Geochemistry of hydrothermal ore deposits [3 ed.] 047157144X

Citation preview

GEOCHEMISTRY OF HYDR OT HER MA LOR E DE P.O SITS Third Edition

Edited by Hubert Lloyd Barnes Ore Depo sits Research Section Th e Pennsylvania State University University Park , Pennsylvania

DEDALUS - Acervo - IGC

1111 1111111111111111111111

30900006945

New York

•

Chichester

•

Weinheim •

Brisbane •

John Wiley & Sons, Inc. Singapore • Toronto

....---------

This text is printed on acid-free paper. Copyright © 1997 by John Wilcy & Sons, Inc.

All rights reserved, Published simultaneously in Canada. Reproduction or translation of any part of this work beyond that permitted by Section 107 or 108 of the 1976 United States Copyright Act without the permission of the copyright owner is unlawful. Requests for permission or further information should be addressed to the Permissions Department, John Wiley & Sons, Inc., 605 Third Avenue, New York; NY 10158-0012

'"

This publication is designed to provide accurate and authoritative information in regard to the subject matter covered. It is sold with the understanding that the publisher is not engaged in rendering legal, accounting, or other profcssional services. If legal advice or other expert assistance is required, the services of a competent professional person should be sought.

Library of Congress Cataloging ill Publication Data: Geochemistry of hydrothermal ore deposits. - 3rd ed. / edited by Hubert Lloyd Barnes. p. cm. Includes bibliographical references and index. ISBN 0-471-57144-X (cloth: alk. paper) 1. Hydrothermal deposits. 2. Geochemistry. I. Barnes, Hubert Lloyd. QE390.5.G43 1997 553- 0.5, arm = 0.25aim exp{(6 .52 - 2667/T)[ln (1 - X~~) + X~ + 0.193]}

(3.10)

where arm and aim refer to the activities of component i in hydrous and anhydrous equivalent melts, respectively, and T is in kelvins (Burnham, 1981 a). Equations (3.9) and (3.10) express the effects of dissolved H20 on the equilibrium properties of aluminosilicate melts and it is now useful to incorporate them into a more comprehensive thermodynamic model of crystal-melt equilibrium in hydrous magmas. Equation (3.3), when rewritten in terms of component i in both melt, m, and crystalline, s, phases and the expression for (s) subtracted from that for (m) yields the fundamental cryoscopic equation of crystal-melt equilibrium:

(3.11 ) where ail is the activity of component i in the melt, whether hydrous (hm) or anhydrous (am); ai is the activity of component i in the coexisting solid (ai = 1.0 for solid of pure i composition); ..1G~li is the l-bar, T (K) standard state Gibbs free energy of fusion (melting) of pure crystalline i; II ..1 Vmi dP,

72

Magmas and Hydrothermal Fluids

which is the integral of the vo lume change upon melting with respect to pressure (P in bars), carries the standard state from 1 bar and T up to the p ressure of interest; R is the gas constant, and T is temperature in kelvins. Equation (3.11) is arranged so that the properties of pure substa nce i appear on the right side and the partial molal properties of i in so lution, whether melt (m) or solid (s), appea r on the left side. Thus the terms on the right side yield unique values at a given T and P, as shown in Figures 3.4 and 3.5, whereas those on the left side yield values that are also dependent on solution composition and the mixing behavior of i in so lution, either (m) or (s), or both. The cryoscopic relations presented in Figures 3.4 and 3.5 for the feldspars and quartz have been calcu lated from the corresponding cryoscopic equations presented in Burnham and Nekvasil (1986), whereas those for diopside have been calculated from equations der ived using the experimenta l data of Eggler and Rosenhauer (1978) and Huebner and Turnock (1980). T he cryoscopic (freezing point) curves in Figure 3.4 for a lbite and anorthite pass through zero at their congruent melting temperatures at atmospheric pressure, whereas the curves for sanidine and q uartz pass through zero at the ir metastable melting points of 1513 K and 1744 K, respectively. The cryoscopic curve for diopside passes through zero at 1664 K, despite the apparent very slight incongruency at this temperature (Kushiro, 1972) that has been ignored .here. The metastability in the melting temperature of quartz is due to lower-temperature phase transitions that make cristobalite the stable crystalline phase of silica at 1744 K and 1 bar, but not at the temperatures and pressures of interest here. The metastability in the melting temperature of sanidine, on the other hand, is of present interest and is due to the incon-

8 7

6

...a:

5

-e

°E

0.108 CaAh5i20s + 0.446 Ca2Al l.3335i20s

(anorthite)

(an)

(gr)

+ 0.372 AI3.25i1.60S + 0.074 5i40s (as)

(3.19a)

(qz)

Thus, according to the stoichiometry of this reaction, the speciation model predicts that X~~ =a~~l =0.074. The cryoscopic equation for quartz, which was derived entirely independently from that for anorthite, yields a~~l = 0.072 at this invariant point! This remarkably precise agreement between the stoichiometrically derived X~~l and the cryoscopically derived a~~l at this high pressure is surprising, but not unique. Equally precise agreement is obtained at the 33.3-kbar, 136rC invariant point on the melting curve of albite, where the stoichiometry of the dissociation reaction, NaAl5i3 Os (albite)

¢:::>

0.771 NaAl5i3 (ab)

+0.0575i40S

o, + 0.172 Na 1.333A11.3335i2.6670S (jd)

(3. 19b)

(qz) and the cryoscopic equation for quartz yield identical values for X~~l and 3 0l a qz •

Melt Speciation and the Quasicrystolline Model

79

The precise agreement between the stoichiometrically derived mole fractions and the cryoscopically derived activities of speciation reaction products carries with it the important implications that (1) the accuracy of the cryoscopic equations involved is very high; (2) the ideal mixing behavior of reaction products assumed by the quasicrystalline model is verified; and (3) the nature of the speciation reaction (identity of the reaction products) is confirmed. In this latter connection, the appearance of corundum, co, on the anorthite liquidus at 904 kbar, coupled with its disappearance at approximately 32.1 kbar (S2, Figure 3.7), might appear to cast doubt on the speciation reaction as written in equation (3.17). The preliminary cryoscopic relations for corundum (Als.333 Os), however, indicate that a~~ at the 904-kbar singular point (0.05) is less than half that required to lower the anorthite liquidus temperature by 60 °C (Figure 3.7) and the corresponding a~~ to 0.717. It is therefore presumed that a secondary speciation reaction, 0.083 A13.2Si1.60S 0.05 Als.3330S + 0.033 Si40s (as) (co) (qz)

(3.20)

produces co and qz from the intermediate product as. As pressure is increased to 32.1 kbar, however, primary reaction (3.19) apparently produces enough qz to drive reaction (3.20) back to the left and reduce a~~l below the approximately 0.02 required to sustain corundum on the liquidus. The values of a~~ in melts of pure anorthite composition as a function of pressure along the melting curve of anorthite, calculated from equation (3.11), may be substituted into equation (3.18) to yield the first two pressure terms in equation (3.21), as well as curve (S-L-X) in Figure 3.7: x:~

-.a~~l

= X~~ [3.11 x 10 - 2P - 1.09 X 10- 4p 2

+ X~~/X~(sat)(3.04 x 10- 2 - 4.42 x 1O- 2P + 8.6 x 10 - 3 p2 - 2.0

X

10- 4 p3 - 7.1 x 10 - 6 p4

+ 1.98 x 10- 7 PS) + Xf~1(1.86 - 5.2 x 1O- 2 P)]

(3.21)

where P is in kilobars. It should be emphasized that this equation includes terms only for the effect of pressure (dissociation), interaction with H20, and interaction with the fo (Mg4Si20S) melt component. Interaction occurs to some extent with nearly all other nonaluminosilicate components in igneous rock melts, but their inclusion in equation (3.21) is beyond the scope of the present discussion and would merely serve to confuse the issue. Saturation of anorthite melts with H20 lowers the melting temperature along the experimentally determined S-L-V-X curve in Figure 3.7. This curve, which is based on the data of Yoder (1954, 1965), Stewart (1967), Boettcher (1970, 1971), and Erikson (1979), passes through a singular point (S1) near 9.0 kbar, where corundum appears, then through a second singular point (S2) at 14 kbar, where zoisite appears, and finally terminates at the

80

Magmas and Hydrothermal Fluids

14 .8-kbar, 880°C invariant point (1) with the assemblage anorthite, zoisite, corundum, kyanite, melt, and " va po r" (Hj Ovrich fluid). It will be observed that, unlike in the anhydrous system, corundum persists to this invariant point, because a~~l is much too low to stabilize quartz. Also, as with all other hydrous minerals to be discussed later, a zoisite-like stoichiometric un it (Ca I.2sA II.92Si1.92 O s' 0.32 H20) is no longer regarded as a melt component (d. Burnham, 1981 a), but as a hydrous precipitate resulting from the inverse of an incongruent melting reaction involving the gr, as, and qz speciation products in the proportions 24 : 12 : 1. In comparison with the calculated S-L-V-C curve in Figure 3.7, which was calculated from the cryoscopic equation (3.11) using values of a~~ from the first two P terms in equation (3.21) and values of a~~n from equations (3.9) and (3.10), the experimentally determined S-L-V-X curve lies at lower temperatures at all pressures above approximately 2 kbar and the difference increases markedly with increasing pressure up to the 14.8-kbar invariant point (1). These relationships clearly indicate that addition of large amounts of H20 substantially increases the extent of dissociative speciation, presumably by increasing the number of non bridging oxygen sites in the melt and thereby promoting the IV Al => VI Al shift, as in the formation of gr and as (VIAl) from an (IV AI). Under Hj Ovundersaturared conditions [X~/X~(sat) < 1.0, equation (3.21 )], however, anorthite-melt ("vapor" absent) equilibrium persists to higher pressures, the actual values of which are dependent on X~~. Thus there is an entire family of curves between the S-L-X and S-Lv-x curves in Figure 3.7, the pressure coordinates of which are nonlinearly dependent on the X~~ ratio in equation (3.21). It was noted by Burnham and Nekvasil (1986) that the addition of qz (Si4 O s) to anorthite melts suppresses the simple dissociation reaction (3.19a) via the mass action effect, in favor of an interactive type speciation reaction of the form 1.0 CaAI2Si20S + 0.416 Si40s ~ 0.50 Ca2AI1.333Si20s (an) (qz) (gr) + 0.916 AI1.4SSSi2.9I Os

(3.22)

(dpy) where dpy is a dehydroxylated pyrophyllite-like species in which Al is in fivefold coordination (v AI). As also noted by these authors, the dpy species is produced by interaction between qz and all three feldspar-like species (an, ab, and or). The fact that the extent of dissociation of an with pressure is much larger than that of ab, as indicated in equations (3.19a) and (3.19b), provides a simple explanation for the experimental observations of Green (1969) and others more recently that the liquidus plagioclase in mafic melts of fixed bulk composition becomes more albitic with increasing pressure. This effect of enhancing th e albite content of liquidus plagioclase as a result of the differential lowering

Melt Speciotion ond the Quosicrystolline Model

81

of a ~~n with respect to X ~~l has important implications for the origin of massif anorthosites, which commonly contain An50-5 5 plagioclase. More important, perhaps, when combined with the effects of speciation by interaction with the olivine-like components of mafic melts, as discussed below, the extensive speciation of the an component carries very important additional implications for the origin of tholeiitic and alkali olivine basalts, which in turn carry important implications regarding the sources ofI-type (White and Chappell, 1977) orebearing magmas, especially those parental to "porphyry type" deposits. Application of the cryoscopic equation (3.11) for forsterite, which was derived mainly from the experimental work of Roeder and Emslie (1970), to the experimentally determined liquidus relations on the joint forsterite--anorthite (Anderson, 1915; Kushiro and Yoder, 1965; Presnall et aI., 1978) yields an am (093 + 1 1 x 10 - 2 P)] where again P expression for X am fa - aam fa = Xam[x fa an " , is in kilobars. A comparison with the last parenthetical term in equation (3.21) reveals that, at atmospheric pressure, the extent of speciation of the fo component along this join is just half that for the an component; hence the speciation reaction may be written 1.0 CaAI2Si20s + 0.50 Mg4Si20S ~ 0.50 Mg2A140s + 0.25 Si40s (an) (fo) (sp) (qz) + 0.75 Ca1.333Mg1.333Si2.6670S

(3.23)

(di) where sp is a spinel-like species. As a consequence of the extensive pressure effect on the dissociative speciation of the an component, as compared with none on the fo component, the pressure coefficient in the above expression is greater than half the corresponding coefficient in equation for X (3.21). In other words, in the presence of fo the extent of dissociative speciation with pressure actually appears to be diminished slightly. Confirmation of the stoichiometry of reaction equation (3.23) comes from the fact that a liquidus field for spinel extends from An55 to AnS5 along the forsterite-anorthite join and its thermal maximum occurs essentially at An67 (2 An: 1 Fo; Anderson, 1915), where the maximum amount of sp (0.21 mol) is produced. Upon increasing pressure, the extent of sp production increases slightly, reaching only 0.26 mol at 20 kbar. As this pressure is approached, however, the nature of the products, but apparently not the stoichiometry of the reactants, changes such that a pyrope-like (py) species is formed. In this simple binary system, the speciation reaction may be written

r: -ar:

1.0 CaAl2Si20s + 0.50 Mg4Si20g ~ 0.75 Mg2Al1.333Si20S (an) (fo) (py) + 0.375 CaI.333Mg1.333Si2.6670S + 0.375 Ca1.333Ah.667Si1.3330S (3.24)

(di)

(cts)

82

Magmas and Hydrathermal Fluids

wh ere cts is a calcium Tschermaks pyroxene-like species in which half of the Al is in sixfold coordination. In more complex natural basaltic systems, the ab (NaAlSi30g) component contributes Si [ef. equation (3.19b)] to the py seecies, thereby producing jd (Na 1.333AII.333Si2.6670g), a jadeite-like species ( I AI), and greatly reducing the cts content of the melt. Accompanying the formation of jd, both by this reaction and the pressure-induced dissociation of the ab component, the coexisting diopsidic pyroxene tends to become omphacitic at these high pressures. The net result of precipitation upon cooling of silica-bearing pyropic garnet and clinopyroxene from the natural olivine basalt analogue of the melt represented in reaction equation (3.24), as elaborated on by Yoder and Tilley (1962), is to drive the residual melt into the nepheline-normative field, thereby yielding the family of alkali basalt magmas. Conversely, precipitation of the nonsilicate spinel and clinopyroxene from the same olivine basalt melt at pressures less than approximately 20 kbar, in accordance with reaction equation (3.23), drives the residual melt into the hypersthene-normative field, thereby yielding the family of tholeiitic basalt magmas that appear to be parental to the subduction-related Cu-Au porphyry magmas, as in the Andes and southwestern Pacific (e.g., Bougainville). Finally, if this same olivine basalt melt is cooled in the vicinity of 20-25 kbar, both spinel and pyrope-rich garnet plus omphacitic (jadeitic) clinopyroxene may precipitate, thereby giving rise to the more shoshonitic, trachytic, and peralkaline trends.

Melt Species and Coexisting Crystalline Phases A majority of melt species are represented by crystalline phases with the same stoichiometry, but several of them are not. In the preceding section it was mentioned that the dpy species is an ubiquitous product of interaction between qz and all of the feldspar-like components, but neither it nor its hydrous equivalent, pyrophyllite, appears as a solid phase at magmatic temperatures. The dpy component is especially prevalent in felsic peraluminous (corundum-normative S-type granite) melts, from which it may precipitate as sillimanite (rarely kyanite or andalusite) plus quartz or topaz plus quartz in F-rich magmas. In hydrous peraluminous melts, dpy combines with the or component to precipitate muscovite plus quartz under appropriate P-T conditions. Herein, incidentally, lies a ready explanation for the almost ubiquitous coexistence of quartz with primary book muscovite in peraluminous granite pegmatites. In hypersthene-normative peraluminous melts, on the other hand, co(AIs.3330g) first combines with equal molal amounts of hy [(Mg, Feh.667Si2.6670S], a hypersthene-like species, and qz to produce cd [(Mg, Fe)o.89A I I.7S Si2.22 Os], a cordierite-like species (T. Chacko, unpublished experimental results, 1985; Burnham, 1992), which may precipitate as cordierite, if the hy component is Mg-rich, or as almandine-rich garnet plus sillimanite,

Melt Speciation and the Quasicrystalline Model

83

if hy is Fe-rich. In hydrous melts of equivalent composition, cd also combines with or to precipitate aluminous biotite plus quartz. Also, in F-rich melts, which are commonly hydrous and peraluminous, F appears to produce both fl, a fluorite-like component (Ca4Fs), and cr, a cryolite-like component (Na4A11.333 Fg), in proportions that depend on the an: ab ratio in the melt (Burnham, 1992). In contrast to dpy, which has no counterpart in the crystalline phases that precipitate from igneous rock melts, none of the hydrous minerals in igneous rocks appear to have analogous melt species counterparts. The precipitation of zoisite, muscovite, and biotite by the inverse of incongruent melting reactions has already been discussed. Amphibole (mainly hornblende), the remaining especially important hydrous mineral in igneous rocks, also forms by the same type of reaction. For each mineral, however, the conditions of, and constraints on, precipitation differ. The overriding constraint on the precipitation of hydrous minerals from magmas is the H20 content (X~~) of the melt , which from equations (3.4) and (3.5) is reflective of a~~n, the activity of H20 in the hydrous melt as well as in the hydrous mineral (hrni). The minimum H20 content of the melt required for the precipitation of a given hydrous mineral is, within the relatively narrow limits of the multivariancy of the reactions, the same as that required to produce fluid-absent partial melting. Thus the H20 contents indicated along the fluid-absent beginning of melting P-T curves in Figure 3.8 are essentially the minimum amounts required to precipitate the particular hydrous mineral from the melt. Owing to recent refinements in the quasicrystalline model, as discussed above, the H20 content indicated along the fluid-absent beginning of the melting curve of muscovite in Figure 3.8 is 7.4 wt % H20, not the 8.4 wt % (X~~) = 0.59) reported in Burnham and Ohmoto (1980, Figure 1). No corrections in H20 content are required at the beginning of melting of biotite-bearing rocks or hornblende-bearing basaltic amphibolites. In addition to the Mu-S curve in Figure 3.8, another major constraint on the precipitation of primary muscovite from granite magmas is the activity (mole fraction) of the dpy component in the melt. Although the actual value of this critical activity is not known with certainty under all relevant conditions, the experimental data of Voigt (1983) and Joyce (1985) indicate that it is very close to that at melt saturation in an aluminum silicate such as sillimanite, which is approximately 7.7 mol % dpy equivalent on an anhydrous melt basis (Burnham, 1992). The remaining constraint on the precipitation of primary muscovite, of course, is the activity of the or component (a~l~l), which appears to be close to that for the precipitation of K-feldspar. Like muscovite, the constraints on the precipitation of biotite from granitic magmas include a~~l in addition to X~~, but whether or not saturation in K-feldspar must occur before biotite is stabilized remains to be determined. Unlike muscovite, given appropriate values for X~~ and a~~\ precipitation of biotite may be promoted by substantial activities of either the hypersthene-

84

Magmas and Hydrothermal Fluids

90 80 70 60 50 ~ -

-

.c Co

. a:::

40

~

30 20 10

An95

800

1000

1200

Temperature. °C Figure 3.8 Pressure-lemperature prajections of melting relotions in muscovite- ond biotite-beoring schists and gneisses,

and hornblende-beoring basoltic amphibolites (BA). The curves labeled GD-S and BA-S (a.. '" 1) ore the H2 (}saturated solidus for overage granodiorite and tholeiitic basalt, respective~ ; the curves labeled MiJ"S (7.4%H20), B~S (3.3%H20), and BA-S (2.7%H20) represent the divoriont, fluidiJbsent, beginning of melting of nonporous schists, gneisses, and amphibolites. The percentage H20 figures ore the H20 contents, inweight percent, of the first-formed melts. The crystal/melt equilibrium points A-D and the equilibrium curves emanating from Aand Bore as described in the text; the temperatures of the Hy-L and Aug-l curves ore minima, because no allowance has been made for the higher temperature precipitation of plagioclase. [Adopted from Figure 1 of Burnham and Ohmoto (1980)].

forming components [en (enstatite-like, Mg2.667Si2.6670S) and fs (ferrosilitelike, Fe2.667Si2.6670S)I or cd [(Mg, Fe)o.s9AI1.7sSi2.22 as], th e cordierite-like component, or a combination of the two in some mildly peraluminous, metasedimentary source S-type granite magmas, such as the Dartmoor granite discussed below. In metaluminous I-rype (igneous source) granitic melts (W hite a nd Chappell, 1977), cd is either very minor in amount or absent; hence th e biotites precipitated are much less aluminous than those precipitated from cd-bearing S-type granite melts. Owing to the generally more oxidized nature of l-rype magmas, moreover, th eir biotites also tend to be substantially more magnesian. In passing from muscovite through biotite to hornblende, the constraints

Generatian and Emplacaement of Hydrous Magmas

85

on pr ecipitation, in addition to X~~ , not only become more nu merous but also more complex. From th e compilat ion s of Leak e (196 8), Burnham (1979) observed that the average hornblende o f diorites and of qu artz d iorites is composed of ten melt co mponents-as then defined-five of which lab , an , di, w, and 01 (olivine-like) or mt (magnetite-like)] were regard ed as essent ial to the formation of hornblende in melts of pr esent int erest. As melt components are presently defin ed , this list sho uld be modified to inclu de a b, an, di, hd , w, en, and fs, wh ere hd , en, a nd fs are hedenberg ite-, ensta tire-, and ferrosilite-lik e melt components. The activity relat ions of eac h of th ese me lt components (except w ) required to pr ecipitat e horn blende, however, a re no t well known at the present time. Petrographic evidence ind icates that the melts are saturated, or very nearly saturated, with respect to clinopyroxen e (di a nd hd) and plagioclase (an and ab). In fact, the ab sence of hornblende in per aluminous S-type magmas, which by definition lack di and hd , suggests that near saturation with respect to clinopyroxene (augite) is required for the precipitation of hornblende.

GENERATION AND EMPLACEMENT OF HYDROUS MAGMAS Following Burnham (1979, 1981b) and Burnham and Ohmoto (19 80 ) regarding the partial melting process, the source rocks of hydrous felsic magmas-whether in mafic oceanic crust of subduction zones, in th e subco ntinental upper mantle, in amphibolitic to biotitic rocks of the mid- to low er continental crust, or in metasedimentary rocks of the latter region-are here regarded as too low in porosity to harbor a significant am ount of aq ueo us fluid. Moreover, what pore fluid may be present is likely to be C02 -rich; hence the amount of pore H 20 available to promot e partial melting above the H 20 saturated solidii in Figure 3. 8 is minu scule (d. Burnham, 19 79, pp. 92ff) and henceforth will be ignored. Accordingly, fluid-absent partial melting of a muscovitic source rock do es not occur within th e continental crust until the temperatures of th e Mu-S (7.4 % H 20) curve ar e exceeded. Similarly, partial melting of a biotite-bearing, but muscovite-free, so urce rock does not occur until the temperatures of the Bi-S (3.3 % H 20) curve in Figure 3.8 are exce eded. Fluid-absent melting of hornblende-rich basaltic amphibolite, on th e other hand, does not occur anyw here with in the region bounded by the BA-S (2.7% H 20) curve and th e H j Oesaturared solidus, BA-S (aw "" 1), but only a bove it in either temperature or pressure. M etasedimentary source rocks th at give rise to S-typ e granitic magmas commonly contain both mus covite and biotite; as a co nsequence, melti ng will begin at the Mu-S (7.4 % H20) curve, and some biotite will dissolve in this melt and may disappear at a temperature below, or slightly above , the Bi-S (3.3 % H20) curve, dep ending on the initial muscovite and biotite contents of the source rock and its bulk composition. The positions of th ese

86

Magmas and Hydrothermal Fluids

three sets of beginning-of-melting curves are based primarily on ex pe rime ntal data rep orted by Allen et a!. (1972 ), Allen and Boettch er (19 78 ), Bo yd ( 1959), Eggler and Burnham (1973), H amilton et a l. (196 4), H elz (1973), Hill and Boettch er (1970 ), Holloway (1973), Holloway a nd Burnham (19 72 ), and Yod er and T illey (1962) for the basaltic amphibolite, Blenco e (1 974) a nd Burnham ( 1967) for th e muscovitic meta sediment, and Piw inski i (19 68) fo r th e bior iric so urce rocks. In add ition, th erm od ynamic ca lcula t io ns usin g a pr eliminary version o f th e qu asicrystalline model, as described above, wer e employed in large part to assess th e H20 contents of th e first -form ed melts.

Magmas Generated in Subducted Basaltic Amphibolites Considerable attention will be devoted here to the fluid-abs ent partial m elt ing of arnphibo litized tholeiitic basalt, because the magmas parenta l to most , if not all, porphyry copper-gold and porphyry mo lybd enum deposits appear to have been deriv ed th erefrom by partial melting, coupled with crystal/melt differentiati on/fraction ati on and, in the case o f Clim ax type molybdenum porphyries, with partial remelting. A very common setting for porphyry type Cu- Au depo sits is in volcanic arcs overlying subduction zones, whether in the Andes (e.g., E1 Salvad or, Chil e; Gustafson and Hunt, 1975) or on th e island of Bougainville (e.g., Panguna; Mason and McDonald, 1978) in th e sout hwestern Pacific. In this setting, the amphibolitic so urce rocks for th e calc-alka line porphyry magmas ar e, on average, tholeiite in composition. In the continental sett ing, such as th at of the southwestern North Amer ica porph yry Cu- Au depo sits, on th e other hand, unaltered samples of some orerelated porphyries from th ese deposits, wh ere avail abl e, ind ica te by th eir low FeO: Fe203 ratios and other features that the ir pa rent magmas, in a majority of cases, also were derived initially by partial melting of su bd ucted tholeiitic arnphib olires, but und er somewhat different pressure cond itions. The conditions of fluid-absent partial melting of a hydrated (amphibolirized ) subducting slab of the " normal" tholeiitic basalt of Nockolds (1954 ) (d. Figure 3.9a), th e average Hawaiian Islands tho leiitic basa lt (Macdonald and Kats ura, 1964) (d. Figure 3.9a), as indicated in Figure 3.8, or even of the ave rage unalt ered ocea nic crust (d. Alt and H onno rez, 1984), are critica lly dependent on th e therma l regime in the upper part of the subducting slab and the H20 content of the amphibo lite. Under conditions of a low geothermal gradient of 9.0°C/km (d . Figure 3.8), for exa m ple, melting o f this basaltic amph ibolite will begin at approximately 21.5 kbar, 660°C, where th e geotherm intersects the Hj Oesaturared solid us . At thi s point, th e first-formed melt contains approximately 27 wr % H 20; hence a basaltic amphibo lite containing 1.3 wt % H20 (-65% hornblende), which is that of th e average hornblendite (N ocko lds, 1954), can produce only 4 .8 % melt. This melt " never sees th e light of day," because even if su ch a minuscule amount of melt could separate from the slab and ascend , it would react with

Generation ond Emplocoemenl of Hydrous Mogmos

87

upper mantle peridodite, pass back into the stability field of hornblende, and resolidify. Similarly, melts produced by fluid-absent partial melting of this same amphibolite in Figure 3.8 at approximately 22.5 kbar along an II °C/km subduction slip-zone geotherm, which is that proposed by Turcotte and Schubert (1973), would contain 9.8 wt % H20, amount to only 13 vol % of the total mass, and also would be incapable of ascending vertically through the "hornblende barrier" (Burnham, 1981b) without resolidifying. In the latter case, however, the melt/crystal mush produced likely would remain with the slab to modestly greater depths, higher temperatures, and thus larger proportions of melt before it and more or less of its restite rises diapirically into the hotter overlying mantle wedge, as indicated schematically by the arrow in Figure 3.8. Accordingly, our starting point for the following discussions will be when this latter melt/crystal mush circumvents the "hornblende barrier," a requirement if it is to reach the near-surface environment, and reaches (arbitrarily) the oceanic geotherm of Ringwood et al. (1964) in the overlying mantle wedge, as indicated by point A in Figure 3.8 at 21.5 kbar (-74 km) and 1120°C. The preceding view of magma generation in subduction zones, first proposed by the writer in the second edition of this volume (Burnham, 1979), is at odds with the views of some more recent workers in the field of thermal modeling. Many of the findings to be presented later in this chapter, however, are critically dependent on the specific mode of origin of the magmas and can most readily be explained by a subducted tholeiitic amphibolite source material. . At point A, calculations using the quasicrystalline model described above and in Burnham and Nekvasil (1986) and Burnham (1992) indicate that tholeiitic amphibolites, the bulk compositions of which-except for the assumed 1.15 wt % H20-are shown in Figure 3.9a for the "normal" tholeiite (NR-THL) of Nockolds (1954, Table 7) and the average Hawaiian tholeiite (AV-THL) of Macdonald and Katsura (1964, Table 9), will yield 38 mol % (z vol %) melts that contain 2.9 wt % H20 and are in equilibrium with An88 and Ann plagioclase, respectively, as well as with enstatitic hypersthene, highly aluminous (Jd 27) omphacitic clinopyroxene, magnetite, ilmenite, apatite, and quartz. The experimental results of Allen er al. (l972) indicate that garnet also is present at this temperature and pressure, but its composition is not presently known and its influence on the compositions of other coexisting phases has therefore not been considered. The equilibrium melts at point A, the compositions of which are presented in Figure 3.9b in terms of the eight oxygen atoms per formula weight convention described previously, contain 6.5 mol % qz (Si408) component, which is all that is required to achieve saturation with quartz at 21.5 kbar and 11 10°C. This very small percentage of the qz component required to saturate the melt in quartz places severe constraints on the depths at which felsic magmas can be generated. Thus the derivation of granitic melts containing more than

88

Magmas and Hydrothermal Fluids

32

(a)

28

-

24

I?ll

NR-THL

100)

AV-THL

:.~:~

c: 20 Gl

....CJ Gl

c.. '0

:i: 8

ab

or

qz

an di hd en Rock Component

Is

mt

i1+ap (b)

% Melt

i: Ql ....CJ Ql c.. '0

~ NR-THL

38

[ill

38

AV-THL

:i:

or

qz

an

di hd en Melt Component

mt

il+ap

% Melt

c: Ql

....QlCJ

20

~ AV-AND

94

[ill

AV-THL

40

~ NR-THL

46

c.. 16 '0

:E 12 8 4 0

an di hd en Melt Component

mt

i1+ap

Figure 3.9 (0) Bulk compositions of the "normal" tholeiite (NR·THl) of Nockolds (1954) and the overage Hawaiian tholeiite (AY·THl) of Macdonald and Kotsura (1964), based on the eightilxygen molecular norm. (b) Compositions of the melts (less 2.9 wt %H20) from the "normal" (NR·THl) and overoge (AY·THl) tholeiites of Figure 3.90 that are in equilibrium with reslite at point Ain Figure 3.8. Melt components are as defined in the text. (c) Compositions of melts (less 2.9 wt %H20) from overage (AY·THl) and "normal" (NR·THl) tholeiite sources, in equilibrium with restite at point Bin Figure 3.8, as compared with that of the overage andesite (Nockolds, 1954). Note the close similarities in the or and qz contents of each melt.

