rock-forming minerals, such as feldspars. These days, however, such elaborate and indirect optical methods are replaced by the more accurate, quantitative, direct in situ analysis with the aid of an electron microprobe.

Most of the rock-forming minerals are not cubic and are thus optically anisotropic. They, therefore, display different indices of refraction with a maximum value and a minimum value. The difference between these principal refractive indices is the maximum birefringence, The three-dimensional distribution of all possible refractive indices of a crystal is described by an index ellipsoid with the axes X, Y, Z, referred to as the optical indicatrix (in short: indicatrix). The shape of the indicatrix can take three principal forms: 1. For optically isotropic crystals, that are those with cubic symmetry, the indicatrix is a sphere because the refractive index is the same in all directions. 2. For optically uniaxial crystals, that are those with trigonal, tetragonal and hexagonal symmetry, the indicatrix is an ellipsoid with one circular cross section (ellipsoid of revolution). It can have either a prolate or an oblate shape, depending on whether the axis of rotation, i.e., the direction of the extraordinary wave nε is parallel to Z or X, respectively (. Fig. 1.26a, b). Perpendicular to the axis of rotation, there is a circular section, in which the refractive index is equal in all directions, representing the ordinary wave nω. Consequently, when looking down onto such a crystal in the direction of the axis of rotation, the crystal appears optically isotropic. Therefore, the axis of rotation—coinciding with the trigonal, tetragonal or hexagonal c-axis of the crystal— is referred to as optic axis. It should be noted that the optic axis of a crystal is not always parallel to its elongation. A crystal is described as optically positive if nε is greater than nω in which case the shape of the indicatrix is prolate. An example is quartz (. Fig. 1.26a). If nε is smaller than nω, the shape of the indicatrix is oblate and the crystal is optically negative, as is the case with the mineral calcite (. Fig. 1.26b). 3. In optically biaxial crystals with orthorhombic, monoclinic or triclinic symmetry, two optic axes exist and the indicatrix takes the form of a triaxial ellipsoid with the principal axes Z, Y, and X. The refractive index along Z (= nγ) is always the greatest, that along X (= nα) is the smallest, whereas nβ is perpendicular to the nγ-nα plane: nγ ≥ nβ ≥ nα (. Fig. 1.27a). In contrast to the uniaxial ellipsoid of revolution, triaxial ellipsoids display two circular sections that intersect in the Y-axis (= nβ). Two optic axes that are normal to these sections are located in the nγ-nα plane normal to Y (= nβ), the optic plane (. Fig. 1.27b). Looking along the direction of the optic axes, the crystal appears optically isotropic because in the circular section, all refractive indices

23

1.5 · Optical Crystallography

. Fig. 1.26  The indicatrix for optically uniaxial (trigonal, tetragonal and hexagonal) crystals has the shape of an ellipsoid of revolution; it is either a prolate in uniaxially positive crystals (optic axis Z||nγ = nε) or b oblate in uniaxial negative crystals (optic axis X||nα = nε). Looking in the direction of the optic axis, the crystal appears optically isotropic as, in this line of sight, all refractive indices are equal, depending on whether the radius of the circle is either nω = nα or nω= nγ (simplified after Nesse 2004)

are equal to nβ (. Fig. 1.27b). The angle enclosed by the two optic axes is referred to as the optical angle. In optically positive crystals, Z (= nγ) is the acute bisectrix of the optic angle, in optically negative crystals it is X (= nα) (. Fig. 1.27c, d). The difference between the principal refractive indices

�n = nγ − nα

(1.6a)

is the maximum birefringence, which can be observed only in sections that are cut parallel to the nγ − nα plane. In any other, arbitrarily cut sections, the indices of refraction are intermediate between these maximum and minimum values: nγ ≥ nγ′   ≥ nα′  ≥ nα and, consequently, the values of birefringence Δn = nγ′  − nα′ are correspondingly lower. Note that, by analogy, in arbitrarily cut uniaxial crystals, the birefringence observed corresponds to the absolute difference between nε and nω′ which is smaller than the maximum birefringence

�n = nε − nω .

(1.6b)

. Fig. 1.27  The indicatrix for optically biaxial (orthorhombic, monoclinic and triclinic) crystals has the shape of a triaxial ellipsoid. a The three principal indices of refraction nα||X, nβ||Y and nγ||Z; the optic axial plane is ⊥ to nβ. b In a triaxial ellipsoid there are two circular sections that intersect in the Y-axis (= nβ); the two optic axes are perpendicular to the circular sections. Consequently, along these axes, the crystal appears optically isotropic. c Optic axial plane for the biaxial positive indicatrix; the optical angle is 2Vγ, which means that Z (= nγ) forms the acute bisectrix. d Optic plane for the biaxial negative indicatrix; the optical angle is 2Vα, which means that X (= nα) forms the acute bisectrix (simplified after Nesse 2004)

Together with the main refractive indices n α, nβ, nγ (nε, nω in uniaxial crystals), the maximum birefringence Δn and the optic angle, 2Vγ or 2Vα, are diagnostic tools for the identification of minerals in thin section, providing clues on the chemical composition and/or the structural state of solid solutions, e.g., the feldspars (7 Sect. 11.6.2). In orthorhombic crystals, the main axes of the indicatrix Z (= nγ), Y (= nβ) and X (= nα) are parallel to the crystallgraphic axes c or b or a (. Fig. 1.28a). In monoclinic crystals, the indicatrix is inclined in one direction with respect to the crystal form whereas in triclinic crystals, it is inclined in two directions (. Fig. 1.28b, c). On passage through an optically anisotropic, doubly refracting crystal, light splits into two transverse waves that propagate at different velocities. This fact is convincingly demonstrated by using a clear crystal of calcite, a mineral with high double refraction (. Fig. 1.29). In former times, beautiful crystals of calcite were mined on Iceland and, therefore, were called Iceland crystal or Iceland spar. The faster wave conforms to the lower index

1

24

Chapter 1 · Crystals

1

. Fig. 1.28  Position of the indicatrix relative to the crystal form. The optic plane is shaded in grey. a Orthorhombic crystals: the indicatrix axes X, Y and Z coincide with the crystal axes a, b, and c; in the example given is a||Y, b||X and c||Z. As explained below, the extinction is parallel in sections bc and ac, whereas in sections ab, it is parallel with respect to {010} and symmetrical to {hk0}; b monoclinic crystals: the indicatrix is inclined with respect to one crystal form. Thus one of the indicatrix axes coincides with the crystal axis b, whereas the other indicatrix axes are not parallel the crystal axes a and c, except by chance. In the example given is b||Y (= nβ); thus the optic plane is || (010) and the extinction angles are c/Z = +15°, a/X = −5°; c triclinic crystals: the indicatrix is doubly inclined with respect to the crystal form; thus none of the indicatrix axes are parallel to the crystal axes, except by chance (after Nesse 2004)

of refraction nα′ , the slower wave to the higher refractive index nγ′ . In optically uniaxial crystals with trigonal, tetragonal or hexagonal symmetry, an ordinary wave nω and an extraordinary wave nε are distinguished. For nω Snell’s Law is valid whereas, for nε, this is not the case: 5 nε conforms to the optic axis; it is parallel to nγ in optically positive, parallel to nα in optically negative crystals 5 nω conforms to the radius of the circular section; it is parallel to nα in optically positive, parallel to nγ in optically negative crystals

It is proportional to the birefringence Δn and the thickness of the crystal d:

Ŵ = d · �n

(1.7)

If the thickness of the thin section is kept constant, i.e., close to 25 µm, one can calculate the birefringence of a mineral from the amount of retardation.

In optically biaxial crystals with orthorhombic, monoclinic and triclinic symmetry, none of the two waves conform to Snell’s Law. z Interference Colours

When a ray of light propagates through air, the wave-vectors of the vibration are statistically distributed around the axis of the propagation direction. As the ray of light enters a crystal, however, it becomes polarised, either linear or plane polarised. At the same time it will split into two rays of light, an ordinary and extraordinary ray, giving rise to the double refraction or birefringence described above for calcite. The ordinary ray appears stationary whereas the extraordinary ray deviates by a certain distance from the path of the ordinary one. This explains why one of the images (letters) beneath the calcite cleavage rhomb in . Fig. 1.29 would appear stationary and the other to move around when rotating the cleavage rhomb. The two waves move through the crystal at different velocities and vibrate in two planes that are perpendicular to each other. The delay of one of the rays relative to the other when exiting the crystal is referred to as the retardation Γ (given in nm).

. Fig. 1.29  Transparent cleaved fragment of calcite, rhombohedal ¯ The clearly visible duplication of the letters cleavage after {1010}. beneath the calcite (“optical crystallography is fun”) illustrates its very high birefringence (Δ = 0.1719)

1.5 · Optical Crystallography

Working with a petrographic microscope, polarised light is used at the outset, which is produced by a polarising filter, the lower polar, placed below the stage of the microscope. In the 19th and early 20th centuries, a Nicol prism (named after William Nicol, 1768–1851) was used to generate polarised light. For this purpose, an inclusion-free and transparent cleavage rhomb of calcite (. Fig. 1.29) was prepared, which lets through only the plane polarised ordinary rays of light, whereas the extraordinary rays are deflected and absorbed. While the lower polar is generally a fixed installation, a second polarising filter, the upper polar or analyser, the vibration direction of which is perpendicular to that of the lower polar, may be placed for certain purposes at the upper end of the optical path between objective and eye lens. If the analyser is put in, one speaks of crossed polars or, in remembrance of older designs, of crossed Nicols (in short +Nic). Entering a crystal, plane polarised light is split into two plane polarised waves of different velocity that vibrate perpendicular to each other and interfere on entering the upper polar (. Fig. 1.30). The resulting interference phenomena, when interpreted correctly, can provide great help in the optical mineral identification. Let’s first consider the situation in monochromatic light, i.e., with light of a specific wavelength λ. In case, the retardation between the two waves is Γ = iλ, where i is an integer, both waves vibrate in opposite directions, i.e., they are out of phase. When entering the upper polar, addition of vectors leads to a resulting wave S that vibrates perpendicular to the vibration direction of the upper polar and, as a consequence, is obliterated (. Fig. 1.30a). If, however, the retardation is Γ = (i + ½λ) both waves vibrate in the same direction, i.e., they are in phase. The resulting wave S vibrates parallel to the vibration direction of the upper polar and thus will be amplified (. Fig. 1.30b). In . Fig. 1.31, the interference pattern of a wedge-shaped piece of quartz is shown. As the birefringence of this crystal Δn is constant, the retardation Γ merely depends on the respective thickness d of the quartz wedge. If the retardation is an integer multiple of the wavelength Γ = iλ, light transmission attains a minimum leading to total extinction, whereas retardation of Γ = (i + ½λ) leads to maximum transmission of light. In contrast to monochromatic light, white light is composed of an infinite number of waves with different wavelengths. Upon entering a crystal, the rays of light split into two, which propagate at different velocity and vibrate perpendicular to each other. This implies interference. For a certain thickness of the crystal, the birefringence and, therefore, the retardation will be about the same for all wavelengths. However, as wavelengths are different in white light, some will reach the upper polar in phase, some out of phase and, therefore, will be either obliterated or transmitted. The combination of the different wavelengths that penetrate the upper polar results in different interference colours which, in a thin section of constant thickness d, reflect the birefringence Δn. With increasing

25

. Fig. 1.30  Interference phenomena in double-refracting crystal plate. a The retardation is one wavelength Γ = λ. b The retardation is one half wavelength Γ = ½λ. For explanation see text (after Nesse 2004)

. Fig. 1.31  Interference pattern formed by the propagation of monochromatic light through a quartz wedge. a If the retardation is an integer multiple of the wavelength, Γ = iλ, the fast and the slow wave destructively interfere at the upper polar and thus are obliterated, producing a dark band. If the retardation is Γ = (i + ½λ), both waves interfere constructively at the upper polar, and light passes with maximum intensity. b Transmission through the upper polar (in %) depending on the retardation and assuming ideal optical conditions (after Nesse 2004)

1

26

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Chapter 1 · Crystals

retardation, the interference colours change from dark grey via light grey, white, yellow, to red of the first order. With increasing orders of retardation, the colour sequence blue → green → yellow → red repeats itself several times until the colours appear progressively more washed-out. Various interference colours are beautifully documented in the microphotograph of a crystal of zircon, Zr[SiO4], with impressive growth zonation (. Fig. 11.6). At a thickness d of 25 µm, quartz with Δnmax = 0.009 shows a grey of first order, whereas the olivine end-member forsterite, Mg2[SiO4], with Δnmax = 0.033 displays a green of second order. Of course, optically isotropic crystals appear black under +Nic, because Δn and, consequently, Γ are zero. Anomalous interference colours are shown by crystals that display strong dispersion of birefringence, i.e., when birefringence and retardation are markedly different for different wavelengths of light. In minerals with overall low birefringence, e.g., in many chlorites (7 Sect. 11.5.5), subnormal interference colours are observed, whereas in minerals with overall high birefringence and strong dispersion, e.g., in epidote (7 Sect. 11.2), the interference colours are abnormally high. z Extinction Angle

In addition to the retardation Γ and the wavelength λ, the amount of light T that leaves the upper polar depends on the angle τ between the indicatrix and the vibration direction of the lower polar:   180◦ Ŵ T = − sin2 · · sin 2τ · sin 2(τ −90◦ ) · 100 

(1.8)

It is easily recognised that maximum light transfer is achieved if τ = 45°, 135°, 225°, 315°, whereas the amount of light passing is T = 0 whenever τ = 90°, 180°, 270°, or 360°. Consequently, turning the crystal through 360°, one observes four times complete extinction and—in diagonal position—four times maximum transmission (brightness). In monoclinic and triclinic crystals, the indicatrix is tilted to the crystal form in one or two directions, respectively (. Fig. 1.28). Accordingly, extinction takes place if the crystal is tilted at a certain angle against the vibration direction of the lower polar. This angle is referred to as the extinction angle. For instance, in the (010) section of the monoclinic crystal shown in . Fig. 1.28b, extinction angles c/Z = +15° and a/X = –5° are recorded. On the other hand, one observes straight extinction in the (100) section of the same crystal and symmetrical extinction, relative to the {110} faces as the XZ plane bisects the angle between the faces (110) and (11¯0). Prominent examples are the {110} cleavage planes of pyroxenes and amphiboles (. Figs. 11.27, 11.28). Straight extinction is also seen in relevant sections of orthorhombic (. Fig. 1.28a) and optically uniaxial crystals. Measuring the angle of extinction is only possible if the morphological directions are clearly indicated by crystal faces, cleavage planes or twin planes.

z Accessory Plates

An accessory plate is used to determine the position of ′ ′ the respective higher or lower refractive index, nγ or nα. This plate is typically made of gypsum of equal thickness (hence “gypsum plate”). Its retardation is exactly Γ = 551 nm, corresponding to first-order red and is placed, in diagonal (or 45°) position between the objective and the ′ upper polar. If nγ of the gypsum plate coincides with nγ ′ or nα of the mineral, both retardations are either added or subtracted, respectively (. Fig. 1.32). In the case of addition, the first-order grey shown by quartz (Δnmax = 0.009), for example, changes to second-order blue. In the case of subtraction, that is the wavelength of the gypsum plate minus that of first-order grey, results in first-order yellow. The gypsum plate can also be used to determine whether nγ or nα are oriented parallel to the long axis of a crystal (. Fig. 1.32), which is then described as either positive or negative elongation, respectively. The character of elongation is, however, not to be confused with the optical character, which does not have to be necessarily the same. z Conoscopic Interference Figures

All of the above observations refer to the behaviour of light when passing through birefringent crystals under orthoscopic illumination (Greek όρθὀς = straight, σκοπέω = examine, inspect), that is when the rays of light enter the thin section at right angles. Further characteristics of a crystal, such as its optical character, can be determined under conoscopic illumination (Greek κώνος = cone), that is, when only the central light ray enters the thin section under a right angle but all other rays enter at an oblique angle that deviates progressively from 90° with increasing distance from the central path. To this effect a high-power objective (with 45× or 50× magnification) and a retractable auxiliary condenser lens, both with high numerical aperture, are used. Thus a cone of strongly converging light is produced. When inserting a small lens, named after Emile Bertrand (1844– 1909), into the optical path just below the ocular, a range of different interference figures can be observed, which can be used to deduce the optical character as well as the optic axes angle 2V in the case of biaxial minerals. For further details

. Fig. 1.32  Determination of the elongation of a crystal by means of a gypsum plate (red of first order). a Subtraction, i.e. nα is || to the longest axis of the crystal: negative elongation. b Addition, i.e. nγ is || to the longest axis of the crystal: positive elongation (after Müller and Raith 1976)

1.5 · Optical Crystallography

on conoscopic methods and their theoretical background, the reader is referred to special textbooks on optical mineralogy (e.g., Nesse 2004). For the conoscopic investigation of an optically uniaxial mineral, a grain is selected that, on orthoscopic illumination, remains virtually dark under +Nic, indicating that this grain is cut nearly perpendicular to its optic axis. Under conoscopic illumination, the typical uniaxial interference figure is a dark Maltese cross, the centre of which, the melatope (Greek µ´εoς = black, τ´oπoς = place), corresponds to the exit point of the optic axis. The four beams of the cross, called isogyres (Greek ίσος = equal, γυρός = rounded, curved, crooked), correspond to the four directions of extinction in normal position (. Fig. 1.33a). In the four sectors between the isogyres, i.e., in diagonal position, one observes coloured segments of a circle, the isochromes, with interference colours whose order increases outwards. According to Eq. (1.7), the number of isochromes depends on the retardation Γ, i.e., on the thickness d of the crystal and its birefringence Δnmax. In normal thin sections with d = 25 µm, quartz (Δnmax = 0.009) displays only a few, but calcite (Δnmax = 0.172) a great number of isochromes. The optical character of a uniaxial mineral can be determined by inserting a gypsum plate, as this will invariably lead to addition and subtraction of interference colours in adjacent

. Fig. 1.33 Conoscopic interference figures a for optically uniaxial, b for optically biaxial crystals (left: in diagonal position, right: in normal position); c–e determination of the optical character using a gypsum or quartz red I plate (left: optically positive, right: optically negative crystals); c optically uniaxial crystal, cut ⊥ to the optic axis; d optically biaxial crystal cut ⊥ to the acute bisectrix; e optically biaxial crystal cut ⊥ to one of the optic axes (simplified after Müller and Raith 1976)

27

quadrants. In optically positive crystals, addition is noted in the upper right and bottom left quadrants and subtraction in the other two quadrants, whereas optically negative crystals show the exact opposite effect, that is subtraction in the upper right and bottom left quadrants (. Fig. 1.33c). In sections that are cut oblique to the optic axis, the Maltese cross does not show up in the centre of the field of view but somewhere on the side. On rotation of the microscope stage it moves on a circle, the diameter of which increases with the inclination of the optic axis. If the section is cut at a high angle to the optic axis, the Maltese cross will not be visible any more and instead the horizontal and vertical beams of the cross will be seen to move through the field of view parallel to the vibration directions of the lower and upper polar. For conoscopic investigation of an optically biaxial mineral, a section is selected that is cut as close as possible to the acute bisectrix of the optical angle and, therefore, shows relatively low interference colours. In diagonal position, the isogyres that conform to the positions of extinction display the shape of hyperbolas; their vertexes correspond to the position of the two optic axes (. Fig. 1.33b, left). On rotation of the microscope stage, the hyperbolas change in shape until, in normal position, a dark cross appears that is, however, different from the Maltese cross described above for

1

28

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Chapter 1 · Crystals

optically uniaxial crystals (. Fig. 1.33b, right). The shape of the isochromes in diagonal and in normal position is shown in . Fig. 1.33b (left and right, respectively). In sections cut perpendicular to one of the two optic axes—as indicated, in orthoscopic illumination, by complete extinction under +Nic—the vertex of one hyperbola is situated exactly in the centre of the field of view, whereas the second hyperbola is not visible (. Fig. 1.33e). Determination of the optical character of biaxial crystals by using a gypsum plate, is demonstrated in . Fig. 1.33d, e. z Natural Colour

In general, the amplitude of a light wave, passing through a crystal, is reduced due to absorption. The extent of absorption varies for different wavelengths. Thus some wavelengths are totally absorbed, others partly and others not, resulting in natural proper colours (not to be confused with interference colours!), which can be an important feature for mineral identification in a thin section. The intensity of absorption depends on the thickness of the crystal and its chemical composition, with atoms of the transitional metals, especially Fe, but also Mn, Ti, Cr, V, Co, Ni, Cu, playing a particularly important role. Minerals devoid of Fe, especially feldspars and quartz, are colourless in thin section, whereas those richer in Fe, like amphiboles (7 Sect. 11.4.3), biotite (7 Sect. 11.5.2), chlorite (7 Sect. 11.5.5), tourmaline (7 Sect. 11.3) and epidote (7 Sect. 11.2), show colours of variable intensity. In optically isotropic crystals the colour is the same in any section. In anisotropic crystals, however, absorption of light of different wavelengths is anisotropic as well and, therefore, the colour changes depending on orientation. The strongest differences in colour and/or in colour intensity are recorded along the indicatrix axes X, Y and Z. Thus optically uniaxial crystals are often dichroitic, biaxial ones pleochroitic (Greek πλέον = more, χρώμα = colour). In the following, possible absorption schemes are given for a uniaxial and a biaxial mineral example, respectively: 5 tourmaline (variety schorl): X (bluish grey) ≪ Z (olive brown) 5 hornblende X (light yellowish green) 2.5 Ga: Archaic cratons; 2.5–1.6 Ga: Lower Proterozoic cratons; 1.6–0.8 Ga: Upper Proterozoic crystalline complexes; 200 µm, placed in a very fine-grained matrix. Framesites contain typical minerals of the Earth’s mantle such as rose-coloured, Cr-rich pyrope or almandine garnet (7 Sect. 11.1), emerald-green, Cr-rich clinopyroxene (7 Sect. 11.4.1) and chromite, FeCr2O4, but no olivine, besides Cr-poor, Ca-rich garnet, typically present in eclogites (Sects. 26.3.1, 28.3.9), but no omphacite (7 Sect. 11.4.1). Together with diamond, framesite is extracted from kimberlite in the Premier and Venetia mines, South Africa, the Orapa and Jwaneng mines, Botswana, and the Mir Mine, Siberia. Most researchers agree that framesite was formed by rapid crystallisation of carbon in limited parts of the Earth’s mantle. Carbonatite magmas, deeply subducted carbonate rocks or contentrations of organic carbon may have served as starting materials (Heaney et al. 2005). Diamond synthesis  Almost synchronously, in 1955, the Swedish ASEA company and General Electric in USA succeeded in the artificial transformation of graphite to diamond at pressures of 50–60 kbar (5–6 GPa) and temperatures of 1400 °C. By this process, molten metals, generally Ni and/or Fe, also Co, were used as catalysts. Although diamond synthesis is extremely energy-consuming, a high amount of industrial diamonds is being synthesised, as the natural resources are insufficient to meet the demand. In 2015, China was the leading producer of synthetic diamonds (>4 × 109 carats). Outside its stability field, diamond is precipitated from hot gases. By this Chemical Vapour Deposit (CVD) method, pure single crystals of diamond were grown, reaching a thickness of up to 4.5 mm (Hemley et al. 2005). This technique is used, e.g., for coating technical materials, production of compound powders, of diamond windows for space shuttles and diamond-anvils for high-pressure– high-temperature experiments (7 Sect. 29.3.4).

z Stability Fields of Diamond and Graphite

From the much higher density and the by far closer packing of C atoms in the diamond structure, it is clear that

4

83

4.3 · Non-metals

diamond is the high-pressure modification of carbon, a fact confirmed by ultra-high pressure experiments (Bundy et al. 1961; Berman 1962). These show that, at room temperature, graphite is the stable C modification up to pressures of about 20 kbar (=2 GPa), whereas diamond is metastable. The metastable coexistence of the two C modifications at low atmospheric pressure and room temperature is due to the fact that the structural transformation diamond → graphite is extremely slow and requires a high activation energy. As shown in the stability diagram of carbon, the stability curve of the reaction

diamond ⇋ graphite

[4.1]

has a positive slope, which implies that with increasing temperature, the transformation pressure of this reaction increases as well (. Fig. 4.15). Following the continental geothermal gradient, i.e., the increase of temperature with depth below Earth’s surface, estimated by geophysical and petrological methods, the equilibrium curve of Reaction [4.1] will be reached at a pressure of about 40 kbar, corresponding to a depth of about 140 km (point E in . Fig. 4.15). In oceanic areas, the temperature increase with depth is even higher. Therefore, following the oceanic geothermal gradient, the equilibrium curve of Reaction [4.1] will be reached at even greater depth (. Fig. 29.15).

. Fig. 4.15  Pressure-temperature diagram showing the stability curve graphite/diamond, as determined by high- and ultra-high pressure experiments. Also shown is the sub-continental geothermal gradient, i.e. the increase of temperature with depth below continental areas of the Earth, the course of which in the Earth’s crust and upper mantle has been estimated by geophysical and petrological methods. Following this gradient, the graphite ↔ diamond transition is reached at point E, at a pressure of about 40 kbar, at a depth of about 140 km (after Bundy et al. 1961 and Berman 1962, from Ernst 1976)

84

Chapter 4 · Elements

4 . Fig. 4.16  Crystal forms of α-sulfur. a–c Crystals with rhombic-dipyramidal habit; d distorted crystal approaching disphenoidal habit

{111} (. Fig. 4.16d). Much more frequently, native sulfur occurs in fine-grained masses or coatings. Monoclinic native sulfur, β-S, is formed, below its melting point of 119 °C, as fine-grained crystal coatings on volcanic craters. Upon cooling below 95.6 °C, it readily transforms into the rhombic α -sulfur and thus is rare in nature. Physical properties Cleavage

barely visible

Fracture

conchoidal, brittle

Hardness

1½–2

Density

2.0–2.1

Colour

sulfur-yellow or orange, due to small amounts of selenium, Se; in places, brown bitumen coatings (. Fig. 4.17)

Lustre

adamantine on crystal faces, greasy to resinous on fractures, translucent in thin splinters

Streak

white

Chemical composition  S may be substituted by minor Se.

. Fig. 4.17  Crystals of natural sulfur, partly with brown coatings of bitumen, that grew onto calcite; Agrigento, Sicily, Italy; Width of view is c. 9 cm

z Sulfur α-S Crystal form, habit and stability relations  Beautiful crystals

of orthorhombic native sulfur, α-S, crystal class 2/m2/m2/m, are common. In many cases they show two rhombic dipyramids, of which the steeper one, {111}, predominates, combined with a prism, e.g., {110}, and/or the basal pinacoid, {001} (. Figs. 4.16a–c, 4.17). In some cases of lower symmetry, the dipyramids are replaced by the rhombic disphenoid,

Crystal structure  The unit cell of rhombic sulfur contains 16 circular S8 molecules thus amounting to as much as 128 S atoms per formula unit (. Fig. 4.18b). Within these S8 rings, covalent bonds predominate with a considerable overlap of the electron shells (. Fig. 4.18a), whereas the rings are attracted to each other by weak van der Waals bonds only. These structural properties are responsible for the poor electric and thermal conductivity, the low melting and sublimation temperatures as well as the low hardness and density of native sulfur. Occurrence  Native sulfur is deposited from volcanic

exhalations and hot springs. Volcanic sublimates of sulfur are occasionally mined. Up to the end of the 20th century, however, sedimentary sulfur, formed by reduction of anhydrite or gypsum aided by sulfur bacteria was economically much more relevant. Historically important deposits were mined on the island of Sicily, southern Italy. Nowadays elementary sulfur is mainly produced during purification of

85 References and Suggestions for Further Reading

. Fig. 4.18  a The S8 rings of orthorhombic sulfur. b Elementary cell of orthorhombic sulfur, consisting of 16 ring-shaped S8 molecules that ¯ are mutually arranged, in a rouleau-like fashion, in the [110] and [111] directions of the structure. The dominant zones in the crystal forms of sulfur mirror these structural directions (from Klein and Hurlbut 1993)

Ernst WG (1976) Petrologic Phase Equilibria. Freeman, San Francisco Foote AE (1891) A new locality for meteoritic iron with a preliminary notice on the discovery of diamonds in the iron. Am J Sci 42:413–417 Geim AK, Kim P (2008) Wunderstoff aus dem Bleistift. Spektrum der Wissenschaft, August 2008, pp 86–93 Hammond CR (2004) The Elements. Handbook of Chemistry and Physics, 81st edn. Boca Raton Kroto HW, Heath JR, O’Brian SC et al (1985) C60: Buckminsterfullerene. Nature 318:162–163 Medenbach O, Wilk H (1977) Zauberwelt der Mineralien. Sigloch edn. Künzelsau Thalwil Salzburg Ramdohr P, Strunz H (1978) Klockmanns Lehrbuch der Mineralogie, 16th edn. Enke Verlag, Stuttgart Schmitt RT, Lapke C, Lingemann CM et al (2005) Distribution and origin of impact diamonds in the Ries Crater, Germany. In: Kenkmann T, Hörz F, Deutsch H (eds) Large Meteorite Impacts III. Geol Soc America Spec Paper 384:1–16 Shcheka GG, Lehmann B, Gierth E et al (2004) Macrocrystals of Pt-Fe alloy from the Kondyor PGE placer deposit, Khabarovskiy Kray, Russia: trace element content, mineral inclusions and reaction assemblages. Can Mineral 42:601–617 Shumilova TG, Isaenko SI, Divaev FK, Akai J (2012) Natural carbon nanofibers in graphite. Mineral Petrol 104:155–162

Further Reading

natural gas and petroleum by removing sulfur-containing contaminants. At present, the labour-intensive and highly dangerous mining of sulfur sublimates takes place at the active volcanic vent of Ijen Volcano, eastern Java, Indonesia. On the island of Volcano, southern Italy, recovery of volcanic sulfur dates back to antiquity and reached a maximum of 240 t per year in 1873–1876 when 450 workers, mostly convicts, worked on the active Fossa Volcano. After the last phase of volcanic activity on the island had ceded in 1888–1890, mining on Volcano virtually ceased.

Economic significance  Native sulfur is used for production of sulfuric acid, for volcanising natural rubber, in pulp industries, for production of pesticides and herbicides, and in pyrotechnics.

References and Suggestions for Further Reading Berman R (1962) Graphite-diamond equilibrium boundary. 1st Internat Congr Diamonds in Industry. Ditchling Press, Sussex, pp 291–295 Bundy FP, Bovenkerk HP, Strong HM, Wentorf RH Jr (1961) ­Diamond-graphite equilibrium line from growth and graphitization of diamond. J Chem Phys 35:383–391 El Goresy A, Gillet P, Chen M et al (2001) In situ discovery of shock-induced graphite-diamond phase transition in gneisses ­ from the Ries crater, Germany. Am Mineral 86:611–621 El Goresy A, Dubrovinsky LS, Gillet P et al (2003a) A novel cubic, transparent and superhard polymorph of carbon from the Ries and Popigai craters: implications for understanding dynamic-induced natural high-pressure phase transitions in the carbon system. Lunar Planet Sci 34 (CD-ROM) El Goresy A, Dubrovinsky LS, Gillet P et al (2003b) A new natural, superhard and transparent polymorph of carbon from the Popigai impact crater, Russia. CR Geoscience 335:889–898

Anthony JW, Bideaux RA, Bladh KW, Nichols MC (2000) Handbook of Mineralogy. Diamond. Mineralogical Society of America Anthony JW, Bideaux RA, et al (1990) Handbook of Mineralogy, vol I: elements, sulfides, sulfosalts. Mineral Data Publ., Tucson Cabri LJ, Harris DC, Weiser TW (1996) The mineralogy and distribution of platinum group minerals (PGM) in placer deposits of the world. Explor Mining Geol 5:73–167 Gurney JJ, Helmstaedt HH, le Roux AP et al (2005) Diamonds: crustal distribution and formation processes in time and space and an integrated deposit model. In: Hedenquist JW, Thompson JFH, Goldfarb RJ, Richards JP (eds) Economic geology. One Hundreth Anniversary Volume, pp 143–177 Harlow GE (ed) (1998) The Nature of Diamonds. Cambridge University Press, Cambridge Harlow GE, Davies RM (2005) Diamonds. Elements 1:67–69 Heaney PJ, Vicenzi EP, De S (2005) Strange diamonds: the mysterious origins of carbonado and framesite. Elements 1:85–89 Hemley RJ, Chen Y-C, Yan C-S (2005) Growing diamonds by chemical vapor deposition. Elements 1:105–108 Hough RM, Butt CRM, Fischer-Bühner J (2009) The crystallography, metallography and composition of gold. Elements 5:297–302 Huss GR (2005) Meteoritic nanodiamonds: messengers from the stars. Elements 1:97–100 Klein C, Dutrow B (2007) Manual of Mineral Science, 23rd ed., Wiley & Sons, Hoboken, New Jersey Klein C, Hurlbut CS Jr (1985) Manual of Mineralogy (after James D. Dana), 20th edn. Wiley, New York Klein C, Hurlbut CS Jr (1993) Manual of Mineralogy (after James D. Dana), 21st ed. Wiley, New York Levinson AA, Gurney JJ, Kirkley MB (1992) Diamond sources and production: past, present, and future. Gems Gemol 28:234–254 Luque FJ, Huizenga J-M et al (2014) Vein graphite deposits: geological settings, origin, and economic significance. Miner Deposita 49:261–277 Mungall JE, Meurer WP (eds) (2004) Platinum group elements: petrology, geochemistry, mineralogy. Can Mineral 42:241–694 Ogasawara Y (2005) Microdiamonds in ultrahigh-pressure metamorphic rocks. Elements 1:91–96 Rumble D (2014) Hydrothermal graphitic carbon. Elements 10:427–433

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Chapter 4 · Elements

Stachel T, Brey GP, Harris JW (2005) Inclusions in sublithospheric diamonds: glimpses of deep Earth. Elements 1:73–78 U.S. Geological Survey (2020) Mineral commodity summaries 2020: U.S. Geological Survey, 200 p. 7 https://doi.org/10.3133/mcs2020

4

Walshe JL, Cleverley JS (2009) Gold deposits: Where, when and why. Elements 5:288–295 Yang J-S, Robinson PT, Dilek Y (2014) Diamonds in ophiolites. Elements 10:127–130

87

Sulfides, Arsenides and Complex Sulfides (Sulfosalts) 5.1 Metal Sulfides with M:S>1:1 (Generally 2:1) – 88 5.2 Metal Sulfides and Arsenides with M:S = 1:1 (. Table 5.2) – 90 5.3 Metal Sulfides, Sulfarsenides and Arsenides with M:S = 1:2 – 96 5.4 Arsenic Sulfides – 102 5.5 Complex Metal Sulfides (Sulfosalts) – 102 References – 104

© Springer-Verlag GmbH Germany, part of Springer Nature 2020 M. Okrusch, H. E. Frimmel, Mineralogy, Springer Textbooks in Earth Sciences, Geography and Environment, https://doi.org/10.1007/978-3-662-57316-7_5

5

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Chapter 5 · Sulfides, Arsenides and Complex Sulfides (Sulfosalts)

Introduction

5

This mineral class comprises most of the ore minerals. Many of them are opaque, even in thin sections that are not more than 20–30 µm in thickness, and show metallic lustre with different shades of colour. Non-opaque sulfidic ore minerals are translucent in thin sections or along thin edges and display a very high refractivity of light, some of them showing adamantine lustre. Sulfide minerals tend to have a distinct streak colour that can be helpful in mineral identification. In their crystal structures, most sulfide minerals exhibit mixed bonds containing metallic, atomic and ionic components. These may be accompanied by van der Waals bonds, especially in layer structures. The sulfide minerals are subdivided into groups with different metal-sulfur (M:S) ratios (e.g., Strunz and Nickel 2001): 1. Metal sulfides with M:S > 1:1 (generally 2:1) 2. Metal sulfides and arsenites with M:S = 1:1 3. Metal sulfides, sulfarsenides and arsenites with M:S = 1:2 4. Arsenic sulfides 5. Complex metal sulfides (sulfosalts)

transforms to trigonal low-temperature digenite, ~Cu1.8S, crystal class 3¯m. In a temperature range of 435–93 °C hexagonal high-temperature chalcocite, Cu2S, crystal class 6/m2/m2/m, is stable. Tetragonal low-temperature chalcocite, Cu2S, crystal class 422, crystallises at temperatures below 103 °C. In general, members of this group form massive ore, with crystal faces being only rarely developed (. Table 5.1). Djurleite, Cu1.96S, and anilite, Cu1.75S, are stable at temperatures of  1:1 Mineral/Polymorph

Formula

Metal content

Temperature (°C)

Crystal class

High-T

Cu9S5–Cu2S

Variable

>75

¯ 4/m32/m

Low-T

Cu1.8S

Cu ~72%

103

6/m2/m2/m

Low-T

Cu2S

Cu 79.8%

265

¯ 4/m32/m

Low-T

Cu5FeS4

Cu 63.3%

173

¯ 4/m32/m

Akanthite

Ag2S

Ag 87.1%

 1:1 (Generally 2:1)

Stability relations, crystal form and habit  Cubic argentite,

z Bornite Cu9FeS4 Crystal forms  High-temperature bornite is cubic with crystal class 4/m3¯2/m, low-temperature bornite orthorhombic with 2/m2/m2/m. Euhedral crystals are rare, in places aggregates of distorted cubes can be observed. In most cases, however, bornite occurs as massive ore. Physical properties Cleavage

rarely distinct

Fracture

conchoidal

Hardness

3

Density

5.06–5.08

Optical properties

opaque; reddish bronze-coloured on fresh fracture surfaces; tarnish multi-coloured, red or blue, finally black; metallic lustre

Streak

greyish-black

Crystal structure and stability relations  The crystal struc-

ture of cubic high-temperature bornite is similar to that of sphalerite. Upon cooling below about 265 °C, high-temperature bornite transforms into a metastable, cubic modification known as intermediate bornite which, at about 200 °C, converts to the stable, orthorhombic lowtemperature modification. Low-temperature bornite displays complex superstructures with many structural defects that result in a range of Cu:Fe:S ratios. Chemical composition  At higher temperatures, bornite

shows a broad compositional range in the system Cu–Fe–S and forms a complete solid solution series with digenite, to a lesser extent with chalcopyrite. Upon cooling, these solid solutions exsolve to form lamellar digenite-bornite intergrowths, in places with small grains of chalcopyrite. Occurrence  As a primary ore mineral, bornite crystallises

from hydrothermal solutions, but also forms as a secondary mineral during weathering of Cu deposits, being concentrated in metal-enriched zones of gossans. Moreover, bornite is an important ore mineral in black shales, especially in the Kupferschiefer (7 Sect. 25.2.11). Weathering of bornite leads to the formation of chalcocite and covellite, CuS, and, finally, of the Cu-carbonates azurite and malachite (7 Sect. 8.4). Economic relevance  Bornite is an important Cu-ore min-

eral (see chalcopyrite).

z Argentite and Acanthite Ag2S

Ag2S, crystal class 4/m3¯2/m, is the high-temperature polymorph of Ag2S. It may show cubic crystal forms, frequently with cubic habit, i.e., with dominant cube {100}, whereas the octahedron {111} and other forms are subordinate. Upon cooling below 173 °C, argentite transforms readily into fine-grained lamellar twins of monoclinic acanthite, Ag2S, crystal class 2/m, the stable polymorph of Ag2S. These aggregates preserve the outer crystal forms of the initial cubic argentite which, therefore, in fact is a paramorph of polycrystalline acanthite after argentite. At temperatures below 173 °C, primary formation of acanthite may lead to monoclinic crystals, in many cases with spiky or dendroidal habit, elongated parallel to c. In general, however, primary or secondary acanthite is anhedral and forms massive or powdery ore, the latter being known as “earthy silver” or “earthy argentite”. Acanthite and native silver may replace each other mutually. Physical properties Cleavage

absent

Fracture

ductile, sliceable with a knife, in ancient times used to emboss coins

Hardness

2–2½

Density

7.3

Optical properties

opaque; on fresh fracture surface, leaden grey with metallic lustre; weathering leads to dim coating and black tarnish and, finally, to decomposition to a dark leaden grey powder

Streak

Dark leaden grey with metallic lustre

Crystal structures  In both Ag2S polymorphs, [4]-coordinated S atoms constitute a body-centred cubic array. In acanthite, Ag[3]S[4] 3 triangles are interconnected to form sheets nearly parallel to (010) that are mutually linked by Ag[2] atoms. In argentite, however, the Ag atoms are not restricted to fixed positions but display a disordered, “liquid-like” distribution, giving rise to a very high ionic and electronic conductivity (e.g., Kashida et al. 2003). Occurrence  As a primary ore, argentite is precipi-

tated from hydrothermal solutions, frequently forming tiny inclusions in galenite. Secondary transformation of argentite leads to the low-temperature Ag2S polymorph acanthite which, in addition, may be formed by weathering of Ag deposits and concentrated in metal-enriched zones of gossans.

Economic relevance  Argentite and acanthite are impor-

tant Ag-ore minerals, especially as silver-carriers included in galena. For use of Ag as metallic raw material see 7 Sect. 4.1.

5

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Chapter 5 · Sulfides, Arsenides and Complex Sulfides (Sulfosalts)

z Pentlandite (Ni,Fe)9S8 Crystal form  The cubic mineral, belonging to crystal class 4/m3¯2/m, does not show crystal faces, but typically occurs as flame-like exsolution lamellae in pyrrhotite, Fe1−xS. Moreover, pentlandite may form individual or granular aggregates together with pyrrhotite and chalcopyrite (. Fig. 21.7). Physical properties

5

Cleavage

distinct after {111}

Fracture

brittle

Hardness

3½–4

Density

4.6–5.0

Optical properties

opaque; bronze-yellow with metallic lustre

Streak

black

Characteristic property

in contrast to pyrrhotite, pentlandite is non-magnetic

Crystal structure and chemical composition  Pentlandite crystallises in a spinel-type structure (. Fig. 7.2) with

cubic close packing of the S atoms and the metal atoms in the tetrahedral and octahedral interstices. The Ni:Fe ratio is close to, but not exactly, 1:1, minor amounts of Co are common. Occurrence  In most cases, pentlandite forms by second-

ary exsolution from an initial high-temperature pyrrhotite-pentlandite solid solution that crystallised, together with chalcopyrite and pyrite, from sulfide liquids in layered magmatic intrusions, e.g., in the world-renowned Ni deposits of Sudbury, Ontario, Canada, and Bushveld, South Africa (7 Sect. 21.3.1). In some cases, separate grains of pentlandite and pyrrhotite may crystallise from the sulfide liquid (. Fig. 21.7).

a higher durability against corrosion. Moreover, nickel is used, e.g., as an alloy metal for high-temperature materials in power-stations, for constructing guide-blades in turbines and chemical instruments, for electro-plating and coating as well as for catalysts. For coinage, nickel has been used since antiquity. The name “nickel” was used by the Saxonian miners as a swear word, referring to an evil mountain troll, when they recovered nickel sulfides instead of the much more valued silver ore, in the deeper parts of hydrothermal Bi–Co–Ni–Ag–U veins in the western Erzgebirge, Germany.

5.2  Metal Sulfides and Arsenides

with M:S = 1:1 (. Table 5.2)

z Galena (Galenite) PbS Crystal form and habit  Crystal class 4/m3¯2/m; well-

developed crystals of considerable size are common in many localities worldwide. Predominant crystal forms (. Figs. 5.1, 5.2) are the cube {100} and the octahedron {111} alone or in combination that are called cuboctahedron if the faces of both forms display about the same size (. Fig. 5.1a). Besides, the dodecahedron {110}, the trisoctahedron {221} and other forms are commonly developed. Generally, however, “galena” is anhedral, forming granular or sparry, in many places fine-grained to dense ore sheets which, if strongly deformed, are known as striate galena. Deformation twins are abundant. Physical properties Cleavage

{100}, perfect

Hardness



Density

7.4–7.6

Optical properties

opaque; leaden grey, sometimes dim tarnish colours; metallic lustre on fresh fracture surfaces

Streak

greyish black

Economic relevance  Pentlandite is the most important

ore mineral of nickel, a widely used steel-refining metal. Nickel-steel, containing 2½–3½% Ni, is stronger and has

. Fig. 5.1  Crystal forms and habits of “galena”; a so-called cuboctahedron, i.e. a combination of cube {100} and octahedron {111}, both being similar in size; b octahedron {111} dominant, cube {100} subordinate; c, d various combinations of cube {100}, octahedron {111}, dodecahedron {110}, trisoctahedron {221} and icositetrahedron {322}

91

5.2 · Metal Sulfides and Arsenites with M:S = 1:1

. Fig. 5.3  Crystal structure of “galena”

Chemical composition  Next to the principal cation Pb, galena can contain small amounts of Ag, reaching in total 0.01 and 0.3, in rare cases up to 1 wt%. Much of this Ag is, however, present as minute inclusions of different Ag minerals, predominantly of Ag-rich fahlore (freibergite, see below), polybasite, (Ag,Cu)16Sb2S11, proustite–pyrargyrite (see below), native Ag, argentite and acanthite. Occurrence  Together with sphalerite and other sulfide

. Fig. 5.2  Crystals of galena displaying the typical combination of cube {100} and octahedron {111}, grown together with calcite, Dalnegorsk, Primorsky Krai, Eastern Siberia; width of view is c. 4 cm . Table 5.2  Metal sulfides and arsenides with M:S = 1:1 Mineral

Formula

Metal content

Crystal class

“Galena”

PbS

Pb 86.6%

¯ 4/m32/m

Sphalerite

α-ZnS

Zn 67.1%

¯ 43m

Wurtzite

β-ZnS

Zn 67.1%

6mm

Chalcopyrite

CuFeS2

Cu 34.6%

¯ 42m

Enargite

Cu3AsS4

Cu 48.3%

mm2

Niccolite

NiAs

Ni 43.9%

6/m2/m2/m

Pyrrhotite

FeS–Fe5S6

6/m2/m2/m

Covellite

CuS

Cu 66.4%

6/m2/m2/m

Cinnabarite

HgS

Hg 86.2%

32

minerals, galena occurs worldwide in hydrothermal veins (. Fig. 23.9). In replacement deposits, it is formed by reaction of hydrothermal fluids with limestones or dolomites. Among these, the stratabound Pb-Zn replacement deposits of Mississippi Valley type (MVT) are of high economic interest. The same holds true for stratabound sedimentary-exhalative Pb-Zn deposits (SEDEX).

Economic relevance  Galena is the most abundant and most important Pb ore mineral. Moreover, owing to the wealth of noble Ag-sulfides, frequently included in galena as tiny grains, it is also the most important Ag ore. Lead as metallic raw material  Plates of lead are used in

accumulators or for shielding against radioactive or X-rays. Moreover, Pb is used for cables and as alloy metal. Up to the early 2000s, tetraethyle lead (TEL) was used as antiknock agent in gasoline, but has been banned, in nearly all industrialised countries, because of its high toxicity.

z Sphalerite α-ZnS Crystal form and habit  Crystal class 4¯ 3m, i.e., compared

Crystal structure  The Pb[6]S[6] structure (. Fig. 5.3) con-

sists of two cubic, face-centred lattices of Pb and S, respectively, that are placed into each other after shifting at a half lattice constant ½a, i.e., along the edge of the cube. Consequently, each Pb atom is bound to 6 S atoms in octahedral coordination and vice versa, i.e., both Pb and S are [6]-coordinated. The bonds between Pb and S are predominantly metallic. Geometrically, the PbS structure conforms to the structure of halite Na[6]Cl[6] which, however is a good example for pure heteropolar bonding (. Fig. 6.2).

to galena, the cubic symmetry is lower (hemiedry). Crystal faces are common, in many cases in tetrahedral habit (. Fig. 5.4a), where the positive tetrahedron {111} and the negative tetrahedron {11¯1} can be distinguished by their different lustre and by means of etch figures. Another important crystal form is the dodecahedron {110}, often in combination with the tetrahedra {111} and {11¯1} (. Fig. 5.4b) as well as the positive and the negative tristetrahedron {311} and {31¯1}, respectively. Due to repeated twinning at {111}, determination of the crystal forms is often difficult (. Fig. 5.4c). In most cases, sphalerite is

5

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Chapter 5 · Sulfides, Arsenides and Complex Sulfides (Sulfosalts)

5 . Fig. 5.4  Crystal forms and habits of sphalerite; a tetrahedral habit with dominant positive tetrahedron {111} and subordinate negative ¯ b rhombohedral habit with dominant dodecahedron {110} and minor tetrahedra {111} and {111}; ¯ c lamellar twins intergrown tetrahedron {111}; parallel to the (111) face

anhedral, forming granular or sparry ore sheets. An interesting structural variety is “fibrous blende”, colloform, layered intergrowths of columnar to fibrous sphalerite and/or wurtzite (see below) ± “galena” ± pyrite ± marcasite that have crystallised, at relatively low temperatures, possibly from a colloidal solution. Physical properties Cleavage

perfect in {110}, brittle

Hardness

3½–4

Density

3.9–4.1

Optical properties

transparent to translucent, at least in thin splinters, never fully opaque; white (rare), yellow, red, oily-green or black; lustre resinous to adamantine, especially on cleavage planes

Streak

yellowish to dark brown but never black

Crystal structure  The sphalerite structure (. Fig. 5.6)

consists of two cubic, face-centred lattices of Zn and S, respectively, that are placed into each other after shifting at ¼ of the space diagonal. Consequently, each Zn atom is bound to 4 S atoms in tetrahedral coordination and vice versa, i.e., both Zn and S are [4]-coordinated. The sphalerite structure can be derived from the diamond structure (. Fig. 4.11a) in substituting half of the C atoms by Zn, the other half by S. Thereby, the symmetry of diamond, crystal class 4/m3¯ 2/m, is reduced to 4¯ 3m of sphalerite. In contrast to diamond, the binding forces in sphalerite are not purely covalent (atomic) but contain a considerable amount of ionic (heteropolar) character conforming to the polar covalent bond type. The differences in crystal structure and bond type explain that the cleavage planes in sphalerite are parallel to {110} as compared to diamond, in which they are parallel to {111} (. Fig. 5.5).

. Fig. 5.5  Polished section of “fibrous blende”, colloform, layered intergrowths of sphalerite (violet grey), wurtzite (light brown) and pyrite (golden yellow), Olkusz, Poland

Chemical composition  In general, sphalerite contains Fe, thus forming solid solutions between ZnS and FeS with a maximum of 26 mol% FeS. With increasing Fe content sphalerite becomes darker in colour. Besides, sphalerite can contain some Mn and Cd, and traces of the rare metals In, Ga, Tl and Ge.

93

5.2 · Metal Sulfides and Arsenites with M:S = 1:1

Crystal structure  Like in the sphalerite structure, Zn

is bound to four S and vice versa, i.e., Zn and S are both [4]-coordinated. However, the respective sequences of the closest packed layers of S atoms is different in both structures, i.e., cubic close packing in sphalerite and hexagonal close packing in wurtzite.

. Fig. 5.6  Crystal structure of sphalerite

Stability relations  Although wurtzite is the high-temperature polymorph of ZnS, it can crystallise metastably at low temperatures from acid solutions, its formation being favoured by high Cd contents. Occurrence  In most cases, wurtzite is a constituent of

Stability relations  At atmospheric pressure and room

temperature, sphalerite is the stable ZnS modification, whereas the high-temperature ZnS polymorph wurtzite is metastable.

Occurrence  Together with galena, sphalerite occurs in hydrothermal veins (. Fig. 23.9) and in hydrothermal

replacement deposits formed at the expense of limestone or dolomite. Moreover, it is found in stratabound MVT and SEDEX type Pb–Zn deposits. The present-day submarine-exhalative precipitation of sphalerite can be observed in the so-called black smokers that are being discharged at the sea floor along mid-oceanic ridges (. Figs. 23.10a, 23.11). Economic relevance  Sphalerite is the most important and most abundant ore mineral of Zn. In addition, the accompanying metals mentioned above, especially Cd, can be extracted during smelting of sphalerite. As Cd is extremely poisonous, zinc works should be run with extreme care, obeying strict charges for pollution control. Zinc and cadmium as metallic raw materials  As an

anti-corrosive, zinc is applied to iron or steel surfaces by hot galvanising or electroplating. Moreover, Zn is an important alloy metal, especially with Cu to produce brass, and is used in galvanic elements. In smaller amounts, pulverised sphalerite serves as a pigment in lithopones. Cadmium, extracted as a by-product of sphalerite smelting, is used in Ni–Cd or Ag–Cd batteries, as component in fusible alloys, e.g., Wood’s metal with a melting point of 70 °C, in nuclear technology, as anti-corrosive and as pigment.

fibrous blende, formed at low temperatures in hydrothermal veins and in Pb–Zn replacement deposits, e.g., in Joplin, Missouri, USA, and Upper Silesia, Poland.

Economic relevance  Locally important Zn ore mineral.

z Chalcopyrite CuFeS2 Crystal form and habit  Crystal form and habit Crystal class 4¯ 2m; in many cases, crystals show the positive and negative tetragonal disphenoids {111} and {11¯1} (. Fig. 5.7a) or the steeper disphenoid {772}, often in combination with the tetragonal scalenohedron {212} (. Fig. 5.7b). Highly distorted or twinned crystals, intergrown at (111), are common (. Fig. 5.7c). In most cases, however, chalcopyrite occurs as compact ore sheets of variable grain size. Physical properties Cleavage

absent

Hardness

3½–4½, distinctly lower than that of pyrite FeS2

Density

4.1–4.3

Optical properties

opaque with metallic lustre; greenish-yellow to dark yellow, brasscoloured, often with multi-coloured tarnish

Streak

black to greenish-black

z Wurtzite β-ZnS Crystal form and habit  Crystal class 6mm; frequently tufted or radial-fibrous aggregates, often intergrown with sphalerite to form fibrous blende (. Fig. 5.5). Pyramidal, short prismatic crystals are imperfect and relatively rare. Physical properties  are similar to sphalerite. Cleavage

perfect parallel to the prism face {101¯0}, distinct parallel to the base-pinacoid {0001}

. Fig. 5.7  Crystal forms and habits of chalcopyrite; a dominant tetragonal disphenoid {111}; b combination of the steep tetragonal disphenoid {722} with the tetragonal scalenohedron {212}; c twin similar to the spinel law

5

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Chapter 5 · Sulfides, Arsenides and Complex Sulfides (Sulfosalts)

5

. Fig. 5.8  Crystal structure of chalcopyrite

Crystal structure  The crystal structure of chalcopyrite (. Fig. 5.8) can be derived from that of sphalerite by duplicating the unit cell and substituting two Zn cations by Cu and Fe. The close structural affinity between chalcopyrite and sphalerite leads to orientated intergrowths of both minerals, visible under the microscope. In part, these textures testify to exsolution from primary high-temperature solid solutions as revealed by experimental investigations in the system CuS–FeS–ZnS (Kojima and Sugaki 1984, 1985). Occurrence  Chalcopyrite is the most common of all

Cu minerals. Together with pyrrhotite, pentlandite and pyrite, it crystallised from sulfide melts in layered intrusions, such as Sudbury, Canada, or Bushveld, South Africa. Chalcopyrite occurs in hydrothermal veins and impregnations, especially in porphyry copper ores (7 Sect. 23.2.4), as well as in volcanogenic massive sulfide (VMS) deposits (7 Sect. 23.5.2). Chalcopyrite is being formed together with other sulfide ore minerals in so-called black smokers by submarine-exhalative activity on the seafloor along midocean ridges (. Figs. 23.10b, 23.11). Moreover, chalcopyrite is locally concentrated in black shales, e.g., the Kupferschiefer of Germany and Poland (7 Sect. 25.2.11). Upon weathering, chalcopyrite decomposes to form inhomogeneous secondary Cu-minerals, such as ores consisting of azurite, Cu3(CO3OH)2, malachite, Cu2CO3(OH)2, cuprite, Cu2O, covellite, CuO, bornite, goethite FeOOH and others. Economic relevance  Next to the minerals of the chalcoc-

ite-digenite group, chalcopyrite is the most important Cu ore mineral. Copper as metallic raw material  Utilisation of copper as a

metallic raw material dates back to about 9000 BC. Initially it was processed by cold working of rare native copper, but not much later, annealing, smelting and the lost wax method were invented, thereby making it possible to process Cucompounds, especially the Cu-sulfides described in this

chapter. Cu–Sn alloys known as bronze were produced about 4500 BC by the Vinča culture on the Balkan Peninsula, but general utilisation in the same area started only between 3700–3200 BC, at the dawn of the Bronze Age. In contrast, the Cu–Zn alloy brass, although known in ancient Greece, has become a common metallic material, besides bronze, since the times of the Roman Empire, and is still widely in use. Nowadays, not more than about 5% of the worldwide Cu production is processed in Cu alloys, especially brass for component parts of engines, musical instruments and tongues of organ pipes, while cupronickel, Cu75Ni25, is a coining metal alloy. Moreover, Cu alloys may serve as fungicides applied in agriculture or as antibiotics used in medicine. Owing to its excellent electric conductivity, about 60% of the annual production is used as pure Cu metal for electrical wires and cables and ca. 15% for machinery, especially electric motors. As Cu is durable, corrosion-resistant and weatherproof, ca. 20% of the total production are still used for conventional roofing and plumbing, an application dating back to antiquity. A thin coating of green patina, a mixture of Cu carbonates and sulfates formed by natural weathering, make copper roofs even more resistant and thus popular. Cu oxides and carbonates cause green or brown colours in glasses and ceramics. z Enargite Cu3AsS4 Crystal form and habit  Crystal class mm2; the orthor-

hombic, pseudo-hexagonally shaped prismatic crystals are elongated parallel to c and show vertical striation. Others are tabular parallel to (001) and may be intergrown to form multiple twins. Commonly, radial, sparry or granular aggregates of enargite form massive ore sheets. Physical properties Cleavage

{110} perfect, {100} and {010} distinct

Fracture

rugged, brittle

Hardness

3

Density

4.45

Optical properties

opaque; grey to greyish-black; semi-metallic lustre

Streak

black

Crystal structure  Similar to wurtzite, whereby Zn is

replaced at ¾ by Cu, at ¼ by As. This leads to a decrease in symmetry from hexagonal (6 mm) to orthorhombic (mm2).

Chemical composition  Up to about 6% As can be replaced

by Sb, whereas some Fe and Zn may substitute for Cu.

Occurrence  In hydrothermal veins, replacement and

impregnation ores.

Economic relevance  Locally, enargite is an important Cu

ore mineral.

5.2 · Metal Sulfides and Arsenites with M:S = 1:1

95

z Niccolite (Nickeline) NiAs Crystal form  Crystal class 6/m2/m2/m. Niccolite shows rarely hexagonal crystal forms and, commonly, occurs as granular, massive ore. Physical properties Cleavage

¯ and {0001}, rarely visible ∥{1010}

Fracture

conchoidal, brittle

Hardness

5–5½

Density

7.78

Optical properties

opaque; light copper-coloured with darker tarnish; metallic lustre

Streak

brownish-black

Crystal structure  The NiAs structure is an important

structural type. It can be described as hexagonal packing of the As atoms, the octahedral interstices of which being filled with Ni atoms. To a limited extent, As can be replaced by Sb.

Occurrence  In hydrothermal veins. Economic relevance  Locally, niccolite is an important Ni

ore mineral.

. Fig. 5.9  Superimposed crystals of pyrrhotite together with quartz (rock crystals), Primorsky Krai, Eastern Siberia; Width of view is c. 3 cm

z Pyrrhotite FeS–Fe5S6

Crystal structure  Pyrrhotite has the NiAs structure, where

Stability relations, crystal form and habit  The high-

temperature modification of pyrrhotite, crystal class 6/m2/m2/m, is stable above ~300 °C. At lower temperatures, numerous hexagonal, orthorhombic and monoclinic varieties (polytypes) exist, depending on temperature and Fe content. Well-developed crystals are rare. They may occur as hexagonal plates with dominant basal pinacoid {0001} and small hexagonal prism {112¯ 0}, in many instances arranged to aggregates resembling rosettes or book-stacks (. Fig. 5.9). In most cases, however, anhedral, granular or flaky pyrrhotite is interspersed or forms massive ore. Physical properties Cleavage

¯ visible in coarse∥{0001} and {1120}, grained samples

Fracture

conchoidal, brittle

Hardness

3½–4½

Density

4.6–4.9

Optical properties

opaque with metallic lustre; light bronzecoloured, with dim brown tarnish

Streak

greyish-black

Additional property

usually ferromagnetic

S takes the As, and Fe the Ni positions. In general, pyrrhotite is Fe deficient, relative to S, ranging in composition between Fe10S11 and Fe5S6. The Fe atoms missing leave vacancies in the crystal structure. In troilite, FeS, observed only in meteorites and lunar rocks, the Fe positions are fully occupied. Chemical composition  At high temperatures, a com-

plete monosulfide solid solution (mss), (Fe,Ni)S, exists as shown by experimental investigations in the system Fe–Ni–S (Fleet 2006, and references therein). Upon cooling the contents of Ni and minor Co exsolve to pentlandite, (Ni,Fe,Co)9S8, lamellae within a pyrrhotite host (. Fig. 21.7).

Occurrence  Pyrrhotite is a common sulfide ore constit-

uent and commonly found as accessory phase in igneous and metamorphic rocks. Higher concentrations of nickel-bearing pyrrhotite crystallised, together with chalcopyrite, from sulfidic melt droplets in layered mafic to ultramafic intrusions, e.g., in the well-endowed Ni deposits of Sudbury, Ontario, Canada and Bushveld, South Africa. Pyrrhotite can also occur in metamorphosed ore deposits of igneous origin, in pegmatites and in hydrothermal veins.

5

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Chapter 5 · Sulfides, Arsenides and Complex Sulfides (Sulfosalts)

Economic relevance  Pentlandite-bearing pyrrhotite is a very important Ni-ore, whereas pure, Ni-free pyrrhotite is hardly mined as an iron ore, but is used, still to day, to produce Fe-sulfate and jeweller’s rouge, utilised for polishing of gemstones.

z Covellite CuS

5

Crystal form and habit  Crystal class 6/m2/m2/m. Tabular crystals with hexagonal crystal faces are rare. Generally, covellite is anhedral, forming fine-grained or sparry aggregates. Physical properties Cleavage

{0001}, perfect

Hardness

1½–2

Density

4.60–4.76

Optical properties

transparent in thin flakes; semi-metallic lustre; bluish black to indigo blue. Owing to an extremely strong dispersion of refraction, covellite changes its colour when immersed in liquids: It appears violet blue in water, reddish violet in cedar wood oil and scarlet to orange-red in liquids with higher refractive indices (7 Sect. 1.5.3)

Streak

bluish black

Spionkopite, Cu1.12S, and yarrowite, Cu1.39S, intermediate in composition between covellite and digenite, display crystal structure and optical properties similar to covellite but do not change their colour on immersion.

The crystal structure of covellite is made up by double sheets of CuS4 tetrahedra parallel to (0001), linked by S–S bonds parallel to c. These units alternate with Cu3S sheets, in which Cu and S are mutually connected in triangular coordination. Occurrence  Covellite is a secondary Cu-mineral, com-

monly formed in subordinate amounts, by secondary alteration of Cu-sulfide ores, but is not principal ore mineral in Cu deposits.

z Cinnabarite HgS Crystal form and habit  Crystal class 32; distinct rhombo-

hedral or tabular crystals are rare. In general, cinnabarite occurs in granular, dense or earthy masses, impregnating country rocks.

Physical properties Cleavage

¯ rather perfect {1010},

Hardness



Density

8.10

Optical properties

transparent in thin splinters; adamantine lustre; red, but often clouded by inclusions of bitumen and earthy matter, an example being the “hepatic cinnabar” from Idrija Mine, Slovenia

Streak

red

Crystal structure  The crystal structure can be derived from the PbS structure that is deformed along the space diagonal and in which Hg takes the position of Pb. Occurrence  Cinnabarite occurs in the form of strata-

bound, hydrothermal impregnations in tectonically deformed country rocks, prime examples being the former mines of Almadén in southern Spain, Idrija in Slovenia and New Almadén and New Idria in California (7 Sect. 23.5.3).

Economic relevance  Cinnabarite used to be the most important Hg ore mineral. In the past, from Neolithic times, it was widely used as a pigment. Through the centuries, both these applications played an important role, starting in Roman times. In Central America, cinnabarite pigments became popular in the ancient Olmec and Maya civilisations, a prominent example being the Tomb of the Red Queen in Palenque, Mexico (600–700 AD). Cinnabarite was also used in the famous Chinese lacquerware that came into fashion during the Song Dynasty (960–1279).

Technical use of mercury  Due to its high toxicity, the tech-

nical application of mercury in physical instruments (e.g., thermometers), in electric and electronic instruments, in medicine, agriculture, and gold mining has become severely limited or prohibited. Controversial is the use of the Hg alloy Hg–Cu–(Ag,Au) amalgam as dental filling. All these health risks led effectively to the end of commercial cinnabarite mining.

5.3  Metal Sulfides, Sulfarsenides

and Arsenides with M:S = 1:2

z Stibnite Sb2S3 Crystal form and habit  Crystal class 2/m2/m2/m; orthorhombic crystals, elongated parallel to c with spike- or needle-shaped habit, commonly show different faces,

97

5.3 · Metal Sulfides, Sulfarsenides and Arsenides with M:S = 1:2

Occurrence  In hydrothermal veins. Economic relevance  Stibnite is the principal ore mineral

of Sb, used as an alloy metal, especially in lead and tin alloys, e.g., for characters, solders and pellets as well as for accumulators. Pure antimony is applied in semiconductor techniques.

z Molybdenite MoS2 Crystal form and habit  Crystal class 6/m2/m2/m; hexagonal, though imperfect plates, in most cases, however, in flaky or scaly aggregates. Physical properties Cleavage

∥(0001) perfect, forming ductile, inelastic cleavage flakes

Hardness

1–1½

Density

4.62–4.73

Optical properties

opaque; lead-grey; metallic lustre

Streak

dark grey

Additional characteristic

greasy and losing colour on touch

Crystal structure  Hexagonal layer structure with MoS2 lay. Fig. 5.10  Stibnite, acicular habit combination of dominant rhombic prism {110}, rhombic pinacoid {010}, both elongated ∥c, and two different rhombic pyramids

especially prism {110}, rhombic dipyramid {131} and pinacoid {010} (. Fig. 5.10), mostly with longitudinal striation. Stibnite crystals are often bent or twisted parallel to b, slightly turned around c, and may form tangled, tuft-like or radial aggregates. Most commonly, stibnite occurs as granular or dense masses (. Table 5.3).

ers parallel to (0001) the valencies of which are saturated. The individual layers are attached to each other by weak van der Waal bonds, resulting in the perfect cleavage parallel to (0001).

Chemical composition  Molybdenite, virtually pure MoS2,

can contain small amounts of rhenium, up to 0.3 wt%.

Occurrence  In hydrothermal impregnation deposits (7 Sect. 23.2.3), especially in porphyry molybdenite ores

but also in pegmatite dykes.

Cleavage

perfect ∥{010}. Commonly, translation in (010) parallel to the c-axis leads to horizontal striation on the slightly undulating cleavage planes

Hardness

2

Density

4.52–4.62

Economic relevance  Molybdenite is the principal ore mineral of Mo, widely used as steel-refining metal, especially in high-temperature resistant alloys, but also in cast iron. Moreover, it is applied in electrical engineering. Owing to its high melting point of 2620 °C, Mo is an important metal used for construction of missiles and nuclear power plants. Due to its very low hardness and perfect cleavage, molybdenite serves, either on its own, or as a component, in dry lubricating materials.

Optical properties

opaque with strong metallic lustre; lead-grey with blackish to bluish tarnish

z Pyrite FeS2

Streak

dark lead-grey

Physical properties

Chemical composition  Stibnite can contain small amounts of Au, Ag, Fe, Pb and Cu. Crystal structure  Double chains parallel to c, conforming

to the elongated habit.

Crystal form and habit  Crystal class cubic didodecahedral 2/m3¯; commonly as well-developed crystals, showing a variety of different crystal forms, especially cube {100}, pentagonal dodecahedron (pyritohedron) {210}, octahedron {111} and didodecahedron {321}, often in combinations (. Figs. 5.11, 5.12). In most cases, the cube faces

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Chapter 5 · Sulfides, Arsenides and Complex Sulfides (Sulfosalts)

. Table 5.3  Metal sulfides, sulfarsenides and arsenides with M:S ≤ 1:2

5

Mineral

Formula

Metal content

Crystal class

Stibnite

Sb2S3

Sb 71.4%

2/m2/m2/m

Molybdenite

MoS2

Mo 59.9%

6/m2/m2/m

Pyrite

FeS2

Fe 46.6% S 53.8%

2/m3¯

Marcasite

FeS2

Fe 46.6% S 53.8%

2/m2/m2/m

Arsenopyrite

FeAsS

Fe 34.3% As 46.0%

2/m

Cobaltite

(Co,Fe)AsS

Co + Fe 35.4%

23

Löllingite

FeAs2

As 72.8%

2/m2/m2/m

Safflorite

CoAs2

Co 28.2%

2/m2/m2/m

Rammelsbergite

NiAs2

Ni 28.2%

2/m2/m2/m

Skutterudite

(Co,Ni)As3

Co up to 24%

2/m3¯

Ni-Skutterudite (chloanthite)

(Ni,Co)As3

Ni up to 28%

2/m3¯

{100} of pyrite show a growth striation, due to a multiple alternation of the {100} and {210} faces. This indicates that, within the cubic system, pyrite displays a relatively low symmetry with twofold rather than fourfold rotation axes parallel to a1, a2, a3. Penetration twins with a twin axis [100] are common (. Fig. 5.11f). In many places, pyrite forms granular aggregates, in which crystal faces are not or only poorly developed. Physical properties Cleavage

{100}, indistinct

Fracture

conchoidal, brittle

Hardness

6–6½, unusually hard for a sulfide (important for the distinction from chalcopyrite!)

Density

4.95–5.02

Optical properties

opaque with metallic lustre; bright brass yellow, in cases with multi-coloured tarnish

Streak

greenish to brownish black

Distinction from native gold

gold is much softer and ductile, it displays a golden streak and a golden to whitish-golden lustre, depending on the Ag content

Chemical composition  Fe can be replaced by small

amounts of Ni or Co.

Occurrence  Pyrite is by far the most common sulfide

mineral. It has a wide stability field and thus may occur wherever the geochemical prerequisites are given. Pyrite is a constituent of most sulfide deposits and may form huge accumulations, especially massive sulfide deposits. Although pyrite cannot crystallise from a melt, it is a common secondary accessory mineral in many igneous and also in metamorphic rocks. Pyrite may occur also in sedimentary rocks, such as in black shales (7 Sect. 25.2.11), where it is deposited, as impregnations or concretions, under anaerobic, oxygen-free or oxygen-poor conditions. In many cases, the thiospinel greigite, Fe2+Fe3+S4, serves as precursor mineral. When subjected to atmospheric weathering, pyrite decomposes, via various

Crystal structure  The pyrite structure (. Fig. 5.13) is very similar to that of galena, PbS (. Fig. 5.3). In pyrite, the Pb

positions are taken by the Fe atoms, whereas the centres of dumb-bell shaped S2 groups are placed in the S positions of the PbS structure. The axes of the S2 dumbbells are parallel to the threefold axes but in different orientations. Therefore, the symmetry is lower than in the crystal class of galena 4/m3¯2/m. Each Fe atom is neighboured by 6 S atoms at the same distance. Within the S2 dumbbells, atomic bonding is predominant, whereas the Fe–S2 bonds are metallic.

. Fig. 5.11  Crystal forms and habits of pyrite; a cube {100} with stri¯ b pentagonal dodecahedron ation indicating lower symmetry 2/m3; {210}; c combination pentagonal dodecahedron {210} and cube {100}; d, e combination octahedron {111} and pentagonal dodecahedron {210}; f penetration twin with twin axis [001]

99

5.3 · Metal Sulfides, Sulfarsenides and Arsenides with M:S = 1:2

gold ore, as it may contain tiny inclusions of native (“invisible”) Au. z Marcasite FeS2 Crystal form and habit  Marcasite is the orthorhombic

modification of FeS2 with crystal class 2/m2/m2/m. In general, the habit of individual crystals is tabular after {100} (. Fig. 5.14a), more rarely prismatic parallel to c. In most cases, marcasite crystals are multiply twinned to form cockscomb- or spear-shaped groups (. Fig. 5.14b, c). Moreover, marcasite abundantly occurs as radially structured aggregates, as incrustations covering other minerals or as dense masses. Physical properties

. Fig. 5.12  Crystal group of pyrite with combination of pentagonal dodecahedron {210} and cube {100} leading to the characteristic striation, Pasto Bueno, Peru; Width of view is c. 9 cm

Cleavage

{110}, imperfect

Fracture

uneven, brittle

Hardness

6–6½

Density

4.89, slightly lower than for pyrite

Optical properties

opaque with metallic lustre; as compared to pyrite, marcasite is more greenish yellow with greenish tarnish

Streak

greenish to blackish grey

Crystal structure  Similar to the pyrite structure but with

lower symmetry.

Occurrence  Marcasite is formed in low-temperature

hydrothermal environments, such as lead-zinc replacement deposits of MVT type (7 Sect. 23.6.2) and gold-pyrite replacement deposits of Carlin type (7 Sect. 23.3.4), in bituminous shales, e.g., the Permian Kupferschiefer in Germany and Poland and the Central African Copper Belt (7 Sect. 25.2.11), and as concentrations in marl and lignite. Upon weathering of pyrrhotite, marcasite crystallises as an intermediate product, in part forming pseudomorphs. . Fig. 5.13  Crystal structure of pyrite

intermediate compounds, to Fe3+-oxyhydrates such as goethite α-FeOOH, and lepidocrocite, γ-FeOOH, which are the main constituents of a fine-grained mixture called limonite (7 Sect. 7.5). Commonly, these FeOOH-minerals form pseudomorphs after pyrite. Economic relevance  Per se, pyrite is not an iron ore, but is mined for production of sulfurous acid, an extremely important starting compound in various branches of chemical industry. However, the roasted pyrite residues can be smelted to produce iron and, in addition, may serve as polishing powder or pigment. Locally, pyrite is mined as

So far, the stability relations between pyrite and marcasite are not well-understood. Experimental investigations indicate that marcasite is metastable above ~150 °C and is transformed into pyrite above ~400 °C. Accordingly, marcasite is formed in nature at relatively low temperatures, preferably from acid solutions.

. Fig. 5.14  Crystal forms and habits of marcasite; a single crystal; b, c cockscomb-shaped multiple twins

5

100

5

Chapter 5 · Sulfides, Arsenides and Complex Sulfides (Sulfosalts)

z Arsenopyrite FeAsS

z Cobaltite CoAsS

Crystal form and habit  Crystal class 2/m; the monoclinic

Crystal form and habit  Crystal class mm2 (orthorhombic-pyramidal); well-developed crystals show pseudocubic forms: entagondodecahedron (pyritohydron) {210}, commonly combined with the octahedron {111}, more rarely with the cube {100}. Similar to pyrite, cube faces show striation. In most cases, cobaltite occurs in compact, granular aggregates.

(pseudo-rhombic) crystals are elongated parallel to c or, less commonly b or a (. Fig. 5.15a, b) and, in the simplest case, may show a combination of the prism faces {110} and {014} only, but additional forms may be present. A characteristic feature is striation on {014}. Crystals are commonly twinned, whereas triplets are rarer. In many cases, arsenopyrite forms granular aggregates. Physical properties Cleavage

{110}, somewhat distinct

Fracture

uneven, brittle

Hardness

5½–6

Density

6.07

Optical properties

opaque with metallic lustre; pewter white, with multi-coloured or dark tarnish

Streak

black

Crystal structure  The crystal structure of arsenopyrite can be derived from the marcasite structure where half of the S atoms are replaced by As. This leads to a reduction of the symmetry from orthorhombic to monoclinic. Chemical composition  The As:S ratio deviates from the

ideal composition. Moreover, Fe can be partly replaced by Co or Ni. Like in pyrite, arsenopyrite may contain tiny inclusions of native (“invisible”) gold.

Occurrence  In hydrothermal veins. Economic relevance  Arsenopyrite is the most important

ore mineral for As.

Physical properties Cleavage

{100}, not always distinct

Fracture

uneven, brittle

Hardness



Density

6.33

Optical properties

opaque; metallic lustre; silvery with metallic reddish shade and reddish-grey tarnish

Streak

greyish black

Crystal structure  Similar to the pyrite structure but lower

symmetry as the dumbbell-shaped S2 groups are replaced by As–S groups.

Chemical composition  The theoretical Co content is

35.4%. However, Co can be partly replaced by Fe, reaching values of up to 10%.

Occurrence  Cobaltite can occur in hydrothermal veins

and replacement deposits, such as skarn ores.

Economic relevance  Cobaltite is an important ore mineral of cobalt, widely used as an alloy metal to produce Co-steel for abrasion-resistant tools and heat-resistant permanent magnets as well as high-temperature alloys, especially for jet engines. In chemical industries, Co compounds serve for manufacturing catalysts and the cobalt blue pigment for glasses, glazes and lacquers. The name “cobalt” is derived from the German word “Kobold” (=goblin) used by the Saxonian miners who were deeply disappointed to recover cobaltite, instead of silver ore, in the deeper parts of hydrothermal Bi–Co–Ni–Ag–U veins in the western Erzgebirge. Only after the invention of cobalt blue, cobaltite and other Co ore minerals were highly treasured, especially in the 17th and 18th centuries.

z Löllingite FeAs2 . Fig. 5.15  Crystal forms and habits of arsenopyrite; a simple combination of the prism faces {110} and {014}; b habit displaying different faces, elongated parallel to c

Crystal form and habit Crystal class 2/m2/m2/m; löllingite may form prismatic crystals but generally occurs in compact, granular or columnar aggregates.

101

5.3 · Metal Sulfides, Sulfarsenides and Arsenides with M:S = 1:2

Economic relevance  Co-ore mineral.

Physical properties Cleavage

{001}, distinct

Fracture

uneven, brittle

Hardness

5, softer than arsenopyrite

Density

7.0–7.4

Optical properties

opaque; metallic lustre; colour similar to arsenopyrite but somewhat brighter on fresh fracture surfaces, grey tarnish

Streak

greyish black

Crystal structure  Löllingite, safflorite and rammelsbergite

display marcasite structure.

Chemical composition  In contrast to the ideal formula,

the Fe/As ratio is variable. Fe is partly replaced by Co and Ni, As by S and Sb. Gold contents can be elevated due to tiny inclusions of native (“invisible”) gold. Occurrence  In hydrothermal veins and replacement

deposits.

Economic relevance  Löllingite ores are mined for As.

z Safflorite CoAs2 Crystal form and habit  Crystal class 2/m2/m2/m; locally

in tiny, orthorhombic crystals that may form star-shaped triplets, well visible under the ore microscope in reflected light; more commonly, safflorite occurs as compact, granular or radially structured aggregates. Physical properties Cleavage

hardly visible

Fracture

uneven, brittle

Hardness

4½–5½, depending on Fe content

Density

6.9–7.3

Optical properties

opaque; metallic lustre; pewter white, with darker tarnish

Streak

black

Chemical composition  Co may be partly replaced by Fe,

but hardly by Ni, which means that there is no solid solution between safflorite and rammelsbergite.

Occurrence  In hydrothermal veins; safflorite is much

more widespread than assumed in former times, as it was often confused with skutterudite (“smaltite”).

z Rammelsbergite NiAs2 Crystal form and habit  Orthorhombic, crystal class 2/m2/

m2/m, small crystals; under the ore microscope, a fine lamellar structure and additional twinning are visible but no star-shaped triplets as in safflorite.

Physical properties  Similar to safflorite. Chemical composition  Ni is replaced by some Fe but

hardly by Co.

Occurrence  Like safflorite in hydrothermal veins. Economic relevance  Ni-ore mineral.

z Skutterudite—Nickelskutterudite (Chloanthite) (Co,Ni)As3, (Ni,Co)As3

Complete solid-solution.

Crystal form and habit  Cubic, crystal class 2/m3¯; crystal forms may be visible, predominantly cube {100} and octahedron {111}, less frequently rhomb-dodecahedron {110} and pentagonal dodecahedron {210}, or combinations of them. In most cases, the skutterudite solid solutions form massive, dense to granular ore. Physical properties Cleavage

missing

Fracture

uneven, brittle

Hardness

5½–6

Density

6.4–6.8

Optical properties

opaque with metallic lustre; pewter white to light steel grey, dark grey tarnish colours

Streak

greyish black to black

Chemical composition  Some of the Co and Ni are always

replaced by Fe.

Occurrence  In hydrothermal veins. Alteration  On initial weathering, coatings of peach-blossom coloured erythrite (“cobalt bloom”), Co(AsO4)2·8H2O, or green annabergite (“nickel bloom”), Ni(AsO4)2·8H2O, may form, depending on the initial Co/Ni ratio. Economic relevance  Important ore minerals for Co and Ni.

5

102

Chapter 5 · Sulfides, Arsenides and Complex Sulfides (Sulfosalts)

5.4  Arsenic Sulfides

5

Physical properties

z Realgar As4S4

Cleavage

(010), perfect, cleavage laminae flexible but not elastic, sliceable

Hardness

1½–2

Crystal form and habit  Crystal class 2/m; small, monoclinic, short prismatic crystals with vertical striation; generally in granular aggregates, also as thin coatings, frequently together with orpiment (. Table 5.4).

Density

3.49

Optical properties

translucent; lemon-yellow; resinous lustre, on cleavage planes mother-ofpearl lustre

Streak

pale yellow

Physical properties Cleavage

(010) and (210), rather perfect

Fracture

conchoidal, brittle

Hardness

1½–2

Density

3.48

Optical properties

translucent to transparent on edges; red to orange; adamantine-like lustre to greasy shine; decomposes under the influence of light

Streak

orange yellow

[2] Crystal structure  Cradle-shaped groups of As[3] 4 S4 , similar to the puckered S8 rings of sulfur (. Fig. 4.18), are

mutually attracted by weak van der Waals bonds whereas, within the As4S4 rings, covalent bonding predominates.

Chemical composition  70.1 wt% As, 29.9 wt% S. Occurrence  Together with orpiment, realgar is formed

in low-temperature hydrothermal veins or impregnations, deposited from hot springs and as sublimation product from volcanic gases, or is formed by weathering of As- and S-containing ore minerals.

Economic relevance  Owing to its toxicity, realgar is only

rarely used in pyrotechniques, as herbicide and insecticide, or in tannery. In the Roman Empire and during the Renaissance period, it frequently served as pigment for red paint.

z Orpiment As4S6 Crystal form and habit  Crystal class 2/m; the monoclinic

crystals are generally small, either tabular after (010) or of short prismatic habit; predominantly in massive ore sheets or as powdery efflorescence product; commonly together with realgar.

Crystal structure  Chains of AsS3 pyramids parallel to c are linked, by shared S atoms, to form As2S4 layers parallel to (010). These are mutually attracted by weak van der Waals bonds whereas, within the layers, covalent bonding is prevalent. Chemical composition  Ideally 61 wt% As, 39 wt% S, As

may be partly replaced by Sb, in maximum amounts of 2.7 wt%.

Occurrence  Orpiment

and realgar occur typically together in the same environments. Moreover orpiment is formed as an alteration product of realgar.

Economic relevance  Up to the 19th century, orpiment

was widely used as the so-called royal-yellow pigment. Nowadays, it is applied in the production of IR-permeable glasses, of photoconductors and semiconductors. 5.5  Complex Metal Sulfides (Sulfosalts)

General formula:

Ax By Sn with A = Ag, Cu, Pb etc., B = As, Sb, Bi. The complex metal sulfides, the so-called sulfosalts, form a relatively large group of ore minerals, but only few of them are economically important. In contrast to the sulfides so far described, As and Sb form metallic constituents in their respective crystal structures (. Table 5.5) whereas, in arsenides and antimonides, As and Sb take the positions of the S atoms. z Proustite—Pyragyrite Ag3AsS3, Ag3SbS3

. Table 5.5  Important sulfosalts Mineral

Formula

Metal content

Crystal class

. Table 5.4  Arsenic sulfides

Proustite

Ag3AsS3

Ag 65.4%

3m

Mineral

Formula

Metal content

Crystal class

Pyrargyrite

Ag3SbS3

Ag 59.7%

3m

Realgar

As4S4

As 70.1%

2/m

Tennantite

Cu12As4S13

¯ 43m

Orpiment

As4S6

Sb 61.0%

2/m

Tetrahedrite

Cu12Sb4S13

¯ 43m

103

5.5 · Complex Metal Sulfides (Sulfosalts)

. Fig. 5.16  Crystal forms and habits of pyrargyrite; hexagonal prism ¯ ¯ trigonal pyramids {1011} ¯ and {0112}, ¯ ditrigonal pyramids {2131} {1120}, ¯ and {123¯1}

Both minerals, also known as light and dark ruby silver, crystallise in the crystal class 3m and in the same structural type but do not form a solid solution. They show similar ditrigonal-pyramidal crystal forms, display similar physical properties and occur in similar settings. Crystal form and habit  Crystal class 3m; occasionally beautiful ditrigonal-pyramidal crystals, rich in different faces and forms, are found, especially of pyragyrite (. Fig. 5.16). Their habit is generally prismatic with dominant hexagonal prism {112¯ 0}, in other cases apparently scalenohedral or rhombohedral, due to the formation of ditrigonal {213¯1} or trigonal {101¯1} pyramids, respectively (. Figs. 5.16, 5.17). Twins are common. In many cases, however, proustite and pyrargyrite are found as massive aggregates, as vein fillings and as coatings.

A reliable macroscopic distinction between proustite and pyrargyrite is not always possible. Crystal structure  The proustite and pyrargyrite structures

display infinite spiral Ag–As or Ag–Sb chains parallel to c, interconnected by AsS3 and SbS3 pyramids, respectively, all of which point in the same direction, resulting in a polar c-axis.