Generation ond Emplocoement of Hydrous Magmos

89

15% normative quartz (maximum at 12 kbar) directly from the mantle of the Earth is not a viable process. The calculated melt compositions at point A, Figure 3.8, as presented in Figure 3.9b, are of special importance here because of their general similarities to the compositions of many andesites around the world. The P-T-X conditions at point A will therefore be used as a starting point in the development of the following scenarios: (1) the ascent of the melt phases produced at point A from both tholeiitic sources to the near-surface environment, from which they might erupt to produce anorthite megacryst-bearing andesites; (2) the diapiric ascent of the entire magma bodies produced from both tholeiitic sources, without extensive crystal/melt fractionation, to point B at a depth of approximately 41 km (12 kbar), (3) the separation of the melt phase produced from "normal" tholeiite at point B from its crystalline residua (restite) and emplacement in the subvolcanic environment (-7.5 krn): and (4) the emplacement and solidification of andesitic magmas and their differentiates at a midcontinental crust depth of approximately 18 km (5 kbar), where they later undergo partial remelting and emplacement in the near-surface environment. The term "restite," because it will be used frequently in the ensuing discussions of the generation of magmas, should be clarified at this point. As defined by White and Chappell (1977) and elaborated on by Chappell et al. (1987), the term "restite" refers to that mineral fraction of a rock that remains solid during partial melting. Thus application of the quasicrystalline model to several suites of Australian granites (Burnham, 1992) revealed that melts of uniform composition within a given suite, in equilibrium with quartz and plagioclase of a fixed composition at the same temperature and pressure, could be extracted from each member of a given suite by removing up to 65% of the bulk rock as unmelted crystalline residue-that is, "restite." This 65% maximum is close to that value beyond which van der Molen and Paterson (1979) found the melt phase in quantities too small to segregate from granitic magmas. Application of the quasicrystalline model to many other suites of igneous rocks from around the world strongly supports the concept that the vast bulk of variations in composition within a given suite of granitic rocks, or even within a single pluton, are the result of variations in the extent of segregation of melt from restite. The fact that the amount of restite in a given suite with fixed melt and feldspar compositions ranges from approximately 10 to 65 vol % suggests that the melt/restite segregation processes, such as filter pressing, are highly variable in efficiency. SCENARIO 1

In the first scenario, the 38% melts produced at point A, Figure 3.8, from amphibolitized "normal" (Nockolds, 1954) and average Hawaiian (Macdonald and Katsura, 1964) tholeiites of Figure 3.9b are presumed to effectively separate from the bulk of their restite and ascend directly to the near-surface

90

Magmas and Hydrothermal Fluids

environment. During this ascent, the negative TIP slope of the plagioclase liquidus, as shown contrasted with the positive TIP slopes of the crystal-melt equilibrium curves for the other phases in Figure 3.8, induces precipitation of anorthitic plagioclase, the amount of which depends on the thermal regime of the ascending magma body. In the interior of an ascending body, effectively insulated from heat loss to the wall rocks, the cooling effect of adiabatic expansion upon decompression from 21.5 to 2.0 kbar, for example, is more than compensated by the latent heat released from the crystallization of 15% anorthitic plagioclase from the "normal" tholeiite-derived melt or of 18 % from the average Hawaiian tholeiite-derived melt . These amounts of plagioclase crystallization are required to reestablish plagioclase-melt equilibrium at H20 contents in the melt that are 15-18% higher (3.4-3.45 wt % ) and corresponding temperatures that are 30-45°C higher than initially at point A (1120 °C). The end-products of this process, which subsequently may be eru pted, are anorthite megacrysts of Ann ("normal ") and An95 (average Hawaiian) compositions set in a largely aphyric, vesicular groundmass, as described from many circum-Pacific localities, notably in Japan (d. Kuno, 1950). A specimen of largely aphyric, vesicular andesite in the writer's possession from Arenal volcano, Costa Rica, contains euhedral anorthite megacrysts that are as much as 3.5 em across and An93-95 in composition. This extremely close correspondence between the calculated and observed compositions of these anorthite crystals lends strong support to a megacrystic (autolithic)-not xenocrystic (xenolithic)-origin, to the quantitative predictability of the quasicrystalline model, on which the calculations are based, and to the subducted tholeiite origin of the megacrystic andesite magmas. SCENARIO 2

In the second scenario, essentially the entire bodies (melt plus restite) of average Hawaiian and "normal" tholeiite magmas at point A, Figure 3.8, are presumed to continue diapiric ascent to point B at a depth of approximately 41 km (12 kbar). Upon equilibration at point B, the andesitic melts from both sources contain 2.9 wt % H20 and are in equilibrium with AnS5 plagioclase at a temperature of 1130 ± 5°C. The compositions of these melts, after removal of 60% restite from the average Hawaiian tholeiite (AV-THL) and 54 % resrire from the "normal" tholeiite (NR-THL) compositions, ar e presented in Figure 3.9c, where they may be compared with melt having the bulk composition of Nockolds' (1954) average andesite (AV-AND) at the same temperature (l130°C), plagioclase composition (AnS5), and H20 content (2.9 wt % ). The correspondence between the composition of the melt from Nockolds' average andesite and that derived by 40% partial melting of the average Hawaiian tholeiite (Macdonald and Katsura, 1964), close as it is, has several important implications, not the least of which is the confirmation of the widely held view of many petrologists (d. Boettcher, 1973; Carmichael et a!.,

Generation ond Emplacoement of Hydrous Magmas

9J

1974) that orogenic andesites are derived from tholeiitic parents, most commonly by partial melting in subduction zone s. Another important implication of this correspondence is, again, confirmation of the quantitative predictability of the quasicrystalline model in defining equilibrium relations in magmas. Yet another, and immediately more relevant, implication is that this origin of orogenic andesites is a starting point in the development of a quantitatively rigorous model for the derivation of ore-bearing magmatic systems. There are other important implications of the compositional relations in Figure 3.9c. For example, the difference between Nockolds' (1954) average andesite melt and that derived by partial melting of his "normal" tholeiite is not as small as for that derived from average Hawaiian tholeiite. The differences, which are mainly in 3-4 % higher contents of or and qz, and a correspondingly lower content of an in the latter, are readily attributable to the "normal" tholeiite having undergone minor anorthitic plagioclase fractionation at low pressures, as compared with the average Hawaiian tholeiite. In this respect, the An91 plagioclase-melt equilibrium temperature at 2.0 kbar for the "normal" tholeiite is more than 70°C above the first appearance of the second crystalline phase, augite; hence there is an ample temperature interval over which plagioclase alone may fractionate. Despite these minor differences, however, all three andesitic melts have in common generally low or: ab ratios « 0.3) and or: qz ratios close to unity. These or: ab and or: qz ratios, coupled with low FeO : Fe203 ratios, appear to be characteristic features of andesitic rocks from island arc and continental margin orogenic settings, as well as of the intrusive stocks with which the porphyry copper-gold deposits are associated in these regions. Within the continental cratons, on the other hand, this uniformity of andesite compositions no longer prevails: (1) the or: ab ratios tend to increase, leading to more latiric (monzonitic) compositions, as in the San Juan Mountains of Colorado (Larsen and Cross, 1956); (2) the or: qz ratio commonly becomes much greater than unity; and (3) the FeO: Fe203 ratio in unweathered rocks tends to increase markedly. The first two of these differences appear to go hand-in-hand, as would be expected if the melts segregated from their tholeiitic to more alkalic source magmas at deeper levels in the upper mantle than that found above for the average andesite. Owing to the effects of pressure on increasing the jadeite (1.33 ab) content of the restite clinopyroxene solid solution and on decreasing the equilibrium qz content of the melt, both the or: qz ratios are increased. Moreover, this latter ratio is enhanced at higher pressures by the increased stability of silica-rich garnet, as contrasted with the greater stability of silica-free spinel at lower pressures [equation (3.23) and (3.24)]. The third difference, in the FeO: Fe203 ratio, is presumably due to the fact that these subcratonically derived magmas have never been exposed to surface oxidation, such as on the seafloor prior to subduction. As an example, the molar FeO : Fel03 ratio in the Proterozoic Sudbury, Ontario, norite (Naldrett and Hewins, 1984), as discussed further below, is 7.4 (very reduced), whereas that in Nockolds' (1954) aver-

92

Mogmos ond Hydrothermol Fluids

age andesite is 3.5 and that in the Cu-Au porphyry, the Kav erong quartz diorite at Panguna, Bougainville (Mason and McDonald, 1978), is only 1.6.

SCENARIO 3 In the third scenario, the andesitic melt that was derived by melting 46 % of " no rmal" tholeiite at point B (Figure 3.8) in Scenario 2 above (Figure 3.9c, NR-THL) is presumed to segregate from its restire and ascend to a subvolcanic depth of 7.5 km (2 kbar) . Upon emplacement at this depth and reestablishment of equilibrium betw een the melt and all major crystalline phases, the temperature will have decreas ed to 970 °C and 31 % of the total mass, 24 % of which is An73 plagioclase, will have precipitated. The composition of the melt under these conditions-less 3.65% dissolved H20-iS shown in Figure 3.10a (NR-THL), where it may be compared at the same temperature, pressure, and H20 content with that of the mineralization-related Kaverong quartz diorite (KAV-QD) in the Panguna, Bougainville, porphyry copper-gold deposit (Mason and McDonald, 1978, Table 2). This comparison is made after removal of 36 % restite plus differentiate from the Kaverong quartz diorite, to bring the melt also into equilibrium with hypersthene, as well as An73 plagiocla se, augite , magnetite, ilmenite, and apatite. The correspond ence between these two melts and their equilibrium plagioclase compositions is again impressive and strongly suggests that the Kaverong quartz diorite was derived by parti al melting of amphibolitized " no rmal" tholeiite in a subduction zone and underwent segregation of th e melt from restite at a shallow upper mantl e depth, as well as 31 % crystal fractionation , probably at a shallow crustal level. Thus the final quartz diorite melt (N R-T H L) represents only 32 % of the origin al subducted thol eiite, but it contains most of the "incompatible elements," such as Cu. Th e term "incompatible elements " is placed within quotation marks here because the temperatures ar e substantially above the stability fields of minerals that sequester Cu , such as pyrrhotite and biotite (Candella and Holland, 1986). The similarity in compositions of the mineralization-related Kaverong quartz diorite at Panguna with differentiates of andesitic magmas produced by partial melting of "normal" tholeiitic basalt is not unique, as will now be demonstrated for the mineralization-related " L" porphyry at El Salvador, Chile (Gustafson and Hunt, 1975). Here again, the 46 % andesitic melt produc ed from the "normal" tholeiitic source (Figure 3.9c, NR-THL) at point B, Figure 3.8 , is again presumed to separate from its restite, ascend to a depth of 7.5 km (2 kbar) , and undergo differentiation. In this case, however, crystallization-differentiation is presumed to continue down to 910°C, the highest temp erature at which melt of the average least altered " L" porphyry co mposition (Gustafson and Hunt, 1975, Table 1, columns 6 and 9), less 3.7 wr % H20 and 21 mol % An63 plagioclase, is in eq uilibrium with plagioclase of the same composition and augite, hypersthene, magnetite, ilmenite, and ap at ite. The amount of differentiate required to be removed from th e

Generation and Emplocoemenl of Hydrous Magmas

93

(a) ~

!Zl

NR-THL

[ill KAV-CD

44

32 64

r-----.----.--.....---.-----;--~--.-----,---.--_,

40

~

36

'E ell

...u

32

~

NR-THL

28

lEI

"L"-POR 79

28

ell

C.

(5

:E

Figure 3.10 (0) Composition of melt derived from "normal" tholeiite (NR·THl) at 970°C, 2.0 kbor, and 3.65 wt % H20, as described in the text, compored with that derived hom Koverong quartz diorite, Ponguno, Bougoinville (KAV-QO), under the some conditions. (b) (ompcsincn of melt derived from "normal" tholeiite (NR·THl) at 910°, 2.0 kbor, and 3.7 wt % H20, as described in the text, compared with that derived hom the "L" porphyry, EI Salvador, Chile ("L"·POR) under the some conditions. (c) Composition of melt derived from "normal" tholeiite (NR·THl) at 870°C, 2.0 kbor, and 3.8 wt % H20, as described in the text, compared with that derived from the granodiorite porphyry of the Hanover Stock, Santo Rita, New Mexico (HAN-GOP), under the some conditions.

94

Magmas and Hydrothermal Fluids

"normal" tholeiite-derived andesitic magma (Figure 3.9c, NR-THL) to reach this equilibrium temperature is 39.5 mol %, 28% of which is An63 plagioclase, and the melt composition, less 3.7 wt % H20, is as shown in Figure 3.IOb (NR-THL). Also shown in Figure 3.10b is the composition of the average least altered "L" porphyry melt ("L"-POR), less the H20 and plagioclase mentioned above, from which it is apparent that this melt is slightly richer in qz and correspondingly poorer in hd (the hedenbergite-Iike component) than the "normal" tholeiite-derived melt. These minor differences can readily be attributed to the higher oxidation state of the" L" porphyry, which results in conversion of the ferrous silicate components into mt plus qz. Despite these minor differences, however, the similarity between the two melt compositions is great and serves to confirm an essentially "normal" tholeiite as the original source also for the "L" porphyry magma. Moreover, upon attaining equilibrium at 910°C and 7.S-km depth, the melt derived from the "normal" tholeiite source and that of the" L" porphyry represent only 28 % of the original volume of "normal" tholeiite source rock. Hence the "L" porphyry melt is presumed to have undergone an even greater enrichment in "incompatible elements," such as Cu, than did the Kaverong quartz diorite melt at Panguna. Upon moving farther onto the continents, magmatic evolution, not surprisingly, becomes more complex. The granodiorite porphyry of the Hanover Stock at Santa Rita, New Mexico (Jones et aI., 1967), for example, fits the above "normal" tholeiite-derived melt model well, as may be observed in Figure 3.10c (HAN-GDP), where removal of 78 mol % total restite plus differentiate from " normal" tholeiite (NR-THL) at 2.0 kbar provides good correspondence with the Hanover granodiorite porphyry after removal of 23 % restite (21% AnS8 plagioclase) from it at 870°C and an H20 content of 3.8 wt % . As with the "L" porphyry of EI Salvador, the low hd and fs contents, coupled with the higher qz content, are presumed to result from the higher oxidation state of the Hanover porphyry. Virtually these same compositional, phase equilibrium, and temperature relationships are shared by the Ruby Star granodiorite in the Sierrita porphyry copper deposit, Arizona (Anthony and Titley, 1988, Table 1, RS-MR-S), after removal of 18% Anss plagioclase. Thus the proposition that porphyry Cu-Au magmas were derived ultimately from subducted tholeiitic amphibolites is further strengthened. The Bingham quartz monzonite at Bingham, Utah (Moore, 1978), on the other hand, does not rigidly conform to the above model. Instead, the Bingham quartz monzonite, as shown in Figure 3.11, closely fits a variation of this model in which the melt derived from amphibolitized "normal" tholeiite (NR-THL), with 2.9 wt % dissolved H20, is segregated from restite at 24 kbar (-82 km) and 1l3SoC, emplaced at a depth of 7.5 km, and reequilibrated with AnS? plagioclase and hypersthene, but not augite. The Bingham quartz monzonite formed from a melt of the composition also shown in Figure 3.11 (BNG-QM) that is in equilibrium with the same minerals (also excluding augite) at the same temperature, pressure (2.0 kbar), and H 20 content, after removal of only 2.0 mol % hypersthene. Again, the correspon-

Generation ond Emplacoement of Hydrous Magmas

9S

36 r - - - - . - - - . . . , - - - - , , - -.......- - . , . - - - - . - - - . . . , - - - - , , - -.......- - - - ,

32

% Melt

28

C

24

~

20

[ill BNG-OM 98

NR-THL 18

CII

...u

CII Q.

"0 :E

8 4

ab

or

qz

an di hd Melt Component

en

fs

mt

il+ap

Figure 3.11 Composition of melt derived hom "normal" tholeiite (NR·THl) at 1000°C, 2.0 kbar, and 4.9 wt %H20, as described in the text, compored with that of the Binghom quartz monzonite (less 2.0%hypersthene), Bingham, Utah (BNlTOM), under the some conditions.

dence is remarkably close, especially considering the fact that only 32 mol % melt was produced at 82 km and 43 % of that was differentiated out upon cooling to 1000°C during emplacement, leav ing a melt that contained 4.9 wt % H20 and represented only 18 % of the original " normal" tholeiite. With the resulting approximately fivefold enrichment in " incompatible elements " such as Cu and Au in this melt, small wonder that the Bingham Mine ha s been such a large producer. The effects of doubling the pressure, hence depth, at which melt/restite segregation takes place is illustrated by a comparison of the "normal " tholeiite-derived composition (NR-THL) in Figure 3.11 with that in Figure 3.10a. Thus increasing the pressure from 12 to 24 kbar at essentially the same temperature of 1130°C increases the jd : Di ratio in the equilibrium clinopyroxene from less than 0.2 to 0.7. As a consequence, approximately four times as much of the ab melt component is removed and precipitated as jadeite component in restite omphacitic pyroxene, thereby increasing the or: ab ratio in the high-pressure melt. As a further consequence of this reaction, the di content of the melt required to attain equilibrium with the omphacitic pyroxene is greatly reduced. Hence the melt becomes grossly undersaturated with respect to AI-poor augite upon emplacement of the melt at shallow depths, which is precisely as observed above for the Bingham quartz monzonite melt (Figure 3.11 , BNG-QM). Also, upon increasing the pressure to 24 kbar at 1130°C, the solubility of quartz in the melt is limited to only 6 mol % (-5.5 wt %) ; hence the or: qz ratio increases markedly above unity as restite segregation occurs, again just as observed in the Bingham quartz monzonite melt. Furthermore, the reduction in ab content, coupled with enhanced speciation of the an component, results in a greatly increased

96

Magmas and Hydrothermal Fluids

an: ab ratio (more anorthitic plagioclase), once again as observed in the Bingham quartz monzonite melt. It would be instructive at this point to briefly examine the mineralizationrelated rocks of some stockwork molybdenum deposits. As recognized by Mutschler et al. (1981), there are two groups of stockwork molybdenite deposits-those associated with quartz diorite, granodiorite, and quartz monzonite stocks, such as the Kitsault molybdenum deposit, British Columbia (Steininger, 1985), and those associated with granite and rhyolite porphyry stocks and plugs, such as at Climax and Urad-Henderson, Colorado. The more mafic rocks of the former group, such as the East Lobe at Kitsault, resemble the Bingham quartz monzonite in composition, except they generally do not exhibit as large an enhancement in their or: ab and or: qz ratios. Nevertheless, their similarities strongly suggest a common parent, which was an amphibolitized "normal" tholeiite that probably was subducted and underwent rnelt/restite segregation at somewhat shallower depths than the parent magma of the Bingham quartz monzonite (82 km) , but greater depths than the parent magmas of the volcanic arc copper-gold porphyries (41 km). The magmas that produced the granite and rhyolite porphyry stocks and plugs of the Climax type, on the other hand, are so highly fractionated that their parentage is not as readily discerned. Their low FeO: Fe203 ratios suggest that they also originated from subducted tholeiitic amphibolites, but their high degrees of fractionation, their tectonic settings, and their lower Cu: Mo ratios suggest that more than one stage of partial melting may have been involved in their evolution, as now discussed. SCENARIO 4

In this fourth scenario, the same andesitic melt that was derived from Nockolds' (1954) "normal" tholeiite composition at point B in Figure 3.8 (Figure 3.9c, NR-THL) is again presumed to segregate from restite, but to ascend only to a midcrustallevel of approximately 18 km (5.0 kbar). During ascent and emplacement, this melt is also presumed to reequilibrate with all anhydrous crystalline phases, except quartz, by precipitation and removal of 13 mol % AnS3 plagioclase. This melt is then presumed to cool and crystallize, only to be partially remelted at a later time. Upon later partial remelting at the same pressure (point C, Figure 3.8), probably in response to lower crustal intrusion of mantle-derived mafic magmas in an extensional tectonic regime, equilibrium between melt (2.8 wt % H20), An29 plagioclase, augite, hypersthene, magnetite, ilmenite, apatite, and, simultaneously, infinitesimal amounts of quartz and Or60 K-feldspar is reached at 850°C. At this temperature, 45 mol % melt will have been produced, 93% of which consists of ab, or, and qz in the weight proportions shown projected onto the haplogranite ternary diagram in Figure 3.12 (THLS). Also shown by the enclosed area in Figure 3.12 is the compositional field of granite and rhyolite porphyries associated with the Climax-

Generation and Emplacaement of Hydrous Magmas

97

Qz

Ab

L..-_

Or

__'''_ _=_----'''''--_~:----'''---_=_=_-....Jo -4 kbar), sillimanite (rarely andalusite) at lower press ures, o r topaz from F-rich magmas, such as that of the Bodmin M oor gra ni te (F = 0. 34 wt % ). Topaz, like the hydro us minera ls, does no t form a co mpone nt in th ese silica- rich melts; instead , F in th e melt resi des in fluo rit e-like (f l, Ca4FS) and cryolite-like (cr, Na4A I1.333Fs) components (Burn ham, 1992), in propo rtions tha t are dependent on th e total F con tent an d the an : ab ratio, as indicated by the following equilibrium relationship tha t is based mainly on the experimental data of Luth and Muncill (1989): x amx am cr an 472 X XamXam = . HsFs f1

(3.25)

ab

in which XHsFs - 0 .0 17 (wt % F). The eq uili b rium melt compositions presented in Figure 3.15 for a ll four

104

Magmas and Hydrothermal Fluids

plutons, which represent the average bulk compositions minus restite, were calculated at a pressure of 3.0 kbar and a melt H20 content of 3.1 wt % (X~~~ = 0.33) mainly in order to facilitate comparisons among the four plutons and to minimize the amount of restite quartz that must be removed to attain equilibrium with melt and K-feldspar. Surprisingly, only 2.9 mol % restite quartz had to be removed from the bulk composition of only one pluton, the Dartmoor granite, to bring quartz and K-feldspar into equilibrium with melt at a temperature of 805°C. For the other three plutons, only 0.5 (Lands End), 2.0 (Bodmin Moor), and 5.0 (Carnmenellis) mol % or had to be removed in 01'70- 79 restite K-feldspar to achieve the same equilibrium state at temperatures of 795 °C (Carnmenellis), 800 °C (Bodmin Moor), and 810°C (Lands End). Moreover, all melts, except those from the Lands End pluton, also have An40- S0 plagioclase in the equilibrium assemblage. To bring plagioclase also into equilibrium with K-feldspar and quartz in the Lands End melts (Figure 3.15) would require the removal of 4.0% qz, an additional 6.0 % or, and 1.0 % cd, which would, in turn, lower the equilibrium temperature to approximately 800°C. These components were not removed, because the standard deviations from the mean for all three Lands End samples, as indicated by the error bars in Figure 3.15, are already smaller than those for the other plutons and indicate that no abnormally large amounts of restite quartz are present. Attention has been devoted here to these seemingly minor compositional features of the Cornubian batholith, because of the evidence they provide against crystal/melt fractionation as a significant process for enriching the batholith in Sn and other elements. Instead, when coupled with the additional observation that the centrally located Bodmin Moor and Carnmenellis plutons are exceptionally peraluminous, they provide evidence in support of a high degree of partial melting of a metal-enriched, but regionally heterogeneous , metasedimentary protolith. Finally, it is noteworthy that these same compositional features are shared by other Sn-rich granites elsewhere in the world, such as Tasmania (Groves, 1972).

EVOLUTION OF MAGMATIC AQUEOUS PHASES The following discussions on the evolution of the magmatic aqueous phases depend largely on the initial treatment of the subject by Burnham (1979) and further development by Burnham and Ohrnoto (1980) and Burnham (1981b, 1983, 1985). Hence Figure 3.5a,b in Burnham (1979) and Figures 1 and 2 in Burnham (1985), somewhat modified, are presented here as two pairs of illustrations (Figures 3.16 and 3.17) and will serve as reference illustrations for the following discussions on the derivation of the magmatic aqueous phases and its impact on the evolution of hydrothermal ore-forming fluids. In the evolution of the magmatic aqueous phases, two sequential processes play critical roles: first, the process of second (retrograde, resurgent) boiling, in

Evolution of Magmatic Aqueous Pnases

lOS

accordance with the reaction Hj Ovsaturared melt => crystals + aqueous fluid and second, the process of decompression upon the magma following frac ture failure in the roof rocks of the magma chamber. Any granitic magma that contains more than approximately 0.6 wt % H20, which is close to the maximum amount of H 20 that can structurally be bound in hydrous minerals (hornblende and biotite) under magmatic conditions, must undergo the process of second boiling prior to reaching th e Hj Ovsaturated solidus, as indicated by the P-T projection of the divariant 5-L-V surface for the Bingham, Utah, quartz monzonite (BNG-QM, Figure 3.11) in Figure 3 .16b. The precise stage of cr ystallization at which H20 saturation is reached and this process commences, however, depends critically on the initial H20 content of the magma and the depth (pressure) at which crystallization takes place, because the solubility of H 20 in shallow-seated melts (generally less than 10 -km depth), as indicated by the F~v curve in Figure 3.17b, is strongly pressure dependent. To illustrate these relationships, the second boiling process will be applied to the Bingham quartz monzonite magma, Bingham, Utah (Moore, 1978) under the initial assumptions that: (1) the magma, consisting of 90 % melt, 6.0 % Anso plagioclase, and 4.0 % hypersthene, was emplaced at an initial temperature of 1020°C a nd solidified to a minimum depth of approximately 1.9 km (51 in Figure 3.16a); (2) the initial H20 content of th e melt phase was 2. 7 wt % (F~ =0.027), which is the saturation value at 0.5 kbar; and (3) the H20saturated carapace along the 5 I solidus boundary in Figure .L'l ea was essentially impervious to the loss of H20, except very near the margins. Thus, at the stage depicted, the magma body inside 5 I has become essentially closed to fluid escape and a separate aqueous phase exists everywhere inside 5 I where the temperature is below the 5-L-V projection in Figure 3.16b. These 5-L-V temperatures of the Bingham quartz monzonite, it should be noted, are relatively high in comparison with those of other porphyry-Cu porphyries, such as the Hanover granodiorite porphyry (HAN-GDP) discussed above (d. Figure 3.10c) and shown as the dashed 5-L-V projection in Figure 3.16b, as a result of its higher an and much lower qz contents (d. Figures 3.10c and 3.11).

The Second Boiling Process As cooling and crystallization proceed from the walls and roof inward and the solidus 5 I approaches 52 in Figure 3.l7a, the entire volume of magma between 51 and 52 must undergo second boiling and exsolve all of its initial H20 content, except the small a mo unt bound in hydrous minerals. This process of second boiling takes place with an increase in volume of th e entire magma body undergoing second boiling, as may be seen by dividing equa-

o

0-

El

...!ol

I

.- -

c

e::\ -y-

Distance. km

o

,., ~ .~

,. -Rrt.:~; '_ '. '~' .~ 1 · c:::::::=:

700

..I vv i i

900

I

~

Temperature, °C

800

i i i

I i

FIGURE 3.16 (0) Conceptuol model of 0 porphyry stock ot the stoge when the mogmo inside surfoce S1hes become isoloted from the woll rocks ond the mogmo in the H2D-soturoted caropoce, os ot point A, is undergoing second boiling. (b) P-T projection of the crysteljmelt ond crysteljmeltj"vopor" (SH) equilibrium surfoces inthe Binghom quortz monzonite (solid curves) ond the Honover stock (doshed curves) mogmos; the processes occurring ot points Aond A' are discussed in the text.

J:;::>' ~

I

1100

I

~

J.

.&:l

2.0

1.5 ~

J.

~

0.027, because the depth (pressure) at which second boiling commences is directly dependent on F~v' Application of equation (3.26a) to the Bingham quartz monzonite model

Evolution of MogmaticAqueous Phases

109

depicted in Figure 3.16, in which F~v = 0.027, F~ = 0.90, a nd F~~ = 0 .003 at P > 0.5 kbar and near-solidus temper atures, yields th e P, t. Vt relati o ns depicted in Figure 3.17b. These en ergi es represent th e maximum work don e on the surroundings upon complet e cry st alli za tion o f th e Bingh am qu a rtz monzonite magma , provided th e tensile strengt h of the w a ll/roc ks is high enough to absorb this work with out deforming to fra cture fa ilure. That this restriction is not fulfill ed at depths of less th an 4-5 km is evide nt in Figure 3.17b, where P, t. Vt exceeds th e maximum tensile st rength (MTS) o f the ign eous wall rock s (Shaw, 1980; Schock and Loui s, 1982). This maximum tensile strength, which is eq uivale nt to 100 bar s, is for a homogen eous, unfractured rock with a density of 2.65 g/ cm', a large mass o f whi ch would be difficult to find in a subvolcanic intrusive syste m. For th e present, however, it will be assumed that MTS pr evails and that BNG-QM magma at point A in Figures 3.16 a nd 3. 17 (T = 920 0 e = 1193 K, P, = 0.6 kba r = 6 x 1Q8ergs/cm3) lies on the S-L-V projection a nd hence is just saturated with H20(F~ = 0.030) at an F~l of 0.81. Upon cooling and further cry stallization at point A in Figure 3.1 6a, second boiling produces an aqueous phase th at results in an increase j n volum e and concomitant increase in internal pr essure in th e BNG-QM magma th at is confined by rigid wallrocks. Upon reaching po int A' in Figure 3.16b, where the internal pressure has incr eased 100 bars to overcom e th e maximum tensile strength (MTS) in Figure 3.17b, eq uations (3.26a-c), in co njunctio n with a calculated equilibrium temperature from th e qu asicrystalline model of 820 o e, indicate that F~v, th e ma ss fraction of H 20 in th e " va po r" phase, will have reached 0.015, which is approximately 63 % of th e o rigina l H 20 content of th e magma (F~v = 0.024); henc e th e amount of melt remaining, in which F~v = 0.034 at 0.7 kb ar, represents approx imat ely 26 % of th e total magma. For comparison, in order for the second boiling process to ge ne rate the 100-bar ov erpressure nece ssary to overcome MTS a t 1.0-kbar Ph a total of 88 % of the magma must be cr ystalline-hen ce th e proximity o f th e P, t. V t curve to MTS in Figure 3.17b at thi s pressure. Increasing the initial H20 content of the melt to 4 .0 wt % ( F~v = 0.040) , which is th e saturation value at 1.0 kbar, increases the pr essure o f int ersectio n of the Pt t. Vt curve with MTS in Figure 3.1 7b to approximately 1.9 kb ar, thus placing an approximate upper limit on the dep th at which th e second boiling process can produce extensive fracture failure, primarily by lat eral stre tching in th e roof ro cks (d. Burnham, 1985). Beyond thi s de pth, incidentally, the second boiling process is commonly expressed in th e occurrenc e of miarolitic cavities and in pegmatites.

The Decompression Process Failure of the roof rocks of th e H j Ovsaturared carapace pr oduces myr iad steeply dipping fractures , because th e axis o f least princip al stres s in thi s

11 0

Magmas and Hydrothermal Fluids

sha llow crustal environment generally lies essentially in the horizontal plane. Many of these fractures may extend downward short distances into the upper part of the largely cry stalline Hj Ovsaturated carapace, as well as upward into the rigid roof rocks, where fluid pressures equivalent to at least one-third lithosratic pressure (Poisson's ratio = 0.25) are required to prevent the fractures from closing (Turcotte and Schubert, 1982). To prevent the fractures from extending farther toward the surface, moreover, the fluid pressure cannot exceed approximately one-third of the lithosraric pressure plus MTS (100 bars) at the top of the fracture. Following fracture failure at point A' in Figure 3.16b and return of P, temporarily to lirhostatic pres sure , the decompressional exsolution of more H 20 from the interstitial melt causes rapid crystallization (" pressure quench"; Jahns and Tuttle, 1963) of a small additional amount of melt. Of much greater importance, however, is the decompressive escape of the magmatic aqueous pha ses into the fracture system, where the fluid pressure (Pf) can be no greater than 300 bars (0.33 x 600 + 100 = 300). That a fluid pressure of 300 bars in the fracture system is reasonable is attested to by the common occurrence of fluid inclusions in hydrothermal quartz of several breccia pipes that record entrapment pressures of 300 bars and higher (Sillitoe and Sawkins, 1971; Reynolds and Beane, 1985). Thus, in the vicinity of the pos tulated incipient fracture at point A in Figure 3.17a, where the initial P-T conditions are as shown at point A' in Figure 3.16b, the expansion of already exsolved H20 bubbles into the fracture releases approximately 4.3 x 10 10 ergs/kg of magma, in accordance with the relation [ef. Burnham, 1985, equation (4)1

(3.27) of magma, where Pfn and Pin are the final (fn) and initial (in) pressures, respectively, in ergs/cm-', .1 V, is the volume expansion of H20 (v, steam) upon decompression from Pin = 7 x 10 8 to Pfn = 3 x 108ergs/cm3, F~vin = 0.015 is the mass fraction of H20 initially present as bubbles, and T is tem perature in kelvins. The coefficient, 4.61 x 10 9 ergs/kg ' K, is the gas constant (R) for H20 [ef. equation (3.26a)]; impurities such as alkali and metal chlorides, at levels found in fluid inclusions, may lower this coefficient; hence th e amount of energy released may be reduced by as much as 15 %, but other irnpuritiesv s uch as C02 and S02, may raise it by a corresponding amount. In addition to the energy released into the fracture at point A, Figure 3.1 7a , by expansion of already exsolved H20, decompression of the magma in the vicinity of the fracture to 300 bars carries the magma across the H20saturated solid us, as indicated by the arrow emanating from point A' in Figure 3.16b. As a consequence, the decompressed magma undergoes complete " pressure quench " and concomitant exsolution of all remaining H2 O. The maximum amount of energy that can be released by this latter process can

..

Evolution of Magmatic Aqueous Phases

111

be closely estimated from equation (3 .26a ) by replacing F~l with F m (0.26 ) and neglecting the last term, which is applicable only in the case of equilibrium crystallization during the second boiling process. The P ~ V energy thus calculated to have been released in this process is 4.1 x 10 10 ergs/kg of magma and, when combined with that released in the expansion of already exsolved H20, yields a total of approximately 8.4 x 10 10 ergs/kg of magma that is carried into the fracture by only 24 grams of H20. Altogether, then, a maximum of approximately 1.2 x 1011 ergs of P ~ V energy has been released (work done on the surroundings) by just 1 kg of magma at point A, Figure 3.17a, from the onset of second boiling. As an aid in placing these energies into more familiar geological perspective, the 8.4 x 10 10 ergs/kg of magma, which may be released into the fracture in a matter of seconds (Burnham, 1983), is approximately twice the estimated average kinetic energy expended, per kilogram of material erupted, in explosive volcanic eruptions (Burnham, 1972). It is also more than twice the energy released, per kilogram of magma erupted, in the Mount St. Helens blast of May 18, 1980 (Eichelberger and Hayes, 1982; Burnham, 1983, Figure 3). Perhaps more in the perspective of the present discussions, this amount of energy is sufficient to lift an equivalent mass of rock 900 m in the absence of frictional resistance. Thus the emplacement of 80-ton blocks of rock almost 500 m above their original positions in the breccias of the Warren (Bisbee) district, Arizona (Bryant, 1968), is not difficult to envisage. Following the initial failure of the wall and roof rocks, with all of the attendant energy release and resultant fracturing just discussed, the Hj Ovsaturated carapace retreats downward, as in Figure 3.17a. This retreat pauses temporarily, however, when the fluid pressure in the downward penetrating fractures becomes less than one-third the confining lithostaric pressure and the system once again becomes essentially closed to the escape of volatiles. At this stage, the second boiling process again builds up internal overpressure, but the amount of overpressure required to produce wallrock failure in the already weakened carapace is much less than the MTS in Figure 3.17b. As a consequence, fracture failure is confined in large part to reopening preexisting fractures, into which effervescing magma may be intruded as intramineral dikes (D2 in Figure 3.17a), as at Bingham, Utah (Wilson, 1978), and to the formation of breccia pipes (BP2 in Figure 3.17a), which are characteristic features of most porphyry systems. This episodic buildup of internal pressure, followed by renewed fracturing, may occur several times, as attested to by the very common occurrence of cross-cutting veins of differing mineralogy in many porphyry systems. This downward extension of fractures eventually terminates, however, when the increasing lithostatic load pressure exceeds approximately three times the fluid pressure in the fractures, as mentioned previously, Coincident with the termination of fracturing, of course, is effectively the termination of the mineralizing potential of a given porphyry system, unless fresh magma intrudes the system to depths shallower than 7 or 8 km. The end product of

112

Magmas and Hydrothermal Fluids

these rep eated pr ocesses of seco nd boiling, fracturing, a nd decompression is most co mmo nly a chimney- like fracture system th at tends to be cente red on th e porphy ry stock and serves as th e co nduit through which th e magm at ic aqueo us ph ase ascends, produces submagmatic hydrothermal al te ra tio n, a nd, under th e right co nditions, ultimately pr oduces a porph yry-type ore dep osit. Thus, to close thi s cha pter, some of the major features of th e magmati c aq ueo us ph ases wilI now be briefly exa mined.

Important Major Features of the Magmatic Aqueous Phase(s) The generatio n of a magm atic aqueo us ph ase or ph ases by th e foregoing processes is acco mpanied by partitioning of all elements in th e syste m such that th e chemical potenti al or fugacity of each chemical species is th e sa me in all phases at equilibrium. Th e volatile element chlorine, which, contra ry to ea rlier sta tements (Burn ha m and Ohmoto, 1980, p. 5), appears to dissolv e in silicate melts mainl y as alkali chl oride complexes, is partitioned strongly into the aqu eous phase in metaluminous to peraluminous magmas (Kilinc and Burnham, 1972; Webster and HolIoway, 1990) because: (1) it is unable to sequester much of the alka li metals from th e ab and or aluminosilicate melt compo nents, and (2) in aqu eou s solutions at magmatic temperatures and low to moderate crustal pressures, it forms highly stable, neutral chloride com plexes with hydrogen, alk ali metals, alkaline earths, and heavy met als. Fluorine also forms stable neutral fluorid e complexes in th e magmatic aqueous phase , but it also forms stable fl and cr melt species (see section entitled M elt Species and Coexisting Cr ystalline Phases) as well. As a consequence, F partitions approximately 5: 1 from th e aqueous solution toward th e melt (Burn ham , 196 7; Hards , 1976) and into fluorite, topaz, a nd th e micas, whereas CI partitions approximately 40: 1 toward the aqueous phase from th e melt (Kilinc and Burnham, 1972) und er the conditions of int erest here. In peralkalin e magmas, on the other hand, a hali te-like melt component appears to be mod erately stabl e and, thu s, to substantialIy lower th e preference o f Cl for the magmatic aqueous phase. Throughout alI of th ese discussions on hydrous magmas and th eir exsolved aqueous pha ses, the assumption has been tacit that the volatile contents of the magmas were obtained by th e initial fluid-absent partial melting o f th e magma source rocks . Also, inasmuch as a ll of the porphyry magm as ana lyzed in the earli er parts of this chapter wer e found to have gen etic ties to th oleiitic amphibolites, the assumption is implicit that th eir volatile m akeup a lso ori ginated in th ese rocks . Hence, from scores of analyses of hornblende pr esented in Leake (1968) and elsewhere, the average mass fraction of CI is approximately 0.03 F~, where F~ is the mass fra ction in th e melt, F of H20 in th e melt. This average Fe) is less than 50 % gr eater than that in sea wa ter a nd, considering th e range in tabulated valu es, suggests that th e so urce of th e volatiles in the Cu-Au and Mo porphyry ma gmas co nside red

el,

Evolution of Magmatic Aqueous Phases

113

here was seawater, which is consistent with a subducted tholeiite so urce for these magmas. Multiplication of th e 0.03 coeffi cient a bove by 40, wh ich is roughly the fluid/melt partition coefficient {Of/ mc il for Cl in po rph yry melts (Kilinc and Burnham , 1972), and conversion to Na CI mass equivalents, yields the simple relation F~a Cl = 2.0 F~ . Thus, in the magmatic aqueo us pha ses of our previous example of th e Bingh am qu artz mon zon ite magma at po int A in Figure 3.16b, where F~~ = F~v = 0.030 (the beginn ing of seco nd boi ling), the bulk F~a CI = 0.06 or 6.0. wt % N aCI equivalent. Th is co mposition, tog ether with the temperature and pre ssure of po int A, Figure 3.16 b, also is sho wn as point A in the isobaric T-X projection of th e system NaCl-H20 in Figure 3.18, as adapted from Bodnar et al. (1985, Figure 10 ). In Figure 3.18, under th ese P-T-X co nditions, second boiling produces two aqueous solutions, which, according to the lever principle, consist of 7.0% brine co ntai ning 78 wt % NaCl eq . and 93 % "vapor" containing only 1.0 wt % NaCl eq . . It wi ll also be observed in Figure 3.18 (point B) that upon initiation of second boiling at 1.0 kbar (Figure 3.16b, SlO°C) , F~a CI has increased to 0.085 (8.5 wt % NaCleq . ) and two fluid phases are again produced-11.5% of the aqueous fluids containing 53 wt % Na Cl eq . and th e

900

00

...eli::::l ...

Cl:I Cll Q.

800

700

E

e

600

500

50

100

Weight Percent NaCI T- X projections of the system NaCI-H20 at pressures between 500 and 1750 bors. Point Arepresents the total CI content of the fluid phesets). in NaCI equivalents, that exsalves from the melt at point A(600 bersl in Figures 3.16 and 3.17; point Brepresents the some Iotal CI that exsalves ot 1000 bors, and point Cthat which exsolves at 2000 bors. Tie lines connect coexisting brines and H2O-rich ·vopar" phases.

Figure 3.18

114

Magmas and Hydrothermal Fluids

remaining 88. 5% " va por" containing 2.0 wt % N aCl cq . . At 2.0 kb ar, 765°C (po int C, Figure 3.18) , on the other hand, second boiling produces on ly a single phase th at contains approximately 12.5 wt % NaCl cq . . Accordingly, th e information contained in Figur e 3. 18 suggests that th e pr essure boundary between a single and tw o-phase separa tion of the magmatic aqueous phase(s) from the Bingham quartz monzonite magma lies at approximately 1.4 kb ar, but this ass umes that the mix ing properties of other chloride complex es, espec ially those of th e heavy metals, are closely similar to those of NaCI. The ch ances are th at the se heav y metal chloride complexes, which can account for as much as 20 % or mor e of the total CI (Burnham and Ohmoto, 1980), would rai se the pres sure of immiscibility, but by precisely how much is not now kn own. This caveat notwithstanding, the brin e-"vapor " immis cibility relations depicted by the isobaric tie lines in Figure 3.18 for th e Bingham quartz monzonite magmatic aqueous phases agree with the compositions of coexisting brine-rich a nd " vapo r" -rich fluid inclusions in quartz from th e "co re" of th e pluton, some of which hom ogenize at T> 700 °C (Roedder, 1971). Further decompression of alr eady exsolved fluids following second boilinginduc ed fracture failure at point A' in Figure 3.16b promotes boiling o f the brin e and furth er enrichment of metal chloride complexes in it. Also, the brine component of the fluids that exsolve from th e residual melt at A' upon pressu re quench are simila rly enriched in these same chloride complexes. Just how rich these brin es can become in heavy metal chlorides, such as those of Cu, depends mainly on: (1) the initial metal content of th e ma gma source rock; (2) th e degree of partial melting of the source rock and the sm all amounts of the metal remaining in the restite; (3) the extent of cr ystal fractionation that has occur red pr ior to high-l evel emplacement of the magma; and (4) the degree of crystallinity of the magma at th e time second boiling co mmences. To place the factors affecting metal partitioning in perspective, in terms of th eir impact on th e end result, onc e again the Bingham , Ut ah , porphyry system will be called upon , although it should be recognized th at enrichment factors in the Bingham quartz monzonite magmas may be somewhat larger th an average. Thus choosing Cu as an example, the original so urce rock for the Bingham quartz monzonite magma, presumably a subd ucted a m phibo litized th oleiite, passes through five stages of enrichment. The initial tholeiite co ntai ns an average of - 75 ppm Cu (Engel et aI., 1965); hen ce melting of 32% of " norma l" thol eiite (NR -THL, Figure 3.9a) at 24 kbar yields, during th e first stage of enr ichment , a melt that contains -230 ppm C u, assuming a ll Cu is partitioned into the melt . Ina smuch as 43 % of th e melt is lost to cr ystal fractionation during ascent and emplacement of thi s magma, th e Cu content of the melt of Bingham quartz monzonite composition in th e second stage of enr ichment will have reached -400 ppm. Then, upon co oling and 19 % furth er crystallization of th e melt to reach H20 saturation and commencement of second boilin g at point A in Figur e 3.16 b, th e Cu content of th e melt , whi ch now repre sents only 15 % of th e original subd ucted " nor-

Evolution of Mogmotic Aqueous Phases

115

mal " thol eiit e, has increased in the thi rd stage of enrich ment to a to ta l of - 490 ppm. As second boi ling commences at this point, a partition coeffi cient of 9 : 1 in favo r of the magma tic aqueous phase(s) (Candella and H o Iland, 198 4) places, in the fourt h stage of enrichment, -4400 ppm (0.44 wt %) Cu in th e aqueous fluids, more than 90 % of which is concentrated in the imm iscible, more sa line 7.0 % bri ne, thereby yielding, by th e fifth stage of enric hmen t, a so lution tha t contains - 6.0 wt % Cu . Th is Cu content is much grea ter than tha t fo un d in any fluid inclusions from porphyry copper deposits (R. J. Bodnar, person al co mmunica tion, 1984); hence it may be reaso na bly co ncl ude d th at th e efficiency of one or more of these co ncentration pro cesses is less than 100 % . The fact th at approximately 20 % jadei tic pyroxene plu s hypersthene a nd 5. 0% magnet ite plus ilmenite were left in the so urce restite at 24 kbar suggests tha t much of the initial Cu co ntent may have been removed in th ese min eral s. N evertheless, the implicat ion here is th at the Bingham quartz monzonite magma conta ined amo unts of Cu fully adeq uate to mak e perhaps th e world 's largest porphyry Cu-Au deposit. Moreover, th ese sa me large enrichment factors doubtless co ntributed substa nt ially to making the Bingham Mine th e largest producer of Au in the Unite d Sta tes during World War II (J. A. No ble, person al co mm un icatio n, 1953) . Iron chlorides acco unt for 20 % or mo re of the tota l aq ueous ch loride conte nt at point A in Figu re 3. 16b (1.1 molal BCl f), based on the experimental results discu ssed by Burnham (19 79) and grap hed in Burnham and Ohmoto (1980, Figur e 6). T his percentage is calculated from an observed bulk stoichio me try of FeC12.67 at 750 °C, 2.0 kbar, an d an Io, in the upper part of th e magn et ite field. At the somewha t lower f0 2 values that are perhap s more ap pro priate for most porphyry systems, the stoichiometry like ly wo uld be closer to FeCI2 and th e percentage of CI complexed with Fe would be co rres pond ingly lower. In either case, Fe constitu tes the thi rd most abun dant ca tionic con st itu ent in the magmat ic aq ueo us phase(s)-afte r NaC lv a nd KClf-at a tota l Fb > - 0.030 (3.0 w t % CI). Small wo nde r, then, that a ma jority of po rphyry-type o re dep osits are dom inat ed by pyrite (FeS2), th e deposition of which releases CI fro m the Fe-ch loride co mp lexes and H from the S-bearing complexes to produce H CI, whi ch in turn produces the spatially ass oc iated, but low er temperature, ph yllic (sericitic) alte ra tion. Corroborating evide nce of th e Fe- CI-H reacti on s is found in on e o f th e most pyritic porphyry Cu deposits observed by the writer- Santa Rita, N ew M ex ico-which is also one the most sericitica lly altere d. In associated reactions, th e bulk of th e primary cha lcopyrite-bo rn ite ore at Santa Rita occ urs inward o f th e seric itic (phyllitic) zone, nea r th e diff use bo un dary with the porassically altered zo ne, w here the su bmag ma tic hydrolysis of S02,

(3.28) produces much of th e sulfide sulfur necessa ry to precipitate th ese ore miner-

116

Mogmos ond Hydrothermol Fluids

al s, The precipitation of both sulfides and sulfate (anhydrite) from t he subv magmatic Cu-, Fe-, and Ca-chloride solutions also produces the HCl that is a major cause of sericit ic alteration. In contrast to the brine phase, which carries most of the chlorides of the a lkalies, alka line earths, precious meta ls, and most of the base metals, the much more voluminous "vapor" phase contains most of the Mo as molybdate (Ca nde lla a nd Holland, 1984) and, by analogy, most of th e W as tungstate, as well as th e bulk of the silica as H4Si04, and sulfur mostly as S02 in relatively ox id ized I-type porphyry Cu-Au-Mo magmas and as H2S in the much more reduced S-type Sn-W magmas. This "vapor" phase also contains essentially all of the C02 in the system, which, although not abundant in mo st porphyry systems, promotes brine-"vapor" immiscib ility at somewhat higher pressures than depicted in Figure 3.18. The occurrence of Mo in the " va po r " phase, if mainly as H2Mo04, implies that the pa rtition coefficient for Mo (D%,~), w hich is approxi mately 3.0 at 1.4 kbar (Candella and Holland, 1984), should be, like si lica (H4Si04 ), Hj Ovpr essure dependent. Thus fracture failure and decompression of the exs ol ved " va po r " ph ase in a Climax-type porphyry Mo system should result in coprecipiration of quartz and molybdenite, as it apparently does. The highly silicified nature of the ore-bearing porphyry at Climax, Colorado, as co m pa red with th e av erage porphyry Cu ore body, for example, has long been recognized. It also must be recognized, however, that the Climax porph yry magma was more felsic (silicic) initially. The comp lexing of CI w it h Fe hav ing been discussed above, it is now appropriate to examine the distribution o f C I a mong NaCI, KCI, a n d HCI in magmatic fluids, mainly as a function of th e magmatic phase assemblage present (d. Burnham, 1979, pp. 120ff). Thus in the Bingham quartz monzon ite magma at temperatures up to -50°C below point A in Figure 3.16b, wh ere 1.1 molal total elf fluids are in equilibrium with plagioclase, hyper= 0.49 a nd sthe ne, magnetite, and melt, in the melt X ~'i'/( x~~n + X a X ~C1/(X~C1 + X ~aCI ) (w here the superscript f refers to all fluid phases at eq uilib rium) also eq ua ls 0.49 fKilinc, 1969; Burnham, 1979 ). These same fluids also have 0.1 molal HCI , most of which is in rhev va po r " phase and capable of causing substantia l aluminum si licate a lteration should it escape into the wa llrocks. Upon further cooling of the Bingham magma and t he appearance of horn blende, in which Na " and OH - are essential constituents, a~a CI and af.,C' becom e " buffered" at substantially lower values than previously. As a consequen ce? X kcl increas es to satisfy the total chloride stoichiometry of the fluid s and X K CI/(X ~CI + X ~aCI) reaches as high as 0.75 (Kilinc, 1969 ). By these processes, then, th e fluids that almost certainly produce the inn er ore-r elated potassic alteration at Bingham, Utah, and in porphyry Cu-Au deposits the world over a re produced. In S-type magmas associated with Sn-W (wolfram ite, huebnerite) minera lization, the princip les a re t he same, b ut th e res u lts are d ifferent . O w ing to

b)

In RetlOspect

117

their peraluminous nature, di and hd are not melt components; hence, instead of hornblende, biotite precipitates upon cooling of these magmas. Thus at temperatures above the first appearance of biotite, the controls are essentially the same as in the l-type magmas, such as that of the Bingham quartz monzonite, where X~CI/(X~CI+Xf aCI)::: X ~~l/(~~l+X~ bl) and HCl f ::: o.irci", again promoting the formation of topaz. Within the stabili7 field of biotite, however, a~CI' ins tead of a f scr- becomes buffered and X NaCi increases to satisfy the total chloride stoichiometry of the fluids; hence X~ C1/(X~ C1 + X~aCI) decr eases to as low as - 0.25 . As a consequence, fluids emanating fro m t hese magm as produce albitization even of K-feldspar at submagmatic tempera t ures, th ereb y beco me enric hed in xci', and eventually cool into the fi eld of sericite stability. The end-pro duct of these processes, then, is greisenizatio n, w hich is very preva lent in th e tin-bea ring granites of the Cornubia n batholith discu ssed ea rlier in th is cha pter. Another important differ ence betwee n th e l-type porp hyry Cu-A u a nd Mo, a nd th e S-t yp e Sn-W magmatic aq ueous ph ases is the oxida tio n sta te, which in turn affects the speciation of sulfur in th e mag matic aqueous phases. In relativ ely oxidize d l-rype magma tic fluids, S resides mainly in S02, which is concent ra ted in the "vapor" phase and can coexist with metal-rich br ines at magmatic temper atur es, but not at lower temperatures, as discussed previously. In th e more red uced S-type magmatic fluids, however, S resides largely in H2S, which promotes the stability of sulfides (mostly pyrrhotite) at the lowe r magmatic tempe ratures (Figure 3.8, point D) and effectively purges much of the sulfur from the cassiterite- and wolframite-bearing S-type Sn-W systems. Concomitantly, the Fe-chloride content of the magmatic aqueous phases is greatly redu ced, thereby yielding the Fe-poor tin greisens, and even the China clays, of the Cornubian batholith systems.

IN RETROSPECT Despite th e fact th at space has restricted th e pursuit in de tail of th e many to pics discu ssed in thi s chapter, it is the hop e of the a utho r th at enough has been p rese nted on th e critica l issues to guide th e read er into a mo re th orough investi gation o f th ose issues of inte rest. Appl icat ion s of th e q uasicrysta lline mo del to th e orig in of min eral izat ion-relat ed magmas have been develop ing up to submiss io n o f th e ma nuscript of th is chapter a nd th e implica tions of th e res ults a re still under assessment.

ACKNOWLEDGMENTS I especia lly wish to thank Professor H. Nekvasil for development of the origina l Fortran program that made possible t he initial application of the qua-

118

Magmas and Hydrathermal Fluids

sicrysta lline model to calculation of the equilibrium relations in magmas; I am also grateful to Professor J. R. Holloway for expanding and greatly reducing the run time of this program. A debt of gratitude is owed to Dr. B. W. Chappell for generously providing analytical data on the S-type granites of the Cornubian batholith prior to publication . This presentation has benefitted greatly from th e critical comments of Professors H. L. Barnes and J. R. Holloway.