Occurrence  Both minerals are found, together with other

Physical properties Cleavage

¯ distinct {1011},

Fracture

conchoidal, brittle

Hardness

2–2½

Density

5.57 (proustite), 5.85 (pyrargyrite)

Optical properties

translucent to nearly transparent with adamantine lustre; colour: proustite scarlet-vermilion (. Fig. 5.17), the surface becomes darker on exposition to light; pyrargyrite dark red to greyish-black in reflected, red in transmitted light

Streak

. Fig. 5.17  Crystal of proustite with dominant ditrigonal pyramid ¯ Charanacillo, Chile; Width of view is c. 2 mm {2131},

proustite scarlet-vermilion, pyrargyrite cherry-red

Ag minerals, such as argentite, acanthite, Ag-rich tetrahedrite (freibergite) and native Ag, in low-temperature hydrothermal veins (7 Sect. 23.4.2).

Economic relevance  Pyrargyrite is an important and relatively frequent Ag-ore mineral, it is more common than the associated proustite.

z Tennantite—Tetrahedrite Cu12As4S13, Cu12Sb4S13 Crystal form and habit  Crystal class 4¯ 3m; in most cases, the cubic crystals show tetrahedral habit with dominant tetrahedra {111} or {11¯1} or tristetrahedra {211} or {21¯1}, or

5

104

Chapter 5 · Sulfides, Arsenides and Complex Sulfides (Sulfosalts)

may substitute for As and Sb. Fe is always present and can reach as much as 13 wt%, Zn up to 8 wt%. In many cases, Ag content ranges between 2 and 4 wt% and can reach as much as 18% in freibergite, whereas schwazite can contain up to 17 wt% Hg. Occurrence  Tetrahedrite is the most common sulfosalt

5

. Fig. 5.18  Crystal forms and habits of tetrahedrite and tennantite; a combination of tetrahedron {111} and tristetrahedron {211}; b tristetrahe¯ with dron {211}; c combination of two different tetrahedra {111} and {111} cube {100}, visible as thin ridges only, of which face (001) is indicated

combinations of these (. Fig. 5.18). Penetration twins are common. In many cases, members of the tennantite-tetrahedrite series occur as granular aggregates or veins.

mineral, whereas tennantite is less common. They predominantly occur, together with other Cu, Ag, Pb and Zn sulfide minerals, in low- to medium-temperature hydrothermal veins (e.g., 7 Sect. 23.4.4).

Economic relevance  Tetrahedrite

and tennantite are important ore minerals for Ag and Cu, locally also for Hg (schwazite).

References

Physical properties Cleavage

missing

Fracture

conchoidal, brittle

Hardness

3–4½ (tennantite harder)

Density

4.6–5.1 (tetrahedrite denser)

Optical properties

sallow metallic lustre; in thin splinters not completely opaque; steel-grey, with greenish to bluish tint, tetrahedrite is generally darker than tennantite

Streak

greyish-black on tetrahedrite, reddish-grey to reddish-brown on tennantite

Crystal structure  The structure of both minerals can be derived from the sphalerite structure (. Fig. 5.6), in which

one quarter of the metal atoms are replaced by As or Sb. These are linked to 3 S atoms to form trigonal Cu3As or Cu3Sb pyramids. This leads to omission of 3 S atoms per unit cell. Chemical composition  Tetrahedrite and tennantite form

a complete solid solution series, in which part of the Cu atoms are replaced by Fe, Zn, Ag or Hg. Moreover, some Bi

Fleet ME (2006) Phase equilibria at high temperatures. In: Vaughan DJ (ed) Sulfide mineralogy and geochemistry. Rev Mineral Geochem 61:365–419 Kashida S, Watanabe N, Hasegawa H et al (2003) Electronic structure of Ag2S, band calculation and photoelectron spectroscopy. Solid-State Ionics 158:167–175 Kojima S, Sugaki A (1984) Phase relations in the central portion of the Cu–Fe–Zn–S system between 800° and 500 °C. Mineral J Japan 12:15–28 Kojima S, Sugaki A (1985) Phase relations in the Cu–Fe–Zn–S system between 500° and 300 °C und hydrothermal conditions. Econ Geol 80:158–171 Further Reading Anthony JW, Bideaux RA et al (1990) Handbook of Mineralogy, vol I: Elements, sulfides, sulfosalts. Mineral Data Publ, Tucson, Arizona Bowles JFW, Howie RA, Vaughan DJ, Zussman J (2011) Rock-forming Minerals, vol 5A: Oxides, hydroxides and sulphides, 2nd edn. Geological Society, London Kiseeva S, Edmonds M (eds) (2017) Sulfides. Elements 13:81–122 Ramdohr P (1976) The Ore Minerals and Their Intergrowths. Pergamon, Oxford Strunz H, Nickel EH (2001) Strunz Mineralogical Tables, 9th ed. Schweizerbart, Stuttgart Vaughan DJ (ed) (2006) Sulfide mineralogy. Rev Mineral Geochem 61 Vaughan DJ, Corkhill, CL (2017) Mineralogy of sulfides. Elements 13:81–87

105

Halides References – 109

© Springer-Verlag GmbH Germany, part of Springer Nature 2020 M. Okrusch, H. E. Frimmel, Mineralogy, Springer Textbooks in Earth Sciences, Geography and Environment, https://doi.org/10.1007/978-3-662-57316-7_6

6

106

Chapter 6 · Halides

Introduction The minerals of this class (. Table 6.1) consist of the large, electronegative halogen ions Cl−, F−, Br− and J−, coordinated with low-valence cations, such as Na+, K+, Ca2+ and Mg2+, some of which are relatively large as well. The bond character is predominantly ionic. In general, the halides are colourless or allochromatic, i.e., their colours are produced by foreign ions or mineral inclusions. Halides typically have low densities, low refractive indices and a weak lustre. Some of them are soluble in water.

Physical properties Cleavage

{100}, perfect; translation on {110}

Fracture

conchoidal, brittle

Hardness



Density

2.16–2.17

Optical properties

transparent; vitreous lustre; colourless; red or yellow colours are due to inclusions of haematite or limonite, whereas inclusions of clay lead to grey, of bitumen to black, colours. Blue colouration is due to various lattice defects, so called colour-centres. These are caused by irradiation, the radiation source being presumably the radioactive isotope 40K, derived from nearby K minerals

Streak

white

Additional characteristics

easily soluble in water, salty taste, flame colouration deep yellow

z Halite

6

NaCl Crystal form and habit  Crystal class 4/m3¯2/m; well-developed crystals typically in cube form {100} (. Fig. 6.1). Commonly

in massive, granular or sparry, in places fibrous aggregates, known as rock salt. In some cases, clusters of clay form pseudomorphs after halite.

Chemical composition  Halite in evaporite rocks, formed at low temperatures, is virtually pure NaCl with KCl contents of distinctly less than 1 mol% (Chang et al. 1996). However, at higher temperatures, the contents of KCl in halite and of NaCl in coexisting sylvite increase, and at >500 °C, a complete solid solution exists between NaCl and KCl (Waldbaum 1969). Cl in halite can be replaced by up to about 200 ppm Br (Chang et al. 1996). Crystal structure  Similar to galena (. Fig. 5.3), the Na[6]Cl[6] structure (. Fig. 6.2) consists of two cubic, face-centred

lattices of Na+ and Cl−, respectively, that are placed into each other after shifting for a half lattice constant ½a, i.e., along the edge of the cube. Consequently, each Na+ cation is bound to 6 Cl− anions in octahedral coordination and vice versa, i.e., both Na+ and Cl− are [6]-coordinated. Halite is a good example for pure heteropolar (ionic) bonding. However, the electrostatic Coulomb attraction between the . Table 6.1  The most important halides

. Fig. 6.1  Crystal group of halite. Due to skeletal crystal growth, the cube faces {100} are not perfectly developed but indicated by crystal edges; Koehn Dry Lake, California, USA; Width of view c. 5 cm

Mineral

Formula

Crystal class

Halite

NaCl

¯ 4/m32/m

Sylvite

KCl

¯ 4/m32/m

Fluorite

CaF2

¯ 4/m32/m

Carnallite

KMgCl3•6H2O

2/m2/m2/m

107 Introduction

Crystal structure  Isostructural with halite, pure ionic bonding between K+ and Cl−. Chemical composition  In evaporite deposits sylvite con-

tains only small amounts of Na and traces of Rb and Cs, whereas Cl− can be replaced by up to 0.5% Br−.

. Fig. 6.2  Crystal structure of halite, NaCl

Occurrence  Sylvite is an important constituent in evaporites, such as sylvinite, a mixture of halite and sylvite. The mineral is also formed as a volcanic sublimation product. Economic relevance  As constituent of potash salts, syl-

large monovalent ions is relatively weak leading to the low hardness of halite. The perfect cleavage follows the densely occupied {100} faces. Occurrence  Halite is a predominant or exclusive con-

stituent of evaporites, such as rock salt, intercalated with potash salts and anhydrite or gypsum rocks, to form saline deposits (7 Sect. 25.7.2). Moreover halite occurs as efflorescent in deserts and dry steppes, at the margins of salt lakes or salt pans (7 Sect. 25.7.1), and as sublimation product of active volcanoes.

Economic relevance  Rock salt is an important raw material in the chemical industry, used for production of metallic Na, of soda and caustic soda, of chlorine gas or of hydrochloric acid. Moreover, rock salt finds wide application, e.g., in animal feed, for water softening, as preserving agent, and as de-icing salt (road salt). NaCl used as table salt is produced in salines, either by leaching from rock salt or by evaporation of seawater. In 2019, the world production of rock salt is estimated at 293 Mt, the most important producers being PR China (60 Mt), the USA (42 Mt), India (30 Mt), Germany (14 Mt), Australia (13 Mt), Canada (12 Mt), and Mexico (9 Mt) (U.S. Geological Survey 2020).

z Sylvite

KCl Crystal form and habit  Crystal class 4/m3¯2/m; crystal faces {100}, commonly in combination with {111}. As constituent in potash salts, sylvite forms massive, granular to sparry aggregates. Physical properties. Similar to halite. Cleavage

{100} perfect

Hardness

2

Density

1.99, a little lower than for halite

Optical properties

similar to halite

Distinguishing features to halite

Bitter-salty, reddish violet flame colouration

vite is a starting product for high-class fertilisers. Moreover, it serves as a raw material in the chemical industry, especially for production of potassium compounds, and in glass making.

z Fluorite

CaF2 Crystal form and habit  Crystal class 4/m3¯2/m; well-de-

veloped crystals are common, predominantly cubes {100} (. Figs. 6.3, 6.4), in places combined with octahedron {111}, dodecahedron {110}, tetrahexahedron {hk0} or hexoctahedron {hkl} or forming penetration twins on {111} (. Fig. 6.3). More rarely, dominant forms are {111} or {110}. In some cases, fluorite crystals show zoning. Euhedral crystals of fluorite reaching as much as 20 cm in size have been found in the mining area of Dalnegorsk, Primorsky Krai, Eastern Siberia. More commonly, however, fluorite forms massive, sparry or fine-grained aggregates that may show chromatic layering. Physical properties Cleavage

{111}, perfect

Fracture

subconchoidal to uneven, brittle

Hardness

4, reference mineral of Mohs’ hardness scale

Density

3.18

Optical properties

translucent to transparent; vitreous lustre; fluorite displays a great diversity of colours, especially green (. Fig. 2.1), violet, yellow, colourless (. Fig. 6.4). These generally rather pale colours are caused by trace elements or by lattice defects in the crystal structure. The deep blue to blackish violet colour observed in some fluorites is due to radioactive radiation emitted by inclusions of uranium minerals. Thereby, part of the Ca2+ ions is reduced to metallic Ca of µm to nm size serving as a pigment in colloidal distribution. The F− ions are oxidised to F2 gas which, on smashing or grinding fluorite, is released and emits a strong smell, first recorded in the Wölsendorf deposit, Bavaria. Therefore, this variety is kown as “fetid fluorite” or “antozonite”, the “Stinkspat” of the Bavarian miners. When radiated by UV light, fluorite displays variably intense fluorescence, caused by traces of rare-earth elements (REE) substituting for Ca2+

6

108

Chapter 6 · Halides

. Fig. 6.5  Crystal structure of fluorite, described in the text . Fig. 6.3  Two fluorite cubes {100}, forming a penetration twin on {111}

Chemical composition  In yttrofluorite and cerfluorite,

6

Ca2+ is partly replaced by Y3+ or Ce3+, respectively, while the charge balance is attained by additional F− placed on free lattice positions, a fact known as interstitial defect. Occurrence  Fluorite commonly occurs in hydrothermal veins (7 Sect. 23.4.9) and impregnations. In addition,

stratabound fluorite deposits intercalated within sedimentary rocks gain more and more economic significance (7 Sect. 23.6.2). The largest fluorite deposit worldwide, Bayan Obo, Inner Mongolia, northern China, is related to carbonatites. It contains estimated reserves of 130 Mt of CaF2. In 2019, the most important fl ­ uorite-producing countries were PR China with an annual production of 4.0 Mt, Mexico (1.2 Mt), Mongolia (0.67 Mt), Vietnam (0.24 Mt), South Africa (0.24 Mt), and Spain (0.14 Mt) (U.S. Geological Survey 2020).

. Fig. 6.4  Fluorite with cubic habit {100}; Clara mine, Oberwolfach, Black Forest, Germany; Width of view is c. 2 cm

Crystal structure  As shown in . Fig. 6.5, the Ca2+ ions form

a face-centred lattice, in which eight F− ions are placed in the centres of the partial cubes, and thus form a primitive cube with half edge length. This implies that each Ca2+ is surrounded by 8 F− in cubic, and each F− by 4 Ca2+ in tetrahedral coordination. The perfect cleavage after {111} is parallel to the lattice planes that exclusively contain cations or anions, respectively.

Economic relevance  Fluorite, also named fluorspar in industry, is an important raw material with extremely versatile application. Metallurgical grade fluorite (60–85% CaF2) is mainly used in metallurgy as flux melting agent, whereas ceramic grade fluorite (85–95% CaF2) serves for manufacturing glass, enamel and ceramics. The most important application, however, is in the chemical industry, where acid grade fluorite (≥97% CaF2) is used for production of hydrofluoric acid and of various fluor compounds. For electrolytic production of aluminium metal, an artificial kryolite liquid, made from fluorite, is added to the bauxite ore, as the only important deposit of natural kryolite, Na3AlF6, at Ivigtut (Kitaa), Greenland, is totally exhausted. Moreover, colourless, pure fluorite is cut to form lenses for apochromates, i.e., objectives of extremely high resolution, used in light and UV optics. For this application, large synthetic fluorite crystals are grown from a melt in a vacuum furnace with a temperature gradient. Crushed natural fluorite is used as starting material, impurities being scavenged by addition of PbF2. The annual production of synthetic fluorite is at about 200 t. Since the 19th century, purple-blue fluorite, the famous “Blue John”, has been mined at Castleton, Derbyshire, England, for use as semi-precious stone and for ornamental carving.

109 References

Occurrence  Alone or together with halite, kieserite,

z Carnallite

KMgCl3·6H2O Crystal form and habit  Crystal class 2/m2/m2/m; generally in massive, granular aggregates. Physical properties Cleavage

none, conchoidal fracture

Hardness

1–2

Density

1.6

Optical properties

translucent to transparent; greasy lustre; usually reddish due to inclusions of tiny haematite scales that produce a characteristic metallic shimmer, more rarely yellow or milky-white

Additional characteristics

somewhat bitter taste, hygroscopic, easily soluble in water and deliquescent

structure Three-dimensional framework of plane- and edge-sharing KCl6 octahedra. The large cavities between them are filled with Mg(H2O)6 octahedra that are surrounded by 12 Cl− ions. Crystal

Chemical composition  In carnallite, some

replaced by minor amounts of Rb+.

Br−,

and

K+

Cl− can be by traces of

MgSO4·H2O, and/or additional salt minerals, carnallite forms the uppermost layer in evaporite sequences, the so-called carnallite region in evaporite deposits (7 Sect. 25.7.2), e.g., in northern and central Germany. In places, carnallite shows a brecciated structure. Economic relevance  Carnallite is the most important pri-

mary potash-salt mineral that predominantly serves for production of potash fertiliser and, to a lesser extent, of Mg metal.

References Waldbaum DR (1969) Thermodynamic mixing properties of NaCl–KCl liquids. Geochim Cosmochim Acta 33:1415–1427 U.S. Geological Survey (2020) Mineral commodity summaries 2020: U.S. Geological Survey, 200 p., 7 https://doi.org/10.3133/mcs2020 Further Reading Anthony JW, Bideaux RA, Bladh KW, Nichols MC (1997) Handbook of Mineralogy, vol III: halides, hydroxides, oxides. Mineral Data Publ., Tucson Chang LL, Howie RA, Zussman J (1996) Rock-forming Minerals, vol 5B, 2nd edn. Non-silicates: sulphates, carbonates, phosphates, halides. Longmans, Harlow

6

111

Oxides and Hydroxides 7.1 M2O Compounds – 112 7.2 M3O4 Compounds – 112

7.2.1 Spinel Group – 113

7.3 M2O3 Compounds – 116 7.4 MO2 Compounds – 120 7.5 Hydroxides – 124 References – 126

© Springer-Verlag GmbH Germany, part of Springer Nature 2020 M. Okrusch, H. E. Frimmel, Mineralogy, Springer Textbooks in Earth Sciences, Geography and Environment, https://doi.org/10.1007/978-3-662-57316-7_7

7

112

Chapter 7 · Oxides and Hydroxides

Introduction

7

In this mineral class, oxygen forms compounds with one, two or more metals. In contrast to the sulfides, the crystal structures of oxides are marked by predominantly ionic bonds, in part with transition to atomic bonding. According to their metal-oxygen ratio, different divisions of oxides can be distinguished, such as M2O, MO, M3O4, M2O3 and MO2. More complicated oxidic compounds contain two or more metal atoms, such as most minerals of the spinel group, X[4]Y2[6]O4. In common spinel, the tetrahedrally coordinated X position is occupied by Mg2+, the octahedrally coordinated Y position by Al3+. In magnetite, Fe3O4, the X position is taken by Fe3+, while both Fe2+ and Fe3+ occupy the Y position of the spinel structure. The crystal structures of important Mn oxides are characterised by tunnel structures. Anions in the hydroxide division are (OH)−-groups, either alone or in combination with O2−. Important oxide and hydroxide minerals are listed in . Table 7.1.

7.1  M2O Compounds z Cuprite Cu2O

. Table 7.1  Important oxide and hydroxide minerals Mineral

Formula

Metal content*

Crystal class

Cu 88.8%

¯ 4/m32/m

1. M2O compounds Cuprite

Cu2O

2. M3O4 compounds Chrysoberyl

BeAl2O4

2/m2/m2/m

Spinel

MgAl2O4

¯ 4/m32/m

Magnetite

Fe2+Fe23+O4

Fe 72.4%

¯ 4/m32/m

Chromite

FeCr2O4

Cr 46.5%

¯ 4/m32/m

3. M2O3 compounds Corundum

Al2O3

Al 52.9%

¯ 32 m

Haematite

Fe2O3

Fe 69.9%

¯ 32 m

Ilmenite

FeTiO3

Fe 36.8%, Ti 31.6%



Perovskite

CaTiO3

2/m2/m2/m

4. MO2 compounds**

Crystal form and habit  Crystal class 4/m3¯2/m; larger,

well-developed crystals are fairly common, typically with octahedron {111}, rhomb-dodecahedron {110} and cube {100}, often in combination; more common are massive, granular, or powdery aggregates. Physical properties Cleavage

{111}, distinct

Fracture

uneven, brittle

Hardness

3½–4

Density

6.1

Optical properties

translucent to transparent; metallic-adamantine lustre on crystal faces or fresh fractures; massive pieces are metallic grey to reddish brown, transparent fragments ruby-red

Streak

brownish-red

Rutile

TiO2

Ti 60.0%

4/m2/m2/m

Cassiterite

SnO2

Sn 78.8%

4/m2/m2/m

Pyrolusite

MnO2

Mn 63.2%

4/m2/m2/m

Manganates with tunnel structures Uraninite

UO2 to U3O8

U 88.2–84.8%

¯ 4/m32/m

Gibbsite

γ-Al(OH)3

Al 34.6%

2/m

Diaspore

α-AlOOH

Al 45.0%

2/m2/m2/m

Goethite

α-FeOOH

Fe 62.9%

2/m2/m2/m

Lepido­ crocite

γ-FeOOH

Fe 62.9%

2/m2/m2/m

5. Hydroxides

*In nature, the principal metal contents are generally smaller than the theoretical values given here, due to substitution by other metals **SiO2 minerals are described in the silicate chapter (7 Sect. 11.6.1)

7.2  M3O4 Compounds Occurrence  Oxidation product of Cu-sulfide minerals or

native copper. Red copper ore is a reddish-brown, earthy mixture of cuprite and other Cu minerals with limonite (7 Sect. 7.5).

Economic relevance  Although cuprite occurs in minor

amounts in many places, it is mined as a copper ore only locally.

z Chrysoberyl BeAl2O4 Crystal form and habit  Crystal class 2/m2/m2/m; typically as thick platelets. The predominant pinacoid face {010}, commonly showing a distinct striation, can be combined

113

7.2 · M3O4 Compounds

7.2.1  Spinel Group

. Fig. 7.1  Penetration trills of a chrysoberyl fom Esperito Santo, Brazil, and b alexandrite from Novello Claims, Zimbabwe. As the photograph was taken in artificial light, the colour of alexandrite appears reddish-brown; collection Hermann Bank, Idar-Oberstein, Germany

with the {001} pinacoid and the {101} and {012} prisms as well as the rhombic dipyramid {111}, which, in many cases, dominates. Twins on {103} with ~60° angles are common. Penetration of three such twins leads to complex twins (trills) of seemingly hexagonal habit, dominated by {111} faces, resembling hexagonal dipyramids (. Fig. 7.1). Physical properties Cleavage

{001}, distinct

Fracture

conchoidal

Hardness



Density

3.65–3.8

Optical properties

transparent to translucent; lustre; vitreous pale greenish-yellow to green. Cat’s eye chrysoberyl shows a wavy flare, the chatoyancy. The colour of the Cr-bearing variety alexandrite is emerald-green in daylight, but changes to reddish in artificial light (. Fig. 7.1b). This feature is caused by two different absorption bands in the optical spectrum, where yellow and blue wavelengths are absorbed, green and red are transmitted

A great number of oxide minerals, some sulfides (e.g., the thiospinels linnaeite Co3S4 and greigite Fe3S4) and many synthetic compounds crystallise in the spinel-type structure, which is extremely variable and can take up at least 30 different chemical elements with cation-valences ranging from +1 to +6. In this structure, oxygen—or sulfur in thiospinels—forms a cubic close-packed lattice with 32 oxygens in the unit cell, where 8 of the tetrahedral and 16 of the octahedral interstices are occupied by cations (. Fig. 7.2). In normal spinels with the general formula X[4]Y2[6]O4, the 8 tetrahedral sites are filled with divalent cations, the 16 octahedral sites with trivalent cations, in the following indicated by [ ]. Examples are chromite Fe2+[Cr23+]O4, magnesiochromite Mg2+[Cr23+]O4 and hercynite Fe2+[Al23+]O4. In inverse spinels with the general formula Y[4][X[6]Y[6]]O4, the tetrahedral sites are generally occupied by trivalent, the octahedral sites by di- and trivalent cations, examples being magnetite Fe3+[Fe2+Fe3+]O4, magnesioferrite Fe3+[Mg2+Fe3+]]O4 and jacobsite Fe3+[Mn2+Fe3+]O4, whereas ulvöspinel has the structural formula Fe2+[Ti4+Fe2+]O4. The inverse spinel-structure of magnetite can be derived from its magnetic moment of 4.07 μB, extrapolated to absolute zero = 0 K (7 Sect. 1.4.4). We recall that Fe2+ and Fe3+ have magnetic moments of 4 μB and 5 μB, respectively (. Table 1.3). If magnetite were a normal spinel with the structural formula Fe2+[Fe23+]O4, the tetrahedral sites would be occupied by 8 Fe2+, the magnetic moments of which would add to 8 × 4 μB = 32 μB, whereas the octahedral sites would be occupied by 16 Fe3+ leading to 16 × 5 μB = 80 μB but with opposite spin direction. The difference would result in a total magnetic moment of 48 μB or—based on 4 oxygens—of 6 μB, a value that is far too high. If,

Crystal structure  Similar to the structure of olivine (. Fig. 11.3); the small Be2+ ions form BeO4 tetrahedra, the corners of which are linked with those of edge-sharing Al3+O6 octahedra. Analogous to olivine, the chrysoberyl formula could be written as Al2BeO4. Occurrence  Chrysoberyl occurs in Al-rich pegmatites and skarns (Sects. 23.3.1, 26.6.1). Together with emerald (7 Sect. 11.3), alexandrite crystallises in so-called black-

walls, i.e., in nearly pure biotite schists, formed by metasomatic reactions between ultramafic and felsic rocks (7 Sect. 26.6.1), a long-known example being the famous occurrence at Tokovaya River, Ural Mountains, Russia. Chrysoberyl can be locally concentrated in placer deposits.

Economic relevance  Cat’s eye chrysoberyl and alexandrite

are highly priced gemstones.

. Fig. 7.2  Spinel structure made up of cubic close-packed lattice of oxygens, the tetrahedral and octahedral interstices of which are partly filled by cations; after Lindsley (1976)

7

114

7

Chapter 7 · Oxides and Hydroxides

however, magnetite is considered an inverse spinel Fe3+[Fe2+Fe3+]O4, one obtains 8 × 5 μB = 40 μB for the tetrahedral, and 8 × 4 μB + 8 × 5 μB = 72 μB for the octahedral sites. The difference of 32 μB or—based on 4 oxygens—4 μB is very close to the magnetic moment of 4.07 μB, experimentally determined (Lindsley 1976).

Economic relevance  Clear, transparent spinel with ruby-

Presumably, many spinels display a cation distribution transitional between the normal and the inverse structure type. For instance, the structure of common spinel MgAl2O4 conforms at about 87% to the inverse type. Therefore, simplified mineral formulae are given in the following. In natural spinels, the divalent cations Mg2+, Fe2+, Zn2+ or Mn2+ may completely substitute for each other, whereas the mutual substitution between trivalent cations Al3+, Fe3+, Mn3+ or Cr3+ is limited. Due to the manifold alternative substitutions, minerals of the spinel group show considerable differences in their physical properties. In order to achieve or optimise their specific technical properties, the chemical composition of synthetic crystals with spinel structure is purposefully variegated, an approach known as material design. Based on their chemical composition, one can distinguish the following spinel end-members: 5 aluminium spinels, e.g., common spinel, MgAl2O4, hercynite, Fe2+Al2O4, gahnite ZnAl2O4, galaxite MnAl2O4 5 iron spinels, e.g., magnetite Fe2+Fe23+O4, magnesioferrite MgFe23+O4, jacobsite MnFe23+O4, ulvöspinel Fe22+TiO4, franklinite ZnFe23+O4 5 chromium spinels, e.g., chromite Fe2+Cr2O4, magnesiochromite MgCr2O4

z Additional Aluminium Spinels

The densities (D) given in the following refer to the respective pure end-members.

red colour is a valuable gemstone. Single crystals of spinel of any colour are industrially grown using the flame-fusion method of Verneuil (7 Sect. 2.4.3).

Hercynite, Fe2+Al2O4: H 7½–8, D 4.40, colour black, streak dark greyish-green; occurrence: together with titanian magnetite in orthomagmatic ore deposits or with corundum in Al-rich metamorphic rocks, such as emery deposits. Pleonaste, (Mg,Fe2+)(Al,Fe3+)2O4, is a black ferroan spinel, in thin splinters translucent with green colour. Gahnite ZnAl2O4: H 7½–8, D 4.62, colour dark yellowish-brown, dark bluish-green or black, translucent, streak grey; occurrence: in metamorphic rocks and ore deposits, in pegmatites and granites. Galaxite, MnAl2O4: H 7½, D 4.04, black, in splinters and streak reddish-brown; occurrence: in metamorphic rocks and manganese deposits.

z Magnetite Fe2+Fe23+O4 Crystal form and habit  Crystal class 4/m3¯2/m; cubic crystals commonly show the octrahedron {111} (. Fig. 2.9),

more rarely the rhomb-dodecahedron {110}, or combinations of both; twins on the Spinel Law {111}; magnetite is a common accessory constituent in many different rocks. In massive ore, it forms granular aggregates. Martite is a pseudomorph of haematite after magnetite. Physical properties Cleavage

parting on {111} indicated

Fracture

conchoidal, brittle

z Common Spinel MgAl2O4

Hardness

6

Density

5.20

Crystal form and habit  As in all members of the spinel

Optical properties

opaque with metallic to submetallic lustre; black to brownish-black, in part with bluish-grey tarnish; in thin splinters translucent with brown colour

Streak

black

Characteristic property

highly ferromagnetic with a Curie temperature of 578 °C (7 Sect. 1.4.4)

group, the crystal class is 4/m3¯2/m. In most cases, crystals occur as octahedra {111}, more rarely in combination with the rhomb-dodecahedron {110} and/or the trisoctahedron {221}. Twinning on {111}, the Spinel Law, is common. In rocks, spinels form granular aggregates. Physical properties Cleavage

{111}, imperfect

Fracture

conchoidal

Hardness

7½–8

Density

3.55

Optical properties

transparent to translucent; vitreous lustre; common spinel shows many different colours; due to traces of Cr, gem-quality colour varieties are mostly red, rarely blue or green

Occurrence  Predominantly in metamorphic rocks, sec-

ondary enrichment in placer deposits.

Chemical composition  The total Fe content may range up to the theoretical value of 72.4%, but commonly Fe2+ is replaced by some Mg or Mn, and Fe3+ by some Al, Cr, V, Mn3+ or Ti+Fe2+. On cooling of titanian magnetite, lamellae of ulvöspinel are exsolved parallel to {1111} but are generally transformed into ilmenite lamellae by means of the oxidation reaction

3Fe2+ 2 TiO4 +1/2O2 ⇋ 3Fe2+ TiO3 + Fe2+ Fe3+ 2 O4 ulv¨ospinel

ilmenite

magnetite

[7.1]

However, upon very quick cooling, e.g., in volcanic lavas, this exsolution is kinetically inhibited and Reaction [7.1] does not proceed.

115

7.2 · M3O4 Compounds

Occurrence  As a differentiation product of basic magmas,

magnetite, especially titanian magnetite, forms important orthomagmatic ore deposits (7 Sect. 21.2.2); moreover, magnetite is a common accessory constituent in many rock types. It can be concentrated in skarn deposits, formed by metasomatic replacement of carbonate rocks (7 Sect. 23.3.1), or in banded iron formations (BIF), e.g., itabirite, together with, or instead of, haematite (7 Sect. 25.4.2). Crystal chains of magnetite and the thiospinel greigite, Fe2+Fe23+S4, are precipitated in cells of magnetotactic bacteria, thus forming organelles called magnetosomes. These allow the bacteria to align in the Earth’s magnetic field and to move along the magnetic flux lines. Greigite also crystallises in oxygen-free (anaerobic) sedimentary environments and can serve as a precursor of pyrite.

Franklinite, ZnFe23+O4: H 6–6½, D 5.34, colour black, streak reddish brown to black, dull metallic lustre, opaque. Occurrence: In metamorphic Zn–Mn–Fe deposits, such as the type locality at the old mines of Franklin Furnace, New Jersey, USA.

z Chromite FeCr2O4 Crystal form and habit  Crystal class 4/m3¯2/m; in general,

chromite forms massive, granular bodies or is interspersed within ultramafic rocks, whereas small cubic crystals with {111} are very rare. Physical properties Cleavage

absent

Fracture

conchoidal, brittle

important ore minerals of Fe and Ti deposits.

Hardness



Density

FeCr2O4 5.09, MgCr2O4 4.52

Geological relevance  Magnetite in igneous and metamorphic rocks facilitates palaeomagnetic investigations. For instance, the stripe patterns on the ocean floor, documenting repeated inversions of the Earth’s magnetic field, led to the detection of sea-floor spreading, a fundamental evidence of plate tectonics. These patterns are produced by fine-grained ferromagnetic magnetite crystals, dispersed in oceanic basalts.

Optical properties

opaque; greasy metallic to submetallic lustre; black to brownish-black; in thin splinters translucent with brown colour

Streak

dark brown

Economic relevance  Magnetite and titanian magnetite are

When a basaltic lava flow is cooled below the Curie temperature of 578 °C, magnetite is transformed from the paramagnetic to the ferromagnetic state, whereby the four or five unpaired 3d-electrons of the Fe2+ and Fe3+ atoms, respectively, undergo parallel alignment in individual crystal domains (7 Sect. 1.4.4). Under the influence of the Earth’s magnetic field, the magnetic moments of the magnetite grains in the whole lava flow become orientated parallel to each other and to Earth’s magnetic field lines. Due to this thermoremanent magnetisation (TRM), the orientation and inclination of the Earth’s magnetic field can be reconstructed, thus making it possible to determine the geographic latitude at the time of volcanic eruption. It should be noted, however, that this procedure works only if subsequently the lava flow did not undergo renewed heating above the Curie temperature. For instance, in basalt transformed into amphibolite during a metamorphic overprint at temperatures >573 °C (cf. 7 Chap. 26), the orientation of the magnetite crystals reflects the time of metamorphism rather than documenting the age of the preceding, often much older, volcanic event.

In addition, palaeomagnetic investigations can also be performed on sediments or sedimentary rocks that contain magnetotactic bacteria. z z Additional Iron Spinels

Magnesioferrite, MgFe23+O4: H 5½–6½, D 4.52, colour black, lustre metallic, streak dark read. Occurrence: due to fumarolic activity, small octahedra, regularily intergrown with haematite, precipitate on volcanic rocks, e.g., from Vesuvius and Island of Stromboli, Italy. Jacobsite, MnFe23+O4: H 5½–6½, D 4.87, colour black, streak reddish black; occurrence: as an ore mineral in metamorphic manganese deposits, such as the Kalahari Manganese Fields, South Africa, and on the abandoned Mn deposits of Jakobsberg and Långban, Central Sweden. Ulvöspinel, Fe22+TiO4: Occurrence: In most cases, ulvöspinel is formed by exsolution from titanian magnetite and is commonly replaced by ilmenite (+magnetite) by the oxidation Reaction [7.1].

Chemical composition  Variable, with up to 46.5% Cr; complete solid solution with magnesiochromite, MgCr2O4, and the Mg–Cr bearing hercynite, known as picotite, (Fe2+,Mg)(Al,Cr,Fe3+)2O4. Occurrence  Similar to titanian magnetite, chromite crys-

tallises from basic/ultrabasic magmas and forms important orthomagmatic ore deposits of Cr (7 Sect. 21.2.1). Stratiform chromite deposits occur in layered intrusions consisting of norite or gabbro, various ultramafic rocks and anorthosite, where they form chromite-rich layers and segregations, for example in the Bushveld Igneous Complex in South Africa (. Fig. 21.5) Podiform or alpinotype chromite deposits are found dispersed or concentrated in layers, schlieren or cocade-shaped aggregates in ultramafic rocks of ophiolite complexes, i.e., obducted slivers of oceanic lithosphere (7 Sect. 29.2.1). Moreover, chromite can be concentrated, together with platinum-group metals, in placers. Economic relevance  Chromite is the only Cr-ore mineral

that is of economic interest. In 2019, the world production was ca. 44 Mt chromite, the largest producers being South Africa (17 Mt), Turkey (10 Mt), Kazakhstan (6.7 Mt), and (16%), India (4.1 Mt) (U.S. Geological Survey 2020).

Chromium as metallic raw material  Cr, an important steel-re-

fining metal, is generally traded as the semi-finished product ferrochrome, a Cr–Fe alloy, usually with 50–70% Cr, derived from chromite by electric melting. About 85% of the annual Cr production is used for chromium steel, which is characterised by high hardness and resistance to corrosion, whereas stainless steel commonly contains 13–25% Cr plus some Ni. As an anti-corrosive, iron or steel surfaces are covered with Cr metal

7

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Chapter 7 · Oxides and Hydroxides

Cr metal was already used during the Chinese Qin dynasty (221– 206 BC), to coat metal weapons. In the western world, however, Cr was only detected in 1761, when the mineral crocoite, PbCrO4, was used as a red pigment. Metallic Cr was first produced in 1797 by Louis Nicolas Vauquelin (1763–1829).

7.3  M2O3 Compounds In the M2O3 structures, oxygens form a hexagonal (nearly) close-packed lattice, where the cations, such as Al3+, Fe3+ or Ti4+, occupy 2/3 of the octahedral sites.

z Corundum Al2O3

7

. Fig. 7.3  Single crystals of corundum, variety ruby, in a gneiss from Morogoro, Tansania; width of view is c. 8 cm (Photograph: Rainer Altherr, Heidelberg)

Crystal form and habit  Crystal class 3¯2/m; crystals of corundum show prismatic, tabular or rhombohedral habit. Commonly, different hexagonal dipyramids, such as {112¯ 1}, ¯ .3}, taper {224¯ 1}, {224¯ 3}, and the very steep form {14.14.28 to form arched barrel shapes (. Figs. 7.3, 7.4). Most corundum crystals display a prominent horizontal striation, reflecting alternating growth stages or due to lamellar twinning on {101¯ 1} and {0001}. Many large crystals show uneven, rough faces. More commonly, however, corundum occurs in massive, granular aggregates, rock forming in emeries, or as accessory constituent in some SiO2-deficient rocks.

Physical properties

¯ . Fig. 7.4  Corundum, combination of hexagonal dipyramids {1121}, ¯ {2243}, ¯ and {991¯ 82} ¯ and basal ¯ with rhombohedron {1011} {2241}, pinacoid {0001}

by electroplating. Steel for high-speed tools contains 3–5% Cr, whereas the Cr–Ni bearing superalloy inconel, utilised for construction of jet engines and gas turbines, contains 18.6% Cr. High-temperature-resistant chromite-magnesite bricks are used for building blast furnaces. Different Cr compounds serve as pigments in paints and laquers, for leather tanning, preservation of wood, and as catalysts.

Cleavage

¯ none, parting on {0001} and {1011}

Fracture

conchoidal

Hardness

9, extremely hard, reference mineral of Mohs’ hardness scale

Density

3.98–4.02

Optical properties

adamantine to vitreous lustre; common corundum is colourless to yellowish or bluish grey and transparent on thin edges, whereas gem-quality corundum crystals of different colour are translucent to transparent. The red colour of ruby is caused by Cr3+ or V3+, that of blue sapphire by Fe2+ + Ti4+ or Fe2+ + Fe3+, whereas the orange colour of padparadscha is due to Cr3+ and lattice defects. Additional colour varieties of corundum are also called sapphire, e.g., green to yellowish green ones coloured by Fe3+, pink ones by Ti3+ and yellow ones by Fe3+ and lattice defects (Schmetzer and Bank 1981). A colourless variety is leuco sapphire. Star sapphire and star ruby contain oriented inclusions of rutile needles, occasionally also of haematite platelets, which cause asterism, a star-shaped flare pattern, best viewed in stones cut en cabochon, perpendicular to the optical axis

117

7.3 · M2O3 Compounds

and sapphire are situated, e.g., in Thailand and Sri Lanka. Mining of large ruby deposits under the receding ice shelf of Greenland has started in 2007. Economic relevance  Due to its hardness, corundum is

. Fig. 7.5  Crystal structures of corundum, haematite and ilmenite: a the Al–O3–Al or Fe–O3–Fe building blocks, respectively. b The rhombohedral cell of corundum or haematite, respectively. The centres of the building blocks are placed at the corners and in the centre of the rhombohedron. For clarity, the oygen triplets are omitted. c The crystal structure of ilmenite conforms to the corundum or haematite structures but is composed of Fe–O3–Ti units with alternating Fe and Ti positions. The oxygen triplets are omitted (after Lindsley 1976)

Crystal structure  The oxygens form a hexagonal (nearly) close-packed lattice, where Al3+ occupies 2/3 of the octahedral sites. . Figure 7.5a, b shows the Al–O3–Al building blocks that occupy all corners and the centre of the rhombohedral unit cell. Chemical composition  Although, in general, colour-

ing cations substituting for Al3+ are present in only small amounts (950 °C, haematite and ilmenite form a complete solid solution. Upon slow cooling, ilmenite lamellae are exsolved, parallel to (0001), from haematite and vice versa. No solid solution exists between haematite and corundum.

Occurrence  Together with magnetite, siderite and/or . Fig. 7.7  Crystal group of haematite with tabular habit, a so-called iron rose, Ouro Preto, Brazil; Besides the dominant basal pinacoid ¯ are developed; {0001}, slim faces of the steep rhombohedron {1011} width of view is c. 5 cm

by deformation. Twinning on {0001} is less common. The various habits of haematite crystals reflect differences in formation conditions. For instance, crystals of flaky habit are formed at low temperatures. Haematite, crystallised from a gel at even lower temperatures shows reniform or botryoidal shape and spherolitic structure, known as kidney iron ore. Frequently, haematite forms massive, granular, foliaceous, scaly, dense or earthy aggregates.

Physical properties Cleavage

haematite crystals, especially those with flaky or foliaceous habit, called micaceous haematite, shows parting or delamination on {0001} or the ¯ twin plane {1011}

Fracture

conchoidal, brittle

Hardness

5–6

Density

5.254 (or less)

Fe-silicates, haematite forms a main constituent of banded iron formation (BIF), such as itabirite, jaspilite or taconite, and of Phanerozoic iron stones (7 Sect. 25.4.2). Haematite is a widespread accessory mineral in metamorphic, less commonly in igneous rocks and occurs in hydrothermal veins. Haematite is formed by volcanic exhalation and by metasomatic reaction with limestones, e.g., in the famous occurrence on the island of Elba, Italy. By secondary alteration, magnetite is oxidised to form pseudomorphs of lamellar haematite, known as martite. During weathering, haematite eventually transforms into limonite.

Economic relevance  Haematite is an important iron ore

for production of steel and cast iron. Moreover, it is utilised as pigment and jeweller’s rouge. Since prehistoric times, red ochre has been used by several indigenous peoples for body painting as well as for ornamental painting and artistic drawing.

z Ilmenite FeTiO3 Crystal form and habit  Crystal class 3¯; ilmenite can show trigonal-rhombohedral crystals with variable, rhombohedral or tabular habit. Like in haematite, lamellar twins on {101¯1} can be developed, but no scalenohedra or hexagonal

119

7.3 · M2O3 Compounds

dipyramides. In Ti-ore deposits, ilmenite forms massive, granular aggregates. Isolated ilmenite grains are distributed as accessory constituents in many hard rock types but ilmenite can be also concentrated, in places even to ore grade, in placer deposits.

Physical properties Cleavage

absent but, like in haematite, parting along ¯ twins lamellar {1011}

Fracture

conchoidal, brittle

Hardness

5–6

Density

4.70–4.79, increases with higher Fe2O3 contents in ilmenite that formed at higher temperatures

Optical properties

opaque; lustre metallic on fresh fracture faces, otherwise dull, submetallic; brownish black to steel-grey; in thin splinters translucent with brown colour

Streak

black to dark brown

The crystal structure of ilmenite is similar to the structure of corundum and haematite, where 2/3 of the octahedral sites are taken by Fe2+ and Ti4+, alternating in successive layers (. Fig. 7.5c). Therefore, the symmetry is lower than in corundum and haematite. Chemical composition  Although extensive solid solution

exists with geikelite, MgTiO3, (up to 70 mol%) and pyrophanite, MnTiO3, (up to 64 mol%), the Mg and Mn contents are much smaller in most natural ilmenites. Small amounts of Cr3+ and V3+ can substitute for Fe3+. At temperatures >950 °C, ilmenite and haematite form a complete solid solution series. Upon slow cooling, haematite exsolves from ilmenite forming lamellae parallel to (0001). Occurrence  Ilmenite is a common accessory mineral in

many igneous and metamorphic rocks and can undergo secondary enrichment in ilmenite sands in coastal areas. Profitable concentrations of ilmenite are formed by magmatic differentiation, especially in anorthosite massifs, such as at Tellnes, southern Norway, and Allard Lake, Quebec, Canada (7 Sect. 21.2.2).

Economic relevance  Ilmenite is an important Ti-ore mineral. In 2019, ca. 7.6 Mt ilmenite concentrates were produced worldwide (including a small proportion from rutile), the most important producers being China (2.1 Mt), South Africa (0.82 Mt), Canada (0.69 Mt), Australia (0.66 Mt). and Mozambique (0.59 Mt) (U.S. Geological Survey 2020). Titanium as metallic raw material  Due to their high strength/density ratio, excellent corrosion resistance and favourable mechanical properties, Ti metal and Ti alloys with Fe, Al, V, Mo, Ni and Zr are used in a wide range of applications, especially in aerospace, e.g., for construction of jet engines, missiles and spacecrafts, in marine

technology and for manufacturing of high-quality commodities, e.g., sports articles. As Ti metal is extremely biocompatible, it is the first choice for production of surgical instruments, orthopaedic and dental implants. The Fe–Ti alloy ferrotitanium is used in steel-making, e.g., for production of stainless steel. However, 95% of Ti ore mined worldwide is industrially transformed to titanium dioxide TiO2, in most cases as the rutile polymorph, less frequently as anatase, widely used as titanium white pigment (see below). z Perovskite CaTiO3 Crystal form and habit  Perovskite, crystal class 2/m2/

m2/m, forms crystals of cubical or octahedral habit, hardly deviating from the cubic symmetry, but also occur as branching, skeletal crystals. Lamellar penetration twins on {110} and {112} are common. Physical properties Cleavage

{100}, {010}, {001} rather distinct

Fracture

conchoidal

Hardness



Density

3.95–4.84

Optical properties

non-transparent to translucent; adamantine to metallic lustre, occasionally dull; colour black, reddish brown, orange or honey yellow

Streak

greyish white to colourless

Crystal structure  Perovskite represents a highly impor-

tant structural type, not only found in nature but also in a wide range of synthetic compounds applied in technology. In the perovskite structure, attaining a very dense packing, Ti4+ is bound to 6 oxygens in octahedral coordination while the sites between the corner-sharing TiO6 octahedra accommodate the large Ca2+ cations, which are thus coordinated with 12 oxygens. In contrast to the ideal cubic structure shown in . Fig. 7.8, the TiO6 octahedra are slightly tilted leading to orthorhombic symmetry, whereas synthetic compounds can be cubic, tetragonal or orthorhombic, depending on the exact orientation of the octahedra. In these synthetic structures, more than 20

. Fig. 7.8  Idealised structure of perovskite formed by TiO6 octahedra and Ca filling the big gaps (after Náray-Szabo 1943, from Deer et al. 1962)

7

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Chapter 7 · Oxides and Hydroxides

elements, such as Ca2+, Ba2+, Pb2+, K+, can take the position of the large cations X, whereas the small cations Y are represented by nearly 50 elements, such as Ti4+, Zr4+, Sn4+, Nb5+, Ga3+. Chemical composition  In natural perovskite, Ti can be

replaced by some Nb, and Ca by considerable amounts of REE or alkali elements. Occurrence  Perovskite is an accessory mineral in alkaline

igneous rocks as well as in carbonatites, kimberlites and pyroxenites. It also occurs in impure contact metamorphic marbles. Locally, perovskite is concentrated to form profitable Ti deposits, e.g., in Bagagem, Brazil.

7

Geological relevance  In rocks of the Earth’s crust and upper mantle or as Ti ore, perovskite, CaTiO3, attains only local significance. In contrast, more than 70% of the Earth’s lower mantle consists of the mineral (Mg,Fe) SiO3 with perovskite structure, while CaSiO3 perovskite accounts for 7%. Thus silicate perovskites are the most abundant minerals of the whole Earth (7 Sect. 29.3.4). Economic relevance  Perovskites with piezoelectric properties are industrially synthesised in a wide range of different compositions, forming the basis for electrical ceramics. Depending on their chemical compositions, they have been used as non-conductors (insolators), semiconductors, metallic conductors or high-temperature superconductors. The artificial perovskite SrTiO3 is sold as a relatively expensive diamond simulant under the name fabulite.

7.4  MO2 Compounds Very important MO2 compounds are the SiO2 minerals, above all quartz. However, judging from their crystal structure, these rather belong to the framework silicates and therefore are described in 7 Sect. 11.6.1.

Physical properties Cleavage

{110}, perfect

Fracture

conchoidal, brittle

Hardness

6–6½

Density

4.23 (pure TiO2)

Optical properties

subtranslucent to transparent; adamantine to submetallic lustre; dark red, reddish brown to yellowish, rarely black

Streak

pale brown

Index of refraction

nε = 2.6, nω   = 2.9, i.e., higher than that of diamond!

Crystal structure  In the rutile structure, Ti is surrounded, at nearly equal distances, by 6 oxygens in octahedral coordination. The TiO6 octahedra are connected, along diagonal edges, and form infinite chains parallel to c (. Fig. 7.9). Neighbouring chains are rotated by 90° and displaced by c/2, they are mutually linked by corner-sharing octahedra. TiO2 polymorphs  Among the three TiO2 polymorphs,

the high-temperature form rutile is widespread in nature, whereas tetragonal anatase (4/m2/m2/m; H 5½–6; D 3.82– 3.97) and orthorhombic brookite (2/m2/m2/m; H 5½–6; D 4.08–4.18) are less frequent.

Chemical composition  Rutile can contain considerable amounts of Fe2+, Fe3+, Nb5+ and Ta5+, leading to an increase in density to as much as 5.5. The substitution of Ti4+ by Nb5+ or Ta5+ is enabled by similar ionic radii, and compensated for either by complementary addition of Fe2+ or by vacancies in the structure. Moreover, smaller amounts of Sn4+, Cr3+, V3+ and Al3+ may be present. Lunar rutile from the Apollo 12 site contains La3+ and Ce3+ in addition to Nb and Cr (Deer et al. 2013). Occurrence  In numerous rocks rutile, mostly of micro-

scopic size, is present as an accessory mineral. It forms tiny needles in slates and phyllites, but may form larger crystals in rocks of higher metamorphic grade, such as eclogites, or in granite pegmatites. Sands and sandstones often contain

z Rutile TiO2 Crystal form and habit  Crystal class 4/m2/m2/m; com-

monly, the crystals show tetragonal prisms {110} with vertical striation and are terminated by dipyramids {111}. Short primatic, columnar or acicular habits are common, as well as hair-like needles included in quartz crystals. Elbow or cyclic twins on {011} are typical. In cases, rutile occurs as anhedral grains or forms compact aggregates.

. Fig. 7.9  Structure of rutile, consisting of chains of TiO6 octahedra (after Lindsley 1976)

121

7.4 · MO2 Compounds

rutile grains as a heavy mineral that can be concentrated to ore grade. Economic relevance  Rutile is an important Ti mineral and is occasionally used for recovery of Ti metal, by reducing TiCl4 with Mg metal. The compound TiN is an extremely hard ceramic material, commonly utilised for thin coatings (typically 250 t in weight (Rickwood 1981). Metamorphic rocks consisting entirely or predominantly of calcite are calcitic and variably siliceous marble. Calcite or other carbonate minerals can also crystallise from carbonatic magmas, leading to the formation of carbonatites (7 Sect. 13.2.3). Economic relevance  Limestones, predominantly or exclu-

sively consisting of calcite, are raw materials of paramount economic interest. In the building industry, they are widely used for production of non-hydraulic and hydraulic binders, lime mortar and Portland cement. Since ancient times, attractive limestones and marbles have been utilised as building and dimension stone. Special varieties of economic interest are white, granular marble, e.g., from Carrara, Italy, porous travertine or fine-grained lithographic limestone. Moreover, limestone finds a wide range of practical applications, such as in chemical industries, in production of glass

and cellulose, as flux in metallurgy, for sugar refining, as agricultural lime, as filler material or white pigment in paper. In optical industries, there is still a high demand for clear transparent crystals of Icelandic double spar (. Fig. 1.29). z Magnesite MgCO3 Crystal form and habit  Crystal class 3¯2/m; crystals are

simple with rhombohedral habit {101¯1}. Rock-forming magnesite occurs in sparry or granular aggregates or in dense, microgranular aggregates, often with colloidal structures, so-called gel magnesite. Physical properties Cleavage

¯ perfect, but only developed on coarse{1011}, grained, sparry crystals; gliding on {0001}

Fracture

conchoidal, shown by gel magnesite

Hardness

3½–4½

Density

2.98 (−3.48)

Optical properties

transparent to translucent; vitreous to pearly lustre, dull on gel magnesite; colourless, snow-white, greyish to yellowish white, greyish black

Crystal structure  Isostructural with calcite. Chemical composition  Complete solid solution with siderite, FeCO3, but only limited contents of Mn2+ or Ca2+; magnesite with 5–50 mol% of the FeCO3 component is known as breunnerite.

131

8.1 · Calcite Group, 32/m

Occurrence  Sparry magnesite commonly occurs as masses

and veins in serpentinite or ultramafic rocks in which it is formed during metamorphic overprint, either by metasomatic reaction with CO2-bearing fluids or MgCl2 solutions. Moreover, magnesite can be formed by metasomatic replacement of limestone, either during early diagenesis or under the influence of hydrothermal fluids. Sparry magnesite forms irregular bodies within limestone and dolomite of the so-called Greywacke Zone of the Eastern Alps, Austria. Important deposits occur near Radenthein, Carinthia, as well as at Trieben and Veitsch, Styria, where they are mainly mined in open-cast pits. Gel magnesite is a weathering product of serpentinite.

Economic relevance  Magnesite is an important raw

material especially for the production of super-refractory magnesite bricks, which serve as internal lining of kilns, oxygen converters (LD method), blast-furnaces and incinerators. For this purpose, raw magnesite is sintered at temperatures of 1500 °C to form MgO and subsequently mixed with tar and compressed at low temperatures. Before the first charge of steel is being produced, the tar-magnesite lining of the converter is burned. Magnesite bricks of customisable shape, such as nozzle bricks, are heated in special installations before being deposited into the converter. By caustic treatment of raw magnesite at 800 °C, MgCl2 brine is produced that, mixed with filling materials, is processed to magnesia cement, used for refractory building and isolating materials. Mg-metal is rarely produced from magnesite, but commonly from wastes of processing K–Mg salt, such as carnallite or polyhalite (7 Sect. 25.7.2), or directly from sea-water.

z Siderite FeCO3 Crystal form and habit  Crystal class 3¯2/m; well-developed

crystals mostly form saddle-shaped rhombohedra {101¯1} with bent faces, rarely with lamellar twinning on {011¯2} or simple twinning on {0001}; rock-forming siderite occurs in sparry, granular or earthy aggregates. Siderite with globular, botryoidal or reniform shape, presumably derived from a gel, is known as sphaerosiderite. Physical properties

Crystal structure  Isostructural with calcite. Chemical composition  Theoretical Fe content is 48.2 wt%,

however with variable contents of Mn2+, Mg2+ and Ca2+, more rarely of Zn2+; complete solid solutions with rhodochrosite, MnCO3, and magnesite, MgCO3.

Occurrence  Siderite occurs in hydrothermal veins and is

formed by metasomatic replacement of limestone or marble by reaction with Fe-rich hydrothermal fluids. A prime example is the Erzberg (=Ore Mountain) in Styria, Austria (7 Sect. 23.3.5). Moreover, siderite forms a constituent of Phanerozoic ironstones (7 Sect. 25.4.2). Upon weathering, siderite is transformed into limonite.

Economic relevance  Owing to its Mn content and easy

smelting, siderite is a valuable iron ore, known as “spathic iron ore”.

z Rhodochrosite MnCO3 Crystal form and habit  Crystal class 3¯2/m; well-developed,

often saddle-shaped, crystals with rhombohedral habit {101¯1} are generally small. They may also occur in druses (. Fig. 8.5). Commonly, rhodochrosite occurs in granular or sparry aggregates, or as layered incrustations with botryoidal or reniform surface and radial structure. Moreover, it may form larger, dowdy masses with cellular, crusty or earthy appearance. Physical properties Cleavage

¯ perfect {1011},

Fracture

brittle

Hardness

3½–4

Density

3.70 for pure MnCO3 but decreasing with increasing Mg2+ content

Optical properties

translucent to transparent; vitreous lustre; pink, pale-red, raspberry-red (. Fig. 8.5)

Streak

white

Crystal structure  Isostructural with calcite.

Cleavage

¯ perfect {1011},

Hardness

4–4½

erite, calcite and smithsonite, very limited with magnesite.

Density

3.96 for pure FeCO3, but decreasing down to 3.5 with increasing Mn2+ or Mg2+ contents

Occurrence  In hydrothermal veins; as weathering product

Optical properties

transparent to translucent; vitreous to pearly lustre; light greyish-yellow; upon progressive oxidation, yellowish to yellowish brown, eventually dark brown with colourful metallic tarnish with progressive oxidation to limonite

Practical application  As polished ornamental stone used

Chemical composition  Complete solid solutions with sid-

in the oxidation zone of ore deposits. for handicrafts.

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Chapter 8 · Carbonates, Nitrates and Borates

Occurrence  In most cases, smithsonite occurs in oxida-

tion zones of hydrothermal sulfide deposits that are hosted by limestones, in which solutions of Zn sulfate, formed by weathering of sphalerite, react with CaCO3, thereby precipitating ZnCO3. Economic relevance  Calamine, a mixture of smithsonite with hemimorphite, Zn4[Si2O7]·(OH)2·H2O, is an important Zn ore.

8.2  Aragonite Group, 2/m2/m2/m

Minerals of the aragonite group crystallise in the orthorhombic, although pseudo-hexagonal aragonite structure. By analogy to the calcite structure, planar, triangular CO3 groups are arranged parallel to (001). Rotated by 60°, they are superimposed to form pairs linked by [9]-coordinated cations (. Fig. 8.6b), as opposed to [6]-coordinated ones in the calcite structure. This difference is due to the larger cationic radii in the aragonite group (. Table 8.1).

8

. Fig. 8.5  Rhodochrosite with typical rhombohedral habit, on quartz; Pasto Bueno, Peru; width of view is 3.8 cm

z Smithsonite ZnCO3 Crystal form and habit  Crystal class 3¯2/m; smithsonite can

occur in small rhombohedral crystals but in most cases is compact and fine-grained, forming incrustations, reniform or conical aggregates, often friable. Physical properties Cleavage

¯ perfect {1011},

Fracture

brittle

Hardness

4–4½

Density

4.30–4.45

Optical properties

translucent to hazy; vitreous lustre; colourless, yellowish, greenish, brownish, pale violet (due to traces of Co2+), bluish (due to traces of Cu2+)

Streak

white

z Aragonite CaCO3 Crystal form and habit  Crystal class 2/m2/m2/m; the orthorhombic-dipyramidal crystals generally show prismatic habit with {110} prism and {010} pinacoid, elongated parallel to c and truncated by {011} prisms and {111} pyramids or, in pointed crystals, with {061} prisms and {9 12 2} pyramids. Acicular, radial aggregates are common. Aragonite shows simple or lamellar twinning on {110}. More common and typical are cyclic twins on {110} producing pseudo-hexagonal elongated crystals with intergrowth sutures, striations or longitudinal grooves parallel to c (. Fig. 8.7). Aragonite occurs also as compact, fine-grained masses or

Crystal structure  Isostructural with calcite. composition Theoretical Zn content is 52.1 wt%, complete solid solution with siderite, FeCO3, and rhodochrosite, MnCO3, whereas the contents of Ca2+ and Mg2+ are generally small; occasionally, Cd2+ and Pb2+ are present as well as traces of Cu2+ and Co2+.

Chemical

. Fig. 8.6  Fragments of the calcite structure, projected onto (0001) (a) and the aragonite structure (b), projected onto (001). CO32− groups are shown as light blue triangles, Ca2+ ions as dark blue spheres. In calcite, Ca is surrounded by 6, in aragonite by 9 oxygens (after Strunz and Nickel 2001)

8.2 · Aragonite Group, 2/m2/m2/m

133

. Fig. 8.7  Aragonite, cyclic penetration twin on (110) with pseudo-hexagonal symmetry

incrustations. Pisolitic limestones composed of aragonite globules with concentric structure are known as pea grit.

Physical properties Cleavage

{010}, indistinct, {110}, poor and rarely observed

Fracture

conchoidal

Hardness

3½–4, somewhat higher than on calcite, owing to the higher number of Ca–O bonds in the aragonite structure

Density

2.94–2.95, higher than in calcite due to the denser packing of the aragonite structure

Optical properties

transparent to translucent; lustre: vitreous on crystal faces, greasy on fractures; colourless to delicately coloured

Identification  Like calcite, aragonite dissolves easily in

cold dilute HCl with strong effervescence. Upon boiling in Co(NO3)2 solution, powders of aragonite and other orthorhombic carbonates as well as of hexagonal vaterite become violet-coloured, whereas powders of calcite or other trigonal carbonates remain virtually unchanged. Crystal structure  Aragonite structure, as described above. Chemical composition  Most aragonite specimens are rela-

tively pure CaCO3. Some Ca2+ can be replaced by Sr2+ and Pb2+, more rarely by Fe2+ or Mg2+.

Stability field and mode of occurrence  Aragonite is less common than calcite but occurs as a rock-forming mineral as well. Due to its denser structure with Ca[9] instead of Ca[6] as in calcite, aragonite is the high-pressure polymorph of CaCO3, its stability field being situated in the high-pressure, low-temperature range of the P-T diagram (. Fig. 8.8). Consequently, aragonite can only form as a stable mineral under conditions of high-pressure metamorphism in subductions zones, where the geothermal gradient is unusually low (Sects. 26.2.5, 28.3.8).

. Fig. 8.8  Pressure-temperature diagram of the CaCO3 system, showing the stability fields of calcite and aragonite; the experimentally determined aragonite-calcite equilibrium curve is shown as solid line, its extrapolation as dashed line, the range of experimental uncertainty is shaded (after Johannes and Puhan 1971)

Although aragonite is not a stable mineral under the conditions of the Earth’s surface, it can crystallise metastably in vesicles of volcanic rocks, as constituent of dripstones and of sinters, precipitated from thermal springs or geysers. Moreover, aragonite can form metastably by precipitation from seawater and/or by biogenic processes. Thus it is present in vast amounts of calcareous mud, e.g., on the Great Bahama Bank and in the Florida Bay, or in oolitic limestones of different age. Calcareous sedimentary rocks consisting of aragonite are intercalated with gypsum- and halite-deposits in many evaporites (cf. Chang et al. 1996). As a single constituent, or together with calcite or Mg-calcite, aragonite forms shells or endoskeletons of calcareous algae, foraminifers, sponges, corals, snails, ammonites and pearl mussels. The attractive nacre (mother-of-pearl) layer of natural pearls is composed of aragonite, whereas the shell itself consists of calcite. In the presence of a solvent or upon longer grinding, aragonite slowly transforms into the stable polymorph calcite, a reaction that proceeds very rapidly at temperatures at approximately 400 °C. The transformation is monotropic, i.e., irreversible. However, aragonite is not transformed into calcite, if the aragonite structure is stabilised by Sr2+ substituting for Ca2+. Another metastable polymorph of CaCO3 is hexagonal vaterite, precipitated from aqueous solutions at low temperature or formed by biogenic processes. Economic relevance  Polished calcareous rocks consisting

of aragonite, such as travertine, are used as ornamental and dimension stone.

8

134

Chapter 8 · Carbonates, Nitrates and Borates

z Strontianite SrCO3 Crystal form and habit  Crystal class 2/m2/m2/m; similar to aragonite, crystals of strontianite commonly show acicular or columnar habit, frequently forming simple, lamellar or cyclic twins on {110}, and occasionally arranged to form tufted aggregates. More commonly strontianite occurs in massive, columnar or granular aggregates. Physical properties

8

Cleavage

{110}, distinct, {021} and {010}, poor

Fracture

conchoidal

Hardness



Density

~3.76, i.e., distinctly higher than that of aragonite

Optical properties

transparent to translucent; lustre: vitreous on crystal faces, greasy on fractures; colourless, white, grey, light green or light brown

Crystal structure  Isostructural with aragonite. Chemical composition  In strontianite Sr2+ is always

replaced by some Ca2+, in cases up to ~25 mol% CaCO3, whereas Ba2+ and Pb2+ are present in trace amounts only.

. Fig. 8.9  Cerussite, cyclic penetration twin on (110) with pseudo-hexagonal symmetry

Physical properties Compared with aragonite and strontianite, cerussite displays a higher density, a vivid adamantine lustre and distinctly higher refractive indices nγ = 2.079, nα = 1.804. Cleavage

{110}, distinct, {021}, fair

Fracture

conchoidal, brittle

Hardness

3–3½

Density

6.55

Optical properties

transparent to translucent; adamantine lustre; white, grey, yellowish, brownish

Occurrence  As a low-temperature hydrothermal mineral,

strontianite fills fissures and cavities or forms concretions in limestones and marls in which the Sr2+ ions are leached by lateral secretion from the adjacent country rocks. Rarely, strontianite is found in hydrothermal mineral and ore deposits or in carbonatites. Economic relevance  Strontianite is mainly used for pro-

duction of Sr metal by reduction of SrO with Al metal. Moreover, it forms a constituent in special glasses and serves as a raw material for the production of Sr nitrate, which is extensively employed in pyrotechniques, making use of the crimson flame coloration of Sr. In the 19th century, the now obsolete strontian process played an important role in recovering sugar from molasses (treacle). z Cerussite PbCO3 Crystal form and habit  Crystal class 2/m2/m2/m; often as well-developed single crystals or in groups (. Fig. 9.2),

dominant forms being pinacoids {010} and {001}, the orthorhombic dipyramid {111} and orthorhombic prisms {110}, {130} and {021}; typical are cyclic pseudo-hexagonal penetration twins (. Fig. 8.9), or lamellar twins, both on {110}. Cerussite with platy habit on {010} may occur as honeycombed, star- or fan-shaped intergrowths, whereas elongate crystals may form fascicular aggregates. Moreover, cerussite occurs in earthy-powdery masses.

Crystal structure  Isostructural with aragonite. Occurrence  Cerussite typically occurs in the oxidation

zone of lead deposits, commonly formed by weathering of galena, PbS.

Economic relevance  In places, cerussite constitutes a Pb

ore mineral.

z Witherite BaCO3 Crystal form and habit  Crystal class 2/m2/m2/m; well-de-

veloped crystals nearly always occur as circular twins on {110}, forming pseudo-hexagonal dipyramids. Massive witherite occurs as columnar or lamellar intergrowths. Physical properties Similar to aragonite and strontianite Cleavage

{010}, distinct, {110} poor

Fracture

conchoidal, brittle

Hardness



Density

4.30, i.e., higher than that of aragonite and strontianite but lower than that of cerussite

Optical properties

translucent; lustre: vitreous on crystal faces, greasy on fractures; colourless, white, yellowish

135

8.3 · Dolomite Group

Crystal structure  Isostructural with aragonite. Chemical composition  Commonly Ba2+ is replaced by

considerable amounts of Sr2+ but only small amounts of Ca2+ and Mg2+. Occurrence  Witherite is less common than aragonite,

strontianite or cerussite. It occurs in low-temperature hydrothermal veins or fills cavities in limestones and marls.

¯ are symmetrical . Fig. 8.10  Etch figures on the cleavage face (1011) ¯ on calcite with crystal class 32/m (a) but asymmetric on dolomite, due to the lower symmetry of crystal class 3¯ (b)

8.3  Dolomite Group

The mineral dolomite, CaMg(CO3)2, is not a solid solution between calcite and magnesite but a stoichiometric compound, a so-called double salt, with a cation ratio Ca:Mg = 1:1. The structure of dolomite is analogous to that of calcite (. Fig. 8.1) in which, however, Ca2+ and Mg2+ occupy [6]-coordinated sites in alternating layers parallel to (0001). Therefore, dolomite crystals do not contain mirror planes parallel to c as in calcite crystals. Instead, glide planes are developed in the dolomite structure. The differences in symmetry can be nicely documented by etch figures artificially produced on the rhombohedral cleavage planes {101¯1} of calcite and dolomite, respectively (. Fig. 8.10). At temperatures >500 °C, dolomite can accommodate slightly more Ca2+ and the Ca:Mg ratio becomes gradually higher than 1:1. In reverse, calcite coexisting with dolomite becomes progressively enriched in Mg2+ with increasing temperature. Thus, the miscibility gap in the system CaCO3–CaMg(CO3)2 is asymmetric and eventually vanishes at a temperature of ~1080 °C (. Fig. 8.11). The Ca:Mg ratio of calcite, coexisting with dolomite, can serve as a useful geological thermometer to estimate the temperature at which this mineral assemblage has formed. In contrast, magnesite can accommodate only relatively low contents of the CaMg(CO3)2 component, even at high temperatures. z Dolomite CaMg(CO3)2 Crystal form and habit  Crystal class 3¯; in nearly all cases, well-developed crystals of dolomite show the rhombohedron {101¯1} as dominant or exclusive crystal form. Frequently, these crystals are highly distorted and saddle-shaped, comprising a great number of subindividuals. Pressure-induced, lamellar twinning on {022¯ 1} is much rarer than the {101¯1} twinning in calcite, and the twin lamellae are parallel to the short rather the long diagonal of the cleavage rhombohedron {101¯1}. In rocks, dolomite forms granular aggregates.

. Fig. 8.11  Isobaric temperature-concentration diagram for the binary system calcite-dolomite showing the miscibility gap between the two end-members. The CO2 pressure of 50 bars is low but sufficient to avoid decarbonisation (modified after Goldsmith and Heard 1961)

Physical properties Cleavage

¯ perfect {1011}

Fracture

conchoidal

Hardness

3½–4

Density

2.87

Optical properties

transparent to translucent colourless, vitreous, in cases pearly lustre; commonly pale pink, yellowish or brownish, but also white, brown or black

8

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Chapter 8 · Carbonates, Nitrates and Borates

Distinction from calcite  In contrast to calcite, dolomite is hardly attacked by cold, diluted acids, such as HCl, but dissolves easily in hot acids with strong effervescence. Chemical composition  In most cases, the composition of dolomite is close to the theoretical formula CaMg[CO3]2, although some Mg2+ can be replaced by Fe2+, Mn2+ and occasionally Zn2+, thus leading to partial solid solution with ankerite, Ca(Fe,Mg)(CO3)2, (see below) and complete solid solution with kutnohorite, CaMn(CO3)2 and minrecordite CaZn(CO3)2. Occurrence  Dolomite is an important rock-forming

8

mineral, especially in calcareous sedimentary rocks, such as limestone and marl. During diagenesis, dolomite often replaces, or still coexists with calcite (7 Sect. 25.3.5). During progressive evaporation of seawater, the less soluble Ca-carbonates aragonite or calcite are the first to precipitate, followed by dolomite (7 Sect. 25.7.2). Metamorphic overprint of calcareous sedimentary rocks can produce dolomitic marble or dolomite-bearing calc-silicate rocks (7 Sect. 26.3.1). Dolomitic carbonatite can form from crystallisation of Ca–Mg carbonate magma. Dolomite also occurs as a gangue mineral in hydrothermal veins and as a hydrothermal replacement of limestone. Under super-saline anoxic conditions, sulfate-reducing bacteria, such as desulfobulbus mediterraneus, can mediate microbial nucleation of Mg-rich dolomite (Krause et al. 2012).

Economic relevance  In the past, burned dolomite was

mixed with tar and used in the steel industry as a cheap substitute for magnesite stones, for lining of oxygen converters (LD process) or electro furnaces. Moreover, burned dolomite is used as fluxing agent in iron smelting with basic scorification as well as a raw material for production of refractories and building materials.

z Ankerite CaFe(CO3)2

Chemical composition  Broad solid solution between dolo-

mite and ankerite, with up to 70 mol% CaFe(CO3)2, as well as between ankerite and kutnohorite, CaMn[CO3]2.

Occurrence  Ankerite is a constituent of several calcareous

sedimentary rocks, formed during diagenesis or by hydrothermal replacement of limestone. Together with siderite, it occurs in the carbonate facies of banded iron formation (BIF) as well as in some other metamorphic rocks. Ankerite and F ­ e-bearing dolomite can also form by evaporation of saline lake water. Moreover, ankerite occurs as a gangue mineral in hydrothermal veins and in some carbonatites.

Economic relevance  Similar to dolomite, ankerite can serve as fluxing agent in iron smelting.

8.4  Azurite-Malachite Group z Azurite Cu3[CO3(OH)]2 Crystal form and habit  Crystal class 2/m; azurite can occur

in beautiful, columnar or platy crystals, displaying many different faces (. Fig. 8.12). In other cases, fine-grained azurite forms dense or globular aggregates with reniform to botryoidal surface, concretions, earthy masses or coatings. Physical properties Cleavage

{011}, perfect, {100}, distinct

Fracture

conchoidal

Hardness

3½–4

Density

3.77–3.89

Optical properties

transparent to translucent; vitreous lustre; deep azure blue (. Fig. 8.12; name!), light blue in earthy masses

Streak

light blue

Crystal form and habit  Similar to dolomite, crystal class 3¯. Physical properties Cleavage

¯ perfect {1011}

Fracture

conchoidal

Hardness

3½–4

Density

2.93–3.10, i.e., higher than dolomite

Optical properties

transparent to translucent; vitreous, in cases pearly lustre; yellowish-white, by oxidation of Fe2+ changing to brown

Crystal structure  Like dolomite with statistical distribution of Fe2+ and Mg2+ in the Fe–Mg sheets of the structure.

Chemical composition  Generally close to the ideal for-

mula with 55.3% Cu.

Occurrence  Oxidation product of Cu-sulfides, such as

enargite, tetrahedrite and tennantite, occurring in oxidation zones of copper ores, especially in the vicinity of limestones (7 Sect. 24.6.1) or impregnating sandstones. Pseudomorphic replacement of azurite by malachite, uptaking additional water, is common. Spectacular examples are known from the Tsumeb Mine, Namibia (. Fig. 8.12).

Economic relevance  During medieval times and until the

17th century, azurite was used as a pigment for painting.

137

8.5 · Nitrates

Chemical composition  Commonly close to the ideal for-

mula with Cu contents of up to 57.4 wt%. However, Cu2+ can be replaced, at a considerable extent, by Zn2+ and Co2+.

Occurrence  Malachite is a common Cu mineral formed,

together with less abundant azurite, by oxidation of primary copper ores in oxidation zones or as impregnation in sandstones.

Economic relevance  Malachite is locally mined as Cu ore.

Polished malachite serves as an ornamental stone, used for handcrafts.

8.5  Nitrates

Nitrates are salts of nitric acid, composed of isolated anionic complexes [NO3]−. The most important nitrate mineral, nitratine, has a structure analogous to calcite (. Figs. 8.1, 8.6a), whereas nitre displays aragonite structure (. Fig. 8.6b).

z Nitratine NaNO3 . Fig. 8.12  Crystals of azurite (blue), partially replaced by malachite (green), on smithsonite (yellowish) from the former Tsumeb mine, Namibia. The turquoise-coloured, spherical aggregates are rosasite, (Zn,Cu)2CO3(OH)2. The extremely rare otavite, CdCO3, is white. Width of view is c. 3 cm; collection Hartwig Frimmel

Crystal form and habit  Crystal class 3¯2/m; well-developed

crystals with rhombohedral habit are rare. Commonly, nitratine occurs in granular aggregates. Physical properties

z Malachite Cu2CO3(OH)2 Crystal form and habit  Crystal class 2/m; malachite com-

monly displays acicular habit, arranged to form tufts, whereas coarse-grained, well-developed crystals are rare. More commonly malachite occurs as massive or layered aggregates with reniform to botryoidal surface, in earthy masses or as coatings. Physical properties Cleavage

{201} good, {010} distinct

Fracture

conchoidal

Hardness

3½–4

Density

3.7–4.1

Optical properties

translucent; lustre: vitreous, silky in fibrous, or dull in earthy varieties; deep green (. Fig. 8.12), light green in earthy masses

Streak

light green

Cleavage

¯ perfect {1011},

Fracture

conchoidal, brittle

Hardness

1½–2

Density

2.29

Optical properties

transparent; vitreous lustre; colourless, white, yellow, grey or reddish-brown

Solubility

somewhat hygroscopic, easily soluble in H2O

Occurrence  As main constituent of terrestrial evapo-

rites, nitratine exclusively occurs in zones of arid climate, such as the Atacama Desert in northern Chile and Peru (7 Sect. 25.7.1).

Economic relevance  The nitrate deposits of the Atacama Desert, consisting predominantly of nitratine (Chile saltpetre) and subordinate nitre (saltpetre), have been mined since the 1830s and used for production of fertilisers and explosives, as well elementary nitrogen. Nowadays, however, atmospheric nitrogen is the vastly predominant source for N2 and N-compounds (7 Sect. 25.7.1).

8

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Chapter 8 · Carbonates, Nitrates and Borates

z Nitre KNO3

Physical properties

Crystal form and habit  Crystal class 2/m2/m2/m; nitre forms aggregates of acicular crystals, powdery efflorescence or granular incrustations. Physical properties

8

Cleavage

{010}, perfect

Hardness

4–4½

Density

2.42

Optical properties

transparent to translucent; vitreous lustre; colourless to white

Crystal structure  Corner-sharing [BO3OH] tetrahedra and

Cleavage

{011}, perfect

Fracture

conchoidal, brittle

Hardness

2

Density

2.09–2.14

planar [BO2OH] triangles are connected, via corners and edges, to form undulating chains parallel to c (. Fig. 8.13a) as well as chains parallel to a. The Ca2+ cations and the H2O molecules are interspersed between these chains.

Optical properties

vitreous lustre; colourless, white, grey

Occurrence and economic relevance  See borax.

Solubility

non-hygroscopic, easily soluble in H2O

z Borax Na2B4O5(OH)4·8H2O

Occurrence  As subordinate constituent in the nitrate

deposits of the Atacama Desert, Chile. As efflorescence nitre coats surfaces of ground, walls and rock faces, e.g., in limestone caves.

Economic relevance  Mined as K-saltpetre (7 Sect. 25.7.1).

8.6  Borates

In borates, boron is either [3]- or [4]-coordinated with oxygen, thus forming [BO3]3− or [BO2OH]2− triangles as well as [BO4]5− or [BO3OH]4− tetrahedra. In the former case, the anionic radius of boron is 0.10 Å, i.e., somewhat larger than for C[3], in the latter, it is 0.20 Å, i.e., somewhat smaller than for Si[4]. Some of the borate minerals contain [BO3]3− triangles only, such as sassolite, H3BO3, others exclusively [BO4]5− tetrahedra, such as sinhalite that has olivine structure (. Fig. 11.3). Most of the borate minerals, however, contain both building blocks in their structure. Like in the silicates (7 Chap. 11), one can distinguish neso-, soro-, ino-, phyllo- and tecto-borates. Because of their structural diversity, borates are described as a separate class in the classification of Strunz and Nickel (2001).

z Colemanite CaB3O4(OH)3·H2O Crystal form and habit  Crystal class 2/m; well-developed

short prismatic crystals with a variety of faces are common, otherwise colemanite occurs in compact, granular masses.

Crystal form and habit  Crystal class 2/m; prismatic crys-

tals are common, moreover borax occurs in granular or cellular aggregates or as incrustations. Physical properties Cleavage

{100}, perfect

Fracture

conchoidal

Hardness

2–2½

Density

1.7

Optical properties

translucent; lustre: vitreous or dull, due to a thin coating of tincalconite, Na2B4O5(OH)4·3H2O; colourless, white, grey, yellow

Taste

sweetish-alkalescent

Solubility

dissolves readily in HCl

Crystal structure  Two types of undulating chains paral-

lel to c alternate with each other. These are composed of edge-sharing Na(H2O)6 octahedra and of isolated building blocks, each consisting of two BO3OH tetrahedra, sharing corners with two planar BO2OH triangles (. Fig. 8.13b, c). Occurrence  Borax is the most important boron min-

eral. It can form, together with kernite, ulexite and colemanite, in regions with arid climates, predominantly by evaporation of undrained lakes (7 Sect. 25.7.1) and by efflorescence from soils.

Economic relevance  Borax is the most important raw

material for boron and, in addition, an industrial semi-finished product for recovery of boron from other borate minerals. Boron is used in a broad range of industrial applications: Production of optical glass fibres, of china and enamel, as constituent of detergents, drugs and fertilisers, as solvent for metal oxides, flux for metallurgical processes,

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8.6 · Borates

neutron absorber in nuclear reactors, rocket fuel and as additive in motor fuels, also applied in airbags. Owing to its exceptional chemical, mechanical and thermal stabilities, boron-nitride, BN, is used for manufacturing high-T equipment and is applied in nanotechnology, such as for production of nanomeshes and nanotubes. Due to its Mohs’ hardness of 9½–10, BN is also used as abrasive. Another superhard ceramic material is boron carbide, ~B4C, Mohs’ hardness 9½, that is used, e.g., for tank armours and bulletproof vests. The most important boron deposits are situated in western and central Turkey and account for approximately half of the world’s annual boron production, followed by Rio Tinto Borax Mine near Boron, California. z Kernite Na2B4O6(OH)2·3H2O Crystal form and habit  Crystal class 2/m; partly occurring in large crystals; the largest crystal is recorded to have measured 2.44 × 0.9 × 0.9 m and ~3.8 t in weight (Rickwood 1981); rock-forming in coarse-grained, sparry aggregates. Physical properties Cleavage

{001} and {100}, perfect

Fracture

splintery or fibrous ∥b, brittle

Hardness

3

Density

1.95

Optical properties

translucent to transparent, in cases limpid; lustre: vitreous to pearly, on weathering dull, due to a thin coating of tincalconite; colourless or white

Solubility

dissolves slowly in H2O

Crystal structure  Undulating chains, formed of cor-

ner-sharing groups, each consisting of two corner-sharing [BO4] tetrahedra and one [BO2OH] triangle (. Fig. 8.13d). The chains are interconnected by polyhedra of NaO5(H2O) parallel to a and by Na2(H2O)3 parallel to c.

Occurrence  The main occurrence of kernite is in the

lower parts of a giant boron deposit, 6.5 km in length, 1.5 km broad and 75 m thick, intercalated in Tertiary clays of the Mohave Desert, California. There, kernite occurs together with borax, colemanite and ulexite, and was presumably formed by dehydration of borax in course of a weak metamorphic overprint. Another occurrence of kernite is in the three borax mines at Tincalayu, Argentina.

Economic relevance  See borax. . Fig. 8.13  Fragments of crystal structures of borates. a Colemanite, borate chains //c-axis, b-axis is horizontal. b Borax, projection on (010), c-axis vertical. c Borax, projection on (100), c-axis vertical. d Kernite, projection on (001), borate chains //b-axis. Legend for a, b and d: the BO3OH tetrahedra are shown in blue, the BO2OH triangles in green and the cations Ca or Na as red spheres (a, b and d after Strunz and Nickel 2001, c after Klein and Hurlbut 1985)

z Ulexite CaNaB5O6(OH)6·5H2O Crystal form and habit  Crystal class 1¯, triclinic; ulexite commonly occurs as delicate fibres, aggregated to round,

8

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Chapter 8 · Carbonates, Nitrates and Borates

fluffy aggregates known as “cotton balls” or occasionally as “television stone” with parallel arrangement of the fibres to generate properties of fibre optics. Physical properties Cleavage

{010}, perfect

Hardness

2½, however in fibrous aggregates an apparent hardness of 1 only

Density

1.96

Optical properties

transparent to translucent; silky lustre; snowwhite

and of Na(OH)2(H2O)4 octahedra, both ∥c, are linked by borate ions. Crystal structure  Chains of CaO3(OH)3(H2O)2 polyhedra

8

Occurrence and economic relevance  See borax.

References and Suggestions for Further Reading Evans RC (1976) Einführung in die Kristallchemie. Walter de Gruyter, Berlin, New York Goldsmith JR, Heard HC (1961) Subsolidus phase relationships in the system CaCO3–MgCO3. J Geol 69:45–74 Johannes W, Puhan D (1971) The calcite–aragonite transition, reinvestigated. Contrib Miner Petrol 31:28–38

Krause S, Liebetrau V, Gorb S et al (2012) Microbial nucleation of Mg-rich dolomite in exopolymeric substances under anoxic modern seawater salinity: new insight into an old enigma. Geology 40:587–590 Rickwood PC (1981) The largest crystals. Am Mineral 66:885–907 Strunz H, Nickel EH (2001) Strunz Mineralogical Tables, 9th edn. Schweizerbart, Stuttgart Whittaker EJW, Muntus R (1970) Ionic radii for use in geochemistry. Geochim Cosmochim Acta 34:945–956 Further Reading Anthony JW, Bideaux RA, Bladh KW, Nichols MC (2003) Handbook of Mineralogy, vol V: borates, carbonates, sulfates. Mineralogical Society of America Chang LL, Howie RA, Zussman J (1996) Rock-forming Minerals, vol. 5B, Non-silicates: sulphates, carbonates, phosphates, halides, 2nd edn. Longmans, Harlow, Essex Deer WA, Howie RA, Zussman J (1962) Rock-forming Minerals, vol. 5, Non-silicates, Longmans, London Deer WA, Howie RA, Zussman J (2013) An Introduction to the Rock-forming Minerals, 3rd edn. Mineral Soc, London Grew ES, Anovitz LM (eds) (1996) Boron—mineralogy, petrology, geochemistry. Rev Mineral Geochem 33 Klein C, Hurlbut CS Jr (1985) Manual of Mineralogy (after James D Dana), 20th edn. Wiley, New York Lippmann F (1973) Sedimentary Carbonate Minerals. Springer-Verlag, Berlin, Heidelberg Morse JW, Mackenzie FT (1990) Geochemistry of sedimentary carbonates. In: Developments in sedimentology, vol 48. Elsevier, Oxford, New York Reeder JR (ed) (1983) Carbonates: mineralogy and chemistry. Rev Mineral 11

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Sulfates, Chromates, Molybdates, Wolframates 9.1 Sulfates – 142 9.2 Chromates – 147 9.3 Molybdates and Wolframates – 147 Suggestions for Further Reading – 149

© Springer-Verlag GmbH Germany, part of Springer Nature 2020 M. Okrusch, H. E. Frimmel, Mineralogy, Springer Textbooks in Earth Sciences, Geography and Environment, https://doi.org/10.1007/978-3-662-57316-7_9

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Chapter 9 · Sulfates, Chromates, Molybdates, Wolframates

Introduction

9

The fundamental structural unit of the sulfate minerals is the anionic complex [SO4]2−. It forms a slightly distorted tetrahedron in which a central sulfur atom is connected with four oxygens by very strong covalent bonds. In contrast, the bonds between [SO4] and the cations are predominantly heteropolar. In the crystal structures of the H2O-free sulfate minerals baryte, celestine and anglesite, each of the relatively large cations Ba2+, Sr2+ and Pb2+ is coordinated with 12 oxygens at slightly different distances, forming [BaO12] polyhedra (. Fig. 9.3). In the less distorted anhydrite structure, each of the smaller Ca2+ cations is surrounded by 8 oxygens only, at nearly the same distance. The crystal structure of the H2O-bearing Ca sulfate mineral gypsum (. Fig. 9.5) consists of [SO4]2− layers parallel to {010} with strong bonds to Ca2+. These sheets are separated by layers of H2O molecules, linked by weak van der Waals bonds only. This explains the excellent cleavage of gypsum on {010} (. Table 9.1). Examples of important chromate, molybdate and wolframate minerals are given in . Table 9.2. The crystal structures of crocoite, PbCrO4, wulfenite, PbMoO4, and scheelite, CaWO4, are based on tetrahedral anionic complexes, similar to those in the sulfate minerals. In contrast, W in wolframite, (Fe,Mn)WO4, forms an octahedral WO6 complex.