REFERENCES Allen, J. c., and A. L. Boettcher (1978) Amphiboles in andesites and basalts: II. Stability as a function of P-T-fHzo-fo z: Am. Mineral. 63, 1074-1087. _ _ , P. J. Modreski, C. Haygood, and A. L. Boettcher (1972) The role of water in the mantle of the earth: the stability of amphiboles and micas: 24th Int. Geol. COl/g. Section 2, 231-240. AIt, J. c., and J. Honnorez (1984) Alteration of the upper oceanic crust, DSDP site 417: mineralogy and chemistry: Contrib. Mineral. Petrol. 87, 149-169. Anderson , O. (1915) The system anorrhire-forsterire-silica: Am.]. Sci. 39,407-454. Anthony, E. Y., and S. R. Titley (1988) Progressive mixing of isotopic reservoirs during magma genesis at the Sierrita porphyry copper deposit, Arizona: inverse solutions: Geochim. Cosmocbim, Acta 52, 2235-2249. Bateman, P. c., F. C. W. Dodge , and P. E. Bruggman (1984) Major oxide analyses, CIP\X' norms, modes, and bulk specific gravities of plutonic rocks from the Mariposa 10 by 2° sheet , Central Sierra Nevada , California: U.S. Geol. Suru. Open File Rep. 84-162. Blencoe, J. G. (1974) An experimental study of muscovite-paragonite stability regions: Unpublished Ph.D. thesis, Stanford University. Bodnar, R. J., C. W. Burnham, and S. M. Sterner (1985) Synthetic fluid inclusions in natural quartz. III. Determination of phase equilibrium properties in the system H20-NaCl toIOOO°C and 1500 bars : Geochim. Cosmochim. Acta 49, 1861-1873. Boettcher, A. L. (1970) The system CaO-AI203-Si02-H20 at high pressures and temperatures: j. Petrol. 2, 337-379. _ _ , (197 1) The nature of the crust of the earth, with special emphasis on the role of plagioclase: Geoplrys . Monogr. Ser. 14, 875-909. _ _ , (1973) Volcanism and orogenic belts-the origin of andesites: Tectonophysics 17, 223-240. _ _ and P. J. Wyllie (1969) Phase relationships in the system NaAlSi3O g-Si02-H20 to 35 kilobars pressure: Alii. j. Sci. 267, 875-909. _ _ , Q. Guo, S. R. Bohlen, and B. Hanson (1984) Melting in feldspar-bearing systems to high pressures and the structures of aluminosilicate liquids: Geology 12, 202-204. Boyd, F. R. (1959) Hydrothermal investigations of amphiboles: In: Researches in G eochemistry, P. Abelson (ed.). New York: Wiley, pp. 377-396. Brown, P. E., J. R. Bowman, and W. C. Kelly (1985) Petrologic and stable isotope

References

119

con straints on th e so urce a nd evo lutio n of ska rn -forming fluids at Pine Creek, California : Econ, C eol. 80 , 72-95. Bryant, D. G. (1968) Intrusive breccias associated with ore, Wa rre n (Bisbee) mining district, Arizona: Econ, Ceo l. 63 , 1-12. Burnh am , C. Wayn e (196 7) H ydrothermal fluids at the magm at ic stage : In: Ceo chemistry of Hydrothermal Ore Deposits, H. L. Barn es (ed.). New York: Holt, Rin ehart, and Winston , pp. 34-76. _ _ (1972) The energy of ex plosive volca nic eruptions: Earth Miner. Sci., Th e Penn sylvania State University, 41 , 69-70 . _ _ (1974) The th ermodynam ics of melt ing in experimental silicate-volatile systems: Fortschr. Mineral. 52 , 101-118. _ _ (19 75) Water a nd magm as; a mixing mod el: Geochim. Cosmocbim. Ac ta 39 , 1077-1084. _ _ (1979) Magmas and hydrotherm al fluids: In: Geochemistry of Hydroth ermal Or e Deposits, H. L. Barnes (ed.) . N ew York: Wiley, pp . 71-136. _ _ (1981a) The nature of multicomponent aluminos ilicate melts: In: Chemistry and C eoch emistry of Solutions at High Temperatures and Pressures, D. T. Rick ard and F. E. Wickman (eds.) . Phys. Chem, Earth 13/14, 191-227. _ _ (1981 b) Converg ence and mineralization-Is there a relat ion ?: Ceo l. Soc. Am. M em. 154, 761-768. _ _ (1983) Deep submar ine pyroclastic eruptions: Econ. Ceol. Mo nogr. 5, 142-148. _ _ (1985) Energy release in subvolcanic environments : implications for breccia formation : Econ. C eol. 80 , 151 5-1522 . _ _ (1992 ) Calculat ed melt and restite compositions of some Australian granites: Trans. Roy. Soc. Edinburgh Earth Sci. I 83, 387-3 97 . _ _ and N. F. Davis (1971) The role of H 20 in silicate melts: I. P-V-T relat ions in the system NaAIS i30S-H20 to 10 kilob ars and 1000°C: Am. J. Sci. 270, 54-79 . _ _ a nd (1974) Th e role of H 20 in silica te melts: II. Thermod ynamic a nd ph ase relations in th e system N a AISi30S-H20 to 10 kilobar s, 700° to 1l00°C: Am. J. Sci. 274, 902-940. ___ , J. R. Holloway, and N. F. Davis (1969) Th erm od ynami c pr operties of water tolOOO°C and 10,000 bars: Ceol. Soc. Am. Spec. Pap. 132. _ __ a nd H. Nekvasil (198 6) Equ ilibrium propert ies of gra nite pegma tite magm as: Am. Min eral. Jahns Mem. Issue 71, 239-263. _ _ and H . Ohmoto (1980) Late-stage pr ocesses of felsic magm ati sm : Mining C eol. (jpn.) Spec. Issue 8, 1-11. Candella, P. A., a nd H. D . Holland (19 84) Th e partition ing o f copper and mol ybdenum between silica te melts and aq ueous fluids: C eochim. Cosmocbim , Ac ta 48, 373-380. _ _ and (1986) A mass transfer model for co pper a nd mol ybd enum in ma gmatic hydrothermal systems: th e origin of porphyry-typ e ore depo sits: Econ. C eol. 81 , 1-19. Carmichael, 1. S. E., F. J. Turne r, and J. Verh oogen (1974) Igneous Petrology: New York: McGraw-HilI. Carten, R. B., E. P. Gera ghty, B. M. Walker, and J. R. Shannon (1988) Cyclic develop-

120

Magmas and Hydrathermal Fluids

rncnr of igneous features and their relationship to high-temperature hydrothermal features in the Henderson porphyry molybdenum deposit. Colorado: Econ. C eol. 83, 266-296. Chappell, B. W., A. ] . R . White, and D. Wyborn (1987) The importance of residual so ur ce material (resrire) in granite petrogenesis: j. Petrol. 28, 1111-1138. Dingwell, D. B., D. M. Harris, and C. M . Scarfe (1984) The solubility of H20 in the system Si02-Ah03-Na20-K20 at 1 to 2 kbar:]. Ceol. 92, 387-395. Eggler, D. H ., and C. W. Burnham (1973) Crystallization and fractionation trends in the system andesite-H20-C02-02 at pressures to 10 kb: BIIII. Ceol. Soc. Am. 84, 2517-2532. _ _ and M. Rosenhauer (1978) Carbon dioxide in silicate melts: II. Solubilities of C02 and H 20 in CaMgSi206 (diopside) liquids and vapors at pressures to 40 kb: Alii. ]. Sci. 287, 64-94. Eichelberger,] . c., and D. B. Hayes (1982) Magmatic model for the Mount St. Helens blast of May 18, 1980: j. Geopbys. Res. 87,7727-7738. Engel, A. E. ]., C. G. Engel, and R. G. Havens (1965) Chemical characteristics of oceanic basalts and the upper mantle: BIIII. Ceol. Soc. Am. 76, 719-734. Erikson, R. L. (1979) An experimental and theoretical investigation of plagioclase melt ing relations: Unpublished M.S. thesis, The Pennsylvania State University, University Park. Goldsmith , J. R. (1980) The melting and breakdown reactions of anorthite at high pressures and temperatures: Am. Mineral. 65, 272-284. • Green, T. H. (1969) High-pressure experimental studies on the origin of anorthosite: Can. ]. Earth Sci. 6, 427-440. Groves, D. I. (1972) The geochemical solution of tin-bearing granites in the Blue Tier batholith, Tasmania: Econ. Ceol. 67, 445-457. Gustafson, L. B., and J. P. Hunt (1975) The porphyry copper deposit at El Salvador, Chile: Econ. Ceol. 70, 857-912. Hamilton, D. L., C. W. Burnham, and E. F. Osborn (1964) The solubility of water and effects of oxygen fugacity and water content on crystallization in mafic magmas: j. Petrol. 5,21-39. Hards, N.] . (1976) Distribution of elements between the fluid phase and silicate melt phase of granites and nepheline syenites: NERC Rep. Prog. Exp. Petrol. 3, 88-90. Helz, R. T. (1973) Phase relations of basalts in their melting range at PH20 = 5 kb as a function of oxygen fugacity : Part I. Mafic phases: j. Petrol. 14, 249-302. Hill, R. E. T., and A. L. Boettcher (1970) Water in the earth's mantle: melting curves of basalt-water and basalt-water-carbon dioxide. Science 167, 980-981. Holloway,]. R. (1973) The system pargasite-Hj Oi-Cfjj : a model for melting of a hydrous mineral with a mixed-volatile fluid-I. Experimental results to 8 kbar: Ceochim. Cosmochim . Acta 37, 651-666. _ _ and C. W. Burnham (1972) Melting relations of basalt w ith equilibrium water pressure less than total pressure: j. Petrol. 13, 1-29. Huebner, J. S., and A. C. Turnock (1980) The melting relations at one bar of pyroxen es composed largely of Ca-, Mg-, and Fe-bearing components: Am. Mineral. 65,

225-271.

References

171

Jahns, R. H ., and O. F. Tuttle (1963) Layered pegmatite-aplite int rus ives: Mineral. Soc. Am. Spec. Pap . 1, 78-92. Jones, W. R., R. M . Hernon, and S. L. Moore (196 7) General geology of the Santa Rita quadrangle, Grant County, New Mexico: U.S. C eol. Sur o. Prof. Pap . 555. Joyce, D. (1985 ) A phase equil ibr ium study in the system NaAlSi30S-S i02-AI 2SiO sH20 at 2 kilo bar s and petrogenetic implications: Unpublished M.S. thesis, The Pennsylvania State University, University Park. Kennedy, G. c., G. J. Wasserberg, H . C. He ard, and R. C. Newton (1962) The upper three-phase region in the system Si02-H20: Am. J. Sci. 26 0, 501 -521. Kilinc, 1. A. (19 69) Experimental metamorphism and anatexis of shales and graywackes: Unpu blished Ph. D. thesis, T he Pennsylvania State University, University Park. _ _ and C. W. Burnham (1972) Partition ing of chloride betwee n a silicate melt a nd coex ist ing aqu eo us ph ase from 2 to 8 kilo bars: Econ. Ceol. 67, 23 1-235. Kuno, H. (1950) Petrology o f Hak on e volcano and adjacent areas, Jap a n: Bull, C eol. Soc. [p n. 56, 79-8 3. Kushiro, I. (1972) Determinat ion of the liqu idu s relations in synthetic silica te systems: th e system forst erite- diopside-silica at one atmosphere : Am. Mineral. 57, 1260-1271. _ _ and H . S. Yode r, J r. (1965) T he reac tio ns betw een fors terite and a northite at high pressures: Carnegie Institution Washingtoll Yearbook 65, 89-94 . Larso n, E. S., J r., an d W. Cross (1956 ) Geology and petrology of the San Juan region, so uthwestern Colorado: U.S. C eol. Surv. Prof. Pap . 258 . Leake, B. E. (1968) A ca ta log of analyzed calciferous and subc alciferous amphiboles togethe r with th eir nomenclatu re and associated minerals: C eol. So c. Am. Spec. Pap . 98. Lut h, R. W., a nd G. E. Munci ll (1989) Fluorine in alum inos ilicate systems: phase rela tio ns in the system NaAlSi3Os-CaAI2Si2OS-F20 _1 : Geochim. Cosmochim. Acta 53, 1937-1942. Macdonald, G. A., an d T. Katsura (1964 ) Chemical composition of Hawaiia n lavas: J. Petrol. 5, 82-133. Mason , D. R., and J. A. M cD onald (1978 ) Intrusive rocks a nd porphyry coppe r occurren ces of th e Pa pu a New Guinea-Solomon Islands region: a reconnaissa nce study: Econ. C eol. 73 , 857-877. Moo re, W. J. (19 78) Chem ical characteristics of hydr otherm al a ltera tion a t Bingh am , Utah: Econ. C eol. 73, 126 0-1269. Mu ir, T. L., a nd W. V. Peredery (1984 ) Th e O na ping For ma tio n: In: T he Ceology an d O re D eposits of the Sudbury Structure, E. G. Pye, A. J. Na ld ret t, a nd P. E. Giblin (cds .). Onta rio Geo logica l Survey, Spec. Vol. 1, Chap. 7, pp. 139-21 0. M utschler, F. E., E. G. Wright, S. Lud ington, and J. T. Abbott (1981) Granite molybde nite systems: Econ. C eol. 76, 874-897. Na ldrett, A. J., and R. H . Hewins (1984 ) T he main mass of the Sudbu ry igneous complex: In: The Ceology and Ore Deposits of the Sudbury Structure, E. G. Pye, A. J. Na ldrett, and P. E. Giblin (eds. ). Ontario Geological Survey, Spec. Vol. 1. Nockolds, S. R. (1954) Average chemi cal compositions of some igneous rocks: Bull. Ceol. Soc. Am. 65 , 1007-1032.

122

Magmas and Hydrothermal Fluids

Piwinskii, A. ] . (1968) Experimental studies of igneous rock series, central Sierra Nevada batholith, California:]. Geol. 76, 548-570. Presnall, D. c., S. A. Dixon, ]. R. Dixon, T. H. O 'Donnell, N. L. Brenner, R. L. Schrock, and D. W. Dycus (1978) Liquidus phase relations on the join diopside-forsrerirc-anorthire from 1 arm to 20 kbar: their bearing on the generation and cry stallization of basaltic magma: Contr. Mineral. Petrol. 66, 203-220. Reynolds, T. j., and R. E. Beane (1985) Evolution of hydrothermal fluid characteristics at the Santa Rita, New Mexico, porphyry copper deposit: Econ. Geol. 80,

1328-1347. Ringwood, A. E., I. D. MacGregor, and F. R. Boyd (1964) Petrologic composition of the upper mantle. Carnegie Institution Washington Yearbook 63, 147-152. Roedder, E. (1971) Fluid inclu sion studies on the porphyry-type ore deposits at Bingham, Utah, Butte , Montana, and Climax, Colorado: Econ. Geol. 66, 98-120. Roeder, P. L., and R. F. Emslie (1970) Olivine-liquid equilibrium: Contr. Min eral.

Petrol. 29, 275-289. Schock, R. N. , and H. Louis (1982) Strain behavior of a granite and a graywacke sandstone in tension:]. Geophys. Res. 87, 7817-7823. Seal, R. R. II, A. H. Clark, and C. ] . Morrissy (1987) Stockwork tungsten (scheelite)- molybdenum mineralization, Lake George, southwestern New Brunswick: Econ. Geol. 82, 1259-1282. Shaw, H. R. (1980) The fracture mechanism of magma transport from the mantle to the surface: In: Physics of Magmatic Processes, R. B. Hargraves (ed.). Princeton, N]: Princeton University Press, pp. 201-264. Sillitoe , R. H., and F. ]. Sawkins (1971) Geologic, mineralogic and fluid inclusion studies relating to the origin of copper-bearing tourmaline breccia pipes, Chile:

Econ. Geol. 66, 1028-1041. Steininger, R. C. (1985) Geology of the Kitsault molybdenum deposit, British Columbia: Econ. Geol. 80, 57-71 . Stewart, D. B. (1967) Four phase curve in the system CaAI2Si20s-SiOz-HzO between 1 and 10 kilobars: Scbureiz. Mineral. Petrogr. Mitt. 47, 35-39. Turcotte, D. L., and G. Schubert (1973) Frictional heating on the descending lithosphere: j. Geopbys. Res. 78, 5876-5886. _ _ and _ _ (1982 ) Geodynamics. New York: Wiley. Tuttle, O. F., and N . L. Bowen (1958) Origin of granite in the light of experimental studies in the system NaAlSi 30g-KAISi30g-Si02-H20: Geol. Soc. Am. Mem. 74. van der Molen, I., and M . S. Paterson (1979) Experimental deformation of partially melted granite: Contr. Mineral. Petrol. 70, 299-318 . Voigt, D. E. (1983) The solubility of Al2SiOs in the system KAISi30g-Si02-H20 at 2 kbar and its implications for melt speciation: Unpublished M .S. thesis, The Pennsylvania State University, University Park. Webster, J. D., and J . R. Holloway (1990) Partitioning of F and CI between magmatic hydrothermal fluids and highly evolved granitic melts: Geol. Soc. Am. Spec. Pap.

246 ,21-34. White, A. J. R., and B. W. Chappell (1977) Ultra metamorphism and granitoid genesis:

Tectonophysics 79, 7-22. Willis-Richards, ]., and N. J. Jackson (1989) Evolution of the Cornubian ore field,

......

References

123

southwest England: Part I. Batholith modelling and ore distribution: Econ. Ceol. 84, 1078-1100. Wilson, J. C. (1978) Ore fluid-magma relationships in a vesicular quartz lati te porphyry dik e a t Bingh a m, Uta h: Econ. Ceo l. 73, 1287-1307. Wood , B. J . ( 1978) Reaction invo lving a northite and CaA I2S i06 pyroxenes a t high pressures a nd temperatures: A m. [. Sci. 278, 930-942. Yoder, H . S. J r. (1954) The system anorthite-diopside-wa ter: Carnegie Institution Washington Yearbook 53, 106-107. _ _ (1965) Diopside-anorthite-warer at five and ten kilobars and its bearing on ex plosive volcanism: Carnegie Institution Washington Yearbook 64, 82-89. _ _ and C. E. Ti lley (1962) Origin of basalt magmas: an experimental study of natural an d synthetic rock systems : ] . Petrol. 3, 342-532.

....,

Chapter

4 Hydrogeology and Geochemistry of Ore Genesis in Sedimentary Basins Grant Garven Deportment of Earth and Planetary Sciences, Johns Hopkins University Jeff P. Raffensperger Deportment of Environmental Sciences, University of Virginia

Reconstruction of paleohydrologic systems is undoubtedly one of the most fruitful areas for continuing research. Unfortunately it is also one of the most difficult, but the rewards are clearly worth the effort, because the prize will be the ability to pinpoint the location of buried ore deposits. B. J. Skinner (1979)

Hydrothermal ore deposits represent the time-integrated effect of hydrologic systems, many of which are operative over millions of years and involve crustal-scale fluid flow. For example, geologists have long recognized the important role of large-scale groundwater flow as a mechanism for the solution, transport, and deposition of ore minerals within sedimentary basins. In the previous edition of this book, J. S. Hanor provided an excellent discussion on the geochemistry of fluids in sediments, with particular attention given to the process of diagenesis as a mechanism for producing ore-forming brines. He described the physical and chemical properties of subsurface fluids and reviewed general aspects of brine migration. In this chapter, we solely discuss the hydrodynamic regimes responsible for the transport and deposition of epigenetic, sediment-hosted ore deposits. Our interests center on the physical and chemical evaluation of hydro125

126

Hydrogeology and GeochemislTy of Ore Genesis in Sedimentary Basins

logic systems known to occur in sedimentary basins, both past and present. The study of ore formation considered here is really a classical exercise in paleohydrogeology (Baskov, 1987), but modernized to take advantages of recent theoretical developments in groundwater modeling. Numerical modeling of crustal hydrogeology has become an important tool for quantifying geologic relationships of fluid migration, heat flow, and reactive transport as related to ore formation. We are now able to conduct numerical experiments of coupled transport processes in hydrothermal systems, simulate specific ore deposit environments, and compare or contrast competing hypotheses for fluid migration and ore deposition. This progress is due in large part to the remarkable technological advances in computer hardware and software over the past decade. Of course, much of the theory behind these hydrogeologic models of sedimentary systems is not new, only borrowed and revised from an earlier decade when similar advances were developed for flow systems associated with cooling plutons (Norton and Cathles, 1979) and for geochemical mass transfer between groundwater and minerals comprising the porous medium (Helgeson, 1979). Forces causing deep groundwater migration, such as the gravitational head gradients associated with topographic relief, fluid buoyancy associated with temperature and salinity gradients, and changes in stress fields associated with sediment compaction, erosion, and faulting, vary spatially and temporally as the structural and hydrologic framework of a basin evolves (Toth, 1978). Figure 4.1 illustrates several styles of regional groundwater flow, each of which is reflective of a tectonic regime. Gravity-driven flow (Figure 4.1 a) dominates in any continental setting where a topographic gradient exists and climate is moderate (Hubbert, 1940). But other hydrologic systems may precede gravity-driven flow as the basin evolves through stages of rapid sedimentation with compaction-driven flow (Figure 4.1d), subsidence and burial of thick aquifers with thermally driven free convection (Figure 4.1 b), compression and thrusting with tectonically driven flow (Figure 4.1c), and faulting with seismic pumping (Figure 4.1e). Fluid flow would cease in a basin if no hydraulic gradient existed or if the permeability became negligible at depth (Figure 4.1f), conditions we feel are rarely ever present in the Earth's crust. Because groundwater flow is the dominant mechanism for transporting chemical mass in sedimentary basins, knowledge of the hydrodynamics and geochemistry of flow and transport in porous and fractured media is fundamental to understanding sediment-hosted ore genesis. Quantitative calculations can be used to gauge the relative importance of the flow systems outlined above in stratabound ore formation. The aim of this chapter therefore is to concentrate on the basic principles of hydrogeology that bear on sedimenthosted ore formation. We shall discuss hydrologic concepts of ore formation and review mathematical theory for mass and chemical transport, although our primary objective is to cover hydrologic case studies of actual ore districts rather than theory. Our presentation is largely restricted to reviewing

Groundwoter Flow in Sedimentary Bosins

A

127

Fold & Thrust Belt

UPLIFTED FORELAND Ma ximum Flow Rate: 1 - 10 m / yr

j

200

Kilometer s

B

INTRACRATONIC SAC or RIFT

Maximum Flow Rate: 0 .1 - 1 m / yr

c

THRUST TERRANE

M aximum Flow Rate: 0 .1 - 1 m / yr

I

Kilometers

20

FIGURE 4.1 Hydrogeologic ond tectonic regimes for lorge-scole groundwoter flowin sedimentory basins. Arrows depict general flow directions. (0) Grovity- or topogrophyilriven flow in 0 forelond basin. (b) Thermol~ driven free convection in on introcratonic sog or rift bosin. (c) Tectonicolly driven flow in 0 fold-ond-thrust belt. (d) Overpressurir.g during compoction of 0 continental morgin. (e) Seismic pumping of deep fluids in 0 rift. (fl Comportmentalizotion of 0 bosin with no regionol flow (from Gorven, 1995).

hydrologic evaluations of strata bound or e forma tion involving mathematical simulation techniques.

GROUNDWATER FLOW IN SEDIMENTARY BASINS No introduction to the role of groundwater flow in ore genesis would be complete without some discussion of the types of flow systems encountered

-..

128

Hydrogeology and GeochemislTy of Ore Genesis in Sedimenlory Basins

o

E

M aximum Flow Rate: 0 .1 - 1 em /yr

RAPiDlY SUBSIDING MARGIN

Maximum Flow Rate: ) 10 m/yr I

SEISMIC PUMPING

r-J

5 Kilometers 0

Normal Fault

F

No Flow Betw een Compartments

PRESSURE COMPARTMENTS

~' 200

FIGURE 4.1

Kno me te rs

0

(Continued)

in sedimentary basins. We shall leave the physics and geochemical aspects of fluid migration to a later section, so that we can first review basic mechanisms causing flow. .

Mechanisms of Regional Groundwater Flow Sedimentary basins are subjected to several forc es that are known to cause large-scale fluid migration (Garven and Freeze, 1984a; Hanor, 1987; Bethke, 1989):

1. Gravity du e

to

changes in elevation along the water table

Groundwater Flow in Sedimentary Basins

2. 3. 4. 5. 6. 7.

129

Buoyancy due to thermal or salinity gradients Compaction due to sediment loading with subsidence and burial Tectonic loading due to crustal compression and thrusting Dilation due to crustal extension and normal faulting Stress relaxation due to erosional unloading Overpressuring due to diagenetic reactions

Topographic relief is responsible for the dominant form of groundwater flow in continental land masses, both in the shallow and deep subsurface. Figure 4.1a shows the typical profile envisaged by hydrologists for a foreland basin with a water-table slope that is a subdued replica of the landscape. Maximum flow rates of 1-10 mlyr develop in deep aquifers, while much smaller seepage rates occur in aquitards (Garven and Freeze, 1984a,b). Topography, heterogeneity, and anisotropy in permeability and the basin geometry are the primary factors controlling flow patterns. Free convection cells are driven by fluid-density gradients associated with temperature and salinity fields (Figure 4.1 b). Flow rates in convection cells may approach 0.1 m/yr, depending on thickness of the aquifer, fluid-density gradient, and regional permeability (Raffensperger and Garven, 1995a,b). Both topography-driven flow and buoyancy-driven flow will experience transients if the topographic gradient changes through geologic time or if temperature and salinity fields are variable. Although gravity-driven flow (sometimes called forced convection) will overwhelm free convection in uplifted basins, situations could occur in rifted basins where both types of convection coexist. Transient flow fields are more commonly associated with abnormal pressure gradients created through compaction, tectonic dilation, and chemical diagenesis. Tectonic compression and thrusting produce large overpressures near continental margins (Figure 4.1c). Groundwater flow rates in this environment are not well known, but theoretical calculations suggest velocities in the meters-per-year range for deeply buried aquifers, although these will dissipate (Garven et aI., 1993). Rapid sedimentation during subsidence of young basins can produce overpressures that approach the weight of the overburden (Figure 4.1d). Dehydration reactions, pressure solution, and hydrocarbon generation may contribute further to the generation of very high pore pressures (Bethke, 1989). Flow rates generated by compaction are typically much less than 1 cmlyr because of low-permeability shale sequences, unless catastrophic faulting can vent the overpressures (Figure 4.1e). One school of thought also argues that overpressures are maintained with seals or impermeable barriers, thereby separating the sedimentary basin into isolated compartments (Figure 4.1 f). Most hydrologists, however, believe no rock in the Earth's upper crust is impermeable, but rather a wide spectrum of finite permeabilities exists (Freeze and Cherry, 1979). Other coupled processes related to temperature and concentration gradients have the energy potential to cause fluid migration, yet they are generally

130

Hydrogeology ond Geochemistry of Ore Genesis in Sedimentory Bosins

considered to be of secondary or minor importance (Bredehoeft and Norton,

1990).

Basin Evolution and Groundwater Flow It is important to emphasize the role of basin evolution in characteriz ing the hydrologic setting and therefore the ore-forming system, so this will be a recurring theme throughout the chapter. The flow mechanisms de scribed above evolve gradually over geologic time as the sedimentary basin subsides or rifts, becomes structurally compressed and uplifted, and eventually is eroded. Figure 4.2 illustrates the general sequence of hydrologic regimes thought to have developed in the foreland basins adjacent to the Appalachian and Ouachita orogenic belts in the eastern interior of North America. Regional brine migration across the eastern interior region provided one possible source area for Pb-Zn mineralization in the Midcontinent

A. EARLY PALEOZOIC PASSIVE MARGIN

B.

C. POST-ALLEGHANIAN EROSION

FIGURE 4.2. Evolution of flow systems associoted with forelond basins in the ll.S, Midcontinent (from Gorven et 01., 1993). (0) Fluid migrotion is driven by sediment compoction ond density gradients during subsidence of rifted morgin ond plotform. (b) Moximum uplift of Appalochion and Ouochita Mountains accurs in the Permion, post·Alleghonion orogeny. Deep groundwoter flow is driven by grovity across the Midcontinent through regional aquifers. As a result, lorge MVT ore deposits ore formed os brines of vorioble temperature ond solinity dischorge neor basin morgins. (c) Bosin-wide migrotion of ore-forming brines declines in the Eorly Mesozoic becouse erosion reduces topographic drive and creates locol flow systems.

Conceptual Models for Stratabound Ore Genesis

131

(Garven et al., 1993) . Compaction-driven flow and/or free convection probably characterized much of the early Paleozoic hydrogeology, but lat er deep groundwater flow was most strongly driven by topographic gradients created by the closing period of th e Alleghanian orogeny in late Permian time (Figure 4.2b). Uplift of the Appalachian Mountains and foreland platform drove brines out of the Appalachian Basin, over adjacent arches and domes, and across the Illinois sag, with brine discharge focused near the edges of the sedimentary basins. Later, northward migration of brines occurred with uplift of the Ouachita mountain belt. Post-Alleghanian erosion beveled the high relief across all the forelands such that basin-wide discharge of deep brines was greatly reduced or dissected by the appearance of local flow systems, although some deep brine discharge continues to the present (Musgrove and Banner, 1993). Erosional unloading may have created additional transients, which would have persisted for some time, particularly in thick shale sections (Neuzil, 1986, 1993).

CONCEPTUAL MODELS FOR STRATABOUND ORE GENESIS Mineral deposits are formed in a wide variety of sedimentary environments, and many owe their origin to one form of hydrologic system or another. We adopt the modern usage of "groundwater" to include all subsurface water irrespective of the original source, chemical composition, and depth of circulation (Freeze and Cherry, 1979). In some ore-forming environments, therefore, deep groundwater may have originated as shallow meteoric water, mixed with oceanic water or mixed with evaporated brines in continental rifts or dissolved deep evaporite beds to become concentrated brines (Hardie, 1990) . The same fluids may eventually encounter metamorphic or magmatic conditions within orogenic belts as the host sediments are buried, lithified, and metasomatized as part of the crustal rock cycle.

Stratabound Deposits Sharp and Kyle (1988) present a hydrogeologic classification of ore deposits that includes the shallowest soil water of the unsaturated zone involved in weathering and supergene enrichment of sulfide ores to hydrothermal fluids convecting in deeply buried sediments intruded by plutons. In sedimentary basins not associated with volcanic or plutonic processes, stratabound and stratiform concentrations of copper, lead , zinc, barite, fluorite, uranium, and gold can forn{ by large-scale migration of groundwater of variable temperature, density, and composition. This is particularly true for epigenetic, stratabound ore deposits, although in many ways the distinctions between descriptive and genetic classifications such as stratabound or strar-

r 132

Hydrogeology and Geochemistry of Ore Genesis in Sedimentary Basins

iform and syngenetic or epigenetic are blurred. For example, shale-hosted copper deposits such as the Kupferschiefer in Europe and White Pine in Michigan were once thought to be syngenetic in origin but are now viewed as having formed through the cross-formationa l flow of oxidized, Cu-bearing brines (Figure 4.3) that reacted with pyritic shale (Figure 4.4). Mississippi Valley-type (MVT) deposits of lead-zinc-barite-fluorite are known to have formed by brine migrations in continenta l fore land basins of Western Canada and the U.S. Midcontinent. However, carbonate-hosted ore deposits of the Irish Central Mid lands may have involved both subsurface brine migration and "sedimentary exhalations" (sedex) of brines on the seafloor in an active rifting environment (Russell, 1992). Some urani um ore deposits form as redox-controlled roll fronts and tabu lar bodies within relatively shallow sandstone aquifers of Mesozo ic and Cenozoic age . In contrast, the giant unconformity-type uranium deposits of Proterozoic age formed at much greater depths, apparently due to convection of brines within thick basin sandstones and fractured basement. Gustafson and Williams (1981) provide a detailed review of other sediment-hosted deposits not covered in this chapter. We have adopted their philosophy on viewing the formation of sedimenthosted ores as normal products of sedimentary basin evolution, involving several variants of groundwater migration and tectonic regimes.