9.1  Sulfates z Baryte (Barite) BaSO4 Crystal form and habit  Crystal class 2/m2/m2/m; the orthorhombic crystals are often well-developed, in places with many different faces. The habit of baryte is mainly tabular on {001}, but also elongate parallel to b or a with the prisms {101} or {011}, respectively (. Fig. 9.1). The combination of the basal pinacoid {001} and the vertical prism {210}, corresponding to the main cleavage planes, is frequently observed. Moreover, baryte can occur in granular or lamellar aggregates, and tabular crystals can form crested or hemispherical intergrowths (. Figs. 9.2, 9.3).

Physical properties Cleavage

{001}, perfect, {021} very good, {010} good

Hardness

2½–3½

Fracture

conchoidal, moderately brittle

Density

~4.5, i.e., very high for a non-metallic mineral and thus of diagnostic value

Optical properties

transparent to translucent, often dull; lustre: vitreous, pearly on the cleavage plane {001}; colourless, white, blue, yellow, orange or red, commonly in pale shades

Crystal structure  Mirror-symmetric [BaO12] polyhedra, connected on 4 corners and 3 edges with [SO4] tetrahedra, are

. Table 9.1  Some important sulfate minerals Formula

Crystal class

Baryte

BaSO4

2/m2/m2/m

Celestine

SrSO4

2/m2/m2/m

Anglesite

PbSO4

2/m2/m2/m

Anhydrite

CaSO4

2/m2/m2/m

CaSO4·2H2O

2/m

H2O-free sulfates

H2O-bearing sulfates Gypsum

. Table 9.2  The most important chromates, molybdates and wolframatesa Mineral

Formula

Crocoite

Pb[9]Cr[4]O

Crystal class

Wulfenite

Pb[8]Mo[4]O

Scheelite

Ca[8]W[4]O

Wolframite

(Fe,Mn)[6]W[6]O

2/m

4

4/m or 4

4

4/m

4 4

2/m

aCoordination

number against oxygen are given in raised square brackets

arranged in layers parallel to {001}. Strong atomic bonds within these layers give rise to the perfect cleavage parallel to {001}. Chemical composition  Ba2+ is in many cases replaced by

variable amounts of Sr2+, to a lesser extent by Ca2+ and rarely by Pb2+. Complete solid solution exists between baryte and celestine, SrSO4. Occurrence  Baryte, the most important Ba mineral, is

widely distributed. It represents the main constituent of hydrothermal baryte veins (7 Sect. 23.4.9) and occurs as gangue in hydrothermal sulfide veins and in volcanogenic-sedimentary or stratabound hydrothermal sulfide deposits (Sects. 23.5, 23.6). Baryte precipitates, together with anhydrite and sulfide minerals, from the white and black smokers, hydrothermal vents on the seafloor near mid-ocean ridges (. Figs. 23.10, 23.11). Finely dispersed baryte can occur in sedimentary rocks, such as limestones, sandstones and mudstones, e.g., in the pelagic sediments of the Eastern Pacific, often stained by bitumen in greyish black colours. Economic relevance  Baryte is used in a wide range of technical applications. Due to its high density, some 77% of the annual production worldwide is used as drilling mud in oil and gas well drilling, in order to avoid “blowouts”. Moreover, baryte serves as filler in paints and plastics, as constituent in glass ceramics, as barium meal in medicine, for radiation protection in X-ray techniques and

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9.1 · Sulfates

. Fig. 9.1  Crystal forms and habits of baryte. a Platy on {001}, b elongate b, c elongate a

. Fig. 9.3  Fragment of the baryte structure projected on (010); [BaO12] polyhedra (yellow) are connected over their corners and edges with [SO4]-tetrahedra (blue); (from Strunz and Nickel 2001)

Physical properties Cleavage

{001}, perfect, {021}, very good, {010}, good

Fracture

conchoidal

Hardness

3–3½

Density

~3.98

Optical properties

tranlucent to transparent; lustre: vitreous or pearly, greasy on conchoidal fractures; colourless to white, yellow, yellow-brown, orange, reddish, frequently blue, bluish or bluish-green, giving rise to the name “celestine” (from Latin coelestis = sky-blue)

Crystal structure  Like baryte with Sr2+ instead of Ba2+. Chemical composition  Sr2+ can be replaced by Ba2+ or

Ca2+, more rarely by Pb2+. Complete solid solution exists between celestine and baryte.

Occurrence  Celestine is rarer than baryte and predom. Fig. 9.2  Crested baryte (orange) on cerussite (colourless), Mibladen, Morocco; width of view is c. 10 cm

nuclear power stations, especially as a component in heavyweight concrete, for sound reduction in engine compartments, as automobile finishes, and for production of Ba preparations in the chemical industry. Furthermore, baryte forms an important constituent of the white pigment lithopone, applied in textiles, paper and paint.

inantly occurs in sedimentary rocks, such as dolomite, dolomitic limestone and marl, in which it formed directly through precipitation from seawater or by metasomatic exchange of anhydrite or gypsum with Sr-rich waters. A well-known locality is at Yate, near Bristol, England. Moreover, celestine occurs in tension joints or cavities in limestones, as concretions and, more rarely, in hydrothermal veins or in vugs of volcanic rocks.

Economic relevance  See strontianite; celestine is also used as filler in white or coloured paints.

z Celestine (Celestite) SrSO4

z Anglesite PbSO4

Crystal form and habit  Crystal class 2/m2/m2/m; habits

Crystal form and habit  Crystal class 2/m2/m2/m; the small but often well-developed, facetted crystals resemble those of baryte, usually displaying tabular, more rarely elongate or short prismatic, habit.

similar to baryte, tabular on {001} or prisms elongated parallel to a or b, fibrous in tension joints or cavities of limestone, granular or sparry aggregates, occasionally as nodules.

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Chapter 9 · Sulfates, Chromates, Molybdates, Wolframates

Physical properties Cleavage

{001}, perfect, {210}, less perfect

Fracture

conchoidal

Hardness

3

Density

6.2–6.4, remarkably high

Optical properties

transparent to translucent; adamantine lustre; colourless or coloured in pale shades

Crystal structure  Like baryte with Pb2+ instead of Ba2+. Chemical composition  Theoretically, anglesite contains

68.3 wt% Pb which is, however, in most cases replaced by considerable amounts of Ba.

Occurrence  Anglesite occurs as an alteration product of

galena in the oxidation zone of sulfidic Pb deposits.

9

Economic relevance  Locally, anglesite is smelted, together

with the primary ores, to produce lead.

z Anhydrite CaSO4 Crystal form and habit  Crystal class 2/m2/m2/m; rare

well-developed crystals are tabular on {001}, isometric or elongate along a. Lamellar twins on {110} may form in a tectonic stress field. More common is anhydrite as fine- to coarse-grained, or sparry, rock-forming mineral in evaporite deposits Nodular masses of anhydrite with a so-called chicken-wire texture are typical of replacement of gypsum. Physical properties Cleavage

anhydrite displays three different cleavage planes {010}, perfect, {100} very good, {001}, good. These are oriented perpendicular to each other resulting in nearly cubical cleavage bodies

Fracture

brittle

Hardness

3–3½

Density

2.9–3.0

Optical properties

transparent to translucent; lustre: pearly to vitreous on cleavage planes (001) and (010), respectively; the pure mineral is colourless to cloudy white, whereas rock-forming anhydrite is often grey, bluish, violet or reddish

Crystal structure  [CaO8] groups form distorted triangular

dodecahedra. These are connected, via edges, with [SO4] tetrahedra to constitute alternating chains parallel to c.

Chemical composition  Almost ideal but small amounts of Ca can be replaced by Sr. Occurrence  Anhydrite is an important rock-forming mineral of marine salt deposits (7 Sect. 25.7.2) in which it occurs,

often together with gypsum, halite (rock salt) and potassic salts, such as sylvite and carnallite (7 Chap. 6). The primary Ca-sulfate to precipitate from evaporating seawater is gypsum, CaSO4·2H2O, whereas anhydrite is typically formed by subsequent diagenetic dehydration of gypsum. Experiments on the solubility of anhydrite and gypsum indicate that anhydrite can precipitate directly from seawater (average salinity 35‰) when evaporating at temperatures >49 °C whereas, at higher temperatures, it precipitates only from solutions of increasingly higher salinity (Blount and Dickson 1973). High temperatures combined with high salinities are realised in the vicinity of black and white smokers that vent on the seafloor near mid-ocean ridges, where anhydrite is deposited together with baryte and sulfide minerals (7 Sect. 23.5.1: . Fig. 23.13). Moreover, anhydrite can form as efflorescent in deserts and dry steppes, and as sublimation product on active volcanoes (7 Sect. 14.5). Upon weathering, anhydrite slowly re-hydrates to form gypsum with a volume increase of ca. 60%. Economic relevance  Anhydrite is used in the chemical industry, especially for production of sulfuric acid. Moreover, it is an important additive in building materials and is used for the production of fast-acting binders and floor screeds.

z Gypsum CaSO4·2H2O Crystal form and habit  Crystal class 2/m; gypsum may

form well-developed crystals, often arranged in druses. Their habit is frequently tabular with dominant pinacoid {010} (. Fig. 9.4a, b), less commonly prismatic, elongate along c, e.g., with {120} and {1¯20}, rarely fibrous with {120} and {1¯20}. Individual crystals can attain large or even giant sizes, especially in gypsum caves (. Fig. 9.6). Swallowtail twins on {100} are common (. Fig. 9.4c), whereas Montmartre twins on {001}, grown in clay at Montmartre near Paris, are rarer. These show mostly a lense-shaped habit due to the suppression of the vertical prism (. Fig. 9.4d). Penetration twins are common. Rock-forming gypsum occurs in fine-grained or sparry masses. Pure white, finegrained gypsum is called alabaster, fissure-filling, fibrous gypsum with silky lustre is known as satin spar.

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9.1 · Sulfates

. Fig. 9.4  Crystal forms of gypsum with orthorhombic prisms {011}, ¯ {120}, ¯ {120} ¯ {011}, and pinacoid {010}. a Single crystal, platy on {010}; b single crystal, elongate c, traces of cleavage planes are indicated; c genuine swallowtail twin on {100}; d Montmartre twin on {001} Physical properties Cleavage

¯ {010} perfect, {100} distinct, {111} fibrous (. Fig. 9.4b–d); cleavage folia {010} are inelastically pliable

Fracture

conchoidal, brittle

Hardness

2, standard mineral of Mohs’ scale of hardness

Density

2.30–2.37

Optical properties

transparent to translucent; transparent crystals of gypsum displaying beautiful moonshine lustre are known as selenite, colourless; lustre on cleavage planes: pearly on {010}, vitreous on {100}, silky on ¯ {111}; white, yellowish or reddish, due to inclusions of bitumen grey or brown

Crystal structure  Edge-sharing [CaO6(H2O)2] polyhedra

and [SO4] tetrahedra form indefinite chains parallel to c. These constitute double layers parallel to {010} that are connected by hydrogen bonds, giving rise to the perfect {010} cleavage (. Fig. 9.5). Chemical composition  Virtually ideal CaSO4·2H2O, addi-

tional elements are present at best in trace amounts.

Occurrence  Locally, gypsum is an important rock-forming mineral, especially of marine salt deposits (7 Sect. 25.7.2) in

which it occurs, commonly together with anhydrite, halite (rock salt) and potassic salts, such as sylvite and carnallite. Primary syn-sedimentary gypsum is commonly transformed during subsequent burial beneath younger strata to anhydrite which, however, can re-hydrate during exposure to meteoric waters on or near the land surface in humid climates to form secondary gypsum. In continental

. Fig. 9.5  The crystal structure of gypsum, projected on (001) illustrates the perfect (010) cleavage; The parallel double layers consist of edge-sharing [CaO6(H2O)2] polyhedra and [SO4] tetrahedra (yellow with small red spheres) and are connected by (OH) groups; red: oxygen, green: hydrogen; blue: calcium (after Deer et al. 2013)

evaporites (7 Sect. 25.7.1) gypsum can crystallise at the margins of salt lakes, salt pans or sabkhas, often forming concretions in muddy or marly sediments. One variety of this is desert rose, rosette-like intergrowths of gypsum crystals that are rich in sand inclusions. In salt deserts and dry steppes, gypsum occurs as efflorescent from sulfate-bearing solutions, and around volcanic vents, it can precipitate as a sublimation product of sulfurous volcanic gases (7 Sect. 14.5). The possibly most outstanding occurrence of gypsum is the gypsum cave Cueva de los Cristales in the Naica Mine at Santo Domingo, Chihuahua, Mexico. It is lined by the largest known crystals of the gypsum variety selenite, reaching up to 14 m length and up to 1 m thickness (. Fig. 9.6). The crystals show striated prism faces {120}, {140} and {160} as well as {1¯11}. Compared to common tabular crystals, the {010} pinacoid is less well-developed on elongate, and missing on short-prismatic, Naica crystals (García-Ruiz et al. 2007). Naica is an important Zn–Pb–Ag skarn deposit. Mineralisation took place about 26 Ma ago by metasomatic reaction between Lower Cretaceous limestone and highly saline hydrothermal solutions (7 Sect. 23.3.1). During a late hydrothermal stage, oxidation of the sulfide minerals led to the formation of diluted sulfurous acid, H2SO3, that reacted with the host limestone to form geothermal waters rich in Ca-sulfate. From these, lenses of anhydrite were precipitated that are widely distributed below the 240 m level of the mine. The gypsum caves were formed later, 1–2 Ma ago, when hydrothermal water penetrated along tectonic fault zones, thereby creating solution cavities in the surrounding limestone (. Fig. 9.7a). The giant gypsum crystals of the Naica Mine can serve as an excellent example for the influence of dissolution equilibria on crystal growth (García-Ruiz et al. 2007). Solubility of anhydrite markedly increases with lowering of temperature, whereas it changes much less for gypsum. At 59 °C, both solubility curves intersect and both minerals display the same solubility (. Fig. 9.7b). At present, the temperature of the low-salinity, sulfate- and carbonate-rich thermal waters in the Naica Mine range between 48 and 59 °C. As demonstrated by fluid-inclusion studies, the growth of the gypsum crystals in the Cueva de los Cristales took place at a temperature of around 54 °C. This means that the thermal water was slightly

9

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Chapter 9 · Sulfates, Chromates, Molybdates, Wolframates

9

. Fig. 9.7  a Schematic cross section through a tension joint in limestone of the Naica skarn deposit, where lenses of anhydrite are dissolved and giant crystals of gypsum have grown, at a temperature of c. 54 °C, from anhydrite-saturated thermal water; b solubility of gypsum and anhydrite in a concentration (in millimole/litre) versus temperature phase diagram; conditions of gypsum crystallisation shown in red. CA–CG(t) is the difference between the solubilities of gypsum and anhydrite at a given temperature (modified after Forti and Sanna 2010)

. Fig. 9.6  Giant gypsum crystals, Cueva de los Cristales, Naica Mine at Santo Domingo, Chihuahua, Mexico; the crystals show the typical moonshine lustre that gave rise to the name selenite (photograph Javier Trueba, Madrid, Spain, courtesy of Contacto/Agentur Focus, Hamburg, Germany)

undersaturated in anhydrite, whereas gypsum started to crystallise (. Fig. 9.7b). At this temperature, the solubility difference between the two minerals, i.e., the super-saturation in gypsum and, consequently, the nucleation frequency are extremely small. Thus nucleation and subsequent crystal growth were limited to only a few sites from which giant selenite crystals could develop (García-Ruiz et al. 2007). In contrast, much more but considerably smaller crystals were formed at lower temperatures in the nearby Cueva de las Espadas, the “Cave of the Swords”. In the Cueva de los Cristales, the rare situation occurred that favourable conditions for crystal growth, i.e., limited nucleation and high growth rate, remained virtually unchanged for more than a million years. Growth of the giant crystals continued until the 1980s, when the thermal waters were pumped down by the mine management and were cooled by mixing with surface waters. In situ 230Th/234U isotope analyses performed on one sample that was taken ca. 5 cm within a giant crystal yielded an age of 34544 ± 819 years, implying that the giant selenite crystal grew over a period of several 100,000 years. This agrees with experimental data that suggest an annual growth rate of 0.004 mm and a total time span of some 250,000 years for the growth of these giant crystals (Sanna et al. 2011).

Economic relevance  For thousands of years gypsum (plaster rock) has been an important raw material in arts and architecture. Upon heating to 120–130 °C, gypsum loses most of its structurally bound water and transforms to calcined gypsum, approaching the hemihydrate composition CaSO4·½H2O. By stirring with water, the hemihydrate hardens and recrystallises in short time under uptake of water and formation of artificial gypsum. Therefore, calcined gypsum is technically used as moulding plaster, stucco (“plaster of Paris”), for production of plasterboards for drywalls as well as for surgical and dental applications. At temperatures above 180 °C, natural gypsum loses most of its structurally bound water and is transformed into the metastable, hexagonal anhydrite modification γ -CaSO4 (with 250 °C, converts into the totally H2O-free, orthorhombic, “natural” anhydrite βCaSO4. Both modifications have lost their capability for quick uptake of water and can harden only over extended periods. At temperatures >1000 °C, β-CaSO4 partly dissociates to release SO3 and forms a solid solution between CaSO4 and CaO, which is widely used in civil engineering, e.g., for producing plastic and refractory mortars or floor screeds. Besides, calcined gypsum is applied as fertiliser and for the production of sulfuric acid and elementary sulfur. A considerable proportion of the gypsum demand is met by synthetic gypsum, so-called FGD gypsum that is produced in the course of the desulfurisation of flue gas, i.e., by reaction of SO2 with finely ground limestone, in modern coal-fired power stations.

9.2  Chromates

. Fig. 9.8  Crocoite, Dundas, Tasmania; width of view is 4 cm

Crystal structure  Isostructural with monazite (7 Chap. 10) and similar to zircon (. Fig. 11.7), consisting of PbO9 pol-

z Crocoite PbCrO4

yhedra and CrO4 tetrahedra.

Crystal form and habit  Crystal class 2/m; crocoite can form columnar, vertically striated crystals (. Fig. 9.8), up

to several centimetres in size, or smaller ones with spiry or acicular habit. In addition, it can occur as massive or disseminated aggregates, as encrustation or efflorescent.

Physical properties Cleavage

{110}, fairly perfect

Fracture

rough, conchoidal

Hardness

2½–3

Density

5.9–6.1,

Optical properties

transparent to translucent; adamantine to resinous lustre, in some cases greasy; red, yellowish-red, orange (. Fig. 9.8)

Streak

9

147

9.2 · Chromates

orange-yellow

Chemical

composition Theoretical

64.1 wt% Pb, 16.1 wt% Cr.

metal

contents:

Occurrence  Crocoite is formed in the oxidation zone

above hydrothermal Pb deposits that are hosted by rocks containing chromite. In these rare cases, Pb- and Cr-bearing weathering solutions have mixed and reacted with each other.

Economic relevance  Of no economic significance. The element chromium was first detected in the mineral crocoite.

9.3  Molybdates and Wolframates z Wulfenite PbMoO4

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Chapter 9 · Sulfates, Chromates, Molybdates, Wolframates

Crystal form and habit  Wulfenite belongs to the crystal class 4/m and in many cases develops tetragonal-dipyramids. Besides, however, there are also wulfenite crystals with lower symmetry showing forms of the tetragonal-pyramidal class 4 (Hintze and Linck 1930), which leads to diverging descriptions in textbooks. Most wulfenite crystals are of quadratic, tabular habit on {001}, with variable thickness, whereas elongate prismatic or (di-)pyramidal crystals are less common. Moreover, wulfenite may occur in druses, as massive, dense aggregates, as incrustation or as efflorescent. Wulfenite can form pseudomorphs after galena.

9

. Fig. 9.9  Scheelite, combinations of various tetragonal dipyramides

Physical properties

Physical properties

Cleavage

{101} distinct

Cleavage

{011} distinct

Fracture

uneven to conchoidal, brittle

Fracture

conchoidal, brittle

Hardness

4½–5

Hardness

3

Density

Density

6.7–6.9

5.9–6.1, unusually high for a non-metallic mineral and thus a diagnostic feature

Optical properties

transparent to subtranslucent; adamantine to resinous lustre; yellow, orange yellow or reddish orange, rarely olive-green, grey, blue, brown or colourless

Optical properties

Streak

white

Special property

wulfenite is piezoelectric (7 Sect. 1.4.3) indicating a polar c-axis and conforming to the crystal class 4

translucent on edges, rarely transparent; lustre: resinous to adamantine on crystal and cleavage planes, greasy on fractures (similar to quartz); colourless, greyish white, yellowish, greenish, orange or brownish; refractive indices: nω 1.920, nε 1.934, i.e., distinctly higher than those of quartz

Diagnostically important property

bluish-white fluorescence in short wavelength UV radiation

Crystal structure  Isostructural with scheelite. composition Theoretical metal contents: 56.4 wt% Pb, 26.1 wt% Mo; wulfenite forms a solid solution with stolzite, PbWO4.

Chemical

Occurrence  Together with other secondary Pb minerals,

wulfenite occurs in the oxidation zone of hydrothermal Pb deposits.

Economic relevance  Wulfenite is rarely mined as Mo

ore, whereas by far the most important Mo ore mineral is molybdenite, MoS2 (7 Sect. 5.3).

z Scheelite CaWO4 Crystal form and habit  Crystal class 4/m; scheelite displays

nearly octahedral habit because of the dominance of tetragonal dipyramids {111} or {112} (. Fig. 9.9a, b). Frequently, these faces display an oblique striation, caused by the combination with subordinate dipyramids, especially {21¯3}, {101} and {211} (. Fig. 9.9b), thus reflecting the lack of mirror planes parallel to c. The striation on {112} helps to distinguish twinned from simple crystals. Scheelite may occur as single crystals but commonly forms massive or disseminated aggregates. Occasionally, scheelite encrusts crystals of quartz.

Crystal structure  Scheelite displays a distorted zircon structure (. Fig. 11.7), in which isolated [WO4]4− tetrahedra, slightly flattened normal to c, are connected via O–Ca–O bridges, to form a three-dimensional framework. The large Ca2+ cations are coordinated to 8 oxygens. composition Theoretical metal content: 63.9 wt% W, but typically some W is substituted by Mo and partial solid solution exists between scheelite and powellite, CaMoO4.

Chemical

Occurrence  Scheelite occurs predominantly in skarn

deposits, where it is formed by contact-metasomatic replacement of limestone reacting with high-temperature hydrothermal fluids (7 Sect. 23.3.1). Moreover, scheelite may occur in hydrothermal veins, often accompanied by cassiterite, or more rarely in granitic pegmatites.

Economic relevance  Scheelite is, next to wolframite, the

most important ore mineral for tungsten (see below). Due to their high refractive indices and despite of their low hardness, clear, flawless scheelite crystals are cut and polished to be used as facetted gemstones.

149 Suggestions for Further Reading

z Wolframite (Fe,Mn)WO4 Crystal form and habit  Crystal class 2/m; well-developed crystals may attain relatively large sizes. They are commonly tabular on {100}, but others are short or long prismatic on {110}. Faces parallel to c display vertical striation. Wolframite crystals can be twinned on (100). In most cases, however, wolframite occurs in massive, conchoidal, bladed or columnar aggregates. Physical properties Cleavage

{010} perfect; in contrast, otherwise similar sphalerite, ZnS, shows two different cleavage ¯ whereas in cassiterite, planes (110) and (110), SnO2, the cleavages are imperfect or poor

Fracture

uneven to rough

Hardness

4–4½

Density

7.0–7.5, very high and increasing with Fe/Mn ratio

Optical properties

transparent to translucent; submetallic to resinous lustre; black in Fe-rich, brown in Mn-rich members of the solid solution

Streak

nearly black to brownish black depending on Mn content

Chemical composition  Wolframite forms a complete solid solution between the end members ferberite, FeWO4, and huebnerite, MnWO4, which, however are rarer than the intermediate members. The theoretical W contents are 60.5 wt% for ferberite and 60.7 wt% for huebnerite. Crystal structure  Edge-sharing WO6 and (Fe,Mn)O6 octa-

hedra form “zig-zag” chains parallel to c and sheets parallel to (100); judging from this structural arrangement, Strunz and Nickel (2001) classify wolframite no longer as a wolframate, but place it in their class 4, oxides.

Occurrence  Wolframite occurs, together with cassiterite,

SnO2, and molybdenite, MoS2, in quartz-rich pegmatite dykes and in hydrothermal impregnations in altered granite, known as greisens, and in related hydrothermal veins (7 Sect. 23.2.2). Moreover, wolframite can be concentrated in placer deposits. Economic relevance  Wolframite is, next to scheelite, the

most important ore mineral for tungsten, a metal with

an extremely high melting point of 3422 °C. Therefore, it serves as an important constituent in super-alloys with Fe, Ni, Co and Mo. Tungsten steels are implemented, e.g., in rocket nozzles, turbine blades, high-speed drilling equipment and armour-piercing ammunition. Moreover, tungsten is used for filaments and targets in X-ray tubes, as single-crystal filaments of incandescent light bulbs, as electrodes for tungsten inert gas (TIG) welding and for radiation shielding. A high amount of the world production serves for manufacturing of tungsten carbide, WC, standing out by a high melting point of ~2870 °C and a Mohs’ hardness of 9, widely implemented in cutting tools, drilling bits, surgical instruments, trekking and skiing poles as well as armour-piercing ammunition. WC is also used for colouration of glass and china. In 2019, the annual worldwide production of W metal amounted to 85,000 t. The main producing countries are China (70 kt), followed by Vietnam (4.8 kt), Mongolia (1.9 kt), Russia (1.5 kt), Bolivia (1.2 kt), North Korea (1.1 kt), and Rwanda (1.1 kt) (U.S. Geological Survey 2020).

Suggestions for Further Reading Blount CW, Dickson FW (1973) Gypsum–anhydrite equilbria in the system CaSO4–H2O and CaSO4–NaCl–H2O. Amer Mineral 58:323–331 Forti P, Sanna L (2010) The Naica project: a multidisciplinary study of the largest gypsum crystals of the world. Episodes 33:1–10 García-Ruiz JM, Villasuso R, Ayora C et al (2007) The formation of natural gypsum megacrystals in Naica, Mexico. Geology 35:327–330 Hintze C, Linck G (1930) Handbuch der Mineralogie, I.3.2, Sulfate, Chromate, Molybdate, Wolframate, Uranate. de Gruyter, Berlin Leipzig Sanna L, Forti P, Lauritzen SE (2011) Preliminary U/Th dating and the evolution of gypsum crystals from Naica caves. Acta Carsologica, Postojna, Slovenia 40:17–28 Strunz H, Nickel EH (2001) Strunz Mineralogical Tables, 9th edn. Schweizerbart, Stuttgart U.S. Geological Survey (2020) Mineral commodity summaries 2020: U.S. Geological Survey, 200 p. 7 https://doi.org/10.3133/mcs2020 Further Reading Alpers CN, Jambor JL, Nordstrom DK (eds) (2000) Sulfate minerals— crystallography, geochemistry and environmental significance. Rev Mineral Geochem 40 Anthony JW, Bideaux RA, Bladh KW, Nichols MC (2003) Handbook of Mineralogy, vol V: Borates, carbonates, sulfates. Mineralogical Society of America Chang LL, Howie RA, Zussman J (1996) Rock-forming Minerals, vol 5B, 2nd edn. Non-silicates: sulphates, carbonates, phosphates, halides. Longmans, Harlow Essex Deer WA, Howie RA, Zussman J (2013) An Introduction to the Rock-forming Minerals, 3rd edn. The Mineralogical Society, London

9

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Phosphates, Arsenates, Vanadates Suggestions for Further Reading – 155

© Springer-Verlag GmbH Germany, part of Springer Nature 2020 M. Okrusch, H. E. Frimmel, Mineralogy, Springer Textbooks in Earth Sciences, Geography and Environment, https://doi.org/10.1007/978-3-662-57316-7_10

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Chapter 10 · Phosphates, Arsenates, Vanadates

Introduction This class of minerals comprises a particularly large variety of different mineral species because of a wide range in ionic substitution possibilities. All members contain tetrahedral anionic complexes, such as [PO4]3−, [AsO4]3− and [VO4]3−, as structural units, whereby P5+, As5+ and V5+ can substitute for each other. The cations are coordinated with 9 oxygens. Apatite, by far the most important and most widespread mineral of this class, contains F−, Cl− and (OH)− as additional anions that can mutually replace each other. The phosphate pyromorphite, the arsenate mimetite and the vanadate vanadinite display crystal structures similar to apatite, in which Pb2+ occurs instead of Ca2+ (. Table 10.1).

z Monazite

CePO4 Crystal form and habit  Crystal class 2/m; monazite develops tabular or prismatic crystals, abundantly twinned on {001}, or forms granular aggregates.

10

Physical properties Cleavage

{100} moderate, {001} variable

Fracture

conchoidal, brittle

Hardness

5

Density

5.0–5.3

Optical properties

translucent; resinous to vitreous lustre; light yellow to dark reddish brown, in cases nearly white

Crystal structure  The most important structural ele-

ments are linear chains along c consisting of alternating [PO4] tetrahedra and edge-sharing [CeO9] polyhedra. These are cross-linked, parallel to a, by “zig-zag” chains of edge-sharing [CeO9] polyhedra. The monazite structure resembles that of zircon (. Fig. 11.7).

Chemical composition  Ce can be replaced by La, Nd and Th. Monazite-(Ce) is characterised by Ce  >  (La  +  Nd), monazite-(La) by La > (Ce + Nd) and monazite-(Nd) by Nd > (La + Ce). The ThO2 contents are highly variable and can reach nearly 20 wt% (Chang et al. 1996). Occurrence  Monazite is the most common mineral of the

rare-earth elements (REE) and occurs as accessory phase in a variety of rocks, such as granites, rhyolite, gneisses and micaschists. Higher concentrations of monazite can be found in phosphate pegmatites as in Iveland, Norway, or in Madagascar. The most important sources of monazite and other REE minerals, however, are related to carbonatites (7 Sect. 21.4). Monazite is also found in hydrothermal veins. Sedimentary transport leads to concentration of monazite in placer deposits as well as in coastal or fluvial sands.

Economic relevance  Due to their excellent electronic, magnetic, optical and catalytic properties, the rare earth

. Table 10.1  Important phosphates, arsenates, vanadates Mineral

Formula

Crystal class

Monazite

CePO4

2/m

Xenotime

(Y,Yb)PO4

4/m2/m2/m

Apatite

Ca5(PO4)3(F,Cl,OH)

6/m

Pyromorphite

Pb5(PO4)3Cl

6/m

Mimetite

Pb5(AsO4)3Cl

6/m

Vanadinite

Pb5(VO4)3Cl

6/m

elements (REE) form a special group of chemical elements, which are indispensable for a wide range of modern technologies. Therefore, the REE are considered “critical” raw materials for the global economy and of strategic relevance (e.g., Hatch 2012; Chakhmouradian and Wall 2012; Simandl 2014). Industrially produced REE compounds are applied in two groups of industrial processes (Hatch 2012). They can serve as 5 process enablers, such as fluid-cracking catalysts (FCCs) used in petroleum-refining industry, as catalytic converters in automobile industry, and as additives to improve the polishing of glass planes, mirrors, TV screens, computer displays and wafers for production of silicon chips; 5 technology building blocks, e.g., of highly efficient permanent magnets, battery cells for energy storage, and phosphorescent materials in plasma screen and liquid crystal displays (LCDs), light-emitting diodes (LEDs) and compact fluorescent lamps (CFLs). Among the primary REE occurrences, the giant deposits of Bayan Obo, Inner Mongolia, northern China, and Mountain Pass, California, both related to carbonatites, play an exceptional role. Bayan Obo contains about 43% of the world resources in REE, mainly bound to bastnäsite, (Ce,La,Nd,Y)CO3(F,OH), and monazite, with total amounts of ca. 57 Mt REE2O3. In addition, monazite is extracted from coastal placer deposits in Australia, Brazil, India, Malaysia and Florida. In 2019, the PR China produced c. 132,000 t REE2O3 or nearly 63% of the annual worldwide supply of c. 210 kt, followed by USA (26 kt), Myanmar (22 kt), Australia (21 kt), India (3 kt), Russia (2.7 kt), and Madagascar (2.0 kt) (U.S. Geological Survey 2020). At present China holds the monopoly in processing REE ores (Hatch 2012). In geochronology, monazite is successfully used, similar to zircon (7 Sect. 11.1), for radiometric dating with the U–Pb method (7 Sect. 33.5.3). z Xenotime

(Y,Yb)PO4 Crystal class 4/m2/m2/m; crystal structure similar to that of monazite; cleavage {100} perfect; hardness 4–5; density 4.5–5.1; vitreous to resinous lustre; colour yellowish, reddish, pale brown, white.

153 Introduction

z Apatite

Ca5(PO4)3(F,Cl,OH) Crystal form and habit  Crystal class 6/m; well-developed

crystals are common and can attain large sizes. They typically display the hexagonal prism {101¯0} combined with the hexagonal dipyramid {1011¯} and the basal pinacoid {0001}, their relative sizes defining the short or more elongate prismatic habit of the crystals (. Figs. 10.1a, b, 10.2). Transparent, short prismatic crystals, grown on fissures or in druses, mostly show additional dipyramid forms, especially {112¯ 1} and {213¯1} (. Fig. 10.1a). Apatite needles or prisms of microscopic size occur as accessory phase in a great variety of rock types. Phosphorite is a sedimentary rock containing granular, dense or cryptocrystalline apatite, commonly mixed with amorphous (Ca-)phosphates and carbonates (7 Sect. 25.6). Incrustations of phosphorite, derived from amorphous, colloidal matter or precipitated from organisms, often display grapelike, reniform or stalactitic surfaces.

. Fig. 10.1  Crystal form and habit of apatite; a short-prismatic habit, ¯ rich in different forms, such as various hexagonal dipyramids {1011}, ¯ hexagonal prism {1010} ¯ and {2131}, ¯ and basal pinacoid {0001}; {1121} The combination of these forms indicates that there is a mirror plane perpendicular to the c axis, but no mirror planes //c; b simple prismatic ¯ subordinate {1011} ¯ as well as {0001} habit with predominant {1010},

Physical properties Cleavage

¯ poor {0001} and {1010}

Fracture

uneven to conchoidal

Hardness

5, reference mineral of Mohs’ scale of hardness

Density

3.1–3.53

Optical properties

transparent to translucent; vitreous to subresinous lustre on some of the crystal faces, greasy on conchoidal fractures; apatite can be colourless or white, but commonly occurs in different colours, such as yellowish or bluish green, blue or brown

Note: apatite can be easily confused with other minerals, a fact indicated by its name, derived from the Greek word απατάω = to bluff, to cheat. structure Corner-sharing [CaO9] polyhedra, forming chains along c, are connected via edges and corners with [PO4] tetrahedra in hexagonal arrangement; the additional anions F, Cl and (OH) are situated in wide channels, oriented parallel to c (. Fig. 10.3). Crystal

Chemical composition  F, Cl and (OH) can replace each other.

Depending on the dominance of either of these additional anions, fluorapatite, the most widespread variety, chlorapatite and hydroxylapatite are distinguished. Partial   coupled substi tution (OH)− /PO4 3− ⇋ O2− /CO3 2− leads to the formation of carbonate apatite. Moreover, [PO4]3− can be replaced, to a limited extent, by [SO4]2− and, simultaneously, by [SiO4]4−, whereby charge balance is retained by the coupled substitution P5+ ⇋ S6+ and P5+ ⇋ Si4+. Commonly, the cation Ca2+ can be partially replaced by Sr2+, to a lesser extent by Ba2+, Pb2+, Mn2+, Na+ and REE (see Chang et al. 1996). The robust crystal structure of apatite can accommodate more than half of the long-lived chemical elements known, giving rise to over 40 different phosphate, arsenate,

. Fig. 10.2  Crystals of apatite displaying the combination of the ¯ and the hexagonal dipyramid {1011} ¯ on a predominant prism {1010} matrix of calcite from Sljudjanka, Siberia; Width of view is c. 5 cm

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Chapter 10 · Phosphates, Arsenates, Vanadates

ordered bio-apatite with decreasing [HPO4]2− but increasing (CO3OH)3− contents eventually develops during an extended process of maturation and recrystallisation (Boskey 2007). Over extended periods of time, bio-apatite is concentrated predominantly in phosphorite deposits where it often serves as petrifying mineral of fossil bones and excrements, known as guano. Economic relevance  Apatite is the main phosphate in inor. Fig. 10.3  Crystal structure of apatite, projected along the c axis onto the (0001) plane; the large [CaO9] polyhedra are shown in light blue, the [PO4] tetrahedra in dark blue and the (OH,F,Cl) anions as light blue circles

vanadate, sulfate and silicate minerals (Hughes and Rakovan 2015).

10

Occurrence  As an accessory mineral, apatite is widely distributed in various rock types, not only on Earth but also in other planetary bodies. The F/Cl/(OH) ratios of apatite in lunar rock samples and meteorites, in Martian meteorites and chondritic meteorites derived from the asteroid belt provide useful information on volatile abundances and processes throughout the solar system and during its history (McCubbin and Jones 2015). Apatite yields information on the behaviour of volatile and trace elements in terrestrial or extraterrestrial magmas (Webster and Piccoli 2015) and is used for isotopic age dating, e.g., by the U–Pb method (Chew and Spikings 2015). In rare cases, inorganic apatite occurs as a main constituent, such as in carbonatites (Sects. 13.2.3, 21.4), e.g., at Phalaborwa, South Africa, in orthomagmatic magnetite-apatite deposits, e.g., at Kiruna, northern Sweden (7 Sect. 21.3.3), in phosphate pegmatite (7 Sect. 22.3) as well as in hydrothermal veins and impregnations. Together with collagen and other matrix proteins, bio-apatite is the most important constituent of human and vertebral teeth and bones (e.g., Pasteris et al. 2008; 7 Sect. 2.5.1, . Fig. 2.14). In bones and in dentine, crystals of bio-apatite attain lengths of 20–50 nm and thicknesses of 12–20 nm (1 nm = 12−6 mm), whereas in dental enamel, they are about 10 times longer and thicker. In contrast to inorganic hydroxylapatite and fluorapatite, formed by geological processes, bio-apatite displays a highly disordered crystal structure, has a non-stoichiometric chemical composition, and contains a high amount of [CO3OH]3−. Therefore, its Ca/P ratio is markedly higher than the theoretical value of 1.67. Other features characteristic of bio-apatite are a distinct deficiency in (OH) and vacant sites in the crystal structure. These properties, together with a small grain size, cause a high free surface energy. Therefore, bio-apatite is easily soluble and displays a high reactivity, e.g., with medicines. In human or animal tissues, crystals of bio-apatite can grow over short periods of days, weeks or months, initially with a highly disordered structure and a high amount of free [HPO4]2−-ions. More

ganic nature. Apatite and phosphorite are important raw materials for the chemical industry, especially for the production of synthetic phosphoric fertilisers, such as superphosphate, ammonium phosphate or “nitrophoska”, of phosphoric acid and elementary phosphor as well as a diversity of technical products (Rakovan and Pasteris 2015). For instance, natural or synthetic apatite as well as artificial products with apatite structure display a considerable microporosity, mainly caused by the channels parallel to c in the apatite structure (White et al. 2005). These render manifold processes of ion exchange that might be technically usable in the future, e.g., for phosphoric acid fuel cells, for photocatalysis or for storage of radioactive waste (Oelkers and Montel 2008).

z Pyromorphite

Pb5(PO4)3Cl Crystal form and habit  Crystal class 6/m; simple prismatic

crystals showing the hexagonal prism {101¯0} with barrelshaped curvature, combined with the basal pinacoid {0001} are common, whereas acicular crystals are rare. In most cases, pyromorphite crystals grow in groups on fissures or in druses, form globular or reniform aggregates, incrustations or effloresents. Physical properties Cleavage

none

Fracture

uneven, conchoidal

Hardness

3½–4

Density

7.04

Optical properties

translucent to subtransparent; lustre: adamantine on crystal faces, resinous to greasy on fractures; colour commonly green, by traces of Cu (“green lead ore”) or brown (“brown lead ore”), but also yellow, grey or colourless, more rarely orange red

Crystal structure  Isostructural with apatite. Chemical composition  In pyromorphite, [PO4] is partly replaced by [AsO4] and there is a complete solid solution with mimetite, Pb5(AsO4)3Cl. Moreover, Ca can replace Pb to a certain extent. Occurrence  Pyromorphite forms as a secondary mineral

in oxidation zones of Pb sulfide deposits.

155 Suggestions for Further Reading

z Mimetite

Pb5(AsO4)3Cl Crystal form and habit  Crystal class 6/m; crystal shapes similar to those of pyromorphite. Physical properties Cleavage

none

Fracture

uneven, conchoidal

Hardness

3½–4

Density

~7.1

Optical properties

translucent; lustre: adamantine on crystal faces, greasy on fractures; yellow, brown, green, less commonly grey to colourless

Crystal structure  Isostructural with apatite. Chemical composition  Complete solid solution with pyro-

morphite.

Occurrence  In the oxidation zone of ore deposits that

contain Pb and As minerals, such as galena, PbS, in association with tennantite, Cu12As4S13.

z Vanadinite

Pb5(VO4)3Cl Crystal form and habit  Crystal class 6/m; prismatic crystals

with dominant prism {101¯0}, in combination with basal pinacoid {0001} and hexagonal dipyramids {101¯1} and {213¯1},

can show columnar, rounded or barrel-shaped habit. Moreover, tabular crystals are found with dominant {0001} and subordinate {101¯0} (. Fig. 10.4). Vanidinite can also occur in grape-like or reniform aggregates or in sturdy masses. Physical properties Cleavage

none

Fracture

uneven, conchoidal, brittle

Hardness

3

Density

6.9

Optical properties

translucent to transparent; resinous to subadamantine lustre; ruby red (. Fig. 10.4), orange yellow, yellowish brown

. Fig. 10.4  Crystal group of vanadinite from Mibladen, Morocco; tabular habit with prominent basal pinacoid {0001} and small ¯ width of view is 1 cm hexagonal prism {1010};

Occurrence  Vanadinite occurs in the oxidation zone of Pb

deposits, associated with carbonate rocks, where it can be concentrated to form mineable deposits of V ore.

Economic relevance  Vanidinite is one of the major

sources of vanadium, used as an alloy metal to form special steels, such as ferrovanadium. Already small amounts of 0.15–0.25% V increase the strength of a high-carbon steel considerably, making it suitable for production of bicycle frames, axles, gears or crankshafts. High-speed tool steels (HSS) used for surgical instruments and tools contain 1–5% V. Titanium alloys with Al and V, such as 6Al–4V, characterised by high strength and high T-stability, are used in jet engines, high-speed air frames and dental implants. Various V compounds find a wide range of technical applications, e.g., in ceramic and glass industries, as catalysts, superconductors and superconducting magnets.

Crystal structure  Isostructural with apatite.

Suggestions for Further Reading

Chemical composition  [VO4] can be partly replaced by

Anthony JW, Bideaux RA, Bladh KW, Nichols MC (1997) Handbook of Mineralogy, vol IV: Arsenates, phosphates, vanadates. Mineralogical Society of America Boskey AL (2007) Mineralization of bones and teeth. Elements 3:385– 391

[PO4] and small amounts of [AsO4]. The theoretical metal contents of pure vanadinite are 73.2% Pb and 10.8% V.

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Chapter 10 · Phosphates, Arsenates, Vanadates

Chakhmouradian AR, Wall F (2012) Rare earth elements: minerals, mines, magnets (and more). Elements 8:333–340 Chang LL, Howie RA, Zussman J (1996) Rock-forming Minerals, vol 5B, 2nd edn. Non-silicates: sulphates, carbonates, phosphates, halides. Longmans, Harlow Essex Chew DM, Spikings RA (2015) Geochronology and thermochronology using apatite: time and temperature, lower crust to surface. Elements 11:189–194 Deer WA, Howie RA, Zussman J (2013) An Introduction to the Rock-forming Minerals 3rd edn. The Mineralogical Society, London Elliott JC (1994) Structures and Chemistry of Apatites and Other Calcium Orthophosphates. Elsevier, Amsterdam Hatch GP (2012) Dynamics of the global market for rare earths. Elements 8:341–346 Hughes JM, Rakovan JF (2015) Structurally robust, chemically diverse: apatite and apatite supergroup minerals. Elements 11:165–170 Kohn MJ, Rakovan J, Hughes JM (eds) (2002) Phosphates—geochemical, geobiological and materials importance. Rev Mineral Geochem 48 McCubbin FM, Jones RH (2015) Extraterrestrial apatite: planetary geochemistry to astrobiology. Elements 11:183–188 Nriagu JO, Moore PB (1984) Phosphate Minerals. Springer, Berlin, Heidelberg

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Oelkers EH, Montel J-M (2008) Phosphates and nuclear waste storage. Elements 4:113–116 Pasero M, Kampf AR, Ferraris C et al (2010) Nomenclature of the apatite supergroup minerals. Eur J Mineral 22:163–179 Pasteris JD, Wopenka B, Valsami-Jones E (2008) Bone and tooth mineralization: why apatite? Elements 4:94–104 Rakovan JF, Pasteris JD (2015) A technological gem: materials, medical, and environmental mineralogy of apatite. Elements 11:195–200 Simandl GJ (2014) Geology and market-dependent significance of rare earth element resources. Mineral Depos 49:889–904 Strunz H, Nickel EH (2001) Strunz Mineralogical Tables, 9th edn. Schweizerbart, Stuttgart U.S. Geological Survey (2020) Mineral commodity summaries 2020: U.S. Geological Survey, 200 p., 7 https://doi.org/10.3133/mcs2020 Valsami-Jones E, Oelkers EH (eds) (2008) Phosphates and global sustainability. Elements 4:83–116 Webster JD, Piccoli PM (2015) Magmatic apatite: a powerful, yet deceptive, mineral. Elements 11:177–182 White T, Ferraris C, Kim J, Madhavi S (2005) Apatite—an adaptive framework structure. In: Ferraris G, Merlino M (eds) Micro- and mesoporous mineral phases. Rev Mineral Geochem 57:307–401

157

Silicates 11.1  Orthosilicates (Nesosilicates) (. Table 11.1) – 160 11.2  Disilicates (Sorosilicates) – 169 11.2.1  Melilite Series (. Table 11.2) – 169

11.3  Ring Silicates (Cyclosilicates) – 172 11.4  Chain Silicates (Inosilicates) – 177 11.4.1  Pyroxene Family – 178 11.4.2  Pyroxenoids – 183 11.4.3  Amphibole Family – 184

11.5  Sheet Silicates (Phyllosilicates) – 189 11.5.1  Pyrophyllite-Talc Group – 192 11.5.2  Mica Group – 193 11.5.3  Hydromica Group – 195 11.5.4  Brittle-Mica Group – 195 11.5.5  Chlorite Series – 195 11.5.6  Serpentine Group – 196 11.5.7  Clay Minerals – 198 11.5.8  Apophyllite Group – 200

11.6  Framework Silicates (Tectosilicates) – 201 11.6.1  SiO2 Minerals – 201 11.6.2  Feldspar Family (. Table 11.11) – 216 11.6.3  Feldspathoids (Foids) – 228 11.6.4  Cancrinite Group – 231 11.6.5  Scapolite Group – 231 11.6.6  Zeolite Family – 232

References and Suggestions for Further Reading – 237

© Springer-Verlag GmbH Germany, part of Springer Nature 2020 M. Okrusch, H. E. Frimmel, Mineralogy, Springer Textbooks in Earth Sciences, Geography and Environment, https://doi.org/10.1007/978-3-662-57316-7_11

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Chapter 11 · Silicates

Introduction Silicate minerals, including quartz, play a pivotal role in the make-up of planet Earth. From direct observation as well as from geophysical and petrological modelling it is clear that silicates form the principal constituents of the Earth’s crust and mantle, which together constitute about 67.3 wt% of our planet. There is also plenty of direct and indirect evidence that the Moon, the terrestrial planets and the asteroids are predominantly of silicatic composition. Moreover, silicate minerals and quartz are of great technical and economic importance.

Fundamental Structural Principles and Classification of Silicates Silicates have certain structural principles in common, which form the basis of their classification (. Fig. 11.1).

11

. Fig. 11.1  The fundamental principles of the silicate structures: a Orthosilicates; b disilicates; c−e ring silicates: c 3-membered ring; d 4-membered ring; e 6-membered ring; f, g chain silicates: f single chain; g double chain; h layered silicates; i framework silicates: sodalite cage as example

1. In silicate structures, silicon, Si4+, is always linked to 4 oxygens in tetrahedral coordination. This fact is valid independent of the respective Si:O ratios as expressed by the mineral formulae reflecting different anionic complexes, such as SiO3, SiO4, SiO5, Si2O7, Si3O8, Si4O10 or Si4O11. The 4 oxygens take the corners of a nearly ideal tetrahedron and, owing to their large size of 1.27 Å, touch each other, thus leaving only a tiny space in the centre of the tetrahedra for the small Si4+ with an ionic radius of 0.34 Å. In fact, the outer electron shells of the Si and O atoms overlap, because the small, but highly charged silicon is able to polarise the large oxygen atoms. This leads to a polar covalent bond type, intermediate in character between atomic and ionic bonding. The strong bonds within the SiO4 tetrahedron correspond to the four sp3 hybrid orbitals (see . Fig. 1.15). 2. A characteristic property of silicate structures is the fact that a given oxygen atom can belong to two different [SiO4]

159 Silicates

tetrahedra. Thereby, a variety of anionic building blocks can be formed in addition to isolated [SiO4] tetraehedra (. Fig. 11.1): Double-tetrahedra [Si2O7]6–; circular groups of different composition, such as 3-membered rings [Si3O9]6–, 4-membered rings [Si4O12]8– and 6-membered rings [Si6O18]12–; infinite, one-dimensional chains [Si2O6]4– and double chains [Si4O11]6–; infinite, two-dimensional layers [Si4O10]4– and infinite, three-dimensional frameworks. 3. Another important crystal-chemical principle is the double role played by trivalent aluminium in silicate structures. Al3+ can be linked either to six oxygens, i.e., in [6]-coordination, thereby attaining an ionic radius of 0.61 Å, or to four oxygens, i.e., in [4]-coordination leading to an ionic radius of 0.47 Å. Thus, in alumosilicates, Al3+ replaces some or all Si4+ in the tetrahedral site, forming an [AlO4] tetrahedron. On the other hand, in aluminosilicates, Al3+ occupies also the slightly larger octahedral site with 6 oxygens as nearest neighbours, and substituting for cations of similar ionic radius such as Mg2+ (0.80 Å), Fe2+ (0.69 Å) or Fe3+ (0.63 Å) etc. When ions with different valency are substituted for each other, such as Si4+ ⇋ Al3+ or Mg2+ ⇋ Al3+ , charge balance is achieved by coupled substitution (coupled valence equalisation). Prominent examples are the plagioclase solid-solution series with the end-members albite, Na[AlSi3O8], and anorthite, Ca[Al2Si2O8], in which the charge balance is retained by the coupled substitution Na+ Si4+ ⇋ Ca2+ Al3+, or the various solid solution series in pyroxenes, e.g., between diopside, CaMg[Si2O6], and jadeite, NaAl[Si2O6], with the coupled substitution Ca2+ Mg2+ ⇋ Na+ Al3+ . In silicate structures, Si4+ can be substituted by Al3+ only to a maximum Al[4]:Si[4] ratio of 1:1, which means that there is no transition from alumosilicates to aluminates. Without the knowledge of this double role of aluminium, a sensible classification and subdivision of the silicate minerals was not possible and many minerals could not even be described by a suitable chemical formula. The issue was complicated by the wealth of different solid-solution series among silicate minerals, leading to apparently unintelligible chemical compositions. Thus, the initial approach to describe the silicate minerals as salts of different silicic acids proved futile. Only after the advent of X-ray diffraction methods, crystal structures of the most important silicate minerals were determined, thus providing deeper insights into their internal crystal-chemical arrangements and their chemical affinities. First suggestions for a modern classification of silicate minerals, made by William L. Bragg (1890–1971) and Felix Machatschki (1895–1970) at the end of the 1920s, still form the basis of today’s understanding of the crystal chemistry of silicates. Current classification of silicate minerals is based on increasing polymerisation of the Si–O complex and the mutual linkage of [SiO4] tetrahedra. For clarity, the anionic building blocks are set in square brackets. The following structural types are distinguished (. Fig. 11.1): 5 Orthosilicates (or nesosilicates) with isolated [SiO4]4– tetrahedra (7 Sect. 11.1); examples are forsterite Mg2[SiO4]; olivine, (Mg,Fe)2[SiO4]; zircon, Zr2[SiO4]. In some of the orthosilicates. additional anions, such as F– or (OH)– are present, an example being topaz, Al2[SiO4](F,OH)2. 5 Disilicates (or sorosilicates) with finite groups, essentially double-tetrahedra of the composition [Si2O7]6–, in which two [SiO4] tetrahedra are connected via one common oxygen; as this bridging oxygen belongs by half to each of

the two tetrahedra, the Si:O ratio is 2:7. Example: melilite with the end-members åkermanite, Ca2Mg[Si2O7], and gehlenite Ca2Al[6][Al[4]SiO7] (7 Sect. 11.2). 5 Ring silicates (or cyclosilicates) with separate, closed 3-, 4or 6-membered rings of [SiO4] tetrahedra; in these circles each [SiO4] tetrahedron shares 2 of its 4 oxygens with the 2 adjacent [SiO4] tetrahedra and, consequently, the following anion complexes result for the 3-, 4- or 6-membered rings, respectively: [Si3O9]6–, [Si4O12]8– and [Si6O18]12–. An example is schorl, a common end-member of the tourmaline group, NaFe32+Al6[Si6O18](BO3)3 (OH)3(OH,F) (7 Sect. 11.3). 5 Chain silicates (or inosilicates) with one-dimensional, infinite chains or double chains of [SiO4] tetrahedra; in single chains, each [SiO4] tetrahedron shares 2 of its 4 oxygens with the neighbouring tetrahedron in the direction of the chain and, therefore, the Si:O ratio is 1:3 (like in the cyclosilicates). The most important examples are the minerals of the pyroxene family (7 Sect. 11.4.1), in which the one-dimensional linkage of [SiO4] tetrahedra leads to a basic anionic group of the composition Si2O64–. Examples are orthopyroxene, (Mg,Fe)[6]2[Si2O6], and diopside, Ca[8]Mg[6][Si2O6]. The infinite double chains are formed by connecting two simple chains of [SiO4] tetrahedra by one bridging oxygen. Therefore, compared to the single chain, every second tetrahedron shares an oxygen with a tetrahedron of the neighbouring chain, which leads to an anion complex of [Si4O11]6–. The double chain contains cavities that can be filled with additional anions, such as (OH)− or F−. Prominent examples are members of the amphibole family (7 Sect. 11.4.3), such as anthophyllite, (Mg,Fe)[6]7[Si8O22] (OH)2, or tremolite, Ca[8]2(Mg,Fe)[6]5[Si8O22](OH,F)2. 5 Sheet silicates (or phyllosilicates) are composed of two-dimensional, infinite layers of [SiO4] tetrahedra, in which each [SiO4] tetrahedron shares 3 bridging oxygens with the 3 adjacent tetrahedra. This leads to a Si:O ratio of 2:5, and the basic anionic group has the composition [Si2O5]2– or [Si4O10]4–. The silicatic layers, too, contain cavities that can be filled with additional anions, mainly (OH)– or F–. Examples described in 7 Sect. 11.5 are: 5 pyrophyllite, Al2[Si4O10](OH)2, 5 talc, Mg3[Si4O10](OH)2, 5 muscovite, K+Al[6]2[Al[4]Si3O10](OH)2, 5 phlogopite, K+Mg3[Al[4]Si3O10](OH,F)2. In the micas muscovite and phlogopite, ¼ of Si[4] sites are occupied by Al[4]. Therefore, the anionic complex attains a single negative charge that is compensated by incorporation of monovalent cations such as K+ or Na+. The formula of muscovite is derived from the pyrophyllite formula, that of phlogopite from the talc formula. 5 In framework silicates (or tectosilicates) all the [SiO4] tetrahedra are connected to each other over all their four corners, forming a three-dimensional framework. Thus only 4/2 oxygens belong to each Si, leading to SiO2, the formula of the silicon dioxide quartz, an electrostatically neutral structure. Therefore, real tectosilicates are only possible if Si4+ is partly substituted by Al3+, whereby the structure attains one or more negative charges that can be compensated by incorporation of one or more cations. As the three-dimensional framework contains a lot of open spaces, large cations such as K+, Na+, Ca2+ etc., can be accommodated in these spaces and alumosilicates, such as feldspars (7 Sect. 11.6.2) and feldspathoids (7 Sect. 11.6.3), are formed.

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In some cases, large, additional anions, such as Cl–, SO42– etc., or H2O molecules are incorporated in the open spaces within the framework. The H2O molecules are connected with the silicate structure by extremely weak bonds. Consequently, they can easily escape upon heating without destruction of the structure that, in reverse, again takes up

11.1  Orthosilicates (Nesosilicates)

(. Table 11.1)

. Table 11.1  Important orthosilicates Mineral

z Olivine (Mg,Fe)2[SiO4] Crystal form and habit  Crystal class 2/m2/m2/m; the ort-

11

water in an atmosphere saturated which H2O steam. These H2O-rich tectosilicates belong to the large, technically important zeolite family (7 Sect. 11.6.6). As a result of the lose packing in their structures, tectosilicate minerals are characterised by low densities and relatively low refractive indices and birefringence.

horhombic-dipyramidal crystals commonly show vertical prisms {110} and {120}, combined with the prisms {021} and {101}, the dipyramid {111} and the pinacoid {010} (. Fig. 11.2). Euhedral or subhedral phenocrysts commonly are present in volcanic rocks. In magmatic peridotite and its metamorphosed equivalents, anhedral olivine occurs as granular aggregates. These aggregates can be found as so-called olivine nodules (. Figs. 2.8, 13.5b) in the form of xenoliths in basaltic rocks. Anhedral olivine is also present in some siliceous marbles and high-grade metamorphic ultramafic rocks.

Formula

Crystal class 2/m2/m2/m

Olivine

(Mg,Fe)2[SiO4]

   Forsterite

Mg2[SiO4]

   Fayalite

Fe2[SiO4]

Zircon

Zr[SiO4]

4/m2/m2/m

Garnet group

X32+Y23+[SiO4]3

¯ 4/m32/m

   Sillimanite

Al[6]Al[4][O/SiO4]

2/m2/m2/m

   Andalusite

Al[6]Al[5][O/SiO

2/m2/m2/m

   Kyanite

Al[6]Al[6][O/SiO4]



Topaz

Al2[SiO4](F,OH)2

2/m2/m2/m

Staurolite

Fe2Al9[(SiO4)4]O6(O,OH)2

2/m

Chloritoid

(Fe,Mg,Mn)Al2[SiO4]O(OH)2

1¯ and 2/m

Titanite

CaTi[SiO4](O,OH,F)

2/m

Al2SiO5 group 4]

Physical properties Cleavage

{010}, distinct to imperfect, {100}, weak

Fracture

conchoidal

Hardness

6½–7

Density

3.222 for pure forsterite, 4.392 for pure fayalite

Optical properties

transparent to translucent; lustre: vitrous on crystal faces, greasy on fractures; olive green, also yellowish brown to reddish brown, depending on fayalite content

Crystal structure  The olivine structure can be described as

nearly hexagonal close packing of oxygen parallel to (100) with the sequence 121212 … (. Fig. 11.3). Si is situated in

. Fig. 11.2  Crystal forms of olivine; orthorhombic dipyramid {111}, dominant prisms {110} and {021}, additional prisms {120} and {101}, pinacoid {010}

. Fig. 11.3  Crystal structure of the olivine end-member forsterite, projected along the a-axis onto the (100) plane. Between the isolated [SiO4] tetrahedra (Si atoms not shown), the Mg[6] cations occupy the octahedral sites, i.e. each Mg is linked to 6 oxygens as nearest neighbours. The oxygen atoms form a nearly hexagonal close packing (after Bragg and Bragg, from Evans 1976)

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11.1 · Orthosilicates (Nesosilicates) (Table 11.1)

the small tetrahedral sites between 4 O, whereas the cations Mg2+ and Fe2+ take the somewhat larger, octahedral sites between 6 O as nearest neighbours. At very high pressures of ≥70 kbar, the olivine structure is transformed into the even closer packed crystal structures of wadsleyite, β-(Mg,Fe)2[SiO4], and ringwoodite, γ -(Mg,Fe)2[SiO4], which are similar to the spinel structure (. Fig. 7.2). Presumably, these minerals are the major constituents of the transition zone between the Earth’s upper and lower mantle (7 Sect. 29.3.3). Chemical composition  Olivine forms a complete solid

solution series with the end-members forsterite, Mg2[SiO4], and fayalite, Fe2[SiO4] (. Fig. 18.14). In common rock-forming olivine is typically magnesian with the forsterite component ranging between 90 and 70 mol%. Characteristically, olivine contains minor amounts of Ni2+ replacing Mg2+, and also of Mn2+, replacing Fe2+, the latter substitution being mainly realised in fayalite-rich olivines.

Occurrence  Olivine is an important rock-forming mineral, especially in ultramafic rocks (. Fig. 13.5b), and also occurs in basalts (. Fig. 13.8a), forming zoned phe-

nocrysts with Mg-rich cores. Most importantly, olivine is the main constituent in rocks of the Earth’s upper mantle (7 Sect. 29.3.1). It also occurs in meteorites, especially in chondrites (7 Sect. 31.3.1). In the presence of (hydrothermal) water, olivine alters to serpentine minerals under uptake of H2O (. Fig. 26.13b). Brownish iddingsite is a common alteration product of olivine and comprises very fine-grained aggregates of montmorillonite, chlorite, goethite, haematite etc. (. Fig. 13.11b).

Economic relevance  Dunite, a nearly monomineralic ultramafic rock consisting essentially of forsterite-rich olivine, is an important raw material for production of highly refractive dunite bricks. Transparent, olive-green crystals of olivine, known as chrysolite or peridote, are highly valued gemstones.

z Zircon Zr[SiO4] Crystal form and habit  Crystal class 4/m2/m2/m; crys-

tal form and habit strongly depend on the conditions of crystallisation. Well-developed crystals tend to be prismatic, showing simple combinations of the tetragonal prisms {100} and/or {110} with the tetragonal dipyramid {101}, but additional faces, e.g., the tetragonal dipyramid {112} and/or the ditetragonal dipyramid {211} may be present as well (. Fig. 11.4). Crystals with pyramidal habit show either only {101} or in combination with {301} (. Fig. 11.5). Frequently, zircon crystals show extensive zonation, visible under the microscope (. Figs. 11.6,

. Fig. 11.4  Crystal forms of zircon: tetragonal prisms {110} and {100}, ditetragonal dipyramid {211}, tetragonal dipyramids {101} and {112}

33.16b). In clastic sediments, zircon can occur as rounded

detrital grains.

Physical properties Cleavage

{110}, imperfect, {111} poor

Fracture

conchoidal

Hardness



Density

4.6–4.7, i.e., relatively high

Optical properties

non-transparent to translucent, rarely transparent; adamantine shows resinous or greasy lustre; commonly brown, also colourless, yellow, orange red, more rarely green

Crystal structure  Analogous to monazite, isolated [SiO4] tetrahedra are connected, via corners and edges, with zig-zag chains of edge-sharing [ZrO8] polyhedra, thereby forming a three-dimensional framework (. Fig. 11.7). Radiation emitted from traces of Th and U, which substitute for Zr, can cause variable degrees of radiation damage up to total destruction of the crystal structure. This results in a so-called metamict state, which is marked by lower density, hardness and double refraction (. Fig. 11.6). Chemical composition  To a certain extent, Zr can be

replaced by Hf, Th and U. In addition, zircon contains a wide range of trace elements, such as REE and P. The element hafnium, Hf, was first detected in zircon from Norway by Coster and Hevesy (1923).

Occurrence  Zircon is a common accessory mineral in

many igneous and metamorphic rocks, especially those of felsic mineralogy, and it can reach larger grain size in nepheline syenite and related pegmatite (. Fig. 11.5; 7 Sect. 22.3), in which it can be concentrated to economic grade. Furthermore, zircon is found as a heavy mineral in clastic sedimentary rocks, especially in sands and sandstones as well as in placer deposits. The so-called pleochroic

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Chapter 11 · Silicates

11 . Fig. 11.5  Zircon crystal with two different tetragonal dipyramids {101} and {301} from a pegmatite, Hunza Valley, Cashmere; width of view is c. 2 cm

haloes around tiny crystals included in coloured minerals, predominantly biotite, result from radiation emitted by Th and U.

. Fig. 11.6  Photomicrograph of a zircon crystal from a leucogranite, near Dannemora, Adirondack Mountains, State New York, U.S.A; section //c; length of the crystal is 360 μm (=0.36 mm); thickness of the thin section is 30 μm; crossed polars (+Nic.); Most of the crystal shows primary zonation and underwent moderate radiation damage that caused a distinct reduction of the double refraction, visible by interference colours of 2nd order; In contrast, the rounded, uranium-poor core mainly shows the high interference colour pink of 3rd order, typical of zircon without considerable structural defects (Photograph: Lutz Nasdala, Vienna, from Nasdala et al. 2005)

Economic relevance  Zircon is an important raw material

for the production of the elements Zr and Hf as well as Zr compounds (e.g., Watson 2007). Zirconium metal is used for cladding of fuel rods in nuclear reactors and forms a component of Fe–Zr and Zr–Nb alloys, the latter being used in supraconductors. Glasses of Zr (and Hf) fluorides display an extremely high infrared (IR) permeability and are therefore used in optical fibres. At a very high temperature of 1660 °C, zircon decomposes into ZrO2, known as zirconia, and SiO2. As ZrO2 has an extremely high melting point of 2700 °C, zirconia crucibles and slip-cast bricks consisting of polycrystalline zircon are mechanically resistant, acid-proof highly refractory materials. Porous, ZrO2-based ceramics are excellent thermal insolators used for furnace linings. Containers of zirconia can serve for melting of high-temperature glasses or metals, such as platinum. Tooth crowns, dental implants and artificial hip joints are made of ZrO2. Y-stabilised ZrO2 is employed in fuel cells and as ionic conductor in lambda oxygen sensors for measuring O2 fugacities (c.f.

. Fig. 11.7  Crystal structure of zircon, projected onto the (100) plane; Edge-sharing [ZrO8] polyhedra (yellow) form zig-zag chains and are connected, via corners and edges, with isolated [SiO4] tetrahedra to constitute a three-dimensional framework (modified after Zoltai and Stout 1984)

163

11.1 · Orthosilicates (Nesosilicates) (Table 11.1)

7 Sect. 26.1.4). Other Zr compounds are used in glazes in ceramics and glasses. Transparent, beautifully coloured zircon grains are highly esteemed gemstones, such as the green variety or the brownish to reddish-orange variety known as hyazinte. The intense blue colour of many facetted zircon crystals is, however, in most cases artificial and caused by thermal treatment. At present, the by far most important producers of zircon are Australia (39% of global production, totaling 1.4 Mt) and South Africa (26%), followed by USA (7%), China (6%), and Senegal (5%) (U.S. Geological Survey 2020). Geochronology  Based on its contents in U and Th, zircon has been used, for a long time, for radiometric age determinations of igneous and metamorphic rocks, especially with the uranium-lead (U-Pb) method (7 Sect. 33.5.3, . Fig. 33.15). U-Pb dating of detrital zircon grains in sedimentary rocks can provide a maximum age limit of sedimentation and important information on the sediment provenance. Thus it can give useful hints for plate tectonic reconstructions of old cratons and orogens. An essential methodical break-through was achieved by the in-situ isotope analysis of isolated single zircon grains, even allowing for the age determination of different growth stages or generations within a single grain (. Fig. 33.16b, c; cf. Harley and Kelly 2007).

¯ Crystal form and habit  Crystal class 4/m32/m ; well-de-

veloped crystals show the rhomb-dodecahedron {110} or, less commonly, the trapezohedron {211}, either alone or in combination (. Figs. 11.8, 2.10), whereas combinations with {hkl} faces are rare. Commonly rock-forming garnet is intergrown with associated minerals, thus occurring as subhedral or chubby grains or aggregates. In many cases, garnet is zoned. Physical properties Cleavage

occasionally, parting on {110} is indicated

Fracture

conchoidal, brittle

Hardness

6–7½, depending on the chemical composition

Density

ranging from 3.58 for pure pyrope to 4.32 for pure almandine

Optical properties

coulour varies with composition. Pyroperich garnet is deep red, almandine-rich garnet is brownish-red, spessartine-rich garnet yellowish to brownish red, grossular-rich garnet light to yellowish-green or brownish to reddish yellow. A red grossular variety is known as hessonite (. Fig. 11.32), a green one, coloured by traces of Cr3+ and V3+ as tsavorite. Andradite-rich garnet is brownish to black, the varieties topazolite and demantoid are yellowish-green. Melanite, a Ti-bearing andradite, is deep black in colour, in thin section it is dark brown and translucent. Uvarovite-rich garnet is dark emerald-green

Lustre

commonly vitreous to greasy; demantoid shows adamantine lustre

z Garnet Group X2+3Y3+2[SiO4]3

In the structural formula of natural garnets, the X2+ and Y3+ positions are occupied by the following cations:

X2+ = Mg2+ , Fe2+ , Mn2+ , Ca2+ Y3+ = Al[6] , Fe3+ , Cr3+ , V3+ End-members of the so-called pyralspite series are 5 pyrope Mg3Al2[SiO4]3 5 almandine Fe3Al2[SiO4]3 5 spessartine Mn3Al2[SiO4]3 End-members of the so-called ugrandite series are 5 uvarovite Ca3Cr2[SiO4]3 5 grossular Ca3Al2[SiO4]3 5 andradite Ca3Fe23+[SiO4]3 In addition, numerous end-members of the garnet group have been synthesised, most of which play a very limited role in nature but can be of high technical relevance.

Crystal structure  Alternating, corner-sharing [SiO4] tetra-

hedra and [Y3+O6] octahedra form zig-zag chains that are parallel to the three edges of the cubic unit cell, thereby forming a three-dimensional framework with distorted hexahedral cavities. These are filled by the [8]-coordinated X2+ cations (. Fig. 11.9).

Chemical composition  Within

the pyralspite series, complete solid solutions exist between the end-members almandine—pyrope and almandine—spessartine, within the ugrandite series between grossular and andradite. Pyralspite garnets can contain up to 30 mol% of grossular + andradite. In melanite, the chemical valence is adjusted by the coupled substitutions 2AI3+[6] ⇌ Ti4+[6]Fe2+[6] or AI3+[6]Si4+[4] ⇌ Ti4+[6]Fe2+[4]. In this case,

11

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Chapter 11 · Silicates

. Fig. 11.8  Crystal forms and habits of garnet with a simple rhomb-dodecahedron {110}; d trapezohedron {211}; b, c combinations of both, b {110} dominant; c {211} dominant

extracted from placer deposits at Podsedice, Czech Republic, and the so-called Cape ruby, a pyrope-rich by-product from the diamond mines of South Africa. Much rarer is the yellowish green demantoid, valued for its adamantine lustre. The same holds true for the green grossular variety tsavorite that was discovered, in 1967, in marbles of the Neoproterozoic Mozambique Belt at the border between Tanzania and Kenya, and named after the Tsavo National Park. The orange-coloured spessartine from Ramona, San Diego County, California, is also cut as a gemstone (Rossman 2009). Synthetic non-silicates with garnet structure display special magnetic, optical, lasing and ion-conducting properties that give rise to a wide variety of technical applications (e.g., Baxter et al. 2013; Geiger 2013).

11 . Fig. 11.9  Garnet structure, parallel to (100); [SiO4] tetrahedra (blue), sharing corners with [AlO6] octahedra (yellow); both structural elements share corners and edges with distorted XO8 hexahedra (green); oxygens are shown as pink spheres (modified after Zoltai and Stout 1984)

a replacement of Si4+[4] by Fe2+[4] is possible to a limited extent. Occurrence  Garnets are important rock-forming miner-

als that are common constituents in a variety of metamorphic rocks in the Earth’s crust and in garnet peridotite of the Earth’s upper mantle. Field evidence and experimental investigations clearly indicate that the formation of important members of the garnet group is favoured by high to very high pressures. This holds true especially for pyroperich garnets formed under P-T conditions of the Earth’s upper mantle (7 Sect. 29.3.1). In contrast, continental plutonic and volcanic rocks only rarely contain garnet. For instance, melanite predominantly occurs in alkaline magmatic rocks, whereas topazolite is confined to hydrothermal veins. Moreover, garnet can be found as heavy mineral in sands and sandstones or concentrated in placer deposits. Economic relevance  Clear, beautifully coloured garnets are sought after gemstones (e.g., Galoisy 2013). The most prominent examples are the pyrope-rich Bohemian garnets

z Al2SiO5 Group

The crystal structures of the three Al2SiO5 polymorphs sillimanite, Al[6]Al[4][O/SiO4], andalusite, Al[6]Al[5][O/ SiO4], both orthorhombic, and kyanite, Al[6]Al[6][O/SiO4], triclinic, contain chains of edge-sharing [AlO6] octahedra as a common structural element. However, the kind of linkage of these chains and, consequently, the coordination numbers of the second Al in the formula are different. In the sillimanite structure, the chains are connected, via corners, by isolated alternating [AlO4] and [SiO4] tetrahedra (. Fig. 11.10a), whereas in andalusite the linkage is achieved by pairs of edge-sharing [AlO5] polyhedra alternating with pairs of isolated [SiO4] tetrahedra (. Fig. 11.10b). A closer packing is present in the kyanite structure, in which two chains of edge-sharing [AlO6] octahedra are connected, via edges, to form bands parallel to c. Neighbouring bands are linked with each other by [SiO4] tetrahedra (. Fig. 11.10c). By these structural differences, the cleavages {010} of sillimanite, {110} of andalusite and {100} and {010} of kyanite can be explained as well as the distinct anisotropy of hardness typical of kyanite (. Fig. 1.20). The phase relations of the three Al2SiO5 polymorphs are depicted in a P-T diagram (. Fig. 27.2). Andalusite, the polymorph with the lowest density, is restricted to

165

11.1 · Orthosilicates (Nesosilicates) (Table 11.1)

. Fig. 11.10  Al-silicate structures, projected onto (100): a The sillimanite structure consists of chains //c, formed by edge-sharing [AlO6] octahedra (yellow) that are connected, via corners, with alternating [AlO4] tetrahedra (green) and [SiO4] tetrahedra (blue); b The andalusite structure is also formed by chains of edge-sharing [AlO6] octahedra (yellow) that are alternately connected, via corners, with pairs of edge-sharing AlO5 polyhedra (green) and isolated [SiO4] tetrahedra (blue); c In contrast, the kyanite structure, displaying a much closer packing, is formed by bands of edge-sharing [AlO6] octahedra (yellow). Laterally attached [SiO4] tetrahedra (blue) establish connections to the neighbouring [AlO6] band (after Papike 1987, from Kerrick 1990)

the lowest pressure range. With increasing pressure it transforms to kyanite at lower, and to sillimanite at higher temperatures. Sillimanite transforms into kyanite at very high pressures. At one distinct P-T condition only, all three Al2SiO5 polymorphs can stably coexist with each other. This invariant triple point was experimentally determined at ~4 kbar and ~520 °C (e.g., Bohlen et al. 1991; Holdaway and Mukhopadhyay 1993). Based on their stability relations, the Al-silicates provide important clues as to the pressure-temperature conditions at which a metamorphic rock was formed.

z Sillimanite Al[6]Al[4][SiO5] Crystal form and habit  Crystal class 2/m2/m2/m; long

prismatic to acicular habit in metamorphic rocks; as fibrolite in fibrous bunches, matted aggregates or knots. Physical properties Cleavage

{010}, perfect, with transverse parting ⊥c

Fracture

splintery

Hardness

6½–7½

Density

3.23–3.27

Optical properties

transparent to translucent; lustre: vitreous; silky in fibrous aggregates; white, yellowish white, grey, brownish or greenish

Chemical composition  Some Al is commonly replaced by

small amounts of Fe3+.

Occurrence  Characteristic constituent of aluminous, that

is originally clay-rich, metamorphosed sedimentary rocks, so-called metapelites, such as micaschists, paragneisses and metapelitic granulites, formed at temperatures above ~520 °C.

z Andalusite Al[6]Al[5][SiO5] Crystal form and habit  Crystal class 2/m2/m2/m; euhedral, columnar crystals showing dominant orthorhombic prism {110} and basal pinacoid {001}, are common. Additional prisms {101} and {011} can be present. Nearly square cross sections perpendicular to c can be easily recognised, particularly under the microscope. The variety chiastolite is distinguished by the concentration of carbonaceous pigment in the centre and specific sectors of the crystal, which leads to the appearance of a dark cross in sections normal to {001}; Moreover, rock-forming andalusite can occur in radial or granular aggregates. Physical properties Cleavage

{110}, occasionally distinct

Fracture

uneven, conchoidal

Hardness

6½–7½

Density

3.15

Optical properties

transparent to non-transparent; vitreous lustre; grey, reddish, dark pink or brownish

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Chapter 11 · Silicates

Chemical composition  Andalusite can contain considerable amounts of Fe3+ and Mn3+. Viridine is a Mn-rich andalusite with up to 19.6 wt% Mn2O3 and 4.8 wt% Fe2O3. Occurrence  Rock-forming andalusite is stable at pressures

 Y[6], and thus have monoclinic symmetry. In orthopyroxenes, however, the cations are nearly similar in size and are both [6]-coordinated, X[6] ∼ = Y[6], leading to higher, orthorhombic, symmetry. The increase in symmetry is achieved by submicroscopic twinning parallel

to (100) with duplication of the unit cell. As a first approximation, many clinopyroxenes can be regarded as members of the 4-component system CaMgSi2O6– CaFeSi2O6–Mg2Si2O6–Fe2Si2O6, depicted in the pyroxene quadrilateral (. Fig. 11.29a, b). Monoclinic Mg–Fe pyroxenes of the solid-solution series Mg2[Si2O6]– Fe2[Si2O6] (clinoenstatite-clinoferrosilite) are unusual in terrestrial rocks, and the orthorhombic high-T polymorph protoenstatite is unknown in nature. The Ca-poor clinopyroxene pigeonite is stable only at low pressures, as demonstrated in the experimentally determined, pseudo-binary phase diagram protoenstatite-diopside at 1 bar pressure (. Fig. 18.16b). The Ca2[Si2O6] end-member forms the chain silicate wollastonite that structurally differs from the pyroxenes and, therefore, has been classified as pyroxenoid (7 Sect. 11.4.2). The crystal structures of the pyroxene family are characterised by two types of infinite single chains parallel to c (. Figs. 11.1f, 11.30). These consist of corner-sharing [SiO4] tetrahedra that result in anion complexes of [SiO3]2– or [Si2O6]4– and are connected, via corners, with edge-sharing [Y2+[6]O6)4– or [Y3+[6]O6]3– octahedra at the M1 site, e.g., with Y[6]=Mg2+ in diopside or Y[6]=Al3+ in jadeite. In clinopyroxenes, the larger X cations Ca2+ and Na+ are [8]-coordinated. They fill the large M2 gaps in the crystal structure and have somewhat weaker bonds with oxygen. In orthopyroxenes the M1 and M2 sites are similar in size and, therefore, are occupied by the divalent cations Mg2+ and Fe2+, both in [6]-coordination with oxygen, thereby forming octahedra or deformed (pseudo-) octahedra on the M1 and M2 sites, respectively.

The Subcommittee on Pyroxenes of the Commission on New Minerals and Mineral Names (CNMMN) of the International Mineralogical Association (IMA) proposed to abandon the names bronzite, hypersthene and ferrohypersthene, used in the past to describe members of the orthopyroxene solid solution series enstatite–ferrosilite (Morimoto et al. 1988). As these traditional names are commonly used in rock names (e.g., bronzitite) and in meteorite nomenclature (e.g., hypersthene chondrite), we decided to still list these orthopyroxene members in . Table 11.5 as well as the members salite and ferrosalite of the diopside–hedenbergite solid solution.

The positions of the most important pyroxenes in the pyroxene quadrilateral diopside (Di)—hedenbergite (Hd)— enstatite (En)—ferrosilite (Fs) are shown in . Fig. 11.29a, b, whereas the solid solutions between the end-members jadeite (Jd), aegirine (Aeg) and the Ca–Mg–Fe pyroxenes (Quad = Wo + En + Fs) are depicted in . Fig. 11.29c. The important pyroxenes are listed in . Table 11.5.

179

11.4 · Chain Silicates (Inosilicates)

crystals are rare. In most cases rock-forming Mg–Fe pyroxenes occur as massive, granular or laminated aggregates. Physical properties Cleavage

distinct along the vertical prism {210}, with cleavage angle ~88°, typical of all pyroxenes, i.e.,  to the weakest bonds along the [SiO3] chains (. Fig. 11.27); Commonly, translation on {100} leads to parting, kinking or wave-like bending along this face

Hardness

5–6

Density

3.21–3.96, increasing with Fe content

Optical properties

translucent to non-transparent; colourless to greyish green (enstatite) to dark brown or green (hypersthene); lustre: dull on cleavage planes {210}; on parting plane {100} bronzite shows bronze-like, hypersthene copper red schiller, due to exsolution of minute ilmenite platelets, oriented on this plane

Chemical composition  The orthopyroxenes form a complete solid-solution series between the end-members enstatite, Mg2[Si2O6], and ferrosilite, Fe2[Si2O6] (. Fig. 11.29b), ranging between nearly En100 to En10Fs90. So far, pure

. Fig. 11.29  Classification of a Ca–Mg–Fe clinopyroxenes, b orthopyroxenes and c solid solutions between jadeite (Jd), aegirine (Aeg) and Ca–Mg–Fe clinopyroxenes (Quad) after Morimoto et al. (1988); jadeite and omphacite typically occur in high-pressure metamorphic rocks (Sects. 26.3.1, 28.3.8, 28.3.9)

Mg–Fe Pyroxenes z Enstatite Mg2[Si2O6] z Ferrosilite Fe2[Si2O6] form and habit Predominantly crystal class 2/m2/m2/m (orthopyroxenes), whereas monoclinic Fe–Mg pyroxenes of crystal class 2/m are uncommon. Well-developed

Crystal

. Fig. 11.30  Crystal structure of jadeite, projected approximately in the direction of the a-axis. Corner-sharing [SiO4] tetrahedra (blue) and edge-sharing [AlO6] octahedra (M1 sites, yellow) form infinite chains //c that are connected with each other via corners. The large cations Na+ (pink spheres) are positioned in the more spaceous M2 sites of the structure (modified after Burnham et al. 1967)

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Chapter 11 · Silicates

. Table 11.5  Important pyroxenes Minerals

Formula

Solid-solution

Crystal class

Enstatite (En)

Mg2[Si2O6]

En100Fs0–En90Fs10

2/m2/m2/m

Bronzite

(Mg,Fe)2[Si2O6]

En90Fs10–En70Fs30

2/m2/m2/m

Hypersthene

(Mg,Fe)2[Si2O6]

En70Fs30–En50Fs50

2/m2/m2/m

Ferrohypersthene

(Fe,Mg)2[Si2O6]

En50Fs50–En30Fs70

2/m2/m2/m

Ferrosilite (Fs)

Fe2[Si2O6]

2/m2/m2/m

Pigeonite

c. Ca0.25(Mg,Fe)1.75[Si2O6]

Mg–Fe pyroxenes

2/m

Ca pyroxenes Diopside (Di)

CaMg[Si2O6]

Di100Hd0–Di90Hd10

2/m

Salite

Ca(Mg,Fe)[Si2O6]

Di90Hd10–Di50Hd50

2/m

Ferrosalite

Ca(Fe,Mg)[Si2O6]

Di50Hd50–Di10Hd90

2/m

Hedenbergite (Hd)

CaFe[Si2O6]

Di10Hd90–Di0Hd100

2/m

Na pyroxenes Jadeite

NaAl[Si2O6]

2/m

Aegirine (Acmite)

NaFe3+[Si2O6]

2/m

Omphacite

jadeite-augite solid solution

2/m

Aegirine augite

Aegirine augite solid solution

2/m

LiAl[Si2O6]

2/m

Na-Ca pyroxenes

11

Li pyroxene Spodumene

ferrosilite has not been observed in nature. By definition, the content of the wollastonite component, Ca2[Si2O6], in Mg–Fe pyroxenes is ≤5 mol%. Moreover, small amounts of Al, Fe3+, Mn, Ti, Cr and Ni may be present. Occurrence  Orthopyroxenes are common in mafic igne-

ous rocks, such as norite, picrite, basalt and can be present even in andesite (7 Sect. 13.2.1). Due to a wide miscibility gap between the diopside–hedenbergite and enstatite–ferrosilite solid solution series, orthopyroxenes abundantly coexist with clinopyroxenes. In such rocks, either orthopyroxene develops exsolution lamellae of clinopyroxene or the other way around (. Fig. 13.5a). Mg-rich orthopyroxenes are common constituents of ultramafic magmatites and their metamorphic equivalents. Moreover, orthopyroxenes are typical of high-grade metamorphic rocks, especially pyroxene granulite (7 Sect. 28.3.5). A considerable amount of orthopyroxene occurs in peridotite that constitutes much of the Earth’s upper mantle

(7 Sect. 29.3.1), fragments of which can be brought to the surface as xenoliths during volcanic eruptions. Furthermore, orthopyroxene is present in many primitive meteorites known as chondrites (7 Sect. 31.3.1). Members of the monoclinic solid solution series clinoenstatite–clinoferrosilite are extremely rare in nature, e.g., occasionally as skeletal crystals in volcanic rocks. Clinoenstatite has been observed in some meteorites (. Table 31.2) and as circumstellar dust particles (. Table 34.1).

Ca Pyroxenes z Diopside CaMg[Si2O6] z Hedenbergite CaFe[Si2O6]

181

11.4 · Chain Silicates (Inosilicates)

Crystal form and habit  Crystal class 2/m; commonly, diopside, salite, ferrosalite and hedenbergite form granular aggregates. Well-developed, long prismatic crystals of diopside show dominant pinacoids {100} and {010}, giving rise to nearly rectangular cross sections perpendicular to c, whereas the prisms {111} and {110} are more subordinate (. Fig. 11.32). Physical properties Cleavage

{110}, imperfect to variably distinct, the cleavage angle of ~87° is typical of all clinopyroxenes. In the variety “diallage”, translation (100) leads to parting along this face

Fracture

conchoidal, brittle

Hardness

5½–6½

Density

3.22 (for pure diopside) to 3.56 (for pure hedenbergite)

Optical properties

transparent to translucent; vitreous lustre; diopside is white, grey or greyish green, chrome-diopside emerald green, hedenbergite greenish black

Chemical composition  Diopside and hedenbergite form a complete solid solution series (. Fig. 11.29a) in which density and refractive indices increase almost linearly with increasing Fe/Mg ratio. To a certain extent, Si can be replaced by Al. Additional constituents are Fe3+ in ferrian and Cr3+ in chromian diopsides, while Zn2+ can be present in zincian diopside and hedenbergite. Complete solid solution exists between diopside-hedenbergite and the end-member johannsenite, CaMn[Si2O6]. Occurrence  Diopside

is a constituent of mafic or ultramafic igneous rocks and of peridotite nodules, derived from the Earth’s mantle. It also occurs in dolomitic marble, in diopside amphibolite and pyroxene granulite, whereas hedenbergite can be a constituent of Fe-rich, contact metasomatic calc-silicate rocks, known as skarn.

Physical properties Cleavage

{110}, good; translation on (100) leads to parting along this face

Fracture

conchoidal, brittle

Hardness

5½–6

Density

3.19–3.56, depending on composition

Optical properties

lustre: dull, more rarely vivid lustre on cleavage planes and crystal faces, rarely translucent; common augite is green to brownish black, the Fe-, Ti-rich basaltic augite pitch-black

Chemical composition  Compared to the diopside–hedenbergite solid solution series, the chemical composition of augite is much more complicated, mainly due to the coupled substitutions [6]  [6]  ⇋ Na+[8] Al, Fe3+ , and 5 Ca[8] Mg, Fe2+ 5 Ca[8] Mg, Fe2+ )[6] Si[4] ⇋ Al[6] Al[4] .

By the latter exchange, the Ca- and Mg-Tschermak’s molecules (Ca-Ts), Ca[8]Al[6][Al[4]SiO6], and (Mg-Ts), Mg[6]Al[6][Al[4]SiO6], are formed which, however, are not known as pyroxene end-members. Augite with 3–5 wt% TiO2, referred to as titanian augite, displays characteristic, hour-glass shaped growth structures (. Fig. 13.12a). Additional minor constituents are Cr, predominantly in Mg-rich, and Mn in Fe-rich augite. Occurrence  Augite is a widespread rock-forming min-

eral. Alone or together with orthopyroxene and/or olivine it forms the mafic constituent of the plutonic rock gabbro, whereas basaltic augite and titanian augite are its counterparts in basalt, the most common volcanic rock type.

z Augite (Ca,Na)(Mg,Fe2+,Al,Fe3+,Ti)[(Si,Al)2O6] Crystal form and habit  Crystal class 2/m; augite is commonly short-prismatic with dominant vertical prism {11 0} and longitudinal prism {111}, combined with pinacoids {100} and {010} (. Fig. 11.31a). This applies especially to well-developed crystals of basaltic augite that can occur as phenocrysts in volcanic and pyroclastic rocks (. Figs. 2.6, 3.2). Augite crystals can form simple or multiple twins on {100} or penetration twins on {101}.

. Fig. 11.31  Crystal forms and habits of clinopyroxenes: a short-pris¯ ¯ matic augite turned around c by 180°, with pinacoids {100} and {010} and prisms {110} and {111}: b acmite, elongate //c with prisms {110}, {310}, {661} and {221} as well as pinacoid {100}

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Chapter 11 · Silicates

Alkali Pyroxenes z Aegirine (acmite) NaFe3+[Si2O6] Crystal form and habit  Crystal class 2/m; acicular crystals with steep terminal prisms {661} and {221}, vertical prisms {110} and {310} as well as pinacoid {100} are common (. Fig. 11.31b) as is simple or lamellar twinning on {100}. In many cases, aegirine needles form tufted aggregates. Physical properties

11 . Fig. 11.32  Crystals of diopside (light green) and grossular (variety hessonite, red) from Mussa Alp, Piedmont, Italy; Width of view is c. 1 cm

Cleavage

{110}, more distinct than in other pyroxenes, parting on {100}

Hardness

6

Density

3.55–3.60 (aegirine), 3.40–3.60 (aegirine augite)

Optical properties

translucent; vitreous to resinous lustre; dark green, greenish or reddish brown to black

Chemical composition  Due to the coupled substitution  Na+ Fe3+ ⇋ Ca2+ Mg, Fe2+ aegirine forms a solid solution series with augite, known as aegirine augite, which is more common than the pure aegirine end-member. Moreover, the substitution Fe3+ ⇋ Al leads to solid solutions between aegirine and jadeite as well as aegirine augite and omphacite (. Fig. 11.29c). Occurrence  Aegirine and aegirine augite are common

Augite is also an important constituent of ultramafic peridotite nodules, derived from the Earth’s upper mantle. In addition, augite can occur in high-grade metamorphic rocks, such as pyroxene granulite. z Pigeonite Ca0.25(Mg,Fe)1.75[Si2O6]

In contrast to the orthopyroxenes, the Ca-poor monoclinic clinopyroxene pigeonite contains about 5–20 mol% of the Ca2Si2O6 component (. Fig. 11.29a). It forms phenocrysts of brown, greenish-brown or black colour and with long-prismatic habit in basaltic rocks. Its hardness is 6, its density ranges from 3.17 to 3.46, depending on composition. Distinction from other pyroxenes is possible only by microscopic methods, X-ray diffraction and/or electron microprobe analysis. Commonly, pigeonite crystallises as early segregate from hot, basaltic lavas that undergo rapid cooling. Upon slow cooling, however, a complicated exsolution process sets in by which augite lamellae develop parallel to (001) in a host crystal of Ca-depleted pigeonite or orthopyroxene, known as inverted pigeonite. This process can be described by the pseudo binary system Mg2Si2O6 (En)–CaMgSi2O6 (Di) (. Fig. 18.16b).

mafic constituents in alkaline igneous rocks, such as phonolite (7 Sect. 13.2.2, . Fig. 13.11a), typically containing zoned phenocrysts with a core of aegirine augite and a rim of aegirine. Na–Fe3+-rich clinopyroxenes are also formed in metamorphic rocks. z Jadeite NaAl[6][Si2O6] z Omphacite (Ca,Na)(Mg,Fe2+,Fe3+,Al)[Si2O6] Crystal form and habit  Crystal class 2/m; commonly, jadeite

shows fibrous habit and forms matted aggregates, whereas prismatic jadeite is rare. In eclogite, omphacite constitutes, together with garnet, a principal rock-forming phase. Physical properties Cleavage

{110}

Hardness

6 (jadeite), 5–6 (omphacite); tufted jadeite aggregates are very tough

Density

3.24–3.43 (jadeite), 3.16–3.43 (omphacite)

Optical properties

translucent; lustre: vitreous, pearly on cleavage planes, silky lustre on tufted aggregates of jadeite; jadeite is pale green to deep green, also colourless, omphacite light to dark green

183

11.4 · Chain Silicates (Inosilicates)

Chemical composition  Al[6] can be replaced by Fe3+, thus

forming solid solutions with aegirine. Omphacite is a solid solution between augite and jadeite (. Fig. 11.29c).

Occurrence  Jadeite is a typical high-pressure mineral that

occurs in blueschist and jadeite gneiss. It forms during high-P and ultra-high-P metamorphism by the reaction

Physical properties Density

3.03–3.23

Optical properties

transparent to translucent, in part clouded; lustre: commonly dull, but faces of limpid crystals can show vitreous lustre; colourless, light grey, pale pink, yellowish, pale green

Chemical composition  To a certain extent, Li+ can be

[6] Na[AlSi[4] 3 O8 ] ↔ NaAl [Si2 O6 ] + SiO2 albite jadeite quartz

[11.1]

(. Fig. 26.1), whereby the framework-silicate structure of albite with Al[4] is replaced by the denser chain-silicate structure of jadeite with Al[6]. Exceptionally high pressures at low to moderate temperatures are realised in subduction and continental collision zones. Pure jadeite rocks, known as jade (or jadeitite), are also formed in subduction zones. Worldwide, many jade occurrences are known, especially in Japan, but also in Myanmar, the Polar Urals, Russia, on the island of Syros, Cyclades archipelago, Greece, on the Caribbean Islands and in California. A common companion of jadeite rocks is serpentinite (7 Sect. 11.5.6). Nearly monomineralic jadeite rocks are formed by two different, but spatially and genetically related, processes (e.g., Tsujimori and Harlow 2012; Meng et al. 2016): (1) Jadeite precipitates directly from Na– Al–Si-rich hydrothermal solutions thereby forming dykes and fissures filled with jade. (2) The jadeite-saturated solutions replace the adjacent country rocks, a process referred to as Na-metasomatism.

Omphacite constitutes, together with garnet, the rock eclogite, which is typical of high-pressure metamorphism of basaltic protoliths (7 Sect. 26.3.1, 28.3.9). Economic relevance  Beautifully coloured or pure white jade

is a highly valued ornamental stone for producing objects of handicraft, especially by the traditional artisans of China. In prehistoric times, jade was a much demanded raw material for the production of weapons and tools, making use of its excellent mechanical properties, such as high toughness.

z Spodumene LiAl[6][Si2O6] Crystal form and habit  Crystal class 2/m; spodumene can occur as well-developed crystals, the faces of which are commonly rough, etched or striated parallel to c. Giant crystals, up to 90 t in weight, have been described (Rickwood 1981). Rock-forming spodumene occurs in coarsegrained, sparry or radiant aggregates. Physical properties Cleavage

{110}

Hardness

6½–7

replaced by Na+, and Al3+ by Fe3+.

Occurrence  Spodumene is a typical mineral of Li-rich pegmatites (7 Sect. 22.3). Hydrothermal recrystallisation

can lead to clear, colourless or beautifully coloured crystals of gem quality, especially the lilac variety kunzite or the emerald-green hiddenite. Economic relevance  Spodumene is an important raw

material for the production of lithium, which is used, e.g., in lithium-ion batteries, for alloys with high strength/ weight ratio as well as in glass and ceramics of high thermal resistance. Kunzite and hiddenite are cut and polished as gemstones.

11.4.2  Pyroxenoids

The general chemical formula of the pyroxenoids is M[SiO3] or a multiple of it, where M predominantly stands for Ca, Mg, Fe and Mn. As in the pyroxenes, the pyroxenoid structures are built of infinite single chains of [SiO4]-tetrahedra. However, the lattice constants c0, defined by the distance between each two [SiO4]-tetrahedra of identical orientation, are longer (Liebau 1959, 1985). Whereas the pyroxene structure is built of a 2-periodic single chain, in which each second tetrahedron points in the same direction, the pyroxenoid structures consist of 3-periodic, 5-periodic or 7-periodic single chains. Prominent examples are wollastonite, Ca3[Si3O9], rhodonite, (Mn,Ca,Fe)5[Si5O15], and pyroxferroite, (Ca,Fe)(Fe,Mn)6[Si7O21], respectively (. Fig. 11.33). Depending on the respective chain type, the [6]-coordinated cations show different arrangements in the individual structures.

z Wollastonite Ca3[Si3O9], Simplified Ca[SiO3] Crystal form and habit  Wollastonite occurs in various

modifications: low-temperature polytypes are the triclinic wollastonite-Tc of crystal class 1¯ and the monoclinic wollastonite-2M, also known as para-wollastonite, of crystal class 2/m. Above 1150 °C, the triclinic pseudo-wollastonite is stable, its structure being made up of three-membered

11

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Chapter 11 · Silicates

elements and as a substitute for asbestos because of its high melting point of 1540 °C. z Rhodonite (Mn,Ca,Fe)5[Si5O15] Crystal form and habit  Crystal class 1¯; prismatic or tabular crystals are rare. Rhodonite commonly occurs in pink to flesh-coloured masses that are transsected by black veins of Mn-oxides. Physical properties Cleavage

{100} and {001}, perfect

Hardness

5½–6½

Density

3.4–3.7

Optical properties

transparent to translucent; vitreous lustre, pearly on cleavage planes; light incarnate to rose coloured or brownish-red

. Fig. 11.33  Single chains of [SiO4] tetrahedra. a 2-periodic single chain: pyroxenes; b 3-periodic single chain: wollastonite; c 5-periodic single chain: rhodonite; d 7-periodic single chain: pyroxferroite

11

rings [Si3O9], which led to the term cyclo-wollastonite for this type. Commonly, wollastonite forms massive, fibrous, acicular or stalk-like aggregates, whereas well-developed, tabular or acicular crystals are rare.

Chemical composition  Rhodonite is a complex solid solu-

tion with variable contents of Mn ≫ Ca > Fe2+ ≷Mg .

Occurrence  Predominantly

deposits.

in

metamorphosed

Mn

Physical properties

Economic relevance  Rhodonite is used as dimension

Cleavage

{100} and {001}, perfect

Hardness

4½–5

Density

2.86–3.09

Optical properties

transparent to translucent; lustre: vitreous, pearly on cleavage planes, silky on fibrous aggregates; commonly white, also faintly coloured

stone and for artifacts.

Chemical composition  Wollastonite can contain consider-

able amounts of Mg, Fe and Mn.

Occurrence  Wollastonite is a typical contact metamor-

phic mineral, formed by thermal overprint of siliceous limestones, especially by the reaction

CaCO3 + SiO2 ⇋ Ca[SiO3 ] + CO2 Calcite quartz wollastonite

[11.2]

Pseudo-wollastonite is found in volcanic ejecta, affected by pyrometamorphism at very high temperatures (7 Sect. 28.3.7). Economic relevance  Wollastonite is a raw material for

high-temperature ceramics. Moreover, it is used as filler in plastics, stains, glues, insulating materials, construction

11.4.3  Amphibole Family

The chemical composition of amphiboles can be expressed by the general formula A0–1B2C5[T8O22](OH,F)2. The various sites in the amphibole structure can be occupied by the following cations: 5 A = Na+, more rarely K+, ⃞ (vacant site) 5 B = Na+, Ca2+, Mg2+, Fe2+, Mn2+ 5 C = Mg2+, Fe2+, Mn2+, Al3+, Fe3+, Ti4+ 5 T = Si4+, Al3+ In the C position, substitution of Al3+ by Fe3+ as well as between Ti4+ and other cations is limited. The same holds true for the Si4+ ⇋ Al3+ exchange in the T position. In the amphibole structure, corner-sharing [TO4] tetrahedra form infinite double chains parallel to c, which resemble a rope-ladder composed of an infinite sequence of 6-membered rings (. Fig. 11.36). Their composition is [(Si,Al)4O11]. The C cations form infinite ribbons parallel to c of edge-sharing [C(O,(OH)6] octahedra that are connected, via corners, with the [TO4] double chains to form alternating, sandwich-like units

185

11.4 · Chain Silicates (Inosilicates)

By analogy to the pyroxene family, monoclinic clinoamphiboles are clearly predominant. The ratio between [6]- and [8]-coordinated sites, occupied by the C and B cations, respectively, is 5:2, as compared to 1:1 in clinopyroxenes. In the orthorhombic Mg–Fe amphiboles, known as orthoamphiboles, all the cationic positions are [6]-coordinated, as in the orthopyroxenes. The amphibole nomenclature of the IMA Commission on New Minerals and Mineral Names (Leake et al. 1997) is given in . Table 11.6.

referred to as “I-beams”. Depending on their respective sizes, the [6]-coordinated C cations can be placed in three different positions, M1, M2 and M3. At the margins of the octahedral ribbons, unshared oxygens are replaced by the relatively large anions of second order, commonly (OH)– or, to a minor extent, F–. This gives rise to the spacious M4 position that, in the Ca-, Na–Ca and Na-amphiboles, is mainly occupied by the large [8]-coordinated B cations Ca2+ and Na+. The even larger A sites, situated between the centres of neighbouring 6-membered rings in adjacent double chains, are fully or partly occupied by [10]- or [12]-coordinated Na+ but can also remain vacant (⃞) (. Fig. 11.36).

Mg–Fe Amphiboles z Anthophyllite—Ferroanthophyllite (Mg,Fe2+)7[Si8O22](OH)2

. Table 11.6  Important amphibole end-members Mineral

Formula

Crystal class

Anthophyllite-ferroanthophyllite

⎕(Mg,Fe2+)7[Si8O22](OH)2

2/m2/m2/m

Gedrite-ferrogedrite

⎕(Mg,Fe2+)

5Al2[Al2Si6O22](OH)2

2/m2/m2/m

Cummingtonite-grunerite

⎕(Mg,Fe2+)

7[Si8O22](OH)2

2/m

Mg–Fe–Mn amphiboles

Li amphiboles Holmquistite-ferroholmquistite

⎕Li2(Mg,Fe2+)3Al2[Si8O22](OH)2

2 /m2/m2/m

Clinoholmquistite-clinoferroholmquistite

⎕Li2(Mg,Fe2+)3Al2[Si8O22](OH)2

2/m

Tremolite-actinolite-ferroactinolite

⎕Ca2(Mg,Fe2+)5[Si8O22](OH)2

2/m

Magnesiohornblende-ferrohornblende

⎕Ca2

(Mg,Fe2+)

2/m

Tschermakite-ferro-/ferritschermakite

⎕Ca2

(Mg,Fe2+)

Edenite-ferroedenite

NaCa2(Mg,Fe2+)5[AlSi7O22](OH)2

2/m

Pargasite-ferropargasite

NaCa2(Mg,Fe2+)4Al[Al2Si6O22](OH)2

2/m

Ca amphiboles

Magnesiohastingsite-hastingsite Kaersutite-ferrokaersutite

4

(Al,Fe3+)[AlSi

3

(Al,Fe3+)

7O22](OH)2

2[Al2Si6O22](OH)2

NaCa2

(Mg,Fe2+)

Fe3+[Al

4

NaCa2

(Mg,Fe2+)

4Ti[Al2Si6O23](OH)

2Si6O22](OH)2

2/m

2/m 2/m

Na–Ca amphiboles Richterite-ferrorichterite

NaCaNa(Mg,Fe2+)5[Si8O22](OH)2

2/m

Magnesiokatophorite-katophorite

NaCaNa(Mg,Fe2+)4(Al,Fe3+)[AlSi7O22](OH)2

2/m

Magnesiotaramite-taramite

NaCaNa(Mg,Fe2+)3AlFe3+[Al2Si6O22](OH)2

2/m

Winchite-ferowinchite

⎕CaNa(Mg,Fe2+)

2/m

Barroisite-ferrobarroisite

⎕CaNa(Mg,Fe2+)

4

(Al,Fe3+)[Si

3

AlFe3+[AlSi

8O22](OH)2 7O22](OH)2

2/m

Na amphiboles Glaucophane-ferroglaucophane

⎕Na2(Mg,Fe2+)3Al2[Si8O22](OH)2

2/m

Magnesioriebeckite-riebeckite

⎕Na2(Mg,Fe2+)3Fe23+[Si8O22](OH)2

2/m

Eckermanite-ferroeckermanite

NaNa2(Mg,Fe2+)4Al[Si8O22](OH)2

2/m

Magnesioarfvedsonite-arfvedsonite

NaNa2

(Mg,Fe2+)

4

Fe3+[Si

8O22](OH)2

2/m

11

186

Chapter 11 · Silicates

z Gedrite—Ferrogedrite (Mg,Fe2+)5Al2[Al2Si6O22](OH)2 Crystal form and habit Crystal class 2/m2/m2/m; anthophyllite and gedrite form subparallel or radiating aggregates of columnar to acicular crystals along c or asbestiform fibres. Physical properties

11

Cleavage

{210}, perfect, {010} and {100} poor; ¯ ~54½°; cleavage angle (210):(210) transversal parting ⊥c is common

Hardness

5½–6

Density

2.85 (anthophyllite) to 3.57 (ferrogedrite)

Optical properties

anthophyllite transparent to translucent; vitreous lustre, pearly on cleavage planes, bronze-like schiller on (010); gedrite: translucent, transparent on thin edges; vitreous to silky lustre; colourless, yellowish grey to yellowish brown, green, clove-brown, dark brown, depending on Fe content

Chemical composition  The orthorhombic Mg-Fe amphi-

boles form a complete solid solution series that extends from nearly pure anthophyllite, Mg7[Si8O22](OH)2, to ferroanthophyllite with up to 65 mol% Fe72+[Si8O22](OH)2. Mg-rich members are much more common than Fe-rich ones.

In contrast, the orthorhombic gedrite-ferrogedrite series covers the full range of Mg ⇋ Fe2+ substitution, and (Mg,Fe2+)[6]Si[4] is partly replaced by Al[6]Al[4], similar to the Tschermak’s substitution in pyroxenes. The limit between anthophyllite and gedrite has been arbitrarily set at Si7, i.e., halfway between the two end-members. At low temperature, solid solution between anthophyllite and gedrite is incomplete. Upon cooling, exsolution of fine lamellae takes place and causes iridescence on {010}. The large A site in the gedrite structure can accommodate some Ca2+ indicating a limited solid solution with Ca-amphiboles. Moreover, some Na+ can be present on the A site due 3+[4] to the coupled substitution ⃞[12]Si4+[4] ⇋ Na+[12] + Al .

Crystal form and habit  Crystal class  2/m; crystals of

fibrous to acicular habit form radiating or tufted aggregates. Simple or lamellar twinning on {110} is common. Physical properties Cleavage

{110}, good, cleavage angle (110): ¯ ~55° (110)

Hardness

5–6

Density

3.10–3.60, increasing with Fe content

Optical properties

light green to greyish green, beige, brown

Lustre

silky

Chemical composition  The monoclinic Mg–Fe amphi-

boles form a complete solid solution series between the virtually pure end-members cummingtonite, Mg7[Si8O22] (OH)2, and grunerite, Fe72+[Si8O22](OH)2, the limit being set at Mg:Fe2+=1:1. Rarely, monoclinic Mg–Fe amphiboles with considerable Mn2+ contents occur, with the pure end-members manganocummingtonite, Mn2Mg5[Si8O22] (OH)2, and manganogrunerite, Mn2Fe52+[Si8O22](OH)2. Low contents of Ca and Na in these testify to limited solid solutions with Ca-amphiboles. Occurrence  Cummingtonite and grunerite are relatively

common in metamorphic terrains, where they occur, together with hornblende, in amphibolites or, together with anthophyllite, in metamorphosed ultramafic rocks. Grunerite and manganogrunerite, coexisting with quartz, magnetite and/or haematite, are constituents of metamorphosed banded iron formation (BIF).

Economic relevance  For decades, fibrous grunerite, known as amosite, had been mined in great quantities as asbestos by the AMOSA company (Asbestos Mines of Southern Africa) in Mpumalanga, South Africa. However, with growing awareness of the extreme health hazards of amphibole asbestos, these operations were closed down long ago.

Occurrence  Orthoamphiboles occur in Mg-rich meta-

morphic rocks, such as anthophyllite-cordierite gneiss, or form by metasomatic reaction between ultramafic bodies, e.g., serpentinites, and their country rocks. Rarely they can be found also in magmatic rocks.

Economic relevance  Fibrous anthophyllite had been mined as asbestos, especially in two large deposits at Tuusniemi, Finland (1918–1975), and Matsubase, Japan (1883– 1970), until the toxicity of asbestos was recognised.

z Cummingtonite—Grunerite (Mg,Fe2+)7[Si8O22](OH)2

Ca Amphiboles z Tremolite Ca2Mg5[Si8O22](OH)2 z actinolite—ferroactinolite Ca2Fe52+[Si8O22](OH)2 Crystal form and habit  Crystal class 2/m; prismatic, stalk-

like or acicular crystals, with predominant vertical prism {110} and subordinate {121} (. Fig. 11.34a), are typically

187

11.4 · Chain Silicates (Inosilicates)

Physical properties

. Fig. 11.34  Amphibole forms, combination of different faces such as vertical prism {110}, vertical pinacoid {010}, monoclinic prisms {011}, ¯ pinacoid {001}. a Actinolite; b, c hornblende, where crystal {121}, {111}, c is rotated, against crystal b, by 90° around the c axis

Cleavage

¯ ~56° {110}; cleavage angle (110):(110) transversal parting ⊥c is common

Hardness

5–6

Density

2.99–3.48, increasing with Fe content

Optical properties

transparent to translucent; lustre: commonly dull, but faces of limpid crystals can show vitreous lustre; tremolite is colourless, light grey, pale pink, yellowish, actinolite pale to dark green, ferroactinolite dark green to black

Chemical composition  Most amphiboles of this group are

tremolite and actinolite with Mg/(Mg+Fe) ratios of >0.9 and 0.9–0.5, respectively, whereas ferroactinolite with Mg/(Mg+Fe) ratios of 0.5, Ca in B position >1.5 and Ca in A position 1:5 and conforming Al:Si ratios of >3:5 are rare. Intermediate members, rich in Mg and Fe, respectively, used to be referred to as meroxene and lepidomelane but these terms have become obsolete. Occurrence  Biotite is a very widespread rock-forming

mineral that commonly occurs in metamorphic rocks, such as micaschist (. Fig. 26.12a) and gneisses as well as in felsic plutonic rocks, such as granite or granodiorite (. Fig. 13.4a, b) and related pegmatite. However, due to its limited thermal stability, especially at low pressures, biotite is less abundant in volcanic rocks (. Fig. 13.8b) and, if initially formed, is commonly altered to fine-grained, opaque mixtures rich in magnetite, known as “opacite”.

z Lepidolite K(Li,Al[6])3[(Si,Al[4])4O10](F,OH)2 z Zinnwaldite K(Fe2+,Li,Al,⃞)3[(Si,Al)4O10](OH,F)2 Crystal form and habit  The polytypes 1M and 2M2 of

crystal class 2/m are most common, whereas the monoclinic polytype 2M1 (2/m) and the trigonal polytype 3T (32) or mixtures of both occur more rarely. Lepidolite is commonly arranged in flaky or scaly, in cases hemispheric aggregates, whereas well-developed crystals are rare. In contrast, zinnwaldite can form tabular crystals with pseudo-hexagonal outline, commonly grown in fan-shaped aggregates.

Cleavage

{001} perfect, flakes are flexible and elastic

Physical properties

Hardness

2–3

Cleavage

{001}, perfect

Density

2.7–3.3, depending on Fe content

Hardness

2½–4

Optical properties

translucent to transparent, non-transparent when rich in tiny magnetite inclusions (opacite); pearly lustre on cleavage planes; dark green, brownish green, light brown, dark brown to blackish brown

Density

2.80–2.90 (lepidolite), 2.90–3.02 (zinnwaldite)

Optical properties

transparent to translucent; pearly lustre; lepidolite is white to pale pinkor peach-blossom coloured, due to traces of Mn2+; zinnwaldite is pale violet, silver-grey, yellowish, brownish or even almost black

Crystal structure  Trioctahedral 3-layer silicate. Chemical composition  Compared to phlogopite, a varia-

ble proportion of Mg2+ in the octahedral layer is replaced by Fe2+ as well as Fe3+, Al[6] and Ti4+, to a minor extent also by Mn and Li. Charge balance is achieved by the

Crystal structure and chemical composition  Due to its small ionic radius, Li+ does not occupy the [12]-coordinated interlayer sites but replaces [6]-coordinated Al3+ in the octahedral layers. A Li:Al ratio of exactly 1:1 would result

195

11.5 · Sheet Silicates (Phyllosilicates)

in perfect charge balance and in ideal trioctahedral occupancy. In nature, however, the Li:Al[6] ratio is highly variable. Therefore, charge balance is achieved either by a corresponding variation of the Si:Al[4] ratio in the tetrahedral layer or, alternatively, by vacancies in the octahedral layer. This leads to transitions between tri- and dioctahedral occupancy. Occurrence  Lepidolite, the most common Li-mineral, and

zinnwaldite occur, together with other Li-bearing silicates, such as amblygonite, LiAlPO4(F,OH), or spodumene, LiAl[Si2O6], in granitic pegmatites. Zinnwaldite also forms under high-temperature hydrothermal conditions and can be accompanied by cassiterite, topaz, fluorite and quartz. Economic relevance  Lepidolite and zinnwaldite are mined

as raw materials for producing the light metal lithium, used in Li and Li-ion batteries, in special alloys with high strength/weight ratios, in special glasses and ceramics of high thermal resistance, in Li-salts, and pyrotechnical material.

11.5.3  Hydromica Group z Illite (K,H3O)Al2[(Si,Al)4O10](H2O,OH)2 Crystal structure  The hydromica illite (hydromuscovite)

is a dioctahedral, more rarely trioctahedral, 3-layer silicate with mica-structure (. Fig. 11.37) in which K+ is partly replaced by H3O+. The disordered illite-1Md is prevalent, whereas disordered versions of the polytypes 3T and 2M1 are less common. Grain size  Commonly, illite is very fine-grained (30 gt−1, and extreme contents of 1703 and 1886 gt−1 (Wagner 1929; Scoon and Mitchel 2010, 2011).

21.3.2  Pyrrhotite-Pentlandite-Chalcopyrite

Deposits in Komatiites

Komatiite is an ultramafic volcanic rock that crystallised from melt with an extremely high liquidus temperature. Its temporal distribution is restricted essentially to the Archaean. Due to its high age, komatiite typically shows evidence of metamorphic overprint, turning it into greenish rocks, typical constituents of greenstones belts. Its effusion as lava flows or intrusion as dykes clearly indicate that the geothermal gradient in the Archaean must have been distinctly higher than in Proterozoic and Phanerozoic times. Liquid exsolution of sulfide melts caused considerable concentrations of Ni- and Cu-ore, containing pyrrhotite, pentlandite, chalcopyrite and pyrite as dominant ore minerals, whereas the amount of PGE is low. At the base of the lava flows, a zone with massive ore bodies can be developed, which is sharply delineated against an overlying zone of ultramafic rock, penetrated by a network of sulfide ore. This zone is overlain, again with sharp boundary, by ultramafic rock with minor or missing ore contents (Barnes et al. 2017; Stone et al. 2004). In many cases, the komatiite-hosted ore became metamorphosed, which led to the decomposition of pyrrhotite to form magnetite + pyrite according to the oxidation reaction

12FeS + 6O2 → 3Fe2+ Fe3+ 2 O4 + 3FeS2 + 3S2 Pyrrhotite magnetite pyrite

[21.1]

and, consequently, to relative enrichment of Ni (Stone et al. 2004). Economically important examples are the Alexo Mine at Timmins, Ontario (e.g., Houlé et al. 2012), the Kambalda Ore Field (Stone et al. 2004) and the Ni deposit Mount Keith in the Yilgarn Craton (Barnes et al. 2011), both situated in Western Australia, as well as some ore deposits in Zimbabwe and South Africa, e.g., at the Komati River. The

sulfide ore deposits of Petsamo (Petchenga), Karelia, Russia, and of Lynn Lake, Manitoba, Canada, are related to lavas of picritic tholeiite. 21.3.3  Magnetite-Apatite Deposits

Oxidic melts can be separated from silicatic melts in the liquid state and can form individual intrusions or even lava flows when high proportions of volatile components, such as H2O, CO2, F, Cl, P2O5 and B2O3 are present. Thus oxide ore deposits can form, most notably massive magnetite-apatite deposits. A prime example is Europe’s largest iron ore deposit at Kiruna, northern Sweden, for which an orthomagmatic origin has been suggested (Jonsson et al. 2013). However, it is still under discussion, whether the Fe oxide-rich melt was intrusive or extrusive (7 Sect. 23.5.4). The ore bodies at Kiruna are intercalated within felsic to intermediate (sub)volcanic rocks that have been deformed and metamorphosed. The biggest ore body Kirunavaara has an extension of 40 km and is 80–90 m thick. The remaining resource at Kiruna is huge (1012 Mt) with an average grade of 48% Fe and up to 5% P. The typically Ti-poor magnetite at Kiruna contains isolated grains or layers of fluorapatite, Ca5(PO4)3F. The magnetite is partly altered to secondary haematite, a process known as martitisation. Under the influence of a fluid phase, scapolite (7 Sect. 11.6.5) formed together with albite and/or tourmaline. Starting from the ore bodies, the country rock is progressively scapolitised and albitised. Additional important iron-ore deposits of this type are Grängesberg, Sweden, Pea Ridge and Iron Mountain, Missouri, USA, Durango, Mexico, and Savage River, Tasmania, as well as several deposits in the Bafq district in central Iran Other economically relevant magnetite-apatite deposits are related to carbonatites, as described in the following section. 21.4  Carbonatite- and Alkaline-Magmatic

Rock-Hosted Deposits

Carbonatites and foid-bearing alkaline igneous rocks commonly occur together to form intrusive (ring) complexes as well as dyke swarms or volcanoes. The tectonic setting in which they form are typically intra-continental rift zones or major fault zones in stable cratons. In many cases, a spatial and genetic relationship to a hot-spot in the underlying mantle seems indicated. There is no doubt that carbonatite magmas form by liquid exsolution of mantle-derived melts (Lee and Wyllie 1998; Chakhmouradian and Zaitsev 2012). By this process and following fractional crystallisation, a wealth of economically interesting minerals can be concentrated, such as apatite, magnetite, pyrochlore, (Ca,Na)2Nb2(O,OH,F)7, xenotime, YPO4, the Zr minerals zircon and baddeleyite, ZrO2, as well as the REE minerals monazite, (Ce,La,Nd)PO4, bastnäsite, (Ce,La,Nd,Y)CO3(F,OH),

21

368

21

Chapter 21 · Orthomagmatic Mineral Deposits

parisite, Ca(Nd,Ce,La)2(CO3)3F2, burbankite, (Na,Ca)3(Sr,Ba,Ce)3(CO3)5, and loparite, (Na,Ce,Ca)(Ti,Nb)O3. Besides, igneous complexes of carbonatite and alkaline magmatic rocks can contain industrial minerals, such as fluorite, baryte and strontianite in economic amounts. One of the largest REE deposit worldwide is situated in the Mountain Pass carbonatite, California, USA. It is surpassed, however, by the giant deposit Bayan Obo, Inner Mongolia, northern China, containing 70% of the world reserves on REE minerals, mainly bastnäsite, parisite and monazite, amounting to 48 Mt REE2O3 as well as 2.2 Mt Nb2O3. Details of the petrogenesis of this deposit, situated in the North-China Craton, are still under discussion (e.g., Laznicka 2006; Kynicky et al. 2012). Although the stratabound mineralisations form layers, lenses, veins and impregnations in metamorphosed sediments, mainly dolomite marbles of Proterozoic age, the metal content is derived, in all probability, from carbonatitic magmas that are documented by discordant carbonatite dykes, occurring in close vicinity (e.g., Le Bas et al. 1992; Yang et al. 2011).

The famous magnetite-apatite deposits of Kovdor, Chibiny and Lovozero, Kola Peninsula, Russia, and of Sokli, Finland, are hosted by complexes of carbonatite and alkaline magmatic rocks as well. In these deposits, apatite contains economically relevant amounts of SrO, REE2O3 and Re2O3. In the 2060 Ma old Phalaborwa Igneous Complex, Limpopo Province, South Africa, consisting of alkaline magmatic rocks and carbonatite, concentrations of Cu-ore minerals, such as chalcopyrite, cubanite, CuFe2S3, bornite, Cu5FeS4, and valleriite, (Fe,Cu)S·0.75(Mg,Al,Fe) (OH)2, are mined. The sulfide mineralisation took place at an early stage of the magmatic evolution by separation of a copper-sulfide melt by liquid immiscibility. As by-products, magnetite, apatite, Au, Ag, PGE, U, Zr etc. are also extracted. A nearby open-cast mine has been opened in a 2062 Ma apatite-rich pyroxenite, also known as phoscorite, so far representing the largest orthomagmatic apatite deposit worldwide (Evans 1993; Wu et al. 2011). In addition, the industrial mineral vermiculite, an expandable clay mineral (7 Sect. 11.5.7), is mined at Phalaborwa. Mining took place, from 1966 to 2002, in the second largest opencast pit on Earth, attaining a diameter of almost 2000 m and a depth of 230 m below the sea-level, and is now continued by underground operations.

References and Suggestions for Further Reading Ballhaus CG, Stumpfl EG (1986) Sulfide and platinum mineralization in the Merensky Reef: evidence from hydrous minerals and fluid inclusions. Contrib Miner Petrol 94:193–204 Barkov AY, Fleet ME (2004) An unusual association of hydrothermal platinum-group minerals from the Imandra Layered Complex, Kola Peninsula, Northwestern Russia. Can Mineral 42:455–467 Barnes SJ, Fiorentini ML, Fardon MC (2011) Platinum group element and nickel sulphide ore tenors of the Mount Keith nickel deposit, Yilgarn Craton, Australia. Miner Deposita 47:129–150

Bateman JD (1951) The formation of late magmatic oxide ores. Econ Geol 53:404–426 Bowen NL (1928) The Evolution of Igneous Rocks. Dover Publications, New York (Reprint 1956) Dare SAS, Barnes S-J, Pritchard HM (2010) The distribution of platinum group elements (PGE) and other chalcophile elements among sulfides from the Creighton Ni-Cu-PGE deposit, Sudbury, Canada, and the origin of palladium in pentlandite. Miner Deposita 45:765–793 Dietz RS (1964) Sudbury structure as an astrobleme. J Geol 72:412–434 Distler VV, Kryachko VV, Yudovskaya MA (2008) Ore petrology of chromite-PGE mineralization in the Kempirsai ophiolite complex. Mineral Petrol 92:31–58 Houlé MG, Lesher CM, Davis PC (2012) Thermochemical erosion at the Alexo Mine, Abitibi greenstone belt, Ontario: implication for the genesis of komatiite-associated Ni-Cu-(PGE) mineralization. Miner Deposita 47:105–128 Irvine TN (1977) Origin of chromitite layers in the Muscox Intrusion and other stratiform intrusions: a new interpretation. Geology 5:273–277 Jonsson E, Troll VR, Högdahl K et al (2013) Magmatic origin of giant ‘Kiruna-type’ apatite-iron-oxide ores in Central Sweden. Nature Sci Rep 3:1644. 7 https://doi.org/10.1038/srep01644 Jourdan F, Reimold WU, Deutsch A (2012) Dating terrestrial impact structures. Elements 8:49–53 Keays RR, Lightfoot PC, Hamlyn PR (2012) Sulfide saturation history of the Stillwater Complex, Montana: chemostratigraphic variation in platinum group elements. Miner Deposita 47:151–173 Le Bas MJ, Keller J, Tao K et al (1992) Carbonatite dikes at Bayan Obo, Inner Mongolia, China. Mineral Petrol 46:195–228 Lightfoot PC (2016) Nickel Sulphide Ores and Impact Melts: Origin of the Sudbury Igneous Complex. Elsevier, Amsterdam Locmelis M, Melcher F, Oberthür T (2010) Platinum group element distribution in the oxidized Main Sulfide Zone, Great Dike, Zimbabwe. Miner Deposita 45:93–109 Mukwakwami J, Lafrance B, Lesher CM et al (2014) Deformation, metamorphism and mobilization of Ni–Cu–PGE sulphide ores at Garson Mine, Sudbury. Miner Deposita 49:175–198 Naldrett AJ, Wilson A, Kinnaird J et al (2009) PGE tenor and metal ratios within and below the Merensky Reef, Bushveld Complex: implication for its genesis. J Petrol 50:625–659 Naldrett AJ, Wilson A, Kinnaird J et al (2012) The origin of chromitites and related PGE mineralization in the Bushveld Complex: new mineralogical and petrological constraints. Miner Deposita 47:209–232 Oberthür T, Davis SW, Bleckinsop TG, Höhndorf A (2002) Precise U-Pb mineral ages, Rb-Sr and Sm-Nd systematics fort he Great Dyke, Zimabwe—constraints for crustal evolution and metallogeny of the Zimbabwe Craton. Precambr Res 113:293–306 Polovina JS, Hudson DM, Jones RE (2004) Petrographic and geochemical characteristics of postmagmatic hydrothermal alteration and mineralization in the J-M Reef, Stillwater Complex, Montana. Can Mineral 42:261–277 Prichard HM, Barnes S-J, Maier WD, Fisher PC (2004) Variations in the nature of platinum-group minerals in a cross-section through the Merensky Reef at Impala Platinum: implications for the mode of formation of the reef. Can Mineral 42:423–437 Schneiderhöhn H (1958) Die Erzlagerstätten der Erde. Vol I: Die Erzlagerstätten der Frühkristallisation. Gustav Fischer, Stuttgart Scoates JS, Friedman RM (2008) Precise age of the platiniferous Merensky Reef, Bushveld Complex, South Africa, by the U-Pb zircon chemical abrasion ID-TIMS technique. Econ Geol 103:465–471 Scoon RN, Mitchel AA (2010) The principal geological features of the Onverwacht platiniferous dunite pipe, eastern limb of the Bushveld Complex. South African J Geol 113:155–168 Scoon RN, Mitchel AA (2011) The principal geological features of the Mooihoek platiniferous dunite pipe, eastern limb of the Bushveld Complex, and similarities with replaced Merensky Reef at the Amandelbult Mine, South Africa. South African J Geol 14:15–40

369 References and Suggestions for Further Reading

Seabrook CL, Prichard HM, Fisher PC (2004) Platinum-group minerals in the Raglan Ni-Cu-(PGE) sulfide deposit, Cape Smith, Quebec, Canada. Can Mineral 42:485–497 Stanton RL (1972) Ore Petrology. McGraw-Hill, New York Stone WE, Beresford SW (2004) New frontiers in research on NiS-PGE mineralization: introduction and overview. Mineral Petrol 82:179–182 Stone WE, Heydari M, Seat Z (2004) Nickel tenor variations between Archean, komatiite-associated nickel sulphide deposits, Kambalda ore field: the metamorphic modification model revisited. Mineral Petrol 82:295–316 U.S. Geological Survey (2020) Mineral Commodity Summaries 2020, U.S. Geological Survey, 200 p., 7 https://doi.org/10.3133/mcs2020 Wagner PA (1929) The Platinum Deposits and Mines of South Africa. Oliver and Boyd, Edinburgh Wu F-Y, Yang YH, Li QL et al (2011) In situ determination of U-Pb ages and Sr–Nd–Hf isotopic constraints on the petrogenesis of the Phalaborwa carbonatite complex, South Africa. Lithos 127:309–322 Yang K-F, Fan H-R, Santosh M et al (2011) Mesoproterozoic carbonatite magmatism in the Bayan Obo Deposit, Inner Mongolia, North China: constraints for the mechanism of super accumulation of rare earth elements. Ore Geol Rev 40:122–131 Yudovskaya MA, Kinnaird JA (2010) Chromite in the Platreef (Bushveld Complex, South Africa): occurrence and evolution of the chemical composition. Miner Deposita 45:369–391 Zeh A, Ovtcharova M, Wilson AH, Schaltegger U (2015) The Bushveld Complex was emplaced and cooled in less that one million years— results of zirconology, and geotectonic implications. Earth Planet Sci Lett 418:103–114 Zieg JM, Marsh BD (2005) The Sudbury igneous complex: viscous emulsion differentiation of a superheated impact melt sheet. Geol Soc Am Bull 117:1427–1450 Further Reading Arndt NT, Lesher CM, Czamanske GK (2005) Mantle-derived magmas and magmatic Ni-Cu-(PGE) deposits. In: Hedenquist JW, Thompson JFH, Goldfarb RJ, Richards JP (eds) Economic geology, One Hundreth anniversary volume, pp 5–23 Ashwal LD (1993) Anorthosites. Springer-Verlag, Berlin, Heidelberg, New York, Tokyo Barnes S-J, Lightfoot PC (2005) Formation of magmatic nickel-sulfide deposits and processes affecting their copper and platinum group element contents. In: Hedenquist JW, Thompson JFH, Goldfarb RJ,

Richards JP (eds) Economic geology, One Hundreth anniversary volume, pp 179–213 Barnes SJ, Holwell DA, Le Vaillant M (2017) Magmatic sulphide ore deposits. Elements 13:89–95 Cawthorn RG, Barnes SJ, Ballhaus C, Malitch KN (2005) Platinum group element, chromium, and vanadium deposits in mafic and ultramafic rocks. In: Hedenquist JW, Thompson JFH, Goldfarb RJ, Richards JP (eds) Economic geology, One Hundreth anniversary volume, pp 215–249 Chakhmouradian AR, Zaitsev AN (2012) Rare earth mineralization in igneous rocks: sources and processes. Elements 8:347–353 Evans AM (1993) Ore Geology and Industrial Minerals, 3rd edn. Blackwell Science, Oxford Guilbert JM, Park CF (1986) The Geology of Ore Deposits, 4th edn. Freeman, New York Keays RR, Lightfoot PC (2004) Formation of Ni-Cu-platinum group element sulfide mineralization in the Sudbury impact melt sheet. Mineral Petrol 82:217–258 Kruger FJ (2005) Filling the Bushveld magma chamber: lateral expansion, roof and floor interaction, magma unconformities, and the formation of giant chromitite, PGE and Ti-V-magnetite deposits. Miner Deposita 40:451–472 Kynicky J, Smith MP, Xu C (2012) Diversity of rare earth deposits: the key example of China. Elements 8:361–367 Laznicka P (2006) Giant Metallic Deposits—Future Sources of Industrial Metals. Springer-Verlag, Berlin, Heidelberg, New York Lee W-J, Wyllie PJ (1998) Processes of crustal carbonatite formation by liquid immiscibility and differentiation, elucidated by model systems. J Petrol 39:2005–2013 Maier WD, Groves DI (2011) Temporal and spatial controls on the formation of magmatic PGE and Ni-Cu deposits. Miner Deposita 46:841–857 Marsh BD (2006) Dynamics of magmatic systems. Elements 2:287–292 Naldrett AJ (2005) A history of our understanding of magmatic Ni-Cu sulfide deposits. Can Mineral 43:2069–2098 Naldrett AJ (2010) From the mantle to the bank: the life of a Ni-Cu(PGE) sulfide deposit. South African J Geol 113:1–32 Pirajno F (2004) Hotspots and mantle plumes: global intraplate tectonics, magmatism and ore deposits. Mineral Petrol 82:193–216 Pohl WL (2011) Economic Geology—Principles and Practice. Wiley-Blackwell, Chichester Oxford

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Pegmatites 22.1 Theoretical Considerations – 372 22.2 Field Relations, Petrography and Petrogenesis of Pegmatites – 373 22.3 Pegmatites as Sources of Economic Minerals – 375 22.4 Geochemical Classification of Granitic Pegmatites – 376 References and Suggestions for Further Reading – 377

© Springer-Verlag GmbH Germany, part of Springer Nature 2020 M. Okrusch, H. E. Frimmel, Mineralogy, Springer Textbooks in Earth Sciences, Geography and Environment, https://doi.org/10.1007/978-3-662-57316-7_22

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Pegmatites are intrusive igneous rocks of exceptionally coarse grain size, in which individual crystals can reach as much as tens of metres in size. Examples of giant crystals reported from pegmatite are a microcline measuring 49.4 × 36 × 13.7 m and weighing nearly 15,000 tons or a phlogopite plate with a size of 10 × 4.3  m (Rickwood 1981). Pegmatites form from residual silicate melts that are enriched in H2O, OH−, CO2, HCO32−, CO32−, SO42−, PO43−, H3BO3, F−, Cl− and other volatile or semi-volatile components. In principle, any plutonic rock can display pegmatitic structure, an example being very coarse-grained varieties of mafic to ultramafic rocks of the Merensky Reef in the Bushveld Igneous Complex (7 Sect. 21.3.1). The overwhelming majority of pegmatites are, however, of granitic composition that closely approaches the ternary minima or eutectics in the system Qz–Or–Ab–(An)–SiO2. In the following the term pegmatite will be understood as describing a very coarse-grained felsic plutonic rock, whereas other rocks with pegmatitic structure will be referred to as pegmatoid. Essential constituents of pegmatite are the same as in granite, that is quartz, feldspars, mainly microcline and albite, and possibly muscovite, with subordinate biotite, garnet, tourmaline, apatite and Fe–Ti oxides. Much scarcer, although of great economic relevance, are chemically complex granitic pegmatites that contain high amounts of rare elements, such as Li, Rb, Cs, Be, Sc, Y, REE, Nb, Ta, W, Sn

and U. Due to their large ionic radii and/or their high field strengths, these elements are not incorporated, or are present only in traces, in the crystal structures of the rock-forming minerals. Consequently, these incompatible elements are concentrated in residual aqueous melts, thus making granitic pegmatite an important source for a wide range of raw materials. Also of high economic significance are syenitic pegmatite and nepheline-syenitic pegmatite that consist of alkalifeldspar, alkali pyroxene, biotite, amphibole and nepheline or minor quartz plus a variety of rare minerals. 22.1  Theoretical Considerations

In order to understand the development of a granitic magma down-temperature to the pegmatitic stage, we may start with the simplified granite system Qz–Or–Ab–(An)– H2O (7 Sect. 20.2). Fundamental is the solubility of H2O in the granitic melt. The T-X diagram granite-H2O shown in . Fig. 22.1 was experimentally determined by Whitney (1975) in a synthetic, Fe-Mg-free model granite consisting of 26.5% Qz, 34% Or, 32% Ab and 7.5% An at a constant total pressure of 2 kbar. Displayed on the left side of this diagram are the liquidus curve A–B–C, the solidus curve D–E, the solubility curve for H2O in the melt F–B, and the border line B–D between H2O undersaturation and H2O oversaturation. The maximum solubility of H2O in the granitic melt,

. Fig. 22.1  a T-X diagram for the system granite−H2O at Ptot = 2 kbar; for explanations see text (after Whitney 1975); b H2O contents in melt inclusions in pegmatitic quartz from Ehrenfriedersdorf, Saxonian Erzgebirge (Ore Mountains), Germany, melt A displays prograde, melt B retrograde solubility for H2O; above the critical point at 712 °C only one homogeneous, supercritical liquid phase exists, rich in H2O (after Thomas et al. 2000)

22.2 · Field Relations, Petrography and Petrogenesis of Pegmatites

reached at 840 °C (at Ptot = 2 kbar), is about 6.5% (point B). With rising temperature, the solubility of H2O decreases, as indicated by the negative slope of the curve B–F, a fact known as retrograde solubility. On the other hand, the H2O content of the melt cannot be quantified from the border line B–D because the diagram does not describe a true binary system but only a binary section through the multi-component system granite–H2O. Schematically presented is also the retrograde solubility of the non-volatile components of granite in water vapour G-C-E which decreases with increasing temperature. Both curves intersect at the critical point of the system, above which the melt and vapour phase cannot be distinguished anymore, but only one homogeneous fluid phase exists (7 Sect. 18.1). However, experimental results in the simple model system Ab–H2O clearly indicate that, in the Earth’s crust, super-critical behaviour of granitic melts in the pure model system Qz–Or–Ab–(An)–H2O can be expected only at unrealistically high temperatures of far above 1500 °C. Not until pressures conforming to depths of the Earth’s upper mantle are attained, the critical temperature falls below 1000 °C (e.g., Paillat et al. 1992; Sowerby and Keppler 2002).

When a magma devoid of H2O reaches the liquidus curve at 1180 °C (point A), plagioclase crystallises, followed by alkali feldspar and quartz until at 700 °C (point D) the solidus curve is crossed. The same crystallisation sequence applies to a granitic magma with 2% H2O but with a lower liquidus temperature of 1030 °C (. Fig. 22.1a). Upon further cooling, the H2O content of the melt increases and, depending on the exact composition of the melt, the line B–D is crossed and a vapour phase V rich in H2O is liberated to form small bubbles in the melt and fluid inclusions within growing crystals (7 Chap. 12). If the line B–D is crossed at lower temperature (and/or pressure), retrograde boiling takes place—analogous to the effect seen when opening a bottle of soda water (and thus decreasing the pressure of the liquid). Below the line B–C, Pl + L + V coexist and eventually become accompanied by Akf and Qz upon further cooling (. Fig. 22.1a). Below the solidus curve D–E, the last melt would have disappeared, while a vapour phase is left from which late minerals can crystallise. When cooling below the critical temperature, e.g., 374 °C for pure H2O, the vapour condensates to form a hydrothermal solution. According to Jahns and Burnham (1969), granites with pegmatitic texture can form in the lower part of the liquidus curve A–B, where the magma is already very rich in H2O. However, the formation of the typical giant crystals cannot be explained satisfactorily by means of the simplified granite system Qz–Or–Ab–(An)–H2O alone. Of critical importance is the presence of additional volatile and semi-volatile components, such as F−, Cl−, H3BO3, CO2, HCO32−, CO32−, SO42−, PO43−, H3BO3, as well as rare elements such as Li, Rb, Cs or Be, acting as potential flux components. These are either dissolved in the residual granitic magma or concentrated in a separate flux-enriched aqueous silicate liquid, causing drastic changes in the following physical parameters (e.g., London and Morgan 2012; Thomas et al. 2012):

373

5 Liquidus and solidus temperatures decrease considerably. For instance, experiments on the crystallisation of a granitic melt with higher contents of Li, F, H3BO3 and PO42− at 2 kbar revealed a shift of point A to ca. 950 °C, point B to ca. 700 °C and point D to ca. 450 °C (London 1992). Consequently, the T interval between B and D would roughly conform to the temperature range that has been traditionally assumed for the formation of most pegmatites. 5 The solubility of H2O in a granitic melt increases, e.g., at a pressure of 2 kbar, to a maximum of 11.5% at point B. Previously underestimated fluxing agents such as carbonates or bi-carbonates of Li or Na can increase the solubility of H2O to an extraordinary extent (Thomas et al. 2012). 5 The critical temperature, above which melt and vapour homogenise into a single phase, is drastically decreased. Consequently, the existence of super-critical fluids can be suspected at the end of the pegmatitic stage (e.g., Thomas et al. 2000, 2003; Sowerby and Keppler 2002). Super-critical and subcritical behaviour are strongly influenced by load pressure and chemical composition of the fluid. 5 During cooling, the viscosity of pegmatite-forming melts increases from ca. 10−4 poise, conforming to a waterlike or even super-critical behaviour, up to ca. 10 poise, similar to a gel-like substance (Thomas et al. 2012). (For definition of the unit poise see 7 Sect. 16.4). For instance, investigation of melt and fluid inclusions performed by Thomas et al. (2000, 2003) on minerals of a pegmatite from Ehrenfriedersdorf, Saxonian Erzgebirge (Ore Mountains), Germany, revealed two different silicate liquids that coexisted, in equilibrium, during pegmatite formation. At a pressure of 1 kbar and a temperature of 500 °C, melt A contained ca. 2.5%, melt B ca. 47% H2O. With rising temperature, the H2O content of melt A increased (prograde solubility), but decreased in melt B (retrograde solubility). Both solubility curves meet at the critical point at 712 °C and 21% H2O, above which only a single, super-critical melt phase existed (. Fig. 22.1b). The H2O-rich melt was enriched in H3BO3, whereas F− and PO42− partitioned preferentially into the H2O-poor melt (Rickers et al. 2006).

The evidence for two coexisting silicate melts with different H2O contents and strong fractionation of rare elements broadens the model of Jahns and Burnham (1969) and provides realistic hints at the ways of pegmatite formation in nature. Moreover, giant crystals, typical of pegmatite, can only crystallise from a considerably supercooled silicate liquid, an important prerequisite that will be discussed in the following section (e.g., Simmons and Webber 2008). 22.2  Field Relations, Petrography

and Petrogenesis of Pegmatites

Pegmatitic magmas intrude into plutons, from which they are derived, and adjacent country rocks, filling opening fractures or larger cavities, forming dykes or stocks, partly of considerable size. Pegmatite dykes are variably developed,

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5 Specific mineral assemblages can form zones that are oriented parallel to the contacts. 5 Monomineralic zones can be developed in many cases. The most common example are core zones of pure quartz, finally growing as euhedral rock crystal into the remaining miarolitic cavities (. Fig. 22.2).

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Almost all of these features can be also present in hydrothermal veins that crystallised from a predominantly aqueous fluid (e.g., . Fig. 23.8). The only texture exclusively present in pegmatite and thus of high petrogenetic significance is graphic granite. Its formation was explained by two contrasting models: (1) Simultaneous crystallisation of quartz and alkali feldspar from a residual melt, or (2) selective replacement by reaction with hydrothermal fluids (. Fig. 22.3). Although an igneous origin according to model (1) is preferred by most petrologists, it should be noted that, in many graphic granites, the quartz:feldspar ratio is lower than experimentally determined for the cotectic lines in the simplified granite system Qz–Ab–Or–H2O at low pressures (. Fig. 20.3).

. Fig. 22.2  Pegmatite dyke, Mursinka, Urals, with miarolitic cavities in its central zone; thickness of the dyke about 2 m (after Betechtin 1953, from Schneiderhöhn 1962)

commonly with pinch-and swell structure, of bulgy or lensoid, rarely tabular shape (. Fig. 3.9). Commonly, they penetrate their wall rock discordantly. In other cases, concordant and discordant contacts alternate, adapting to preexisting structures of the country rock. Many pegmatite dykes are located in the marginal part, especially the roof, of a granite pluton and its surroundings. Their maximum distance from their parent pluton can be as much as tens of kilometres (Linnen et al. 2012). Many larger pegmatite dykes and stocks display a well-developed zonal arrangement, in which the younger minerals of the inner zones can grow between or replace the older ones of the outer zones, but never in reverse. The following structural and compositional features are characteristic (e.g., London and Morgan 2012): 5 In many cases, the marginal parts are fine-grained, forming outer zones of aplite, followed by graphic granite, an intimate intergrowth of quartz and microcline or microcline-perthite resembling the cuneiform script of the ancient Orient, giving rise to the name pegmatite (Greek π´ǫγνυµι = to bind together) (. Fig. 22.3). 5 The grain size coarsens in inward direction. 5 Elongate crystals occur perpendicular to the contact of the pegmatite body.

Fenn (1986) was the first to show by dynamic crystallisation experiments that graphic quartz–feldspar intergrowths can form when H2O-undersaturated granitic liquids are cooled well below their liquidus temperatures, for instance, at an undercooling ΔT of 145–165 °C at liquidus temperatures of 700–750 °C and pressures of 3–4 kbar. With increasing ΔT, the driving force for crystal formation from the melt also increases, whereas crystal nucleation and attainment of equilibrium are inhibited, due to the increasing viscosity of the melt. This impedes the diffusion of the chemical components. Upon cooling, the effects of supersaturation versus increased viscosity compete with each other, resulting in a slight delay in nucleation and thus in the formation of stable crystals (. Fig. 22.4). Under these conditions, the growth rate is much faster than for individual crystals because the graphic quartz-feldspar intergrowths are close in composition to the granitic liquid, although not necessarily in exactly cotectic proportions. In summary, graphic granite forms from flux-bearing but not flux-rich liquid of very high viscosity in a rapidly advancing crystallisation front (London and Morgan 2012). In contrast, the inner zones of pegmatites are dominated by exceedingly coarse-grained blocky textures. According to the model of London (2005, 2009), these crystallised from flux-enriched aqueous melts of lower viscosity and higher degrees of undercooling. Similar to the metallurgical process of constitutional zone refining, fluxing components as well as rare elements are concentrated in a boundary layer of liquid that advances along the front of crystal growth towards the inner zones of the pegmatite body. Constitutional zone refining or zone melting is an important process that considerably improves the purity of crystals synthesised on experimental or industrial scale. A melting zone is produced by a mobile induction furnace and moved through the crystal, thereby pushing all the impurities towards its upper or lower end.

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22.3 · Pegmatites as Sources of Economic Minerals

. Fig. 22.3  Pegmatite with graphic structure formed by intergrowths of microcline (light pink) and quartz (medium to dark grey), resembling the cuneiform script of ancient Egypt; Kniebreche ravine near Glattbach, Spessart, NW Bavaria, Germany

Consequently, the growth of giant crystals of K-feldspar, mica, beryl or spodumene is due to exceptionally favourable conditions of seed-crystal selection and crystal growth. In this regard it is interesting to note that, in contrast to former assumptions, relatively short times are estimated now-

. Fig. 22.4  Relationship between undercooling below thermodynamic crystallisation temperature and delay in crystal nucleation, i.e., time interval (in hours) between the onset of undercooling ΔT and the first appearance of macroscopic crystals; as an example, the first crystallisation of K-feldspar from a H2O-saturated granitic melt (~Ab38Or29Qtz33) at 2 kbar is shown by the solid curve; also indicated are the contours for various volumetric proportions of crystals as a function of time (London and Morgan 2012)

adays for cooling and crystallisation of pegmatites, which would explain the undercooling of the liquid (e.g., London 2005, 2009 ; Simmons and Webber 2008; London and Morgan 2012). For instance, a short period of not more than 3–5 months has been estimated for the time between

intrusion and cooling below the solidus temperature in the case of the 20 m thick Harding pegmatite in New Mexico (London 2005). For an only 2 m thick pegmatite dyke at Ramona, California, a cooling period of not more than 25 days was estimated, and the famous, 30 cm thick gemstone-rich pegmatites at the Himalaya Mine, San Diego County, California, probably did not take more than a week to cool down to solidus temperature. High-grade metamorphic rocks commonly contain rock domains that resemble pegmatite in appearance but mostly lack sharp boundaries with the surrounding country rock. They are unrelated to granite plutons or any other igneous intrusion but represent felsic constituents of migmatites, known as leucosomes. They consist of quartz and feldspars in cotectic proportions (. Fig. 20.3) whereas biotite and other metamorphic minerals are subordinate. Leucosomes are formed by anatexis, that is partial melting, during highgrade metamorphism. 22.3  Pegmatites as Sources of Economic

Minerals

Many pegmatite bodies are of considerable economic significance. They can serve as sources for important industrial minerals, such as feldspars and micas (e.g., Černý 1982; London and Kontak 2012; Glover et al. 2012) as well as for gemstones (Simmons et al. 2012). Moreover, during formation of pegmatite, rare elements can be concentrated, such as Li, Rb, Cs, Be, Sr, Ba, B, Sc, Y, REE, Nb, Ta, Zr, Hf, P, Th and U as well as Sn, Mo and W, in cases even Cu and Au. Pegmatite provinces are regional accumulations of pegmatite deposits with specific mineral content. Based on industrial minerals or metallic raw materials produced from pegmatites, the following types can be distinguished (output numbers are from the U.S. Geological Survey 2020): 5 Most common and mined worldwide are feldspar pegmatites, also known as “ceramic pegmatites”. They are “barren”, i.e., virtually devoid of characteristic accessories such as gemstones or rare metal minerals. About 85–90% of all feldspar mined worldwide serves as raw material for various ceramics, including chinaware, and of glass, whereas the rest is used for a variety of technical applications, such as PVC plastics, paint, abrasive fillers, and welding rods. From more than 50 feldspar-producing countries, Turkey (7.5 Mt), Italy (4.0 Mt) and India (4.0 Mt) rank first, followed by China (2.0 Mt), and Thailand (1.6 Mt), together adding up to about 73% of the world production of 26 Mt in 2019. 5 Mica pegmatites contain large thin flakes of muscovite or phlogopite, which are perfectly flat, flexible and tough. Famous deposits are located in the Black Hills, South Dakota, at Bruth Pine, North Carolina and in the Petaca District, New Mexico, all in USA, in the Uluguru Mountains of Tanzania, in Bengal, northeastern India,

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and in Sri Lanka. Due to their excellent physical properties, muscovite is used, either as flake or finely milled, in electric and electronic industries or as fillers, in coatings, lubricants and cosmetics. In addition muscovite and phlogopite serve as fluid sealant in petroleum drilling (Glover et al. 2012). For 2019, the world production of scrap and flake mica is estimated at 380 kt, of which China alone produced 100 kt, followed by Finland (64 kt), USA (38 kt), Madagascar (36 kt), Canada (23 kt), France (22 kt), and India (16 kt). Natural micas are partly supplemented by synthetic ones. In the past, transparent plates of muscovite, known as “isinglass”, served as window panes in residential buildings or, making use of its thermal resistance, as observation windows in stove doors. In the USA, muscovite, known as “the mineral that won Wold War II”, was used as the first inert, high-temperature resistant, electrically insulating material for manufacturing electrical condensers and vacuum tubes (Glover et al. 2012). 5 Lithium pegmatites are rich in spodumene, LiAl[Si2O6], locally occurring as giant crystals of up to 16 m in size, and/or the lithium micas lepidolite and zinnwaldite (7 Sect. 11.5.2). Some pegmatites, such as those of the Black Hills, South Dakota, and the Kings Mountains, North Carolina, USA, also contain amblygonite, LiAlPO4(F,OH). The light metal Li is used for production of glass and ceramics, as flux in aluminium metallurgy, for production of Li-based solid-state batteries and chemical Li-compounds. One of the biggest spodumene deposits is the Greenbushes pegmatite, Western Australia. Based on this deposit, Australia produced 42 kt of Li in 2019, followed by Chile (18 kt), China (7.5 kt), Argentina (6.4 kt), and Zimbabwe (1.6 kt) (U.S. Geological Survey 2020). As by-products, the metals Be, Rb, Cs, Zr, Ni, Nb, Ta and Sn can be derived from Li pegmatites. 5 Beryllium pegmatites are rich in beryl (. Fig. 11.20), forming giant crystals of up to 16 m in length (Rickwood 1981). In 2019, 170 t Be, about 65% of the world's production, were produced in the USA, although from non-pegmatitic deposits as well. The light metal Be has many technical applications (7 Sect. 11.3). 5 Gemstone pegmatites can contain the precious varieties of beryl, especially aquamarine, of tourmaline (. Figs. 11.24, 11.25), topaz, spodumene, rose quartz and the alkali feldspars amazonite (. Fig. 11.77) and moonstone. Clear crystals than can be cut to gems are obtained, almost exclusively, from miarolitic cavities. Major gem-pegmatite districts are in Brazil (Minas Gerais, Parabaiba), Madagascar, DR Congo, Kenya, Tanzania, Mozambique, Zambia, Zimbabwe, Namibia, Nigeria, USA (California, Colorado, New England), Italy (Island of Elba), Finland, the Ukraine, Russia (Ural, Transbaikalia), PR China (Xinjiang Uygur, Yunan), Mayanmar, Vietnam, India, Afghanistan and Pakistan (Simmons et al. 2012). 5 Pegmatites with uranium and thorium minerals contain predominantly uraninite, U3O8, and thorianite, ThO2. The most important uranium pegmatites currently known are in the Bancroft district, Ontario, Canada. 5 Niobate-tantalate pegmatites, crystallised from residual melts of alkali-granite magmas, contain columbite solid solutions (Fe,Mn)(Ta,Nb)2O6, especially the Ta-rich end member tantalite, as well as REE minerals (see below). The worldwide production in tantalum has enormously increased since the turn of the millennium and has reached about 1800 t in 2019. Ta metal is of high strategic significance because it is used

for the production of miniaturised tantalum capacitors that are inserted in mobile phones, laptops and motor vehicles. The largest Ta deposit of the world has been pegmatites at Wodgina Pan near Greenbushes in Western Australia, which made the country the most important Ta producer for many years. More recently, most of the annual Ta production comes from the African countries DR Congo (740 t), Rwanda (370 t), Brazil (250 t), Nigeria (210 t), and China (100 t). In 2019, the annual production of niobium amounted to 74,000 t, of which 65,000 t were produced in Brazil and 7600 t in Canada, where the most important deposit is the Tanco pegmatite at Bernic Lake, Manitoba. More than 50 mica pegmatite mines in the Petaca District in New Mexico, USA, produce Be, Nb, Ta, Bi, U, Th and REE as by-products. In some areas plagued by civil war, such as the Republic of Congo and adjacent central African countries, the columbite-tantalite concentrate “Coltan” has been extracted by small-scale, informal miners from intensely weathered pegmatites, but also from placer deposits (7 Sect. 25.2.7), to raise funds also for war activities and thus have become known as “blood coltan” (Melcher et al. 2008). 5 Rare-earth element pegmatites are spatially and genetically related to peralkaline A-type granites. In contrast to carbonatites, heavy rare-earth elements are typically concentrated in these pegmatites, which consequently are gaining increasing economic significance because of the comparatively high demand for heavy REE. Important REE minerals in these pegmatites are monazite, CePO4, gadolinite, (Y,Ce,REE)2Fe2+Be2[(SiO4)2]O2, fergusonite, (Y,Ce,Nd)NbO4, euxenite, (Y,Ca,REE)(Nb,Ta)2(O,OH)6, samarskite (Y,Fe3+,U4+) (Nb,Ta)O4, and the complex Zr-silicates eudyalite and elpidite as well as the Y mineral gagarinite NaCaYF6, named after the famous Russian cosmonaut Yuri A. Gagarin (1934–1968; Chakhmouradian and Zaitzev 2012). 5 Zirconium-titanium pegmatites, especially connected with nepheline syenites, contain elevated levels of zircon, titanite CaTi[SiO4]O, and many rare minerals. Well-known are the occurrences at the Langesund Fjord, southern Norway, at Miask, Ural Mountains, Russia, and several places in Greenland. 5 Phosphate pegmatites contain apatite, amblygonite, (Li,Na) AlPO4(F,OH), triphylite, Li(Fe,Mn)PO4, monazite and numerous rare phosphate minerals. An important occurrence is the Varuträsk pegmatite in Sweden. The world-famous, extremely mineral-rich phosphate pegmatite body of Hagendorf, Bavaria, Germany, was mined from 1894 to 1989. 5 Tin pegmatites contain cassiterite, SnO2, wolframite, (Fe,Mn) WO4, and molybdenite, MoS2, in different proportions. Important occurrences are located in the Black Hills, South Dakota, and in Maine, USA, as well as in northwestern Namibia and northern Portugal.

22.4  Geochemical Classification of Granitic

Pegmatites

From a geochemical point of view, three major families of granite pegmatites can be distinguished (Černý and Ercit 2005; Černý et al. 2005, 2012; Martin and De Vito 2005): 1. Pegmatites of the NYF family are enriched in Nb > Ta, Y and F, furthermore in Be, REE, Sc, Ti, Zr, Th and U. Important minerals are topaz, beryl, allanite and xenotime, YPO4, as well as the REE minerals mentioned above

377 References and Suggestions for Further Reading

(Černý et al. 2012). The NYF pegmatites are differentiation products of subaluminous A- and I-type granites that were typically emplaced in an anorogenic environment with extensional tectonics. 2. In contrast, pegmatites of the LCT family are characterised by high concentrations of Li, Cs and Ta as well as Rb, Be, Sn, B, P and F. Predominant minerals are topaz, tourmaline (elbaite), spodumene, petalite, LiAl[Si4O10], lepidolite, amblygonite and other phosphate minerals as well as columbite (Černý et al. 2012). LCT pegmatites are mainly related to peraluminous S-type, more rarely I-type, granites that were emplaced syn- to late orogenically, along convergent plate margins above subduction zones. 3. The mixed NYF + LCT family displays characteristics of both groups.

References and Suggestions for Further Reading Betechtin AG (1953) Lehrbuch der Mineralogie. Verlag Technik, Berlin Černý P (ed) (1982) Short course in granitic pegmatites in science and industry. Mineral Assoc Canada, Winnipeg, Canada Fenn PM (1986) On the origin of graphic granite. Am Mineral 71:325– 330 Jahns RH, Burnham CW (1969) Experimental study of pegmatite genesis. I. A model for derivation and crystallization of granitic pegmatites. Econ Geol 64:843–864 London D (1992) The application of experimental petrology to the genesis and crystallization of granitic pegmatites. Can Mineral 30:499–540 London D, Morgan GB, Hervig RL (1989) Vapor-undersaturated experiments with Macusani glass + H2O at 200 MPa and the internal differentiation of pegmatites. Contrib Miner Petrol 102:1–17 Melcher F, BGR-Gruppe Coltan (2008) Herkunftsnachweis von “Blutcoltan” aus Zentralafrika. GMIT, Geowiss Mitt 31:18–20 Paillat O, Elphick SC, Brown WL (1992) The solubility of water in NaAlSi3O8 melts: a re-examination of Ab-H2O phase relationships and critical behaviour at high pressures. Contrib Miner Petrol 112:490–500 Rickers K, Thomas R, Heinrich W (2006) The behaviour of trace elements during the chemical evolution of the H2O-, B- and F-rich granite-pegmatite-hydrothermal system at Ehrenfriedersdorf, Germany: a SXFR study of melt and fluid inclusions. Miner Deposita 41:229–245 Rickwood PC (1981) The largest crystals. Am Mineral 66:885–907 Schneiderhöhn H (1962) Die Erzlagerstätten der Erde, vol II: Die Pegmatite. Gustav Fischer, Stuttgart

Sowerby JR, Keppler H (2002) The effect of fluorine, boron and excess sodium on the critical curve in the albite-H2O system. Contrib Miner Petrol 143:32–37 Thomas R, Webster JD, Heinrich W (2000) Melt inclusions in pegmatite quartz: complete miscibility between silicate melts and hydrous fluids at low pressure. Contrib Miner Petrol 139:394–401 Thomas R, Förster H-J, Heinrich W (2003) The behaviour of boron in a peraluminous granite-pegmatite system and associated hydrothermal solutions: a melt and fluid-inclusion study. Contrib Miner Petrol 144:457–472 U.S. Geological Survey (2020) Mineral commodity summaries 2020: U.S. Geological Survey, 200 p. 7 https://doi.org/10.3133/mcs2020 Whitney JA (1975) The effects of pressure, temperature, and XH2O on phase assemblages in four synthetic rock compositions. J Geol 83:1–27 Further Reading Chakhmouradian AR, Zaitsev AN (2012) Rare earth mineralization in igneous rocks: sources and processes. Elements 8:347–353 Černý P, Ercit TS (2005) The classification of granite pegmatites revisited. Can Mineral 43:2005–2026 Černý P, Blevin PL, Cuney M, London D (2005) Granite-related ore deposits. In: Hedenquist JW, Thompson JFH, Goldfarb RJ, Richards JP (eds) Economic geology. One Hundreth Anniversary Volume, pp 237–370 Černý P, London D, Novák M (2012) Granitic pegmatites as reflections of their sources. Elements 8:289–294 Evans AM (1993) Ore Geology and Industrial Minerals, 3rd edn. Blackwell Science, Oxford Glover AS, Rogers WZ, Barton JE (2012) Granitic pegmatites: storehouse of industrial minerals. Elements 8:269–273 Linnen RL, Van Lichtervelde M, Černý P (2012) Granitic pegmatites as sources of strategic minerals. Elements 8:275–280 London D (2005) Granitic pegmatites: an assessment of current concepts and directions for the future. Lithos 80:281–303 London D (2009) The origin of primary textures of granitic pegmatites. Can Mineral 47:697–724 London D, Kontak DJ (2012) Granitic pegmatites: scientific wonders and economic bonanzas. Elements 8:257–261 London D, Morgan GB VI (2012) The pegmatite puzzle. Elements 8:263–268 Martin RF, De Vito C (2005) The patterns of enrichment in felsic pegmatites ultimately depend on tectonic setting. Can Mineral 43:2027–2048 Simmons WB, Webber KL (2008) Pegmatite genesis: State of art. Eur J Mineral 20:421–438 Simmons WB, Pezzotta F, Shigley JE, Beurlen H (2012) Granite pegmatites as sources of colored gemstones. Elements 8:281–287 Thomas R, Davidson P, Beurlen H (2012) The competing models for the origin and internal evolution of granitic pegmatites in the light of melt and fluid inclusion research. Mineral Petrol 106:55–73

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Hydrothermal Mineral Deposits 23.1  Basic Principles – 380 23.2  Hydrothermal Impregnation Deposits – 382 23.2.1  Granite-Related Sn–W Deposits – 382 23.2.2  Porphyry Cu-(Mo-,Au-)Deposits – 384 23.2.3  Impregnations with Native Copper (Lake Superior Type) – 386

23.3  Hydrothermal Replacement Deposits – 386 23.3.1  Skarn Deposits – 386 23.3.2  Mesothermal Cu–As Replacement Deposits – 387 23.3.3  Hydrothermal Pb–Ag–Zn Replacement Deposits – 388 23.3.4  Hydrothermal Gold-Pyrite Replacement Deposits (Carlin Type) – 388 23.3.5  Metasomatic Siderite Deposits – 388

23.4  Hydrothermal Vein-Type Deposits – 388 23.4.1  Orogenic Gold-Quartz Veins – 389 23.4.2  Epithermal Au and Au–Ag Veins (Subvolcanic) – 390 23.4.3  Mesothermal Cu Ore Veins – 391 23.4.4  Pb–Ag–Zn Ore Veins – 391 23.4.5  Sn–Ag–Bi Ore Veins in the Bolivian Tin Belt – 392 23.4.6  Veins of Bi–Co–Ni–Ag–U Ore – 393 23.4.7  Telethermal Stibnite-Quartz Veins – 394 23.4.8  Hydrothermal Siderite and Haematite Veins – 394 23.4.9  Non-metallic Hydrothermal Veins – 394 23.4.10  Quartz Veins – 395 23.4.11  Mineralisation in Late-Orogenic Tension Joints – 395

23.5  Volcanogenic-Sedimentary Ore Deposits – 395 23.5.1  Ore Formation by Hydrothermal Activity in the Deep Sea: Black Smokers – 395 23.5.2  Volcanic-Hosted Massive Sulfide-Ore Deposits (VMS Deposits) – 397 23.5.3  Volcanogenic Massive Hg Deposits – 399 23.5.4  Magmatogenic Oxide Ore Deposits – 399

23.6  Non-magmatic Stratabound Hydrothermal Deposits – 400 23.6.1  Sedimentary Exhalative Pb–Zn Deposits (SEDEX Deposits) – 400 23.6.2  Carbonate-Hosted Ore Deposits (MVT) – 401

23.7  Unconformity-Related Uranium Deposits – 401 References and Suggestions for Further Reading – 402 © Springer-Verlag GmbH Germany, part of Springer Nature 2020 M. Okrusch, H. E. Frimmel, Mineralogy, Springer Textbooks in Earth Sciences, Geography and Environment, https://doi.org/10.1007/978-3-662-57316-7_23

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Many mineral deposits do not form from a magma but from hydrothermal fluids, which can bear a genetic relationship to magmatism or can be completely unrelated to any magmatic activity. While the term hydrothermal simply means “hotter than the surroundings”, the term fluid is used by most geoscientists in a rather loose way and encompasses both super-critical fluids and subcritical solutions, depending on temperature, pressure and composition. Hydrothermal mineral deposits can form within the Earth’s crust, at various plate tectonic positions, in the following modes: 5 hydrothermal impregnations, 5 hydrothermal-metasomatic replacements, or 5 hydrothermal ore and mineral veins, which, in many cases, can occur in various combinations. In addition, hydrothermal deposits can form 5 at the sea floor as submarine-volcanogenic or syn-sedimentary exhalative ore deposits, 5 at the Earth’s surface as (subaerial) fumarolic products or precipitates from thermal springs. 23.1  Basic Principles

As we have seen in 7 Sect. 22.1, a H2O-rich vapour phase is released during crystallisation of H2O-bearing granitic magma, a process known as retrograde boiling (. Fig. 22.1a, curve B–D). Moreover, a saline hydrous fluid can form that contains additional volatile components, such as F, Cl, H3BO3 as well as rare chemical elements, especially heavy metals. Type and amount of the chemical components dissolved are of critical importance on the specific processes that lead to different hydrothermal mineral deposits. These react sensitively to changes in pressure and temperature, especially on variations in the ratio of lithostatic vs. hydrostatic pressure, such as in the roof of a crystallising granite pluton or a subvolcanic intrusion. Sudden pressure release caused by tectonic movements can lead to spontaneous retrograde (secondary) boiling.

The term hydrothermal describes nothing else but a solution or super-critical fluid that is hotter than its surroundings, irrespective of its origin. Precipitation of ore minerals from such fluids can lead to the local concentration of these minerals to form ore deposits. In many cases, the ore-forming fluid is derived from a cooling magma. As the vapour phase occupies a distinctly higher volume than a melt, hydraulic fracturing of the country rock can take place when H2O, which behaves in an incompatible way, becomes progressively concentrated in the rest melt. Consequently, vapour can penetrate the marginal parts of the intrusion or the country rock along fissures. Condensation of vapour leads to precipitation of less soluble chemical compounds, such as cassiterite, SnO2, wolframite, (Fe,Mn)WO4, scheelite,

CaWO4, molybdenite, MoS2, chalcopyrite CuFeS2, haematite, fluorite, tourmaline, Li-mica or quartz, all of which can form hydrothermal veins. In their vicinity, marginal parts of granite plutons and the adjacent country rocks can be impregnated or replaced by these minerals. In most cases, these mineralisations are restricted to a limited area, commonly not more than a few hundred metres in vertical extension. Thus underground mining in these mineral deposits is restricted to rather small vertical extents. The above processes are related to igneous activity and take place at relatively high temperatures of >400 °C. Consequently, mineral deposits thus formed have been referred to as late-magmatic hydrothermal deposits (Guilbert and Park 1986). Hydrothermal processes can proceed over a wide range in temperature from some 700 down to 50 °C (Evans 1993). In cases, hydrothermal veins contain banded mineral aggregates displaying spherical, cockade-like or garland-shaped arrangement (. Fig. 23.1). These colloform textures testify to precipitation from hydrothermal solutions of relatively low temperature, though not necessarily of colloidal state (Roedder 1968). According to the temperature of their formation, Lindgren (1933) classified hydrothermal deposits as hypothermal (400–300 °C, in part distinctly higher), mesothermal (300–200 °C), epithermal (200–100 °C) and (tele)thermal deposits ( Ag > Hg > Cu > Pb >  Co > Fe > Zn > Mn. If, for instance, metal-bearing solutions meet sulfide ores, the nobler metals are precipitated as they are more electropositive in character. Source and circulation of hydrothermal solutions  Analyses of oxygen and hydrogen isotope ratios in hydrothermal minerals and their fluid inclusions have revealed that hydrothermal solutions/fluids are not juvenile. Juvenile H2O would have never taken part in the water cycle, implying that it emanated entirely from fractionation of primitive mantle-derived magmas. Most hydrothermal fluids involve H2O that was previously stored in sedimentary rocks, generated by metamorphic dehydration reactions or magmatic differentiation, or was derived directly from seawater or meteoric waters, such as percolating surface waters, groundwater or lakes.