Hydrologic Mechanisms of Ore Formation The central role of deep and shallow groundwater circulation in ore genesis has been recognized for over a century. Daubree (1887) concluded that hot water was the most important agent in the formation of many types of ores and that these waters cou ld be meteoric fluids that were heated at depth. His theory presupposes the presence of large-scale groundwater flow systems. The association of sed iment-hosted ores with regional paleoaquifers also implies large volumes of gro undwater were involved in ore formation. No single groundwater flow system can be held accountable for all sedimenthosted ore types, but the mechanisms of flow in basins are well enough known to weigh the pros and cons of competing hypotheses . We shall summarize conceptual mode ls encountered in the literature for a select few of the ore deposit settings. We admit, however, that many models are based on geo logic interpretation alo ne and have yet to be quantified hydrologically.

LEAD ANDZINC Daubree's theory is perhaps best appreciated for the carbonate-hosted lead-zinc deposits of North Amer ica. For example, it is now we ll established from petrographic and geochemical studies that MVT ore deposits of the Midcontinent region formed from warm Na-Ca-CI meta l-bearing basinal brines some time after the lithification of their host rocks (Sverjensky,

W

w

-

;-'

~

tJ

s

~

~

0

~l

2-

\

\

T 20

TNONESUCH

: .r

I

10 0

, -.

. " " ",~~:~~E:~;~:,

COPPER ~

FIGURE 4.3 (oncepruol f10Vl pot1emsinthe Midcontinent Rift at the time ollOlmotion 01 the White Pine copper deposit, Michigan (from GalVen andFreeze, 1984bl.Original dolo on copper content 01 Nonesuch Shale isfrom White (l97l) .

-

I 10

Uphfled Region

:L\

6

NORTH

SOUTH

Hydrogeology ond Geochemistry of Ore Genesis in Sedimentory Bosins

134

0 .0 -

, , , I I REACTION OF CU-ORE FORMING FLUID WITH PYRITIC SHALE ( REL. RATE PY/C=IO) Cu

- 2 .01- 4 .0 '

E d

.2

Pb

~c - 8.0 !lc

3

:c 0.

-12 .0-140

60t

/ Fe

/~ .Zn

Fe

- 60 -

.§. -100 -

"

Zn

.....

I I I I

LOG fa 2 (Q)

-,

r-

~-

Cu - --Pb

- -44.0

--- ....

----.i ---.3 "!!

I

I 1:!..2~

>

- - 4 3.0

- -45.0

/

- -46.0

pH

- -47.0

50 -

0 .0

MAX . MICROCLINE MUSCOVITE

+

QUARTZ

PYRITE

~

~ u c

...'e 0

11

'0

s 3

-2.0

-1.0

logs

FIGURE 4.4 Chemicol evolufion of on oil field brine reamng with pyrific shole in 0 closed system, os computed in 0 reamon-poth E03/E06 simulofion (from Sverjensky, 1987), Microdine, muscovite, quartz, pyrite, and graphite comprise the shale bed, with 0 relative role of reorfion for graphite to be ~ of the other inifiol minerals. The recdian progress variable ~ represents moles of shole reacted. The sequence ofsulfide minerals is typicol of the zonofion observed in Cu-rich deposits like at White Pine, Michigon.

1986). Abundant geologic evidence documents a pervasive and widespread migration of basinal brines and relates MVT ore formation with Late Paleozoic tectonism in the central and eastern United States (Oliver, 1986; Bethke and Marshak, 1990). Calculations by Garven et al. (1993) evaluate likely patterns of groundwater flow and mechanisms for fluid migration in the Midcontinent Basins. Chamberlin (1882) and others before him wrote about the disseminated

.

Conceptual Models for SllOlabound Ore Genesis

135

nature of metals with in th e Paleo zoic section of the Midwest a nd argued that percolating groundwater could ha ve fo rmed ore dep osits in the carbonate strata of so uthwes tern Wis con sin. Similar geol ogic noti on s were explored by Van Hise and Bain (1902 ), Cox (19 11), and Sieb enthal (19 15) in th eir studies of lead-zinc ores of th e M ississippi Valley and Tr i-State regions. Fluid inclusion data , however, ind icated th at deep artesian groundwater flow appeared to be a pr erequisit e (New house, 193 3 ). Quest ion s regard ing the occurrence of deep groundwater flow and geochemistry of tr ansport put the " meteoric hypothesis " in disfavor for most of the first hal f of th is century, as sentiment swung back to even older theories assoc iat ed with warm fluids of igneous derivation (Ohle, 1959). The role of " latera l secretion" or " artesia n flow" in stratabound ore formation had not, however, been totally abandoned. Pelissonnier (1967) reasoned that deep groundwater could circulate through fractured basement under the hydraulic head created by topographic relief to form MVT ore deposits. He proposed that faults along basement highs would guide th e deep fluids upward into the sedimentary basin. Brine migration driven by sediment compaction became a popular model for MVT ore formation after the articles by Noble (196 3), Ja ckson and Beales (196 7 ), and Dozy (1970 ). According to this concept, steady subsidence in foreland and intracratonic sags would drive metal-bearing pore fluids out of compressible muds into more permeable sandstone and carbonate beds with lateral flow toward the basin margin (Figur e 4.1d ). H owever, White (1971 ), Cathles and Smith (1983), and Bethke (1985 ) demonstrated ma ssbalance limitations imposed by compaction-d riven flow as a means for ore formation. Furthermore, flow rates generated by compaction appear to be too small to elevate temperature near the edge of th e basin , as reflected in fluid inclusion data. Several variants of th e compaction story involve episodic venting of overpr essured aquifers. Sharp (1978) built a finite-difference model to simulate overpressuring in the Arkoma Basin and showed how rapid flow rates (1-10 rn/yr) might be achieved in adj acent aquifers if the sed imentary pile had become overpressured to lirhosratic values and suddenly ruptured by growth faulting. Fifty or more dewatering pulses, one every million years or so , were postulated by Cathles and Smith (1983) to explain the color banding observed in sphalerite of th e Upper Mississippi Valley. Fowl er and Anderson (1991) dismiss the concept of long-distance transport of brines and instead conjecture proximal venting of geopressured zones, using th e Gulf of Mexico as a modern analogue. Cathles (1993) also suggests th at venting may be aided by natural gas gen eration in the sedimentary pile. Deloule and Turcotte (1989 ) advanced a more complicated flow scenario invol ving a co m bination of buoyancy-driven hydrofracturing in tandem with 5 0 or more dew at ering pulses from th e Illin o is Basin. They postulat ed turbulent fl ow of hot brines up through fra ctures into sha llowe r strata in order to pr eserve fluid heat content.

136

Hydrogeology ond Geochemistry of Ore Genesis in Sedimentary Basins

Large overpressures probably form ed in the development of the Arkoma Basin, but sedimentation rates were too slow in the Illinois Basin to generate overpressures of th e typ e needed for compaction-venting arguments (Bethke, 1986). Some ha ve used th e Gulf of Mexico as a modern analogue for MVT o re-formation hydrodynamics, but Harrison a nd Summa (199 1) demonst rate that overpressuring is restricted to the very deepest parts of th e basin well beyond th e shelf margin. Episodic venting to the shelf margin may be pla usible, but the spatia l mech anic s, thermal consequences, a nd p al eohydrology of such dewatering processes remain to be quantified . However, the evidence linking the role of fractures to MVT mineral ization is tenuous. A thick Devonian evaporite bed underlies the carbonate-hosted Pb-Zn ore deposits at Pine Point. Granitic basement and Cambrian sandstone and dolomite aquifers are pervasively fractured in southeast Missouri lead districts, but major faults are virtually barren of mineralization. Some dating studies have placed Allegha nian ages on mineralization events in the Midcontinent (Symons and Sangster, 1991; Brannon et al., 1991), a period of final thrusting and subaerial uplift. It would be difficult to satisfy these ages, therefore, if overpressuring during sedimentation with periodic venting of deep aquifers is invoked as the sole hydrologic mechanism for ore formation. The role of tectonic compression , thrusting, and crusta l dilation in MVT ore formation may be more viable as the timing of th ese processes is at least compatible with general ages of mineralization. Oliver (1986) proposed that deformation of the Appalachian and Ouachita orogenic belts in the Late Paleozoic would have driven fluids away from the thrust belt toward the foreland platform to creat e both oilfields and ore districts (Figure 4.1c). Duane and de Wit (1988) appealed to Oliver's mechanism as an explanation of Pb-Zn ore formation within th e northern Caledonides. In contrast, Kesler and van der Pluijm (1990) concluded the MVT-type ore mineralization in eastern Tennessee was formed long before the final stages of Alleghanian thrusting and uplift. Hydromechanical modeling of tectonically driven brine migration (Ge and Carven, 1989, 1992) suggests relatively slow rates of flow (less than 1 m/yr). As a result of the small fluid discharges, the effect of forced convection on the geothermal gradient in the adjacent foreland basin is limited. It has also been conjectured that heat may be released catastrophically during orogenesis as a result of free convection of fluids in deeply fractured continental basement (Deming, 1992). A similar theory was originally proposed by Etheridge et a\. (1983) as a flow mechanism for regional metamorphism. But regional permeabilities in excess of 10- 15 m 2 (k > 1 millidarcy) are required for even slugg ish free convection, conditions th at are unlikely to be duplicated over continental scales in th e deep crystalline basement (Va lley, 1986; Ferry and Dipple, 1991 ). Seismic pumping of fluids along faults is another flow mechanism ca lled up on by some (Clendenin and Duane, 1990), yet there is no direct evidence linking ore mineralization to faulting in the giant ore districts such as south-

Conceptuol Models for SITOlobound Ore Genesis

137

east Missouri (Leach and Rowan, 1991) and Pine Point (Garven, 1985). Major faults in these MVT districts are barren of mineralization. The hydromechanics of seismic pumping has been analyzed with both analytical calculations (Nur and Booker, 1972; Sibson et al., 1975; Rudnicki and Hsu, 1988) and numerical simulations (Carrigan er aI., 1991; Muir-Wood and King, 1993): flow volumes generated by seismic dilation can be on the order of 5 x 10 6 rn-' and fluid pressure changes can cause centimeter to meter scale excursions of the water table. Although sedimentary basins are pervasively fractured, it is difficult to envision how seismic pumping could explain the widespread ore mineralization and pervasive sediment alterations observed throughout the U.S. Midcontinent. Garven et al. (1993) concluded that fracture networks are more likely to have affected regional and local permeability of the Paleozoic section than to have provided a driving force for regional brine migration. Some carbonate-hosted ores occur in rift environments where faults complicate the hydrothermal history. Deposits in Carboniferous limestones of the Central Midlands Basin of Ireland probably formed from thermally driven convection cells. Russell (1992) proposed a transient flow system with expansion and downward propagation of the convection cells through the rift during crustal extension. Expulsion of brines at the seafloor during faulting and/or hydraulic "jacking" of the sedimentary pile due to geopressuring of the rift was envisaged by Lydon (1986), as some of the Irish deposits have more similarity to sedex type rather than MVT-type of ore deposits (LeHuray et a!., 1987). Classical sedex deposits such as the giant lead-zinc deposits at Mount Isa and McArthur River, Australia, were also formed during the early stages of continental rifting, a period of high heat flow and faulting. Solomon and Heinrich (1992) conceptualized deep convection cells of the Russell type for the Mid-Proterozoic rifts of northern Australia, with high heat flow generated by radioactive decay in underlying granitic basement rocks. Recognizing the possibility of other fluid-flow mechanisms, Garven and Freeze (1982, 1984a,b) proposed a new hydrologic model for fluid and heat transport in MVT ore formation. They called for groundwater circulation at the scale of an entire basin in which the deepest, saline pore water would be driven updip toward the shallow edge of the foreland platform as a result of emerging topographic relief alone (Figure 4.1a). Hydrogeologic models of generic foreland basins were used by Garven and Freeze (1984b) to characterize a wide range of hydrologic, geologic, and geochemical conditions favorable and unfavorable for ore genesis. A fundamental result from this work was the demonstration that basin-wide aquifers allowed for the necessary focusing of brine and heat required for ore genesis. Flow rates of meters per year could be sustained over millions of years in the hydrothermal system, although this would gradually decrease as erosion of the landscape reduced the gravitational force driving brines across the foreland basin. Forced convection by gravity-driven flow lowers heat flow at the recharge end of the

Hydrogeology ond Geochemistry of Ore Genesis in Sedimentary Basins

138

foreland basin but elevates heat flow at the discharge end of the basin, such that warm brines (80-150°C) precipitate MVT ores at relatively shallow depths (less than 1500 m). Variability in depositional temperatures could be attributable to long-term transients in the thermal and hydraulic flow fi elds (Figure 4.5). The only limitation in their hydrologic model is the availability of saline water. Gravity-driven flow can easily deplete the entire salt source of the pore fluid, even obliterating bedded salt beds , due to meteoric flushing if strong hydrodynamic gradients are imposed for time scales in excess of a few million years (Bethke, 1986). Mathematical modeling of gravity-driven flow was first applied to a specific MVT are district by Ga rven (1985). He docu mented the paleohydrogeology of are genesis at Pine Point near the eastern edge of the Western Canada Basin wit h a two-dimensional finite-element model. Garven related

E a r l y - - - - - - - - - - - - - 1•• Late

TEMPERATURE cu 00

'::l

::l .C

u

•

~

::l -

e

~8.. oe uE

::Cl-

SALINITY u

o

OL...--.-----.----r---..-----.---......-_~----J FWORITE

MAIN STAGE CALCITE

BA RI T E

FIGURE 4.5 fluid temperature and salinity inferred from fluid inclusions os a function of paragenesis in the rennel Tennessee ore district (adapted from Gratz and Misra, 1987). Patterns such as this may reflect thermal and salinity tlOnsients in Mvr ore-forming groundwater flow systems.

Conceptual Models for Stratabound Ore Genesis

139

basin-wide brine migration a nd associated heat transport to a regional flow system that probably form ed Pine Point in less than a milli on yea rs, du e to th e Cretaceous-Tertiary eme rgence o f th e Rocky Mountains. Deep br ine flow was focu sed through a Devoni an dolomite-reef aq uifer (Figure 4.6). Bethke (1986) applied a finite-difference mod el to sim ulate th e regional flow that formed the Upper Mis sissippi Valley district in southern Wisconsin during a period when deep groundwater may hav e been dr iven no rthward across the Illinois Basin as a res ult of uplift of the Pascola Arch in Permian to Cretaceous tim e. Another ca lculatio n by Bethke and Marshak (1990) affords estimates on flow in the Arkoma Basin after uplift of th e Ouachita foldbelt, which may have implications for th e Tri-St at e ore district of Kansas-Oklahoma-M issouri. Garven et al. (1993) used a variety of numerical simulations to demonstrate how brine migration occurred throughout the early Paleozo ic in the U.S. Midcontinent region as for eland basins and intracratonic sags compacted and later were subjected to thrusting. But this compaction-driven flow was only a harbinger of the massive flow systems to follow. Tectonic uplift of the Appalachian and Ouachita foldbelts and adjacent platforms in the Late Paleozoic provided the necessary topographic relief for driving continentalscale brine migration that resulted in MVT ore formation . Basinal brines moved updip and out of th e Appalachian Basin to th e west and northwest through regional aquifers at rates of meters per year to form o re deposits in central Tennessee, central Kentucky, southeast Mi ssouri, and so uthe rn Wisconsin. Brines were later driven northward out of the Arkoma Basin to form ore deposits on the Ozark Dome and along the Reelfoot Rift to form deposits in southern Illinois and Kentucky. The largest flow rates were su stained for at most a few million years: fluid velocity gradually declin ed as erosion caught up with uplift such that the continental-scale flow systems were too weakened and d issected by local flow to generate large ore d istricts. Regional flow systems driven by topographic relief appear to place mo st of the pieces in the puzzle of MVT ore gen esis in their correct positions. For example, deep flow rates of 1 rn/yr can be sustained for millions of years, albeit th e rates will wane over hundreds of thousand s of years as erosion wears down th e elevated landscape. Because of the elevated groundwater flow rates, regional heat flow becomes elevated near th e edges of foreland platforms and along intracratonic arches, thereby expla ining the av erage deposition temperatures reflected by fluid inclu sions, at lea st in the sh allow platforms. Large-scale patterns of petroleum accumulation, cementation, and diagenesis associated with basin evo lution also fit well under a gen eral theory of gravity-driven flow. No other hydraulic th eory merges a ll of these pieces of the puzzle together so well. Sandstone-hosted lead deposits (Bjorlykke and Sangster, 1981) may share a similar hydrologic origin with MVT or es. Rickard et al. (1979) proposed

"'"'

0

"w

>

.g w

;;;

til

e

;;;

.5

..

:i

i

- 30 0 0

- 2000

-' 000

0

1000

~

II

•

Lia rd Range

West

1 II \H~ I h1r ' ll

\

20 00..,

I

~ ,

1~

,

Vert ical Ex aggera tion - 60

Hor izontal Sca le in Kilomet ers

o,

P::=-

1p V . (pq) = - a;-

j

L

=

k

Xij Xik

rj>

p,

po

p

Vi

P.O p. ,

p.

v·

ai'

XikMklsl

Dynamic viscosity of groundwater Reference dynamic viscosity Dynamic viscosity ratio = P.O/p. Stoichiometric reaction coefficient for the ith species G roundwater density Reference groundwater density Relative groundwa ter density =(p - pol/ po Porosity Composition coefficient (moles of i per formula weight j ) Composition coefficient (moles of i per formula weight k)

II,

K=

k

L

XikMklsl

XijN1jloq)

XijN1jloql +

=

L

j

=L

Moss Action in C hemical Subsystem:

M r, i

M I, ilsl

M I. iloql

Moss Balance in Chemical Subsystem:

Summary of Equations Gaverning Groundwater Flow, Heat Tronsport, and Reactive Mass Transport in the Subsurface

Conservation of Groundwater M oss :

Tobie 4.1

14.61

14.5c)

14.5bl

14.50)

146

Hydrogeology ond Geochemisliy of Ore Genesis in Sedimentory Bosins

more convenient for modeling rapid compaction of basins, as 5 s is a nonlinear function of porosity (Bethke and Corbet, 1988). The last two terms are often neglected for gravity-driven flow because total changes in stress and temperature arc relatively small. Mathematical solutions to equations (4.1) through (4.6) provide a transient picture of the flow rates, temperature, chemical composition of the fluid, and masses of minerals in the sedimentary basin. Each of the equations, however, arc implicitly coupled to each other and so mathematica l solutions are nontrivial, even for problems that might be simplified to one spatial dimension. For example, the Darcy flow rate q in equation (4.3) controls th e rate of heat convected thro ugh the sediment, but it must be obtained from solu tions to equations (4.1) and (4.2), wh ich depend on fluid properties th a t in turn must be ob tai ned fro m so lving equation (4.4 ), subject to reactive m ass tra nsfer in equations (4.5 ) and (4.6). Even mor e co mp lica tions arise if mi nera lization substa ntia lly affects porosity and permeab ility because th is feedback will alte r flow patterns dynamically and therefore patterns of chemical mass tra nsport (Norton, 1988). Similar nonlinear feedbacks may occur between the stress and flow fields because small strains may have a large influence on the permeability of fracture networks. Few hydrogeologic studies of ore formation have considered multicomponent reactive flow due to the complexity of this mathematical problem. Ague and Brimhall (1989) and Lichtner and Biino (1992a) simulated supergene enrichment of sulfide ore bodies with one-dimensional models; Lichtner and Biino (1992b) discuss a one-dimensional reactive flow simulation of MVT mineralization in limestone. Raffensperger and Garven (1995a,b) studied unconformity-type uranium ore formation with a twodimensional reactive transport model of free convection cells in a sandstone basin.

Numerical Methods N umerica l solutions of equations (4.1) through (4.6) are necessary to preserve the ir physical coupling and to mak e predict ion s of flow pa tterns in geologically inte resting flow domains. Detailed exa mp les of n umerica l formu lations can be foun d in the references already cited above. Both finite-d ifference and finite-element methods have found app lication in ore deposit research related to sedimentary basins. In either approach, the flow region must be discretized into a finite number of blocks or elements, within which fluid properties, such as density and viscosity, and rock properties, such as porosity, permeability, bulk compressibility, and thermal conductivity, are speci fied. Boundary conditions for pressure, temperature, and concentration also need to be assigned. The partial differential transport equations (4.1) through (4.4) are typically solved in a sequential-iterative approach with matrix inversion methods and then marched through time steps. The algebraic equations

Mathematical Modeling of Ore Formation in Basins

147

for th e che mica l submodel requ ire a Newton-Raphson iteration sol utio n d ue to th eir highly nonlinear for m. T he rmodynamic databases for a ll reactions suppleme nt th e ch emi cal eq uilibria calcu lations.

Regional Permeability and the Role of Faults The best way to treat groundwa ter migra tion in a frac tured medium is still a n open qu est ion in th e hyd rogeo logic mod eling of sedimen ta ry basins. It depends on the density of fracturing in relati on to th e sca le of th e gro undwater flow syste m and on th e relative permeab ility o f the sediments con tai ni ng the fracture systems . Two main approach es have been applied to problem s of groundwater flow (Fr eeze and Cher ry, 19 79 ; De M ar sily, 19 86): (1) modeling the fluid transport through each indi vidual fracture in th e network or (2) modeling the transport regim e as an equivale nt po rou s and anisotro pic continuum. In th e first a p proach, bulk fl ow rat es of a single-phase fluid throu gh a crystalline rock with multi ple fra cture sets can be com puted by integrating a N av ier-Stokes typ e eq ua tion to get (Snow, 196 9) N

q = - -Pg "L..J f,D,b,(I - n,n, )'Vh Jl

(4.8)

,=1

wh ere q is th e specific disch arge o r Dar cy velocity vector, p is fluid density, g is th e gravity co nst ant, Jl is fluid viscosi ty, f is a frac ture ro ughness facto r, D is fracture density (inverse of fracture spacing), b is the fract ure a perture, I is th e identity tensor, n is a un it vecto r orien ted norm al to th e fracture plane, and h is hydraul ic hea d. Eq uatio n (4.8) sums ove r N fracture se ts, with each set assumed to be co mposed o f per fectly par allel fracture planes th at ar e infinite in extent, w ith co nstant spaci ng and a perture. Some n ume rical models built along th ese lines use sta tistical mod els to gene rate th e fracture networks mathem atically and relax some o f the geo me try assumptions inherent in eq uation (4.8). Data regard ing o rienta tion, co ntinuity, and aperture of ea ch fracture mu st be o btai ned at the outcrop, analyzed by television logging of open boreholes, o r mapped via bor ehole tom ogr aphy. Fo rt unately, th e exte nt of basin aq uifers and density of fracturing normall y p reclude th e need to consider individual faults within a math emat ical domain o f a sedi mentary ba sin . In the seco nd approach, th e fractu red med ium is viewed essentially as a porous medium. In thi s case the permeabi lity becomes an aniso tro pic co ntin uum parameter th at characterizes the ability of a forma tion to transmit fluids th rough a vo lume of roc k large enough to encompass ma ny fracture sets. The flo w rate per unit area o f frac ture d roc k can be comp uted:

148

Hydrogeology ond Geochemistry of Ore Genesis in Sedimentory Bosins

q = _ kpg Vh

(4.9)

p.

where k is the anisotropic permeability tensor for some representative volume of rock. This form of Darcy's Law appears in most hydrologic studies of ore formation, but modified for variable-density fluid flow. In the case of fractured porous media, the effect of the fracture sets is to increase the bulk hydraulic conductivity of the formation at the regiona l scale, as graphed in Figure 4.8. Such phenomena are well known in the case of carbonate aquifers (Kiraly, 1975) and even thick sha le aq uitards (Bredehoeft et al ., 1983; Bethke, 1989; Capuano, 1993). A major q uest ion to be addressed here rela tes to the sca le o n w hic h a fractured rock mass behaves as a uniform, anisotropic mediu m. G a rven (1994) reviews data fro m hydrologic field experiments th at indicate it is reason able to view most frac tured format ions in basi ns as eq uiva lent a n isotropic porous media for any length sca le of fluid migration exceeding 10-100 meters. Many stratabound ores such as MVT lead- zinc deposits are hosted by formations that are porous and fractured or so-called dual-porosity media . For example, Gregg et al, (1993) report a primary porosity average around 19% based on petrographic observations for the ore-bearing dolomite beds in southeast Missouri. T hey showed that porosity and permeability were facies controlled at the onset of mineralization but today are controlled mainly by fractures and breccias due to the reduction of primary porosi ty by mineralization.

t

Effect of Karst and Regional Fracture Networks

_ l_ •

k = 1 darcy

Effect of Macroscale Fracture Sets

10°

10'

10 2

10 3

Effect of Primary Porosity and Microfractures

Scale of Measurement (m)

FIGURE 4.8 Effect of scale of measurement on the hydroulic conductivity of carbonote rocks in centrol Europe (odopted from GOIVen, 1986). Originol dolo ore from Kiro~ (1975). The open circles denote overoge permeobility doto os determined from care plugs, borehole tests, ond calibrotion of regionol flow model. Bulk permeobility oppeors to grow with the scole of flow because carbonote oQu~ers ore pervasively noctured ond korstic in noture.

Mathematical Modeling of Ore Formation in Basins

149

Over 90 % of the ore mineralization in southeast Mis souri is disseminated in dolomite porosity or replaces host dolom ite (Gregg and Gerdemann, 19 89). It would therefore seem that an eq uivalent porous medium approach is justified when dealing with large-scale flow in sed imenta ry basins co mpo sed o f clastics and carbonate strata . No consensus currently exists o n how best to tr eat isolated fractures o r sparsely fractured formation s at th e bas in scal e. Som e hydrologic stud ies hav e modeled groundwater flow through discrete fracture networks, but at sca les below a few hundred meters (Cacas et aI., 1990). Limited hydraulic field data exist for individual fractures: single faults can sometimes hav e high permeability and sometimes have low permeability (For ster and Evans , 1991 ). Unless aquifers are structurally offset by hundreds of meters, groundwater moves easily through both primary porosity associated with lithology and seco nda ry porosity associated with fracture networks, solution features, and unconformities. For the present, the most practical way to understand the regional role of these fractured porous sediments is to treat them as hydrostratigraphic units with enhanced bulk permeability (Figure 4.8) . Common sense dictates that kilometer-scale offsets across normal faults and thrust planes filled w ith go uge may indeed impede groundwater flow, so these ought to be tr eated as low-permeability zones (Bredehoeft et aI., 1992; Garven et aI., 1993). It may be desirable to consider discrete fractures in order to underst and geochemical transport and ore mineralization at flow length sca les below 100 meters (Arnold and Bahr, 1992). Discrete fracture modeling may be practical if all fracture sets can be mapped through mine workings and outcrops and the bulk permeability of individual fractures measured in situ through packer tests in boreholes. This typ e of analysis is impractical and fortunately unnecessary at the regional scal e, wh ere topography, hydrostratigraphy, and basement structure are th e dominant factors influenc ing regional fluid flow in sedimentary basins.

Transient Versus Steady Flow Flow systems within sedimentary basins rarely if ever reach true steady state as transients always will exist due to th e compressibility of the rock media, the low permeability of aquitards, diagenetic permeability modification, and the dynamic conditions set by tectonism and erosion. However, so me hydrologic systems may approach a state of dynamic equilibrium ov er geologic time periods such that the time-dependent terms in eq uations (4.1) and (4.2) may be suffi ciently small to ignore in mathematical so lutions. The response time r required for relaxation of a transient pr essure disturbance can be estimated as (4.10)

1SO

Hydrogeology and GeochemislIy of Ore Genesis in Sedimentary Basins

wh ere L is th e representative length for the fl ow system and K = KIS s is the basin hydraulic diffu sivit y (Domenico and Schwartz, 1990; Phillips, 1991). Th e length scale L for a basin might be ch osen to cha racte rize tim e sca les for fl ow relaxati on along the length of an aquifer subjected to tectonic compression , or alt ern at ively across th e thickness of a sha le aq uita rd for maintaining overpressures in a confined sandstone aquifer. The sa me exp ression (4.10) can also be used to estima te th ermal relaxati on times, which will be much longer th an hydraulic values because of the smaller thermal diffusivity parameter associated with heat transport (Ga rven, 1989). Most hydrogeologic models "Of gravity-driven flow and thermally driven flow in basin s utilize the steady-state assumption, wh ereas studies of compaction and tectonic compression necessitate a transient model because the pore space is constantly deforming. T he equations of multicomponent mass tra nsport with reactions (4.4) generally demand a tra nsient solution, even if the groundwater and therma l flow fields have reached a steady state, because conc entrations are ever cha nging by flow and diffusion.