Fluid transport is mainly concentrated along highly permeable zones in a rock complex. Note that the hydraulic permeability coefficient of rocks can be extremely variable and with differences of as much as 13 magnitudes (Cathless 1997). Principally, two types of fluid systems can be distinguished (see Kesler 2005): 5 Unconfined (open) fluid systems develop close to the Earth’s surface, e.g., in sedimentary basins, where precipitations seep into relatively cold, unconsolidated or brittly deformed rock sequences. As the pressure in these containments is nearly hydrostatic, they can be refilled at any time and thus held active. The fluids are renewed either directly by precipitations or by inflow from topographic highs. If such systems are heated at greater depth, the fluids can rise within zones of tectonic weakness, e.g., along faults, thus forming thermal springs, from which calcareous or siliceous sinter can precipitate (7 Sect. 14.5). 5 Confined (canalised) fluid systems are located at greater depths of the Earth and, due the overlying rock column, are subjected to higher lithostatic pressures and higher temperatures, and are situated within the regime of ductile deformation. They are driven by supply of magmatic or metamorphic fluids or by tectonic or sedimentary processes, which create isolated fluid reservoirs. As soon as the fluid pressure exceeds the lithostatic pressure or a fluid reservoir is penetrated by a fault, the fluid is driven into the neighbouring open system (Cox 2005). Commonly, the pore space used by hydrothermal fluids becomes filled with quartz, calcite or anhydrite. Whereas quartz displays prograde solubility at low temperatures, calcite

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and anhydrite generally display retrograde solubility at hydrothermal conditions. Thus, on refilling of open fluid systems, calcite and anhydrite are precipitated at the downward branch of the convection circle, i.e., in the region of higher temperature, whereas quartz crystallises in the ascending branch, within which the fluids cool down (e.g., Rimstidt 1997). Depending on their genesis, hydrothermal solutions/fluids can have a wide range of temperatures and salinities (e.g., Kesler 2005). With a temperature range of 700–350 °C, residual solutions of igneous origin are rather hot and display also highly variable salinities of as much as 70 wt% NaCl eq. in some brines. In contrast, hydrothermal solutions released during metamorphic dehydration reactions (7 Sect. 27.2.2) are usually considerably less saline with up to 6% NaCl eq. and lower temperatures typically in the range of 250–400 °C. Fluid systems escaping on the sea floor of mid-ocean ridges or in back-arc regions have temperatures of 200–400 °C and salinities similar to those of the sea water, i.e., around 3.5% NaCl eq. These hydrothermal solutions have produced volcanogenic massive sulfide deposits (VMS deposits, Sects. 23.5.1, 23.5.2) not only in the geological past but still today. Fluids that formed sedimentary-exhalative deposits (SEDEX, 7 Sect. 23.6.1) commonly have similar temperatures of 200–300 °C but higher salinities of up to 15 wt% NaCl eq. Distinctly cooler (100–200 °C) and highly saline (commonly 15–25% NaCl eq.) are basinal brines forming deposits of the Mississippi-Valley type (MVT, 7 Sect. 23.6.2). Heat source  Whereas solutions of igneous or metamor-

phic origin bring along their own heat content, percolating meteoric water is heated within the Earth’s crust, thus forming convection cells that allow the ascent of hydrothermal solutions. Such heating processes can take place in regions with enhanced geothermal gradient, such as in the roof of igneous intrusions or of magma chambers as well as in regions of extensional tectonics. Moreover, decay of radioactive elements, especially U, Th and K, contained in granites and gneisses can act as heat source.

Sources of metal content  There is convincing evidence that the metals in hydrothermal ore deposits are not always derived from a crystallising magma but can be leached from the country rock. For instance, traces of Pb in feldspar or of Zn and Cu in biotite can be mobilised and concentrated in the hydrothermal solutions. Nowadays, even the Au contents in some of the orogenic-type gold deposits, such as the huge gold-quartz vein system of the Mother Lode in California, USA, or the gold deposit Yellowknife, Canada, are assumed to be derived from metamorphic dehydration of underlying strata (Goldfarb and Groves 2015). Of course, such extensive leaching processes must involve rock volumes much bigger than the hydrothermal deposits derived from them.

In cases of a probable magmatic derivation, plutonic and subvolcanic settings at shallow crustal levels can be distin-

guished among hydrothermal mineral deposits. Volcanic hydrothermal deposits are less frequent but, where present, can be of high economic significance. Others can be hosted by volcanic rocks but did not form by a magmatic process, prime examples being VMS deposits (Sects. 23.5.1, 23.5.2). Commonly, a spatial zonation can be observed that reflects systematic changes in temperature away from the central site of hydrothermal discharge. 23.2  Hydrothermal Impregnation Deposits 23.2.1  Granite-Related Sn–W Deposits z Cassiterite

Tin is dissolved in high-T hydrothermal fluids as hydrous complexes with different ligands such as Cl−, F− and (OH)−. At decreasing temperature or salinity as well as at increasing pH value or redox potential cassiterite crystallises according to the following reactions, at which Sn2+ is oxidised to Sn4+ (Lehmann 1990):

SnCl+ + H2 O + 1/2 O2 ⇋SnO2 + 2H+ + Cl− [23.1] cassiterite SnF+ + H2 O + 1/2 O2 ⇋SnO2 + 2H+ + F− cassiterite

[23.2]

Analogous reactions can be formulated for the crystallisation of quartz from the fluid phase: + − SiCl+ 3 + 2H 2 On ⇋ SiO2 + 4H + 3Cl

[23.3]

+ − SiF+ 3 + 2H2 On ⇋ SiO2 + 4H + 3F

[23.4]

Commonly, the replacement of alkali feldspar in granite by cassiterite is linked to the formation of quartz and white mica by the reaction:

SnCl+ 3 + 3(Na, K)[AlSi3 O8 ] + 2H2 O ⇋ SnO2 + KAl2 [AlSi3 O10 ](OH)2 +

+

+ 6SiO2 + 2Na + 4H + 3Cl

[23.5]



Moreover, the acids thus released, especially HCl and HF, cause the alteration of igneous feldspar to form topaz, quartz and fluorite:

CaAl2 Si2 O8 + 4F− + 4H+ ⇋ Al2 [SiO4 ]F2 + SiO2 + CaF2 + 2H2 O

[23.6]

In addition, tourmaline and Li-mica can form by reaction with BO3- and Li-bearing hydrothermal solutions.

23.2 · Hydrothermal Impregnation Deposits

383

Rocks that formed by transformation of granite in the “pneumatolytic” range transitional between the pegmatitic and hydrothermal stage are called greisen (from German “Greis” = old man), a name traditionally used by Saxonian miners. The process is called greisenisation (=ageing).

Among the primary tin deposits, tin greisen and cassierite veins play the predominant role. Cassiterite typically occurs together with quartz, topaz and/or tourmaline, Li-mica (lepidolite or zinnwaldite) and wolframite. Common accessory minerals are apatite, fluorite, scheelite, molybdenite and haematite. In most cases, tin greisen are located in the roof of granite plutons, in fact representing the youngest generation of a granite province, enriched in SiO2, alkalis and other incompatible elements. Two textural varieties are distinguished, the greisen s.str. is coarse-grained, whereas the fine-grained variety was named “zwitter” (from German = “hybrid”). During greisenisation, feldspars are replaced by topaz or tourmaline, quartz and cassiterite, primary mica by Li-mica. The historic areas of tin mining in Germany, located in the eastern and central regions of the Saxonian Erzgebirge (Ore Mountains), always contain topaz greisen, whereas in Cornwall, SW England, already mined for tin in Roman times, tourmaline greisen dominates. The same holds true for the tin deposits of northern Portugal and NW Spain, presently the richest of its kind in Europe (. Fig. 23.2). In the tin deposit of Altenberg, eastern Erzgebirge, the bulged upper apex of the granite body was largely transformed, down to a depth of 250 m, into fine-grained greisen, the Altenberg Zwitterstock (. Fig. 23.3). It consists of a dense, clustered joint system filled with fine-grained cassiterite, separating diffuse zones of cassiterite impregnations (. Fig. 23.4).

. Fig. 23.3  Cross section of a granite stock at Altenberg, Erzgebirge, Germany, inner parts of which have been transformed into greisen; (after Schlegel, modified by Bauman et al. 1979)

. Fig. 23.4  Impregnations of cassiterite, known as zwitter bands, in very fine-grained greisen of Altenberg, Saxonian Erzgebirge, Germany (after Beck 1903)

. Fig. 23.2  Granite-related hydrothermal mineralization in Cornwall, SW England, UK; zonal arrangement of Sn → Cu → Pb–Zn → Fe ores with increasing distance to the granite contact (from Evans 1993)

The chemistry of the mineralising hydrothermal system is not only indicated by the observed ore paragenesis but could be directly determined by in situ laser ablation ICP-MS analyses (see 7 Chap. 12) of fluid inclusions in vein quartz of Cinovec. These revealed the presence of Fe, Na and elevated Sn concentrations in the high-temperature (400– 370 °C) aqueous fluids (Graupner et al. 2005) (. Fig. 23.5). Tin greisen deposits are widespread across Europe and have been the focus of mining in the Saxonian Erzgebirge, in Cornwall, the French Massif Central and the Iberian Peninsula. They are all related to syn-orogenic Variscan granites. Elsewhere, important tin deposits were formed in the back-arc region of Late-Palaeozoic and Mesozoic active continental margins. Examples are deposits in the Peruvian and Bolivian Andes (see below), the Yukon District in NW Canada, in Korea, in the Jangxi Province, PR China, and

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Chapter 23 · Hydrothermal Mineral Deposits

z Wolframite

23

. Fig. 23.5  Hydrothermal quartz vein with cassiterite, wolframite, scheelite, fluorite and a marginal zone of lepidolite, transecting a greisenised granite; Cínovec, Czech Ore Mountains, Czech Republic (from Beck 1903)

in New South Wales, Australia. Of high economic significance is the Southeast Asian tin belt within a Late-Cretaceous to Early-Tertiary island arc, extending from Myanmar via Thailand and Malaysia to the Indonesian Archipelago, contributing 37% of the world tin production. Most famous are the tin deposits of the Malaian Peninsula and the Indonesian islands of Bangka and Billiton, known as the “tin islands”. At present, however, mining is focussed on secondary placer deposits (7 Sect. 25.2.7). The cassiterite-bearing veins of the Bolivian Andes are described in 7 Sect. 23.4.5. In 2019, the anual world production of tin amounted to 310,000 metric tons, of which nearly 88% came from six countries: PR China (85 kt), Indonesia (80 kt), Myanmar (54 kt), Peru (18.5 kt), Bolivia (17 kt), and Brazil (17 kt) (U.S. Geological Survey 2020).

Wolframite, (Fe,Mn)WO4, is a constant companion of cassiterite in many tin deposits, especially in hydrothermal veins that display relatively simple mineral assemblages with quartz, cassiterite, wolframite and minor tourmaline of the dark schorl variety, closely resembling wolframite in hand specimen. In contrast to cassiterite, wolframite displays columnar habit. Many tin deposits of global economic relevance contain sufficient amounts of wolframite to make them workable for production of W. In the Saxonian and Czech Ore Mountains and in Cornwall, the traditional Sn deposits and their slag heaps have been mined for W since the end of the 19th century. Among the Sn–W deposits in northern Spain and northern Portugal, Panasqueira is one of the largest tungsten mines in the world. On the other hand, hydrothermal tungsten veins devoid of cassiterite are known as well. The mineralogy of the veins is much simpler, commonly consisting of quartz and wolframite. The most important tungsten deposits are located in PR China, which in 2019 produced 70.0 kt W (= 82% of the world production of 85 kt W), followed by Vietnam (4.8 kt), Mongolia (1.9 kt), Russia (1.5 kt), Bolivia (1.2 kt), North Korea (1.1 kt), and Rwanda (1.1 kt) (U.S. Geological Survey 2020). 23.2.2  Porphyry Cu-(Mo-,Au-)Deposits z Porphyry Cu-Ore Deposits, with Variable Amounts of Mo, Au and Ag

High-temperature hydrothermal (hypothermal) impregnation deposits of Cu ore can be located within the upper parts of igneous intrusions, commonly I-type granite, granodiorite, tonalite, more rarely monzonite or diorite. In many cases, these are the subvolcanic or volcanic extensions of the plutons

. Fig. 23.6  Schematic cross section of a porphyry copper deposit; a Alteration zones; b Mineralisation zones (after Sillitoe 1973 and Lowell and Guilbert 1970, from Evans 1993)

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23.2 · Hydrothermal Impregnation Deposits

(. Fig. 23.6), especially trachyandesite, dacite and andesite with typical porphyritic texture. For the latter this style of deposit is referred to as porphyry copper deposit. The Cu ore occurs dispersed, either filling an intricate joint system, known as stockwork ore, or impregnating the pore space of the rock as disseminated ore. In addition, associated volcanic breccia zones can be mineralised as well. The stockwork system can be explained by hydraulic fracturing due to secondary boiling (7 Sect. 23.1), whereas the porosity is a result of late magmatic alteration processes, typical of these deposits (see below). Consequently, a clear petrogenetic relationship exists with the crystallisation of an underlying magma. The similarity of this deposit type with greisen deposits is obvious. Results of geochronological investigations and thermal modelling suggest that commonly individual hydrothermal mineralising “events” took place over time spans on the order of 50,000–500,000 years. However, some of the large porphyry Cu-deposits are the result of a series of individual mineralising events at the same site over several million years (Seedorff et al. 2005). Porphyry copper ores form the largest Cu deposits that provide more than half of the present world production of copper and, consequently, are of highest economic significance. Some ore deposits of this type also produce significant amounts of Mo and Au, both of which can be in places the main metals, and subordinately Ag.

The main ore minerals are chalcopyrite, CuFeS2, enargite, Cu3AsS4, molybdenite, MoS2, and pyrite, FeS2. The deposits are relatively low-grade, with Cu contents of 0.3–2%, and up to 0.15% Mo, 4.3 g t−1 Ag and 1 g t−1 Au. The total tonnage of ore is, however, typically huge, in the range of 500–5000 Mt, and thus these deposits represent vast metal resources (e.g., Fontboté et al. 2017). As a selective extraction of stockwork ore, disseminated ore and country rock is impossible in these deposits, the deposit is typically mined as a whole by huge open-cast mining operations, leaving behind the largest man-made holes on the Earth’s surface, exemplified by the Bingham Mine in Utah, USA.

Characteristic of porphyry copper deposits is the strong late-magmatic alteration of the pluton and its country rock, due to the influence of hydrothermal solutions. To a minor extent, this process is accompanied by replacement phenomena. According to the predominant alteration minerals formed, the following zones can be distinguished from the related igneous body outwards (. Fig. 23.6a): 5 K-feldspar zone with hydrothermal K-feldspar, biotite and chlorite; 5 sericite zone with quartz, sericite, pyrite as well as subordinate chlorite, illite and rutile; 5 argillic alteration zone with kaolinite, montmorillonite and pyrite; 5 propylitic alteration zone with chlorite, epidote, pyrite and calcite.

The Cu-ore can occur within the intrusive body and/or the country rock but is commonly concentrated in the border region between the K-feldspar and sericite zones (. Fig. 23.6b). Pyrite is dominant in the sericite zone and gradually decreases via the argillic towards the propylitic alteration zone. Fluid inclusion and stable isotope studies (. Fig. 33.11d) have revealed that the ore forming solutions are partly derived from the crystallising magma itself, partly from heated meteoric water that convected in the vicinity. Chalcopyrite is precipitated exclusively in the mixing zone of both fluid systems. The magmatogenic hydrotherms are very hot (hypothermal) with temperatures between 700 and 550 °C and contain high amounts of Cl− ions that are of crucial importance for the transport of metals. Most investigators assume that the metal content of the hydrothermal solutions is predominantly juvenile, i.e., not derived from the country rock but from the crystallising magma. Porphyry copper ores are formed at convergent plate margins above subducted oceanic lithospheric plates (. Figs. 29.17, 29.18). The metal combination Cu–Ag–Au is commonly, but not always, related to island arcs, whereas the combination Cu–Mo is related to orogenic zones of continental margins. The number of deposits known ­world-wide is very high. With reserves of 200 Mt Cu Chile contains 23% of the world reserves of copper and thus is the most important Cu producer, contributing 28% of the 2019 production (in total 20 Mt), followed by Peru (2.4 Mt), PR China (1.6 Mt), DR Congo (1.3 Mt), USA (1.3 Mt), Australia (1.0 Mt), Zambia (0.8 Mt), Mexico (0.77 Mt), Russia (0.75 Mt), and Kazakhstan (0.7 Mt) (U.S. Geological Survey 2020). With total reserves of 4,381 Mt at 0.8% Cu and 1,934 Mt at 0.69% Cu, respectively, the Chilean deposit El Teniente and Chuquicamata belong to the most important base-metal deposits in the world. The biggest Cu and Au deposit in the USA is Bingham, Utah, with a total proven and probable resource of 619 Mt at 0.42% Cu, 0.17 g/t Au, 0.035% Mo, and 2.04 g/t Ag. Other Cu deposits of high economic relevance are Morenci and San M ­ anuel-Kalamazoo in Arizona, Butte, Montana (see also 7 Sect. 23.4.3), Santa Rita, New Mexico, Lornex and Valley Copper, Canada, Cananea, Mexico, Cerro Colorado, Panama, Panguna, Papua New Guinea, Sar Cheshmeh, Iran, Kounrad, Kazakhstan, and numerous deposits in North Xinjiang, NW-China (Chen et al. 2012). Important porphyry Cu-deposits in SE-Europe are Recsk in Hungary, with reserves of 10 Mt, as well as Maidan Pek and Bor in Serbia, the latter two being related to subvolcanic intrusions of propylitised andesite. z Porphyry Au–Ore Deposits

Some of the large porphyry copper deposits contain high reserves in gold (Bierlein et al. 2006; Frimmel 2008; Tosdal et al. 2009). Most prominent is the gold deposit of Grasberg in Indonesia with proven and probable reserves of 5 Mt at 2 g/t Au, 3.7 g/t Ag and 1.3% Cu. Thus forming the second largest known gold concentration on Earth. Also worth mentioning are the deposits Ok Tedi (ca. 1130 t Au) and Porgera (ca. 1110 t Au) in Papua New Guinea (e.g., Garwin et al. 2005), Boddington in the Australian Yilgarn

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Chapter 23 · Hydrothermal Mineral Deposits

Craton (ca. 1280 t), Kalmakyr in Uzbekistan (ca. 1300 t), Bingham in Utah (ca. 1000 t), Cananea in Mexico (ca. 1270 t), as well as numerous gold deposits in the South American Andes such as Bajo de la Alumbrera in Argentina, Chuquicamata and La Escondida in Chile, and the Cajamarca District in Peru (e.g., Sillitoe and Perelló 2005). Many of these deposits are related to K-rich calc-alkaline volcanic rocks, e.g., shoshonite (Müller and Groves 2000).

world (. Fig. 4.2). Most researchers assume that hydrothermal solutions rich in Cu sulfide ascended from deeper levels of the Earth, reduced haematite flakes, Fe2O3, finely dispersed in the consolidated lava flow, thus causing crystallisation of native Cu.

z Porphyry Mo-Ore Deposits

Hydrothermal replacement deposits are formed in highly reactive rocks such as limestone and dolostone or their metamorphosed equivalents, that is marble. Reaction of these carbonate rocks with hydrothermal solutions can lead to a ´ μα = body) metasomatic exchange (Greek με´τα = after, σω with hydrothermal solutions. By this process, calcite or dolomite are replaced by ore minerals of Sn, W, Mo, Fe, Mn, Cu, Pb, Zn, Ag, Co, As, Bi, Hg, e.g., according to a reaction of the type

Commonly, molybdenite, MoS2, is present, as accessory ore mineral, in Sn-W greisen and porphyry Cu-ore deposits, from which Mo can be extracted as a by-product. In some places, Mo is, however, the main or even the only metal that can be extracted economically. These are known as porphyry Mo-ore deposits, as they resemble porphyry Cu-ores and are also related to subvolcanic rocks that commonly display porphyritic texture in a ­supra-subduction setting. A prime example is the Climax deposit in Colorado, USA, with a proven reserve of 160 Mt at 0.15 % Mo. In this deposit, the outer zones of a big intrusion of granodiorite porphyry was transformed into a very f­ine-grained greisen, in which an intricate joint system became filled with quartz and impregnated by molybdenite, thus forming a typical stockwork ore. Moreover, cassiterite and wolframite can be present in subordinate, and pyrite in larger amounts. The central parts of the granodiorite porphyry are totally silicified. Similar in texture and petrogenesis is the nearby Henderson deposit with ore reserves of 67 Mt and average contents of 0.17% Mo. Thanks to these two deposits, the USA ranked first as a Mo producer until 2008. Of increasing economic significance, however, became the porphyry Mo ore deposits in the recently detected Xilamulun metallogenetic belt along the northern margin of the North China Craton (e.g., Wu et al. 2011), turning the PR China into the top Mo producer that in 2019 provided 130 kt or 45% of the global production, followed by Chile (54 kt), the USA (44 kt), and Peru (28 kt) (U.S. Geological Survey 2020). 23.2.3  Impregnations with Native Copper

(Lake Superior Type)

Fundamentally different from the porphyry copper ores is an impregnation deposit on the Keweenaw Peninsula in the Lake Superior region, Michigan, USA. There, a thick Proterozoic sequence of consolidated basaltic lava flows with brecciated and scorial surface became hydrothermally impregnated with native copper, chlorite, epidote, various zeolites, prehnite, Ca2Al[AlSi3O10](OH)2, pumpellyite, Ca2(Mg,Fe2+)(Al,Fe3+)2[SiO4/Si2O7](OH)2H2O, quartz and calcite. This process led to high local concentrations of native copper in the form of spectacular dendritic crystal aggregates, some of which reached >20 t in weight, samples of which being exposed in mineral museums all over the

23.3  Hydrothermal Replacement Deposits

SnCl+ 3 + CaCO3 + H2 O ⇋ SnO2 + Ca2+ + CO2 + 2H+ + 3Cl−

[23.7]

These replacement processes can affect large rock volumes but are commonly irregularly distributed and the amount of ore is difficult to estimate. Hydrothermal replacements related to hydrothermal veins are described in 7 Sect. 23.4. 23.3.1  Skarn Deposits

The majority of skarn deposits are formed at the contact between an igneous intrusion and adjacent carbonate rock. This special case of contact metasomatism (7 Sect. 26.6.1) is commonly accompanied by extensive metasomatic exchange of chemical components, leading to an enrichment of the carbonate rocks mainly in Si, Al, Fe and Mg as well as minor and trace elements. During this process, Ca-rich silicates are formed, such as grossular-andradite garnet, diopside-hedenbergite, wollastonite, tremolite-actinolite, epidote, vesuvianite and other silicates that may contain also F, Cl and B. Replacement textures and reaction rims are frequently observed in the skarn minerals that can attain considerable grain sizes. The hard and tough calc-silicate rocks thus formed are called skarn, an old Swedish mining term. Mineral equilibria and fluid inclusions point to initially high formation temperatures of 500–650 °C as well as high salinities of >50 wt% NaCl eq., with mineralising fluids that become progressively cooler (250 million m3 of rock slid down into a huge water reservoir that had been built for energy generation, despite of the warning by geologists. The flood wave that had been produced by this event destroyed the subjacent village of Langarone and killed ca. 2000 inhabitants.

Glaciers play an important role for denudation and transport of in many cases huge rock fragments in high mountains and in polar regions. As displayed in the one-component system H2O (. Fig. 18.1), the melting temperature of water decreases with increasing pressure. Consequently, the bottom of glaciers is commonly molten because of the load of the overlying mass of ice. This facilitates the downhill flow that can reach several metres per year. The tough mush formed by the mixture of water, ice and glacial drift erodes the bedrock not only downwards but also sideways. Thus, the initial V-shaped profile of a valley incised by fluvial action is gouged out to form a U-shaped trough valley, typically seen in mountainous areas of the world that have been glaciated in the past (especially during the last ice ages). The eroded rock material transported by ice can accumulate upon melting of the ice sheet to form big moraine deposits. The most important transport medium is flowing water as in rivers. It collects and transports rock debris produced by subaerial weathering, erosion and sheet wash, until it is deposited in continental or marine catchment basins. The transport distance covered by the clastic components is essentially controlled by their grain size. Whereas the finegrained silt and clay particles suspended in water commonly reach the open ocean or larger inland lakes, most of the coarser grained particles of sand or gravel size are deposited along the way on river beds or in inland depressions. During transport, gradation, sorting and concentration processes take place that account for the composition and relative frequency of sediments. From the sedimentary load, transported along the bottom of the rivers, predominantly

rudites and arenites are formed, whereas argillites originate from fine-grained clay suspensions. Precipitation of ions or ionic complexes dissolved in water give rise to chemical or, indirectly, to biochemical sediments. The debris released by source rock weathering becomes exposed to mechanical and chemical destruction during transport. Depending on the composition of the rock and the transport distance, coarser gravel moving along the bottom of the river predominantly undergoes mechanical grain-size reduction and rounding, a process that is stronger in the beginning but diminishes with further transport. Of course, fragments of harder rock, such as quartzite, have to be transported over a longer distance to become rounded than those of softer pelitic rock such as shale. Minerals of lower hardness and strong cleavage are ground more easily and thus are concentrated in fractions of lower grain size. Additional transport processes take place in marine or terrestrial catchment areas until the transported material reaches its site of final deposition. 25.2.2  Chemical Alteration During Sediment

Transport

When deposited in the ocean, detrital particles undergo considerable alteration until the sediment package is covered by younger deposits and thus becomes shielded from seawater. This chemical alteration, which has been described as subaquatic weathering (Niggli 1952), resembles subaerial weathering, but the alteration minerals formed on the seafloor can be different. For instance, submarine weathering in the continental shelf regions can lead to crystallisation of green grains of the dioctahedral sheet silicate glauconite, ~(K,Na)(Fe3+,Mg,Fe2+)[(Si,Al)4O10](OH)2, which is usually absent in terrestrial deposits. Precipitation of the chemical components dissolved in seawater can involve a multitude of steps and can be strongly influenced by biological and biochemical processes. Of high economic relevance are sedimentary deposits of Fe- and Mn-ore, of base-metal sulfide ore or of phosphate deposits all of which have formed on the seafloor. 25.2.3  Grain-Size Distribution of Clastic

Sediments

Transport and deposition of clastic sediments lead to characteristic grain-size distributions. These are determined, depending on the respective dominant grain size, by macroscopic measuring, sieving, settling, measuring under the binocular microscope or the scanning electron microscope (SEM), or by pipetting. The different grain-size fractions are depicted as simple histograms or as frequency curves, in which the grading fractions are displayed along the x-axis on a logarithmic scale, the respective frequency along the y-axis (. Figs. 25.1, 25.2). A useful kind of presentation is the summation curve that is constructed, starting with the

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environments. A perfectly rounded particle has a roundness of 1.0. The interested reader is referred to textbooks of sedimentary petrology (e.g., Friedman and Sanders 1978; Blatt 1982; Tucker 1981; Pettijohn et al. 1987) for further details. 25.2.4  Diagenesis of Clastic Sediments

The term diagenesis comprises all surface-near processes by which a loose, subaquatically deposited sediment is transformed, at low pressure and temperature, into a solidified sedimentary rock.

25

. Fig. 25.1  Presentation of the grain-size distribution as histogram; the frequency curve is shown as light line, the summation curve as bold line (from Müller 1964)

smallest grain size, by gradually adding the amounts of the next coarser grain-size classes (. Fig. 25.1; Müller 1964; Friedman and Sanders 1978). The summation curve represents the integral of the frequency curve. Each inflection point conforms to a maximum in the frequency curve. The quartile deviations indicate the points at which each 25, 50 and 75% of the grains display a smaller grain size than indicated by these points, which are designated as Q1, Q2 = Md (median) and Q3 (e.g., Friedman and Sanders 1978).

These curves, which illustrate grain-size distribution, make it possible to draw first conclusions as to the transport and depositional conditions of a sediment. For instance, frequency curves of river and dune sands display a good gradation whereas this is much poorer in glacial tills transported by ice, and deposited at the base of inland glaciers (. Fig. 25.2). Besides the grain-size distribution, the roundness of clastic grains, ranging from angular via subangular, subrounded, rounded to well-rounded, is a useful textural feature that can help in deciphering depositional . Fig. 25.2  a Grain-size distribution of a river sand (_____) and a dune sand (- - - - - -); b grain-size distribution of various glacial tills (from Correns 1939)

Diagenesis starts gradually during the deposition of the sediment and changes, without sharp boundary, into metamorphism at increasing pressure and temperature (7 Chap. 26). In the different groups of sediments, varying processes of diagenesis take place and, consequently, the different stages of diagenesis are hardly comparable. All the relevant processes of diagenesis start from the pore space initially present in the sediment. Thus not only the solid mineral and rock particles of the detritus are involved but also the pore solutions and gases contained in the pore space. Upon subsidence of a sediment pile, its pore space is more and more reduced due to the load of younger overlying sediment. During this process, the pore solutions tend to migrate upward. The grains of the sediment come in closer contact to each other thus attaining a denser package and the sediment becomes progressively solidified by compaction. In addition to these mechanical processes, chemical reactions take place between pore solutions and the mineral and rock fragments that are partly dissolved or more or less replaced by newly formed minerals. Remains of the pore space can be filled by newly formed authigenic minerals. In the subsiding sediment pile, clay minerals undergo chemical exchange with the pore solution. With increasing load pressure, pressure solution plays an important role (. Fig. 26.26), provided sufficient resid-

25.2 · Clastic Sediments and Sedimentary Rocks

ual pore space is left to enable the circulation of pore solutions. Pressure solution is higher at grain contacts that are orientated perpendicular or at high angle to the direction of the load pressure, rather than in pressure shadows. Consequently, mineral grains are dissolved selectively along the boundaries that experience the highest stress. The dissolved material is re-deposited in the adjacent pore space where new minerals can grow. At a later stage of diagenesis, the porosity and thus also the permeability of argillaceous sedimentary rocks becomes considerably reduced. Eventually, the pore space vanishes almost completely at greater depths. The reactions, characteristic of diagenesis, are gradually replaced by metamorphic reactions that begin along grain boundaries. As long as these are covered by a thin film of fluid, pressure solution is still possible at metamorphic conditions (see 7 Sect. 26.4.3). Sandstone  In many sandstones, the clastic quartz grains are overgrown by clear quartz crystals, some of which display well developed faces (. Fig. 25.3). The pores can be filled by fine-grained quartz aggregates precipitated from SiO2-oversaturated pore solutions. In numerous sandstones, K-feldspar with typical adularia habit or albite occur as authigenic phases that grew during diagenesis (7 Sect. 11.6.2), part of which form overgrowths on detrital feldspar grains. Moreover carbonate minerals, such as calcite or dolomite, can fill the former pores between detrital quartz grains. In many cases, the CaCO3 content is derived from organic relics such as shell fragments initially embedded in the sand. After their dissolution, CaCO3 can precipitate to form a fine-grained calcareous cement filling the pore space. Commonly, clay minerals, especially kaolinite, are present in sandstone. These are formed during diagenesis, either by direct precipitation from K- and Na-bearing pore solution . Fig. 25.3 Photomicrograph of a sandstone with silicic cement; the rounded detrital grains of quartz, delineated by circles of tiny opaque grains, continued to grow during diagenesis thus virtually closing the initial pore space; middle part of the Buntsandstein Group, Lower Triassic; Quarry south of Marktheidenfeld, Spessart, Germany; Crossed polars, width of view 1.5 mm (Photograph: Joachim A. Lorenz, Karlstein am Main, Germany)

423

or by, occasionally pseudomorphic, transformation of detrital feldspar grains. In many sandstones, trioctahedral or dioctahedral chlorite forms during diagenesis. Different zeolite minerals can form in deeply buried sedimentary rocks of appropriate composition. During diagenesis, pore solutions of suitable chemical composition can precipitate anhydrite, baryte or sulfides and corrode grains of accessory heavy minerals (7 Sect. 25.2.6) that are otherwise resistant against weathering. Argillite  In clays, the compaction, i.e., the compression

caused by load pressure, is more pronounced than in sandstones. Due to their platy habit, clay minerals are initially arranged in a card-house manner (. Fig. 25.5) resulting in a loose open packing and a distinctly larger pore space than found in sands, a structure maintained by electrostatic surface energy. As this is overcome at increasing load pressure, the clay-mineral platelets become more and more oriented parallel to each other and thus undergo much stronger compression than rounded sand grains. Chemical processes during diagenesis of mud predominantly affect the clay minerals, some of which are consumed and substituted by new ones due to alteration or authigenic growth. Especially kaolinite, montmorillonite and other smectites with expandable layer structures become more and more replaced by illite and chlorite. Recrystallisation of poorly ordered illite (or other clay minerals) leads to an increase in grain size and crystal order (crystallinity).

By all of these processes, loose argillaceous sediments are consolidated to form solid sedimentary rocks, such as siltstone or mudstone.

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As a result of diagenesis, many sedimentary rocks of argillaceous grain size or carbonaceous composition (see below) contain concretions. These are hard, compact masses of nodular or flattened-lensoid, more rarely of irregular shape, which consist of mineral aggregates that formed by precipitation from an aqueous pore solution. Some of them grew around a nucleus that can be a fossil remnant, e.g., a shell or bone. Concretions are preferentially formed in sediments displaying a distinct compositional variability and thus can be concentrated in specific beds of a sedimentary rock package. The substances necessary for the growth of concretions are derived from the immediate vicinity. Commonly, concretions in argillites consist of calcite, dolomite, siderite (as constituent of clay ironstone), apatite (in phosphorite), gypsum (forming aggregates of well-developed crystals), pyrite or marcasite. 25.2.5  Classification of Rudites and Arenites z Classification of Rudites

Rudites are classified according to their degree of roundness ρ (rho), as defined by the ratio between the average radius of individual edges and corners (ri), determined by N measurements, and the radius of the maximum inscribed circle R (. Fig. 25.4):

ρ=



(ri /R) N

(25.1)

ρ is described by the following terms: well-rounded (ρ ≅ 1) → rounded → subrounded → subangular → angular → sharp edged or rugged. A loose sediment that consists at >50% of angular mineral and rock fragments with average grain diameter of >2 mm is classified as detritus, and if consolidated, as breccia, whereas a loose sediment consisting of rounded mineral and rock fragments, known as pebbles, is called gravel, and if consolidated,

. Fig. 25.4  Idealised cross section of sediment particles showing radii of individual edges and corners (r1, r2, r3…) and of the maximum inscribed circle (R) defining the degree of roundness P (rho) according to equation (25.1) (after Krumbein 1940, from Friedman and Sanders 1978)

. Table 25.2  Clastic sediments and sedimentary rocks Loose sediment

Diagenetically consolidated sedimentary rock

Rudite

Debris → breccia Gravel → conglomerate

Arenite

Sand → sandstone Dust (dry) Mud (wet)

Argillite

Silt → siltstone Clay → claystone, including mudstone and shale

conglomerate (. Fig. 25.10). As individual coarse components can display different degrees of roundness, the limits between conglomerate and breccia are not sharp. Moreover, monomict and polymict rudites are distinguished depending on whether the clasts comprise only one or more minerals or rock types. For instance, a quartz or a granite conglomerate consists of quartz or granite pebbles, respectively, as predominant constituents. A polymict conglomerate of Miocene age, forming an important member the Alpine molasse is known by the Swiss name “Nagelfluh” (from German Nagel = nail, Fluh = mass of rock). The rock fragments constituting a rudite sediment can be useful for palaeogeographic reconstructions, as they can provide evidence of their source area, whereas their degree of roundness can hint at the transport distance (. Table 25.2). z Classification of Arenites

Sands with average grain size between 2 and 0.02 mm are classified by the type of detrital particles and the proportion of argillaceous matrix and/or cement. Occasionally, the size of sand grains can exceed 2 mm, which leads to transitions between ruditic and arenitic sediments. Most rudites consist of detrital particles of different size and type among which quartz is by far predominant. Many sands consist almost totally of quartz but in other cases, considerable amounts of feldspar and white mica can occur as well. Additional constituents are present in subordinate amount and, commonly, can be only identified under the microscope or after pre-concentration in the laboratory. Arenites with large proportions of rock fragments are distinguished by the pre-fix “lithic” (. Fig. 25.5). The ternary diagram Q (quartz and lydite)–F (feldspar and kaolinite)–M (mica and chlorite), originally suggested by Krynine (1948) has proved useful for a first-order classification of arenites in the field. It was subsequently expanded by Dott (1964) and Pettijohn et al. (1987) to the ternary diagram Q (quartz)–F (feldspar)–R (lithic fragments) with variable amounts of mainly clay-sized matrix being depicted along an additional axis normal to the

25.2 · Clastic Sediments and Sedimentary Rocks

425

. Fig. 25.5  Classification scheme of sandstones in the triangle quartz—feldspar—rock fragments with different amounts of matrix (2.9 g cm−3, can become concentrated from normally accessory amounts to economic levels in so-called placer deposits. Similar to quartz, heavy minerals in sands are resistant to chemical weathering. Thus they can provide information on the . Fig. 25.8  a Greywacke, consisting of constituents with angular grain shapes: quartz (light), and feldspar (clouded) besides rock fragments and small pebbles; Harz Mountains, Germany; width of view c. 4 mm. b Oolitic limestone with concentric and conchoidal calcooids within a calcite matrix; Harliberg near Vienenburg, Germany; width of view c. 2 mm

The mechanical force of flowing (river) water or strong waves and tidal motion in littoral environments leads to further disintegration of sediment, which is followed by grading and sorting of the minerals and rock fragments according to their grain size and density. Thus heavy minerals can be concentrated to form economic mineral deposits, so-called placer deposits (cf. Garnett and Bassett 2005). To form placer deposits, minerals need to have not only a high density, high chemical resistance and high hardness but also a lack of strong cleavage. Heavy minerals in placer deposits can be both precious-stones and ore minerals. Alluvial or fluvial placers are formed by rivers, littoral placers by wave and tidal action and aeolian placers by strong winds. Eluvial placers are deposited close to their source rocks. The cut-off grade (that is the minimum concentration of a given commodity in order for the material to be economic) in placer deposits is typically much lower than in corresponding primary deposits in hard rocks, because they can be extracted at relatively low cost by open-cast mining. Heavy ore minerals of economic significance are: native gold (density ρ = 16–19 g cm−3, depending on the Ag content), native platinum and Pt-alloys (ρ = 17–19), cassiterite, SnO2 (ρ = 6.8–7.1), ilmenite, FeTiO3 (ρ = 4.5–5.0), magnetite, Fe2+Fe23+O4 (ρ = 5.2) and columbite-tantalite (“coltan”),

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(Fe,Mn)(Ta,Nb)2O6 (ρ ≅  5.2–7.8, increasing with Ta/Nb ratio). z Placer Deposits with Precious Metals

25

Gold is highly malleable and this is also reflected in the mode of occurrence of native gold in placers, where it can occur as thin platelets that were shaped by indentations of surrounding sand grains or gravel. More rarely, placer gold forms rounded grains, so-called nuggets, commonly of pea- or nut-size, in isolated cases even up to ~70 kg in weight (7 Sect. 4.1). The latter are, of course, far too big to represent merely mechanically eroded and abraded fragments of primary gold in whatever source rock and must be products of secondary growth within the sedimentary environment. Placer gold is always poorer in Ag than so-called “rock gold” in primary Au deposits. The systematic decrease in Ag in individual gold grains downstream is due to preferred dissolution and diffusion of Ag out of the gold structure. Microscopic investigations reveal that gold nuggets can consist of a detrital core that is surrounded by authigenic overgrowths of gold, thus testifying to some mobility of Au at near-surface, usually oxidising conditions, probably in the form of Au complexes with Cl−, Br−, CN− as ligands. Experimental studies have shown that Au is dissolved in water more easily in the presence of humic acids (derived from the decomposition of humus in soils). In many cases, MnO2 can serve as oxidant. Fluvial gold placers can be present downstream from primary gold deposits. Large placer goldfields with total production of as much as 150 t can be found in the foothills of orogenic belts around the Pacific Ocean, where primary deposits of hydrothermal gold-quartz veins were formed during orogenic processes in the Tertiary (7 Sect. 23.4.1). Subsequently, continuing uplift enabled on-going erosion of the source and thus delivery of placer gold in depositories at lower levels (Evans 1993). Prominent examples, famous for historic gold rushes, are the gold districts of Fairbanks, Alaska, Klondike in the Yukon District, NW Territory and in British Columbia, Canada, in California, USA, in Columbia and neighbouring countries in South America, at the upper reaches of the Yenisei and Lena rivers in Siberia, Russia, and in New Zealand. PGE-rich gold placers are being mined in Primorsky Krai, in the easternmost part of Russia. Fluvial placers of platinum alloys near Nizhny Tagil in the Ural Mountains of Russia have been a major source of Russian PGE production. The largest known PGE placer deposit is at Kondjor in the Republic of Yakutia, Russia. With an annual production of some 5 t PGE it ranks second only after the giant primary Norilsk deposit (7 Sect. 21.3.1; Shcheka et al. 2004). Further occurrences of this type are on the Russian Island of Sakhalin, in Alaska, USA, and Columbia.

z Fossil Gold Placers

Ancient placers, genetically unrelated to current sedimentary cycles, are called fossil placers. The economically most important gold province with placer deposits of this type is the Mesoarchaean Witwatersrand Basin in South Africa (e.g., Frimmel 2014). Since 1886, these deposits produced some 53,000 t of Au, by far more than any other gold province in the world. Although the South African gold production has been declining for years, resources of estimated 44,000 t Au remain at great depth. Thus the Witwatersrand palaeoplacer deposits constitute some 30% of the total gold endowment (past production, known reserves and resources combined) of the Earth’s crust (Frimmel 2002, 2014). The Witwatersrand Basin fill comprises a thick succession of predominantly siliciclastic metasedimentary rocks that range in age from 2780 ± 3  Ma (Frimmel et al. 2005). Elevated Au concentrations are restricted to pyritic quartz pebble conglomerate beds (. Figs. 25.9, 25.10). The quartz-rich matrix contains detrital, syn-sedimentary and post-depositional pyrite and finely dispersed gold. Locally, gold is highly concentrated within so-called carbon seams, which consist largely of kerogen that represents remnants of some of the earliest primitive organisms on Earth. In extremely rare cases, gold is visible to the unaided eye (. Fig. 25.10) and occurs in two different morphological varieties: (1) Rounded gold particles, 0.1– 0.2 mm in size, evidently mechanically deformed by wind or water (. Fig. 25.11, left), and (2) aggregates of irregularly shaped, skeletal gold crystals (. Fig. 25.11, right). While the former represents detrital gold particles, the latter were formed by precipitation from post-depositional hydrothermal fluids. From the spatial distribution of the ore bodies, which are without exception stratiform and bound to conglomerate beds, and the lack of cross-cutting ore bodies, it is inferred that the source of the hydrothermal gold were detrital gold particles that were remobilised within their host rocks over very short distances. This notion leads to the hydrothermally modified palaeoplacer model, which is by now accepted by most workers. It forms the best explanation for one of the largest known geochemical anomalies on Earth, that is the by far largest concentration of gold anywhere in the Earth’s crust. The gold-rich conglomerate beds and associated arenitic rocks represent alluvial delta deposits (. Fig. 25.12), fluvial channel fills and aeolian deflation surfaces along the margin of a former foreland basin (Minter et al. 1993; Frimmel et al. 2005). The extraordinary concentration of gold was favoured by intensive chemical and mechanical weathering under extremely hostile climatic conditions of a land surface that lacked any vegetation. The primary source of the huge amounts of detrital gold in the Witwatersrand Basin has been an unresolved matter of debate for decades. Simple mass balance calculations preclude the derivation

25.2 · Clastic Sediments and Sedimentary Rocks

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. Fig. 25.9 Gold-bearing conglomerate of the Witwatersrand Basin from the West-Driefontein gold mine, South Africa; most of the pebbles consist of quartz and are cemented by a quartz-bearing, weakly metamorphosed matrix containing minor amounts of very fine-grained invisible gold, together with detrital pyrite and uraninite

. Fig. 25.10  Native gold at the bottom of the 2.8 Ga Basal Reef of the Witwatersrand Supergroup, defining wellpreserved cross-bedding with accumulation of detrital gold along foresets and bottomset; Free State Geduld mine, Welkom goldfield, Witwatersrand Basin; length of scale bar = 1 cm; collection of Hartwig E. Frimmel (Cover Photograph of Mineralium Deposita, volume 50, 2015)

of the huge amounts of gold in that basin from pre-existing gold deposits in the hinterland. The currently favoured explanation assumes the entire Archaean landmass at the time as potential source. Under the given environmental conditions, that is, an atmosphere rich in CO2 and CH4 as well as H2S, and essentially free of oxygen (Frimmel 2005), the solubility of Au in surface waters would have been four orders of magnitude higher than in modern meteoric waters, implying a huge Au flux off the Archaean land

(Frimmel 2014). Colonies of first oxygenic photosynthesisers (presumably cyanobacteria) in near-coastal wetlands might have provided a redox trap for the Au-rich waters and thus fixed large amounts of gold on their surfaces. These delicate microbial structures would then have been easily eroded, which in turn would have released finegrained gold to be re-deposited as detrital micronuggets further downstream or in littoral sediments during regressive stages (Frimmel 2014; Frimmel and Hennigh 2015).

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concentrated as heavy mineral in placer deposits (Frimmel 2005; Sverjensky and Lee 2010). Prime examples of Archaean conglomerates that are rich in uraninite are the 3060 Ma Dominion Reef near Johannesburg, the gold-bearing conglomerates in the 2.9–2.7 Ga Witwatersrand Supergroup and the 2.45 Ga Matinenda Formation at Elliot Lake, Ontario, Canada. The latter produced until the closure of the last mine in 1996 a total of 165,000 t U3O8 at an average ore grade of 0.11%. In the Witwatersrand ores (and particularly mine dumps) close to 100,000 t of U3O8 remain, which constitutes ca. 9% of the world’s U resources (Frimmel and Müller 2011).

25 . Fig. 25.11  Contrasting morphological types of native gold from the sample shown in . Fig. 25.10, derived by dissolution of the silicate content in hydrofluoric acid; Left: rounded to disc-shaped particles, transported as sedimentary detritus. Right: hydrothermally remobilised gold; SEM photograph; scale 0.2 mm (from Frimmel 2002)

z Cassiterite Placers

Due to its mechanical and chemical properties, cassiterite, SnO2, is a typical placer mineral that can be concentrated in the vicinity of primary tin deposits. A large proportion of the global Sn production stems from placer deposits. Commonly, cassiterite is accompanied by additional resistant heavy minerals that are also derived from the primary magmatogenic hydrothermal assemblage formed at relatively high temperatures. Important occurrences are in various regions of Southeast Asia, especially in Indonesia, in PR China, in Nigeria and DR Congo. Fluvial placers were mined for tin in Cornwall and the Saxonian Erzgebirge (Ore Mountains) during the Early Middle Ages and, presumably, already in prehistoric times. z Gemstone Placers

. Fig. 25.12  Block diagram illustrating the concentration of placer gold within the thick delta deposits of two Archaean river systems, named after the Steyn Reef and the Basal Reef of the Witwatersrand Basin (from Frimmel 2014)

Of additional economic significance are Os-rich PGE alloys and huge amounts of uraninite present in the Witwatersrand palaeoplacers. Witwatersrand-type deposits are not unique to the Kaapvaal Craton in South Africa, where they occur at several stratigraphic levels ranging in age from 2.9 to 2.6 Ga but are known from almost all cratons, such as the Pilbara Craton of Western Australia, the Dharwar Craton in India, the Sao Francisco Craton in Brazil, the Superior Province in Canada and the West African Craton in Ghana (Frimmel 2014). The known gold endowment in all these other areas has been so far, however, much smaller. z Uraninite Placers

Under reducing conditions, which prevailed in the Archaean world, uraninite is insoluble and thus resistant to weathering, can survive sediment transport and become

Compared to their primary deposits, gemstones of high quality can be concentrated in placer deposits, because fractured and inclusion-rich stones tend to disintegrate during transport. Especially the gem varieties of corundum, ruby and sapphire, are recovered from fluvial placers. Ruby deposits of highest economic significance are present in Myanmar, for example in the region of Mogok, east of the upper reaches of the Irrawaddy River. Other important deposits occur in Sri Lanka, Thailand and Afghanistan. Of considerable economic interest have been diamond placers along the southwestern coastline of Namibia. Since their discovery by a railway worker in 1908 they have been an important source of high-quality stones. While the onshore reserves are largely exhausted, mining focus has shifted towards exploiting deposits on the seafloor off the coast near Oranjemund in southwestern Namibia, where reserves of 1500 million carats have been delineated (1 carat = 0.2  g). 25.2.8  Red Bed Deposits

Deposits of Cu, Ag or U-Ra-V ores of red-bed type occur as stratified impregnations in clastic sedimentary rocks originally deposited in arid climates. Presumably, the elevated metal content is due to leaching of older metal concentrations in the vicinity by groundwater and subsequent

25.2 · Clastic Sediments and Sedimentary Rocks

precipitation in the subsurface. Fossil plants can play an important role as redox trap for this precipitation. Ore minerals in copper ores of red-bed type are low-T α-chalcocite, Cu2S, and similar Cu-sulfides, bornite Cu5FeS4, covellite, CuS, as well as cuprite, Cu2O, malachite Cu2CO3(OH)2 and other secondary Cu-minerals. The silver ores comprise acanthite, Ag2S, native silver and chlorargyrite, AgCl, the uranium-vanadium ores carnotite, K2(UO2/ VO4)2·3H2O, derived from uraninite. Type localities for red-bed deposits are the numerous Cu and related Ag deposits in the southwestern USA. Uranium-radium-vanadium deposits of this type are hosted by sandstone formations of various ages. The main deposits, situated on the Colorado Plateau, are the most important uranium providers of the USA. Of similar origin are U deposits of the roll-front type. They are hosted by sandstone in fluvial channels. The U is derived from the leaching of U-bearing country rocks, such as granitic basement rocks, by oxidising waters. These can be migrating groundwater or orogenic waters that are expelled from an emerging mountain belt into the foreland. Uranium is then transported in its highly soluble form of U6+ until it comes into contact with a redox barrier, which can take the form of a sediment rich in organic matter, sulfides or oil/gas. The typical tongue shape of the redox front, resembling a roll in cross section, is due to higher groundwater flow rate in the centre of a given aquifer (highly permeable rock unit) relative to the marginal domains. Upon contact with the above reducing agents and reduction to U4+ low-temperature uraninite precipitates as UO2. In Europe, the most important Cu deposits of the redbed type are hosted by Lower Permian sedimentary rocks (Rotliegend), which are distributed in depressions north and south of the eastern Sudet Mountains in Poland and the Czech Republic. Of the same type is the Permian “copper sandstone” in the western foreland of the Ural Mountains in Russia. Currently the world’s largest producer of U is Kazakhstan (28%). There, the ore is hosted by Upper Cretaceous to Tertiary sandstones of the Chu-Sarysu and Syrdyra basins, both of which formed originally a single basin before it became divided in the course of the Pliocene uplift of the Karatau Mountain Belt. The U contents are generally very low (0.03–0.05%) but the total resources are enormous. Worldwide sandstone-hosted U deposits account for 19.5% of known resources and thus are second to only the IOCG deposit of Olympic Dam (7 Sect. 23.5.4). 25.2.9  Classification of Argillites

Argillites are sediments consisting of very fine clay particles, deposited from water bodies or by vapour transport. The clay particles can be either relics from weathering and erosion of source rocks or, more

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commonly, authigenic. In addition, more or less decomposed organic substances, relics of carbonaceous or siliceous fossil fragments as well as syn-sedimentary or early diagenetic minerals, such as pyrite or marcasite, can be present. A proper mineralogical classification of argillitic sediments is not possible without estimating the proportion of minerals present by means of X-ray diffraction. For instance kaolinitic, illitic or montmorillonitic clays can be distinguished, based on the relative amounts of these clay minerals. Of additional interest are the contents of quartz, feldspar and other detrital minerals (. Fig. 25.13).

Aeolian dust sediments  can form in regions, where loose or solid rock is exposed to corrosion and deflation by action of wind, especially in deserts, periglacial zones or flood plains of large rivers. From these areas, fine-grained material is continuously blown off by dust storms and transported over distances of thousands of kilometres. Today the most important source regions of desert dust are situated on the northern hemisphere (e.g., Engelbrecht and Derbyshire 2010; Gieré and Querol 2010; Gieré and Vaughan 2013). Particularly well-known are the dust storms of the Sahara Desert that often cause dust falls on the Canary and Cape Verde Islands (. Fig. 25.14), in the Mediterranean or even in Central Europe. The mineral composition of these dusts is controlled by their respective source regions. Desert dust consists predominantly of clay minerals and quartz. Moreover, aeolian dusts can contain a wide variety of other airborne particles such as volcanic glass, sea salt, and materials of biogenic or anthropogenic origin, e.g., combustion-derived carbonaceous particles such as soot. As absorption by airborne particles critically modifies the budget of solar and terrestrial radiation, dust has considerable direct and indirect effects on climate and ecosystems and, consequently, is of high relevance for human health (Gieré and Vaughan 2013).

Loess  is the most important fossil dust sediment, forming extensive blanket deposits, usually less than 30 m thick, giving rise to highly fertile soils. It is commonly unstratified, poorly consolidated, porous and friable. The grain size is similar in all occurrences and ranges from clay to fine sand with a maximum in the silt fraction (. Table 25.1). Loess consists of well-sorted mineral grains, predominantly quartz and feldspar, with considerable amounts of calcite, mica and clay minerals. The yellowish colour is caused by minor contents of Fe-hydroxides. Upon weathering, decalcified loess loam of reddish to dark brown colour is formed. Commonly, loess contains irregularly shaped concretions consisting of calcite, known as loess kindchen or loess dolls. Loess was formed in the Pleistocene by deflation of dust from cold deserts and unconsolidated glacial or glacio-fluvial sediments exposed in periglacial regions. It is widely distributed on the northern hemisphere, such as the Midwest of the USA, Central Europe and around the Hoangho River in China, where it is redeposited by dust storms still today.

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. Fig. 25.13  Driven by the hot, northeasterly trade wind known as Harmattan, a dust cloud is blown off the Sahara Desert for more than 1600 km across the North Atlantic Ocean including the Canary and Cape Verde islands; Satellite photograph taken on 2 March, 2003 (courtesy of NASA, 7 http://visiblearth.nasa. gov/view.php?id=65292; from Gieré and Querol 2010)

Mud  is a slimy, sticky or slippery mixture of water with

clay or silt that has been deposited in inland lakes or the open ocean after transport as suspended load, by fluvial water or wind. In addition, muds can be deposited from turbidity currents. These are density currents, formed by variable amounts of loose material suspended in water. Triggered by storm surges, tsunamis, earthquakes or sediment overload, clay and silt are dispatched mainly in shelf areas at continental margins. When the suspended sediment settles, the sedimentary load is deposited in the sequence of its grain size: sand → silt → clay. By this process, graded bedding is achieved, a structural feature typical of turbidites.

The cyclically developed turbidite structures, known as Bouma cycles (. Fig. 25.15), are of high significance for the interpretation of sedimentation processes, the sedimentary environment, also known as sedimentary facies, and its geotectonic position, and, consequently, are of great interest in petroleum exploration (Bouma 1962).

The suspended terrigenous sediment load, deposited as marine or lacustrine muds, consists predominantly of clay minerals, quartz, feldspars, carbonate minerals and organic components. Terrigenous muds cover about 20% of the

25.2 · Clastic Sediments and Sedimentary Rocks

. Fig. 25.14  Vertical cross section through a Bouma sequence, deposited from a turbidity current. Ideally, five divisions can be distinguished from top to bottom: A graded sandstone with systematic grain-size increase from top to bottom, deposited at high speed from a relatively fluid liquefied cohesionless particle flow. The coarse-grained basal mudstone layer forms a sharp contact with the underlying argillaceous layer E′ (modified after Friedman and Sanders 1978). B Parallel-laminated sandstone deposited at low speed of the upper-flow regime. C Fine-grained to very fine-grained sandstone with ripple cross lamination deposited at low speed in the lower-flow regime of the turbidity current; D faintly laminated mudstone; E argillaceous layer at the top of the sequence, deposited at low velocity from the tail of the turbidity current. At the contact with the overlying sandstone A′ of the next cycle, abundant sole marks can be formed, due to sediment burden

seafloor. Recent mud deposits are found on tidal flats, such as the Cape Cod Bay, Massachusetts, USA, the Moreton Bay, Australia, the Yellow Sea between China and Korea, the Bridgewater and Morecambe bays, UK, and the Wadden Sea along the coasts of Denmark, Germany and the Netherlands. Hemipelagic deep-sea sediments such as the blue and green mud, coloured by Fe sulfide and organic matter or chlorite and glauconite, respectively, cover continental shelf regions and the continental slopes down to depths of 2 km. The pelagic red clay is deposited at depth below ca. 3500 m and contains a relatively high amount of windblown particles such as meteoric and volcanic dust, of ice-rafted debris, manganese nodules (7 Sect. 25.4.4) and fossils. Its reddish-brown colour is caused by Fe- and Mn-oxides, stable only at oxidising conditions. However, the largest proportion of the seafloor is covered by non-detrital, biogenic mud such as radiolarian, diatom or Globigerina ooze. 25.2.10  Diagenesis of Argillites

Diagenetic consolidation of loose dust and mud sediments leads to formation of mudstone, a general term that includes

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the silt and mud grain size fractions, ranging from 0.004 to 0.052 mm (. Table 25.1), or describes an argillitic sediment of variable grain size. Shale is a mudstone displaying a fine lamination that imparts a fissility roughly parallel to bedding. In contrast, slate is a foliated mudstone that has been distinctly deformed during incipient metamorphic overprint leading to a slaty cleavage that, in many cases, is oriented transversally to the bedding. Diagenetic transformations are mainly dependent on the composition of the sediment, the chemistry of its pore water and the thickness of the overlying sediments. The load executed by the overlying package eventually leads to compaction of the mud, during which the porosity is more and more reduced, the water content is squeezed out toward the top and the sheet silicates become aligned and orientated parallel to bedding. The water circulation is slowed down and the pore solution starts to react with the minerals present, especially the clay minerals. Moreover, crystallisation of new authigenic silicate minerals can take place. A prominent example is authigenic alkali-feldspar which, despite of the low temperature of formation, can display all transitions from metastable disordered to totally ordered Si–Al distributions: albite (high) → albite (low), adularia → microcline. Depending on the additional phases present, calcareous, siliciclastic and bituminous mudstones are distinguished: Calcareous mudstones  Among this group, marls, i.e., mixtures of lime and clay, are most frequent. Marly sediments can contain a detrital carbonate content, washed in together with detrital silicates. However, carbonates of biological origin, such as planctonic lime shells, or formed by (bio-) chemical precipitation are far more common. Marls and marly sediments or artificial lime-clay mixtures of variable composition are important raw materials, e.g., for production of Portland cement. Bituminous mudstones  such as black shales and oil shales are well-stratified sediments of dark grey to black colour that always contain pyrite and a considerable amount (usually > 5 vol%) of organic carbon. Examples are the Lower Cambrian black shales along the margin of the Yangtze Platform in South China, the Palaeozoic graptolite shales in the Lake District and southern Scotland, Sweden and Iran, the Ordovician oil shale (kukersite) in northern Estonia, the Liassic posidonia shale of the Swabian and Franconian Alb in Germany, and the Eocene oil shale in Hessonia, Germany. Bituminous mudstone forms under anoxic conditions, similar to those presently prevailing in the Black Sea. During sedimentation, the detrital mineral particles and dead plankton subside down to deeper water levels, where a lack of water circulation and the quick consumption of O2 by decomposing organic material lead to overall anoxic conditions in an H2S-rich environment. At these anaerobic conditions, slow biochemical decay and transformation of organic matter, mainly of planktonic origin, takes place. By bacterial activity, SO42− dissolved in the seawater and in the pore solution as well as the sulfur content of proteins is reduced to form H2S that reacts with Fe-bearing detrital minerals. By this process, small circular aggregates of

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pyrite, FeS2, can form as fillings of the cells of former bacteria. These framboids (from French framboise = raspberry) can be finely dispersed in the sediment or concentrated as concretions. In neutral or weakly acidic seawater, marcasite, FeS2, crystallises instead of pyrite. Besides, additional heavymetal sulfides can be formed, such as sphalerite, galenite or chalcopyrite. Under anaerobic conditions the preservation of fossils is facilitated. Consequently, some of the best preserved fossils are found in black shales. Famous examples are Lower Jurassic black shales at Messel in Hessonia, Germany, and the Cambrian Burgess shale in the Canadian Rocky Mountains, both UNESCO World Heritage Sites. Oil shales can be of considerable economic interest as alternative source rock for the extraction of crude oil, although its production is more expensive and raises a number of environmental problems such as waste disposal, waste-water management, air pollution and emission of greenhouse-gases. Among the numerous occurrences worldwide, the largest deposits are part of the Eocene Green River Formation situated in Colorado, Utah and Wyoming, USA, constituting 62% of the world’s resources. However, because of low oil price and higher expenses, their production has been abandoned in 1982 and only restarted in 2003 in the course of the U.S. oil-shale development program. In 2008, the world production amounted to 930,000 t of shale oil, of which 98% was produced by China (375,000 t), Estonia (355,000 t) and Brazil (200,000 t). The last decade has seen a tremendous growth in oil production in the USA, mainly from shale, which now rivals with 7 million barrels per day that of Saudi Arabia and Russia.

25.2.11  Base-Metal Deposits in Black Shales

Stratabound base-metal deposits in some carbonaceous and siliciclastic black shales formed in restricted sedimentary basins, in which large convective hydrothermal fluid systems were active (e.g., Hitzman et al. 2005; Selley et al. 2005). The black shales contain Cu- and Cu-Fe sulfides as well as pyrite and marcasite, finely dispersed or in small veins. The metals were derived either from underlying redbed deposits (7 Sect. 25.2.8) or from the surrounding continents and transported by fluids of moderate to high salinity and low to moderate temperature. The salinity necessary for the transport of most metals probably originates from salts of overlying evaporites (7 Sect. 25.7), deposited on the seafloor or in inland lakes. The source of sulfur is the bacterial reduction of SO42− in seawater. Deposits of this type are widespread, but only a few of them have attained global significance. Together three of such supergiant deposits account for 23% of the global production of Cu and of additional base metals, such as Co and Ag: 5 the Permian Kupferschiefer in northern Germany and southern Poland (Silesia) 5 the Neoproterozoic Zambian Copper Belt in Central Africa 5 the Palaeoproterozoic Kodaro-Udokan Basin in Siberia Somewhat smaller Cu deposits of this type occur in the Paradox Basin of Utah and Colorado, USA. They are located in clay- and siltstones deposited at the turn Jurassic/Cretaceous and are overlain by thick marine evaporites.

z Kupferschiefer

In the Kupferschiefer (copper shale), a thin, bituminous argillaceous marl at the base of the Werra Cycle, the lowest cycle of the Upper Permian evaporitic Zechstein succession, sulfides of Cu, Pb and Zn are concentrated, together with remarkable concentrations of V, Mo, U, Ni, Co, Cr, Ag and many other metals. Ore minerals are chalcocite, Cu2S, chalcopyrite, CuFeS2, bornite, Cu5FeS4, covellite, CuS, tennantite Cu12As4S13, galenite, PbS, sphalerite, ZnS, and pyrite, FeS2. The Kupferschiefer was deposited in the very shallow Zechstein Sea that about 255 Ma ago extended over the peneplained Variscan basement in central and northern Europe. In many areas, the argillaceous sediments form a thin veneer on top of post-orogenic conglomerates and sandstones of Lower Permian age, the so-called Rotliegend (= “red footwall”). The Kupferschiefer is covered by dolomite, limestone and evaporites, mainly anhydrite, of the Werra Cycle (Lower Zechstein). Although a synsedimentary first concentration of many metals is certain, ore grades were achieved only later, during diagenesis and subsequent circulation of hydrothermal fluids in the course of Mesozoic extension in the region and locally again as far-field effect of the Alpine orogeny further south (Bechtel et al. 2001; Borg et al. 2012). The most important mining areas of the Kupferschiefer are situated in the area around Rudna and Lubin in southern Poland. Over the past years, Poland extracted on average some 420 kt Cu annually from these deposits, and thus is Europe’s largest Cu producer. Where the Kupferschiefer continues into Germany, the former mining district of Mansfeld in Saxony-Anhalt has one of the longest continuously documented mining histories, starting in AD 1199 until closure due to exhaustion in 1990. The total Cu endowment (past production + reserves) in both the German and Polish deposits amounts to >60 Mt Cu in >3066 Mt ore, with more than half of it still waiting to be mined from the Polish deposits. z Central African Copper Belt

The Neoproterozoic Central African Copper Belt in Zambia and Katanga, DR Congo, represents one of the largest Cu provinces in the world and at the same time hosts the largest known Co resource on Earth. Many of the deposits in this belt show similarities with the Kupferschiefer but they experienced a greater degree of metamorphic overprint (Hitzman et al. 2005). Both syn- and epigenetic mineralisation can be distinguished, with the latter being essentially related to remobilisation of earlier, syngenetic ores by syn-orogenic saline hydrothermal fluids that were expelled towards the forelands in the course of Pan-African orogeny. In addition, further upgrading of ­ the Cu-ores took place in deep-reaching oxidation zones, especially in the Congolese part of the belt. In 2019 the DR Congo produced 1.3 Mt, Zambia 790 kt Cu, thus taking the 4th and 7th rank, respectively, worldwide.

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25.3 · Chemical and Biochemical Sediments and Sedimentary Rocks …

. Fig. 25.15  Classification of limestones according to their fabric (based on Folk 1962)

25.2.12  Transition from Diagenesis to Low-

Grade Metamorphism

During progressive diagenesis, porosity and permeability of a given sediment, especially argillites, decreases more and more, thus hindering the contact between mineral grains and pore solution. Consequently, solution and precipitation reactions become more and more difficult. Upon further increase of temperature and pressure, a gradual transition takes place from diagenesis to metamorphism. Mineral reactions become concentrated to the grain boundaries, thereby more and more approaching thermodynamic equilibrium (7 Chaps. 26 and 27). The transition of poorly ordered illite to white mica involves a progressively better ordered crystal structure, known as illite crystallinity. This feature, which is largely controlled by temperature, can be quantified by the sharpness of the main reflections in X-ray powder diffractograms. The illite crystallinity, defined by the half width of the 10 Å (001) peak of illite in relation to that of the (101¯1) reflection of quartz, is a good measure for the degree of diagenesis and incipient metamorphism. At the onset of low-grade metamorphism, the sheet silicate pyrophyllite, Al2[Si4O10](OH)2 can be formed at the expense of kaolinite. Moreover, the reflectivity of solid hydrocarbons (coal), observed under the reflected-light microscope, can be used as measure of the degree of diagenesis because it increases with temperature. 25.3  Chemical and Biochemical Sediments

and Sedimentary Rocks1

Weathering solutions containing Ca2+, CO32− and HCO3− ions are transported by rivers, thereby reaching inland lakes or the ocean. Evaporation leads to super-saturation of these chemical components and to purely inorganic precipitation of carbonate minerals and, consequently, chemical carbonate sediments. More

1

With contributions by Gerd Geyer, Würzburg

frequently, however, carbonate precipitation is achieved through mediation by organisms, leading to formation of biochemical carbonate sediments. Lacustrine or limnic carbonate sediments are formed in inland lakes, marine ones in the ocean. Depending on their mode of formation, sedimentary carbonate rocks display a considerable structural and textural diversity. They are distinctly rarer than siliciclastic sediments.

The overwhelming majority of carbonate rocks has been deposited in relatively shallow water of stable marginal seas, where they have formed by chemical or biochemical, in part also by clastic sedimentation. Predominant sedimentary minerals are calcite, aragonite and dolomite, occasionally accompanied by minor amounts of quartz, alkali feldspar and clay minerals. Carbonate sediments consisting of siderite, FeCO3, are much rarer but can be of high economic relevance. 25.3.1  Classification of Sedimentary

Carbonate Rocks

Grain-size classification According to their predomi-

nant grain size, carbonate rocks are classified as calcirudite (>2 mm), calcarenite (2 mm–62 µm) and calcilutite (0.02 mm) sparite. For the particles, the following prefixes are used: Bio for skeletal fragments, oo for ooids (see below), pel for peloids, i.e. particles of lensoid shape, intra for intraclasts such as pebbles or fragments of irregular shape. Combining one or two of these prefixes with the suffixes micrite or sparite results in the respective rock name, e.g., biosparite or bio-oomicrite (. Fig. 25.15). A biolithite is a carbonate rock formed in situ, such as a stromatolith or a reefal

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Chapter 25 · Sediments and Sedimentary Rocks

25 . Fig. 25.16  Classification of limestones according to depositional texture (based on Dunham 1962 and additions by Embry and Klovan 1971)

limestone. A dismicrite is a limestone with window structure, consisting of micrite and cavities filled with sparite. A petrogenetic classification, based on depositional processes is given in . Fig. 25.16. A compositional classification is based on the dolomite/ calcite ratio: 5 limestone: 90% dolomite 25.3.2  Solubility and Precipitation

Conditions of Carbonates

At equilibria between solid CaCO3, aqueous solutions and a CO2-bearing gas phase, the following ions are involved: Ca2+, CO32−, HCO3−, H+, OH−. Of special significance is the partial pressure of CO2, PCO2, in the gas phase, in equilibrium with the adjacent solution. Most of the CO2 is dissolved physically in water, subordinately as H2CO3 according to the reaction

H2 O + CO2 ⇋ H2 CO3

[25.1]

The solution of a calcareous sediment or a limestone in water with low contents of dissolved CO2 can be described by the reaction

H2 O + CO2 ↑↓ CaCO3 + 2H+ + CO3 2− ⇋ Ca2+ + 2HCO3 −

[25.2]

In this reaction, part of the HCO3− anions are derived from Reaction [25.1], the other from the dissociation of H2CO3:

H2 CO3 ⇋ H+ + HCO3 −

[25.3]

The principal Reaction [25.2] describes the process that causes dissolution of CaCO3 during weathering of limestone, frequently leading to karst and cave formation. The

retrograde process refers to the precipitation of CaCO3 from sea- or fresh water, as cement in calcareous rocks or during the growth of stalactites and stalagmites in karst caves. Each of these processes that leads to an increase in CO2 enhances the solubility of CaCO3, whereas decrease in CO2 causes precipitation of CaCO3. Moreover, Reactions [25.2] and [25.3] can explain the influence of the concentration of hydrogen ions, expressed by the pH value. At high pH values, Reaction [25.2] proceeds towards the left side leading to precipitation, at low pH towards the right, triggering dissolution of CaCO3. H2CO3 is a stronger acid than HCO3–. In contrast to most other salt species, the solubility of CaCO3 decreases with rising temperature. Moreover, similar to other gases, CO2 is less easily dissolved in warmer than in cooler water. With increasing pressure the solubility of CaCO3 in water increases slightly, independent of the pressure influence on the solubility of CO2. Only at great depths, the solubility of carbonates is so high that they cannot be stable any more. The carbonate compensation depth (CCD) is the level in the ocean, below which the rate of carbonate dissolution is higher than its precipitation rate.

The CCD is extremely variable. In oceans of tropical climates it is about 3500–5000 m for calcite but about 1000 m deeper for aragonite. Carbonate is precipitated from a solution saturated in CaCO3 if the temperature increases or the PCO2 in the gas phase decreases. On the other hand, limestone dissolves upon temperature decrease and/or PCO2 increase. For example, CO2 can be assimilated by plants, thus leading to an overall decrease in PCO2 and consequently to the precipitation of CaCO3—in this case forming encrustations around plant parts known as calc tufa. Decrease of the partial pressure of CO2, decrease of the concentration of CO2 dissolved in water, and/or increase in temperature all lead to oversaturation and thus the precipitation of CaCO3.

25.3 · Chemical and Biochemical Sediments and Sedimentary Rocks …

25.3.3  Inorganic and Biochemical

Carbonate Precipitation in Seawater

Inorganic precipitation of CaCO3  takes place predominantly in shallow seawater. In addition to precipitated minerals, most marine carbonate sediments contain material of biogenic origin. It has been well known for a long time that the near-surface water of the ocean is saturated in CaCO3 and can even be oversaturated in tropical and subtropical regions. Nevertheless, precipitation of CaCO3 from such saturated and even oversaturated solutions is only possible if certain prerequisites are fulfilled, especially the presence of finely ground calcareous shell particles of marine animals acting as nuclei for crystallisation. In shallow marine environments, calcareous ooids can be precipitated from seawater saturated in CaCO3. These calcite aggregates are spherical to oval in shape and display a concentrically layered, radial structure, attaining diameters of 0.25–2 mm, in most cases 0.5–1 mm. A rock composed of ooids is called oolite (. Fig. 25.8b) and, depending on the grain size of the surrounding calcareous matrix is classified either as oosparite or oomicrite, respectively (. Fig. 25.15). While growing, the ooids float in seawater until they have attained a certain size and settle down to the sea floor forming an oolitic limestone. The concentrically layered structure of the ooids reflects movement of the particles on the shallow seabed due to wave action. Moreover, the outer surface of the ooids indicates later mechanical abrasion during transport and deposition, and some of them can be fractured. Due to continuous supply of seawater saturated in CaCO3 by means of ocean currents, marine carbonate deposits can attain considerable thickness.

Primary inorganic precipitation of dolomite, CaMg(CO3)2, takes place during the formation of evaporite sequences (7 Sect. 25.7.2) or by mixing of salt water and fresh-water in various coastal environments. Biochemical carbonate formation  Formation of carbonate

sediments in shallow seas or on continental shelfs commonly involves a great variety of organisms that build up shells or skeletons of CaCO3. The most important carbonate-producing organisms include prokaryotic calcimicrobes (cyanobacteria), various eukaryotic calcareous algae, and numerous animal groups, such as foraminifera, calcareous sponges, corals, bryozoa, bivalves (. Fig. 25.17), gastropods, echinoderms, and other invertebrates. Predominant reef-forming organisms include calcareous algae, calcareous sponges, corals, and oysters. The earliest organo-sedimentary structures in Earth’s history are so-called stromatolites (. Fig. 25.18), multi-layered calcareous bodies of considerable dimension that project above the sea floor. They are formed in shallow, warm seawater by metabolic activity of cyanobacteria. Stromatolitic build-ups are the oldest known organic remains, reported from as early as Palaeoarchaean times, 3.6 Ga ago. An intensified rate of stromatolite formation occurred towards the end of the Archaean, about 2.6 Ga ago, under rapidly changing atmospheric and

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consequently also hydrospheric conditions. A modern example of stromatolite growth is in the Shark Bay, northwestern Australia (. Fig. 2.11). Due to the reaction of CO2 with the Ca2+ ions dissolved in the seawater, the CO2 content of the Earth’s atmosphere decreased drastically during this period while oxygen was produced in large quantities by the first oxygenic photosynthesising microbes, i.e., cyanobacteria. This enabled not only the first massive precipitation of marine carbonates but also that of thick successions of banded iron formation due to the reaction of the newly released oxygen and large amounts of Fe2+ dissolved in the contemporaneous seawater (7 Sect. 25.4.2). Reefs  are ridge-like or mound-like structures built up predominantly by calcareous exoskeletons of invertebrates, either living in colonies or separately, commonly in combination with calcareous algae, forming a framework that is resistant against wave motion (e.g., Tucker 1981). Reef carbonates can serve as important reservoirs for petroleum and natural gas. In the course of the Earth’s history, nearly all invertebrates contributed at some stage or the other to the formation of reefs, such as stromatoporoids from Ordovician to Devonian times, tabulate “corals” from Silurian to Carboniferous, phylloid algae during the Carboniferous and Permian, modern coral groups since the Triassic, sponges during Triassic and Jurassic times, rudists (bivalves) in the Cretaceous as well as corallinacean algae in recent times. From a functional point of view, the following groups of reef-forming organisms can be distinguished:

5 framework formers, such as corals 5 framework binders that cover and reinforce the reef with a crust, e.g., calcareous algae or bryozoae 5 inhabitants of reefs, such as green algae, annelids, boring clams, gastropods, or echinoderms By far most reefs form in warm seawater, that is at low latitudes. Modern examples are the Great Barrier Reef off eastern Australia, the Red Sea and the Bahamas. A number of stony corals (Scleractinia) are able to live also at greater depths and lower temperatures below 20 °C but rarely form reefs. However, Lophelia forms cold-water reefs of up to 2 km in length and 50 m in height, known from along the European continental fringe from the Iberian Peninsula to the North Cape of Norway at depths of 200–600 m.

Based on their geometry, one can distinguish small, circular patch reefs, conical pinnacle reefs, elongate barrier reefs that are separated from the coast by a lagoon or narrow sea, edge reefs that stretch along the coast, and more or less circular atolls, around extinct oceanic volcanoes. For a modern reef to form, a number of conditions need to be met, and this was in all likelihood also the case in the geological past (e.g., Tucker 1981): 5 high water temperatures: the optimum for reef growth is at 25 °C; 5 a low water level: optimum growth is achieved in the top 10 m of the water column;

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. Fig. 25.17 Biosparitic limestone from the Upper Muschelkalk, Middle Triassic, Rottershausen, Lower Franconia, Germany; Tempestite of a fossil seafloor that was stirred up by storm-induced turbulence on the sea bed; both living mussels and shells of died-off mussels, especially of plagiostoma striatum, were flushed out by this disturbance, tilted over and spread out to form a pavement of shell detritus; Mineralogical Museum of Würzburg University, Germany

5 a small tolerance range for salinity; 5 reef growth is favoured by intensive wave motion and a paucity in the supply of terrigeneous silt and clay particles. Consequently, reefs can best develop in shelf regions at the margins of warm epicontinental seas where they form the outer borders of carbonate platforms. Reefs are subdivided into three different zones (. Fig. 25.19): 1. The fore reef has a steep outer slope and a slope toe consisting of coarse reef debris. Towards the open ocean this debris passes into calciturbidite intercalated with deep-sea carbonate ooze; 2. reef core; 3. back reef consisting of reef debris, carbonate sands and oolites. The back reef passes into the carbonate platform with lagoons, where calcareous sands and oozes are deposited.

Thus, in addition to CaCO3, formed by inorganic or biochemical precipitation, marine limestones can contain, or consist predominantly of, clastic components, such as mechanically reworked shells of fossils and/or erosion products of older limestone. During sedimentation, inorganic precipitation of CaCO3 may take place that is not easily distinguished from diagenetic processes.

For thousands of years, limestones have been used as a popular dimension stone (. Fig. 25.20).

Chalk  is a very fine-grained biogenic limestone that formed from ooze accumulated at the bottom of the sea from submicroscopical particles of so-called coccoliths, sclerites of microorganisms known as coccolithophoridae (or coccolithophores) (. Fig. 2.12). It can contain concretions of microcrystalline quartz, known as chert or flint, that were formed during diagenesis from sponge spiculae and other siliceous skeletal parts of organisms and concentrated in bedding-parallel layers. Prominent occurrences of chalk are the Cretaceous limestone cliffs exposed at the coasts of southern England, Normandy and part of the Baltic countries. The spatially vast distribution of Late Cretaceous chalk deposits as a fairly uniform facies is explained by the perhaps most extensive flooding of continents at that time. The similar appearance of those chalk deposits results indirectly from the generally great thicknesses and flat bedding. However, chalk cliffs in the Baltic Sea area, such as on the Danish islands Møn and Fyn and the island of Rügen, Germany, show distinctly tilted and partly inverted bedding that resulted from ice-induced deformation during Pleistocene glaciation.

25.3.4  Formation of Terrestrial Carbonate

Rocks

Terrestrial carbonate rocks can occur as calcrete formed in arid and semi-arid climates, as calcareous sinter precipitated from springs or rivers, and as deposits in continental lakes. Calcrete  generally forms in arid and semi-arid regions,

such as in the deserts of southern Africa, Arabia, India,

25.3 · Chemical and Biochemical Sediments and Sedimentary Rocks …

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. Fig. 25.18 Stromatolite, alternating sequence of cm-thick dolomite beds with thinner layers of chert, forming harder ridges of darker-red colour; about 2.6 Ga Neoarchaean Malmani Subgroup of the Transvaal Supergroup; Lone Creek Falls near Pilgrim’s Rest, Mpumalanga, South Africa: a overview, b detail of the right lower part of the steep face; note that these carbonate rocks represent the oldest welldeveloped carbonate platform in Earth’s history (Photograph: Wolfgang Hermann, Würzburg)

and Australia, but is also known from the High Plains and the Sonora desert in North America. Calcrete, also known as caliche, hardpan or duricrust, is either formed when minerals are leached from the upper soil layer (A horizon) and accumulate in the layer below (B horizon) at depths of approximately 1–3 m below the surface. The mobilised calcium carbonate first forms small grains that with time accumulate and grow into more or less thick solid crusts and beds. An alternative way to form calcrete exists when subsurface water rises through capillary action. In arid regions, ground water often carries dissolved minerals that accumulate as calcrete near the surface. Plant roots may contribute

to the processes by removing water through transpiration, leaving behind dissolved calcium carbonate, which precipitates to form calcrete. Calcrete formation is generally a slow process with its rate depending on the availability of moisture. Rests of calcrete beds are a common component of fossil pedogenic horizons, though generally as eroded particles and/or subjected to more or less severe diagenetic overprint. This diagenetic alteration leads to cloudy structures or cauliflower-type microfacies. Stratified accumulations of calcrete particles can form more or less continuous beds in Mesozoic and Cenozoic playa deposits where they highlight the importance of terrestrial plants in the formation of the calcrete crusts. Locally they are also termed ‘stony marl’ (Stein-

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. Fig. 25.19  Principal marine and littoral areas of carbonate sedimentation and their characteristic facies zones

. Fig. 25.20  Medieval bridge across the Main River in the City of Würzburg, Germany, one of the earliest stone bridges in Central Europe, was constructed of a Middle Triassic biosparitic limestone (Muschelkalk Group); the first building, erected before 1200 and destroyed by a flood, is documented by big, rectangular blocks at the base, whereas considerably smaller dimension stones were used for the second building, erected between 1473 and 1543. The baroque sculptures of saints, decorating the bridge, consist of Middle Triassic (Keuper) sandstone (Photograph: Eckart Amelingmeier, Würzburg)

mergel) in the Central European Keuper (upper Triassic) beds (not to be confused with the homonymous lake deposits of similar age in adjacent formations; see below). Calcareous sinter  occurs, in contrast to calcrete, in regions with more substantial precipitation, in which Ca2+ and HCO3− ions dissolved in groundwater are moved into springs, streams and rivers. When groundwater reaches the surface as a spring or when river water is sprayed as a cascade, the CO2 dissolved in the water becomes released and CaCO3 is precipitated according to Reaction [25.2], a

process that is facilitated by simultaneous warming up of the water. The term calcareous sinter is generally applied to rocks thus formed without any biological component. The name travertine is often used as synonymous to calcareous sinter, but strictly speaking it refers to a porous limestone that is formed primarily from geothermal springs. In many cases it is associated with siliceous sinter. Aquatic plants, bryophytes, algae, and cyanobacteria usually grow on the surfaces of travertine. Their remains are generally preserved in various forms in the deposits and contribute to the distinctive porosity of the rocks (tufa). Travertine is used as

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25.3 · Chemical and Biochemical Sediments and Sedimentary Rocks …

dimension stone, especially for panelling of façades. In Roman times, the travertine of Tivoli near the City of Rome, Italy, was a popular decoration stone and can be seen in many historic buildings, such as Bernini’s colonnades of the St. Peter’s Square and Michelangelo’s ribs in St. Peter’s cathedral in Rome. Other famous travertine occurrences are the Pamukkale cascades in Turkey and Mammoth Hot Springs of Yellowstone National Park, USA. The sites of Pamukkale and the Plitvice Lakes National Park in Croatia are renown for natural dams formed by the calcareous precipitates. Some of the springs in which calcareous sinter is formed have such high temperatures that plants cannot live there. The resulting deposits are then generally less porous. In these cases thermophilic microbes can be important, however, and stromatolitic structure may develop. Lacustrine limestone and marlstone  commonly develop

under various non-humid warm climates. A fall in the lake water table can result in deposition of a thin shallow-water massive limestone or marlstone that may be followed by exposure and desiccation of the lake floor. They may be overlain by shale when the lake becomes flooded again at a later stage. The resulting rhythmic alternations are typical of playa deposits, and the typical whitish, uniform limestone or marlstone beds are a simple indicator of rhythm boundaries. Fossil lacustrine limestone beds are a frequent component of thick playa sequences, such as in the Central European Keuper deposits, where they are known as ‘stony marl beds’ (Steinmergel; not to be confused with the homonymous caliche remains; see above).

Lake chalk  is a fine-grained calcium carbonate or limerich mud with variable amounts of clay and silt that was precipitated in continental lakes. In this case, precipitation of CaCO3 is favoured by a lush aquatic vegetation, such as macrophytes, moss and algae, and cyanobacteria or moss. Such deposits are particularly well-known from continental Europe and known under the German term Seekreide.

25.3.5  Diagenesis of Limestone

Recent unconsolidated calcareous shallow marine sediments consist predominantly of metastable orthorhombic aragonite, CaCO3, and Mg-rich calcite. In contrast, limestone of older, especially pre-Tertiary geological formations consists exclusively of common, Mg-poor calcite. This leads to the conclusion that, during solidification of loose, impure calcareous mud to form lithified limestone, the two metastable phases aragonite and Mg-rich calcite become dissolved and common calcite precipitates from pore solutions, provided the sediment did not consist a priori of pure calcite. Organic components, such as humic acids, have considerable influence on diagenetic processes in limestone. Authigenic crystallisation of various silicate minerals, most frequently of albite or alkali feldspar, can take place not only in clastic sediments but also in limestone. As mentioned

above, diagenesis of carbonate sediments can be accompanied by formation of siliceous concretions. Dolomitic limestone is formed during the early- to late-diagenetic stage by reaction of Mg-bearing pore solution with limestone, initially deposited as calcareous mud. The process of dolomitisation can take place as long as sufficient pore space is available. Early diagenetic formation of dolomite is observed in coastal regions and shallow marine settings, for example at the coast of Florida, the nearby Bahama Bank, the Caribbean Sea or the Persian Gulf (e.g., Tucker 1981). Under the P-T conditions of sedimentation and diagenesis, a large miscibility gap exists between calcite and dolomite. Although, judging from the phase diagram . Fig. 8.11, dolomite should be nearly ideal CaMg(CO3)2 at these low temperatures, the proportion of Ca in natural early diagenetic dolomite typically exceeds the stoichiometric ratio (Füchtbauer 1988). As demonstrated by numerous Palaeozoic Ca-dolomite occurrences, the Ca-excess can be preserved over geological time spans. In reverse, new crystallisation of calcite can take place during diagenesis of dolomite, a process called de-dolomitisation. 25.3.6  Diagenetic Magnesite Deposits

A rarer but economically significant diagenetic process is the stepwise replacement of Ca2+ in limestone by Mg2+ to form large deposits of sparry magnesite, MgCO3, with dolomite as an intermediate product. While this metasomatic process has been ascribed by some authors to reaction of calcite with Mg-bearing orogenic aqueous fluids, isotopic and geochemical evidence speaks for a diagenetic origin of this type of magnesite mineralisation by the circulation of evaporite-derived brines (Azim Zadeh et al. 2015). Magnesite deposits of high economic value form irregular bodies within Devonian limestone and dolomite of the Greywacke Zone in the Eastern Alps of Austria. Magnesite is an ore mineral for the extraction of the light metal magnesium. However, more important is the production of sinter magnesite for the manufacturing of fire bricks used for lining of oxygen converters (LD process) and blast furnaces in steel plants. In addition, caustic magnesite is burned at 800 °C to remove CO2 and to obtain MgO for production of Sorel cement, MgCl2·5Mg(OH)2·6H2O, used as a binder in lightweight building materials. 25.4  Iron- and Manganese-rich Sediments

and Sedimentary Rocks

25.4.1  Stability Fields of Fe-Minerals

The most important minerals in Fe-rich sediments are goethite, α-FeOOH, haematite, Fe2O3, magnetite, FeFe2O4, siderite, FeCO3, and chamosite, a Fe-rich chlorite, and in special cases pyrite, FeS2.