HY 0 R0 GE0 LOG I ( SI M ULA T ION S: (A S EST U DIE S 0 F

BASINS In recent years th e genetic relationship between stratabound ore form at ion and regional groundwater flow has become better understood through paleohydrologic modeling of sedimentary basins. This progress is best discussed through an examination of case studies where th e physics and mathem atical theory described above have been put into application. Two major ore types have received con siderable hydrologic study: (l) Mi ssissippi Valley-type lead- zinc and (2) unconformity-type uranium.

Mississippi Valley-Type Lead-Zinc Deposits Carbona te-hoste d lead- zinc deposi ts constitute the best studied stratabound ores in North America. Summaries of th eir geologic and geochemica l featu res are covered elsewhere (Anderson and Macqueen, 1982; Sangster, 1983; Sverjen sky, 1986). Hydrologic features include an association with the undeformed, platform margins of foreland basins, the presence of basin-wide aquifers overlain by shal e aquitards, and migration of brines through very permeable dolom ite strata at temperatures that varied between 80 °C and 200°C.

WESTERN CANADA BASlIl The Pine Po int deposit is located at the eastern edge of the Western Canada Basin (Figure 4.9). A Middle Devonian barrier complex , th e Keg River For-

Hydrogeologic Simulations: Case Studies of Basins

1\

1\

1\

1\

1\

1\

1\

1\

151

1\

1\

1\

1\

1\

1\

55 0 118 0

1140

110 0

FIGURE 4.9 Middle Devonian paleageolagy and presentiloy hydraulic head mop in the Keg River borrier complex of western Canada (from Garven, 1986). The corbonate unit (brick pattern) comprises the Keg River Formation, on aquifer that focused southwest-northeast brine migration into Pine Point. Presentiloy hydraulic conductivities ore on the order of 3000 m/yr in the Keg River Formation. Hydraulic head contours ore lobeled in feet above mean sea level.

rnation, hosts about 100 million tons of lead-zinc ore, most of which is found in a karstic facies (Kyle, 1981). Ore deposition occurred during the migration of Na-Ca-CI brines wi th temperatures of 51-99°C and sa linities of 10-23 wt% eq uiva lent NaCI (Roedder, 1968). A gentle sou thwest to northeast flow pattern presently ex ists in the basin, a relict of a muc h stro nger hydrod ynamic regime created during the cru sta l relaxation and up lift phase of the Laramide orogeny (Beaumont et aI., 1985). Garven (1985) speculated that up lift of the Rocky Mo untains in Late Cre taceous to Early Tertiary time created a gravity-driven flow system strong enough to transport metal-bearing brines out of th e fore land basin toward Pine Point (Figure 4.6 ). Forced convection was call ed upon as a mechanism for ex plaining the relatively warm th ermal conditions for ore genesis at the basin margin. Figure 4.10 shows the mathematical cross section used to test the gravity-driven model and a sample simulation. The base of the model is impermeable to fluid flow but conductive to bas ement heat flow. The left

Hydrogeology and Geochemistry of Ore Genesis in Sedimentary Basins

152

HYDROSTRATIGRAPHY (EARLY CRETACEOUS)

NE Prne Po,",

o

o

100

400

300

200

I ~oo

Kliomerers

LEGEND

[£] [2]

Snore and Sond'tone (Lo.er Cre'o(oou,) Shale and limestone (Miss issippian)

CD Shale (Upper Devonionl Q]

Dolomil. (Mlddl. DevonIan)

STREAM FUNCTION K, =500m/yr K2=20 m/yr

TEMPERATURE J = 70 mW/m 2

100

200

300

400

500

Kilometers FIGURE 4.1 0 Finit~lemenl simulation of grovityi!riven flow hom the Pine Point case study of GalVen (1985, 1986). See Figure 4.9 for fine of section.

boundary is a hydraulic divide near the leading thrust fault in the deformed belt, and the right boundary is another hydraulic divide, arbitrarily set near the erosional edge of the basin where groundwater would be forced to discharge. Both sides of the model are insulated boundaries for heat flow. The upper surface represents the water-table boundary where pore pressure is atmospheric and temperature is assigned a mean annual air temperature. The stratigraphic profile of Figure 4.6 was represented by four hydraulic units, with Unit 1 portraying the Middle Devonian carbonate aquifer. It was as sumed to have a horizontal hydraulic conductivity K I = 500 m/yr, which is at least ten times greater than overlying mudstones of Late Devonian to Cretaceous age. Steady-state solutions to equations (4.1) and (4.3) define the paleoflow lines and temperature pattern for the basin simulation (Figure 4.10). Ground-

..

Hydrogeologic Simulofions: Cose Studies of Basins

1S3

water discha rge through each streamtube is 50 m 3 / yr, per meter width of basin no rma l to the section. Most of the regional brine migration is focused by the Devonian aquifer unit, where the Darcy flow velocity reaches a maximum of 0.9 rn/yr near the basin margin. Forced convection by gravity-driven flow eleva tes the deep brine temperature to about 80°C near Pine Point, ass uming only 800 me ters of burial. The flow rates indicated here are conservative: in situ field measurements of the host rock permeability indicate Kma x = 5000 m /yr, Mass balance calculations for metal transport in the flow system show a minimum of 100,000 years is required to form the Pine Point district assuming reasonable values of formation properties and metal concentrations (Figure 4 .11 ). Garven (1985, 1986) discusses the results of other numerical experiments and the effects of changing model parameters on regional flow. No reactive-flow simulations have been conducted for Pine Point, but rele-

1000 -.------.......--- - - - , - -- - - - - -- , - - -- - -- - - - ,

PINE POINT MODEL

0N

::c

bo 100

~

10 4 yr

E

0$

OIl

c

0c'"

.9

~c u

o

10

c

8u c

N

I +-~-,.-..:~"""'..,..,.,r====r=::::;="",.,.'Fi""r_rrr_--.------,M....._rrn_1

0 ,1

1.0

10.0

100. 0

Specific Discharge, m3tm2· yr FIGURE 4.11 Relafionship between specific discharge (Darcy velocity), zinc mass precipitated, and durafion of mineralizafion as computed for the Pine Point model (from GalVen, 1985). Boxe(Hn area represents most likely parameter ranges far stratabound ore lormafion.

154

Hydrogeology and GeochemislTy of Ore Genesis in Sedimentary Basins

vant reaction-path calculations with EQ3/EQ6 software (Wolery, 1979) are presented elsewhere by Garven and Freeze (1984b) and Anderson and Garven (1987). Other hydrologic regimes besides gravity-driven flow existed in the Western Canada Basin, but none were as important in MVT-type ore genesis. Compaction-driven flow was too weak to generate the flow rates or temperatures required for Pine Point. Thermally driven free convection is a better candidate for brine migration during the passive or rifted margin stage of basin evolution. However, numerical calculations by Carven (1989) demonstrated centimeter-per-year maximum flow rates and negligible mass transport of brines to the basin margin (Figure 4.12). Tectonic compression, thrusting, and gravity-driven flow developed in two main periods of orogenesis-(1) Late Jurassic to Early Cretaceous and (2) Late Cretaceous to Paleocene. The last phase of deformation and uplift is recorded as a subdued relict in presentday hydrodynamic patterns. Although lead isotope dates for Pine Point are ambiguous (Cumming et al., 1990), the massive accumulation of 200 km 3 of oil in the Lower Cretaceous Mannville Formation of the Athabasca Tar Sands deposit indirectly points to an analogous Cretaceous-Tertiary hydrologic origin for the Pb-Zn mineralization at Pine Point (Garven, 1989).

U.S. MIDCONTINENT BASINS Large carbonate-hosted Pb-Zn-Ba-F districts occur in the Mississippi Valley region (Figure 4.13). In the southeast Missouri ore districts, well over 500 million tons of ore have been mined from Upper Cambrian and Lower Ordovician dolomite strata that blanket the Precambrian ridges and knobs of the St. Francois Mountains and Ozark Dome (Hagni, 1989). Rich deposits of lead and zinc are confined to a dolomitic reef facies of the Cambrian Bonneterre Formation. Pinchouts of the underlying Lamotte Sandstone against the basement and collapse brecciation trends in the Bonneterre Formation appear to control mineralization patterns (Anderson, 1991). Fluid inclusion data indicate a wide range of sphalerite and dolomite precipitation temperatures in the Viburnum Trend. Temperatures mostly range between 90°C and 120 °C, with a mean of 110°C (Leach and Rowan, 1993; Shelton er al., 1993). Salinity varies up to 35 wt% equivalent NaCL Fluid inclusion and trace-element data have been inferred by some to indicate multiple period s of brine mixing coeval with dolomitization and Pb-Zn ore formation (Shelton et al., 1992). The age of ore formation has not been established unambiguously, but most agree that mineralization was genetically associated with the Alleghanian-Ouachita orogeny of Late Pennsylvanian-Early Permian time (Sverjensky and Garven, 1992). Garven et al. (1993) considered several scenarios for regional fluid migration and ore formation in the Midcontinent Basins. Space constraints limit our review to twO hydrogeologic systems for the ore districts in southeast Missouri.

V'I V'I

(J)

6

re:

o

4

6

81

o

2

52

-J

0

~

o

2

~

o

t J___

a:

(J)

~

-J

0

~

W 4

~

W

a:

I

8 ....

I

100

150

I

150

I

I

20"c

250

I

300

200

250

300

I i i

40

200

I

I

400

I

400

KILOMETERS

350

I

350

I

450

1

B.

450

600

I

500

550

I

600

I

LlT =200(

I

650

I

650

TEMPERATURE

550

I i i

500

a regional Devonianaquifernear the base, but flow rates do not exceed 1cmfyr andthe patternsofcirculationdonot allow for metal transport to the basin margin.

FIGURE 4.12 Simulation of thermally driven free convection in the Western Canada Basin (from GalVen, 19B9). Thisprofile contains

50

I

100

i i i

50

£11/1=1.5 x 10-3 m2 / yr

700

I

700

I

A. STREAM FUNCTION

I

750

750

156

Hydrogeology ond GeochemislTy of Ore Genesis in Sedimentory Basins

NC

OK

Location of two hydrogeologic section models for the Midcontinent case study of Gorven et ol (1993). indicate location of mojor MVT ore districts.

FIGURE4.13

Mine

symbo~

The Arkoma Basin, Mississippi Valley Graben, and Black Warrior Basin comprise the southernmost sources of mineralizing brines for ore districts in southwestern Missouri, southeastern Missouri, and southern Illinois-Kentucky, Section A-A' in Figure 4.14 cuts through the eastern half of the Arkoma Basin and extends northward over the Ozark Dome. In Figure 4.14, Cambrian through Mississippian age sediments on lap th e Precambrian St. Francois Mountains. Five hydrostratigraphic units are represented in this simulation: Unit 1 (in black) is the crystalline basement of the St. Francois Mountains with a very low hydraulic conductivity K I = 10- 5 mlyr and a porosity cPt = 0.01; Unit 2 represents a combined ba sal Lamotte sandstone and permeable Bonneterre dolomite that eventually thins downdip in the Arkoma Basin. This unit is assigned a hydraulic conductivity along bedding planes of K2 = 600 mlyr and porosity cP2 = 0.25. A less permeable Ozark Aquifer, Unit 3, conformably overlies the basal units, and K3 = 300 mlyr with cP3 = 0.20. Unit 4 contains low-permeability carbonates and shale of Ordovician to Mississippian age, which are overlain by a thick wedge of Pennsyl vanian-Permian clastics of Unit 5. For both model units, K4 = Ks = 10 mlyr and cP4 = cPs = 0.10. All of the sedimentary units within Figure 4.14 are assumed to maintain a vert ical hydraulic conductivity that is -r&o of th e along-bedding valu es cited above. None of these hydraulic conductivity values should be compared directly to core measurements of permeability. Rather, they should be viewed as simulation parameters representing estimates of hydraulic conductivity for heterogeneous strata at the basin sca le. The ba se of the model is taken to be impermeable to groundwater flow, and a constant heat flow of 70 mW 1m 2 is assigned along this basement sur-

Hydrogeologic Simulations: Case Studies of Basins

157

HYDROSTRATIGRAPHY SECTION A -A', VERSION 1

150

200

250

300

350

400

STREAM FUNCTION

8

AlJ.' -50 m 2/yr

o-t""'--- - . - - ------r- --

o

50

100

-.--150

-

--,- -- , . - -200

250

--,- - - - . - - -----., 300

350

400

8

CI)

II:

TEMPERATURE

6..,------ --20.C

100000 YEARS

W I-

W4

::::ii

o ..J ;;;:2 o-+-O:::::"---.----

o

50

----r- -100

,.------,150

--'--- , - - ----,-

200

250

300

-

--,------, 350

400

KILOMETERS

Hydrogeologic model for sectionA-A', across the Arkoma Basin and Ozork Dome (fromGalVen et ol, 1993). Flow is from south to north.

FIGURE 4.14

face. The fro ntal th ru st of the Ouachita Moun tai ns provides a near-vertica l hydra ulic d ivide at the so uthern end of the section an d an arbitra ry divide is po sitioned at th e thi n northern end. The water-tabl e profile defines th e top surface of th e flow region , wh ere it is ass umed to follow a subdued repli ca of th e Permian topogr aphy. Temp erature along th e water tabl e is fixed at 20° C. Prior to th e fina l stages of th e subaeria l emergence of th e foreland basin and platform, we assume the basi n profi le is submerged or near sea level and under hydrosta tic conditions. After erosional unload ing of th e mo unta inous

158

Hydrogeology and Geochemislly of Ore Genesis in Sedimentary Basins

O uac hita hinterl and , th e foreland basin a nd platform are upl ift ed to th e gentle relief depicted in Figure 4.14 . Th is uplift occurs instantaneou sly in our num er ical mod el, bu t new flow patterns a nd temper ature fields eme rge as th e ba sin adjusts to the new boundary co ndi tio ns. It takes a bo ut 10 5- 106 yea rs for both fl ow fields to ad just to a new steady sta te (G arven , 1988). For sectio n A-A' , regional brine migration is stro ngly controll ed by th e deep Ca mbria n- O rdov ician aquifers in whi ch fl ow is focused . The st rea m line map indic at es th at most o f th e section north of th e Ark ansas-Missouri border (nea r x = 200 km) wa s a regional discharge a rea, a ltho ug h most o f th e fl ow was forced upward near th e pinchouts aga inst th e St. Fra ncois ba sem ent, wh ere Darcy fl ow rates are th e lar gest at abo ut 5 mjyr. The temperature map at 100 ,000 yea rs after upl ift suggests a rather ge ntle rate of cooling for brines movin g out of th e foreland sag . H owever, a very large temperature gradient is pr edicted nea r th e edge of th e aquifer pinchout a n d within the basement high in thi s hydrogeologic scena rio . Therm al histories ar e shown in Figure 4.15 for one refer en ce site near

160 ..--....,..,.....",r-"'"'T'T'.....,--....,..,.,""""'...................,...................,.--.."'T'T'.......

w

a:

:::> ~ et

a: w

0.

:E

w

~

701=-

-e;.==:::::---

K=30 m1yr

TIME (YEARS)

FIGURE 4.15 Thermal history for brine discharge in southeast Missouri, near point 2 in sedien A-A' (see Figure 4.14). The graph shows temperature versus tme for four choices of hydraulic conductlvity assigned to the regional aquifer. Transients such as these may explain variable temperature and salinity observed in fluid indusion data such as those illustrated in Figure 4.5.

Hydrogeologic Simulotions: Case Studies of Basins

159

the basin margin. This graph illustrates how transient thermal pulses might advance through a near-steady groundwater flow system (Garven, 1988, 1989). Four numerical experiments are displayed to show the effect of aquifer hydraulic conductivity on the magnitude of the temperature anomaly propagating through the basin. We have assumed the same hydrologic properties as before but uniformly thickened the sedimentary wedge by 400 meters. The thermal pulse occurs due to the rapid rate of uplift and subsequent development of gravity-driven, forced convection. Hydraulic and thermal diffusivities [equation (4.10)] define the rates at which hydraulic and thermal transients move through the flow system. The extra heat transported through regional aquifers by gravity-driven flow results in elevated temperatures for about 100,000 yr, until conduction can dissipate the heat and temperatures stabilize at a new steady state for the up lifted foreland basin. Based on this model, brine temperatures range between 80° and l30°C in the vicinity of ore d istricts in northern Arkansas and southeast Missouri. Section E-E'. in the eastern interior region considers flow from east to west across the Appalachian and Illinois Basins (Figure 4.16) . The cross section starts near the frontal thrust of the Appalachian Mountains in Virginia, straddles the Kentucky-Tennessee border over the Cincinnati Arch, and extends for about 100 km west of the Ozark Dome into southeast Missouri. Unit 1 is the Lamotte/Mr. Simon Sandstone, which laterally is replaced by hydrostratigraphic Unit 2 over the Nashville Dome and adjacent Appalachian Basin. A basal Cambrian clastic sequence is also represented by Unit 1 in the easternmost part of the model, within the foredeep in front of the thrust belt. The only regional aquifer characterized in the section is Unit 3, the dolostones and limestones of the karstic Knox Group. For Unit 3, K3 = 200 m/yr and ¢3 = 0.20. Ordovician-Silurian shale and carbonates cover the Knox Group (K4 = 1 rn/yr, ¢4 = 0.10) . Mississippian shale caps the shelf sequence as Unit 5 in the model (Ks = 2 m/yr, ¢s = 0.15). Unit 6 represents undifferentiated clastics of the Pennsylvanian-Permian sequence with a nonlinear topographic profile (K6 = 5 rn/yr, ¢6 = 0.10). The St. Francois Mountains in southeast Missouri are barely submerged by about 200 m of Permian sediments. Lateral continuity of the Knox Group (Ozark Aquifer) and Mr. Simon/Lamotte Sandstone a llows for a significant degree of flow focusing as brines are driven downward and across the Nashville Dome. Upward seepage of deep fluids appears to be prevalent all across the Illinois Basin, although most of the flow is driven laterally through the Cambrian-Ordovician aquifers and forced to discharge near the eastern edge of the buried St. Francois Mountains. A relatively small volume of brine migrates over the St. Francois high, even if the Lamotte and Bonneterre beds provide a continuity in the flow path. Maximum Darcy flow rates of 3 m/yr are predicted in the Knox Group aquifer, while minimum flow rates of 10 - 3 m/yr occur in the deepest Cambrian shale. Forced advection of heat in southeast Missouri creates a threefold increase in surface heat flow above the ambient basement heat flux of 70 m \X'/ m 2 . The lateral thermal gradient in the Knox paleoaquifer reaches a maximum

Hydrogeology and GeochemisfTy of Ore Genesis inSedimentary Bosins

160

8

HYDROSTRATIGRAPHY

~6

.

SECTION E-E'

W ~

@

W 4 ::E

0

...J

~2

0

200

400

600

800

8

STREAM FUNCTION

~8 W ~

W ::E4

0

...J

:'::2

0 0

200

400

600

800

8

TEMPERATURE

(/)6

100000 YEARS

a: w

~

W4 ::E

0

...J

~2 0

o

200

400

600

600

KILOMETERS

Hydrogeologic model for section H' across the Appalachian and Illinois Basins of northern Tennessee and southern Kentucky (from Garven et ol, 1993). Flow is from east 10 west. FIGURE 4.16

of 3.2 x 10- 3oe / m near the eastern edge of the St. Francois high (near x = 600 km, Figure 4.16). The maximum vertical temperature gradient is about 1.0 x 10- Joe/m in the same area. Temperatures in the Knox Group aquifer hover around 100°C after 100,000 years of flow, except much warmer conditions persist in the western portion of the modeled section. Transient thermal histories of fluid passing through the Knox Group on the Nashville Dome show temperature rises to a maximum of 148°C after 100,000 years but later drops off to a value of 115°C (Figure 4.17). On the Ozark Dome in southeast Missouri, temperature rises dramatically to a maximum at 143 °C and then levels off to a steady 129°C. Both sets of predicted temperatures for the Late Permian are compatible with fluid inclusion data from are mineraliza-

0-

..

~

S

C-

E u

:l

u~

U

0

110

120

130

140

150

16 0

10

1

~

o-I

::::E

W

I-

w

a:

en

10

2

Time, yr

10·

10 5 10 e 10

7

10 2

10 3

Time, yr

10·

FIGURE 4.17 Thermol history for reference points in sectionH ' (from GOlVen et 01., 1993).

10 3

~

S

o.

E u

:l

0'

.

U

0

10 5

SECTION E-E'

HYDROSTRATIGRAPHY

10 6

10 7

162

Hydrogeology ond Geochemistry of Ore ~enesis in Sedimentary Basins

tion in cent ral Tennessee (100- 140°C) and so utheast M issouri (90- 120 °C), a ltho ug h th e results a re no nuniq ue. O t he r flow mech a nisms proba bly ca used regional br ine m igrati on in the M idcontinent Basins, lon g before uplift of th e Appalachi an-Ou ach ita M oun ta ins. Some of th ese hav e been quantified . Compacti on-d riven flow systems have been simulated for th e Illinois Basin (Bethke, 1986; Bethke et a I., 1991) and th e Ark oma Basin (Bethke and Ma rsha k, 1990 ; Ga rve n er al. , 1993 ). Figur e 4.18 shows one so ut h- no rt h flow sim ulation fo r th e Illin ois Basin. T he bottom and right side of th e mod el a re no-flow boundari es, w hile th e top a nd left side a re ope n to discharge. Small ove rpress ures develop a fter 280 million yea rs of burial becau se of th e slow rat e of sedimentation (3 x 10- 5 mj yr ). Gro un dwate r fl ow rat es reach a maximum of a bo ut 2 x 10- 3 m j yr in the deeper aqu ifers, rat es tha t are too sma ll to elevate th e conducti ve tem perat ure field near the edge of th e basin. Much lar ger overpress ures probab ly formed in th e Arko ma Basin because of th e rapid sedi me ntatio n rate th ere (abo ut 1 x 10- 3 m j yr ), bu t ma xim um fl ow rates of 7 x 10- 3 m jyr we re too sma ll to tran sport sufficient quantities of heat o r aq ueo us ma ss for o re for mation on the Oz ark Dom e (Gar ven et a I., 199 3). However, lateral co mpress ion and thrusting dur ing the bu ilding of th e Ouach ita and Appal achian oroge ns played a mor e impo rta nt role in basin al br ine migration (Oliver, 1986; Ge and Garven, 199 2). Figure 4.1 9 depi cts a simulatio n across th e leading edge of the defor med belt, near th e eas tern Tenn essee ore district. Several hydrost ratigraphic units co mprise th e permeab ility and poroelastic property fields:

.~'Ol ~

.0 2

,~~~~~d3~~~§~~·- .04 .·0 053 ~~

.0 6

50 km

.0 6

.05 .04

End of Permian (22 5 Ma BPl

FIGURE 4.18 CampoctiorHlrivenflow simulotion of the l1Iinois Bosin (fromBethke et ol, 199]). Flaw is fromsouth to north. Allaws represent Dorcy velocity field after deposition of 1.5 kmof Permian shale. Contour lines represent the excess pressure (relative to atmospheric) created by sediment compaction. Contour labe~ denote the mechanical energy per unit fluid volume, =p + pg z. in MPa, where p is pore pressure, p is density, and z is elevation above a datum (Hubbert, 1940).

Hydrogeologic Simulotions: Cose Studies of Bosins

163

HYDROSTRATIGRAPHY 7

6

f! 5 Q)

Q;

4

.§

3 2

~

O+---r-----r-----.----.-----,------.---y----:=-r-~=_L,

HYDRAULIC HEAD 160 years 7

6

f!

5

Q)

Q;

4

.§

3

~

2

O-t-- - -y-----,-- - - - .- - - ..,.-- - -,-- - - -.-- - - y--- - -,-- ----'=__,

FLUID VELOCITY VECTORS Largest Velocity = 3.6 m/yr 7

6

..

5

Q;

4

rn

Q)

E g 3 ~

2

0 0

5

10

15

20

25

30

35

40

Kilometers

Hydromechonicol flow model for the Pine Mountoin thrust systemin eostern Tennessee (from GOIVen et 1993), 160 yeors ofter 0 tectonic force of 800 MPo wos opplied to the right margin to initiote deformotion olong the thrust plone (Unit 7). Flow is from southeost to northwest, with focusing in oquners of Combrion sondstone (Unit 1) ond Ordovicion carbonotes (Unit 4). FIGURE 4.19

01.,

45

164

Hydrogeology and GeochemislTy of Ore Genesis in Sedimentary Basins

Unit 2 is a Cambrian sandstone aquifer (K2 = 100 m/yr) and Unit 4 is the Ordovician Knox dolostone aquifer (K4 = 50 rn/yr), while other units are low-permeability aquitards. Unit 7 is the Pine Mountain thrust fault zone (K7 = 0.01 m/yr), along which slip occurs as a tectonic stress is rapidly applied to the right side of the model. The numerical model uses a poroelastic stress formulation to predict two-dimensional strain and groundwater flow. Large overpressures develop in the foot wall and hanging wall of the thrust sheet at the onset of thrusting, with groundwater driven away from the areas of highest pore pressure with focusing through deep aquifers. Darcy flow rates as high as 100 mlyr occur at the onset of the seismic event near point A (Figure 4.19), but this decays to about 3 m/yr after 160 years, and less than 0.5 mlyr after 500 years of pressure dissipation. Episodic flows such as these may help explain mineralization near fold and thrust belts. But the flow volumes generated by thrusting are too small and the transient effects too diminished to influence ore formation at the huge districts situated hundreds of kilometers away from the orogenic belts.

Reactive Flow. Basin-scale flow modeling is a prerequisite to understanding the processes governing ore formation in a district and at the deposit scale (Figure 4.20). Some stratabound ore districts such as the Viburnum Trend are regional by nature, extending at least 50 km over the Ozark Dome. To quantify geochemical mass transport at this scale is desirable, but unfortunately impractical with the current level of computer hardware available to most researchers. Fully coupled reactive-flow simulations (Table 4-1) require several days of computational time on even the fastest workstations. Our efforts to date have therefore been confined to smaller-scale simulations of coupled brine flow, heat transport, and multicomponent mass transport. Figure 4.21 illustrates a simulation of brine migration over a buried basement high of the type found along the Viburnum Trend of Missouri (Anderson, 1991, Figure 12). The groundwater flow field and heat transport are assumed to be near steady state so that effects of flow transients could be ignored as a first pass at the problem. Our reactive flow field is discretized with three basic rock types: Lamotte Sandstone, Bonneterre Dolomite, and Davis Shale. Hydraulic conductivity along bedding is K ss = 300 rn/yr, Kdol = 150 rn/yr, and K sh = 0.15 m /yr. Vertical permeability across bedding is assumed to be 1/20 of the horizontal values for each layer. Initial porosity is 0.25, 0.20, and 0.10 for the sandstone, dolomite, and shale, respectively. The base of the Lamotte Sandstone is assumed to be impermeable to fluid flow and so flowlines are forced to parallel the basement unconformity. The left side of the model is an inflow boundary while the top and right side are surfaces of groundwater discharge. We have applied a regional hydraulic gradient of 0.005 m/rn to drive brine flow from left to right as shown. Heat enters the flow field at the base as a prescribed flux of 70 mW1m 2 , while the top of the numerical mesh is assigned the 110°C isotherm. Darcy flow rates reach a maximum of 3.0 mlyr near the pinchout of the Lamotte Sandstone and over the crown of the _

Hydrogeologic Simulations: Case Studies of Basins

165

BASIN 5 r--~ 4 (f)

a:

~

...J

2

52

o

o

100

300

200

400

500

KILOMETERS

1000 -r-----------&..--------""T""'1 800

DISTRICT

~600'1I1111. ~400 1

20

10

20 30 KILOMETERS

40

50

200-=======~======

DEPOSIT o+:.-:...:..;..--:.:.;--:....;....--:.:..:.....r-'--""..:........:'-'r'-......:.:..:-:...:..;..,.....:.:..:.....;..:...~

o

2 3 KILOMETERS

4

FIGURE 4.20 Visualization of three scales for regional flow relevant to MVT ore formation in sedimentary basins. Simulation of reactive flow is best done at the smaller scale, where heterogeneities in permeability, lithologic facies, and are mineralization (black patches) can be mopped and discretized in finite-element models (from Gorven, 1995).

. ,,

5

Hydrogeology and Geochemistry of Ore Genesis in Sedimentary Basins

166

en

100

[jj :E 50

o

500

Stream Function

8

150 - r -- - - --

~ w

100

1500

1000 METER S

- - - - --

- --

-

-

- -- - - - --

2000 -,

4 - - - - - - - - -......- - -

[jj :E 50

" 1 - - - - - - - - -....-

1000 METERS

C

150

1500

e .:....: m..:.!p:...:e:....:.r..:. a ..:. t u=.:r:....:e:.-.-,rT ..:.

2000

--.

en 100 +======..:===--===:...--=-~­

II:

w

[jj :E 50

=l-- -- - - - - - -- -

o

1000 METERS

1500

2000

FIGURE 4.21 Southeast Missouri reactive-flow Simulation I. Moss-basedsneomfunction in (b) shows flowlines and is contoured with the interval of 10,000 kg/m-yr (from Gorven, 1995).

gra nite kno b. Flow pattern s in th is scena rio ind icated a bo ut 4° C of cooling a long the flow path as brine movement is up and over the basement high. A cooling of lOoC is possible in other flow scenarios, but Darcy flow rates would need to do uble or basement relief be much more pronounced than conside red in this simulatio n (see Figure 4. 16 fo r th e latter case) . Assuming local chemica l equilibrium, we can calculate patterns of mineralization for various initial and boundary conditions, so as to compare and contrast possible geochemical mechanis ms of MVT ore for mation (Sverjens ky, 1986). Cooling of a metal-sulfid e satura ted fluid has bee n pro posed as one mechanism of ore formation. Figure 4.22 shows the results from a reactive-flow experiment in which meta l-bea ring brine, initiall y in equ ilib-

Hydrogeologic Simulations: Case Studies of Basins

Galena Concentration

167

100,000 years

(/) 100 a: w

tu

:E

50

o

500

1000

1500

2000

METERS

5.0E-04

0.00

moles/L bpm

150 (/) a: w

100

:E

50

tu

100,000 years

Sphalerite Concentration

o -l.--.-.-----.----.- .. --.--,..-=;=--.-----.----r--,,---.,---.-- ,1500 1000 500 o

.,-...----.----,---j

2000

METERS

0.00

1.0E·02

2.0E·02

moles/L bpm

FIGURE 4.22 Southeast Missouri reactive-flow Simulation Ishowing mineralization ponerns for galena and sphalerite, oher 100,000 years of flow. Mixing by transverse dispersion, pHshift, and cooling ore the primary reasons formineralization along the sandstonEHfolomite cantoct. Concentration unit> ore moles per liter of bulk porous media Ibpm).

rium with galena and sphalerite, cools as it moves over th e basem ent high. Th e chemical eq uilibrium ca lculations co nside r mass trans fer between 15 chemical components, 68 secondary species and co mplexes, and 19 min eral phas es, in a numeric al grid with 20 rows and 17 columns of finite eleme nts . Sulfide min erals precipitate mostl y a long th e sa ndsto ne- dolomite co ntact (Figure 4.22). We assumed th e sa nds tone aquifer initi ally contain ed a bo ut 32 moles of quartz per liter of bulk porous media (bpm). After 500,000 yea rs of flow, about 0.5 mol/L bpm are pr ecipitat ed downstream in th e sands to ne, with an orde r o f magnitude sma ller concentration deposited in th e dolomite

168

Hydrogeology and Geochemistry of Ore Genesis in Sedimentary Basins

bed. Because the fluid must be very close to saturation with respect to metal sulfides for simultaneous transport of metals and sulfide in the aquifer, hydrodynamic dispersion across the contact boundary produces galena and sphalerite precipitation with dolomite dissolution because of the assumed contrast in fluid composition. This dispersive flux includes metals and hydrogen moving from the sandstone into the dolomite, and sulfide and hydroxide moving from the dolomite into the sandstone, with both mechanisms causing precipitation. As a result, insufficient metal remains in solution to concentrate as an ore deposit where the fluid is focused over the basement high and further cooled by conduction. The saturation concentrations in the sandstone are [Pb] = 1.1 x 10- 5 M (2.3 ppm), [Zn] = 2.8 x 10- 4 M (18.3 ppm), and lH 2S] = 2.4 x 10- 5 M (0.8 ppm). Initial concentrations of 5.0 x 10- 7 M were assumed for both metals in the dolomite bed. For this trial simulation, the bulk of metal transport in the aquifers is attributable to chloride complexing due to the briney composition of the fluid, low pH of 4.4-4.5, and reduced oxidation state with (02 = 10- 55. In another numerical experiment, we examined the effect of reacting a metal-bearing brine with an iron sulfide-bearing lithology, such that ore is precipitated by the replacement of earlier formed, diagenetic iron disulfides. Figure 4.23 shows a pyrite-bearing dolomite facies fringing the basement high, viewed for a west-east section across the Viburnum Trend. The dolomite only contains 2.4 x 10- 3 mol of pyrite per liter of bulk porous medium, or about 0.01 vol % . We have assumed the aqueous concentrations of lead (2.3 ppm) and zinc (18.3 ppm) are the same as in the previous example, but in the sandstone only. Much smaller metal concentrations are assigned for the carbonate beds (such that the Lamotte Sandstone is the sole source of base metal) and [HsS] =2.4 x 10- 7 M (0.008 ppm) is present in the sandstone. The hydraulic and thermal properties and boundary conditions are the same as in the previous case, with the limestone bed assigned Kl s = 15 m/yr and porosity cP =0.20. Temperature patterns are nearly the same as before. Groundwater flow is focused more through the Lamotte Sandstone now, assuming the limestone is not a regional aquifer. Darcy flow rates near the sandstone pinchout are at a maximum of 1.1 m/yr in this case. The reactive-flow patterns in Figures 4.24 and 4.25 show that appreciable quantities of both galena and sphalerite are precipitated within the dolomite facies, at the expense of pyrite, which is completely dissolved after 50,000 years of flow (Figure 4.26). Sphalerite precipitates first and is later replaced by galena, but both replace pyrite, and galena appears to dissolve less quickly after long flow periods (Figure 4.24). However, once the stationary pyrite source is depleted, continued reaction and transport cause the ore mineral fronts to propagate through the flow field. Regional flow would destroy the ore field in this geochemical scenario if mass transport is allowed to operate for time scales greater than a few million years. The replacement process creates about 1.75 x 10 8 mol (4.2 x 10 4 tonnes) of galena and 1.7 x 10 9 mol (1.7 x 105 ronnes) of sphalerite over 50 km of the Vibrunum Trend, but

Hydrogeologic Simulations: Case Studies of Basins

A 150

Hydrostratigraphy

o

B

169

1000

500

1500

2000

METERS

Stream Function

1 50 -r:...----....:.....~....:.....------------------~

en 100

a::

w

tW ~

50

o

C

500

Temperature

150 --.----=-----

1000

1500

2000

METERS

- - - - -- - -- - - - - - -- --...,

"1-- -- - - - - 111 °C - - - --

en 100 a:: w tW

~

-=!==========.. : . . .:= ===.J-- - -- - - 11 3°C -

-

-

-

-

50

o

500

1000

1500

2000

METERS

FIGURE 4.23 Southeast Missouri reactive-flow Simulation II. Mass.!Josed sneom function in (b) shows flowlines and is contoured with the interval of 10,000 kg/m-yr (from Garven, 1995).

these masses are much sma ller th an the tens of millions of tonnes of me ta l sulfides observed in the district. The ore masses computed here are limited by th e low content of pyrite (0.01 % by volume, or [FeS21 = 8.75 x lOS mole) in the dolomite bed, which is the assumed source of reduced sulfur. We could estimate the time required to account for total lead mineralization in the Viburnum Trend, assuming an unlimited amount of reduced sulfur was avai lable to precipitate about 2 ppm of Pb transported in a brine. About 3 ro nnes/ yr of lead meta l would be deposited over th e entire length of the Viburnum Trend, ass uming a 25-m thick by 50-km wid e min era lization zone in the flow calculatio n an d a I-rn/yr lateral flow ra te. If the Vibu rnum

Hydrogeology and Geochemistry of Ore Genesis in Sedimentary Basins

170

150

Gal ena Concentration

10,000 years

a: 100

CJ)

w

Iii :::E

50 0

I 0

5 00

1000

150 0

2000

METERS

150,00 0 years

150 CJ)

a: 100 w

IW

:::E

50 0

I 500

0

I 1000

I 15 0 0

I 2000

METERS

150 CJ)

a:

--

--

1,000,000 years

---

100

w

IW

:::E

50

0 -, 0

500

1000

1500

2 000

METERS

_.. - .~

_. - -

0.00

1.0E ·03

moles/L bpm

FIGURE 4.24

Southeast Missouri reactive-flow Simulafion II showing mineralizafion pattern for galena at three fime

steps.

Trend cont a ins 30 million tonnes of lead (Anderson, 1991), then at least 10 million years of flow would be needed to acco unt for t he mass of minera lization , unless a larger concentration of lead and reduced sulfur were available or larger ra tes of flow were possible. No do ubt our calcu lat ion is conservative, as larger flow rates and metal concentrations are indeed geologically feasible for basinal brines (Figure 4.11). For example, increasi ng both flow

Hydrogeologic Simulations: Case Studies of Basins

0 .00

5.0E -03

171

1.OE-02

moles/L bpm

FIGURE 4.25 Southeast Missouri reactive-flow Simulation II showing minerolization pattern for sphalerite at three time

steps.

rate and metal co ncentra tion by one order of magnitude (10 ml yr and 20 ppm Pb ) would imply only 100,000 years of flow is required to account for Pb minera lizatio n in the Viburnum Trend. Th is flow period may represent a minimum duration , as ep isodes o f ore deposition pro ba bly occ urred over l OS-10 6 yea rs. Alth ough pr eliminary in nature, th ese two calc ulations serve to illus-

Hydrogeology and Geochemistry of Ore Genesis in Sedimentary Basins

172

Pyrite Concentration

150 en a:

100

w

IW

:;:

50 0

'-

1- '- ' 500

- ..-.

I

1000

1500

I

2000

METERS

10,000 years

150 en a:

100

w

IW

:;:

50 0

'-

0

I

500

1000

2000

METERS

30,000 years

150 en a:

100

w

I-

W

:;:

50 0-, 0

500

,

I

1000

1500

2000

METERS

0.00

2.5E-03 moles/L bpm

FIGURE 4.26

Southeast Missouri reactivlHlowSimulation II showing mineralization pattern for pyrite at three time

steps.

rrate the nature of two-dimensional transport-controlled reactions at the ore-deposit scale. Ultimately, fur ther modeling of th e co upled processes of grou ndwa ter fl ow, heat tra nsport, an d reactive tran sport wi ll help provide additional insight into the mechanisms producing st ratabo un d deposits, not previously realized thro ugh traditiona l reaction -path calculations or on edimensional mod els.

Hvdrogeologic Simulotions: Case Studies of Basins

173

Unconformity·Type Uranium Deposits Middle Proter ozoi c un conformity-type uraniu m deposits accou nt for mo re than 25 % percent of the w orld's prove n reserves. T he majority of these deposits are found w ith in th e Ath ab asca Basin, no rt hern Saskatchewan, Canada, and th e McArthur Basin , N orthern Terr ito ry, Austra lia. Economic deposits range in size fro m a bo ut 10 ,0 00 to 210,000 to nnes as UO ) . High-grade uraninite dep osits are associated w ith unconformabl e co ntacts between Archean-Early Pr oter ozoic metasedimentary roc ks and overlying Middl e Proterozoic sandstones. The geo logy and geochemistry of the uranium deposits and th eir unique associa tio n with the Precam brian have been well studied. There is no gene ra l ag reeme nt, however, regardi ng their genesis, du e to incomplet e kn owledge regard ing possible hydrol ogic systems that may have ex isted at th e tim e of th eir for mation. Ma ny work ers favor a diagenetic-hydrothermal origin for th e formati on of the dep osits in which warm basinal brines enco unter red ucing condition s at or near th e uncon formity as fluids circulate within th e basin (Figure 4. 7). PROTEROZOIC SANDSTONE BASINS

Th e gro un dwa ter flow syste m envisaged in the diagenetic-h yd rothermal theory has received little hydrodynamic stu dy, desp ite the key ro le it plays in transpo rting uranium to th e site o f dep osition. Severa l importa nt q uestions remain un answer ed , such as th e extent, du rat ion , and dr iving mec ha nism behind th e flow system. Darnley (198 1) pr op osed that ur an ium ore fo rmation was associ ated with high heat fl ow ca used by radioac tive decay with in granitic intrusion s (Fehn et a I., 1978). The high heat fl ow would appa rently drive hydrothermal con vect ion cells for hundreds of million s of yea rs (Darnley, 1981). M agm atic cooling following intrus ion is a lso ca pa ble of driving free convecti on, but fo r sho rte r period s o f time (Norto n and Knight, 1977). Hoeve and Quirt (19 84) suggest th at localization of the base ment heat flux by more th er mall y co nd uc tive gra phitic layers may have co ntro lled th e positi on and shape of indiv idu al co nvective cells with in the th ick sa ndsto ne a quifer. Raffensp erg er (19 93) exa mined the po ssib ility of bu oyan cy forces d riving groundwater flow within Proter ozoic basins, using co upled mathem atical models of groundwater flow and heat tran sport. His results suppo rt th e idea that free convecti on is a viable mech ani sm for driving region a l-scale gro undwater flow with in sa ndsto ne- rich sedimenta ry basins. But his mod el only requires somewhat average basement heat fl ow, witho ut additio na l heat provided loc ally by intrusion of pluton s. Figure 4.27 illustra tes a typical model showing free-convect ion cells develop ed across th e ent ire basin. Severa l kilometers o f crystalline basem ent (not show n) form th e base of th e numerical mesh, a nd a relatively low crusta l heat fl ow of 60 mW 1m2 is uni formly applied ac ro ss th e base o f th e mod el. Unlike the hyd rogeol ogic sim ulations of MVT o re form ation in foreland basins, there is no to pogra phic slope to

174

Hydrogeology ond GeochemislTy of are Genesis in Sedimentory Bosins

A Hydrostratigraphy "" ' .1-' 01 '. "

, j' l

"· " 1

,,"

.L

, I'

'~ ,

"

"I I I

.7 'F~nl arged

1. ,,·- ," "' 1

be low

B Stream Function 10r-- - - - - - - - - - - - - - - -- - - ----, 9·

5

4 !, 400

....~..__----_....j

,.....::!"""'~.L..:::.:>::::.l~:.L..::.,;;:;;;::.z....:::

500

700

800

C Temperature

10.,-- -- - - - - - - -- - -- - - - - - - - --, 9

5 4!. 400

.,........,..........w......>.L...LLI...JJ.,>-....l.-L.J..Jl.L.-1.l...l.....L-. . _ _ - - - - - - - l 700 500 800

Buoyoncyilrivenfluid convection simulo~ on for 0 Proterozoic bosin. (0) Stpple pottern represents sondstone oquifer ond lined ond brick potterns represent 10W'permeobility morine shole ond corbonote. (b) Moss-bosed sneom functionshowing flowlines contoured ot on intervol of 50,000 kg/m-yr. (c) Temperoture contours with on intervol of 200 ( (hom Roffensperger, 1993). FIGURE 4.27

the top surface of this basin profile so gravity-driven flow is excluded as a driving mec hanism in this hydrologic model. A hydraulic conductivity of K.", =100 rn/yr is assigned to th e sa nds tone laye r, with a porosity cj> =0.20. Permeabi lity is a hu nd red tim es lower in th e overlyi ng forma tio ns of marine sha le and ca rbona tes. All of th e beds were assigned a vertical permeabi lity of the along-bedding values. Darcy flow rates attain a maximum of 0 .28 m/yr in the sandstone sequence, where the effective Rayleigh number Ra = 19 . Horizontal hydraulic conductivities as low as 50 m/yr were shown to be sufficient to drive centimeter-per-year fluid convection to depths of 3-6 km. T he most important factors influencing flow patterns are basin geometry and hyd ro stratigraphy. Red ucin g tota l basin th ickn ess below 3 km effectively shuts off the free convection system, indicating the importance of both basin dept h and sandstone thickness. The permeability and anisotropy of the sandstone were found to control flow rate and convective cell geometry, respectively. Base-

nk

Hydrogeologic Simulations: Case Studies of Basins

17S

ment heterogeneities in permeability and th ermal co nductivity appea r to have little influenc e on basin-scale pattern s of free convect ion. M ost imp ortant , the small variations in thermal conductivity du e to th e presence of gra phite in the basement rocks does not appear to affect th e location of individua l cells. With horizontal hydraulic conductivities of 50 mlyr or mor e, free convec tion is a viable mechanism for long-term and large-scale fluid movement and mass transport in the Proterozoic Athabasca and McArthur Basins . Given the scale and lithology of th e bas al sa ndsto nes within these basins, a horizontal hydraulic conductivity of 300 mlyr (1 darcy ) or greater is not unre asonable (Garven, 1986; Davis, 1988 ). Furthermore, th e results ind icate that centimeter-per-year flow rates can develop, even if vert ical hydraulic conductivity in the sandstone is as low as 0.5 mlyr (less than 2 md ).

Reactive Flow. Several lines of geologic and geoch emical reasoning support a model of free convection in Proterozoic basins. Alteration halos surrounding the deposits, which usually include quartz dissolution, mimic the pattern of isotherms associated with upwelling limbs of free convection cells. For example, the thick Manitou Falls Formation sandstone in the Ath aba sca Basin contains abundant quartz cement (Ramaekers, 1990). This suggests fluid flow by free convection, because flow through temp eratu re gradients will preferentially dissolve and precipitate qu artz. Finally, the similarity of isotop ic measurements between deposits throughout the Athabasca Basin (Wilson and Kyser, 1987; Kotzer and Kyser, 1990a,b) suggests the circulation of a bas inaltype brine that was regional in extent. Does free convection of uranium-bearing chloride brine, under mod erately oxidizing conditions, precipitate ore -grade quantities of uraninite in the vicinity of the unconformity between basin sandstones and basem ent graphite schists? Raffensperger (1993) conducted dozens of reactive-flow simu lation s to answer this qu estion. Each geochemical transport simulation was repr esented by 14 chemical components, 70 secondary aqueous species, and 21 mineral phases . He found that moderate starting concentrations of total aqueous uranium (less than 30 ppm) in th e form of uranyl-ehloride complexes are sufficient to produce a 1-km long or e body with 130,000 tonnes D0 2 within 500,000 years (Figure 4.28). Basem ent permeability and redox potential exert a stro ng influence on the formation of unconformity-type deposits. The contrast in permeability between basement and graphite-rich fractured basement provides for focusing of groundwater flow upward within the graphite unit. This also produces a stationary redox front near the unconform ity. According to the conceptual model of free convection, diag enetic changes accompany ore formation, producing alt eration halo s obs erved to be associated with th ese ores. Mineral s most often associa ted with th ese altera tion halos include muscovite (sericit e), chlorite, qu artz, hematite, and pyrite. Figures 4.29 through 4.31 show th e amounts of alteration products precip itated at three points in time. Chlorite is th e most abundant alt eration pr oduct at the unconformity, forming a halo by 250,000 years, and thi s halo persists for

Hydrogeology ond GeochemislTy of Ore Genesis in Sedimentary Basins

176

6000

I

I-

.

I

I

- - -- - - - - - - - - - - - - -... ... - -- - -- -- - --- ----------... -- - - - --- -- - - " " "

~

I-

/ I

a: w

I- 3000 W

~

2000 1000 0

..

,

~

~~cra;

Sandstone

]

-j1~50,OOO -sl

:::------~~ .. - - - - - _.-" . . "' ' ' \~=iiiiil;-:::\:=-~~~ ~-=-{: : :~O~ : =,: -Basement ~ \-- - ----- ---:~- 500

/

Graphite

/

~

/

I

2000

0

.

-,

~

4000

0.000

.

Shale

5000 -

(J)

.

I

4000

6000

8000

0.001

.

0.002 moles! Lbpm

0.003

-

.

100 00

0.004

0.005

METER S

Unconformity·type reactive-flow simulation, with expanded view of one flow cell near a graphitic shear lone in the basement. Shaded tcne shows uraninite mineralization pattern after 500,000 years. Concentration units ore moles per liter of bulk porous media. SlJeamfundian is contoured at the interval of 25,000 kg/m-yr (100 kg/ m-yr for dashed lines), with flow clockwise for this convectian cell with upwelling centered aver the graphitic shear lone.

FIGURE 4.28

long periods of time (Figure 4.29) . Significant chlori te precipitation occurs in a small zone just below the deposit, unlike muscov ite, which precipitates both within the ore and above it (Figure 4.30) . After 500,000 years, muscovite forms a precipitation ha lo surrounding the uranium deposit, which persists through at least 1 million years of flow. T his halo develops first as a narrow band along the unconformity due to advec tive and dispersive mass transport across the boundary, but then migrates to envelop the ore body. O ther mineralization fronts such as hematite also migrate to form a ha lo surroundi ng the ore deposit (Figure 4.31 ). Hydrogeologic simulations by Raffensperger and Carven (1995a) also eva luate the possib le role o f gravity-driven flow, bur they found th e predicted alteration patterns to poorly fit field observations.

SUMMA RY We ag ree with th e quo tation introduced at the beginning of th is ch apter that hydrogeologic ana lysis pro vides a cha llenging area of st udy. The practical benefits o f using hydrogeologic modeling as an exploration tool have yet to be realized, but hopefully exp lorationists will apply this new technology in the years ahead. Major advances have been made over the past decade, particularly in th e hydrologic sim ulation of ore-forming systems in sed imentary basins and in the quantitative understanding of the forc es driving large -scale fluid migration. We are now ab le to mathematically model coupled transport processes of fluid flow, heat flow, an d chemica l mass transfer. Current appli-

Hydrogeologic Simulations: Case Studies of Basins

6000

en

50 ,o00 years

Ch lorate Precipltation I

I

I

I

I-

-

I-

-

I-

-

2000 I-

-

4000

177

cr: w

ti:i

:::E

-

o

-

~

o

I

I

I

I

2000

4000

6000

8000

10000 0.00

METERS

250 ,000 years

6000

en

I

I

I

I

I-

-

4000 I-

-

cr: W

ti:i

:::E

2000

o

II-

-

-

-

l

-

o

I

2000

I

I

4000

6000

0.10

molesl Lbpm 0.15

0.20

I

8000

0.05

10000

MET ERS

en

0.25

, years 500000

6000

I

I

I

I

-

-

-

I-

-

-

-

2000 I-

-

4000

cr: w

ti:i

:::E "

o

o

I

I

I

I

2000

4000

6000

8000

10000

METERS

FIGURE 4.29 Unconformity-type reactive-now simulation showing amount of chlorite precipitated at three time steps.

Solid lines indicate boundaries of hydrostratigraphic units. The shaded concentrations are expressed as moles per liter bulk porous media.

• Hydrogeology and Geochemistry of Ore Genesis in Sedimentory Bosins

178

6000

Muscovite Precipitation I

50,000 Years

I

I

I

-

-

-

-

-

-

-

-

.

4000

2000

o

o

I

I

I

I

2000

4000

6000

8000

10000

0.0

METERS

250,000 Years

6000 ,--..----r--..----r--.-------,,-----,r-'---,-----,----,

0.1 C/)

4000

a: w Iw

moles! Lbpm

:iE 2000

0.2

4000

6000

8000

10000

METERS 0.3

500,000 Years

C/)

4000

a: w

ti:i

:iE

2000

2000

4000

6000

8000

10000

METERS

FIGURE 4.30 Unconformity-type reoctivfr-flow simulotion showingamount of muscovite precipitated at three time steps. Salid lines indicate boundaries of hydrostratigrophic units. The shaded concentrotions ore expressed as moles per liter bulk porous media.

Hydrogeologic Simulotions: Case Studies of Basins

6000

en a:

Hematite Precipitation

.

50,000 Years

I

I

I

I

-

-

-

-

i-

-

2000 l-

-

4000

w

Iii ~

,-

-

I'

o

179

o

I

I

I

I

2000

4000

6000

8000

10000 0.000

METERS 250,000 Years

6000

I

I

I

I

~

en a:

4000

w

Iii ~

2000

-

-

-

-

-

-

-

-

0.005

0 .0 10

molesl Lbpm

~.

-

o

-

o

I

I

I

I

2000

4000

6000

8000

0.Q15

10000

METERS 0 .020

500 ,000 Years

6000

I

I

I

I

l-

en a: w

i-

-

i-

-

I-

-

2000 I-

-

4000

Iii ~

~ II o

o

-

.

I

I

I

I

2000

4000

6000

8000

1000 0

METERS FIGURE 4.31 Unconformity-type reactive-flow simulation showing amount of hematite precipitated at three time steps. Solid lines indicate baundories of hydrastlOtigraphic units. The shaded concentrations ore expressed as moles per liter bulk porous media.

180

Hydrogeology ond GeochemislTy of Ore Genesis in Sedimentary Bosins

cations with standard computer systems are restricted to somewhat simplified rninera logical systems for two-dimensional flow fields, but this limitation will disappear in the next few years as even faster computer processors appear in laboratory workstations. Hydrogeologic studies have shown that gravity-driven brine migration has indeed the potential to be the most important hydrothermal mechanism for stratabound ore form ation in the Earth's crust. Although compactional processes result in transient overpressures in rapidly subsiding basins and thrust belts, these flow fields are too weak to explain many of the largest Pb-Zn ore districts situated near the margins of foreland basins . This conclusion is particularly relevant to the MVT ore deposits in Phanerozoic strata of North America where orogeny created the topographic relief necessary to develop continenta l-scale flow systems. These giant flow systems were most vigorous in the ea rliest stages after subaerial emergence of the foreland, but flow rates decreased gradua lly over time, la rgely due to erosiona l processes red ucing the topographic gradient driving fl ow. Stratabound deposits situated in rifts may be formed largely as a result of thermally driven free convection, if thick aquifers are present. For example, free convection played a more important role for uranium ore formation in the Precambrian because unusually thick sandstone and conglomerate aquifers were present in the basins. Modeling techniques and hydrogeologic understanding of sedimentary basins have evo lved to a stage now where it is possible to apply this new advance in hydrogeology to critical thinking of ore deposit genesis and to basic strategies in exploration. Good examples of the former were presented here for the ores of the Mississippi Valley-type lead-zinc and for unconformity-type uranium. Many other types of sediment-hosted ores await to be studied hydrologically. Much more work needs to be done .

ACK NOWLEDG MENTS \V/e greatly appreciate Dimitri Sverjensky's comments on the manuscript and Martin Appold's assistance wi th some of the reactive-flow sim u lations . Keith Loague and Chris Neuzi l provided he lpful comments on parts of the chapter deal ing with fract ure hydrology. Ja y Gregg a nd La rry Cath les kind ly provided prepri nts of research papers, w hic h were he lpful. Gordon Hodge deserves credit fo r his skill at preparing many of the illustrations. An d finally, the first a uthor is grateful to Allan Freeze for steering his hydrogeologic research to the problem of ore formation 15 years ago, and to jozsef Torh for his friendship and mentoring in the study of large-scale groundwater flow in sedimentary basins. The work reported in this chapter was supported by grants from the National Science Foundation, the Nuclear Regulatory Commission, and the Petroleum Research Fund of the American Chemical Society. Computational support was provided through industry sponsorship from Amoco, Arco,

References

J8 J

Conoco, Mobil , Texaco, a nd Unocal , a n d through eq u ip ment grants from the N ational Sci ence Foundation.

REFERENCES Ague,]. J., a nd G. H. Brimhall (19 89) Geochemical modeling of steady sta te fl uid flow and chemical reaction during supergene enrichment of porphyry copper deposits: Econ, Geol. 84 , 506-528. Anderson, G. M . (1991) Organ ic maturation and ore precipit ati on in southeast M issouri: Eco n, Geol. 86 , 909-926. _ _ and G. Garven (198 7) Sulfate- sulfide-carbonate associa tio ns in M ississippi Valley-type lead-zinc deposi ts: Econ. Geol. 82, 482-4 88. _ _ and R. W. Macqueen (1982) Ore deposit modcls-6. Mississippi Valley-type lead-zinc deposits: Geosci. Canada 9, 108-11 7. Arnold, B. W., and J. M . Bahr (1992) Spatial distr ibuti on of effective permeability and fluid flow focusing in fractured carbonate strata relat ed to MVT mineralizat ion, southwest Wisconsin: Geol. Soc. Am. Abst. Prog. 24, A282. Baskov, E. A. (1987) Th e Fundame ntals of Paleohydrogeology of Ore Deposits: Berlin: Springer-Verlag. Beaumont, c., R. Boutil ier, A. S. M ackenzie, and ]. Rullkotrer (1985) Isom erizat ion and a ro rnatiza rion of hydr ocarbon s and the paleoth ermometry and buri al histor y of Alberta for eland basin: Am. Assoc. Petrol. Geol. Bull. 69, 546-566. Bethke, C. M. (1985) A numerical model of compaction-driven groundwater flow and heat transfer a nd its applicatio n to the paleoh ydrology of intracr atonic sedimentary basins: J. Geophys. Res. 90B , 681 7-6828. _ _ (1986) Hydrologic constraints on the genesis of the Upper Mississippi Valley district from Illinois Basin brines: Econ. Geol. 81,233-249. _ _ (1989) Modeling subsurface flow in sedimenta ry basins: Geol. Rundscb. 78, 129-154. _ _ and T. F. Corbet (1988) Linear and nonlin ea r solutions for one-dimensiona l compaction flow in sed imentary basins: Water Resour: Res. 24 , 461-467. _ _ and S. Marshak (1990) Brine migrati on s across North America-the plate tectonics of groundwater: Ann. Rev. Earth Planet. Sci. 18, 287-3 15. - - J. D. Reed , and D. F. Oltz (199 1) Long-range petroleum migrat ion in th e Illinois Basin: Am. Ass oc. Petrol. Geol. Bull. 75, 925-945. Bjorlykke, A., and D. F. Sangster (1981) An overview of sandsto ne lead deposits a nd their relation to red-bed copper and carbonat e-hosted lead-zinc deposits: Econ. Geol. 75th Anniv. Vol., 179-2 13. Brannon,] . c., F. A. Podosek, J. G. Viets, D. L. Leach, M . Goldhaber, and E. L. Row an (1991) Strontium isoto pic constra ints on the or igin of o re-forming fluids of the Viburnum Trend, southeast M issouri: Geochim. Cosmocbim, A cta 55 , 1407-141 9. Bredehoeft, ]. D., K. Belitz , and S. Sharp-H ansen (1992) Th e hydr odynam ics of the Big Horn Basin: a study of the role of faults: Am. Assoc. Petrol. Geol. Bull. 76, 530-546.