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. Fig. 25.21  a Eh−pH diagram for the stability fields of dissolved Fe2+ und Fe3+ ions as well of the minerals haematite, magnetite, pyrite and siderite; the field of common Eh−pH conditions in near-surface environments is shaded in blue; total activity of dissolved carbonate = 1 molar, of dissolved S = 10−6 molar, of dissolved Fe = 10−6 molar (modified from Krauskopf 1979); b stability fields of FeS2, FeS, Fe-silicates and Fe2O3 in the diagram Eh versus the activity of HS− ions, dissolved in pore water, during diagenesis, at intermediate pH values; the FeS2 minerals pyrite or marcasite are precipitated from solutions of high or intermediate HS− activity, such as of composition A, whereas Fe-silicates, e.g. chamosite, can only crystallise from a solution of very low HS− activity, such as of composition B (from Taylor and Macquaker 2011)

Redox potential (Eh) and the hydrogen concentration (pH) are the most critical parameters that control the precipitation of Fe-minerals from natural aqueous solutions. In . Fig. 25.21a, an Eh-pH diagram is displayed for solutions with low contents of sulfide (expressed as HS−) and high carbonate concentrations. At high Eh values, i.e., at strongly oxidising conditions, haematite occupies a large stability field, whereas siderite is precipitated only at negative Eh values, i.e., under reducing conditions. The same holds true for pyrite, the stability field of which, however, is largely expanded in solutions with higher HS−/CO2 ratios. Magnetite exists at highly reducing, alkaline conditions, i.e., at negative Eh and high pH values. However, with decreasing HS− and CO2 in solution, its stability field is expanded to almost neutral conditions. Fe-silicates can precipitate only from alkaline solutions of high pH value, with high SiO2 and low CO2 contents. The stability fields of the most important Fe minerals formed during diagenesis can be displayed particularly well in a diagram Eh versus the activity of the HS− ions that are dissolved in the pore water (. Fig. 25.21b). The FeS2 minerals pyrite and marcasite can precipitate, at relatively low Eh values, in a broad stability field ranging from high to intermediate HS− activity, whereas crystallisation of Fe-silicates, such as chamosite, is possible only at distinctly lower HS− activity. At Eh values higher than ca. −0.28 to −0.25 V, haematite and Fe3+-hydroxides occupy a broad stability field (Taylor and Macquaker 2011). Commonly, the Fe content of groundwater is relatively low. At low O2 contents, Fe2+ is dissolved as ferro-com-

pounds, most commonly as carbonate, chloride or sulfate. In O2-rich surface water, these compounds undergo hydrolysis and are oxidised to form Fe(OH)3, part of which is present in a colloidal state. A relatively low amount is transported, in river water, as Fe3+-oxide hydrosol. However, these hydrosols can move over long distances as long as they are stabilised by colloidal organic substances. As these colloids are positively charged, they can be transported a long way without being precipitated, provided the concentrations of electrolytes and negatively charged colloids remain low in the river water. If this is not the case, Fe is eventually precipitated on the way. As soon as the river reaches the open ocean, the high electrolyte content of the seawater forces the Fe-bearing colloids to precipitate. Thus, depending on the redox potential present in the shelf or delta area, precipitation takes place as hydroxide (goethite), carbonate (siderite), silicate (e.g., chamosite) or even sulfide (pyrite). In many cases, the minute flakes of Fe-bearing colloids are attached to mineral fragments that have been whirled up by the coastal waters. Continued attachment of these flakes can lead to concentric overgrowths around these cores. As soon as these ooids attain a maximum size and weight that does no longer enable them to float in the water, they settle on the seafloor. Thus Fe-rich oolitic sediments are formed, in which numerous ooids can be broken during transport or sedimentation, and some of the oolites can display rubble structures. If the concentration of Fe is sufficiently high, marine-sedimentary oolitic iron ore can thus be formed.

25.4 · Iron- and Manganese-Rich Sediments and Sedimentary Rocks

Today’s rivers and seawater contain extremely low amounts of Fe, e.g., 1 mg m−3 Fe in the open ocean, far too low to form significant sedimentary iron ore deposits. In the geological past, the composition of river and seawater was, however, different, thus enabling at times the formation of voluminous sedimentary iron ores, which will be discussed in more detail below. In continental lakes, swamps or moors, plenty of stabilising colloids are available and, moreover, the electrolyte concentration is low. In these environments, precipitation of Fe-minerals is facilitated by bacteria and plants, which generate a reduction of Fe3+ in the Fe3+-oxide hydrosol and thus promote the precipitation of siderite. If there is little dilution by the input of clastic material, the amount of siderite thus formed can reach levels that warranted local mining of Fe-ore in historic times. 25.4.2  Sedimentary Iron Ores

Sedimentary iron ores and Fe-rich sedimentary rocks are divided into the following principal groups (James 1954; Evans 1993): 5 banded iron formation (BIF) 5 Phanerozoic ironstone 5 terrestrial iron ores z Banded Iron Formation

Today banded iron formation (BIF) is the principal iron ore for the global steel industry and represents the largest Fe reserves and resources on Earth. Apart from the internationally most common term banded iron formation various traditional or regional names, such as itabirite, from the Itabira locality in Brazil, jaspilite, haematite quartzite, specularite or taconite, popular in the USA have been in use. By far most of the known BIF (ca. 90%) formed between >3.8 and 1.9 Ga, i.e., during the Early Archaean and the Late Palaeoproterozoic, when at least 1014–1015 tons of Fe-oxide have been deposited. Distinctive peaks of BIF sedimentation at ca. 3.0, 2.7, 2.4 and 1.9 Ga have been connected with episodes of strong volcanic activity (Islay and Abbott 1999). A characteristic structural feature of BIF is a fine bedding due to a layering of iron ore, chert or jasper, and tiger’s eye, a silicified crocidolite, an asbestos of the Na-amphibole riebeckite, in which Fe2+ is oxidised to Fe3+ (. Fig. 25.22). Each of the 0.5–3 cm thick layers is internally laminated again in the range of millimetres to even tenth of millimetres. In many cases, however, the primary sedimentary structures are obliterated by a later, in part high-grade, metamorphic overprint. There is no doubt that BIF represents a chemical or biochemical sediment, i.e., it formed by precipitation from seawater, with or without the mediation by microorganisms, the exact process remains however speculative (Clout and Simonson 2005; Bekker et al. 2010). The source of iron in the seawater was on the one hand an elevated Fe-flux off the Archaean land surface at a time when due to the lack of

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O2 in the atmosphere, iron was easily dissolved as Fe2+ in meteoric waters (. Fig. 25.21a). On the other hand, submarine volcanic exhalations also played a pivotal role in the supply of Fe to the oceans (Poulton and Canfield 2011). A small proportion of BIF deposits were formed again in the Neoproterozoic in the waning stages of global glaciations (Poulton and Canfield 2011). According to the dominating ore minerals the following BIF facies are distinguished (. Fig. 25.21a, b): 5 oxide facies, the most important BIF facies, in which haematite and/or magnetite dominate; 5 carbonate facies with siderite as principal Fe-mineral; 5 silicate facies with Fe-rich sheet silicates, such as greenalite (7 Sect. 11.5.6), chamosite (7 Sect. 11.5.5), glauconite (7 Sect. 25.2.2), minnesotaite, an Fe-rich talc, and stilpnomelane (7 Sect. 26.1.3), the latter is however a post-depositional low-grade metamorphic phase; in addition, the amphiboles grunerite and manganogrunerite can be present; 5 sulfide facies with pyrite as principal Fe-mineral. From a genetic point of view, two major types of BIF deposits have been distinguished: Superior type  The BIF deposits of this type are interlayered with quartzite, black marly mudstone and additional sedimentary rocks, were deposited in shallow seawater on the continental shelf, cut-off marine basins, advancing coast lines or intracratonic basins. A direct connection with volcanism is not evident. Commonly, the ores belong to the oxide, carbonate or silicate facies. Important deposits of this type are situated in the type area of the Superior Province of North America, where they form so-called ranges, especially the Cuyuna, Mesabi and Vermillion Range in Minnesota, the Penokee-Gogebic Range in Wisconsin and Michigan, the Marquette and Menominie Range in Michigan, as well as the Steep-Rock and the Michipicoten District in Ontario, Canada. Among these, the Mesabi Range is the most productive. Additional deposits of high economic relevance belonging to this type are situated in the Labrador Trough (New Quebec Orogen) of Canada, near Krivoi Rog and Kursk in Russia, in the Bihar-Orissa region of India, in the Hamersley Range in Western Australia (. Fig. 25.22), in Minas Gerais in Brazil, and in the southern Kalahari Desert in the Northwest Province of South Africa. Algoma type  Many BIF of this type, which can occur in

oxide, carbonate or sulfide facies, form part of Archaean greenstone belts, but some are of younger Precambrian age. A characteristic feature is the interlayering with greywacke and volcanic rocks, which can be regarded as evidence of a volcanogenic-exhalative source of the iron. Important deposits of this type are situated in the Neoarchaean Abitibi greenstone belt at the Kirkland Lake and near Temagami in Ontario, in the Vermillion Range of Minnesota, in the Yilgarn Block and the Pilbara District of Western Australia, in Liberia, western Africa, as well as in Zimbabwe.

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. Fig. 25.22  Banded iron formation (BIF) of the Superior type; folded intercalation of haematite (dark blue), tiger’s eye (yellow), jasper (red) and crocidolite (light blue); Hamersley Range, Australia, width of view 15 cm

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z Phanerozoic Ironstones

z Terrestrial Iron Ores

In contrast to BIF, which are restricted to Precambrian units, ironstone is a shallow marine Phanerozoic feature that had some economic significance in historic times. Two different types are distinguished: 5 oolitic haematite-chamosite-siderite ore of Clinton type, named after the Clinton deposit in Alabama, USA, with average contents of 40–50% Fe. Most of these deposits are of Lower Palaeozoic age. An important example is the Wabana deposit in Newfoundland, Canada; 5 iron ore of minette type commonly consist of siderite, chamosite, haematite and limonite. Most of these ores display an oolitic structure. For instance, the Middle Jurassic (Doggerian) minette ores in central England, Lorraine, Luxembourg and southern Germany consist of limonite ooids that are surrounded by a matrix of calcite, siderite and chamosite. On average, these ores contain 30% Fe, >20% SiO2 and 5–20% CaO. Due to their high carbonate contents, many of them can be smelted without addition of fluxes. For centuries, many of the minette ore deposits were of considerable economic relevance but today they cannot compete with BIF deposits.

In the past, some of the iron ores formed on land had been sources for local Fe production but are no longer of economic interest. Two types are distinguished: 5 sideritic blackbands and claybands consisting of siderite, carbonaceous material and argillitic detritus. They are typically associated with coal seams; 5 bog iron ore or morass ore (German: "Raseneisenerz") consisting of limonite precipitated subrecently in continental lakes and swamps, predominantly at relatively high (northern) latitudes. Such deposits form earthy masses, encrusted by sand grains, within or above peat seams.

The sedimentary Fe-deposits of the Salzgitter district, situated in the northern foothills of the Harz Mountains, include oolitic Fe-ore of Jurassic (Liassic and Malm) and Cretaceous age as well as Cretaceous “rubble ore” (argillaceous ironstone). They were derived from the mechanical reworking of lateritic soils into coastal environments. This debris was washed together, deposited and oxidised to limonite in depressions that were caused by salt tectonics. The matrix between the debris and the ooids consists to variable extent of clay, marl, calcite, ankerite or siderite. With reserves of up to 4000 Mt of ore, the Salzgitter district used to house the most important Fe-ore deposits of Germany. Between 1937 and 1982, these have produced 340 Mt ore in total that has been smelted in the Salzgitter steelworks. The last underground mining of this district, shaft “Konrad”, is kept open as permanent deposal site for low-level radioactive waste.

25.4.3  Sedimentary Manganese Ores

Iron and manganese are geochemically very similar. Consequently, their spatial separation by sedimentary processes in seawater or fresh water is a major, much debated problem. Presumably, differences in the redox potential are the decisive factor. At the same temperature and pH value, iron is precipitated as Fe(OH)3 at a lower Eh than manganese as MnO2 (Maynard 1983). An important role could play the mixture between the seawater and O2-rich surface water in coastal areas. In many cases, sedimentary manganese deposits have been upgraded by tropical weathering to form mineable deposits, a prominent example being the Nsuta deposit in Ghana (7 Sect. 24.5.4). z Precambrian Manganese Ores

In many regions, sedimentary manganese ores are associated with BIF and, consequently, can have been deformed and metamorphosed. In the so-called gondite, metamorphic minerals, such as spessartine, Mn3Al2[SiO4]3, ­ Mn-pyroxenes, Mn-amphiboles and braunite,

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25.4 · Iron- and Manganese-Rich Sediments and Sedimentary Rocks

Mn2+Mn63+[SiO4]O8, predominate, whereas carbonatic manganese ores mainly consist of rhodochrosite, MnCO3, kutnohorite, CaMn(CO3)2 or manganoan calcite. Initially, the Mn-rich sediments were deposited in shallow-sea environments of intracratonic basins. Some authors assume that the metal content was derived from submarine volcanic exhalations. Important deposits of this type are situated in Odisha (Orissa), India, in Minas Gerais, Brazil, in Bolivia as well in Gabon and Ghana, central and western Africa. Of high economic significance is the Kalahari ore field in the Palaeoproterozoic Transvaal Supergroup in the Northern Cape Province of South Africa, containing braunite, kutnohorite and hausmannite as predominant ore minerals. With reserves of 4400 Mt of ore, this ore field ranks first worldwide. In the largest mine, Mawatwan, a 20 m thick ore seam contains 38% Mn on average. An important European deposit is situated at Jacobeni, in the crystalline complex of the Eastern Carpathians, Romania. In 2019 South Africa produced 5.5 Mt Mn or nearly 30% of the world production, followed by Australia (3.2 Mt), Gabon (2.4 Mt), China (1.3 Mt) and Brazil (1.2 Mt). z Phanerozoic Manganese Ores

The marine-sedimentary Mn-ore deposits of Nikopol at the lower course of the Dnieper, Ukraine, and of Chiatura at the southern slope of the Caucasus Mountains, Georgia, have been deposited in Oligocene times in the shallow-sea of the southern Ukraine Basin. The ore is made up of oolites, concretions and earthy masses consisting of pyrolusite, β-MnO2, and romanèchite, (Ba,H2O)Mn5O10, whereas rhodochrosite, MnCO3, and manganoan calcite are the predominant ore minerals in the carbonate facies. In the past, these Mn deposits were of paramount economic signficance but lost this status due to their relatively low Mn contents of only 15–25% at Nikopol. Today, the Cretaceous Mn-ore deposits of Groote Eylandt, northern Australia, with average contents of 51% Mn and reserves of 300 Mt, are much more important. The largest Mn deposit in northern America with average contents of 25% Mn is situated in the Molango District of Mexico. 25.4.4  Metal Concentrations on the Ocean

Floor

In the future, manganese nodules in deep sea basins of the Pacific and Indian Ocean could form a significant metal resource (. Fig. 25.23). At present, however, both the costs of mining and the ecological risks are far too high. The nodules precipitated as gels where dissolved Mn2+ and Fe2+ ions came into contact with cold, oxygen-rich, seawater upwelling from greater depth. Initially, they consisted predominantly of X-ray amorphous Mn and Fe compounds which, during later diagenesis, crystallised to form complex, H2O-bearing oxide and hydroxide minerals, such as todorokite, birnessite and vernadite (7 Sect. 7.4). These

contain conspicuously high, though variable, concentrations of base metals, particularly of Ni and Cu, but also Co, adsorbed on their surface. For instance, manganese nodules from the nodule belt of the northern Pacific have average contents of 27% Mn, 1.3% Ni, 1.2% Cu and 0.2% Co. Furthermore, manganese nodules can contain elevated concentrations of the high-tech metals Zr and REE. The concentrically layered structure of the nodules indicates that these have grown, on the ocean floor, over geological time spans, at a rate of about 1 mm/106 years. Presumably, the metal contents are derived from debris of continental weathering and/or from submarine volcanic exhalations. 25.5  Siliceous Sediments and Sedimentary

Rocks

These sediments consist of non-detrital SiO2 minerals, such as opal, chalcedony, jasper or macrocrystalline quartz, and they can be either abiogenic or biogenic. During diagenesis, metastable biogenic opal is transformed, via disordered cristobalite/tridymite, into stable α-quartz. This explains why many cherts consist entirely of (micro)crystalline quartz. Whether SiO2 minerals are dissolved or precipitated depends mainly on temperature and pH. Amorphous SiO2, such as opal, dissolves much more readily than the crystalline SiO2 phases, such as quartz, tridymite and cristobalite. In a temperature range of 0–200 °C, the solubility of SiO2 increases continuously in a linear way. As shown in . Fig. 24.2, silica is moderately soluble, as Si(OH)4, up to a pH of about 9. At higher pH, however, the solubility increases drastically, by up to 30–50 times, due to ionisation of Si(OH)4. z Precipitation and Diagenesis of Siliceous Sediments

River water contains some SiO2 in solution but in extremely low amounts. Similarly, the essentially biogenic SiO2 content dissolved in seawater is very low, thus prohibiting precipitation. Unicellular organisms, such as radiolarian or diatoms or the multi-cellular siliceous sponges take up SiO2 and use it for building their skeletal framework, composed of opal. The dead remnants of these organisms settle on the sea floor where they form siliceous sediments, in the form of diatom or radiolarian ooze. In contrast, porous diatomite, also known as diatomaceous earth or kieselguhr, can also form in fresh water. In either case, skeletal framework of organisms constituting these loose sediments constitute X-ray amorphous opal-A besides some tridymite and cristobalite. During incipient diagenesis, these biogenic siliceous sediments are solidified to form tripoli. Dissolution and recrystallisation processes during advanced diagenesis lead to formation of siliceous shale or porcellainite, consisting of crystalline but highly disordered opal-CT, and finally of chert, composed of chalcedony, a very fine-grained (72 °C while, at  σ3 acting on a point in a rock are usually illustrated by the three orthogonal axes of a stress ellipsoid, in which σ3 coincides with the direction of the pressure performed by the overlying rock column Pl. A mean stress value is given by σmean = (σ1 + σ2 + σ3)/3, whereas the differential stress is usually defined as σdiff = σ1 – σ3. Stress exerted on a plane in the rock is a vector that can be resolved into the components normal stress σn and shear stress τ, respectively. Differential stress is the reason for strain in a rock, whereby the shape and volume of a rock is changed. If this process is accompanied by changes in the relative position of minerals in the rock or of domains in the crystal structure, the more general term deformation is used (Passchier and Trouw 2005). Deformation processes 5 essentially emboss the structure and texture of a metamorphic rock; 5 open pathways for transport of fluids; 5 favour mineral reactions due to augmentation of grain contacts, thereby enhancing the reaction rate and lowering the activation energy.

Note, however, that the stability fields of individual minerals or mineral assemblages are not influenced by stress:

459 26.2 · Metamorphism as a Geological Process

Experimental investigations have shown that, at common metamorphic conditions such as medium to high temperatures, presence of H2O and low strain rates, rocks are not firm enough to endure stress differences of more than a few tens, at best a few hundred bar. Above these pressures, the yield strength of a rock would be overstepped. For this reason, a possible tectonic overpressure can attain only very small magnitudes and, by no means, could explain, e.g., the formation of high-pressure minerals in metamorphic rocks. At low temperature and lithostatic pressure and/or high strain rates, brittle deformation of mineral grains is predominant, resulting in cataclastic metamorphism (7 Sect. 26.2.2). On the other hand, mineral grains undergo ductile deformation at higher temperature and lithostatic pressure and/ or lower strain rates, which is generally the case during regional metamorphism in orogenic belts (7 Sect. 26.2.5). Ductile deformation leads to lattice defects in crystal structures such as line defects, deformation twinning, e.g., in feldspars and carbonate minerals, undulose extinction and deformation lamellae, especially in quartz, deflection in mica and kyanite as well as development of subgrain boundaries (7 Sect. 26.4.3). The boundary between ductile and brittle deformation differs widely between mineral species. For instance, at a given (low) temperature, serpentinite, mudstone, limestone, gypsum or other evaporites can recrystallise and thus undergo ductile deformation, while quartz-feldspar rich rocks, such as granites and gneisses, behave in a brittle manner, resulting in cataclastic deformation. At elastic deformation, the changes in grain shape can be reversed and recovery takes place. Static metamorphism occurs at hydrostatic pressure without any indications of deformation in the rock fabric, whereas deformation processes play an important role in dynamic or dynamo-thermal metamorphism. z Fluids and Fluid Pressure

A fluid phase is commonly present in most rocks during metamorphism, be it as intergranular thin fluid film, in pores or in joints and fractures. The dominant volatile components are H2O and/or CO2, but additional volatiles, such as CO, CH4, HCl, HF, H3BO3, O2, H2 and others can be present as well. In the literature, the fluid phase in metamorphic systems is variably designated as steam, gas, liquid, fluid or super-critical fluid. As the critical points of volatile components are located at low P-T conditions, such as PC = 218 bar, TC = 371 °C for pure H2O, and PC = 73 bar, TC = 31 °C for CO2, many metamorphic processes proceed at super-critical conditions, at which no difference can be made between steam (gas) and liquid. Consequently, the terms fluid phase or fluid are appropriate. Generally, the density of the fluid phase decreases with rising temperature and increases with rising pressure. As T and P increase with depth, thermal expansion and pressure-induced compression can nearly cancel each other out. At a geothermal gradient of 15 °C km−1, the density of H2O at a depth of 35 km deviates only slightly from the surface value of 1.0 g cm−3 whereas, at a geothermal gradient of 50 °C km−1, the density of H2O is 0.67 g cm−3 at a depth of 15 km (cf. Best 2003, p. 75ff, 488ff). Among the natural fluids, H2O displays an extraordinary dissolving power for alkalis, but also for SiO2. Thus, compared to pure water, H2O-rich fluids display higher critical values because they always contain some dissolved ions.

The total pressure Pfl that is applied to the fluid in a rock results from the sum of the partial pressures of individual fluid species involved: Pfl = PH2O + PCO2 +… However, at the elevated lithostatic pressures realised at most metamorphic conditions, the fluid species do not behave like ideal gases. Consequently, the fugacities fH2 O , fCO2 , . . . should be used instead of the partial pressures. By analogy to the activity ai (7 Sect. 20.2.3), the term fi = γiPi is valid for fluids and gases, in which the fugacity coefficient γi is dependent on both P and T.

In many cases, especially during low- to medium-grade metamorphism at crustal levels that are too deep to maintain an interconnected pore space, fluids and the solid rock are under approximately the same lithostatic pressure Pl or total pressure Ptot. In this case, the fluid pressure is equal to the lithostatic pressure Pfl ≈ Pl = Ptot or, provided H2O or CO2 are clearly predominant, Ptot = Pfl ≈ PH2 O or Ptot = Pfl ≈ PCO2. If, on the other hand, the amount of fluid is too low to build up a fluid pressure similar to the lithostatic pressure, the condition Pfl  400 °C, reached at a distance of 330 m from the contact, whereas the same temperature, caused by a granodiorite pluton of D = 10 km, is exceeded at a distance of 3.3 km. As the temporal temperature development within a contact

Cataclastic metamorphism takes place within zones of tectonic deformation, such as in normal, reverse and transform faults or thrust planes. It acts upon a rock and its mineral constituents essentially by mechanical strain, whereby differential stress plays the dominant role. At relatively low temperature and lithostatic pressure and/or high strain rates, the strength of a rock is overstepped leading to brittle deformation such as fracturing and grinding of minerals, a process known as cataclasis (Greek κατάκλασις = fracturing). Intergranular movement involving mutual displacement of the individual particles formed is called cataclastic flow. With increasing temperature and/or decreasing strain rate, tectonic fault breccia, kakirite, cataclasite and mylonite are formed. Fault breccia and kakirite  are typically penetrated by a nar-

rowly spaced system of joints, striations and slickensides and thus tend to polyhedral disintegration on the centimetre to decimetre scale. Microscopic inspection reveals that cataclasis is relatively weak within the individual fragments. The amount of matrix is 50 vol.%. Rocks with 10–50 vol.% matrix are termed protocataclasite, extremely deformed rocks with >90 vol.% matrix ultracataclasite. The so-called mortar structure is characterised by larger fragments, known as porphyroclasts, that are surrounded by collars of finer-grained aggregates of the same mineral. This structure is a result of cataclasis, in some cases combined with some recrystallisation. New minerals can be formed only in joints.

Mylonite  (Greek μύλη = mill) forms in ductile shear zones, in which higher temperature, higher lithostatic pressure and lower strain rates lead to crystalloplastic flow and, consequently, mylonitisation. The transition from brittle to ductile deformation critically depends on the crystal structure of the minerals involved. For instance, quartz behaves in a ductile manner already at relatively low temperatures and thus commonly displays oscillatory

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Chapter 26 · Metamorphic Rocks

extinction already in cataclastic rocks. In contrast, garnet and pyroxene can undergo brittle deformation at temperatures as high as 500–600 °C. Mylonite is distinguished from cataclastic rocks by a typical flaser structure in which a schistose matrix displays a mottled striation and encases eye-shaped porphyroclasts, thereby imaging movement trajectories. Micas are aligned to form extended layers. In many cases, a stretching lineation is visible (7 Sect. 26.4.3). Although, in contrast to cataclasite, mylonite has undergone ductile shear, larger porphyroclasts can be disrupted by brittle deformation. In mylonitised granites, fractured K-feldspar is commonly altered to sericite. Ultramylonite is characterised by a matrix content of >90 vol.% and resembles slate, as its smallest fragments display grain sizes of 600 kbar and correspondingly high residual temperatures of >1500 °C, realised in zone IV, the rock was completely melted. In some cases, the melt volumes formed were big enough to cool down so slowly that they solidified to form holo- or hypo-crystalline rocks closely resembling volcanic rocks of equivalent composition, whereas inhomogeneous glasses formed from melts that had cooled down more rapidly. In zone V, the centre of the impact, the rocks became totally vaporised at peak pressures of >800–1000 kbar (>80 GPa) and residual temperatures of >3000 °C. Already during the compression phase of the impact, lasting only for about half a second, the shock metamorphosed material began to move downward towards the crater bottom before it started to be deposited inside the crater and ejected beyond the crater rim as “ejecta blanket”. In the centre of the growing crater, a so-called vapour plume or ejecta plume consisting of vapour and melt and a small fraction of fine-grained solid material was formed moving vertically upward before it collapsed and formed a fallback deposit mainly inside the crater. The excavation process progressed further during the subsequent excavation phase, in which the compacted matter was released to ambient pressure. Still during the impact, a decompression wave propagated through the impactor and the bedrock (. Fig. 26.5c) at the velocity of sound, i.e., much slower than the shock wave that had attained supersonic speed. Consequently, the excavation phase lasted 10,000 times

26.2 · Metamorphism as a Geological Process

longer than the compression phase, i.e., minutes to tens of minutes. Due to the pressure release, the compressed rock matter was ejected in all directions and deposited as a continuous ejecta blanket around the crater up to a radial distance of 2–3 times the crater radius. During ejection molten rock material can be shaped aerodynamically to form characteristic glass bombs, e.g., the so-called “Flädle” of the Ries Crater. These constitute a major part of the suevite, an impact breccia found mainly inside of impact craters and in small amounts outside on top of the ejecta blanket. The components of suevite are derived predominantly from the deeper levels of the impact crater and mainly consist of melt bodies and fragments of shocked lithic and mineral clasts, placed in a matrix of fine-grained mineral clasts and secondary alteration products such as montmorillonite. A rare component is the silica glass lechatelierite formed by melting of quartz. Much more common than suevite are impact breccias formed by rock fragments that have undergone none or only slight impact metamorphism, such as the Bunte Breccia (“multi-colored breccia”) of the Ries, which constitutes the main mass of the ejecta blanket. The name suevite is derived from the historical region Swabia in southern Germany. It was first applied to the characteristic rocks found in the Ries Crater but is now established as an internationally accepted term. From field, structural and textural evidence as well as numeric modelling, Stöffler et al. (2013) have estimated that, prior to erosion, 108–116 km3 of allochthonous impact rocks had been deposited by the Ries event. The total amount of suevite produced is estimated between 14 and 22 km3, the total amount of impact melt as 4.9– 8.0 km3. According to numerical modelling, only a very small proportion of suevite was formed by the primary ejecta plume, whereas the main mass was deposited from a secondary plume that formed by an explosive reaction of the impact melt, filling a temporary melt pool, with water and volatile-rich sedimentary rocks.

All the types of rock described above can be categorised as “proximal impactites” located inside the crater or in the ejecta blanket around the crater (Stöffler and Grieve 2007). In addition, rare “distal impactites” are known and called tektites. These are rounded glass bodies of centimetre size and black to greenish or yellowish colour that were formed by jetting of molten surficial bedrock material during the earliest phase of an impact. They are ejected over long distances up to several hundred kilometres from the crater centre and form extended scatter fields (7 Sect. 31.4) such as the moldavites related to the Ries crater. Impact metamorphism is the predominant metamorphic process that has affected and still affects lunar rocks of the upper crust of the Moon. As the Moon is not surrounded by an atmosphere and thus no chemical weathering takes place, impact craters of very old age are excellently preserved in great number. The same holds true for the rocks affected by subsequent phases of impact metamorphism and forming the essential constituents of the regolith, an extended layer of rock debris covering the Moon’s surface (cf. 7 Chap. 30). By underground nuclear explosions, artificial shock waves have been produced. These cause changes in the adjacent country rock that closely resemble features produced by natural impact metamorphism.

467

26.2.4  Hydrothermal Metamorphism

Hot solutions or steam that migrate along faults and/or a joints can cause alteration of the country rock, leading to the replacement of primary minerals by newly formed hydrothermal minerals (Coombs 1961). These processes are quite common but are generally restricted to narrow zones adjacent to the fissures. Much more extensive alteration of the country rock is achieved, however, in active geothermal fields, i.e., in regions where hot springs or water vapour are emitted on a larger scale—a phenomenon that can be used for generation of energy (Utada 2001). Boreholes sunk into depths of several hundred metres revealed, at temperatures increasing up to 250 °C, the new formation of zeolites (7 Sect. 11.6.6), such as mordenite, (Na,Ca,K)6[AlSi5O12]8·28H2O, analcime, Na[AlSi2O6]·H2O, laumontite, Ca[Al2Si4O12]·4.5H2O, and wairakite, Ca[AlSi2O6]2·2H2O, as well as albite and adularia. Well-investigated are the regions of Wairaki on the northern island of New Zealand, of Onikobe and Hakone on the Island of Honshu, Japan, and the Yellowstone National Park in Wyoming, USA. All these regions are characterised by an exceptionally high geothermal gradient of as much as 1000 °C km−1. Extensive alteration is also produced in connection with submarine hydrothermal activity, caused by black smokers at mid-ocean ridges (7 Sect. 23.5.1) and during formation of hydrothermal ore deposits, such as porphyry copper ores (7 Sect. 23.2.4). 26.2.5  Regional Metamorphism in Orogens

In cratons of Precambrian age and in Phanerozoic orogenic belts, metamorphic rocks occupy regions extending over hundreds to thousands of square kilometres. This means that, in contrast to the occurrences so far described, metamorphic processes take place on a regional scale and are connected with important orogenic events, thus being related to plate-tectonic processes such as subduction or continent-continent collision. In its character, the regional metamorphism is neither purely dynamic nor purely static-thermal. Rather, a complicated interplay prevails between deformation, producing folding and schistosity (. Figs. 26.9, 26.20, 26.21, 26.22, 26.23), and regional heating that leads to metamorphic recrystallisation and neomineralisation. These processes can repeatedly occur within one or more orogenic phases. Typical products of regional metamorphism are foliated rocks, such as phyllite, micaschist, gneisses, amphibolite or granulites. In their structural characteristics, these differ from the non-schistose hornfelses in thermal aureoles, but also from hardly recrystallised cataclasites or mylonites, whereas blastomylonites are schistose. Similar to contact aureoles, mineral zones of increasing metamorphic grade can be mapped in areas affected by regional metamorphism. These are defined by critical minerals or mineral assemblages that are formed by prograde

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Chapter 26 · Metamorphic Rocks

. Fig. 26.9  Large open fold in metamorphosed sedimentary rocks of the Gemsbok-River Formation in the Pan-African Damara Orogen. The outcrop displays an interlayering of yellowish to brownish carbonate rocks with dark turbidites that have been folded and metamorphosed, about 550–500 Ma ago. Rhino Wash, Ugab region, northern Namibia (Photograph: Martin Okrusch)

26

metamorphism and commonly reflect the P-T conditions at the metamorphic peak. The individual mineral zones are delineated by isograds. These are defined as lines, connecting points, at which a critical mineral, also known as index mineral, first appears in the field. In fact, isograds are curved planes dissecting the orogenic body, of which the lines of intersection with the Earth’s surface are mapped. In many cases, partial melting takes place at highest metamorphic grade thus producing migmatites of regional extent (7 Sect. 26.5). Typically, high-grade metamorphic rocks and/or migmatites are connected with plutons of granitic, granodioritic or tonalitic composition that were emplaced subsequently to the regional metamorphism. In many cases, especially at lower crustal levels, it is difficult to discriminate between regional and contact metamorphic overprint, thus giving rise to the term regional contact metamorphism. z Low Pressure and Medium Pressure Metamorphism

For the first time, zones of increasing metamorphic grade were recognised and mapped in the Dalradian Supergroup of the Caledonian Orogen in Scotland by Barrow (1893, 1912) and Tilley (1925). In metapelites, these so-called Barrow zones are characterised by the following mineral assemblages: 1. chlorite zone: phengitic white mica + chlorite ± microcline + albite + quartz 2. biotite zone: biotite + chlorite + muscovite + albite + quartz 3. garnet zone: almandine-rich garnet + biotite + muscovite + albite/ oligoclase + quartz 4. staurolite zone:



staurolite + almandine + biotite + muscovite + oligoclase + quartz 5. Kyanite zone: kyanite  ± staurolite + almandine + biotite + muscovite + oligoclase + quartz 6. sillimanite zone: sillimanite + almandine + biotite + K-feldspar + oligoclase + quartz Critical assemblages in metabasites and calc-silicate rocks display a similar zonation. Since then, well-developed mineral zones have been described in many metamorphic complexes worldwide. Prominent examples are 5 the Palaeozoic metamorphic complex of Vermont and New Hampshire, USA, 5 the Damara and Kaoko belts in Namibia, formed during the Pan-African Orogeny, 5 the Variscan metamorphic complex in the northern part of the Bavarian Forest, Germany, 5 the sedimentary units of Upper Triassic (Keuper) and Lower Jurassic (Lias) age, situated in the northern foothills and the Central Alps of Switzerland, and metamorphosed during the Alpine Orogeny. The mineral zonations recorded do not always conform to the classical Barrovian zones. In many cases, different index minerals or mineral assemblages have been observed that point to regional differences in P-T conditions at the metamorphic peak and to different geothermal gradients during an orogenic event. A convincing example is presented by the Scottish Caledonides themselves. As already observed by Harker (1932), andalusite and cordierite are formed, as

26.2 · Metamorphism as a Geological Process

additional index minerals, in metamorphic rocks of the Dalradian north of Aberdeen. This leads to a different sequence of mineral zones, which Read (1952) defined as Buchan type and contrasted with the classical Barrow type. According to our present knowledge, the mineral zonation of Buchan type is due to a higher geothermal gradient, i.e., a stronger temperature increase with depth, than attained in the area of Barrovian metamorphism. Such regional differences are well-understood when considering the distribution of isotherms within the continental plate above a subduction zone (. Fig. 28.2). A further instructive example of regional metamorphism at medium pressure conditions, i.e., at medium P/T ratio, is the metamorphic complex on the Island of Naxos, Cyclades archipelago, Greece. The Cyclades Metamorphic Complex essentially consists of a Permo-Mesozoic sedimentary sequence with intercalated volcanic rocks, which rests upon a pre-Alpine basement. The entire assemblage . Fig. 26.10 Simplified geological map of the Island of Naxos, Cyclades Archipelago, Greece, displaying metamorphic isograds in mineral zones defined by assemblages in metapelites and metabauxites (modified after Jansen and Schuiling 1976)

469

was subjected to polymetamorphic overprint, followed by a phase of magmatic activity. Apart from the pre-Alpine relics, the metamorphic complex of Naxos essentially consists of clastic sedimentary rocks and weathered limestone that contains bauxite in karst cavities. In Eocene times, the Mesozoic sedimentary sequence experienced high-pressure metamorphism (see below), relics of which are still preserved in the southeastern part of the island. At the Oligocene/Miocene turn, the high-pressure rocks were subjected to a prograde medium-pressure metamorphic overprint that produced a distinct zonal succession of metamorphic index minerals in metapelites and metabauxites (. Fig. 26.10; Jansen and Schuiling 1976; Feenstra 1985): I. In the diaspore-chloritoid zone, diaspore and chloritoid formed in metabauxite, in part accompanied by pyrophyllite or kyanite, the low-T/high-P polymorph of Al2[O/SiO4] (. Fig. 26.11). Metapelites contain the assemblage

26

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Chapter 26 · Metamorphic Rocks

quartz + albite + muscovite ± paragonite+ chlorite ± chloritoid ± garnet.

kyanite ⇋ sillimanite

II. The corundum-chloritoid zone starts with the corundum isograd defined by the first appearance of corundum in metabauxite according to the dehydration reaction:

2 AlOOH ⇋ Al2 O3 + H2 O diaspore corundum

26

[26.1]

Otherwise, none of the mineral assemblages are changed in metabauxites and metapelites. III. The biotite-chloritoid zone is characterised by the first appearance of biotite in metapelites, whereas paragonite disappears. The following assemblage is typical:

quartz + albite + muscovite + biotite + chlorite ± chloritoid ± garnet.

chloritoid + kyanite ⇋ staurolite + quartz + H2 O

[26.2]

Whereas in metabauxite, kyanite is present already in zone I, it crystallised as a new mineral in metapelites of Zone IV, constituting the following assemblage:

quartz + oligoclase + muscovite + biotite + garnet ± staurolite + kyanite/sillimanite ± andalusite Vb.  Sillimanite zone. In metapelitic rocks of this zone, sillimanite is the only Al2[O/SiO4] polymorph present either in fibrolitic or prismatic habit. The important dehydration reaction

whereas the assemblage

corundum + staurolite + margarite + muscovite ± biotite + chlorite is observed in metabauxite. Va.  Kyanite-sillimanite transitional zone. In metabauxites of Zone V, substantial amounts of green spinel grew together with corundum. Moreover, the breakdown of margarite started to form corundum + anorthite due to the dehydration reaction: [26.3]

The high-T polymorph of Al2[O/SiO4], sillimanite, formed as fibrolite variety in the transitional zone by the reaction

[26.5a]

has constituted the critical assemblage sillimanite + K-feldspar, which is of great significance for the highest grade of low and medium-pressure metamorphism. Thus a common assemblage in metapelites of this zone is:

quartz + oligoclase/andesine + K-feldspar (± relic muscovite) + biotite + garnet + sillimanite. Moreover, dehydration melting according to the reaction

muscovite + quartz + H2 O ⇋ sillimanite + liquid

quartz + oligoclase + muscovite + biotite ± garnet ± staurolite ± kyanite

CaAl2 [Al2 Si2 O10 ](OH)2 ⇋ Al2 O3 + Ca[Al2 Si2 O8 ] + H2 O

although part of the kyanite has been preserved metastably. Locally, the reacting Al-silicates can be seen together under the microscope, occasionally accompanied by the low-P/low-T modification andalusite. Thus the typical assemblage in metapelites of this zone is:

KAl2 [AlSi3 O10 ](OH)2 + SiO2 ⇋ Al2 [O/SiO4 ] + K[AlSi3 O8 ] + H2 O

In metabauxite, the brittle mica margarite was commonly formed together with corundum and chloritoid. IV. In the kyanite-staurolite zone, chloritoid has disappeared and staurolite formed in metabauxites and metapelites, by the simple reaction:

[26.4]

[26.5b]

took place from Zone Vb on, leading to partial melting. In the absence of quartz, muscovite is still stable at higher temperatures (. Fig. 27.8, reaction 12) and thus can be found in metabauxites of Zone Vb. The same holds true for staurolite. VI.  The migmatite core. Incipient melting in the metamorphic complex of Naxos is documented by typical migmatite structures in metasedimentary gneisses. Light-coloured, leucocratic domains resembling pegmatite or aplite are developed besides dark, biotite-rich patches. These restites trace the previous schistosity that is otherwise largely destroyed. As migmatites initially contained a considerable amount of liquid, they were plastically deformed thus producing typical flow folds. Mineral assemblages in partially molten metapelites of Zone VI are:

quartz + oligoclase/andesine + K-feldspar + biotite + garnet + sillimanite

26.2 · Metamorphism as a Geological Process

471

. Fig. 26.11  Pressure-temperature (P-T) diagram for a quantitative estimate of regional variation of metamorphic grade on the Island of Naxos, indicated by the medium-blue arrow and the mineral zones I to Vb. The following equilibrium curves, experimentally determined for important reactions, are displayed: 1: diaspore ⇌ corundum + H2O (Haas 1972); 2: chloritoid + kyanite ⇌ staurolite + quartz + H2O (Richardson 1968); 3: margarite ⇌ corundum + anorthite + H2O (Chatterjee 1974); 4a: kyanite ⇌ andalusite; 4b: kyanite ⇌ sillimanite; 4c: andalusite ⇌ sillimanite (Holdaway and Mukhopadhyay 1993); 5a: muscovite + quartz ⇌ andalusite/sillimanite + K-feldspar + H2O (Chatterjee and Johannes 1974); 5b: muscovite + quartz + H2O ⇌ sillimanite/kyanite + liquid (Storre and Karotke 1972). Light-blue shaded: stability field of the assemblage staurolite + garnet + biotite (+muscovite + quartz) (Spear and Cheney 1989)

Metabauxite is absent. The P-T estimate derived from the experimentally determined equilibrium curves of the reactions [26.1]–[26.5] indicates that the peak-metamorphic temperatures increased regionally from ca. 400 °C in the SE part of Naxos Island to >700 °C in the migmatite core (. Fig. 26.11). Judging from the intersection of the equilibrium curves [26.1] and [26.4a], the transformation of andalusite to kyanite, the minimum H2O-pressure attained in the c­hloritoid-corundum zone (II) was at 3 kbar, and rose to 5–7 kbar in zones IV, V and VI. Thus the average geothermal gradient, easily read off . Fig. 26.11, has only slightly changed over a distance of 20 km, from ca. 27 °C km−1 in the southeastern part of the island to ca. 31 °C km−1 in the migmatite core. In some orogenic belts, the geothermal gradient is considerably higher approaching values known from thermal aureoles. Ultimately, these temperature culminations, also described as thermal dome, are triggered by processes in the Earth’s mantle (7 Sect. 26.5.4). According to the concept of plate tectonics, thermal domes are formed at convergent plate margins above subduction zones, i.e., in island arcs or in Andean-type active continental margins (. Figs. 28.2, 29.18a–c), or in orogenic belts formed by continent-continent collision (. Fig. 29.18d). The heat transport is achieved by rising magma derived from partial melting of the mantle wedge above a subduction zone. Due to their lower density, these magmas can ascend from the mantle into the overlying continental crust.

The heat thus transferred upward leads to a bulge-like temperature distribution and consequently to prograde metamorphism and anatexis. By this process, granitic magma is formed in the lower crust that can rise up to higher crustal levels. An example is the granodiorite on the Island of Naxos that was emplaced subsequently to regional metamorphism, about 15 Ma ago, thereby dissecting the concentric zones of the metamorphic index minerals (. Fig. 26.10). Plutonic rocks of Miocene age are widely distributed in the Cyclades Metamorphic Complex thus defining a distinct high-temperature belt (e.g., Altherr et al. 1982). Part of the magmas produced by anatexis in the subducted lithospheric plate is erupted as lava, ignimbrite or volcanic ash, all of calc-alkaline composition. Volcanic rocks of this type are typical of island arcs and Andean-type orogens. However, not all thermal anomalies are related to convergent plate margins. As we have seen, continued submarine basalt volcanism that takes place at mid-ocean ridges formed at diverging oceanic plates. So called hot spots are persistent volcanic centres interpreted as the surface expression of rising mantle plumes. They can be associated with oceanic ridges but can also occur within oceanic or continental plates.

z High-Pressure and Ultra-High Pressure Metamorphism

Regional metamorphism affecting descending lithospheric plates is of distinctly different character. Due to subduction, the relatively cool oceanic sediments as well as the

26

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Chapter 26 · Metamorphic Rocks

mafic rocks constituting the oceanic crust are transported down to great depths, with relatively high velocities of several centimetres per year and are subjected to increasingly higher pressures. At the onset, this process is not connected with a significant temperature increase due to the poor thermal conductivity of rocks, which leads to downward bulging of the isotherms (. Fig. 28.2). By this process, the basalt and gabbro of the subducted oceanic crust are transformed into eclogite, a metamorphic rock of basaltic composition displaying the critical assemblage garnet + omphacite ± kyanite ± zoisite/epidote ± phengite. Basaltic oceanic crust less deeply subducted is transformed into blueschist, which contains the blue amphibole glaucophane together with lawsonite, jadeite or omphacite, phengitic white mica, in part also aragonite as index minerals. Sediments of the accretion wedge situated between the subducted oceanic plate and the overlying continental plate can be affected by high-pressure metamorphism as well. In that case, aragonite is formed in calcareous sediments and e.g., ferrocarpholite, (Fe,Mg)Al2[Si2O6](OH,F)4, in metapelites. Important reactions that can be used to constrain pressure are:

calcite ⇋ aragonite

[26.6]

and

Na[Al[4] Si3 O8 ] ⇋ NaAl[6] [Si2 O6 ] + SiO2 albite jadeite quartz

[26.7]

(. Fig 26.1). Typical blueschist is formed in a P-T range between about 7 kbar at 200–300 °C and 15 kbar at 400– 500  °C, conforming to a geothermal gradient around 10 °C km−1. Thus the absolute pressure value is by far less significant than the P/T ratio: At a temperature of 600 °C achieved in the staurolite zone of Naxos, a pressure of 6 kbar would clearly testify to medium-pressure metamorphism. Mineral assemblages in high-pressure rocks are only preserved if these were lifted quickly due to tectonic movements. Otherwise, subsequent heat supply would lead to an increase of the geothermal gradient and thus to the formation of medium-pressure assemblages. Oceanic basalt and gabbro and associated sedimentary rocks, metamorphosed under high-pressure conditions, are widely distributed in Alpine orogenic belts around the Pacific Ocean, such as the Franciscan Formation in California, the Shuksan Belt in the State of Washington, in Alaska, in Japan or in New Caledonia. High-pressure rocks also occur in the Eurasian orogenic belts, e.g., in the Alps, the Greek Cyclades, in Turkey or the Tian Shan at the borders of China, Kyrgyzstan and Kazakhstan. In orogenic belts of pre-Alpine age, blueschist and eclogite formed from subducted oceanic basalt or gabbro are much rarer. Here, metamorphic rocks of medium- to low-pressure type are clearly predominant. However, according to the principle of uniformity, plate tectonic processes should have governed older orogenies as well. Consequently, if high-pressure rocks had formed in older orogenic belts,

they would have been likely destroyed by later metamorphic overprint at lower P/T ratios. The oldest eclogite, so far known and presumably formed during subduction of an oceanic lithospheric plate has been recorded in the Proterozoic Usagara Belt in Tanzania, radiometrically dated at ca. 2 Ga (Möller et al. 1995). Perhaps, plate tectonic models are no longer applicable for formation of the Earth’s crust in Archaean times as in general, the geothermal gradient was higher during the early history of the Earth. Thus alternative interpretations involving a lower rate of horizontal movements combined with higher uplift rates have been proposed (e.g., Hamilton 1998). The extent to which plate tectonic processes shaped the 2.6–3.6 Ga old greenstone belts is a matter of debate (e.g., Shirey and Richardson 2011).

High-pressure rocks can be formed also during subduction of continental crust, a typical example being the Cyclades Metamorphic Complex, Greece. As mentioned above, the Permo-Mesozoic sediments and intercalated volcanic rocks were deposited onto a basement of pre-Alpine granites and gneisses. As a constituent of the Apulian microplate, this rock assemblage was subducted, in Eocene times, below the European plate and overprinted by high-pressure metamorphism. Blueschist, jadeite-bearing gneisses and eclogite formed at that stage are still preserved on some of the Cyclade Islands, such as Sifnos and Syros (e.g., Okrusch and Bröcker 1990) whereas elsewhere they are largely obliterated due to the subsequent prograde Barrovian metamorphism at the turn of the Oligocene to Miocene (see above). Extreme crustal thickening in the course of continent-continent collision can lead to ultra-high pressures. The presence of coesite, the high-P polymorph of SiO2, recorded in eclogitic rocks e.g., in the Dora Maira Massif, western Alps, first recorded by Chopin (1984), in the Erzgebirge, Germany, in the western gneiss region of Norway and in the Dabie Shan, PR China, indicates minimum pressures of 25–30 kbar conforming to burial of more than 100 km. Still higher pressures are indicated by the presence of diamond, such as in the Erzgebirge, western Norway, the Kokchetav Massif in Siberia, and the orogenic belts of Su-Lu and Dabie Shan, PR China (e.g., Carswell and Compagnoni 2003; Ogasawara 2005). At present, graphite-rich mudstone and carbonate sediments are deeply subducted in the Hindu-Kush zone due to the ongoing collision of the Indian and the Eurasian plates. There diamond- and coesite-bearing ultra-high pressure rocks should be formed today (Searle et al. 2001). z Paired Metamorphic Belts

Miyashiro (1972) was the first to recognise that, in young orogens, two parallel metamorphic belts of nearly the same age but of contrasting character can occur that are due to the same subduction or collision process: 5 a high-pressure belt with blueschist and/or eclogite that represents the subducted plate, and 5 a low- to medium-pressure belt with prograde mineral zones, migmatites, granitic to tonalitic intrusions and volcanic rocks of calc-alkaline character that formed at an active continental margin or an island arc. Examples are the high P/T Sanbagawa Belt and the low P/T Ryoke Belt in Japan, the Franciscan and the Sierra Nevada Complexes in California, USA, the Waikatipu Belt,

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26.3 · Nomenclature of Regional and Contact Metamorphic Rocks

New Zealand, and the Tasman Belt in Australia. The medium-pressure belt in the Cyclades finds its counterpart in the high-pressure belt of the external Hellenides that is documented by typical blueschist, aragonite- and lawsonite-bearing carbonate rocks and ferrocarpholite-bearing metapelites exposed on the island of Crete and the Peloponnese, Greece. This paired belt was formed by northeast-directed subduction at the turn of the Oligocene to the Miocene, an event that should be distinguished from the earlier subduction phase that produced the Eocene high-pressure rocks in the Cyclades (Altherr et al. 1982; Seidel et al. 1982). Currently, a third subduction phase is taking place at the southern margin of the Aegean Sea, south of Crete, where the Hellenic Trench is being formed. In the adjacent Cyclades Archipelago, the volcanic activity lasted until recently, a prominent example being the island group of Santorini. Its catastrophic Plinian eruption, ca. 3600 years ago, largely destroyed the initial, much higher island, formed a large caldera and contributed to the downfall of the Minoan civilisation in the Eastern Mediterranean. The small volcano Nea Kameni, formed during this event in the centre of the caldera, experienced several minor eruptions, the last three of which occurred in the 20th century.

In other cases, the mutual temporal relations do not conform to the paired belt model. For instance, the high-pressure metamorphism in the Western Alps is clearly older than the subsequent Lepontine phase that caused formation of a thermal dome in the Lepontine Alps. 26.2.6  Burial Metamorphism

Coombs (1961) was the first to recognise a special type of regional metamorphism that is due to burial of sediment packages without penetrative deformation and development of a schistosity. Commonly, metamorphic temperatures attained by burial are not very high. Thus metamorphic recrystallisation remains imperfect and relic minerals of the source rock are still preserved, thus making the distinction from diagenetic processes difficult. Metamorphic minerals formed during burial metamorphism are zeolites, e.g., laumontite, Ca[Al2Si4O12]·4.5H2O, and, at somewhat higher P-T conditions, prehnite, Ca2Al[AlSi3O10](OH)2, and pumpellyite, Ca2(Mg,Fe2+)(Al,Fe3+)2[SiO4/Si2O7](OH)2H2O. Burial metamorphism was first described from the southern island of New Zealand, where greywacke of Triassic age had been deposited within an elongate, subsiding trough that was subsequently affected by diagenesis and, finally, by weak metamorphic overprint. Similar occurrences have been recognised in other Phanerozoic orogens, but also in sedimentary troughs or basins of Proterozoic age, e.g., in northwestern Australia. In some rare cases, high-pressure metamorphism in subduction zones has taken place without penetrative deformation and thus is similar in character to burial metamorphism.

26.2.7  Ocean-Floor Metamorphism

The significance of this type of metamorphism has been realised by the surveys of the research vessel Glomar

Challenger as well as the international Deep Sea Drilling Program (DSDP) and Ocean Drilling Program (ODP) (Melson and van Andel 1966; Miyashiro et al. 1970, 1971). During these programs, samples of pillow basalt, basalt from the ­sheeted-dyke complex, and more rarely of gabbro and peridotite were recovered by dredging or submarine drilling. Some of these samples have been affected by metamorphic overprint, although they are undeformed and typically retain magmatic relic structures. With increasing metamorphic grade, the following minerals appear as new phases (Humphris and Thompson 1978; Gillis and Thompson 1993): 1. zeolites, such as analcime, Na[AlSi2O6]·H2O, heulandite, (Ca,Na,K)9[(Si,Al)36O72]·26H2O, natrolite, Na2[Al2Si3O10]·2H2O, mesolite, Na2Ca2[Al6Si9O30]·8H2O, and scolecite, Ca[Al2Si3O10]·3H2O; 2. prehnite, epidote, chlorite, calcite; 3. albite, actinolite to magnesiohornblende, epidote, chlorite, talc, quartz, titanite; 4. plagioclase, actinolite to magnesiohornblende, epidote, chlorite, biotite, quartz, titanite. Due to the limited thickness of the oceanic crust of some 5–6 km, pressures are low during ocean-floor metamorphism. The temperatures necessary for metamorphism are reached only at mid-ocean ridges where the ongoing ascent of basaltic magma leads to a high geothermal gradient (. Fig. 28.3). In contrast, a normal geothermal gradient is attained in most of the oceanic crust with maximum temperatures of only 100– 200 °C, which are too low for proper metamorphic reactions. An important characteristic of ocean-floor metamorphism are distinctive changes in bulk-rock composition, caused by a chemical exchange between the basaltic rocks and circulating, heated seawater. By these processes black and white smokers effuse and hydrothermal ore minerals can precipitate on the ocean floor (7 Sect. 23.5.1). 26.3  Nomenclature of Regional and Contact

Metamorphic Rocks

The nomenclature of metamorphic rocks is based on structure and mineral assemblage. Collective terms such as phyllite, schist, gneiss, granulite or fels are applied to emphasise the structure. These are specified by addition of characteristic minerals, e.g., staurolite-micaschist, and/ or special structural features, e.g., banded gneiss. Due to the interplay of deformation and recrystallisation, most regional metamorphic rocks display a preferred orientation of inequant fabric elements, especially of platy and prismatic minerals. This leads to excellent planar or linear parallel textures, foliation, e.g., schistosity (S), and lineation (B), e.g., fold axes. In contrast, parallel textures are gradually deteriorated by contact metamorphic overprint ultimately producing massive, homogeneous textures, typical of hornfels. Other groups of metamorphic rocks are named according to

26

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Chapter 26 · Metamorphic Rocks

their mineral content notwithstanding their fabric, e.g., amphibolite or quartzite. In contrast to igneous rocks, names of metamorphic rocks are hardly ever derived from localities. Local names are used almost exclusively in a local or regional context, such as beerbachite in the Odenwald, Germany, Bündner Schiefer in the Swiss Alps or Macduff slates in Scotland. A quantitative classification system for metamorphic rocks has been suggested by Fettes and Desmons (2007).

26

26.3.1  Regional Metamorphic Rocks z Slate

Slate is an argillaceous sedimentary rock that was affected by slight regional metamorphism and deformation thus producing a distinct slaty cleavage that can be parallel to or transecting the initial bedding. The extremely ­fine-grained rock splits into thin plates, that can be used as roof tiles, whereas columnar pencil slate results from two cleavage planes intersecting. Principal constituents are sheet silicates, especially sericite and chlorite as well as quartz and/or carbonate minerals. Their respective modal amounts have a strong influence on the technical properties and the practical application of slates. Thus roof slates, mainly used for production of roof tiles, are relatively rich in quartz, whereas the softer (former) school slates contain higher amounts of calcite. In some occurrences, slate contains porphyroblasts of pyrite, e.g., the Ballachulish slate of Scotland. In the past, slate was mined, in large quantities, in the Harz Mountains, the Thuringian Forest and the Rheinisches Schiefergebirge (=Rhenish Slate Mountains) in Germany, in Wales, Cornwall, Cumbria and Scotland, UK, in the Ardennes of Belgium and France, in Liguria, northern Italy, in the Slate Valley of Vermont, in Maine, Pennsylvania and Virginia, USA. At present, the world’s largest slate producer is Spain, followed by Brazil. Huge slate deposits are being mined in China. z Phyllite

The term phyllite comprises a group of fine-grained, schistose metasedimentary rocks, intermediate in grain size and metamorphic grade between slate and micaschist. The sheet silicates are arranged to form a continuous coating on the cleavage surface thus imparting a characteristic silky sheen. Many phyllites have undergone smallscale folding or developed a crenulation cleavage. Apart from individual porphyroblasts, e.g., of albite, chloritoid or carbonates, minerals in phyllite reach a grain size not exceeding 0.1 mm and thus cannot be identified under a magnifying lense. In phyllite proper, sheet silicates, especially the white micas sericite and paragonite, chlorite and/ or, more rarely, biotite, constitute >50 vol.-% of the rock, followed by quartz. Additional constituents can be albite, chloritoid, spessartine-almandine rich garnet, calcite, dolomite, ankerite and others, which can lead to special names such as albite phyllite or chloritoid phyllite. Quartz

phyllite contains 50–80 vol.% of quartz, calcareous phyllite 10–50 vol.% of carbonate minerals. Protoliths of phyllite are pelitic, Al-rich sediments, such as mud, silt or muddy sand, partly with siliceous or calcareous admixtures. z Micaschist

The micaschist group comprises medium- to coarse-grained metapelitic rocks with distinct schistosity and perfectly splitting into mm- to cm-thick slices parallel to the cleavage planes S, or in thin columns along the fold axes B (or f). In contrast to phyllite, most of the individual constituents can be recognised with the unaided eye or with a magnifying lens. Micas, especially muscovite and biotite, more rarely paragonite, are present in proportions of >50 vol.%, followed by quartz. Characteristic minor constituents can lead to special names, such as garnet-micaschist or staurolite-micaschist (. Figs. 2.10, 26.12a). Quartz-micaschist contains 50–80 vol.% of quartz, calcareous micaschist 10–50 vol.% calcite, dolomite or ankerite. Protoliths are the same as in phyllite. In typical micaschists, only small amounts of feldspar are present, generally 30 vol.%, quartz 90 vol.%, commonly together with minor or accessory muscovite or sericite as well as chlorite, garnet, kyanite, sillimanite, tourmaline, graphite and others. In sericite quartzite, garnet quartzite, graphite quartzite, these additional minerals are present in proportions of >10 vol.%, whereas quartz phyllite and quartz-micaschist contain >20 vol.% of sheet silicates. Fine-grained rocks consisting of quartz + spessartine-rich garnet are known as coticule. These were used in the past for sharpening of cutting tools, e.g., at Viel-Salm in the Belgian Ardennes. Quartzite can form during both regional and contact metamorphism, from siliceous sandstone or chert, whereby recrystallisation of the detrital quartz grains and the siliceous matrix leads to an overall increase in grain size and the development of a granoblastic texture with increasing metamorphic grade. z Marble

Marble is a medium- to coarse-grained metamorphic rock consisting to >90 vol.% of carbonate minerals, such as calcite (. Fig. 16.14) and/or dolomite, more rarely ankerite. It results from metamorphism, be it regional or thermal, of relatively pure carbonate rocks. Common accessories are graphite and phlogopite. Metamorphic overprint of marly limestone leads to the formation of silicate marble with >10 vol.% of silicate minerals, such as phlogopite, tremolite, diopside, vesuvianite, grossular, epidote, chondrodite, Mg5[SiO4]2(OH,F)2, and/or forsterite. Some marbles

477

26.3 · Nomenclature of Regional and Contact Metamorphic Rocks

contain spinel and/or corundum, in places of gem quality. Ophicalcite is a marble that also contains serpentine minerals concentrated in layers or patches. Marble has been a very popular dimension stone that can be found in the exterior and interior parts of many historic representative buildings. Due to its high sensitivity to weathering, however, marble is nowadays almost exclusively used for interior walls and floors. Many statues and sculptures are made of marble because of its relatively low hardness. Famous marble occurrences that have been quaried since ancient times are on the Cycladic Island of Paros and the Pentelikon Mountain Range near Athens in Greece as well as near Carrara in the Tuscany region of Italy, where marble is produced e.g., under the trade names statuario and arabescato. z Calc-Silicate Rocks and Calc-Silicate Gneisses

If the content of silicate minerals in an otherwise calcareous rock increases to >50 vol.%, it is referred to as calc-silicate rock if massive or calc-silicate gneiss if layered. Both consist predominantly of Ca- and Ca–Fe–Mg silicates, such as diopside-hedenbergitess, grossular-andradite-dominated garnet, vesuvianite, tremolite, forsterite, wollastonite ± quartz and variable amounts of calcite. Protoliths are impure limestone or marl. Fe-rich calc-silicate rocks are called skarn (7 Sect. 26.3.2). z Amphibolite

Amphibolite is a medium- to coarse-grained metabasic rock that consists predominantly of amphibole, in most cases green or brown hornblende, and plagioclase. Commonly, diopsidic clinopyroxene, garnet, epidote or zoisite, biotite, quartz and titanite are present as minor phases and, at higher proportions, give rise to more specific names, such as epidote amphibolite. Commonly, the amphibole grains have a lattice-preferred orientation and thus define a distinct foliation, leading to slab jointing of the rock. If, in addition, the amphibole grains are predominantly aligned parallel to B, the amphibolite breaks along columnar joints into elongated prism-like fragments. Some amphibolites can be, however, massive, lacking any preferred orientation. Intermediate between amphibolite and gneiss is hornblende-plagioclase gneiss, which contains >20 vol.% quartz. Metabasites poor in, or virtually devoid of, plagioclase are called hornblende fels or hornblende schist. The protoliths of the majority of amphibolite occurrences are basalt or andesite, their related pyroclastites, as well as gabbro. z Greenschist and Greenstone

Greenschist is a collective term for green-coloured, finegrained metabasites with distinct schistosity, which is essentially defined by green (OH)-bearing minerals, such as chlorite, epidote and actinolite, together with albite ± quartz ± carbonate and ± white mica, a mineral assemblage characteristic of low metamorphic grade. Some greenschists may contain porphyroblasts, e.g., of actinolite,

garnet or magnetite (. Fig. 2.9). Non-schistose metabasite of the same composition is termed greenstone. Protoliths of greenstone are mafic volcanic and plutonic rocks that became metamorphosed without much deformation. Greenstones are particularly abundant in Archaean crustal rocks. z Blueschist, Glaucophane Schist

Blueschist is a metabasite of deep blue to greenish blue colour, typically formed during high-pressure, low-grade metamorphism. Characteristic minerals are blue amphiboles, such as glaucophane, ferroglaucophane or crossite, accompanied by pumpellyite, lawsonite, chlorite, phengitic white mica and albite. At higher temperatures, almandine-rich garnet is formed and lawsonite is replaced by epidote (. Fig. 26.13a). At higher pressures, jadeite crystallises at the expense of albite. Due to a strong lattice preferred orientation of glaucophane needles parallel to S, in places also parallel to B, glaucophane schist displays a distinct schistosity leading to slab or columnar jointing. For massive types without preferred orientation the term glaucophanite should be used. Analogous to amphibolite, protoliths are mafic volcanic and plutonic rocks. z Eclogite

Eclogite is a medium- to coarse-grained, massive to banded metabasite that is essentially made up of green omphacite, an augite-jadeite solid solution, and almandine- and pyrope-rich garnet with considerable ­ grossular content. Additional constituents can be quartz, kyanite, zoisite, phengitic white mica and rutile. Eclogite associated with blueschist can contain epidote instead of zoisite, and titanite instead of rutile. Eclogite has a high density of 3.3–3.5 g cm−3 and is derived from oceanic basalt or gabbro (density ca. 3.0 g cm−3) by metamorphism at high-pressure or ultra-high pressure conditions, the latter being defined by the presence of coesite. z Serpentinite and Other Metamorphic Rocks of Ultramafic Composition

Serpentinite is a very fine-grained, massive to schistose metamorphic rock of ultramafic composition and commonly of dark green colour. It consists predominantly of the serpentine minerals lizardite, antigorite and/or chrysotile and commonly contains some magnetite, talc, chlorite, amphibole and/or carbonate minerals. In many cases, serpentinite contains mineral relics of olivine, orthopyroxene, clinopyroxene and garnet, all of which with a highly magnesian composition (. Fig. 26.13b). These relics indicate that serpentinite forms by retrograde metamorphic overprint of ultramafic protoliths, such as (garnet-)peridotite. As serpentine minerals are very hydrous, the process of serpentinisation requires addition of H2O to the ultramafic protolith. When serpentinite undergoes further prograde metamorphism, a series of dehydration reactions take place that eventually produce (OH)-free minerals, such as olivine, ortho- and

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Chapter 26 · Metamorphic Rocks

26

. Fig. 26.13  Photomicrographs of metamorphic rocks. a Garnet glaucophane schist, consisting of garnet (pink), deformed in a pear-shaped way, glaucophane (light blue), epidote (yellowish) and phengite (colourless with cleavage fissures). Northern coast of Samos Island, Greece. Plane polarised light. Width of view ca. 4 mm. b Pyrope serpentinite displaying typical reticulate (mesh) texture, formed by the serpentine minerals lizardite and chrysotile (light grey interference colours). These replace the minerals of the former peridotite except for minor relics of olivine (mostly with high interference colours) and orthopyroxene (rare, with dark grey interference colours, e.g., lower right), whereas pyrope-rich garnet is almost completely preserved. Zöblitz, Saxonian Erzgebirge, Germany. Crossed polars. Width of view ca. 5 mm

479

26.4 · Structure and Texture of Metamorphic Rocks

. Fig. 26.14  Marble, even-grained rock consisting of calcite with ¯ Carrara, Tuscany, Italy. Width of view ca. lamellar twinning ||{0110}. 2.3 mm

clinopyroxene as well as garnet, thus leading to a highgrade metamorphic mineral assemablage that corresponds to that of the igneous protolith but displays a different texture (. Fig. 26.14). As serpentinisation requires access of considerable amounts of H2O from an external source, this alteration is typically accompanied by metasomatism wherever there is a steep chemical gradient between different rock types. Such steep gradients exist, for instance between serpentinite and felsic rocks, such as pegmatite, granite or granite gneiss. Along their contacts, infiltration and diffusion metasomatism involves mass transfer, which in turn can produce monomineralic rocks of ultramafic composition, such as talc, chlorite, actinolite and biotite schists. Such monomineralic shells of biotite and/or amphibole are referred to as blackwall (7 Sect. 26.6.1) and can be important hosts of the gem minerals emerald and alexandrite. Technical application: Cut and polished slabs of serpentinite and serpentinite breccia are used as dimension stone and for artworks. 26.3.2  Contact Metamorphic Rocks z Hornfels

Hornfels s.str. is a massive, very fine- to fine-grained rock that formed by peri-plutonic contact metamorphism (7 Sect. 26.2.1). The German name “hornfels” describes the typical splintering fracturing, at which individual chips are transparent on edges. Due to the thermal influence of an intruding magma, the country rock can become totally recrystallised to form a granoblastic mosaic structure, which is typical of hornfels. The relatively short time span available for contact metamorphism often prevents textural equilibrium to be achieved, resulting in a massive fine-grained rock, with or without foliation/bedding, in whose matrix larger, inclusion-rich contact metamorphic minerals are embedded in the form of poikiloblasts. The

latter commonly consist of andalusite or cordierite and give rise to the rock name “spotted” hornfels. Depending on the metamorphic grade and time available for metamorphic recrystallisation, hornfelses can still contain remnants of primary bedding of a sedimentary country rock or foliations and lineations of a pre-existing regional-metamorphic rock. If the extent of contact metamorphic recrystallisation was low, and primary structures such as bedding are well preserved while newly formed minerals, such as biotite, andalusite (often of the chiastolite variety) or cordierite forming distinct, randomly orientated porphyroblasts or poikiloblasts, the rock can be a referred to as spotted slate (. Fig. 26.3). In a broader sense, the term “hornfels” is used for any rock that formed under medium- to high-grade contact metamorphic conditions. According to the chemical composition of the precursor rock and the minerals formed, the following hornfels types can be distinguished: 5 metapelitic hornfels: with andalusite/sillimanite, cordierite, biotite, muscovite or K-feldspar, quartz, ± plagioclase, e.g., andalusite-cordierite hornfels; protoliths are siltstone, mudstone, shale; 5 calc-silicate hornfels: with diopside, grossular, vesuvianite, wollastonite, epidote, ± plagioclase, ± calcite; protoliths: marly limestone or lime marl; 5 hornblende hornfels: with hornblende, plagioclase, ± biotite, ± calcite; 5 pyroxene hornfels: with orthopyroxene and/or clinopyroxene, plagioclase (. Fig. 26.15b); protoliths for both these hornfels types are basaltic rocks and their pyroclastites, gabbro or amphibolite; 5 ultramafic hornfels, Mg-rich (. Fig. 26.15a): with olivine, talc, amphibole such as tremolite, cummingtonite or anthophyllite, enstatite, ±chlorite, ±spinel; protolith: serpentinite. z Skarn

Skarn is formed by metasomatic reactions, caused by chemical exchange between limestone or dolomite and an intruding, usually acidic magma. Typical skarn is a coarsegrained rock predominantly consisting of Ca–Fe silicates such as hedenbergite, Fe-rich amphibole, andradite-rich garnet and epidote as well as rarer silicate minerals, intergrown with sulfidic, oxidic or other ore minerals. In the USA, skarn is known as tactite. Deposits of skarn ore can be of high economic significance (7 Sect. 23.3.1) 26.4  Structure and Texture of Metamorphic

Rocks

The structural and textural properties of a metamorphic rock are of particular interest because they reflect its, in many instances complicated, formation history. When unravelling the history of a metamorphic rock, information about the protolith, the P-T conditions during metamorphism as well as the interplay between

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26 . Fig. 26.15  a Ultramafic, Mg-rich hornfels, formed by contact metamorphic overprint of serpentinite. Main constituents are finger-shaped porphyroblasts of olivine (dark pattern), talc (light) and some magnetite (opaque). Kirchbühl near Erbendorf, Upper Palatinate, Germany; Plane polarised light; Width of view ca. 12 mm. b Pyroxene hornfels (so called beerbachite), formed by contact metamorphic overprint of an amphibolite; Major constituents are An-rich plagioclase (light), hypersthene and diopside (dark) as well as magnetite and ilmenite (opaque). Magnetsteine near Nieder-Beerbach, Odenwald Metamorphic Complex, Germany; Plane polarised light; Width of view ca. 5 mm

metamorphic recrystallisation and deformation can be read off the rock. 5 By definition, a metamorphic rock is derived from a pre-existing protolith that can be igneous, sedimentary or metamorphic. Structural or textural relics of this protolith can remain preserved during metamorphic overprint and show through the new metamorphic fabric, similar to a palimpsest of an illuminated manuscript, from which the initial handwriting has been abraded and overwritten. 5 During conventional metamorphism, without reaching the stage of partial melting, new minerals form essentially by solid state reactions, leading to metamorphic textures that are distinctly different from igneous or sedimentary ones (Becke 1903). 5 During regional metamorphism, mineral grains commonly grow in a stress field, thus leading to deformation and the development of a lattice preferred orientation of newly formed metamorphic minerals.

26.4.1  Remnants of Protolith Structures

In spite of intensive recrystallisation, many metamorphic rocks display remnants of original structures inherited from their protolith. Their preservation is not so much a function of metamorphic grade but more so of intensity of ­syn-metamorphic deformation and strain. For instance, in sillimanite-bearing metaturbidites of the Pan-African Damara Belt in Namibia, which have experienced metamorphic temperatures of 600–660 °C and H2O pressures of 3–4 kbar, Kukla et al.

(1990) have described a broad spectrum of sedimentary structures, such as graded bedding, cross bedding, current marks and load casts (. Fig. 26.16), all of which permit the reconstruction of a sequence of several Bouma cycles (7 Sect. 25.2.9). Significantly, the sedimentary structures were much better preserved in psammitic beds, forming low strain zones, rather than in more intensely deformed pelitic beds, which acted as high-strain zones. Understandably, sedimentary structures became totally obliterated as soon as incipient melting took place in the higher grade migmatite zone of the Damara Belt.

z Sedimentary Relics

In general, the sequence of interlayered metamorphic rocks of different composition, e.g., alternating micaschist and marble beds, provide conclusive evidence of a sedimentary origin. Extremely resistant against metamorphic overprint are coarse-grained sedimentary structures such as big cobbles in clastic sediments. Instructive examples are the conglomerate gneisses of Mittweida in Saxony, Germany. In rare cases, even fossils initially present in a sedimentary rock can be preserved during metamorphic overprint. For instance, near the Lukmanier Pass in the Swiss Alps, fragments of crinoids, belemnites, bivalve molluscs and corals can still be recognised in Liassic limestone and marl, that have been metamorphosed during the Alpine Orogeny at temperatures around 550 °C and H2O pressures of >5 kbar (Nabholz et al. 1967). Herbal microfossils, such as pterophytes and spores, have been recorded in metasedimentary rocks, e.g., in the Spessart Metamorphic Complex, Germany, and can be even used for biostratigraphic dating of the rock (Reitz 1987). z Magmatic Relics

Relics of primary field relationships between igneous intrusions, e.g., former dykes or sills, and their country rock can provide conclusive evidence of a magmatic origin of the former. Pillow structures, exemplarily exposed in the high-pressure rocks at Zermatt, Switzerland, testify to submarine basalt as a protolith. Orthogneiss derived from

26.4 · Structure and Texture of Metamorphic Rocks

481

. Fig. 26.16  Trough crossbedding and flute casts (on top of the pocket knife) in a stauroliteand sillimanite-bearing metaturbidite of the Kuiseb Formation, Damara orogenic belt, Khomas Highlands, Namibia (Photograph: Peter Kukla, RWTH Aachen)

granite, granite porphyry or rhyolite with porphyritic texture can contain deformed phenocrysts of K-feldspar that are embedded as porphyroclasts in a blastomylonitic matrix. Well-known examples are the porphyroids and the orthogneisses of the Variscan basement of Germany and the so-called central gneisses of the Tauern Window in the Austrian Alps. Coarse-grained structural relics found in metamorphosed gabbro are also convincing evidence of its intrusive origin. Spectacular examples are the metagabbro exposures on the islands of Syros and Samos, Cycladic blueschist belt, Greece. Amphibolite can contain relics of igneous augite phenocrysts or of former ophitic texture derived from basaltic protoliths. 26.4.2  Metamorphic Textures

The Austrian petrologist Friedrich Becke (1855–1931) first recognised that metamorphic rocks display characteristic textures, testifying to solid state reactions, at which minerals grow at the same time in close contact and competing with each other for space. In contrast, igneous textures indicate successive crystallisation of different minerals from a magma, from which the earliest formed minerals could grow without hindrance. Becke (1903) characterised metamorphic textures as “crystalloblastic” (from Greek βλἀστη = sprout, growth) and indicated terms related to metamorphic textures by adding the suffix “blast” (see below). As most metamorphic processes take place at higher lithostatic pressure, metamorphic rocks are always compact and never cavernous or cellular. Moreover, as metamorphism does not involve chilling of magma, nearly all metamorphic rocks are holocrystalline and do not contain skeletal crystals. Only pseudotachylyte, impact breccias, especially suevite, or buchite can contain rock glass.

Most metamorphic rocks, especially regionally metamorphosed ones, display a textural equilibrium assemblage of minerals, that is, all of the various mineral phases occur in mutual contact. If textural equilibrium was achieved, one can assume that the same mineral assemblage also represents a thermodynamic equilibrium, which was typically attained at peak metamorphic conditions.

As said above, metamorphism does not involve a distinct succession in the crystallisation of minerals, as is typical of igneous rocks. Megacrysts of metamorphic origin are not phenocrysts, early crystallised from magma, but have grown as porphyroblasts in a fully crystalline matrix. Many of them contain inclusions of minerals formed as internal relics on the prograde P-T path. Elongate groups of inclusions can indicate older helicitic fabrics, such as bedding, schistosity or microfolding, that provide important evidence of the geological history of a given rock. Along the retrograde P-T path, the minerals that had formed at the metamorphic peak can become replaced by secondary minerals, although this process is commonly incomplete for kinetic reasons and does not lead to thermodynamic equilibrium (7 Sect. 26.1.1). At favourable conditions, comparable mineral sequences are recorded within the same metamorphic rock. These can document one or more additional P-T path(s), passed through at successive metamorphic events. A metamorphic texture is called granoblastic (Latin granum = grain), if the predominant minerals are of isometric shape and display no preferred orientation. Likewise, leaf-shaped, columnar or fibrous minerals as well as textures dominated by them have been designated as lepidoblastic, nematoblastic or fibroblastic in Becke’s nomenclature. However, these terms are dispensible and should be avoided (Passchier and Trouw 2005). The same holds true for the terms idioblastic, subidioblastic or xenoblastic applied to minerals that have developed or not

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. Fig. 26.17  Common metamorphic textures, mainly caused by thermal metamorphism. a–c Granoblastic polygonal: a quartz, b quartz + biotite, c quartz + pyroxene; d decussate: amphibole; e poikiloblastic: albite; f kelyphitic reaction rim of clinopyroxene + spinel around pyrope-rich garnet; garnet peridotite from Gorduno, Ticino, Switzerland (a–e from Spry 1983)

developed crystallographic faces during their growth. Here the simpler terms euhedral, subhedral or anhedral are fully sufficient.

Becke (1903) has already constituted his “crystalloblastic series”, in which metamorphic minerals are arranged in order of decreasing formation energy, i.e., according to their potential to develop their crystal forms against the resistance of a surrounding solid medium (Eskola 1939). A preference for euhedral crystal growth is noted especially in ortho- and ring silicates, such as garnet (. Fig. 2.10), titanite, staurolite (. Figs. 2.10, 11.15), kyanite, andalusite, zircon or topaz, but also in pyrite, magnetite (. Fig. 2.9) or spinel. During their crystal growth, chain and sheet silicates can develop only some of their faces, such as {110} in amphiboles, and {001} in micas or chlorite. Trigonal carbonate minerals can form porphyroblasts showing the rhombohedron {101¯1}. In contrast, most framework silicates such as quartz, feldspars or cordierite commonly crystallise as anhedral grains. Textures of metamorphic minerals are characterised according to their type of intergrowths. ­Granoblastic-polygonal texture (. Fig. 26.17a–c) is frequently developed in hornfels (26.15b), quartzite, marble (26.14), mica-poor gneiss or granulite. If this texture is fully developed, three adjacent grains commonly meet at an angle of nearly 120° thus forming a triple point (. Fig. 26.17a). Decussate texture is observed in randomly oriented intergrowths of elongate minerals, such as amphiboles or micas, (. Fig. 26.17d). In other cases, an intimate penetration of different mineral species is observed. Many porphyroblasts, e.g., albite, garnet or staurolite, contain numerous inclusions of earlier formed minerals, thus producing a sieve-like texture, also known as poikiloblastic (Greek πóικιλος = different) (. Fig. 26.17e). Symplectites are fine-grained lamellar or vermiform intergrowths of two or more mineral phases,

. Fig. 26.18  Rotational laminar slip illustrated by the pages of a book. a Undeformed; b deformed: right: homogeneous, left: inhomogeneous deformation (after Sander from Eskola 1939)

in most cases produced by retrograde reactions. Myrmekite is a symplectite consisting of plagioclase and vermiform quartz developed in contact with K-feldspar. During retrograde breakdown, mineral grains can be surrounded by radial, fine-grained poly-mineralic or mono-mineralic reaction rims, a feature known as corona texture (. Fig. 27.1). A prominent example is kelyphite (Greek κἐλυφος = nutshell), a fine-grained, radial reaction rim consisting of clinopyroxene or amphibole + spinel, formed by the breakdown of pyrope-rich garnet in contact with olivine (. Fig. 26.17f). 26.4.3  Strain-Induced Preferred Orientation

of Metamorphic Minerals

z Basic Principles

Static recrystallisation, typically realised in thermal aureoles, produces textures that lack any preferred orientation. In contrast, regional metamorphism commonly takes place in a differential stress field, which leads to deformation and, due to alignment of syn-tectonic tabular or prismatic mineral grains, to the preferred orientation of these minerals. Under strain, the shape and volume of a fabric element, such as a mineral or a mineral group, is changed, e.g., a sphere is distorted into an ellipsoid, the strain ellipsoid with the principal axes X, Y and Z. The term deformation is used in a broader sense: In addition, deformation processes include translation and rotation of fabric elements (Passchier and Trouw 2005). During deformation of rocks or minerals, laminar slip of thin, coherent lamellae such as slip along shear planes or intracrystalline plastic slip (. Fig. 26.26c) plays an important role. Among these processes, homogeneous and inhomogeneous deformation can be distinguished, as exemplified by . Fig. 26.18: Different symbols are drawn on the section of a thick paperback that has been deformed by turning its spine upwards. In the right part of the paperback, each page has slipped, relative

26.4 · Structure and Texture of Metamorphic Rocks

to the following page, by a certain amount: it has undergone homogeneous deformation. Obviously, this is not the case in the left, upward bent part of the paperback. 5 When subjected to homogeneous deformation, straight, parallel lines or planes remain straight, parallelograms remain parallelograms, parallelepipeds remain parallelepipeds, whereas circles are transformed into ellipses, and spheres into ellipsoids (Sander 1950). 5 In contrast, inhomogeneous deformation has affected the left part of the paperback, whereby its pages were folded, straight lines or planes bent and circles transformed into pear-shaped figures rather than ellipses. Our paperback model can be well applied to natural rocks, in which smooth transitions between both types of deformation can be observed as well. Commonly, inhomogeneous deformation, such as folding by bending or buckling (. Fig. 26.18b, left), affects larger rock volumes than homogeneous deformation. Porphyroblasts started to roll and layers of helicitic inclusions can be used to reconstruct the initial position of the porphyroblast and to determine the axis and the angle of rotation (. Fig. 26.23). During homogeneous deformation, two different types of shear can be distinguished, simple shear and pure shear. Commonly, however, there are transitions between these two extreme types, known as common shear. 5 In the case of simple shear (. Fig. 26.19a), deformation takes place along a plane of constant orientation,

483

. Fig. 26.20  Fabric or reference axes x, y and z in a hand specimen with slight folding along the fold axis b (=y); xy is the foliation plane

forming an angle with the axes of shorting Z and stretching X, respectively. The particles of a fabric element move in directions parallel to shear planes, and the axes X and Z of the strain ellipsoid rotate into the shear direction. Consequently, simple shear is not coaxial to X and Z. Convincing examples are shearing a deck of cards into one direction (. Fig. 26.18b, right). 5 In the case of pure shear (. Fig. 26.19b), the fabric elements are shortened along the Z-axis of the strain ellipsoid and extended along X, i.e., perpendicular to Z, whereby the particles move along curved tracks that proceed symmetrically to the axes of the strain ellipsoid. Stretching along X with X > Y = Z leads to prolate, cigarlike shape, whereas flattening on YX with Z  Ca > Al > K. Moreover, withdrawal of SiO2 (desilication) can lead to decreasing quartz contents, even to the point of SiO2-deficiency and the formation of corundum. Infiltration of mobile minor elements can lead to the crystallisation of apatite and tourmaline in the phlogopite zone. In some rocks of granitic composition, the rare element Be can be concentrated to such an extent that the ideally colourless beryl, Be3Al2[Si6O18], and chrysoberyl, BeAl2O4, form. When mobile Be infiltrates into ultramafic country rock, the green-coloured varieties emerald and alexandrite can crystallise in biotite/phlogopite or chlorite schist. Their colour is then due to the uptake of some Cr derived from the former serpentinite (. Fig. 26.33). Encounter of the two chemical elements Be and Cr, that commonly do not occur together, was first described in 1928 and interpreted by the Russian geochemist Alexander Y. Fersman (1883–1945) in the famous emerald deposit of the Tokovaya, Ural Mountains (. Fig. 26.34). A similar occurrence is the emerald-alexandrite deposit of Novello Claims in Zimbabwe. A typical blackwall zonation is also developed in the Habachtal, Salzburg, Austria, which constitutes Europe’s largest emerald occurrence (Grundmann and Morteani 1989). There Be was derived from garnet-micaschist and biotite-albite gneiss, whereas pegmatite is absent.

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. Fig. 26.34 Metasomatic reaction zones (blackwalls) developed between pegmatite veins and serpentinite. Cross section through the emerald deposit of Tokovaya, Ural Mountains, Russia (modified after Fersmann 1929 from Schneiderhöhn 1961)

26

26.6.2  Autometasomatism

The term autometasomatism comprises all chemical reactions that take place within an igneous body subsequent to its crystallisation. During these processes, aqueous fluids of a wide range of temperatures (as low as 4 Ga zircons. Annual Rev Earth Planet Sci 37:479–505

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Holdaway MJ, Mukhopadhyay B (1993) A reevaluation of the stability relations of andalusite: thermochemical data and phase diagram for the aluminum silicates. Am Mineral 78:298–315 Holland TJB (1980) The reaction albite = jadeite + quartz determined experimentally in the range 600–1200 °C. Am Mineral 65:129–134 Humphris SE, Thompson G (1978) Hydrothermal alteration of oceanic basalts by seawater. Geochim Cosmochim Acta 42:107–125 Jaeger JC (1957) The temperature in the neighborhood of a cooling intrusive sheet. Am J Sci 255:306–318 Jaeger JC (1959) Temperatures outside a cooling intrusive sheet. Am J Sci 257:44–54 Jansen JBH, Schuiling ED (1976) Metamorphism on Naxos: petrology and geothermal gradients. Am J Sci 276:1225–1253 Johannes W (1988) What controls partial melting in migmatites? J Metam Geol 6:451–465 Jung S, Mezger K (2001) Geochronology in migmatites—a Sm-Nd, U-Pb and Rb-Sr study from the Proterozoic Damara Belt (Namibia): implications for polyphase development of migmatites in high-grade terranes. J Metam Geol 19:77–97 Kaur P, Chaudhri N, Hofmann AW et al (2012) Two-stage, extreme albitization of A-type granites from Rajasthan, India. J Petrol 53:919–948 Kresten P, Morogan V (1986) Fenitization of the Fen Complex, Southern Norway. Lithos 19:27–42 Kukla PA, Kukla C, Stanistreet IG, Okrusch M (1990) Unusual preservation of sedimentary structures in sillimanite-bearing metaturbidites of the Damara Orogen, Central Namibia. J Geol 98:91–99 Kukla C, Kramm U, Kukla PA, Okrusch M (1991) U-Pb monazite data relating to metamorphism and granite intrusion in the north-western Khomas Trough, Damara Orogen, Central Namibia. Communs Geol Surv Namibia 7:49–54, Windhoek, Namibia Langenhorst F, Deutsch A (1998) Minerals in terrestrial impact structures and their characteristic features. In: Marfunin AS (ed) Mineral matter in space, mantle, ocean floor, biosphere, environmental management, and jewelry. Advanced Mineralogy, vol 3, pp 95–119 Lippmann F (1977) Diagenese und beginnende Metamorphose bei Sedimenten. Bull Acad Serbe Sci Nat, T LVI, No 15 Melson WG, van Andel TH (1966) Metamorphism in the Mid-Atlantic Ridge, 22° N latitude. Marine Geol 4:165–186 Miyashiro A (1972) Metamorphism and related magmatism in plate tectonics. Am J Sci 272:629–656 Miyashiro A, Shido F, Ewing M (1970) Petrologic models for the Mid-Atlantic Ridge. Deep Sea Res 17:109–123 Miyashiro A, Shido F, Ewing M (1971) Metamorphisn in the Mid-Atlantic Ridge near 24° and 30°. Phil Trans Roy Soc London A268:589– 603 Möller A, Appel P, Mezger K, Schenk V (1995) Evidence for a 2 Ga subduction zone: eclogites in the Usagaran Belt of Tanzania. Geology 23:1067–1070 Nabholz WK, Niggli E, Wenk E (1967) Lukmanier-Pass: Disentis–Biasca, Exkursion Nr. 23. In: Nabholz WK (ed) Geologischer Führer der Schweiz. Schweizerische Geologische Gesellschaft, Wepf & Co., Basel, Switzerland, Heft 5:400–417 Okrusch M, Bröcker M (1990) Eclogites associated with high-grade blueschists in the Cycladic Archipelago, Greece: a review. Eur J Mineral 2:451–478 Olsen SN (1985) Mass balance in migmatites. In: Ashworth JR (ed) Migmatites. Blackie, Glasgow, London Phillips FM, Zreda MG, Smith SS et al (1991) Age and geomorphic history of meteor crater, Arizona, from cosmogenic 36Cl and 14C in rock varnish. Geochim Cosmochim Acta 55:2695–2698 Read HH (1952) Metamorphism and migmatisation in the Ythan Valley, Aberdeenshire. Trans Edinburgh geol Soc 15:265–279 Reitz E (1987) Palynologie in metamorphen Serien: I. Silurische Sporen in einem granatführenden Glimmerschiefer des Vor-Spessart. Neues Jahrb Geol Paläont Monatsh 1987:699–704

Richardson SW (1968) Staurolite stability in a part of the system Fe–Al– Si–O–H. J Petrol 9:467–488 Robyr M, Vonlanthen P, Baumgartner LP, Grobety B (2007) Growth mechanism of snowball garnets from the Lukmanier Pass area (Central Alps, Switzerland): a combined µCT/EPMA/EBSD study. Terra Nova 19:240–244 Rosenberg CL, Handy MR (2005) Experimental deformation of partially melted granite revisited: implications for the continental crust. J Metam Geol 23:19–28 Rosenbusch H (1877) Die Steiger Schiefer und ihre Contactzone. Abhandl Geol Spezialkarte Elsass-Lothringen 1:80–393, Halle/ Saale Sander B (1950) Einführung in die Gefügekunde geologischer Körper, vol 2. Teil: Die Korngefüge. Springer, Wien, Austria Sawyer EW, Barnes S-J (1988) Temporal and compositional differences between subsolidus and anatectic migmatite leucosomes from the Quetico metasedimentary belt, Canada. J Metam Geol 6:437– 450 Scharbert HG (1963) Zur Nomenklatur der Gesteine in Granulitfazies. Tschermaks Mineral Petrol Mitt 3(8):591–598 Scheumann KH (1961) “Granulit”, eine petrographische Definition. Neues Jahrb Mineral Monatsh 1961:75–80 Schmieder M, Schwarz WH, Buchner E et al (2012) Double and multiple impact events on Earth—hypotheses, tests, and problems. Meteorit Planet Sci 45:1093–1107 Schmitt RT, Lapke C, Lingemann CM et al (2005) Distribution and origin of impact diamonds in the Ries Crater, Germany. In: Kenkmann T, Hörz F, Deutsch H (eds) Large meteorite impacts III. Geol Soc America Spec Paper, vol 384, pp 299–314 Schneiderhöhn H (1961) Die Erzlagerstätten der Erde, vol II: Die Pegmatite. Gustav Fischer, Stuttgart Searle M, Hacker BR, Bilham R (2001) The Hindu Kush seismic zone as a paradigm for the creation of ultrahigh-pressure diamond- and coesite-bearing continental rocks. J Geol 109:143–153 Sederholm JJ (1907) On granite and gneiss. Bull Comm Géol Finlande 23:1–110 Sederholm JJ (1913) Die Entstehung migmatischer Gesteine. Geol Rundschau 4:174–185 Seidel E, Kreuzer H, Harre W (1982) A Late Oligocene/Early Miocene high pressure belt in the external Hellenides. Geol Jahrb E23:165– 206, Hannover Shirey SB, Richardson SH (2011) Start of the Wilson cycle at 3 Ga shown by diamonds from subcontinental mantle. Science 333:434–436 Spear FS, Cheney IT (1989) A petrogenetic grid for pelitic schists in the system SiO2–Al2O3–FeO–MgO–K2O–H2O. Contrib Mineral Petrol 101:149–164 Storre B, Karotke E (1972) Experimental data on melting reactions of muscovite + quartz in the system K2O–Al2O3–SiO2–H2O to 20 Kb water pressure. Contrib Mineral Petrol 36:343–345 Tilley CE (1925) Metamorphic zones in the southern Highlands of Scotland. Quart J Geol Soc London 81:100–112 Utada M (2001) Zeolites in hydrothermally altered rocks. In: Bish DL, Ming DW (eds) Natural zolites: occurrence, properties, applications. Rev Mineral Geochem 45:305–322 Vallance TG (1977) Spilitic degradation of a tholeiitic basalt. J Petrol 15:79–96 Voll G, Töpel J, Pattison DRM, Seifert F (eds) (1991) Equilibrium and Kinetics of Contact Metamorphism: The Ballachulish Igneous Complex and Its Aureole. Springer, Berlin, Heidelberg, New York, Tokyo von Engelhardt W (1960) Der Porenraum der Sedimente. Springer, Berlin, Göttingen, Heidelberg von Engelhardt W, Arndt J, Stöffler D et al (1967) Diaplektische Gläser in in den Breccien des Ries von Nördlingen als Anzeichen der Schockwellenmetamorphose. Contrib Mineral Petrol 15:93–107 Werner E (1904) Das Ries in der schwäbisch-fränkischen Alb. Blätter Schwäbischer Albverein 16/5, Tübingen, Germany

497 References and Suggestions for Further Reading Further Reading Ashword JR (ed) (1985) Migmatites. Blackie, Glasgow, London Best MG (2003) Igneous and Metamorphic Petrology, 2nd edn. Blackwell, Oxford Brown M, Korhonen FJ, Siddoway CS (2011) Organizing melt flow through the crust. Elements 7:261–266 Bucher K, Frey M (2002) Petrogenesis of Metamorphic Rocks, 7th edn. Springer, Berlin, Heidelberg, New York Carswell DA, Compagnoni R (eds) (2003) Ultrahigh pressure metamorphism. ENU Notes in Mineralogy, vol 5. Eötvös University Press, Budapest, Hungary Clark C, Fitzsimmons ICW, Healy D, Harley SL (2011) How does the continental crust really get hot? Elements 7:235–240 Coleman RG (1977) Ophiolites—Ancient Oceanic Lithospheres? Springer, Berlin, Heidelberg, New York Coleman RG, Wang X (1995) Ultrahigh Pressure Metamorphism. Cambridge University Press, Cambridge Collins GS, Melosh HJ, Osinski GR (2012) The impact cratering process. Elements 8:25–30 Eisbacher GH (1996) Einführung in die Tektonik, 2nd edn. Enke Verlag, Stuttgart Ernst WG (1976) Petrologic Phase Equilibria. Freeman, San Francisco Evans BW (2007) Metamorphic petrology. Landmark Paper Nr 3. Mineral Society Great Britain and Ireland, London Fettes D, Desmons J (eds) (2007) Metamorphic Rocks: A Classification and Glossary of Terms. Cambridge University Press, Cambridge Grapes R (2011) Pyrometamorphism, 2nd edn. Springer, Heidelberg, Dordrecht, London, New York Holness MB, Cesare B, Sawyer EW (2011) Melted rocks under the microscope: microstructures and their interpretation. Elements 7:247– 252 Jamieson RA, Unsworth MJ, Harris NBW et al (2011) Crustal melting and the flow of mountains. Elements 7:253–260 Johannes W, Holtz F (1996) Petrogenesis and Experimental Petrology of Granitic Rocks. Springer, Berlin, Heidelberg, New York Karato S, Wenk H-R (eds) (2002) Plastic deformation of minerals and rocks. Rev Mineral Geochim 51 Kerrick DM (ed) (1991) Contact metamorphism. Rev Mineral 26 Kornprobst J (2002) Metamorphic Rocks and Their Geodynamic Significance. Kluver, Dordrecht, The Netherlands Langenhorst F, Deutsch A (2012) Shock metamorphism of minerals. Elements 8:31–36

Mehnert KR (1968) Migmatites and the Origin of Granitic Rocks, 1st edn. Elsevier, Amsterdam, New York Mehnert KR (1971) Migmatites and the Origin of Granitic Rocks, 2nd edn. Elsevier, Amsterdam, New York Meschede M (1994) Methoden der Strukturgeologie. Enke, Stuttgart Miyashiro A (1994) Metamorphic Petrology. UCL Press, London Ogasawara Y (2005) Microdiamonds in ultrahigh-pressure metamorphic rocks. Elements 1:91–96 Passchier CW, Trouw RAJ (2005) Microtectonics, 2nd edn. Springer, Berlin, Heidelberg, New York Reimold WU, Jourdan F (2012) Impact!—bolides, craters, and catastrophes. Elements 8:19–24 Sawyer EW, Cesare B, Brown M (2011) When the continental crust melts. Elements 7:229–234 Spry A (1983) Metamorphic Textures. Pergamon, Oxford Stöffler D, Grieve RAF (2007) Impactites. In: Fettes D, Desmons J (eds) Metamorphic rocks: a classification and glossary of terms. Chapter 2.11. Cambridge University Press, Cambridge, pp 82–92, 111–125, 126–242 Stöffler D, Artemieva NA, Wünnemann K et al (2013) Ries crater and suevite revisited—observations and modelling. Part I: observations. Meteorit Planet Sci 48:515–589 Teichmüller M (1987) Organic material and very low-grade metamorphism. In: Frey M (ed) Low Temperature Metamorphism, Chapter 4. Blackie, Glasgow, London, pp 114–161 Treloar PJ, O’Brien PJ (eds) (1998) What drives metamorphism and metamorphic reactions? Geological Society, London, Special Publications 138, 287 pp, London Turner FJ (1981) Metamorphic petrology—mineralogical, field, and tectonic aspects, 2nd edn. Hemisphere, Washington, New York, London White RW, Stevens G, Johnson EJ (2011) Is the crucible reproducible? Reconciling melting experiments with thermodynamic calculations. Elements 7:241–246 Winkler HGF (1979) Petrogenesis of Metamorphic Rocks, 5th edn. Springer, Berlin, Heidelberg, New York Winter JD (2001) An Introduction to Igneous and Metamorphic Petrology. Prentice Hall, Upper Saddle River, New Jersey Yardley BWD (1989) An introduction to metamorphic petrology. Longman, Harlow, Essex Yardley BWD, MacKenzie WS, Guilford C (1990) Atlas of Metamorphic Rocks and Their Textures. Longman, Harlow, Essex

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Phase Relations and Mineral Reactions in Metamorphic Rocks 27.1  Mineral Equilibria in Metamorphic Rocks – 500 27.1.1  Assessment of Chemical Equilibrium – 500 27.1.2  Application of the Gibbs’ Phase Rule – 500 27.1.3  Gibbs Free Energy: Stable and Metastable Equilibria – 502

27.2  Metamorphic Mineral Reactions – 504 27.2.1  Polymorphic Transformations and Solid-Solid Reactions – 504 27.2.2  Dehydration Reactions – 508 27.2.3  Decarbonation Reactions – 511 27.2.4  Reactions Involving Both H2O and CO2 – 512 27.2.5  Redox Reactions – 515 27.2.6  Petrogenetic Grids – 517

27.3  Geothermometry and Geobarometry – 517 27.4  Pressure-Temperature Evolution of Metamorphic Complexes – 520 27.4.1  Pressure-Temperature Paths – 520 27.4.2  Pressure-Temperature-Time Paths – 521

27.5  Graphical Presentation of Metamorphic Mineral Assemblages – 522 27.5.1  ACF and A′KF Diagrams – 522 27.5.2  AFM Projections – 525

References (see also Chapters 26 and 28) – 527

© Springer-Verlag GmbH Germany, part of Springer Nature 2020 M. Okrusch, H. E. Frimmel, Mineralogy, Springer Textbooks in Earth Sciences, Geography and Environment, https://doi.org/10.1007/978-3-662-57316-7_27

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In the preceding chapter we have shown that metamorphism causes far-reaching changes in texture, structure and mineral content of a rock. In the course of prograde and retrograde mineral reactions, new mineral assemblages develop, which reflect gradual adjustment to changing P-T conditions. In most cases textural equilibrium is achieved at the time when the maximum temperature is experienced, i.e., at the peak of metamorphism. Consequently, peak metamorphic mineral assemblages can be regarded as approximating thermodynamic equilibrium assemblages. This offers the opportunity to use thermodynamic principles to quantify P-T conditions from observed mineral assemblages and their compositions.

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27.1  Mineral Equilibria in Metamorphic

Rocks

27.1.1  Assessment of Chemical Equilibrium

Metamorphic rocks are products of a complicated evolution, involving a prograde and a retrograde P-T path. Consequently, in many metamorphic rocks manifold disequilibrium textures are observed under the microscope, which can be interpreted as a result of prograde or retrograde reaction steps. The following observations can be taken as evidence of disequilibrium: 5 Zonation of solid solutions, such as of garnet, amphibole, epidote or plagioclase, either reflect the prograde or retrograde P-T path, but can be due as well to element fractionation, e.g., preferred concentration of Mn in garnet. As garnet is optically isotropic, its zonation can be recognised only by means of electron microprobe analysis (EMPA). 5 Mineral relics of igneous or sedimentary protoliths, of previous metamorphic events or of the prograde P-T path can be preserved as metastable phases, in most cases forming mineral inclusions in porphyroblasts. 5 Reaction textures between two or more mineral species, commonly occurring as symplectites or reaction coronas, can testify to the prograde, but more frequently to the retrograde P-T path (. Fig. 27.1). A clear indication of chemical disequilibrium is the occurrence of incompatible mineral pairs, that is minerals that are thermodynamically not stable together. Prime examples are quartz + forsterite, quartz + corundum, or graphite + haematite. While it is easy to proof disequilibrium, the opposite, that is the proof of a given mineral assemblage indeed representing a thermodynamic equilibrium, is not as straightforward. An important criterion in this regard is the development of textural equilibrium, that is a common grain-to-grain contact between coexisting minerals.

. Fig. 27.1  Reaction corona of cordierite (Crd) + orthopyroxene (Opx) around garnet (Grt), documenting the reaction garnet + quartz (Qz) ↔ cordierite + orthopyroxene due to near-isothermal decompression, still at very high temperatures; Granulite, Epupa Complex, Namibia; Plane polarised light (photomicrograph: Sönke Brandt)

A mineral assemblage that fulfils this criterion is referred to as equilibrium assemblage. For instance, the mineral combination staurolite + garnet + biotite + muscovite + plagioclase + quartz in a metapelite can be regarded only as equilibrium assemblage if careful microscopic study confirmed that each of these minerals occurs in mutual contact with all the others. Achieving chemical equilibrium becomes easier with increasing metamorphic grade, especially with higher temperatures as this facilitates diffusion. A high rate of deformation also facilitates the achievement of chemical equilibrium because a higher number of mutual grain contacts provides more pathways for intergranular fluid flow. Of course, the longer the period of time available for metamorphism, the greater is the chance of attaining thermodynamic equilibrium. In this case, the Gibbs’ Phase Rule, already discussed in Sect.  18.1, becomes applicable and can thus form the basis for a quantitative approach to reconstructing metamorphic P-T-X conditions (e.g., Seifert 1978). 27.1.2  Application of the Gibbs’ Phase Rule

Victor Moritz Goldschmidt (1888–1947) was the first to recognise, while studying hornfelses in the Oslo region, Norway, that a close relationship exists between mineral assemblages and chemical composition in a metamorphic rock and that the Gibbs’ Phase Rule can be applied to metamorphic mineral assemblages. This rule stipulates the maximum number of phases that can coexist, at given PT conditions and bulk rock composition; see 7 Sect. 18.1 where the Gibbs’ Phase Rule has been defined as follows:

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27.1 · Mineral Equilibria in Metamorphic Rocks

In a heterogeneous multi-component system, the phases (Ph) are all homogeneous parts of the system that can be distinguished by physical methods. Phases can be different kinds of crystals, a melt or different immiscible melts, a fluid or a gas phase. A (metamorphic) rock consists of one or more crystalline phases, that is minerals. In nearly all cases, these formed in equilibrium with a fluid phase. In addition, a melt phase can be present during the formation of migmatites. A thermodynamic system is described by components (C), which are defined as the lowest number of independent chemical compounds or elements that are necessary to describe the composition of all the phases in the system. The variance or degrees of freedom (F) of a system is defined by the number of variables that can be changed independently without changing the state of the system. The most important of these variables of state are pressure (P), temperature (T) and the partial pressures or fugacities of different volatile components, such as PH2 O , PCO2 , . . . or fH2 O , fCO2 , . . . (7 Sect. 26.1.4).

If only P and T are taken into account as variables of state, the Gibbs’ Phase Rule has the form

F = C − Ph + 2

(27.1a)

The Gibbs’ Phase Rule is valid for a specific system of finite dimension, e.g., the content of a platinum crucible, a thin section, a hand specimen or a granite pluton. Taking a lens of silicate marble included within gneiss as an example, the following systems could be chosen:

sition of the three phases. In the P-T diagram (. Fig. 27.2), each of the three phases have a certain stability field, in which only one of them can occur as single phase. Consequently, it follows from F = 1  −  1 + 2  = 2 that these fields are divariant, which implies that both P and T can be changed independently without changing the state of the system. Along the three equilibrium curves, two of the polymorphs coexist. These curves are univariant as F = 1  −  2 + 2  = 1, and thus either P or T can be changed independently without changing the state of the system. The equilibrium curves (1) kyanite ⇌ andalusite and (2) kyanite ⇌ sillimanite have a positive slope (. Fig. 27.2). Any increase in the equilibrium temperature would necessitate a dependent increase in pressure in order to maintain the state of the system. Otherwise, kyanite would become unstable and the transformation to andalusite or sillimanite, respectively, would proceed into their respective divariant stability fields. The same holds true if P is decreased at constant T. In reverse, the univariant equilibrium curve of reaction (3) andalusite ⇌ sillimanite has a negative slope. Thus with increasing T andalusite becomes unstable at constant P, or with rising P at constant T. The triple point at which all three Al2SiO5 polymorphs coexist in equilibrium is invariant, because F = 1  −  3 + 2  = 0. Consequently, the two variables of state are fixed and not variable at all. The moment one of them is, or both are, changed, the state of the system would change as well. Analogous considerations can be applied to multi-component systems in which, according to the Phase Rule, a higher number of mineral phases can coexist in equilib-

1. marble lens + surrounding gneiss; 2. marble lens alone; 3. part of the marble lens; or 4. an idealised system formed by the assemblage calcite + phlogopite  + forsterite constituting the silicate marble. In an open system, both energy and mass can be exchanged with the surroundings. Addition or removal of chemical components as in metasomatism (7 Sect. 26.6) changes the bulk rock composition of the system of choice, that is of the rock. Thus the marble lens that shows reaction with the surrounding gneiss would represent an open system. In contrast, a closed system can exchange energy but no mass with its surroundings. Consequently, a metamorphic rock s.str., that is a rock in which no change of bulk rock composition has taken place during metamorphism (except for the loss of volatiles), can be regarded as a closed system.

Let us choose the simple one-component system Al2SiO5 as example to which we can apply the Gibbs’ Phase Rule. In this system, three phases are possible. These are the three Al2SiO5 polymorphs kyanite, andalusite and sillimanite. Theoretically, we could choose the three elements Al, Si, and O as components. Alternatively, we could take Al2O3 and SiO2 as components. As all of the possible phases have, however, the same chemical formula Al2SiO5, it is sufficient to take Al2SiO5 as a single component. It is the lowest number of chemical compounds necessary to describe the compo-

. Fig. 27.2  P−T diagram for the one-component system Al2SiO5 with the stability fields of the polymorphs kyanite, andalusite and sillimanite and their different densities (in g/cm3); The equilibrium curves (1) kyanite ↔ andalusite, (2) kyanite ↔ sillimanite and (3) andalusite ↔ sillimanite and the triple point H are from Holdaway and Mukhopadhyay (1993); B alternative triple point after Bohlen et al. (1991)

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Chapter 27 · Phase Relations and Mineral Reactions in Metamorphic Rocks

rium. As most of the metamorphic mineral assemblages are stable over relatively large P-T intervals, divariant equilibria with Ph = C and F = 2 should be commonly observed, a fact that was recognised by V. M. Goldschmidt as “mineralogical phase rule”. For instance, in the three-component system CaO–MgO–SiO2, the following six mineral phases could theoretically occur, depending on bulk rock composition and P-T conditions: enstatite, wollastonite, diopside, forsterite, periclase, and quartz. The Phase Rule dictates that of these six phases only a maximum of three can occur together in equilibrium. These could be En–Wo–Qz, or En–Di–Qz, or Di–Fo–Per, all of which would have two degrees of freedom and thus occupy divariant stability fields in P-T space.

27

In general, an n-component system contains n + 2 divariant P-T fields, in which n phases can coexist in equilibrium. These fields are separated by n + 2 univariant equilibrium curves, at which n + 1 phases coexist forming reaction equilibria. The equilibrium curves meet at an invariant point, at which the number of coexisting phases reaches a maximum of n + 2. The relative spatial arrangement of the equilibrium curves around the invariant point can be constructed using the approach of F. A. H. Schreinemakers, published in a series of papers (1915–1925) and summarised by Zen (1966). The description of this important method is beyond the scope of this textbook and the reader is referred to special texts on theoretical metamorphic petrology (e.g., Yardley 1989; Will 1998).

In reality, numerous exceptions from the “mineralogical phase rule” have been observed, as many metamorphic rocks contain more mineral phases than expected from the number of independent chemical components. For instance, metapelites can contain two or even three Al2SiO5 polymorphs side by side (see 7 Sect. 26.2.5). There are several reasons for the actual or apparent violation of the Phase Rule, as can be explained by . Fig. 27.2. 5 During prograde formation of sillimanite, either kyanite or andalusite can be preserved as metastable relics, thus indicating disequilibrium. 5 The rock actually represents P-T conditions exactly placed onto one of the equilibrium curves (1), (2) or (3), or even on the invariant point in which case the coexistence of two or three Al2SiO5 phases represents a true univariant or invariant equilibrium. Unfortunately, in many cases it is not possible to discriminate between these two alternatives. However, later on we shall discuss metamorphic textures that represent actual univariant equilibria. If the number of selected components C is too low, the Phase Rule must be violated. For instance, the Al2SiO5 polymorphs commonly contain some Fe2O3, preferably incorporated within the andalusite structure. In this case, C would have to be increased by 1, as relatively Fe-rich andalusite can coexist, over a limited P-T range, with Fe-poor sillimanite thus representing a divariant equilibrium with F = (1 + 1) − 2 + 2 = 2. This illustrates that sometimes it can be difficult to decide on the number of components. When several components are present in solid solutions, as in ferromagnesian minerals, in the same proportion(s), they can be

combined into one, e.g., (Mg, Fe)O, otherwise they must be regarded as two different components, e.g., MgO and FeO. In many metamorphic rocks, e.g., in metabasites or metapelites, the number of phases Ph present is smaller than the number of components C. Clearly, this fact does not violate the Phase Rule but merely increases the degree of freedom F. For instance, the fact that only one mineral phase is present in a specific P-T field of a binary system, results in the equation F = 2  −  1 + 2  = 3. Thus the field is trivariant and an additional variable can be considered, e.g., the Fe/ Mg ratio of the minerals involved. In these cases it can be difficult to derive unequivocal evidence of possible disequilibria from the Phase Rule. Up to now, we have assumed that the load pressure Pl is equal to the fluid pressure Pfl. However, this assumption is not always true as has been demonstrated in 7 Sect. 26.1.4. Especially during high-grade metamorphism, e.g., leading to the formation of granulite, conditions of Pfl  Pfl = PCO2, the Phase Rule would take the form of Eq. (27.1b). 27.1.3  Gibbs Free Energy: Stable

and Metastable Equilibria

The state of thermodynamic equilibrium in which a mineral or a mineral assemblage is situated can be quantitatively described by the fundamental state function Gibbs free energy G, defined by the equation

G = H − TS

(27.2a)

where H is the enthalpy (or heat content) and S the entropy. As the deduction of this equation is beyond the scope of this textbook, the reader is referred to relevant textbooks

27.1 · Mineral Equilibria in Metamorphic Rocks

503

and B coexist with each other. This line divides two divariant fields, in which phase A (left) and B (right), respectively, are stable. When intersecting the potential surface of phase C with those of A and B, two more intersecting lines result, forming additional univariant equilibrium curves. Finally, the common intersection of all three surfaces forms an invariant point, at which all three phases coexist in equilibrium and GA = GB = GC. These relations will be explained below in more detail using the one-component system Al2SiO5 (Seifert 1978).

. Fig. 27.3  Surfaces of the chemical potential of two polymorphic phases A and B in the G-P-T space; The phase with the lower Gibbs free energy G is favoured; The projection of the intersection line onto the P-T plane results in the univariant equilibrium curve between the two phases (after Seifert 1978)

(e.g., Will 1998). Each system strives towards minimising its Gibbs free energy. Thus, in a one-component system, the polymorph attaining the lowest value of G, at given state variables, will be thermodynamically stable. In a multi-component system the same holds true for the phase assemblage with the lowest G. For simplicity, let’s limit the discussion to a one-component system with the phases A, B and C, and the state variables P and T. In the G-P-T space, a potential surface is assigned to each of these phases, e.g., A and B in . Fig. 27.3 (Seifert 1978). As these surfaces are continuous, they intersect in a line along which GA = GB thus forming a univariant thermodynamic equilibrium. In the left part of the diagram, phase A is more stable, in the right part phase B. Projecting the intersection line onto the P-T plane produces a univariant equilibrium curve, along which the phases A . Fig. 27.4  G-T sections through the one-component system Al2SiO5 at different pressures; a P-T diagram for four different pressures P1 to P4; b G-T diagram for the pressure P3; solid line: kyanite, thin line: andalusite, blue line: sillimanite; The lowest level of stability is shown as continuous line, the medium level as broken line, and the highest level as dotted line; the progress of possible stable or metastable reactions are indicated by arrows; c–e G-T diagrams for the pressures P1, P2 and P4 (after Seifert 1978)

In . Fig. 27.4b, the T-dependence of G is displayed at a constant pressure of, e.g., P3 (. Fig. 27.4a). The individual levels of stability intersect at the points of stable or metastable equilibrium. At P3 = const. and relatively low temperature, kyanite is the stable phase, whereas andalusite is stable at intermediate T and sillimanite at highest T. The Ostwald’s Step Rule, also known as the law of intermediate stages, states that commonly a profoundly metastable stage cannot directly pass over into a stable one, because initially phases are formed that conform to an intermediate level of stability. Thus, at low temperature, the least stable sillimanite is not directly replaced by kyanite, but metastable andalusite is first formed as a transitional phase. In turn, kyanite is not directly transformed into sillimanite, the stable phase at high temperature, but via the intermediate stage of andalusite crystallisation. At medium temperature, stable andalusite is formed in the sequence sillimanite → kyanite → andalusite or kyanite → sillimanite → andalusite (. Fig. 27.4b). Analogous reaction sequences can be derived for the constant pressures P1, P2 and P4, respectively (. Figs. 27.4c–e, 27.5, 27.6).

In natural rocks, obvious violations of the Phase Rule can be explained by the Ostwald Rule. Thus, in metapelites that contain sillimanite as stable Al2SiO5 polymorph, andalusite is commonly preserved as metastable relic. In many cases, sillimanite was not formed by direct replacement of andalusite, shown in . Fig. 27.7 but crystallised via more complex reactions from micas. In many cases, acicular sillimanite, known as fibrolite, grew epitaxially on flakes of biotite.

27

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Chapter 27 · Phase Relations and Mineral Reactions in Metamorphic Rocks

27.2.1  Polymorphic Transformations

and Solid-Solid Reactions

27 . Fig. 27.5  Schematic diagram displaying the different levels of the Gibbs free energy G for an unstable, metastable and stable state, respectively; Initially, the ball drops from the highest G level (1) into the potential trough of the metastable G level (2). In order to reach the stable G level (4), a potential barrier (3) must be overcome by expending a certain activation energy E*; At synthesis of sillimanite from a highly reactive starting material (1), initially metastable andalusite is formed (2), the transformation of which to produce stable sillimanite (4) requires the overcoming of the potential barrier (3), by expending an activation energy (from Ernst 1976) Another example of the relations between stable and metastable phases in a one-component system are the SiO2 polymorphs (. Fig. 27.6; cf. Sect 11.6). As shown in . Fig. 11.44, at ambient pressure (1 bar) the sequence of stable phases at increasing temperature is α-quartz → β-quartz → β-tridymite → β-cristobalite. These phases define the states of the lowest Gibbs free energy within the respective temperature intervals (. Fig. 27.6). Upon cooling from temperatures >1470 °C, β-cristobalite should invert at this temperature to β-tridymite. This, however, is a reconstructive transformation and thus involves a large activation energy, as symbolised in . Fig. 27.5. Therefore, unless the cooling rate is very slow, β-cristobalite will persist as a metastable phase past the transformation temperature. At about 700 °C a displacive transformation that involves little or no activation energy leads to another metastable phase, α-cristobalite. While β-cristobalite possesses a stability field on its own, α-cristobalite does not. Similar arguments apply to tridymite. Note, however, that the term α-tridymite refers to at least five different metastable low-temperature forms.

Volatile components, such as H2O or CO2, are not involved in these reactions. Thus variation in fluid pressure or in Pl/Pfl ratio has no influence on the stability of the reacting mineral phases. Nevertheless, even these reactions commonly proceed in the presence of a fluid phase that can serve as a transport medium and speed up the reaction process. This holds true especially for supercritical H2O-rich fluids, in which silicate minerals are readily dissolved. Note that in this case, the presence of a fluid phase is not necessary from a thermodynamic point of view but for kinetic reasons, the presence of a fluid phase, which acts as a catalyst, can determine whether a thermodynamically predictable reaction actually takes place. At first, the one-component system Al2SiO5, already described, will be examined in detail. In the P-T diagram (. Fig. 27.2), three univariant equilibrium curves are displayed for the reactions

kyanite ⇋ andalusite

[27.1]

kyanite ⇋ sillimanite

[27.2]

andalusite ⇋ sillimanite

[27.3]

all of which intersect in the invariant triple point, at which all three polymorphs coexist with each other. The divariant stability field of andalusite, the polymorph with the lowest density of 3.15 is restricted to the lowest pressure range. At

27.2  Metamorphic Mineral Reactions

In the last decades, numerous equilibrium curves of critical metamorphic reactions have been determined by a wealth of experiments, thus contributing in a major way to our knowledge of formation conditions of metamorphic rocks. Nevertheless, it has to be conceded that many of these experimentally determined reactions are only simplified versions of the much more complex metamorphic processes that take place in nature.

. Fig. 27.6  Stability ranges of low pressure SiO2 polymorphs indicated by schematically drawn G-T curves; β-cristobalite orange; α-cristobalite red; β-tridymite violet; α-tridymite blue; β-quartz green; α-quartz yellowish green; The lowest level of stability at a given temperature is shown as continuous line, the medium and highest levels as broken lines R = reconstructive, D = displacive transformations (c.f. Griffen 1992)

27.2 · Metamorphic Mineral Reactions

27

505

increasing pressure, andalusite is transformed into the polymorphs of higher density kyanite (3.65) or sillimanite (3.20). Moreover, sillimanite is the stable high-T polymorph in the system. During the polymorphic transformations, a fluid phase, acting as a catalyst, can be present but does not have to be. As only (OH)-free minerals take part in the reactions, the H2O content of the fluid phase has no influence on the position of the univariant equilibrium curves in the P-T diagram. For an estimate of P-T conditions attained during metamorphism of pelitic rocks, the phase diagram for the Al2SiO5 system is of paramount interest (see also . Fig. 27.8). Unfortunately, the experimental investigations performed by different working groups initially provided quite conflicting results that have strongly shaken the confidence of field geologists in the results of experimental petrology. The main reason for these difficulties was the fact that the three phases display very small differences in their Gibbs free energy G. Consequently, it is essential to avoid metastable formation of the “wrong” polymorph during the experiments. Moreover, the stability of sillimanite is affected by Al–Si disorder in its crystal structure, which increases with rising temperature. At present, the phase diagram prepared by Holdaway (1971) and confirmed, with slight modifications, by Holdaway and Mukhopadhyay (1993) is generally accepted. Their triple point was determined at T = 504  ± 20 °C and P = 3.75 ± 0.3 kbar. Within limits of error it conforms to the triple point determined by Bohlen et al. (1991) at T = 530 ± 20 °C and P = 4.2 ± 0.3 kbar. Comparing these results provides a sound indication for the precision that can be generally obtained by experimental determination of metamorphic mineral reactions. Their reliability can be also tested by thermodynamic calculations. As shown in the preceding section, the Gibbs free energy is equal for two phases, respectively, at their univariant equilibrium curve, i.e., ΔG = 0. Thus, for the case of thermodynamic equilibrium, the P, T dependence of ΔG is given by the simple equation

�GP,T = 0 = �H ◦ − T �S + (P − 1)�V ◦

. Fig. 27.7  a and b Hornfels of the Steinach aureole near Vohenstrauss, Bavaria, Germany; a replacement of andalusite by sillimanite (orange interference colours); upper left: cordierite and biotite; width of view c. 4 mm, crossed polars; b assemblage sillimanite (light yellow interference colours) + K-feldspar with perthite lamellae (upper left) + cordierite (e.g. centre and lower right) + biotite (e.g. lower left and right); Width of view c. 3.5 mm, crossed polars; c corundum-bearing gneiss von Morogoro, Tanzania: Assemblage alkali feldspar with typical lamellar exsolution (mesoperthite)—corundum (high relief, yellowish interference colours)—phlogopite (high interference colours); width of view c. 1 mm, crossed polars (photograph M. Okrusch)

(27.2b)

In this equation, ΔH° is the difference in enthalpy (=heat content), ΔS° the difference in entropy and ΔV° the difference in molar volume of this reaction, each at standard state conditions, that is, at room temperature T = 25  °C (=298 K) and atmospheric pressure P = 1 bar. For the minerals taking part in the reaction, H° and S° can be determined by calorimetric measurements and V° from the density or—much more exactly—by X-ray diffraction analysis. The thermodynamic values for the Al2SiO5 minerals are listed in . Table 27.1. First, we calculate the equilibrium temperatures for the three reactions at a pressure of P = 1 bar. Thus the last term

506

Chapter 27 · Phase Relations and Mineral Reactions in Metamorphic Rocks

27

. Fig. 27.8  P−T diagram displaying equilibrium curves for dehydration reactions in metapelites: (7) kaolinite + quartz ↔ pyrophyllite + H2O (Thompson 1970); (8) Pyrophyllite ↔ andalusite/kyanite + quartz + H2O (Hemley 1967; Kerrick 1968); (9) paragonite + quartz ↔ albite + andalusite/sillimanite + H2O (Chatterjee 1972); (10) paragonite ↔ albite + corundum + H2O (Chatterjee 1970); (11) muscovite + quartz ↔ K-feldspar + andalusite/sillimanite + H2O (Chatterjee and Johannes 1974); (11a) muscovite + quartz + H2O ↔ sillimanite/kyanite + liquid (Storre and Karotke 1972); (11b) muscovite + quartz ↔ sillimanite/kyanite + H2O-free liquid (Storre 1972); (12) muscovite ↔ K-feldspar + corundum + H2O (Chatterjee and Johannes 1974); Coloured tangent: inclination calculated for reaction (12) at PH2 O = 2 kbar and T = 950 K; dotted coloured line: inclination calculated for reaction (12) at PH2 O = const = 1 kbar and T = 950 K; The P−T diagram for the Al2SiO5 polymorphs (. Fig. 27.2) is included for comparison . Table 27.1  Thermodynamic parameters for the Al2SiO5 polymorphs at P = 1 bar and T = 298 K (=25 °C) (after Holdaway and Mukhopadhyay 1993) S° (J/mol and K)

Phase

H° (kJ/mol)

V° (cm3)

Kyanite

82.86

−2593.70

44.08

Andalusite

91.60

−2589.66

51.48

Sillimanite

95.08

−2586.37

49.86

From these data, the following changes in entropy, enthalpy and volume between reaction product and reactant can be calculated for the three possible univariant reactions: ∆S° (J/mol and K)

∆H° (J/mol)

Kyanite ↔ sillimanite

12.22

7330

5.78

Kyanite ↔ andalusite

8.74

4040

7.40

Andalusite ↔ sillimanite

3.48

3290

−1.62

in Eq. (27.2b) is omitted and, with ΔG = 0, the equation can be simplified to

T1bar =

H ◦ . S ◦

Thus for the equilibrium curve [27.1] kyanite ⇌ andalusite, we obtain:

T1bar = 4040/8.74 = 462 K = 189 ◦ C

∆V° (cm3)

By analogy, T1bar can be calculated at 327 °C for the metastable extension of the equilibrium curve [27.2] kyanite ⇌ sillimanite. Both values conform, within error, to the experimentally determined intersections of these curves. In contrast, the value for T1bar = 672 °C, calculated for equilibrium curve of Reaction [27.3] andalusite ⇌ sillimanite, appears distinctly too low. This is explained by the curvilinear trend of this curve that is due to the fact that Al–Si disorder in the sillimanite structure increases with rising temperature.

In a next step, we shall calculate the slope of the three equilibrium curves using the Clausius-Clapeyron Equation (7 Sect. 18.2.1) in the form ◦

S dP = ◦ dT V

(27.3)

× V° J bar−1 the value Because V° cm3 is equal to 10  obtained must be multiplied by a factor of 10 thus resulting in the calculated slopes for the three equilibrium curves kyanite ⇋ sillimanite dP = 10 × 12.22/5.78 = 21.14 bar/K ≈ 2 kbar/100 ◦ C dT kyanite ⇋ andalusite dP = 10 × 8.74/7.40 = 11.81 bar/K ≈ 1.2 kbar/100 ◦ C dT andalusite ⇋ sillimanite dP = 10 × 3.48/−1.62 = −21.48 bar/K ≈ −2.1 kbar/100 ◦ C dT

Looking at . Fig. 27.2, it is easily tested to what extent the calculated slopes conform to the experimentally determined ones! For the calculation of the triple point, Eq. [27.2] is converted into: ◦





T = �H /�S + (P−1)�V /�S T = T1bar + (P−1)



1 dP/dT

(27.2a) (27.2b)

Inserting the relevant values for the three equilibrium curves in this equation, we obtain:

1 21.14 1 = 945 K + (P−1) × −21.48

TKy/Sil = 600 K + (P−1) × TAnd/Sil

27

507

27.2 · Metamorphic Mineral Reactions

As all three equilibrium curves meet at the triple point, these two equations can be equalised and resolved into P, thus obtaining

PTrip = 345/0.094 + 1 = 3671.2 bar ≈ 3.7 kbar When inserting this value into the equations for TKy/Sil, TAnd/Sil and TKy/And, coinciding results are obtained for the temperature of the triple point TTrip = 774  K = 501  °C that well conforms to the value experimentally determined by Holdaway and Mukhopadhyay (1993). In nature, numerous metamorphic complexes have been recognised, in which two or even three Al2SiO5 polymorphs

occur in mutual contact. The question whether these really form univariant or invariant equilibrium assemblages or are preserved as metastable relics can be answered only by careful microscopic investigations. For instance, in the Steinach aureole, Bavaria, clear textural evidence was found for replacement of andalusite by sillimanite (. Fig. 27.8a) where universal stage measurements revealed the crystallographic orientation relationships cAnd||cSil, bAnd||aSil and aAnd||bSil (Okrusch 1969). It should be noted, however, that the stability limits of kyanite, andalusite and sillimanite are slightly shifted if Al is replaced by Fe3+. In this case, two Al2SiO5 polymorphs can coexist within a limited P-T interval, e.g., Fe-richer andalusite with Fe-poorer sillimanite. Thus the univariant equilibrium curves are broadened to form divariant bands, and the triple point is extended to form a small P-T field. However, in the petrological practice, these modifications can be neglected as they are smaller than the error limits of the experimental determinations. At temperatures above 1000 °C and relatively low pressures, sillimanite breaks down to form mullite, ~Al[6]Al1.2[4] [O/Si0.8O3.9] = 5.5Al2O3·4SiO2, plus free SiO2, and the chemical compound Al2SiO5 does no longer exist. Thus mullite is the only Al-silicate than can coexist with a H2O-free melt.

A further solid-solid reaction, i.e., without H2O or CO2 involved, is the model reaction

NaAl[Si2 O6 ] + SiO2 ⇋ Na[AlSi3 O8 ] jadeite

quartz

albite

[27.4]

The equilibrium curve (. Fig. 26.1) of this reaction delineates the upper temperature stability or the lower pressure stability of jadeite in the presence of quartz: Consequently, jadeite-bearing rocks are formed during high-pressure metamorphism, i.e., at high P/T ratios (Sects. 28.2.8, 28.2.9). However, as natural jadeite commonly contains certain amounts of the diopside and acmite components, it can already form at somewhat lower pressures. The same holds true for the eclogite mineral omphacite, a solid solution of augite + jadeite (+acmite). Over the past decades more and more examples of the high-pressure modifications of SiO2, coesite, and even of carbon, diamond, have been discovered in crustal rocks that had undergone ultra-high pressure metamorphism. The reactions

quartz ⇋ coesite (Fig. 11.44)

[27.5]

graphite ⇋ diamond (Fig. 4.15)

[27.6]

are important indicators for these extreme P-T conditions. In the following, we shall describe metamorphic reactions, in which the volatile components H2O and CO2 participate. During prograde metamorphism, these volatiles are commonly released due to dehydration or decarbonisation reactions (Sects. 27.2.2, 27.2.3) but, in some cases, can be consumed as well (7 Sect. 27.2.4). During retrograde metamorphism, however, hydration reactions are common. In redox reactions (7 Sect. 27.2.5), O2 and H2 are important components in addition to other volatiles such as H2O and CO2.

508

Chapter 27 · Phase Relations and Mineral Reactions in Metamorphic Rocks

27.2.2  Dehydration Reactions z Dehydration Reactions at PH2 O = Ptot

27

An increase in temperature typically leads to progressive dehydration reactions in rocks that initially contain H2Oand/or (OH)-bearing minerals. The majority of metamorphic reactions, some of which have been already presented in 7 Chap. 26 (. Fig. 26.11), belong to this group and are thus of high relevance for estimating metamorphic P-T conditions (. Fig. 27.8). In experiments performed in conventional high-pressure autoclaves, H2O commonly is present as an excess phase and thus is a medium of pressure-transfer. Consequently, the partial pressure of H2O (water vapour pressure) is equal to the total pressure: PH2 O = Pfl = Ptot. Experimental arrangements of this type are known as hydrothermal experiments and their results, when having achieved thermodynamic equilibrium, are displayed in PH2 O −T diagrams. During prograde metamorphism of pelitic sedimentary rocks, e.g., the following dehydration reactions can take place with increasing temperature:

Al4 [Si4 O10 ](OH)8 + 4SiO2 quartz

kaolinite

[27.7]

⇋ 2Al2 [Si4 O10 ](OH)2 + 2H2 O pyrophyllite

Al2 [Si4 O10 ](OH)2 ⇋ pyrophyllite

Al2 SiO5 andalusite/kyanite

[27.8]

+ 3SiO2 +H2 O quartz

NaAl2 [AlSi3 O10 ](OH)2 + SiO2

quartz

paragonite

⇋ Na[AlSi3 O8 ] + albite

[27.9]

Al2 SiO5 andalusite/sillimanite

+H2 O

NaAl2 [AlSi3 O10 ](OH)2 ⇋ Na[AlSi3 O8 ] albite

paragonite

[27.10]

+ Al2 O3 +H2 O corundum

KAl2 [AlSi3 O10 ](OH)2 + SiO2 muscovite

⇋ K[AlSi3 O8 ] + K-feldspar

quartz

Al2 SiO5 andalusite/sillimanite

[27.11]

+H2 O

KAl2 [AlSi3 O10 ](OH)2 ⇋ K[AlSi3 O8 ] muscovite

K-feldspar

[27.12]

+ Al2 O3 +H2 O corundum

At the beginning of metamorphism, the clay mineral kaolinite reacts in the absence of quartz and under H2O release to form pyrophyllite (. Fig. 27.8, curve (7)). If, however, K+ ions are present in the system, muscovite is produced instead. In the past, pyrophyllite was overlooked in many metapelites

but has been proved as a rather common mineral in slate and phyllite. At higher temperatures, pyrophyllite is transformed into andalusite or kyanite plus quartz, whereby H2O is released again (. Fig. 27.8, curve (8)). The equilibrium curves for the breakdown of paragonite + quartz and of paragonite alone (. Fig. 27.8, curves (9), (10)) as well as of muscovite + quartz and of muscovite alone (. Fig. 27.8, curves (11), (12)) proceed at significantly higher temperatures. The assemblages sillimanite + K-feldspar (. Fig. 27.7b) and corundum + K-feldspar (. Fig. 27.7c) already document PT conditions of high-grade metamorphism. Towards higher H2O-pressures, the dehydration curves (9)–(12) end up at invariant points, at which H2O-saturated solidus curves, e.g., muscovite + quartz + H2 O ⇋ sillimanite/kyanite + melt

[27.11a]

or curves of dehydration melting, e.g., muscovite + quartz ⇋ sillimanite/kyanite + melt [27.11b]

branch off (. Fig. 27.8). Over typical crustal P-intervals, dehydration curves display a positive slope. This is due to the fact that the sums of the molar volumes and of the entropies of the reaction products are distinctly higher than those of the corresponding educts, because the molar volume and the entropy of the H2O released are much higher than those of the solid phases. Consequently, ΔV and ΔS become positive which, according to the Clausius-Clapeyron Eq. (27.3), results in a positive slope of the equilibrium curve. Thus for dehydration reactions, the Clausius-Clapeyron Equation has the following simplified form:

�S T /1bar solids dP = 10 × + V T ,P H2O ◦ dT �V solids

(27.4)

In this equation �S T /1bar solids is the difference in entropy of the solid phases at a given temperature and P = 1  bar, ΔV°solids the difference in molar volumes of the solid phases at T = 298  K (=25 °C) and P = 1  bar, S T ,P H2 O and V T ,P H2 O the entropy and molar volume of H2O at a given P-T combination. In . Table 27.2, the thermodynamic data are presented for the participants of Reaction [27.12]. Using these values, the slope of the reaction curve can be calculated from Eq. [27.4], e.g., at a point selected at PH2 O = 2 kbar and T = 680  °C:

−58.53 + 154.04 95.51 dP = 10 × = 10 × dT −6.085 + 33.091 27.006 = 35.4 bar/K ≈ 3.5 kbar/100 ◦ C Moreover, . Fig. 27.8 reveals that the dehydration curves display a curved shape that show a flat slope at low pressure but become steeper as pressure increases. This feature can be explained by the fact that the compressibility of H2O

27

509

27.2 · Metamorphic Mineral Reactions

. Table 27.2  Thermodynamic data for the reaction muscovite ↔ K-feldspar + corundum + H2O Phase

Molar volumes (cm3 mol−1)

Entropies (J mol−1 K−1)

Corundum

+V° = 25.575

+S950 K/1 bar = 173.80

K-feldspar

+V° = 109.05

+S950 K/1 bar = 547.15

Muscovite

−V° = 140.71

−S950 K/1 bar = 779.48

Solid phases

∆V°solids = −6.085

Ssolids 950 K/1bar = −58.53

H 2O

+V950 K/2 kbar = 33.091

Ssolids 950 K/2 kbar = 154.04

vapour decreases with rising pressure: At constant temperature, the molar volume of H2O vapour dramatically decreases with increasing pressure whereas, at higher pressures, this is hardly the case (. Table 27.3). Thus, above ca. 3 kbar, the dehydration curves become steeper and steeper, i.e., the reactions are less and less pressure-dependent. This fact is confirmed, e.g., by metamorphic dehydration reactions affecting ultramafic mineral assemblages in the model system MgO–SiO2–H2O (. Fig. 27.9). During prograde metamorphism of serpentinite, the serpentine mineral antigorite breaks down according to the reaction 5Mg6 [Si4 O10 ](OH)8 ⇋ 12Mg2 [SiO4 ] antigorite

forsterite

. Fig. 27.9  P−T diagram showing the equilibrium curves for dehydration reactions in metamorphosed ultramafic rocks in the model system MgO−SiO2−H2O: (13) Antigorite ↔ forsterite + talc + H2O (Evans et al. 1976); (14) talc + forsterite ↔ anthophyllite + H2O; (15) anthophyllite + forsterite ↔ enstatite + H2O; (16) talc ↔ anthophyllite + quartz + H2O and (17) anthophyllite ↔ enstatite + quartz + H2O (Chernosky et al. 1985)

[27.13]

+ 2Mg3 [Si4 O10 ](OH)2 +18H2 O talc

as soon as a minimum temperature of ca. 500 °C is overstepped at relatively low pressure, e.g., at the contact with an intrusive body. If, in reverse, the temperature falls below 500 °C, forsterite or olivine can react with talc and H2O to produce new serpentine minerals. If talc was initially present in excess during this back reaction, it can form a stable assemblage with serpentine, as its broad stability field extends into the P-T range left of equilibrium curve [27.13]. At increasing temperature, the following reactions take place in the ultramafic system (. Fig. 27.9):

9Mg3 [Si4 O10 ](OH)2 + 4Mg2 [SiO4 ] forsterite

talc

⇋ 5Mg7 [Si8 O22 ](OH)2 +4H2 O

[27.14]

anthophyllite

2Mg7 [Si8 O22 ](OH)2 + 2Mg2 [SiO4 ] forsterite

anthophyllite

⇋ 9Mg2 [Si2 O6 ] +2H2 O

[27.15]

enstatite

7Mg3 [Si4 O10 ](OH)2 ⇋ 3Mg7 [Si8 O22 ](OH)2 . Table 27.3  Molar volumes of water vapour at different pressures and temperatures (Kennedy and Holser 1966) P (bar)

T (°C) 300

500

750

1

47534.0

64236.0

58048.0

10

4646.0

6383.0

8481.0

100

25.2

591.1

829.8

1000

21.9

34.1

71.8

2500

19.7

24.6

33.5

5000

17.9

20.5

24.3

10,000

16.3

17.6

16.3

20,000

14.5

15.3

15.1

anthophyllite

talc

[27.16]

+ 4SiO2 +4H2 O quartz

2Mg7 [Si8 O22 ](OH)2 anthophyllite

⇋ 7Mg2 [Si2 O6 ] + 2SiO2 +2H2 O enstatite

[27.17]

quartz

The reactions [27.15] and [27.17] are good examples that illustrate why, at higher pressures, the equilibrium curves of dehydration reactions can attain a negative slope, because V T ,P solids becomes increasingly negative, which cannot be compensated by the positive V T ,P H2 O. In some cases the decrease in the total volume of the solid phases is so large, i.e., V T ,P solids becomes so negative, that the equilibrium

510

Chapter 27 · Phase Relations and Mineral Reactions in Metamorphic Rocks

27

. Fig. 27.10  P−T diagram showing the equilibrium curves for dehydration reactions of zeolites: (18) Analcime + quartz ↔ albite + H2O (Thompson 1971); (20) heulandite ↔ laumontite + quartz + H2O (Cho et al. 1987); (19) laumontite ↔ lawsonite + quartz + H2O, and (21) laumontite ↔ wairakite + H2O (Liou 1971); (22) wairakite ↔ anorthite + quartz + H2O (Liou 1970); upper stability limit of lawsonite (Liou 1971)

curve attains a negative slope even at low pressures. An example is the reaction (. Fig. 27.10):

laumontite

⇋ Ca[Al2 Si4 O12 ] · 2H2 O +2H2 O

Na[AlSi2 O6 ] · H2 O + SiO2

[27.21]

wairakite

quartz

analcime

Ca[Al2 Si4 O12 ] · 4H2 O

[27.18]

⇋ Na[AlSi3 O8 ] +H2 O albite

Ca[Al2 Si4 O12 ] · 2H2 O wairakite

In contrast, some zeolite minerals undergo dehydration reactions, at which the total modal volume as well as the total entropy of the solid phases is very much reduced, leading to −ΔS/−ΔV. Consequently, a positive slope of the equilibrium curve dP/dT results. An example is the hardly T-dependent upper pressure stability of the zeolite mineral laumontite according to the reaction (. Fig. 27.10):

⇋ Ca[Al2 Si2 O8 ] + 2SiO2 +2H2 O anorthite

[27.22]

quartz

The reactions [27.18] to [27.22] are of great significance for an estimate of the P-T conditions attained during lowest-grade regional metamorphism, ocean-floor as well as hydrothermal metamorphism. z Dehydration Reactions at P H2 O < P tot

Ca[Al2 Si4 O12 ] · 4H2 O laumontite

⇋ CaAl2 [Si2 O7 ](OH)2 · H2 O + 2SiO2 +2H2 O [27.19] lawsonite

quartz

Similar to Reaction [27.18], H2O is released at increasing H2O pressure. In spite of their negative ΔS and ΔV values, the equilibrium curves of the prograde dehydration reactions, causing the breakdown of heulandite, laumontite and wairakite, attain the common shape (. Fig. 27.10):

∼ Ca4.5 [Al9 Si27 O72 ] · 24H2 O heulandite

⇋ 4.5Ca[Al2 Si4 O12 ] · 4H2 O + 9SiO2 +6H2 O laumontite

quartz

[27.20]

So far, we considered dehydration reactions, at which the H2O pressure was equal to the total pressure. However this is not always the case in nature. Especially at high metamorphic grade, the condition PH2 O < Ptot is frequently realised. Two different cases are distinguished: 1. P tot >P fl =P H2 O. The H2O pressure is equal to the fluid pressure which, however, is smaller than the total pressure. At these conditions, the Clausius-Clapeyron Equation attains the form of a partial differential (Greenwood 1961):



∂Ptot ∂T



= 10 × PH2 O

�S P/T ◦ �V solids

(27.5)

27

511

27.2 · Metamorphic Mineral Reactions

Extending this expression with ΔVP,T, we obtain: 

∂Ptot ∂T



�V P,T �SP/T × ◦ P,T �V �V solids

= 10 × PH2 O

dP �V P,T = 10 × × ◦ dT �V solids

(27.5a)

Taking the reaction

KAl2 [AlSi3 O10 ](OH)2 ⇋ K[AlSi3 O8 ] K-feldspar

muscovite

[27.12]

+ Al2 O3 +H2 O corundum

as an example and insert the thermodynamic data, listed in . Table 27.2 for T = 950  K ≈ 680  °C and Ptot = 2  kbar, we obtain for PH2 O = 1 kbar:



∂Ptot ∂T



35.4 × 27.006 = −157 bar/K −6.085 PH2 O ◦

= −15.7 kbar/100 C Thus a new equilibrium curve results with a steep negative slope (. Fig. 27.8). It becomes obvious that, at increasing Ptot, the equilibrium temperature of the muscovite breakdown more and more deviates from that at Ptot = PH2 O. 2. P tot =P fl =P H2 O +P CO2 +P CH4 . . .. The fluid pressure is equal to the total pressure as well but the fluid phase consists of several gas species with their relevant partial pressures (or fugacities). In this case, commonly realised during prograde metamorphism of graphite-bearing pelitic or marly sedimentary rocks, the equilibrium curves retain their positive slopes but are shifted towards lower temperatures. For instance, the equilibrium curve for the reaction

KAl2 [AlSi3 O10 ](OH)2 + SiO2 ⇋ K[AlSi3 O8 ] muscovite

+

quartz

Al2 SiO5 +H2 O andalusite/sillimanite

K-feldspar

[27.11]

was experimentally determined by Kerrick (1972) for different molar fractions of XH2 O = H2O/(H2O + CO2). As shown in . Fig. 27.11, the equilibrium temperature decreases by ca. 50 °C if, at a total fluid pressure of 2 kbar, XH2 O is reduced from 1 to 0.5. At higher Pfl, the temperature decrease is even more distinct. In turn, the temperature of the granite solidus increases with reduction of XH2 O, a fact already demonstrated in . Fig. 20.5. The solidus curves of granite and the equilibrium curves of Reaction [27.11], respectively, intersect at invariant points from which the steep equilibrium curves of the following reactions branch off:

. Fig. 27.11  P-T diagram showing the stability of muscovite in the presence of quartz (±plagioclase) after reaction [27.11] and [27.11a] and position of the granite solidus at Ptot = Pfl = PH2 O + PCO2 and XH2 O = H2 O/(H2 O + CO2 ) values of 1.0, 0.7 and 0.5 (modified after Kerrick 1972)

muscovite + quartz + albite + H2 O ⇋ sillimanite/kyanite + liquid

[27.11a]

and muscovite + quartz + plagioclase ⇋ sillimanite/kyanite + K-feldspar + liquid

[27.11b]

These respective reactions describe either the H2O-saturated melting or the dehydration melting of muscovite in the presence of quartz + plagioclase for different XH2 O values. (Reaction [27.11b] is not shown in . Fig. 27.11, cf. . Fig. 27.8). Only at XH2 O < 0.5, the conventional breakdown of muscovite + quartz according to Reaction [27.11] can lead to the formation of kyanite instead of andalusite or sillimanite. Consequently, kyanite-bearing granulite should be generated at relatively “dry” conditions or, alternatively, migmatite structures indicate partial dehydration melting. From these considerations it must be concluded that equilibrium curves of dehydration reactions can be used for safe T estimates only if independent information is available on the total pressure and the H2O content of the fluid phase. 27.2.3  Decarbonation Reactions

During prograde metamorphism of impure SiO2– and/or Al2O3-bearing carbonate rocks, CO2 is released alone or

512

Chapter 27 · Phase Relations and Mineral Reactions in Metamorphic Rocks

In contrast, ΔV°solids is clearly negative because calcite and quartz have much lower densities than wollastonite:

�V



solids

=V



Wo −



V



Cal 3

+V



Qz



= 39.260

−(36.934 + 22.688) cm /mole = −20.362 cm3 /mole Consequently, the equilibrium curve attains a negative slope at higher CO2 pressures. The same holds true for other decarbonation reactions. In contrast to calcite, dolomite and magnesite react with quartz already at lower temperatures. The negative slope of the equilibrium curve for PCO2 = 1 bar in . Fig. 27.11 results from the corresponding application of the Eq. [27.5a].

27

. Fig. 27.12  Equilibrium curves of the reaction [27.23] calcite + quartz ↔ wollastonite + CO2 for XCO2 values of 1.0, 0.75, 0.5, 0.25 and 0.13 at Ptot = Pfl = PCO2 + PH2 O as well as for PCO2 = 1 bar, based on experimental results of Harker and Tuttle (1956) and Greenwood (1967) from Winkler (1979); Phase relations in the three-component system CaO–SiO2–CO2 are demonstrated by tie lines for two different compositions A and B

together with H2O. A familiar good example of a decarbonation reaction is the breakdown of calcite in the presence of quartz to form wollastonite according to the reaction:

CaCO3 + SiO2 ⇋ CaSiO3 +CO2 calcite

quartz

wollastonite

[27.23]

As demonstrated in . Fig. 27.12, the equilibrium curve of this reaction at Ptot = PCO2 has a shape similar to most dehydration reactions, i.e., a positive slope and a distinct curvature in the low-pressure range. The pressure influence on the equilibrium temperature is significant: At PCO2 = 0.5  kbar, T is ca. 550 °C but as high as ca. 780 °C at 3 kbar. Thus it is not surprising that wollastonite forms more commonly in contact aureoles rather than during regional metamorphism, at which the assemblage calcite + quartz can be stable to temperatures of >700 °C. However, analogous to dehydration reactions, the equilibrium curves of decarbonation reactions are shifted toward lower temperatures if the fluid phase contains other volatile species such as H2O in addition to CO2, i.e., Ptot = Pfl = PCO2 + PH2 O . . .. In . Fig. 27.12, the equilibrium curves for Reaction [27.23] are displayed for different CO2/(CO2 + H2O) ratios i.e., for XCO2 of 0.75, 0.50, 0.25 and 0.13 as well as for 1  bar. If, however, P tot > P fl = P CO2, a PCO2 = const. =  steep equilibrium curve with negative slope branches off the curve for XCO2 = 1, e.g., at PCO2 = 1 kbar. By analogy to dehydration reactions, a reaction such as [27.23] can only be used for deriving a T estimate if independent evidence is available for the total pressure as well as the partial pressures of CO2 and H2O. The total entropy increase of Reaction [27.23] has the same order of magnitude as for dehydration reactions.

Using Reaction [27.23] as an example, we shall apply the Gibbs’ Phase Rule to systems to which the fluid phase consists only of one volatile component such as CO2, with Ptot = Pfl = PCO2. In the three-component system CaO– SiO2–CO2 (C = 3), the four phases calcite, quartz, wollastonite and fluid (Ph = 4), coexist along the univariant equilibrium curve of Reaction  [27.23]: F = C − Ph + 2 = 3 − 4 + 2 = 1. (Not regarded is the rare mineral lime, CaO, rarely formed during pyrometamorphism.) Along the equilibrium curve, either T or PCO2 can be varied independently without disturbing the state of the system. On the other hand, a maximum number of three phases can exist in each of the divariant fields: F = 3 − 3 + 2 = 2, in which both T and PCO2 can be independently changed. In . Fig. 27.12, the phase combinations possible in the concentration triangle CaO–SiO2–CO2 are documented, for two different bulk compositions A and B, by the relevant arrangement of tie lines. At the left side of the equilibrium curve, calcite + quartz coexist in both rocks. Due to Reaction [27.23], the tie line calcite–quartz is replaced by the tie line wollastonite–fluid. Consequently, the assemblage quartz + wollastonite (+fluid) becomes stable in rock A and wollastonite + calcite (+fluid) in rock B. If H2O is present as additional volatile component and Ptot = Pfl = PCO2 + PH2 O, the system gains an additional degree of freedom and the Phase Rule takes the form F = C − Ph + 3. With F = 3  −  4 + 3  = 2, the univariant equilibrium curve of Reaction [27.23] is transformed into a divariant equilibrium surface dividing the Pfl −T −XCO2 space (. Fig. 27.13a). Reactions in which H2O and CO2 are involved as volatile components are usually displayed in isobaric T −XCO2 diagrams, i.e., as sections at constant Ptot = Pfl (. Fig. 27.13b) because of all three parameters, P, T and XCO2, P typically has the least influence on the overall topology of the reactions. This type of diagram will be presented and discussed in the following. 27.2.4  Reactions Involving Both H2O and CO2

Reactions of this kind play an important role during prograde metamorphism of impure limestones, especially of marls and siliceous carbonate rocks. The relevant equilibrium curves are presented in T −XCO2 diagrams that form isobaric sections through the T −Pfl −XCO2 volume, thus

27

513

27.2 · Metamorphic Mineral Reactions

. Fig. 27.13  a Equilibrium surface (shaded) of a decarbonation reaction dividing the T −Pfl −XCO2 space; an isobaric section (Pfl = const) generates a univariant equilibrium curve in the T −XCO2diagram. b Schematic T −XCO2 diagram for a fluid phase consisting of H2O + CO2; Five different types of reactions lead to equilibrium curves of characteristic shape (Greenwood 1967); B and D are solid phases of defined composition (modified from Miyashiro 1994)

presenting lines of intersection of the divariant equilibrium surface at a given total Pfl (. Fig. 27.13a). As the total fluid pressure is kept constant, the number of degrees of freedom is reduced by 1, thus leading to F = C − Ph + 2. Consequently, the equilibrium curves in the T −XCO2 section become univariant again. According to Greenwood (1967), the slope of these equilibrium curves can be calculated from a partial differential that is analogous to the Clausius-Clapeyron Equation. As a simplifying prerequisite, we assume that the fugacity coefficients γi of H2O and CO2 are constant and thus the molar fractions XH2 O = H2 O/(H2 O + CO2 ) and XCO2 = CO2/(H2O + CO2) = 1 - XH2 O are an approximation of the respective fugacities:



∂T ∂XCO2



= Pfl,γ

RT × �S



vCO2 vH2 O − XCO2 XH2 O



(27.6)

In this equation, νCO2 and νH2 O are the numbers of moles of CO2 or H2O, respectively participating in the reaction, and R is the ideal gas constant. We recall that, in the calculations, the reactants are taken as negative, the products as positive. According to Greenwood (1967), five different cases can be distinguished. These are schematically presented in the (. Fig. 27.13b)

T −XCO2 diagram for Pfl = const. 1. Reactions, In Which Only CO2 Is Produced: ν H2 O = 0, ν CO2 > 0 Assuming the common case that ΔS is positive, (∂T /∂XCO2 )Pfl,γ becomes positive, i.e., the equilibrium curve has a positive slope in the T −XCO2 diagram and dissects the T-axis at XCO2 = 1 at the maximum possible temperature. With decreasing temperature, the curve approaches the T-axis asymptotically at XCO2 = 0 (. Fig. 27.13b, curve 1, 27.14). Examples: Reaction [27.23], and analogous decarbonation reactions, such as

CaMg(CO3 )2 ⇋ MgO + CaCO3 +CO2 dolomite

periclase

calcite

[27.24]

2. Reactions, In Which Only H2O Is Produced: ν H2 O ≥ 0, ν CO2 = 0 At positive ΔS, (∂T /∂XCO2 )Pfl,γ and, consequently, the slope of the equilibrium curve become negative. Examples are all pure dehydration reactions (. Fig. 27.13b, curve 2). 3. Reactions, In Which Both CO2 and H2O Are Produced: ν H2 O ≥ 0, ν CO2 ≥ 0 The equilibrium curve attains a maximum, in fact at the point at which

XCO2 = vCO2 /(vCO2 + vH2 O ) because, at this point, the term (∂T /∂XCO2 )Pfl,γ and, consequently, the slope of the equilibrium curve becomes zero. This is easily demonstrated by inserting the term for XCO2 into Eq. (27.6), followed by minor transformation. Soon it becomes obvious that the curve attains a positive slope for low XCO2, and a negative one for high XCO2 values. With decreasing temperature, the equilibrium curve approaches the T-axis asymptotically at XCO2 = 0 as well as at XCO2 = 1 (. Fig. 27.13b, curve 3). Some very important reactions in the system CaO–MgO–SiO2–CO2–H2O may serve as examples:

Mg3 [Si4 O10 ](OH)2 + 5CaMg(CO3 )2 dolomite

talc

⇋ 4Mg2 [SiO4 ] + 5CaCO3 +5CO2 + H2 O forsterite

[27.25]

calcite

Ca2 Mg5 [Si8 O22 ](OH)2 + 3CaCO3 + 2SiO2 tremolite

calcite

⇋ 5CaMg[Si2 O6 ] +3CO2 + H2 O diopside

quartz

[27.26]

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Chapter 27 · Phase Relations and Mineral Reactions in Metamorphic Rocks

5Mg3 [Si4 O10 ](OH)2 + 6CaCO3 + 4SiO2 calcite

talc

quartz

[27.27]

⇋ Ca2 Mg5 [Si8 O22 ](OH)2 +6CO2 + 2H2 O tremolite

In Reaction [27.25] the maximum is reached at

XCO2 = 5/(5 + 1) = 0.83, and at

XCO2 = 3/(3 + 1) or 6/(6 + 2) = 0.75

27

for Reactions [27.26] and [27.27], respectively. 4. Reactions, In Which H2O Is Consumed and CO2 Is Produced: ν H2 O ≤ 0, ν CO2 ≥ 0 As the two volatile components are placed on opposite sides of the reaction, the univariant equilibrium curves for reactions of this type display a point of inflection rather than a maximum. By analogy to type 1, they have a positive slope in the T −XCO2 diagram and approach the T-axis asymptotically on the high-T side at XCO2 = 1 as well as the low-T side at XCO2 = 0 (. Fig. 27.13b, curve 4, . Fig. 27.14, curve 28). The point of inflection is situated, where the 2nd derivative of

(∂T /∂XCO2 )Pfl,γ after ∂XCO2 becomes zero. The formation of talc according to the reaction 3CaMg(CO3 )2 + 4SiO2 +H2 O quartz

dolomite

⇋ Mg3 [Si4 O10 ](OH)2 + 3CaCO3 +3CO2 talc

calcite

[27.28]

. Table 27.4  Variation of (nCO2 /XCO2 ) − (nH2 O /XH2 O ) with XCO2 for reaction (27.28) XCO2

(nCO2 /XCO2 ) − (nH2 O /XH2 O )

0



0.1

31

0.2

16.3

0.3

11.4

0.4

9.2

0.5

8.0

0.6

7.5

0.7

7.6

0.8

8.8

0.9

13.3

1.0



may serve as example for this rather common type of reaction. The variation of

vCO2 /XCO2 −vH2 O /XH2 O = 3/XCO2 −1/XH2 O with XCO2 is easily calculated. The results listed in . Table 27.4 indicate that the slope of the equilibrium curve is very steep at low and high XCO2 values but flattens out towards the point of inflection that is, in this case, close to XCO2 ≈ 0.6. This fact is well demonstrated by the equilibrium curve 4 in the schematic diagram (. Fig. 27.13b) but is less visible for Reaction (27.28) in . Fig. 27.14. 5. Reactions, In Which CO2 Is Consumed and H2O Produced: ν H2 O ≥ 0, ν CO2 ≤ 0 In this type of reaction, the two volatile components are placed on opposite sides of the reaction as well. Thus the

. Fig. 27.14  T −XCO2 diagram for Pfl = 1 kbar for various reactions in the calc-silicate system CaO−MgO−SiO2−CO2−H2O; (23) dolomite + quartz + H2O ↔ talc + calcite + CO2; (24) talc + calcite ↔ tremolite + dolomite + CO2 + H2O; (25) tremolite + dolomite ↔ forsterite + calcite + CO2 + H2O; (26) diopside + deolomite ↔ forsterite + calcite + CO2 (after different authors from Miyashiro 1973, modified)

27

515

27.2 · Metamorphic Mineral Reactions

univariant equilibrium curves are characterised by a point of inflection again but, by analogy to type 2, display a negative slope. In nature, however, reactions of this type are rare. 27.2.5  Redox Reactions

In the Earth’s atmosphere, oxygen occurs predominantly in the form of the free molecule O2, its partial pressure being PO2 ≈ 0.21 bar, conforming to its volumetric proportion of 20.8%. In the Earth’s crust and mantle, however, oxygen is commonly a constituent of chemical compounds, predominantly forming silicate minerals, to a lesser extent oxides, carbonates and others. In addition to oxygen-bearing volatiles, such as H2O and CO2, the fluid phase can contain free O2 and H2, both of which were partly released by dissociation of H2O according to the redox reaction

H2 O ⇋ H2 + 1/2O2

[27.29]

At higher P-T conditions of metamorphic or igneous processes, the more appropriate term oxygen fugacity fO2 should be used instead of the partial pressure PO2. As displayed in the T −fO2 diagram (. Fig. 27.15), there is a considerable pressure influence on Reaction [27.29]: At a temperature of 600 °C and atmospheric pressure of P = 1  bar, H2O dissociates already at fO2 = 10−8 bar whereas, at 600 °C and PH2 O = 2 kbar, fO2 has to increase up to 10−6 bar to permit H2O dissociation. Whereas the partial pressure of oxygen is high in the zone of sedimentation and as high as ca. 0.21 bar during atmospheric weathering, the oxygen fugacity fO2 is significantly lower in most metamorphic and igneous processes. For instance, the equilibrium curve of the redox reaction 2+ 3+ 6Fe3+ 2 O3 ⇋ 4Fe Fe2 O4 +O2 haematite

magnetite

[27.30]

that delimits the stability field of magnetite towards higher oxygen fugacities is determined by the following points: 600 °C/10−14 bar and T = 400 ◦ C/fO2 = 10−21 bar, −5 1000 °C/10  bar (. Fig. 27.15). The upper T stability limits of magnetite, defined by the redox reactions

Fe2+ Fe3+ 2 O4 ⇋ 3 Fe2+ O + 1/2O2 magnetite

w¨ustite

[27.31]

. Fig. 27.15  T −fO2 diagram displaying the univariant equilibrium curves, limiting the stability fields of haematite, magnetite, wüstite and native iron; The pressure influence on the solid phases can be neglected; Also displayed are dissociation equilibria for H2O at P = 1 bar, and PH2 O = 2 kbar as well as for CO2 at PCO2 = 10 bar and 10 kbar (after Miyashiro 1973)

matite. Consequently, haematite should be the only Fe-oxide mineral in igneous and metamorphic rocks if the amount of oxygen present in the fluid phase was exclusively controlled by dissociation of H2O according to Reaction [27.29]. However, this is definitely not the case as many igneous and metamorphic rocks contain magnetite and/or ilmenite as principle opaque Fe-mineral(s). Therefore, in many igneous and metamorphic processes, fO2 must have been lower or fH2 higher than would have resulted from Reaction [27.29]. One possible solution for this enigma is the presence of organic matter or—at higher metamorphic grade—of graphite, a constituent not uncommon in metasedimentary rocks. As first approximation, the reaction

CO2 ⇋ C + O2

[27.33]

can be applied, the equilibrium curve of which is situated predominantly within the stability fields of magnetite or wüstite. In addition, the reactions

and

CO2 ⇋ CO + 1/2O2 Fe Fe3+ 2 O4 magnetite 2+



3Fe0 +2O2 native iron

[27.32]

are given by the points 400 °C/10−33 bar; 600 °C/10−22 bar and 1000 °C/10−11 bar (. Fig. 27.15). In contrast to the dissociation reaction of H2O [27.29], the influence of the total pressure on the solid-solid redox reactions is negligible. As revealed by . Fig. 27.15, the dissociation curves of H2O are situated almost completely within the stability field of hae-

[27.34]

and

C + 2H2 ⇋ CH4

[27.35]

can contribute to the control of fO2 and fH2. In fact, methane, CH4, which indicates relatively reducing conditions, has been

516

Chapter 27 · Phase Relations and Mineral Reactions in Metamorphic Rocks

recorded in many fluid inclusions in metamorphic minerals (7 Chap. 12). CH4 can constitute a considerable proportion of the fluid phase present during formation of graphite-bearing metasedimentary rocks. At given P-T conditions, CH4 is dominant at low fO2, H2O at medium fO2 and CO2 at high fO2. With increasing temperature and decreasing pressure, the H2O content in the fluid phase decreases and graphite becomes more and more decomposed due to the reaction

2C + 2H2 O ⇋ CO2 + CH4

27

[27.36]

(Ohmoto and Kerrick 1977). In general, the oxygen fugacity can be experimentally determined, at fixed P and T, by univariant equilibrium curves of oxide-oxide or oxide-silicate reactions. For instance, in the system Fe–O, the two solid phases haematite and magnetite coexist with each other and with a fluid phase along the equilibrium curve of Reaction [27.30]. Of the three state variables T, Ptot and fO2, Ptot is kept constant and thus the Gibbs’ Phase Rule takes the form F = C − Ph + 2, in which C = 2 and Ph = 3. With F = 2  −  3 + 2  = 1, the equilibrium curve [27.30] is in fact univariant, i.e., fO2 is automatically fixed at given T. The following univariant equilibrium reactions are widely used as buffer systems to control the oxygen fugacity in hydrothermal experiments: HM

haematite-magnetite buffer according to Reaction [27.30]

NNO nickel-nickel oxide buffer according to reaction NiO ⇌ Ni + ½O2 FMQ

fayalite-(magnetite + quartz) buffer according to reaction

MW

magnetite-wüstite buffer according to Reaction [27.31]

2Fe3O4 + 3SiO2 ⇌ 3Fe2[SiO4] + O2

. Fig. 27.16  T −fO2 diagram at constant Pfl = 2 kbar displaying the stability fields of almandine (deep blue) and the assemblage quartz + Fe-chlorite ± magnetite (light blue) as well as the equilibrium curves of important buffer systems (after Hsu 1968)

IM

magnetite-native iron buffer according to Reaction [27.32]

IW

wüstite-native iron buffer according to reaction FeO ⇌ Fe0 + ½O2

IQF

(native iron + quartz)–fayalite buffer according to reaction Fe2[SiO4] ⇌ 2Fe0 + SiO2 + O2

In hydrothermal experiments, the control of fO2 is achieved by the so-called double capsule method, developed by Eugster (1957). For determining the equilibrium curve of a reaction between silicate minerals, such as staurolite + quartz⇌cordierite + andalusite + H2O at defined fO2, the starting mixture is placed into a precious-metal capsule which, in turn, is surrounded by a larger precious-metal capsule that contains the buffer mixture, e.g., FMQ, together with H2O. At given P-T conditions of the experiment, the FMQ buffer attains a defined fO2 that, on its part, influences the dissociation equilibrium of H2O

H2 O ⇋ H2 + 1/2O2

[27.29]

and, consequently, produces a defined H2-fugacity fH2. The small H2 molecule is able to diffuse through the precious metal of the capsule and thus can also control the dissociation equilibrium of H2O [27.29] in the inner capsule. Thereby, fO2 defined by the buffer mixture in the outer capsule is adjusted in the inner capsule as well. Thus the buffer mixtures listed above furnish an incremental fO2 scale. Fe2+-bearing silicates, such as almandine-rich garnet or staurolite, are stable, at given T and Ptot or Pfl, only over a limited fO2 range and react with increasing fO2 to form magnetite or haematite. As an example, the stability field of pure almandine is displayed in . Fig. 27.16. At constant PH2 O, the lower stability limit of almandine according to the reaction

quartz + Fe-chlorite ± magnetite

[27.37]

⇋ almandine + H2 O

has a relatively steep slope in the T −fO2 diagram and successively crosses the equilibrium curves for the IQF, IM and FMQ buffers. The equilibrium curve for the reaction

quartz + hercynite + magnetite ⇋ almandine + H2 O

27

517

27.2 · Metamorphic Mineral Reactions

[27.38]

has a significantly shallower positive slope and runs nearly parallel to the curve for the FMQ-buffer. In contrast, the upper stability limit of almandine has a negative slope in the T −fO2 diagram. Whereas during prograde or retrograde metamorphism, H2O and CO2 behave relatively mobile, this is obviously not the case for O2 and H2. Many metasedimentary rocks provide evidence that, in individual beds, primary differences in fO2 have been conserved during metamorphism. Thus in adjacent layers of banded iron formation (BIF), e.g., in the Canadian Shield, either haematite or magnetite can be present as Fe-oxides, some with sharp, some with blurred boundaries. Actinolite coexisting with haematite has a lower Fe2+/Mg ratio than with magnetite, thus constituting a measure for O2 fugacity in different layers of the rock. At given T and Ptot, the simultaneous occurrence of magnetite + haematite in metamorphic rocks or ore bodies testifies to a specific fO2 value. Consequently, Reaction [27.30] represents an O2 buffer. 27.2.6  Petrogenetic Grids

During the last five decades, numerous equilibrium curves for metamorphic mineral reactions have been experimentally determined. However, determining the equilibrium curves of all mineral reactions, theoretically conceivable or only observed in nature, by high-P-high-T experiments would be a hopeless task. As we have seen, an alternative possibility is thermodynamic calculation of the position and slope of equilibrium curves within P-T, Pfl-T, T-X or T −fO2 diagrams. For such an approach, however, we need to know all the thermodynamic data of the mineral phases involved, in the relevant P-T range, especially their molar volumes V, enthalpies of formation H, and entropies S, moreover, the relations between molar fractions Xi and activities ai of chemical elements in solid solutions. These data have been derived from calorimetric measurements, crystallographic parameters, or high-P-high-T experiments. Moreover, thermodynamic parameters for H2O and CO2 are available for a broad P-T range. As a result of these investigations, internally consistent thermodynamic data sets have been compiled (Berman 1988; Holland and Powell 1990; Powell et al. 2005; Powell and Holland 2010) that can be used to construct petrogenetic grids for specific model systems. For instance, the equilibrium curves relevant for the metamorphic evolution of pelitic rocks can be displayed in a PH2 O −T diagram, for the model system K2O–FeO–MgO–SiO2–H2O. For metamorphic rocks with marly composition, an isobaric T −XCO2 for the model system

CaO–MgO–Al2O3–SiO2–CO2–H2O (CMASCH) is useful. Phase relations between Mg–Fe solid solutions, e.g., of staurolite, garnet, biotite and chlorite in a metapelite can be demonstrated in isobaric T-XFe or isothermal P-XFe sections. Due to the wealth of univariant equilibrium curves and invariant points documented in metamorphic assemblages, petrogenetic grids can be extremely confusing, especially for complex model systems. However, one should recall that not all possible reactions are really “seen” by each possible bulk composition. Thus for Mg-rich metapelitic bulk rock compositions, reactions are irrelevant in which the Fe-rich minerals such as chloritoid or staurolite participate. Consequently, out of the P-T or T −XCO2 diagrams only those equilibrium curves are selected that are of relevance for a specific bulk rock composition. This considerably simplified approach, showing only a section through the model system, has been named pseudo-section. Combined with careful microscopic analysis of the mineral reactions that took place in a metamorphic rock, pseudosections make it possible to reconstruct prograde and retrograde P-T or T −XCO2 paths. As an example, we present a P-T pseudosection in the model system K2O–FeO–MnO–MgO–Al2O3–SiO2–H2O for a specific bulk rock composition, documenting the metamorphic evolution of a kyanite-staurolite micaschist from the Pan-African Kaoko Belt in northern Namibia (. Fig. 27.16). The diagram shows that only a few of the reactions calculated are univariant. Thus the prograde evolution in this rock leads from the assemblage garnet + chlorite + staurolite + muscovite + quartz to the assemblage garnet + biotite + staurolite + kyanite + muscovite + quartz across a small divariant field, in which already biotite, but not kyanite, occurs in the paragenesis. Most of the fields are even trivariant, each of which contains equilibrium assemblages of three mineral phases + muscovite + quartz + H2O-fluid, i.e., Ph = 6. As the system consists of seven components, C = 7, the Gibbs’s Phase Rule F = C − Ph + 2 leads to F = 7 − 6 + 2 = 3. For a deeper understanding of this matter, the reader is referred to the pertinent textbooks such as the comprehensive presentation by Spear (1993) or the more compact text of Will (1998). 27.3  Geothermometry and Geobarometry

Geothermometers and geobarometers are based on the distribution of chemical components between coexisting mineral phases, such as Mg and Fe in biotite and garnet. The relevant element contents can be determined in situ by electron microprobe analysis (EMPA). Provided the composition of the adjacent solid solutions represent indeed a chemical equilibrium (at the metamorphic peak) and were not reset by diffusion during retrograde metamorphism, the difference in Gibbs free energy for an exchange reaction, for example, can be expressed in terms of partitioning of components between the phases according to the following equation:

G + RT ln K = 0

(27.7)

518

Chapter 27 · Phase Relations and Mineral Reactions in Metamorphic Rocks

in which ΔG is the difference in Gibbs free energy, which should be zero at equilibrium. The equilibrium constant K is calculated from the activities ai of the relevant end-members of the reacting solid solutions. For instance, the exchange equilibrium for the cations Mg and Fe2+ between garnet and biotite   KMg3 AlSi3 O10 (OH)2 + Fe2+ 3 [SiO4 ]3 almandine

phlogopite 2+

⇋ KFe

  3 AlSi3 O10 (OH)2 + Mg3 [SiO4 ]3

[27.39]

pyrope

annite

is defined as

27

ln K = ln

aBt Ann · aGrt Prp aBt Phl · aGrt Alm

�HP,T + (P − 1)�V RT 2

(∂lnK/∂P)T = −�V /RT

(27.8)

(27.9)

(27.10)

(e.g., Will 1998). These equations clearly indicate that exchange equilibria with high ΔH and a low ΔV are suitable as geothermometers because the pressure influence is limited. In contrast, reactions with high ΔV and low ΔH are considerably dependent on pressure but much less on temperature and, consequently, are well suited as geobarometers. For constant lnK values, the following equations can be deduced: 

The activities ai can be calculated, according to the equation ai = γi·Xi from the molar ratios Xi = Fe2+/(Fe2++Mg), provided, the activity coefficients γi are known. The temperature dependence of lnK at constant pressure is given by the equation

(∂ ln K/∂T )P =

the pressure dependence of ln K at constant temperature by

∂P ∂T



= ln K

�SP,T − R ln K �HP,T + (P − 1)�V = (27.11) �V T �V

that correspond to the Clausius-Clapeyron Equation. It also demonstrates that (∂P/∂T)ln K is high, at high ΔH. Thus isopleths, i.e., lines of different lnK values, display a steep slope and, consequently, can be used as a geothermometer. In contrast, (∂P/∂T)ln K becomes small at high ΔV resulting in isopleths for different lnK values with flat slope that form a useful geobarometer (. Fig. 27.18). The intersection points of the isopleths for a geothermometer and a geobarometer define a specific P-T combination, at which a metamorphic process took place.

. Fig. 27.17  P-T pseudosection for a kyanite-staurolite micaschist from the kyanite zone of the Pan-African Kaoko Belt, Namibia, in the KMnFMASH system; The bulk rock composition of the sample is displayed in the uppermost box; Lines of medium thickness: univariant equilibrium curves; dark blue: divariant fields; white: trivariant fields; light blue: quadrivariant field; Quartz (Qz), muscovite (Ms) and H2O are excess phases and are not separately mentioned in the assemblages shown in the various fields; For the trivariant field garnet (Grt) + chlorite (Chl) + staurolite (St) + muscovite (Ms) + H2O, the isopleths for XFe = Fe/(Mn + Fe + Mg) and XMn = Mn/(Mn + Fe + Mg) in garnet are shown as light and broken lines, respectively; The prograde and retrograde P-T path, based on the P-T combinations (I)–(V) is displayed by a thick black line (from Gruner 2000)

519

27.3 · Geothermometry and Geobarometry

. Fig. 27.18  Schematic position of possible geothermometers and geobarometers in the P-T diagram (after Will 1998) In many cases, e.g., in P-T pseudosections, lines of the same chemical composition of a specific mineral are drawn that are also known as isopleths. As an example, the steep isopleths for XFe and XMn in garnet that has grown in exchange equilibrium with chlorite and staurolite are displayed in . Fig. 27.17. With increasing temperature, XFe increases whereas XMn markedly decreases.

Equilibria involving cation exchange analogous to Reaction [27.39] commonly display a small ΔV and a high ΔH and thus are frequently used as geothermometers. Prominent examples are the mineral pairs garnet–clinopyroxene, garnet–orthopyroxene, garnet–cordierite, garnet–amphibole, garnet–phengite, clinopyroxene–orthopyroxene, magnetite– ilmenite and calcite–dolomite (e.g., Will 1998). In contrast, many useful geobarometers are based on so-called mass-transfer reactions, in which cations underwent a change in coordination number. If a cation attains different coordination numbers in the reactants and products, respectively, such as Al[4] and Al[6], the exchange reaction commonly has a high ΔV. A well-known example is the so-called GASP barometer (from Grossular–Al-silicate–SiO2–Plagioclase) according to the reaction 2Al[6] 2 [SiO5 ] + Ca3 Al[6] 2 [SiO4 ]3 kyanite

grossular [4]

+ SiO2 ⇋ 3Ca[Al quartz

2 Si2 O8 ] anorthite

[27.40]

Additional geobarometers of this type are based on the reaction equilibria grossular-almandine garnet + rutile = ilmenite + anorthite + quartz (GRIPS), almandine garnet + rutile = ilmenite + Al2SiO5 + quartz (GRAIL), cordierite = pyrope-almandine garnet + sillimanite + quartz, albite = jadeite + quartz, and others (e.g., Will 1998). Frequently used for pressure estimates is the phengite geobarometer. It is based on the Si content in white mica, ranging in composition from muscovite to phengite, according to the coupled substitution Al[6]Al[4]⇌Mg[6]Si[4]. A necessary prerequisite is the coexistence of white mica with

. Fig. 27.19  Phengite geobarometer; PH2 O −T diagram displaying the isopleths for Si contents in phengite (p.f.u.) in equilibrium with K-feldspar, phlogopite, quartz and H2O; Upper T stability limit for muscovite + quartz after the reactions [27.11] and [27.11a] (after Massonne and Schreyer 1987)

(1) phlogopite + K-feldspar + quartz or with (2) talc + phlogopite + kyanite in the KMASH system (Massonne and Schreyer 1987, 1989). Natural rocks typically contain also Fe and biotite occurs instead of the rarer phlogopite. In . Fig. 27.19, the Si-isopleths in phengite are displayed for the assemblage (1). It is obvious that the Si contents are strongly influenced by pressure but much less so by temperature. Thus the diagram can be used as a rather sensitive geobarometer to estimate the pressure of a metamorphic process without knowing its precise temperature. The phengite geobarometer has been broadly applied to metamorphosed granites and arkoses consisting of biotite, phengite (muscovite), K-feldspar and quartz. Under favourable conditions, the Si-content in relic phengite can indicate a preceding stage of high-P metamorphism. The exchange of cations between coexisting minerals is a diffusional process whereby the diffusion rate decreases exponentially with decreasing temperature. Below a specific temperature, known as closure temperature, no significant diffusion takes place over geological time scales, and the exchange equilibrium attained becomes frozen. If, however, the peak-temperature of a metamorphic event is higher than the closure temperature of the geothermometer used, diffusion can still take place and the exchange equilibrium can be reset. Consequently, the temperature calculated in such a case would not conform to the maximum temperature reached but rather to some point on the retrograde P-T path. In order to estimate the P-T development of a metamorphic complex, it is useful to apply different geothermometers and geobarometers to one or several metamorphic rocks. It should be noted, however, that the closure temperature for the cation exchange between two

27

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Chapter 27 · Phase Relations and Mineral Reactions in Metamorphic Rocks

coexisting minerals cannot be determined with high accuracy as it is influenced by several parameters of chemical kinetics, such as heating and cooling history, rate of deformation and fluid flow. Under favourable circumstances, exchange equilibria can be used to estimate P-T conditions for points on the prograde path, e.g., using EMPA data of mineral inclusions in zoned garnet porphyroblasts. Based on the compositional zoning of minerals, the Gibbs’ method of differential thermodynamics can be applied for reconstruction of P-T paths, (Spear 1988; Spear et al. 1991; Zeh and Holness 2003). 27.4  Pressure-Temperature Evolution

27

of Metamorphic Complexes

The reconstruction of the regional and temporal pressure-temperature development of metamorphic complexes is of major importance in geological research because it helps to gain an understanding of orogenic processes in the geological past as well as today. In general, crustal-scale plate tectonic processes, such as subduction, continent-continent collision, continental rifting, related plutonic or volcanic activity, as well as the formation of new oceanic crust at mid-ocean ridges, all cause changes of pressure and temperature in space and time. Consequently, these processes lead to a disturbance of the steady-state geotherm in a specific region. As a result, prograde and retrograde mineral reactions take place that can be reconstructed by careful microscopic observations and quantified based on thermodynamic principles.

27.4.1  Pressure-Temperature Paths

Based on detailed petrography under an optical microscope and EMPA of metamorphic minerals, pressure-temperature paths that have been followed by metamorphic rocks can be reconstructed. For a (semi-)quantitative approach, petrogenetic grids, especially pseudosections (. Fig. 27.17), are combined with results from geothermometry and geobarometry. In addition, isochores of fluid inclusions in minerals can be used to set further constraints on P-T conditions (. Fig. 12.4). With the aid of textural observations, it can be possible to relate crystallisation of minerals grown on the prograde or retrograde P-T path to individual phases of deformation, D1, D2, D3, …, Dn, thus making it possible to reconstruct a pressure-temperature-deformation path (P-T-D path). In principle, two different types of P-T paths can be distinguished, one following a clockwise the other a counterclockwise direction when plotted onto P-T space (. Fig. 27.20). These reflect different mechanisms of orogeny but can occur side by side in different parts of the same orogenic belt. In many cases, the temperature peak (A or D) reached during metamorphism does not coincide with the pressure maximum (B or C).

. Fig. 27.20  Two principally different types of P-T paths, one being clockwise (red), the other one counter-clockwise (blue); The points B and C indicate maximum pressures, the points A and D maximum temperatures

Clockwise P-T paths  have been recorded in many metamorphic complexes. As shown by theoretical modelling, first by England and Thompson (1984), they are the result of crustal thickening in the course of subduction and subsequent continental collision. At the early stage a relatively fast transport to great depths is accompanied by a strong pressure increase, while the increase in temperature lags behind because of the poor thermal conductivity of rock-forming minerals (see . Fig. 28.2). Once the subduction comes to a halt, the local thermal gradient gradually adjusts to more normal values and radiogenic heat production, heat conduction and/or advective heat supply by igneous intrusions all lead to regional increase in temperature in the subducted crustal slab without any significant further pressure increase. Regional heating even continues, however, during pressure release caused by the isostatic uplift of the thickened orogen, during which erosion and/or dissecting along low-angle normal faults leads to crustal thinning. After this phase of nearly isothermal decompression, the P-T path approaches a normal geothermal gradient. Such a development is demonstrated by a combined P-T path consisting of the branches I (black) and III (red; . Fig. 27.21). On the P-dominated prograde branch I, high-pressure and ultra-high pressure rocks such as blueschist and eclogite are formed. Upon temperature increase, the high-P assemblages are widely or completely replaced by medium-P assemblages of Barrovian type (branch III) and can undergo even partial melting, provided a H2O-rich fluid phase is present. A better chance for preserving high-P assemblages exists if high-pressure and ultra-high pressure rocks are quickly transported back up to higher levels of the crust, due to tectonic activity such as overthrust of nappes. These processes lead to characteristic hairpin shaped P-T paths in which the prograde and the retrograde branch are nearly parallel to each other. A prominent example is the Dora Maira Massif of the Italian Alps, in which Chopin (1984) first detected the ultra-high pressure assemblage pyrope + coesite (. Fig. 28.7) indicating extreme P-T conditions of about 800 °C and 30 kbar (. Fig. 28.6). The prograde branch I (black) of the P-T path documents an early stage of continent-continent

27.4 · Pressure-Temperature Evolution of Metamorphic Complexes

521

. Fig. 27.21  P−T diagram displaying different P-T paths that can be followed by crustal rocks during deep subduction and subsequent exhumation at variable velocity (see text); Three combined P-T paths are shown, formed by the branches: I (black), II (blue), III (red) and IV (green); Also displayed are the H2O-saturated and the dry solidus curves of an alkaline granite as well several linearshaped geothermal gradients as reference (from Schreyer 1988)

collision and, consequently, follows a very low geothermal gradient of ca. 7 °C km−1. The retrograde branch II (blue) of the P-T path is nearly parallel to the prograde one which implies extremely fast tectonic exhumation. Moreover, the preservation of the ultra-high P assemblage pyrope + coesite is favoured by a lack of H2O and the extremely large grain size of garnet, whereas the adjacent fine-grained metapelite, rich in white mica, has been overprinted by retrograde reactions. Since then, relict ultra-high P assemblages, formed in deeply subducted continental crust, have been recorded in several metamorphic complexes (7 Sect. 28.2.9). Branch IV (green) is the continuation of the subduction branch I to greater depths of the Earth’s crust and mantle down to >200 km. As various solidus and liquidus curves for granitoid rocks are crossed along this P-T path, progressive partial melting of the crustal rocks takes place with increasing temperature, depending on the prevailing H2O fugacity (. Figs. 20.5, 20.6). The melts thus formed approach syenitic composition with increasing pressure (. Fig. 20.3, inset; 7 Sect. 20.2.2) and can react with peridotite of the Earth’s upper mantle, presumably leading to magma-forming processes on a global scale.

Counter-clockwise P-T paths  (. Fig. 27.20) can develop in island arcs or orogenic belts above subduction zones (. Fig. 28.2) due to advective heat transfer from intruding magma. If such intrusions are voluminous and of regional extent, isobaric heating at relatively shallow crustal levels can reach spatial proportions similar to those of regional metamorphism, hence described as regional contact metamorphism. Later-on, pressure increases as a result of crustal thickening, e.g., caused by nappe stacking, followed by nearly isobaric cooling. By this process, low-pressure metamorphic rocks of Buchan-type are formed.

27.4.2  Pressure-Temperature-Time Paths

The duration of metamorphic processes is of fundamental relevance for understanding the evolution of orogenic belts. Theoretical modelling has revealed that contact aureoles, formed due to the thermal overprint caused by small igneous intrusions (900 5 U–Pb dating of garnet: >800 5 U–Pb dating of monazite: 700–650 5 U–Pb dating of titanite: 670–500 5 Sm–Nd dating of garnet: ~600 5 Rb–Sr dating of muscovite: ~500 5 K–Ar dating of hornblende and muscovite: ~450–400 5 U–Pb dating of rutile: 430–380

27

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Chapter 27 · Phase Relations and Mineral Reactions in Metamorphic Rocks

5 K–Ar dating of biotite: ~300

27

Nevertheless, dating of different minerals in a rock or in a rock series using different isotopic systems can put time markers at the metamorphic peak or at different steps of the retrograde branch of the P-T path whereas, naturally, this is hardly possible for the prograde branch. A necessary prerequisite is, of course, that each of the minerals dated crystallised at the same metamorphic event. This is not always the case, especially for zircon with its high closure temperature. Many single zircon grains or inner cores of those reveal age information on a preceding igneous or metamorphic event rather than the actual metamorphic event of interest (so-called inherited ages). In this case, the temperature reached at the metamorphic peak was not high enough to reset the isotope system. For instance, zircon grains in orthogneisses, still displaying typical magmatic morphology with idiomorphic crystal shape, can be used to date the intrusion age of the granitic protolith (. Figs. 33.15, 33.16). In contrast, zircon grains of rounded shape can yield U–Pb ages that reflect the age of an ancient sediment source of a detrital zircon grain that is now present in a paragneiss for example. Such zircon grains had been transported as detrital heavy mineral into a sedimentary basin, the fill of which was subsequently metamorphosed during a later orogeny. Many zoned zircon grains display older cores with primary igneous or detrital rounded particle shape that were later overgrown by metamorphic rims (e.g., Harley and Kelly 2007; Harley et al. 2007; Rubatto and Hermann 2007). Today, such complex age patterns in a single zircon grain can be revealed by in situ isotope analyses, such as sensitive high-resolution ion microprobe (SHRIMP) or a laser ablation-inductively coupled plasma-mass spectrometer (LA-ICP-MS) analyses (7 Chap. 12). . Figure 27.22 displays the temperature-time development of the Adirondack Metamorphic Complex, New York State, USA, which experienced high-grade meta-

morphism with peak temperatures of ca. 750° and pressures of ca. 7.5 kbar. U–Pb dating on garnet yielded an age of 1064 ± 3 Ma for the peak of this metamorphic event, whereas U–Pb dating of monazite, titanite and rutile as well as K–Ar dating of hornblende and biotite revealed increasingly younger ages, which indicate cooling down to 300 °C within a broad time span of nearly 250 million years. Over this period, the cooling rate had decreased from ca. 4 °C Ma−1 down to ca. 1 °C Ma−1 (Mezger et al. 1990). The uplift rate of the Adirondack Metamorphic Complex has been estimated at about 0.05 mm a−1. By comparison, the present uplift of young orogenic belts, such as the Himalaya Mountains, attains considerably higher rates of 0.2– 0.5 mm a−1, in part even of 4 mm a−1. 27.5  Graphical Presentation of Metamorphic

Mineral Assemblages

The overwhelming majority of silicate rocks consists of the twelve major components SiO2, TiO2, Al2O3, Fe2O3, FeO, MnO, MgO, CaO, Na2O, K2O, P2O5, H2O, with additional CO2 in calcareous mica schists or calc-silicate rocks. It can be useful to plot mineral and bulk rock compositions on a diagram to illustrate systematic and predictable phase relations in metamorphic rocks of various bulk rock compositions at different metamorphic grade. The number of components is, however, far too large to be plotted. At best, four components could be displayed as a tetrahedron but even this is difficult to show on a 2-dimensional piece of paper. More readily plotted is a three-component system in the form of a compositional triangle—a so-called chemographic diagram.

Theoretically, any three-component system can be easily shown as a triangular chemographic diagram, such as SiO2–MgO–CaO, approximating the composition of Fe-free ultramafic rocks. Most rocks contain, however, more than three components, and this requires some simplifications in order to show them as triangles. Some of the most popular triangular chemographic diagrams are briefly described below. 27.5.1  ACF and A′KF Diagrams

. Fig. 27.22  Temperature-time diagram for the cooling history of the Adirondack metamorphic complex, New York, USA, based on U–Pb dating of garnet, monazite, titanite, and rutile as well as K–Ar dating of hornblende and biotite. The domains of uncertainty for the closure temperatures of individual minerals are indicated by coloured boxes (after Mezger et al. 1990, from Spear 1993)

The Finnish petrologist Pentti Eskola (1883–1964) established the ACF and A′KF diagrams, the purpose of which is to graphically display the phase relations in different metamorphic rocks of different bulk rock composition. They represent a convenient way of illustrating the relationships between whole rock chemistry, mineral assemblage and metamorphic grade. To plot these diagrams, the following steps are taken:

523

27.5 · Graphical Presentation of Metamorphic Mineral Assemblages

1. The weight percentages of the components obtained by the chemical analysis are divided by the molecular weight and thus converted into mole fractions. 2. Only rocks are represented that are oversaturated in SiO2, i.e., containing free quartz (or any other SiO2 polymorph). In these rocks, only minerals with the highest possible SiO2 content can be plotted, such as enstatite rather than forsterite, or andalusite rather than corundum. Consequently, the SiO2 content in the bulk rock composition has no influence on the phase relations and SiO2 does not need to be considered as a component. In contrast, different phase diagrams must be used for SiO2-undersaturated rocks, such as metamorphosed ultramafic rocks or bauxite, for which SiO2 must be regarded as independent component. 3. H2O and CO2 can be regarded as completely mobile components. Thus their fugacities or partial pressures— fH2 O ∼ PH2 O , fCO2 ∼ PCO2—are considered as external variables of state, analogous to lithostatic pressure P and temperature T. 4. The only rock-forming mineral containing P2O5 is apatite, whereas most of the TiO2 is located either in rutile, ilmenite or titanite. In each case, one extra component is present in one extra phase, and consequently, the degrees of freedom according to the Gibbs’ Phase Rule do not change. These accessory minerals are therefore ignored in graphical representations. If bulk rock analyses are plotted in ACF and A′KF diagrams, corrections are necessary to account for these accessory minerals, the amount of which must be estimated or determined by modal analysis. According to the chemical formula of apatite, an equivalent amount of 3.3 × P2O5 has to be subtracted from CaO. For ilmenite or titanite the amounts equivalent to 1 × TiO2 are subtracted from FeO or CaO, respectively. It should be noted, however, that ACF and A′ KF diagrams are essentially designed to represent phase relations between coexisting minerals rather than projections of bulk rock analyses.

. Fig. 27.23  Projection of common metamorphic minerals in ACF and A′KF diagrams (after Winkler 1979)

5. Of the eight remaining components, FeO + MnO + MgO as well as Al2O3 + Fe2O3 are combined for the sake of simplicity. The calculation is performed in the following way, whereby, e.g., [FeO] means moles FeO.

ACF diagram A = [Al2 O3 ] + [Fe2 O3 ]−([Na2 O] + [K2 O]) C = [CaO] F = MgO + [FeO] + [MnO] A + C + F = 100 A′ KF diagram A′ = [Al2 O3 ] + [Fe2 O3 ]−([Na2 O] + [K2 O] + [CaO]) K = [K2 O] F = MgO + [FeO] + [MnO] A′ + K + F = 100 Examples for calculation are given in Appendix A.2

In both cases, Al bonded with Na in albite or with K in K-feldspar is not considered. In the A′ KF diagram, the same holds true for Al bonded with Ca in anorthite. Not depicting Na2O and combining FeO + MnO + MgO as well as Al2O3 + Fe2O3 are major drawbacks for both diagrams. Apart from that, the most important silicate minerals, present in metamorphic rocks, can be represented in ACF and/or A′ KF diagrams (. Fig. 27.23), with exception of albite, jadeite or paragonite. In . Fig. 27.24, the approximate compositional fields of igneous and sedimentary rocks are displayed, which can serve as common protoliths for metamorphic rocks. By comparing . Fig. 27.23 and 27.24, it becomes obvious that phase relations in calc-silicate rocks or metabasites are well displayed in the ACF triangle

27

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Chapter 27 · Phase Relations and Mineral Reactions in Metamorphic Rocks

a higher MnO content shifts the lower stability limit of garnet to lower temperatures, in which case combining MnO with FeO + MgO to form one component would be wrong. In many metamorphic reactions, the Fe/Mg ratios of the minerals involved continuously change thus making these gliding (or continuous) reactions divariant. Consequently, FeO and MgO should not be combined to form one single component in rocks in which ferromagnesian minerals have different Fe/Mg ratios (as is the case in metapelites). The same holds true for divariant or multivariant reactions with changing Fe2O3/Al2O3 or other oxide ratios.

27

. Fig. 27.24  Approximate fields for the bulk rock compositions of common igneous and sedimentary rocks in an ACF or an A′KF diagram; G granite, Gd granodiorite, Gb gabbro, B basalt, P peridotite

whereas, for mica-rich metapelites, the A′ KF triangle is much more appropriate. In ACF and A′ KF diagrams, binary solid solutions are depicted as bold lines connecting the two end members, whereas ternary solid solutions are displayed as elongate fields (. Fig. 27.23). Coexisting minerals are linked by tie lines, coexisting solid solutions by bundles of tie lines. The tie lines delimit three-phase triangles, in which each three minerals coexist in equilibrium together with quartz and a fluid. Additional excess phases are the albite content in plagioclase ± K-feldspar ± muscovite in the ACF diagram, and the albite content in K-feldspar and/or plagioclase in the A′ KF diagram. Bulk rock compositions that plot onto a bundle of tie lines or within a three-phase triangle can vary arbitrarily without changing their respective mineral assemblage, whereas the modal proportions of the coexisting minerals are altered. A new assemblage is formed as soon as the bulk chemical composition is shifted across a limiting tie line.

A further problem in the construction of ACF and A′ KF diagrams is the fact that EMPA cannot distinguish between FeO and Fe2O3.

The hornfelses in the Oslo region of southern Norway, investigated in detail by V. M. Goldschmidt (1911) and studied again by Eskola (1915) were the first to which the ACF triangle was applied. The contact metamorphic overprint of the different sedimentary protoliths took place essentially isochemically. Their chemical composition varies from carbonate-free and carbonate-poor mudstone to marly limestone and thus covers broad areas of the ACF diagram. Conforming to the different bulk rock compositions, Goldschmidt distinguished 10 so-called hornfels classes that are characterised by the following two- or three-phase assemblages (. Fig. 27.25): 1. andalusite-cordierite 2. andalusite-cordierite-plagioclase 3. cordierite-plagioclase 4. cordierite-plagioclase-hypersthene 5. plagioclase-hypersthene 6. plagioclase-hypersthene-diopside 7. plagioclase-diopside

Due to an over- or under-correction of the accessory minerals mentioned above, a specific bulk rock composition can plot off the “correct” bundle of tie lines or outside the three-phase fields derived from microscopic observations. Such a situation can also be due to a lack of equilibrium.

If more than three phases seem to coexist in a given rock, crossing tie lines are observed in ACF and A′KF diagrams (. Fig. 27.26). Such a situation can have various reasons: 5 The Gibbs’ Phase Rule is violated and the rock represents thermodynamic disequilibrium. In some cases, the problem can be solved by regarding a smaller area of the thin section, in which local equilibrium was reached. 5 A univariant or divariant mineral reaction is frozen in the rock. Either the sample was collected in the field just on an isograd or one or two reactants have survived as metastable relics. 5 The number of components selected is too low. For instance, a higher TiO2 content can shift the upper stability limit of biotite towards higher temperatures and neglecting this component is unjustified. Similarly,

. Fig. 27.25  ACF diagram for the 10 hornfels classes of Goldschmidt (1911), conforming to the pyroxene-hornfels facies of Eskola (1939; see 7 Sect. 28.1); In all the assemblages indicated by tie lines or threephase fields, quartz, K-feldspar and/or biotite can occur as additional phase(s)

525

27.5 · Graphical Presentation of Metamorphic Mineral Assemblages

8. plagioclase-diopside-grossular 9. diopside-grossular 10. diopside-grossular-wollastonite In addition to the above minerals, quartz, K-feldspar and/or biotite can be present. The quartz content decreases with increasing class number and is absent in many hornfelses of class 8–10, which can contain calcite instead. However, note that, by definition, quartz-free assemblages should not be plotted on the ACF triangle. In ACF and A′ KF diagrams, mineral assemblages are presented that have equilibrated, within a specific range of P-T conditions and in protoliths of different bulk-rock compositions. Thus crystallisation paths, displayed in ternary model systems for igneous rocks, such as diopside– albite–anorthite (. Figs. 18.6, 18.7, 18.8), are not shown. An advantage of diagrams such as ACF and A′ KF lies in the fact that phases that cannot be connected by tie lines should not occur together as equilibrium assemblage, e.g., cordierite–diopside or hypersthene–grossular. Other mineral combinations observed under the microscope for the same bulk-rock composition must have formed at different P-T conditions and should not be represented in the same ACF or A′ KF diagram. 27.5.2  AFM Projections

For numerous metamorphic rocks, ACF and A′ KF diagrams are sufficient to display the equilibrium relations of the mineral assemblages based on these three components. They are, however, not particularly useful for many other rocks, especially metapelite. As we have seen, FeO and MgO are combined to a single component in the ACF and A′ KF diagrams. Although nearly unlimited diadochy between Mg and Fe2+ exists in many mafic minerals, the mutual substitution can be restricted towards either of the two end members at specific P-T conditions. For instance, at moderate temperature and pressure, Fe-rich garnet coexists with Mg-rich cordierite. This leads to assemblages of four minerals coexisting with each other in equilibrium, e.g., garnet–cordierite–biotite–K-feldspar(–quartz), which cannot be displayed in the A′ KF diagram without crossing tie lines, thus apparently contradicting the Gibbs’ Phase Rule (. Fig. 27.26). Consequently, for many metamorphic rocks in which the Fe/Mg ratio in coexisting ferromagnesian minerals is different, especially metapelites, FeO and MgO must be treated as separate components. In the AKFM tetrahedron designed by Thompson (1957), the five major components SiO2 (in excess), Al2O3, FeO, MgO and K2O are considered and the phase relations are projected onto a plane that contains the three components AFM at the base of the tetrahedron (. Fig. 27.27a). For all muscovite-bearing rocks, muscovite serves as projection point, whereas K-feldspar is used as projection point for muscovite-free, K-feldspar bearing rocks, especially metapelites of higher metamorphic grade in which Reaction [27.11]

. Fig. 27.26  Presentation of the four-phase assemblage garnet– cordierite–biotite–K-feldspar in an A′KF triangle is not possible without intersection of tie lines, due to the fact that the combination of FeO and MgO to form one single component is not permissible beause the Fe2+/Mg ratios in all three mafic minerals are different

has occurred. Consequently, muscovite and K-feldspar are not depicted in the AFM projection. The same holds true for quartz that, by definition, is present in excess. It should be noted that, in the AFM projection, Fe2O3 and MnO are not combined with Al2O3 and FeO + MgO, respectively, but are neglected. When calculating the position of garnet, an equivalent proportion of Al2O3 must be subtracted that is bound in the end members spessartine to 1/3MnO and grossular to 1/3CaO. By analogy to the ACF and A′ KF diagrams, the AFM projection is calculated, based on moles, as follows: 1. With Muscovite As Projection Point

A = [Al2 O3 ] − 3[K2 O] M = [MgO] F = [FeO] vertical scale:

A A+F+M

horizontal scale:

M M+F

Explanation  In K2O-bearing minerals, an amount of Al2O3 has to be subtracted that is equivalent to K2O. For muscovite, KAl2[AlSi3O10](OH)2 or K2O·3Al2O3· 6SiO2·2H2O, this is 3[K2O]. For biotite, K(Mg, Fe)3[AlSi3O10](OH)2 or K2O·3(Mg, Fe)O·Al2O3·6SiO2·2H2O, negative values are obtained for A/(A + F + M), as 3[K2O] > [Al2O3]. Consequently, biotite plots below the FM edge of the AKFM tetrahedron (. Figs. 27.27a, b, 27.28). In contrast, projection points of K-free minerals, such as chlorite, almandine-pyrope garnet, cordierite, chloritoid or staurolite plot within the AFM triangle. Solid solutions formed by Fe2+-Mg substitution in these minerals are expressed by a line parallel to the F-M edge. If, in addition, Al is substituted for Fe2+ and Mg, the line is widened to form a band or a field. This is especially the case for chlorite and biotite (. Fig. 27.27b). By definition, muscovite cannot be plotted on the diagram.

27

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Chapter 27 · Phase Relations and Mineral Reactions in Metamorphic Rocks

27

. Fig. 27.27  AFM projection; a AKFM  tetrahedron Al2O3–K2O–FeO–MgO with the plane of projection A = [Al2O3] − F = [FeO] − M = [MgO] extending below the F–M side of the triangle; All points within the tetrahedron, such as a X, Y and B can be projected through Ms (= muscovite), as points X′, Y′ and B′ onto the AFM plane where Y′ and B′ plot beyond the line F–M; Of course, all K-free minerals plot within the AFM triangle. b AFM projection of the most important mineral compositions with projection through Ms (after Best 2003)

2. With K-Feldspar as Projection Point

A = [Al2 O2 ] − [K2 O] M = [MgO] F = [FeO] vertical scale:

A A+F+M

horizontal scale:

M M+F

The same result is obtained by a simple calculation based on molar percentages:

A + M + F = 100 Explanation  If K-feldspar serves as projection point, all

minerals are plotted onto and within the AFM triangle. Examples for calculation are given in Appendix A.2.

In principle, the AFM projection is not designed for presentation of bulk rock compositions. If nevertheless this is intended, relevant cor-

527 References (see also Chapters 26 and 28)

rections must be performed, e.g., for the FeO contents in ilmenite, FeO·TiO2, or magnetite, FeO·Fe2O3, accessory minerals that are not displayed in the AFM projection. This leads to F = [FeO]–[TiO2]–[Fe2O3]. As Fe2+ is replaced by Mn2+ in many silicate minerals, [MnO] can be combined with [FeO] to form F of the bulk rock.

The mineral assemblages of three metapelitic hornfelses P, Q and R, stable at P-T conditions of the hornblende-hornfels facies (7 Sect. 28.2.6), are displayed in the AFM projection in . Fig. 27.28. The assemblage andalusite–biotite–cordierite(–muscovite–quartz) formed in a metapelite of bulk composition P, whereas metapelites Q and R contain the assemblage biotite–cordierite(–muscovite–quartz). Coexisting cordierites and biotites of different Mg/Fe2+ ratio are connected by tie lines. By analogy to ACF and A′ KF diagrams, the composition of solid solutions and the resulting orientation of the tie lines change with changing pressures and temperatures. At constant P-T conditions, the endpoints of the three phase triangle Bt1-Crd1-andalusite are fixed, whereas the bulk rock composition can freely vary within the triangle. If point P is shifted towards A, the assemblage would not change but only the modal amount of andalusite would increase at the expense of biotite and cordierite. In contrast, the Mg/Fe2+ ratios in the twophase assemblage biotite–cordierite(–muscovite–quartz) are not only controlled by P and T, but also by the MgO/ (MgO + FeO) ratio of the bulk rock. Thus Bt2 and Crd2 crystallise in the more ferroan hornfels Q, Bt3 and Crd3 in the more magnesian hornfels R. Shifting of the bulk rock compositions along the tie line would only change the modal biotite/cordierite ratio within the respective rock. Phase relations in metapelites are much better represented by AFM projections than ACF and A′ KF diagrams.

. Fig. 27.28  AFM projection of two mineral assemblages in metapelitic hornfelses; Rock P: biotite (Bt1) + cordierite (Crd1) + andalusite (+ muscovite + quartz); Rocks Q and R: biotites Bt2 and Bt3 + cordierites Crd2 and Crd3 (+ muscovite + quartz)

Similar projections can be developed for metamorphic rocks of all sorts of bulk composition, an example being the ACFM tetrahedron for metabasites with plagioclase as projection point (Robinson et al. 1982).

References (see also Chapters 26 and 28) Berman RG (1988) Internally consistent thermodynamic data for minerals in the system Na2O–K2O–CaO–MgO–FeO–Fe2O3–Al2O3–SiO2– TiO2–H2O–CO2. J Petrol 29:445–522 Bohlen SR, Montana A, Kerrick DM (1991) Precise determination of equilibria kyanite ↔ sillimanite and kyanite ↔ andalusite and a revised triple point for Al2SiO5 polymorphs. Amer Miner 76:677–680 Chatterjee ND (1970) Synthesis and upper stability limit of paragonite. Contrib Miner Petrol 27:244–257 Chatterjee ND (1972) The upper stability limit of the assemblage paragonite + quartz and its natural occurrences. Contrib Miner Petrol 34:288–303 Chatterjee ND, Johannes W (1974) Thermal stability and standard thermodynamic properties of synthetic 2M1-muscovite, KAl2[AlSi3O10(OH)2]. Contrib Miner Petrol 48:89–114 Chernosky JV Jr, Day HW, Caruso LJ (1985) Equilibria in the system MgO–SiO2–H2O: experimental determination of the stability of Mg-anthophyllite. Amer Miner 70:223–236 Cho M, Maruyama S, Liou JG (1987) An experimental investigation of heulandite-laumontite equilibrium at 1000 to 2000 bar Pfluid. Contrib Miner Petrol 97:43–50 Chopin C (1984) Coesite and pure pyrope in high-grade blueschists of the Western Alps: a first record and some consequences. Contrib Miner Petrol 86:107–118 England PC, Thompson AB (1984) Pressure–temperature–time paths of regional metamorphism. Part I: heat transfer during the evolution of regions of thickened continental crust. J Petrol 25:884–928 Eskola P (1915) On the relations between the chemical and mineralogical composition in the metamorphic rocks of the Orijärvi region. Bull Comm géol Finlande 44 (English summary pp 109–145) Eskola P (1939) Die metamorphen Gesteine. In: Barth TFW, Correns CW, Eskola P (eds) Die Entstehung der Gesteine—Ein Lehrbuch der Petrogenese. Springer, Berlin, pp 263–407 (Reprint 1981) Eugster HP (1957) Heterogeneous reactions involving oxidation and reduction at high temperatures. J Chem Phys 26:1760–1761 Evans BW, Johannes W, Oterdoom H, Trommsdorff V (1976) Stability of chrysotile and antigorite in the serpentinite multisystem. Schweiz Miner Petrogr Mitt 56:79–93 Goldschmidt VM (1911) Die Kontaktmetamorphose im Kristianiagebiet. Oslo Vidensk Skr, I Math Nat Kl, no 11 Greenwood HJ (1961) The system NaAlSi2O6–H2O–argon: total pressure and water pressure in metamorphism. J Geophys Res 66:3923–3946 Greenwood HJ (1967) Mineral equilibria in the system MgO–SiO2–H2O– CO2. In: Abelson PH (ed) Researches in Geochemistry. Wiley, New York, pp 542–567 Griffen DT (1992) Silicate Crystal Chemistry. Oxford University Press, Oxford Gruner BB (2000) Metamorphoseentwicklung im Kaokogürtel, NW-Namibia: Phasenpetrologische und geothermobarometrische Untersuchungen panafrikanischer Metapelite. Freiberger Forschungshefte C468:221 pp (Freiberg) Harker RO, Tuttle OF (1956) Experimental data on the PCO2-T curve for the reaction calcite + quartz = wollastonite + carbon dioxide. Amer J Sci 254:239–256 Hemley JJ (1967) Stability relations of pyrophyllite, andalusite, and quartz at elevated pressures and temperatures. Amer Geophys Union Trans 48:224

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27

Chapter 27 · Phase Relations and Mineral Reactions in Metamorphic Rocks

Holdaway MJ (1971) Stability of andalusite and the aluminium silicate phase diagram. Amer J Sci 271:97–131 Holdaway MJ, Mukhopadhyay B (1993) A reevaluation of the stability relations of andalusite: thermochemical data and phase diagram for the aluminum silicates. Am Mineral 78:298–315 Holland TJB, Powell R (1985) An internally consistent thermodynamic dataset with uncertainties and correlations: 2. Data and results. J Metam Geol 3:343–370 Holland TJB, Powell R (1990) An enlarged and updated internally consistent thermodynamic dataset with uncertainties and correlations: the system K2O–Na2O–CaO–MgO–MnO–FeO–Fe2O3– Al2O3–TiO2–SiO2–C–H2–O2. J Metam Geol 8:89–124 Hsu LC (1968) Selected phase relationships in the system Al–Mn–Fe– Si–O–H: a model for garnet equilibria. J Petrol 9:40–83 Kennedy GC, Holser WT (1966) Pressure–volume–temperature and phase relations of water and carbon dioxide. Geol Soc Am Mem 97:371–384 Kerrick DM (1968) Experiments on the upper stability limit of pyrophyllite at 1.8 kilobars and 3.9 kilobars water pressure. Amer J Sci 266:204–214 Kerrick DM (1972) Experimental determination of muscovite + quartz stability with PH2O 600 °C, at best reached in the high-grade part of the amphibolite facies. On the other hand, virtually monomineralic, medium- to coarse-grained marbles are produced by the recrystallisation of pure limestone or dolomite. z Metapelites

Essentially based on phase relations in metapelites, the amphibolite-facies could be subdivided into different subfacies. For instance, break-down of muscovite in the presence of quartz to form either andalusite/sillimanite + K-feldspar or kyanite/sillimanite + melt according to the reactions [27.11], [27.11a] and [27.11b] (. Figs. 27.8, 27.11) leads to the distinction between a relatively lowgrade and high-grade amphibolite facies, respectively. These can be further divided into a low, intermediate and high P/T range, based on the stability of the Al2SiO5 polymorphs andalusite, sillimanite and kyanite (. Figs. 27.8, 27.11). At a medium geothermal gradient, i.e., in the medium-pressure facies series, the lower-grade amphibolite facies conforms to Barrow’s staurolite and kyanite zones. From mudstone and greywacke, micaschists and paragneisses are formed, in which commonly muscovite + biotite + almandine-rich garnet ± staurolite ± kyanite/ sillimanite + quartz + plagioclase coexist in equilibrium. The presence of staurolite together with garnet and, depending on pressure, of kyanite or sillimanite is characteristic of the lowest amphibolite facies (. Fig. 28.4c). Staurolite is formed, e.g., by the following reactions (reaction 2 in . Fig. 26.11):

4Fe2+ Al2 [SiO4 ]O(OH)2 + 5Al2 [SiO5 ]

z Metabasites

Mafic rock compositions are represented by medium- to coarse-grained amphibolite, displaying the critical assemblage hornblende  +  plagioclase (commonly An30–50). In addition, small amounts of almandine-rich garnet or diopside, biotite and quartz can be present. Epidote is still stable in the low-grade part of the amphibolite facies but breaks down, at higher temperatures, in favour of the anorthite component in plagioclase or of grossular-andradite garnet. In metamorphosed ultramafic rocks, the assemblage hornblende + anthophyllite and/or + cumminctonite is stable.

28

535

28.3 · Mineralogical Characteristics of Individual Metamorphic Facies



kyanite/andalusite Chloritoid [28.8]   2+ 2Fe2 Al9 (SiO4 )4 O6 (OH)2 + SiO2 +2H2 O quartz staurolite

chloritoid + muscovite + quartz ⇋ staurolite + garnet + biotite + H2 O

[28.9]

and

garnet + chlorite + muscovite ⇋ staurolite + biotite + quarz + H2 O

[28.10]

536

28

Chapter 28 · Metamorphic Facies and Facies Series

Judging from various experimental results, the lower stability limit of the assemblage staurolite + garnet + biotite (+muscovite + quartz) is constrained by the T-P points at ca. 515 °C/3 kbar, 540 °C/5 kbar and 560 °C/8 kbar PH2 O. It should be noted, however, that formation of staurolite requires high Al2O3/(K2O + Na2O + CaO) and FeO/MgO ratios in the bulk rock composition. At amphibolite-facies conditions, garnet and staurolite display distinctly higher Fe/Mg ratios than biotite, thus it is actually not permissible to combine FeO and MgO as one component when displaying phase relationships on chemographic diagrams. Consequently, the assemblage staurolite + garnet + muscovite (+quartz) can only be presented by crossing tie lines in the A′KF triangle (. Fig. 28.4c), a disadvantage that can be avoided using the AFM projection (. Fig. 27.27). Depending on pressure, the equilibrium curves of the reactions

kyanite ⇋ sillimanite

[27.2]

andalusite ⇋ sillimanite

[27.3]

or

conforming to the first sillimanite isograd, are crossed within the P-T range of the low-grade amphibolite facies (. Figs. 26.11, 27.8), which means that sillimanite can still coexist with muscovite + quartz. However, sillimanite is not always formed by direct replacement of kyanite or andalusite (e.g., . Fig. 27.7a) but due to more complex reactions in which muscovite and biotite are involved as evidenced by common tight intergrowths of sillimanite and biotite. Moreover, staurolite breaks down according to the reaction

  6Fe2+ 2 Al9 (SiO4 )4 O6 (OH)2 + 11SiO2 ⇋

quartz staurolite 2+ 4Fe3 Al2 [SiO4 ]3 + 23Al2 [SiO5 ] +6H2 O And/Sill/Ky almandine

[28.11]

the equilibrium curve of which is defined by T-PH2 O points at ca. 585 °C/2 kbar and ca. 660 °C/7 kbar, i.e., well below the upper stability limit of muscovite + quartz by reactions [27.11] and [27.11a, b]. Thus the stability field of staurolite, displayed in . Fig. 26.11 is entirely within the P-T range of the lower amphibolite facies. Due to the breakdown of staurolite with increasing temperature, the tie lines andesine-staurolite or muscovite-staurolite in the ACF and A′KF diagrams, respectively, are broken. Thus the compositional fields, in which Al2SiO5 polymorphs can form and grow in direct contact with almandine, are considerably extended (. Fig. 28.4d). In the lower amphibolite facies, the tie line biotite-muscovite in the A′KF diagram divides two different compositional ranges (. Fig. 28.4c): 1. In metapelites, the Al2[SiO5] polymorphs as well as staurolite or almandine cannot coexist with K-feldspar.

2. In orthogneisses such as metagranite or metagranodiorite but also metaarkose, the assemblage quartz + K-feldspar + plagioclase (An20–30) + biotite + muscovite is stable (. Fig. 28.4c). With the breakdown of muscovite in the presence of quartz, according to the dehydration reaction

muscovite + quartz ⇋ andalusite/sillimanite + K-feldspar + H2 O

[27.11]

or the related melting reactions [27.11a, b], the tie line muscovite-biotite is broken in the A′KF diagram. Consequently, K-feldspar can coexist with almandine and/or sillimanite in metapelites, assemblages typical of the upper amphibolite facies (. Fig. 28.4d). Due to the reactions [27.11] or [27.11a, b], Barrow’s second sillimanite isograd is crossed in the field. At a pressure (Ptot = Pfl = PH2 O) of 2 or 5 kbar, the equilibrium temperatures of Reaction [27.11] are ca. 620 and 690 °C, respectively. At PH2 O = 8 kbar, the H2O-saturated melting Reaction [27.11a] takes place at ca. 730 °C, whereas H2O-free dehydration melting occurs at ca. 750 °C at Pload = 8  kbar (. Fig. 27.8). Besides, the more complex continuous or gliding dehydration reaction can take place: muscovite + biotite1 + quartz ⇋ almandine + biotite2 + sillimanite + K-feldspar + H2 O

[28.12]

in which biotite1 has a higher Fe/Mg ratio than biotite2. Due to these reactions, the feldspar/mica ratio increases in metapelites, which, consequently, evolve from a micaschist to a paragneiss. At a medium geothermal gradient, common assemblages are sillimanite  +  almandine-rich garnet  ± biotite + K-feldspar + plagioclase + quartz or ­almandine-rich garnet + biotite + K-feldspar + plagioclase + quartz (. Fig. 28.4d). At a higher geothermal gradient, conforming to the low-pressure facies series, andalusite appears first instead of kyanite and is replaced by sillimanite at higher temperatures (Reaction 27.3). Both Al2SiO5 polymorphs can coexist with muscovite and quartz. After overstepping the equilibrium curve of Reaction [27.11], muscovite + quartz react to form andalusite + K-feldspar or sillimanite + K-feldspar. The equilibrium curves of the reactions [27.11] and [27.3] that cross at a temperature of ca. 610 °C and PH2 O of ca. 2 kbar define four different P-T fields, in which one of the four different assemblages andalusite + muscovite + quartz, sillimanite + muscovite + quartz, andalusite + K-feldspar, and sillimanite + K-feldspar are stable, respectively (. Fig. 27.8). An important Mg–Fe silicate, typical of metapelites of the low-pressure facies series, is cordierite, (Mg,Fe2+)2[Al4Si5O18], which can be formed by the reaction

chlorite + muscovite + quartz ⇋ cordierite + biotite + andalusite/sillimanite + H2 O

[28.13]

537

28.3 · Mineralogical Characteristics of Individual Metamorphic Facies

Besides andalusite or sillimanite, cordierite can frequently occur together with almandine-rich garnet but rarely with staurolite, the stability field of which becomes smaller with decreasing pressure (. Fig. 26.11). At the transition from the lower to the upper amphibolite facies, cordierite is formed by the reaction

biotite + muscovite + quartz ⇋ cordierite + K-feldspar + H2 O

[28.14]

Moreover, the assemblage sillimanite + biotite disappears due to the reaction

biotite + sillimanite + quartz ⇋ cordierite + garnet + K-feldspar + H2 O

[28.15]

This reaction is documented in the well-known sillimanite-free haloes in cordierite, first described from cordierite gneiss of the Bavarian Forest and the Black Forest in the German Variscides. In these rocks, sillimanite needles can still form inclusions in cordierite but are restricted to its central cores thus avoiding a direct contact with biotite flakes in the matrix. 28.3.5  Granulite Facies

Granulite-facies metamorphic rocks are typically observed in lower continental crustal rocks, such as deeply eroded metamorphic basement complexes, many but not all of which are Precambrian in age (7 Sect. 29.2.2). Metabasites in the granulite-facies are present as mafic pyroxene granulite that contains, if formed at low to medium pressures, the critical assemblage plagioclase + orthopyroxene. Typically, Opx is rich in Al, i.e., it contains a high amount of the Mg-Tschermak’s molecule, MgAl[6][Al[4]SiO6], conforming to the coupled substitution MgSi ⇌ Al[6]Al[4]. At lower temperatures and/or somewhat higher H2O-pressures, hornblende can occur in stable assemblage with ortho-/clinopyroxene and plagioclase forming pyribolite. However, prograde metamorphism leads to the final dehydration of hornblende of different composition according to the continuous or gliding reactions

hornblende1 + quartz ⇋ plagioclase + orthopyroxene + clinopyroxene + H2 O

[28.16]

hornblende2 + quartz ⇋ plagioclase + orthopyroxene + garnet + H2 O

[28.17]

or

whereby hornblende2 is richer in Al than hornblende1. In pyriclasite and garnet pyriclasite thus formed, the assemblages plagioclase + orthopyroxene + clinopyroxene ± quartz and plagioclase + orthopyroxene + pyrope-rich garnet ± biotite ± quartz are stable (. Fig. 28.4e).

At increasing pressure, the tie line plagioclase–orthopyroxene in the ACF diagram is broken by the simplified reaction   2(Mg, Fe)2 Si2 O6 + Ca[Al2 Si2 O8 ] ⇋ plagioclase (An)

orthopyroxene

[28.18]   (Fe, Mg)3 Al2 SiO4 3 + Ca(Mg, Fe) Si2 O6 + SiO2 quartz garnet





clinopyroxene

leading to the assemblage garnet + clinopyroxene + plagioclase + quartz whereas, in granulites of distinctly basic composition, garnet + clinopyroxene + orthopyroxene coexist with each other. Consequently, the granulite facies can be divided into a low-pressure and a high-pressure subfacies. At nearly isothermal decompression of high-pressure granulite, Reaction [28.18] is reversed and spectacular corona textures, such as symplectites of orthopyroxene + plagioclase or orthopyroxene + cordierite can develop around garnet, (. Fig. 27.1). Felsic granulites are derived either from clastic sediments such as mudstone, greywacke or arkose, or from felsic igneous rocks such as granite or rhyolite. Critical assemblages are quartz + alkali feldspar + plagioclase + almandine-/pyrope-rich garnet + kyanite/sillimanite or + Al-rich orthopyroxene (. Fig. 28.4e) whereas at lower pressures cordierite can be stable as well. These granulite-facies assemblages are formed by dehydration reactions already observed at the transition from the lower to the upper amphibolite facies, e.g., [27.11], [27.11a, b], [28.11] and [28.12]. At granulite-facies conditions, prograde biotite is formed only in subordinate amounts and, if present, is always Mg-rich, i.e., close to the phlogopite end-member. In contrast to the upper amphibolite facies, Al-rich orthopyroxene is present in many felsic granulites known as felsic pyroxene granulite or charnockite. It should be noted, however, that many charnockite occurrences are formed by crystallisation of granitic magmas at high pressures (7 Sect. 26.3.1). Alkali feldspar typically contains roughly equal proportions of the Ab and Or components which, upon cooling, exsolve to form lamellar mesoperthite (. Fig. 27.7c). Its composition clearly indicates that its crystallisation took place at high temperatures above the solvus maximum in the two-component system albite–Kfeldspar (. Fig. 11.66, 18.12). Typically, (OH)-bearing minerals are rare or absent in granulites indicating that at granulite-facies conditions the H2O activity was low, i.e., PH2O  θ2. It follows

552

Chapter 29 · Earth’s Interior

29 . Fig. 29.4  Propagation of seismic waves through the Earth’s interior. a Simplified two-layer model, with constant propagation velocity in both shells, mantle and core, represented by straight seismic rays (after Kertz 1970, ©Elsevier GmbH, Spektrum Akademischer Verlag, Heidelberg, Germany); b more realistic three-layer model consisting of mantle, outer and inner core, in each of which the physical properties and the propagation velocities of seismic waves change continuously, represented by curvilinear seismic rays for P- and S-waves (blue) or for P-waves alone (black) (modified after Brown and Mussett 1993)

v1 n2 sin θ1 = = sin θ2 v2 n1

(29.6)

To begin with, we shall follow the paths of the seismic rays with the help of a highly simplified model of the Earth’s interior, consisting of a homogeneous mantle and a homogeneous core (. Fig. 29.4a). A fan-shaped bundle of rays crosses the mantle undisturbed and linearly, whereas rays hitting the mantle-core boundary at lower angles are refracted towards the axis of incidence and away from it when re-entering into the mantle. As a result, the seismic rays are concentrated within a “focal spot”. Between these two bundles of rays, a broad gap is left in which no seismic waves have arrived at all. Indeed, this “shadow of the core” is always observed when recording natural seismic waves. Compared to this simplified model, however, the physical properties and, consequently, the velocities of the P- and S-waves vary within the Earth’s mantle and core. Thus the seismic rays follow curvilinear rather than straight paths in the Earth’s interior (. Fig. 29.4b). Like in geometric optics, the principle of Pierre de Fermat (1601–1665) is valid, which states that, among all possible paths, a light or seismic ray always follows that path that involves the shortest travel time.

Considering a more realistic model with a mantle, an outer and an inner core, in which the seismic rays follow curvilinear paths (. Fig. 29.4b), three different sectors can be distinguished analogous to the simplified model above: 5 0–103° off the hypocentre: The seismic stations, such as A, B and C in . Fig. 29.4b, record P-waves as well as S-waves, some of them having been reflected at the core-mantle boundary, e.g., station D′. 5 103–143° off the hypocentre: Shadow zone, in which virtually no seismic waves are recorded, thus indicating a discontinuity surface at a depth of 2900 km, dividing the Earth’s mantle from the outer core. A slight “brightening” of the shadow can be explained by the following facts: 5 By boundary surface waves that run along the discontinuity; 5 By P-waves that are reflected at a second, inner discontinuity surface at a depth of 5080 m, dividing the outer from the inner core; 5 By an antipodal “brightening” of the shadow of the core that could be due to an exceptional increase of vP (see below). 5 >143° off the hypocentre: At the “focal spot” seismic waves are recorded again, however exclusively P-waves, no S-waves, e.g., at the stations D, E, F, C′, G. This fact leads to the conclusion that below the discontinuity surface at a depth of 2900 km, a material of liquid state with µ = 0 is present that cannot undergo elastic deformation and, according to Eq. [29.4], is not permeable for S-waves, as vS = 0. According to the model displayed in . Fig. 29.4b, two discontinuity surfaces exist within the Earth’s interior. The upper one delimits, at a depth of 2900 km, a solid outer shell, the Earth’s mantle, from a middle liquid shell, the Earth’s outer core, which in turn is separated from the solid inner core of the Earth by a discontinuity at a depth of 5080 km. 29.1.3  Velocity Distribution of Seismic

Waves in Earth’s Interior

Our picture of the Earth’s interior can be considerably improved by plotting the velocity of the P- and S-waves against depth (. Fig. 29.5a, b). As vP and vS as well as the gradients dvP/dz and dvS/dz are directly dependent on the ratios K/ρ and µ/ρ, abrupt or gradual changes of the gradients must be due to corresponding changes of these physical constants, a fact displayed for the density ρ in . Fig. 29.6. This leads to the following subdivision of the Earth’s interior: Earth’s crust, upper mantle, transition zone, lower mantle, outer core and inner core (. Fig. 29.5). Although initially the layered structure of the Earth has been recognised by seismic methods, it can be derived also from mineralogical and chemical data, thus forming the basis for the understanding of the most basic geological processes. This holds

553

29.2 · The Crust

true although direct observations on rocks become more and more scarce towards greater depths, and are impossible in the Earth’s lower mantle and core. It is important to note that a lot of sophisticated geophysical measurements and calculations are necessary to model the changes of vP, vS, K, µ and ρ with depth. Seismograms contain complex information on different types of seismic waves that interfere and can intensify or cancel out each other. Their analysis is an approach similar to “the deciphering of a coded language” (Allègre 1992). The apparently simple picture that is shown in . Fig. 29.5 and based on seismic modelling is one of the paramount achievements of geophysics. Nevertheless, well-founded theories on the mineralogy of the Earth’s interior had been developed only by connecting geophysical findings with direct observations on rocks of the Earth’s mantle, with results of high-pressure experiments and of thermodynamic modelling (cf. Saxena 2010). 29.2  The Crust

. Fig. 29.5  a Layered structure of the Earth, based on seismic evidence (after Ringwood 1979); b variation of P- and S-wave velocities with depths in the Earth’s mantle and core (after Hart et al. 1977)

The Earth’s crust is separated from the underlying mantle by the Mohorovičić Discontinuity, discovered in 1910 by the Croatian geophysicist Andreiji Mohorovičić (1857–1936). This surface, colloquially referred to as Moho, is characterised by a relatively abrupt increase of the P-wave velocity vP from ca. 6.5–7.0 to ca. 8.0–8.3 km s−1 and a corresponding increase in vS (. Fig. 29.5b). The Moho is located at depths of 5–7 km below the ocean floor, but ca. 30–40 km, locally even up to ca. 60 km, under the continents, and up to 90 km below young orogenic belts. Accordingly, the thickness of the Earth’s crust varies depending on the geological setting of a specific region. It should be noted, however, that the Moho is not an absolutely sharp boundary but a gradational transition. This transition zone can reach 1–2 km in thickness below the sea floor but more below the continents. In many orogenic belts, the Moho is poorly developed or may be doubled whereas, below most mid-ocean ridges, it is totally absent. 29.2.1  Oceanic Crust

. Fig. 29.6  Increase of density with depth in the Earth’s interior (after Clark and Ringwood 1964)

Information on structure and composition of the oceanic crust have been gained from seismic measurements and submarine drilling, performed in the framework of the international Deep-Sea Drilling and Ocean Drilling programs (DSDP and ODP), mainly by the U.S. research vessels Glomar Challenger and Joides Resolution. Additional important information can be derived from ophiolite complexes. These large fragments of oceanic lithosphere that form tectonic nappes within orogenic belts can contain the complete rock inventory of the oceanic crust and the underlying mantle of the Earth, or at least parts of it. As compared to boreholes, ophiolite complexes display three-dimensional outcrops and

29

554

Chapter 29 · Earth’s Interior

can disclose deeper levels not penetrated by boreholes so far. Well-studied examples are the Vourinos Complex in northern Greece, the Troodos Complex in Cyprus and numerous further occurrences in the Dinarides and Hellenides, the Semail Complex in Oman, several occurrences in Indonesia and Papua New Guinea as well as the Bay-of-Islands Complex in New Foundland, Canada. The Semail Complex, the largest and best exposed ophiolite complex in the world, extending for more than 500 km and reaching a total thickness of up to 20 km (Lippard et al. 1986), was obducted onto the northern margin of the Arabian Plate during the Late Cretaceous. Within a short period of 96–94 Ma, basalts and gabbros were emplaced during at least two different phases of igneous activity, the first one at a fast-spreading mid-ocean ridge, the second one above a subduction zone.

29

Combining all information available, it has been concluded that the oceanic crust displays a layered structure, in which rocks of basaltic composition, high in Si and Mg, are predominant. Thus it has been designated with the acronym Sima by Alfred Wegener (1880–1930), the father of the continental drift theory. The thickness of individual layers and the P-wave velocities displayed in the schematic profile . Fig. 29.7 are approximate values that can vary regionally. 5 Layer 1: deep-sea sediments: their thickness continuously increases from zero at the mid-ocean ridges towards the continents, and can reach several kilometres on the slopes of passive continental margins; the oldest deep-sea sediments still found in situ were deposited in uppermost Triassic times; 5 Layer 2: submarine pillow lava of basaltic composition (MORB); 5 Layer 3a: sheeted-dyke complex: Basaltic dykes of MORB composition intruded each other along vertically opening fissures; . Fig. 29.7  Schematic profile of the oceanic lithosphere, consisting of the Earth’s oceanic crust and uppermost mantle, based on seismic measurements, deep-sea drilling, and observations in typical ophiolite complexes. The thickness of individual layers and the seismic velocities can vary considerably from region to region (after Brown and Mussett 1993)

5 Layer 3b: gabbro, crystallised within oceanic basaltic magma chambers; 5 Layer 4a: peridotite with cumulate structure, formed by gravitational subsidence of early crystallised dense olivine and pyroxene crystals in the magma chamber; seismic measurements “see” the contact between gabbro and peridotite as apparent crust/mantle border, the seismic Moho; 5 Layer 4b: harzburgite and peridotite of the Earth’s upper mantle, the border of which cannot be recorded by seismic methods and is thus distinguished as petrographic Moho. Mid-ocean ridges represent separate, important petrological provinces in which submarine volcanic activity permanently creates new oceanic crust. Accordingly, these divergent plate margins are constructive in character. Typical is a significant negative Bougier anomaly, i.e., a gravity anomaly that has been corrected for topography, in this case for the low density of seawater. The deficit in gravity is highest above the axial zone and decreases towards the margin of the ridge. Non-consolidated deep-sea sediments are virtually absent whereas pillow basalts of Layer 2 crop out at the sea floor. Layer 3 gradually develops into a mantle of anomalously low P-wave velocities, commonly 7.1–7.3 km s−1. The Moho is poorly developed and can be absent in many regions. The heat flow is high and volcanic activity with effusion of basaltic lava is common. All these facts indicate that the mid-ocean ridges are zones of uprising convection currents, produced by partial melting in the Earth’s mantle (. Fig. 29.17). As an essential part of oceanic lithospheric plates, newly formed oceanic crust moves away from the mid-ocean ridges with velocities of several centimetres per year, a process known as sea floor spreading. At convergent, destructive plate margins, oceanic plates are subducted below less dense, younger oceanic lithospheric plates or continental lithospheric plates, leading to the formation of island arcs and oro-

555

29.2 · The Crust

genic belts of Andean type, respectively (. Figs. 29.17, 29.18). At passive continental margins, the oceanic crust is thicker than usual and starts beyond the continental shelf, i.e., the section of the continental margin inundated by the sea. Oceanic crust of unusual thickness, reaching as much as 20 km, is developed in the vicinity of ocean islands, such as Hawaii, and in large oceanic flood basalt plateaus (7 Sect. 14.1). 29.2.2  Continental Crust

Already early on, the detailed analysis of the velocity of seismic waves suggested that the Earth’s continental crust should also be composed of various layers. These display, however, a much more complex structure and are more variable in composition than in oceanic crust (. Fig. 29.8). z Unconsolidated Sediments and Consolidated Sedimentary Rocks

The sedimentary cover contains sequences of very different thickness and geological age but can also be totally absent, apart from a thin layer of soil. z Upper Continental Crust

Information on the structure and composition of the upper continental crust have been obtained from fieldwork in many magmatic and metamorphic complexes in various regions of the Earth, especially in deeply eroded

continental cratons of Archaean and Proterozoic age such as Fennoscandia, Laurentia and others. According to these studies, dating back to the end of the 19th century, the upper continental crust consists predominantly of metapelites, quartzo-feldpatic gneisses and migmatites as well as granite, granodiorite and tonalite. The average density of these rocks is ca. 2.7 g cm−3, conforming to a P-wave velocity of ca. 6.0 km s−1. Chemically, most rocks of the upper continental crust are characterised by the predominance of Si and Al, which prompted Alfred Wegener to classify this crustal type by the acronym Sial. Volcanic rocks of basaltic composition and higher density are subordinate, apart from the large regions covered by continental flood basalt (7 Sect. 14.1). A crustal low velocity zone, within which vP decreases to 45 kbar, corresponding to a minimum depth of about 140 km, whereas under the oceans the minimum pressure of ca. 55 kbar is reached at a depth of ca. 185 km (. Fig. 29.15). Green and Ringwood (1963) assumed that the Earth’s whole mantle had a chemical composition that results from a mixture of 3 parts dunite and 1 part basalt and named the

29

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Chapter 29 · Earth’s Interior

model rock of this composition pyrolite. By partial melting, not more than 25% of basaltic magma can be derived from such a rock. (7 Sect. 19.2). Later on, the pyrolite model was somewhat refined, e.g., considering the average composition of ultramafic members of ophiolite complexes (Ringwood 1975). . Table 29.2 clearly reveals the close chemical similarity between theoretical model pyrolite and natural garnet lherzolite, although the latter is still somewhat richer in MgO and slightly depleted in the basaltic components Al2O3, FeOtot, CaO and Na2O. Suitably, the pyrolite model conforms to the assumption that the whole Earth has a bulk composition similar to the average composition of chondrite, the most important group of meteorites (7 Sect. 31.3.1). Fundamental experimental investigations by Green and Ringwood (1967b) have revealed that, at different P-T conditions within the Earth’s interior, pyrolite crystallises in different mineral assemblages (. Fig. 29.14). These are essentially distinguished by the mineral that contains the Al2O3 component as detailed below:

29

. Fig. 29.14  Pressure-temperature (P-T) diagram showing the stability fields of plagioclase pyrolite, pyroxene pyrolite, spinel pyrolite and garnet pyrolite, experimentally determined by Green and Ringwood (1967b). The Al2O3 contents in orthopyroxene coexisting with garnet are shown as stippled lines. Solid line: dry pyrolite solidus; dotted line: solidus of pyrolite containing 0.1 wt% H2O. Graphite/diamond equilibrium curve after Bundy et al. (1961) and Berman (1962); oceanic and continental geotherm after Clark and Ringwood (1964)

1. Plagioclase pyrolite consists of the assemblage olivine » orthopyroxene > clinopyroxene >  plagioclase, which is stable only at low pressures of less than ca. 5 kbar, at temperatures defined by the normal continental or oceanic geotherm (. Fig. 29.14). Consequently, this plagioclase-bearing ultramafic rock cannot occur within the subcontinental Earth’s mantle where the Moho is commonly located at depths corresponding to pressures of >10 kbar. Thus plagioclase pyrolite can be expected at best close to mid-ocean ridges characterised by an anomalously high geothermal gradient. At higher pressures, the Al-content of plagioclase will be incorporated in pyroxene as Ca-Tschermak’s molecule (CaTs) or, to a lesser extent, as jadeite component, according to the model reactions

  + Mg2 [SiO4 ] Ca Al[4] Si O 2 8 2   ⇋ CaAl[6] Al[4] SiO6 + Mg2 [Si2 O6 ]

[29.2]

and   Na Al[4] Si3 O8 + Mg2 [SiO4 ]

⇋ NaAl[6] [Si2 O6 ] + Mg2 [Si2 O6 ]

[29.3]

Alternatively, spinel can be produced as an Al phase of its own, e.g., after the model reaction

  Ca Al[4] 2 Si2 O8 + 2Mg2 [SiO4 ] . Fig. 29.15  Schematic cross section through the lithosphere (Earth’s crust + uppermost mantle), the asthenosphere, and deeper parts of the Earth’s mantle, also showing the graphite-diamond reaction curve (modified after Stachel and Brey 2001)

⇋ MgAl[6] Al[4] O4 + CaMg[Si2 O6 ] + Mg2 [Si2 O6 ]

leading to the formation of spinel pyrolite.

[29.4]

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29.3 · The Mantle

. Table 29.2  Theoretically derived chemical composition of model pyrolite as compared with natural garnet lherzolite (after Green and Ringwood 1963; Ringwood 1975; Brown and Mussett 1993) Oxide (wt%)

Average dunite

Average basalt

3 dunite + 1 basalt

Pyrolite Ringwood 1975

Garnet- lherzolite

SiO2

41.3

50.8

43.7

45.1

45.3

Al2O3

0.54

14.1

3.9

4.6

3.6

FeOtot

7.0

11.7

8.2

8.4

7.3

MgO

49.8

6.3

39.0

38.1

41.3

CaO

0.01

10.4

2.6

3.1

1.9

Na2O

0.01

2.2

0.6

0.4

0.2

K 2O

0.01

0.8

0.2

0.02

0.1

Total

98.67

96.3

98.2

2. Spinel pyrolite is made up of the assemblage olivine » orthopyroxene > clinopyroxene > spinel, the stability field of which extends up to maximum pressures of ca. 14–18 kbar at temperatures corresponding to the continental or oceanic geotherm (. Fig. 29.14). Upon isobaric temperature increase, the pyroxenes can incorporate more and more Al in their crystal structure, eventually leading to the disappearance of the independent Al-mineral spinel and formation of pyroxene pyrolite. 3. Pyroxene pyrolite comprises the assemblage olivine » Al-rich orthopyroxene > Al-rich clinopyroxene that is stable at very high temperatures in excess of 1240–1300 °C and pressures ranging from 10 to 30 kbar (. Fig. 29.14). Consequently, pyroxene pyrolite can be expected at an extremely high geothermal gradient such as realised at midocean ridges. At rising pressures and/or decreasing temperatures, the pyroxenes can incorporate less and less Al in terms of the Ca− or Mg-Tschermak’s molecule (CaTs, MgTs) and, alternatively, pyrope-rich garnet is formed. Moreover, the reaction of spinel with orthopyroxene can lead to the assemblage pyrope-rich garnet + olivine according to the equation

MgAl[6] Al[4] O4 + 2Mg2 [Si2 O6 ] ⇋ Mg3 Al2[6] [SiO4 ]3 + Mg2 [SiO4 ]

99.7

99.7

Switzerland, La Charme, Vosges, France, and Åheim in Norway. As shown in 7 Sect. 19.2, basaltic magma is formed by partial melting of fertile pyrolite, leaving behind depleted lherzolite, harzburgite or dunite as restite (e.g., Green and Ringwood 1967c; Jaques and Green 1980). This melting leads not only to mineralogical but also chemical heterogeneities within the Earth’s lithospheric mantle. Due to the formation and eruption of basaltic magma, specific mantle areas become depleted in K, Na, Ca, Al and Si as well as in the incompatible trace elements, such as Be, Nb, Ta, Sn, Th, U, Pb, Cs, Li, Rb, Sr and REE, whereas Mg becomes concentrated. As a result, large parts of the lithospheric mantle no longer consist of fertile pyrolite but of depleted peridotite, especially of harzburgite and lherzolite (. Figs. 29.16, 29.17). The proportion of such depleted peridotite is higher in the

[29.5]

4. Garnet pyrolite, the most common rock type of the Earth’s lithospheric mantle, contains the assemblage olivine (ca. 57%) + orthopyroxene (ca. 17%) + clinopyroxene (ca. 12%) + garnet (ca. 14%). As shown in . Fig. 29.14, the stability field of garnet pyrolite can be reached at depths of ca. 45 km below the continents and at considerably higher depths of ca. 60 km below the ocean floor. As mentioned above, natural garnet lherzolite, having the same mineral assemblage and similar bulk rock composition to theoretical garnet pyrolite (. Table 29.2), frequently occurs as xenoliths in kimberlite and lamproite. Moreover, it can form tectonic slices within orogenic belts. Well-known examples are the Alpe Arami near Bellinzona, Ticino,

. Fig. 29.16  Schematic presentation of chemical inhomogeneities within the Earth’s upper mantle below the oceanic (left) and continental crust of the Earth (right). Harzburgite: olivine + orthopyroxene + chromite; Lherzolite: olivine + clinopyroxene + orthopyroxene + spinel

29

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Chapter 29 · Earth’s Interior

. Fig. 29.17  Petrological model of plate tectonics (modified after Ringwood 1979). Detailed explanation is given in the text

subcontinental than in the suboceanic mantle of the Earth, as the latter is transported, by sea-floor spreading, into subduction zones and further into the deep mantle: There is no oceanic lithosphere that is older than ca. 200 Ma. In contrast, much more time is available in the continental lithosphere to transform fertile into depleted pyrolite, by partial melting and eruption of the basaltic magma thus formed. From the P-T diagram in . Fig. 29.14 it is evident that the solidus curve of H2O-free pyrolite is intersected neither by the continental nor the oceanic geotherm. Thus partial melting of “dry” pyrolite requires exceptionally high temperatures. However, the situation is drastically changed if the Earth’s upper mantle contains a small amount of H2O. For instance, compared to the dry pyrolite solidus, the solidus curve of a pyrolite containing only 0.1 wt% H2O is displaced to markedly lower temperatures and follows particularly low temperatures at pressures between 25 and 50 kbar where it is intersected by the oceanic geotherm. Consequently, partial melting can occur in this P-T range where 0.5–1 wt% of melt can be formed (. Fig. 19.2). Not coincidentally, the low-velocity zone within the Earth’s upper mantle is situated at about the same depth. 29.3.2  The Asthenosphere as Conveyor Belt

of Lithospheric Plates

As early as 1926, Beno Gutenberg found out that over a depth range of ca. 60–250 km, the P- and S-wave velocities are reduced by about 3–6%, an effect that is more pronounced for vS than for vP (. Fig. 29.5b). As the S-waves are still transmitted in this part of the Earth’s upper mantle, known as the low-velocity zone (LVZ), it must be composed essentially of solid material. In different tectonic environments, the LVZ is situated at different depths and, moreover, is developed much better under the oceans than under

the continents, where it can be totally absent. This fact is explained by the P-T diagram (. Fig. 29.14) that shows that, at depths of 100–170 km below the oceans, a pyrolite with 0.1% H2O can form 0.5–1 wt% of partial melt, thus reducing the shear modulus µ and, consequently, vP and vS (Eqs. (29.3) and (29.4)). In contrast, the pyrolite solidus is not intersected by an average continental geotherm. Thus partial melting of the continental mantle demands either a higher H2O content or an uncommonly high geothermal gradient. The asthenosphere (Greek ασθενὁς weak) is defined as a zone, in which the Earth’s upper mantle behaves relatively mobile, i.e., material can flow both vertically (due to isostasy, for example) and horizontally (e.g., seafloor spreading).

At first sight, the asthenosphere might be regarded as the same as the low velocity zone because both are situated at similar depths and related to the same reason, that is reaching the temperature conditions of partial melting. This is, however, an oversimplification. The asthenosphere plays a crucial role in the theory of plate tectonics, according to which rigid lithospheric plates move on top of the relatively ductile asthenosphere. Essential aspects of plate tectonics are schematically illustrated in . Fig. 29.17: 5 Divergent (constructive) plate margin between two oceanic lithospheric plates: These are composed of the Earth’s oceanic crust and lithospheric mantle, consisting of depleted peridotite and underlying fertile pyrolite. New oceanic crust is generated from ascending basaltic magma derived from the underlying asthenosphere and consolidated to form tholeiitic mid-ocean ridge basalt (MORB).

563

29.3 · The Mantle

5 Mantle diapirs, rising from the asthenosphere, supply oceanic intraplate volcanoes with magma of alkali-basaltic but also of tholeiitic composition, a prominent example being the Hawaii Islands on the Pacific Plate. 5 Convergent (destructive) plate margin: An oceanic lithospheric plate is subducted below another lithospheric plate, be it another oceanic or a continental plate, the latter consisting of a thick crust and an underlying mantle of depleted peridotite. Four different types of convergent plate margins can be distinguished (e.g., Frisch et al. 2011; see . Fig. 29.18a–d; . Fig. 14.1): a. By subduction of one oceanic lithospheric plate below another one, an intra-oceanic, ensimatic island arc is formed on oceanic crust of basaltic composition. Examples are the Mariana Islands at the eastern border of the Philippine Plate or the Lesser Antilles near the eastern border of the Caribbean Plate. b. Subduction of an oceanic below a continental lithospheric plate, separated by a back-arc basin from the continental hinterland, leads to the formation of an ensialic island arc on sialic continental crust. Examples are the Islands of Japan at the eastern margin of the Eurasian Plate. c. At an active continental margin, oceanic lithosphere is subducted below a continental lithospheric plate and is marked by a continental magmatic arc in the supra-subduction position without an interposed back-arc basin. Examples are the Cascade Range of North America as well as the Central and South American Andes, along the western margins of the North and South American plates. d. Complete subduction of an oceanic lithospheric plate eventually leads to continent-continent collision and formation of orogenic fold belts, such as the Alps or the Himalayas. The oceanic part of the subducted plate is teared off and vanishes into the Earth’s mantle, a process known as slab breakoff. The topographic expression of subduction zones are deep oceanic trenches, e.g., the Tonga Trench and the Philippine Trench in the western Pacific with maximum depths of 10,882 m and 10,540 m, respectively. In seismic profiles, subduction zones are registered as Benioff zones (Benioff 1955), in which the foci of seismic events are concentrated along a plane, dipping towards the continent at angles ranging from 15 to 85°, in most cases around 45°, and traceable to depths as deep as 700 km. The force diagram within Benioff zones can be calculated from so-called fault-plane solutions that yield compressive stress in direction of the subducted plate, and tensile stress perpendicular to it. However, directly below the deep ocean trench, tensile stress is recorded in direction of the plate, compressive stress perpendicular to it.

During subduction, the H2O-bearing oceanic crust undergoes high-pressure metamorphism leading to the formation of blueschist and eclogite. The volatiles thus released

migrate into the overlying mantle wedge where they lower the melting temperature and trigger anatexis resulting in the production and eruption of calc-alkaline magma forming magmatic (volcanic) island arcs or magmatic orogenic arcs of Andean type (. Fig. 29.18). Typical is the volcanic sequence tholeiitic basalt → andesite → dacite → rhyodacite  → rhyolite as well as plutonic rocks of I-type character, especially granite, granodiorite, tonalite and trondhjemite, known as TTG suite. The igneous activity is associated with regional metamorphism of low- to medium-pressure type (7 Sect. 26.2.5, . Fig. 28.2). Below the asthenosphere, vP and vS increase again, and the Earth’s mantle attains a density (reduced to 1 bar) of 3.3–3.4 g cm−3, conforming to garnet pyrolite (. Fig. 29.19). 29.3.3  The Transitional Zone Between Upper

and Lower Mantle

Seismically, the transitional zone is defined as a mantle domain between depths of about 410 to 660 km (e.g., Frost 2008), in which the P- and S-wave velocities are highly variable but, on average, have supernormal gradients dvP/dz and dvS/dz. During the 1930s, Birch and Presnal were the first to explain these discontinuities by high-pressure transformation of silicate minerals into denser modifications with more tightly-packed crystal structures (Birch 1952). Analogous examples described in previous chapters are the transformation of the relatively loose sheet structure of graphite into the diamond structure, characterised by cubic close packing (. Figs. 4.11, 4.15) or the coesite, Si[4]O2, → stishovite, Si[6]O2, transformation involving the conversion of a framework silicate into the rutile-type structure (. Figs. 7.9, 11.43). Conventional hydrothermal autoclaves and high-pressure apparatus such as piston-cylinders, belt or multi-anvil devices are not sufficient for direct experimental determination of the phase transitions in the lower part of the Earth’s mantle. Thus, experiments with germanates, structurally analogous to silicates, were performed as the change in the coordination Ge[4] → Ge[6] takes place at much lower pressures than that of Si[4] → Si[6]. However, the relatively recent invention of the diamond-anvil cell makes it possible to perform experiments directly on silicates at pressures of as high as 2 Mbars (=2000 kbar = 200  GPa) and temperatures of up to 5000 °C, conforming to P-T conditions down to the Earth’s core (e.g., Boehler 2000). The transition zone recorded at a depth of about 400 km is the 400 km or 20° discontinuity. Compared to the Moho, it represents a much broader zone that is several tens of kilometres thick. For the phase transitions responsible for the increase of vP and vS at these depths, a series of model reactions have been formulated (Ringwood 1975, 1991; . Fig. 29.19). At depths of about 400 km and temperatures around 1400 °C, olivine starts to transform into phases of the same chemical composition but of higher density (. Figs. 29.19, 29.20):

29

564

Chapter 29 · Earth’s Interior

. Fig. 29.18  Plate margins with different types of subduction zones; For explanation see text (after Frisch et al. 2011)

29

29

565

29.3 · The Mantle

α-(Mg, Fe)2 [SiO4 ] → β-(Mg, Fe)2 [SiO4 ] → γ -(Mg, Fe)2 [SiO4 ] oliviness

wadsleyitess

ringwooditess

[29.6a,b]

. Fig. 29.19  Possible mineral assemblages and rock densities (expressed for a pressure of 1 bar) in a modeled Earth’s mantle of pyrolite composition, down to a depth of ca. 850 km, based on high-pressure experiments. Temperature values derived from the geothermal gradient of Brown and Shankland (1981) (after Ringwood 1991)

In the olivine structure, Mg2+ and Fe2+ are placed within the octahedral sites of a nearly hexagonal dense sphere packing of oxygen and thus are [6]-coordinated, whereas Si[4] occupies the smaller tetrahedral sites (. Fig. 11.3). As pressure increases, olivine is transformed into the more tightly packed β-(Mg, Fe)2 [SiO4 ] phase wadsleyitess, thus undergoing an increase in density by ca. 8%. Because of the Fe content in olivine, Reaction [29.6a] is not univariant but represents a divariant equilibrium, which means that, over a pressure range of roughly 130–150 kbar, Mg-dominant oliviness coexists with wadsleyitess slightly richer in Fe. By analogy, the gliding transformation of wadsleyitess into ringwooditess, the γ -(Mg, Fe)2 [SiO4 ] phase with spinel structure (. Fig. 7.2), takes place over a pressure interval of about 130–200 kbar (. Fig. 29.20), thus leading to an additional density increase by ca. 2%. Starting already at depths of about 300 km, ortho- and clinopyroxene become replaced by garnet. This replacement is completed at depths of about 460 km (at ca. 1500 °C) and can be described by the following model reactions:

Mg2 [Si2 O6 ] + MgAl[AlSiO6 ] → Mg3 Al2 [SiO4 ]3 Enstatite

MgTs

pyrope

  2Mg2 [Si2 O6 ] → Mg3 MgSi[6] Si[4] O4 Enstatite

. Fig. 29.20  Two-component system Mg2SiO4–Fe2SiO4 at T = 1600 °C in a pressure range between 40 and 220 kbar (4–22 GPa); Pure fayalite, α-Fe2[SiO4], is transformed into the γ-phase ringwooditess with spinel structure at c. 68 kbar, whereas pure forsterite, α-Mg2[SiO4], reacts to form the β-phase wadsleyitess at c. 148 kbar which, in turn, transforms into the γ-phase ringwooditess at c. 200 kbar. In (Mg,Fe)2[SiO4] solid solutions, the phase transitions α → β and β → γ are effected via divariant fields (after Akaogi et al. 1989)

3

[29.7]

[29.8]

garnet (majorite)

Thus, pure pyrope is formed from Mg-Tschermak’s molecule together with an equivalent amount of enstatite, whereas enstatite alone is replaced by majorite, a pyrope-rich garnet in which the Al[6] position is replaced by Mg[6] + Si[6], i.e., by about 25% of the Si present. The breakdown of CaTs-bearing orthopyroxene or of diopsidic clinopyroxene contributes to the grossular component in garnet. At depths of about 600 km, a pyrolitic Earth’s mantle would consist of ca. 60 vol% ringwoodite + 40 vol% majorite garnet (. Fig. 29.19). Natural majorite was detected by Moore and Gurney (1985) as inclusion in diamond, providing sound proof for its derivation from the transitional zone (cf. Wood et al. 2013). There is geophysical evidence that at depths of about 650 km, i.e., at the lower boundary of the transitional zone, remnants of a subducted oceanic plate appear horizontally deflected, which would signal the end of density-driven sinking of a subducted slab. Seismic tomographic studies have indicated, however, that lithospheric plates can be subducted also to much greater depths, possibly as far down as the core/mantle boundary (Grand et al. 1997; Hirose and Lay 2008).

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Chapter 29 · Earth’s Interior

. Fig. 29.21  Model for subduction of a cool, thick plate of differentiated oceanic lithosphere leading to formation of the garnetite layer overlying the 650 km discontinuity. A “megalite” body consists of a mixture of former mantle harzburgite plus crustal basalt. For explanation see text (after Ringwood 1991)

29

The subsiding lithosphere undergoes phase transformations in which, in contrast to the surrounding mantle, (OH)-bearing minerals are involved. Thus at depths of ca. 550 km, a former harzburgite consists of ca. 90 vol% (OH)-bearing ringwoodite plus subordinate amounts of (OH)-bearing majorite and stishovite, whereas eclogite of the oceanic crust is transformed into garnetite composed of ca. 90 vol% majorite + 10 vol% stishovite. As garnetite has a slightly lower density than the lower mantle, it should be concentrated to form a layer at a depth of ca. 650 km. In contrast, mixtures of former harzburgite and garnetite can form “megalites”, i.e., relatively cool, viscous bodies that extend down into the lower mantle, in which they are gradually dissolved due to convection processes (Ringwood 1991; . Fig. 29.21). Moreover, during transport to greater depths, partial melting of the oceanic crust and chemical exchange with the depleted lherzolite and harzburgite leads to refertilisation of the subducted mantle rocks. This refertilised peridotite can be accumulated above the garnetite layer, thus forming a possible source for mantle plumes and hot spots that trigger intra-plate volcanism (e.g., Pirajno 2004). z Water within the Earth’s Mantle

Via the subduction of lithospheric plates that contain (OH)- and H2O-bearing minerals, water can be transported deep into the Earth’s mantle (e.g., Ohtani 2005).

These minerals are hardly stable at P-T conditions that are common in wide parts of the mantle but can survive within the relatively cool subducted plate, in which an abnormally low geothermal gradient is realised. The first (OH)-bearing phases in subsiding mantle-peridotite are minerals of the chlorite and serpentine groups (Sects. 11.5.5, 11.5.6). These are transformed, step by step, into the so-called 10-Å phase, Mg3[Si4O14]H6, a sheet silicate with a distance of c0 = 10 Å between the [Si4O10] layers. At higher pressures, dense hydrous magnesium silicate (DHMS) phases are formed, first synthesised by Ringwood and Major (1967). In the Earth’s upper mantle, the DHMS phases A, Mg7[Si2O8](OH)6, and E, Mg2.3[Si1.25O6]H2.4, are stable whereas, in the transitional zone and in the upper part of the Earth’s lower mantle, the DHMS phases C, Mg10[Si3O14](OH)4, and D, Mg1.14[Si1.73O6]H2.81, become subsequently stable in subducted mantle peridotite. However, as wadsleyite and ringwoodite can incorporate up to 3 and 1.0–2.2 wt% H2O, respectively, these theoretically H2O-free high-P phases are indeed the most important H2O minerals in the transitional zone, leading to an overall storage capacity of 0.5–1 wt% H2O. In contrast, the Earth’s lower mantle can accommodate merely