182

Hydrogeology and GeachemislTy of Ore Genesis in Sedimentary Basins

_ _ , C. E. N euzil, and P. C. D. Milly (1 983) Region al fl ow in th e Dakota Aquifer: a study of the role of confi ning layers: U.S. C eol. SIIrv. Water Supply Pap . 223 7. _ _ and D. L. Norton (1 990) Ma ss and energy transport in a deforming Earth 's crust. 11/: Th e Role of Fluids ill Crustal Processes . Wa shington , DC: National Acad em y Press, pp. 27-41. Cacas, M. c., E. Ledoux, G. de Marsily, and B. Tillie (1990) M odeling fracture flow with a stoc has tic discret e fracture netw ork: ca libra tio n and validation. 1. Th e flow model: Wat er Res ourc, Res. 26, 479-489. Capuano, R. M . (1993) Evidence of fluid fl ow in microfractures in geopressured sha les: Alii . Assoc. Petrol. C eol. BIIII. 77, 1303-1314. Carrigan, C. R., G. C. P. King, G. E. Barr, and N. E. Bixler (1991) Potential for watertabl e excurs ions induced by seismic events at Yucca Mountain, N evada: C eology 19, 115 7-1160. Cathles, L. M. (1993) A discussion of flow mechani sms responsible for altera tio n and mineralizati on in the Cambrian aquifers of the Ou achita Basin-Ozark system: In: Diagen esis and Basil/ Deuelopment; A. D. Horgury and A. G. Robinson (eds.). Tulsa, OK: Am. Assoc. Petrol. Geol. Mem., pp. 99-112. _ _ , S. Oszczepalski, and E. C. Jowett (1993) Mass balance evaluation of the late diagenetic hypothesis for Kupferschiefer Cu mineralization in the Lubin Basin of southwestern Poland: Econ . C eol. 88, 948-956. _ _ and A. T. Smith (1983) Thermal con str aints on the formation of Mi ssissippi Valley-type lead-zinc deposits and their implications for episodic basin dewatering and depo sit genesis: Econ. C eol. 78, 983-1002. Chamberlin, T. C. (1882) Th e ore depo sits of southwestern Wisconsin. In: C eol ogy of W isconsil/: Wis e. Ceol. Surv. 4, 365-571. Clendenin, C. W., and M . J. Duane (1990) Focused fluid flow and Ozark Missis sippi Valley-type deposits: C eology 18, 116-119. Corbet, T. F., and C. M. Bethke (1992) Disequilibrium fluid pressures and groundwater flow in the Western Can ada sedimentary basin : J. Geophys. Res. 97, 7203-721 7. Cox, G. H. (1911) The or igin of the lead and zinc ore s of the Upper Mississippi Valley district: Econ. C eol. 6, 427-448. Cumming, G. L., J. R. Kyle, and D. F. Sangster (1990) Pine Point: a case history of lead isotope homogeneity in a Mississippi Valley-type district: Econ. C eo!' 85, 133-144. Daubree, A. (188 7) Lcs Eaux Souterrain es, aux Epoques Anciennes et l'Ep oque A ctuelle, 3 vols. Paris: Dunod. Darnley, A. G. (1981) The relationship between uranium distribution and some ma jor crustal features in Canada: Min eral. Mag. 44, 425-436. Davi s, S. N. (1988) Sandstones and shales. In: Hydrogeology, The Ceology o f N orth America, Vol. 02, W. Back, J. S. Rosenshein, and P. R. Seaber (eds.). Boulder, CO: GeoI. Soc. Am. , pp. 323-332. Deloule, E. and D. L. Turcotte (1989) The flow of hot brines in cracks and the formation of ore deposits: Econ. C eol. 84, 2217-2225 . De M arsily, G. (1986) Quantitative Hydrogeology: New York: Acad emic Press.

a

References

183

Deming, D. (1992) Ca tas tro phic release of heat and fluid fl ow in the conti nenta l crus t: Geology 20 , 83-86. de Wit, R., E. C. Gro n berg, W. B. Richards, a nd W. O. Richmond (1973) Ta thlina area, Distri ct of M ackenz ie, In: The Future Petroleum Provinces of Canada-Their Geology and Pot en tial, R. G. McCrossan (ed.). Ca lgary: Ca nadia n Society of Petrol eum Geolo gists, M emoir 1, 187-212. Domenico, P. A., and F. W. Schwa rtz (1990) Physical and Chemical Hydrogeology: New York: Wiley. Dozy, J. ]. (19 70) A geologic model for the genesis of the lead- zinc ores of the M ississippi Valley, USA: Trans. lust, Min. Me tall. 79, BI 63-170. Duane , M. ]., and M. J. de Wit (1988) Pb-Zn ore deposits of the no rthern Ca ledonides: products of co nt inental -scale fluid mixing and tecton ic exp ulsion du ring continental colli sion: Geology 16 , 99 9- 1002 . Etheridge , M . A., V. J. Wall , and R. H . Vernon (1983) Th e role of the fluid phase during regional met am orphism a nd deformat ion: J. Metam orphic Geo l. 1, 205-226. Fehn, U., L. M. Cathles, and H . D. Holland (197 8) Hydr oth ermal con vection and uran ium dep osits in a bno rma lly radioactive pluton s: Econ, Geol. 73, 1556- 1566. Ferry, ]. M ., a nd G. M. Dipple (1991) Fluid flow, mineral reactions, and metasom atism: Geology 19,211-214. Forster, C. B., a nd J. P. Evans (199 1) Hydrogeolog y of thrust faults an d crysta lline thrust sheets: results of combined field and modeling studies: Geopb ys. Res. Lett. 18, 979-982. Fowler, A. D. , a nd M . T. Anderson (1991) Geopressured zones as pro ximal so urces of hydrothermal fluids in sedimentary basins and the origin of M ississippi Valleytype deposits in sha le-rich sequence: Trans. lnst, Min . Metall. 100 , BI4-BI 8. Freeze, R. A., and J. A. Ch erry (1979) G roundwat er: Englewood Cliffs, N]: Prenti ceHall. Galloway, W. E. (19 78) Uran ium mineralizat ion in the coas ta l-plain fluvial aquifer system, Catahoul a Fo rmat ion , Texas: Econ. Geol. 73, 1655-16 76. Garven, G. (1985) The role of region al fluid flow in the genesis of th e Pine Point deposit: Econ. C eol. 80, 307-324 . _ _ (1986) The role of regional fluid flow in the genesis of the Pine Point deposit-a reply: Econ. C eo!. 81, 1015-1020. _ _ (1988) Effects of tr an sient fluid flow on the thermal histor y of MVT min eralizatio n: C eol. Soc. Am. A bst. Prog. 20, A364. _ _ (1989) A hydrogeologic model for the formati on of the gia nt oil sa nds deposits of the Western Can ad a sedimenta ry basin: Am. J. Sci. 289, 105-16 6. _ _ (1994) Gen esis of stra ta bo un d or e deposits in the M idco nti nent Basins of North America . 1. The role of region al gro undwater flow-a reply: A m . J. Sci. 294, 760-775. _ _ (1995) Continental-scale gro undwa ter flow and geo logic processes: Rev. Earth Planet. Sci. 23, 89-117. _ _ , S. Ge, M . A. Person , a nd D. A. Sverjensky (1993) Genesis of stra ta bo und ore deposits in the M idcontinent Basins of North Amer ica. 1. T he ro le of region al groundwater flow: Am. j. Sci. 293, 497-568 . - - and R. A. Freeze (1982) The role of regio nal gro undwa ter fl ow in the fo r-

184

Hydrogeology and GeochemislTy of Ore Genesis in Sedimentary Basins

mari on of ore deposit s in sedi me nta ry ba sin s: a quantitative a na lysis: In: Proceedings National H ydrogeology Co nference, Second Annual M eeting, Winllip cg, Manitoba, G . O zoray (ed.) Edmonton: Can. Sect. Int. Asso c. Hydrogeol. , pp. 60-67. _ _ _ and R. A. Free ze (1984a) Theoretical an alysis o f the role o f gro undwa ter flow in th e gene sis of stra ta bo und o re deposit : 1. M athematical a nd numerical model: Am. J. Sci. 284, 10 85-1124. _ _ and R. A. Freeze (1984b) Theoretical ana lysis of the role of groundwater fl ow in the gen esis o f stra ta bo und are deposits: 2. Quantitative results: Am. J. Sci. 284, 11 25-11 74 . Ge, S., and G. Garven (198 9) Tectonically induced tr ansient gro undwa te r flow in forela nd basins. In: Origin and Evo lu tion of Sedim entary Basins and Their En ergy and Mill cral Resources, R. A. Price (ed.). Am. Ge ophys. Union M onogr. Ser. 48, 145-15 7. _ _ and G. C arven (1992) H ydromechanical modeling of tectonically-driven groundwat er flow with application to the Ark om a foreland ba sin: J. G eopbys. Res. 97, 911 9-9144. Gra tz, J . E, and K. C. Misra (1987) Fluid inclusion stud y of the Gordonsville zinc dep osit, central Tenn essee: Econ . Geol. 82, 1790-1804. Gregg, J. M., and P. E. Gerdem ann (1989) Sedimentary facies , diagenesis, and are distribution in the Bonneterre Formation (Cambrian) , southeast Mi ssouri. In: Field Guide to th e Upp er Cambrian of Southeastern Missouri: Stratigraphy, Sedimentology, and Econ om ic Geology, J. M. Gregg, J. R. Palmer, and V. E. Kurtz (eds.). Roll a: University of M issouri, pp. 43-55. _ _ , P. R. Laudon, R. E. Woody, and K. L. Shelton (1993) Porosity evo luti o n of the Bonn eterre Dolomite (Cam bria n) southeastern M issouri, U.S.A. : Sedimentology 40,1153-1169. _ _ and K..L. Shelton (19 89 ) Minor- and tr ace-element distributions in the Bonneterre Dolomite (Cambrian), southeas t Mi ssouri: evidence for possible multiplebasin fluid sources a nd pathways during lead-zinc mineral ization: Bull. G eol. Soc. Am. 101 ,221-230. Gu stafson, L. B., and N. Will iams (1981) Sediment-hosted stratiform deposits of co pper, lead , and zinc: Econ. Geol . 75th Annivcr, Vol., 139-178. H agni , R. D. (1989) The so ut heast Missouri lead district: a review: In: M ississippi Valley-Typ e Mineralization of th e Viburnum Trend, Missouri, R. D. Hagni and R. M . Cove ney, Jr. (eds.). Soc. Econ. Geo/. Gu idebook Ser. 5, 12-57. Hanor, J. S. (19 79) The sedime nta ry genesi s of hyd rothermal fluids. In: G eo chem istry of H ydrothermal Ore Dep osits , 2nd ed., H. L. Barnes (ed.). New York: Wiley, pp. 137-1 72. ___ (1987) Origin and Migrat ion of Sub surface Sedimentary Brines: Tul sa, OK: Soc. Eco n. PaleontoL Mineral. , Leer, Notes for Short Course No . 2 1. H ard ie, L. A. (1990) The rule s of rifting and hydrothermal CaCI 2 brines in th e or igin o f potash eva po rites : an hypothesis: Am. }. Sci. 290, 43-106. H arrison, W. J . and L. L. Summa (1991) Paleohydrology of th e Gulf o f M exico Basin: Am.}. Sci. 291, 109-1 76. H elgeson , H . C. (1979) M ass tr ansfer among minerals and hydrothermal so lutions.

References

185

In: Geochemistry of Hydroth ermal Ore Deposits, 2nd ed., H. L. Barnes (ed.). New York: Wiley, pp. 568-610. Heyl, A. V. (196 8) The Upper M ississippi Valley base-metal distr ict. In: Ore Deposits of the United States, 1933-1 967 (C raton- Sales volume). New Yor k: Am. Insr. Mining Metall. Petrol. Eng., Vol. 1, Part 3, Cha p. 21, 431-459. Hoeve, j., and D. Quirt (1984 ) M ineralizat ion and host rock alteratio n in relation to clay mineral diagenesis a nd evo lution of the Middl e-Proterozoic, Athabasca Basin, northern Saskatchewan , Can ada: Sask. Res. Coun cil SR C Tech. Rep. 187. Holcombe, C. J. (1985) Paleoflow mod eling for sedimentar y orebod ies: Econ. Ceol. 80, 172-179. Hubbert, M. K. (1940) The theory of ground-water mot ion: J. Ceol. 48, 785- 944. jackson, S. A., and F. W. Beales (1967) An aspect of sedimentar y basin evolution: the concentration of Mississippi Valley -type orcs during the late stages of diagenesis: Bull. Can. Petrol. Ceol. 15, 383-433. jowett, E. C. (1986) Genesis of Kupferschiefer Cu-Ag deposits by convective flow of Rotliegende brines during Triassic rifting: Econ. Ceol. 81, 1823-1837. Kesler, S. E., and B. A. van der Pluijm (1990) Timing of Mississippi Valley-type min eralization: relation to App alachian orogen ic events: Ceology 18, 1115-1118. Kiraly, L. (1975) Rapport sur l'erat actuel des connaissances dan s Ie domaine des caracteres physiques des roches kartiques: Int. Union Ceol. Sci. Ser. B 3, 53-67. Knapp , R. B. (1989 ) Spatial and temporal scales of local equilibriu m in dyn am ic fluid-rock systems: Geochim, Cosmochim, Acta 53, 1955-1964. Kotzer, T., a nd T. K. Kyser (1990a ) Th e usc of stable and rad iogenic isoto pes in the ident ificatio n of fluids and processes associated with uncon form ity-type uranium deposits. In: Modem Exploration Techniques, L. S. Beck and C. T. Harper (eds .). Regina, Saskatchewan: Sask. Geol. Soc. Spec. Pub. 10, 115-131. _ _ and (1990b ) Fluid histor y of the Athab asca Basin and its relation to uranium deposits: Summary of Investigations 1990, Miscellaneous Report 90-4. Regina, Saskatchewan: Sask. Geol. Surv., Saskatchew an Energy and Mines, pp. 153-157. Kyle, j . R. (1981) Geology of the Pine Point lead-zinc district. In: Handbook of Strata-Bound and Stratiform Ore Deposits, Vol. 9, K. H. Wolf, (ed.). New York : Elsevier, pp. 643-741. Leach, D. L. (1973) Possible relationsh ip of Pb-Zn mineralization in the Ozarks to the Ouachita orogeny: Ceol. Soc. Am. Abst. Prog. 5, A269. _ _ and E. L. Rowan (1991) Comment on the paper by C. W. Clendenin and M. j . Duane, 1990, "Focused fluid flow and Ozark M ississippi Valley-type deposits." Ceology 19, 190. _ _ and (1993) Fluid inclusion studies of region ally ext ensive epigenetic dolomites, Bonneterre Dolomite (Cambrian ), southeast Missouri: evidence of mul tiple fluids during dolomitization an d lead-zinc mineralizat ion -alternati ve interpretation: Bull. Ceol. Soc. Am. 105 , 968-978. LeHuray, A. P., J. B. D. Caulfi eld, D. M. Rye, and P. R. Dixon (1987) Basement controls on sediment-hosted Zn-Pb depo sits: a Pb isotope study of ca rbonifero us mineralization in central Ireland: Econ. Ceol. 82, 1695-1 70 9.

186

Hydrogeology ond GeochemislTy of Ore Genesis in Sedimentary Basins

Lichtner, P. C. (1993) Scaling properties of time-space kinetic mass transport equation s and the local equilibrium limit: Am. J. Sci. 293, 257-296. _ _ and G. G. Biino (1992a) A first principles approach to supergene enri chment of a porphyry copper protore: 1. Cu-Fe-S subsystem: Geochim, Cosmochim, Acta 56, 3987-4013. _ _ and G. G. Biino (1992b) Genesis of Mississippi Valley-type (MVT) deposits: Geol, Soc. Am. Abst. Prog. 24, A324. Lydon, J. W. (1986) Models for the generation of metalliferous hydrothermal systems within sedimentary rocks and their applicability to the Irish Carboniferous Zn-Pb deposits. In: Geology and Genesis of Mineral Deposits in Ire/and, C. J. Andrew, R. W. A. Crowe, S. Finlay, W. M. Pennell, and J. F. Pyne (eds.). Dublin: Irish Assoc. Econ. GeoL, pp. 555-577. Mauk, J. L., W. C. Kelly, B. A. van der Pluijm, and R. W. Seasor (1992) Relations between deformation and sediment -hosted copper mineralization: evidence from the White Pine part of the Midcontinent rift system: Geology 20, 427-430. Muir-Wood, R., and G. C. P. King (1993) Hydrologic signatures of earthquake strain. J. Geoplrys. Res. 98, 22035-22058. Musgrove, M ., and J. L. Banner (1993) Regional ground-water mixing and the origin of saline fluids: Midcontinent, United States: Science 259, 1877-1882. Nash, J. T., H . C. Granger, and S. S. Adams (1981) Geology and concepts of genesis of important types of uranium deposits: Econ. Geol. 75th Anniver. VoL, 63-116. Neuzil, C. E. (1986) Groundwater flow in low-permeability environments: Wat er Resourc. Res. 22, 1163-1195. _ _ (1993) Low fluid pressure within the Pierre Shale: a transient response to erosion : Water Resour. Res. 29, 2007-2020. _ _ and D. W. Pollock (1983) Erosional unloading and fluid pressures in hydraulically tight rocks: J. Geol. 91, 179-193. Newhouse, W. H. (1933) The temperature of formation of the Mississippi Valley lead-zinc deposits: Econ. Geol . 28, 744-750. Noble, E. A. (1963) Formation of ore deposits by water of compaction: Econ. Geol. 58, 1145-1156. Norton, D. (1988) Metasomatism and permeability: Am. j. Sci. 288, 604-618 . _ _ and L. M. Carhles (1979) Thermal aspects of ore deposition. In: Geochemistry of Hydrothermal Ore Deposits, H. L. Barnes (ed.). New York: Wiley, pp. 611-631. _ _ and J. Knight (1977) Transport phenomena in hydrothermal systems: cooling plutons: Am. J. Sci. 277, 937-981. Nur, A., and J. R. Booker (1972) Aftershocks caused by pore fluid flow : Science 146, 1003-1010. Ohle, E. L. (1959) Some considerations in determining the origin of ore deposits of the Mi ssissippi Valley-Part 1.: Econ. Geol. 54, 769-789. Oliver, J. (1986 ) Fluids expelled tectonically from orogenic belts: their role in hydrocarbon migration and other geologic phenomena: Geology 14, 99-102. Ortiz, N. V., J. A. Ferenrchak, F. G. Ethridge, H . C. Granger, and D. K. Sunada (1980) Ground-water flow and uranium in Colorado Plateau: Ground- Water 18, 596-606. Palciauskas, V. V., and P. A. Domenico (1980) Microfracture development in

References

187

compacting sediments : rela tio n to hydroca rbon-matura tio n kinetic s: Am. Assoc. Petrol. Geol. BIl Il. 64 , 92 7-93 7. Pelissonnier, J. (1967) Ana lyse paleohydrogeologique des gisement s stratiformes de plomb, zinc, ba ryte, fluo rite du type" Mississippi Valley" : Econ. Geol, Monogr. 3,234-252. Person, M. A. an d G. Garven (1992) Hydrologic constraints on petroleum generation within co ntinenta l rift basins: theory and application to the Rhine Graben. Am. Assoc. Petrol. Geol. BIlIl. 76 , 468-488. Phillips, O. M. (199 1) Flow and Reactions in Permeable Rocks. Cambridge: Cam bridge Univesiry Press. Raffensperger, J. P. (199 3) Q ua nt itat ive evaluation of the hydrologic and geochemical processes invo lved in the for mation of unco nfo rmity -type uranium deposi ts. Unpublished Ph.D. dissertati o n, T he Johns Hopkins University, Baltimore, Maryland. _ _ and G. Ga rven (1995a ) The format ion of unconform ity-type uran ium ore deposits. 1. Coupled groundwat er flow and heat transport mod eling. Am . j. Sci. 295 ,581-636. _ _ and _ _ (1995 b) The forma tion of unconfo rmity -type uraniu m ore deposi ts. 2. Co upled hydrochem ical mod eling. Am . J. Sci. 295, 639-696. Ramaekers, P. (1990) Geo logy of the Atha basca Group (Helikian) in northern Saska tchewan . Saskat chewan Energy and Mines, Sask. Geol. SIIrv. Rep. 195. Rickard , D. T., M. Y. W illden, N. E. Marinder, and T. H. Donnelly (1979) Studies on th e genesis of the l a isvall sandstone lead-zinc deposit, Sweden. Econ . Ceol. 74, 1255- 12 85 . Roedd er, E. (1968) Tem per at ure, salinity, and origin of the o re-forming fluids at Pine Point, Northwest Territories, Can ada, from fluid inclusio n studies. Econ , Ceol. 63, 439-450. Rudnicki, J. W., and T.-C. Hsu (1988 ) Pore pressure change s indu ced by slip on permeable a nd impermeable faults . J. Geophys. Res. 93, 3275 -3285. Russell, M. J. (1992) Plate tecto nics and hydrotherma l ore deposits. In: Understanding the Earth , G. C. Brown, C. J. Hawkeswo rt h, and R. C. L. Wilso n (eds.). Ca mbridge: Cambridge University Press, Chap . I I, pp. 204-221. Sanford , R. F. (1982) Preliminary model of regio nal Mesozoic groundwater flow and uran ium dep osition in th e Co lorado Plateau . Ceo logy 10, 348-352. _ _ (1992) A new mod el for tabular-type uranium deposits: Econ. Geo l. 87, 2041 -2055. Sangster, D. F. (19 83) M ississipp i Valley-type deposits: a geo logical melange: In: Int ernati on al Conference on M ississippi Valley-Type Lead-Zinc Deposits, Proceedings Yolurne, G. Kisvar sanyi, S. K. Gra nt, W. P. Pratt, and J. W. Keon ig (eds .). Rolla: University of M issouri, pp . 7-19. Sharp, J. M . Jr. (19 78 ) Energy an d moment um tr ansport mode l of the O uachita Basin an d its po ssible impact on forma tion of economic minera l deposits: Econ . Ceol. 73 , 1057-1068 . _ _ and J . R. Kyle (1988) The role of ground-water processes in the formation of o re de posits: In: Hydrogeology, The Geology of North America, Vol. 02, W.

+ 188

Hydrogeology ond Geochemistry of Ore Genesis in Sedimentary Basins

Back , J . S. Rosenshein, an d P. R. Sea ber (eds.). Boulder, CO: Geol. Soc. Am., pp. 461-483 . Shelto n, K. L., R. M . Bauer and J. M. Gregg (1992) Fluid incl usion studies of regionally extensive epigenetic dolomites, Bonneterre Dolomite (Cambrian) , so utheast M issouri: evidence of multiple fluids during dol omitization a nd lead-zinc mineralization: BII//. Geol. Soc. Am. 104,675-683. _ _ , _ _ , and _ _ (1993) Fluid inclusion studies of region all y ex tensive epigenetic dol omites, Bonneterre Dolomite (Cam brian), so utheas t Mi ssou ri: evidence of mult iple fl uid s during dolomitizati on and lead-zinc min er al izati on-a reply: BII//. Geol. Soc. Am. 105 , 972-978 . Siebenthal, C. E. (1915) Origin of the zinc and lead deposits o f th e joplin region, Missour i, Kan sas, and Okl ahoma: U.S. Geol. Slim BII//. 606 . Sibson, R. H. , J. McM. Moore , and A. H . Rankin (1975) Seismic pumping-a hydr othermal fluid tran sport mechanism:]. Geo l. Soc. London 131 , 653-659. Skinner, B. J. (1979) The many o rigins of hydr othermal mineral deposits: In: Geochem istry of Hydrothermal Ore Deposits, 2nd ed ., H. L. Barnes (ed .). New York: Wiley, pp. 1- 2 1. Sno w, D. T. (1 96 9) Anisotropic perm eab ility of fractured media: Water Resourc. Res. 5, 1273-1289. Solomon, M. , and C. A. Heinrich (1992) Are high-heat-producing gr an ites essent ial to the origin of giant lead- zinc depo sits at M ount Isa and M cArthur River, Australi a?: Exp10r. Mining Geol. 1, 85-91. Steefel, C. I., and A. C. Lasag a (199 2) Putting tra nsport into wat er-rock inte rac tion models: Geology 20, 680-684. _ _ and _ _ (1994 ) A coupled mod el for transport of multiple chemica l species and kinetic precipitation/dissolution reactions with applica tion to reactive flow in single pha se hydrothermal systems: Am. ]. Sci. 294, 52 9-592. Sverjensky, D. A. (1986) Genesis of M ississippi Valley-typ e lead zinc deposits: Ann. Rev. Earth Planet. Sci. 14, 177-1 99. _ _ (1987 ) Th e role of migrating oil field brines in the formation of sedimenthosted Cu-r ich deposits: Econ, Geol. 82, 1130-1141. _ _ and G. Gar ven (1992) Tracing great fluid migrations: Nature 356, 481-482. Symons, D. T. A., and D. F. Sangster (1991) Paleomagnetic age of the centra l M issouri barite deposits and its genet ic implications: Econ. Geol. 86, 1-12. T orh , j . (1978) Gravity-indu ced cross-formationa l flow of formation fluids, Red Earth region , Alberta, Can ada: ana lysis, patterns, evolutio n: Wa ter Resourc, Res. 14, 805-843. Valley, J. W. (1986) Stabl e isotope geochemistry of met amorphic rocks. In: Stable Isot op es in High Temp erature Geological Processes, J. W. Valley, H. P. T aylor, Jr., a nd j . R. O 'Neil (ed s.) , Min eral. Soc. Am. Rev. Mineral. 16 , 445-489. Van H ise, C. R., and H. F. Bain (1902) Lead and zinc deposits of the Mi ssissipp i Vall ey, USA: A m . ln st, Min. Eng. Trans. 23 , 376-434. Walsh , M . P., S. L. Bryant, R. S. Schecht er, and L. W. Lake (1984) Precipitati on and dissolution of solids attending flow through porous media: Am. lnst , Chem , Eng. ]. 30, 3 17-328.

References

189

White, W. S. (1971) A paleoh ydr ologic mod el for mineralizat ion of the White Pine copper dep osit , northern Mi chigan: Econ. Ceo l. 66, 1-13. Wilson , M . R., and T. K. Kyser (1987) Sta ble isotope geochemistry of alteration associat ed with th e Key Lak e uran ium deposit, Ca na da : Econ. Ceol. 82, 1540-1557. Wolery, T. J. (1979) Computer sim ula tio n of chemica l equi librium betwee n aq ueo us solutions and miner als: th e EQ3/EQ6 softwa re package: Livermor e, CA: University of California Lawrenc e Liverm or e Lab or ator y, Repor t UCR L-52658. Wood , J. R., a nd T. A. Hewett (198 2) Fluid convection and mass transfer in po rou s sandsto nes-a th eoret ical approach: Ceochim. Cosmochim, Acta 46 , 1707-1713.

Chapter

s Thermal Aspects of Ore Formation lAWRENCE M. CATHlES III Deportment of Geologicol Sciences, Cornell University

As the title of this book suggests, understanding the geochemistry of hydrothermal ore deposits involves understanding the interactions between hydrologic and thermal processes. Perhaps the most important thermal interaction is the advection of heat by water moving through a permeable porous or fractured lithic matrix in which there are temperature gradients. Water movement through temperature or other physical gradients (such as pressure or salinity) drives chemical change. For example, silica may be deposited if pore waters move from areas of higher to lower temperature because the solubility of silica is temperature dependent. Thus there can be a direct connection between fluid movement through temperature gradients and ore deposition. How much ore is deposited depends on how fast the thermal gradients move through the subsurface. Pore-fluid movements cause temperature gradients to migrate in the direction of fluid flow. The pore waters may be driven by the temperature gradients themselves (convection) or by an independent process. The important point is that fluid movements cause the subsurface temperature distribution to change if it is not uniform. Some of the strongest constraints on hydrothermal ore deposition are thermal. The thermal aspects of ore formation can be investigated in a variety of ways. One way is to construct coupled numerical computer models of temperature and fluid flow and simulate the evolution of geologic systems of 191

192

Thermol Aspects of Ore Formotion

int erest. A great deal can, and has been, learned from such simulations, and more will be learned as chemical change is coupled with temperature and fluid flow. Simulations are very useful in checking more intuitive methods. Examples of recent reviews of this approach are available from a number of authors (Carhles, 1977, 1983, 1990; Norton and Knight, 1977; Norton and Taylor, 1979; Garven and Freeze, 1984a,b; Norton, 1984, 1988; Bethke, 1986; Bethke and Marshak, 1990; Bethke et aI., 1991; Burrus et aI., 1991; Furlong er aI., 1991; Hanson, 1992; Person and Garven, 1992). Computer simulations have the disadvantage, however, that most geologists have neither ready access to computers nor the deta iled knowledge of the software that is required to make such simulations. Also, the computer simulations are often complex enough in their own right that fundamenta l insights and understanding are difficult to extract. Ano ther approach that can help interpret eit her ore deposits or computer simulations is to constrai n importa nt aspects of th e ore-forming process with simple calc ulat ions that ca n be carried out wi th paper an d pencil. Calculations of this sort are often included in computer simu lation papers, a re exceptionally useful, and in some cases can be geologically more realistic than large-scale simula tions beca use the ana lysis can be focused specifically on critical aspects of the problem. Hand calculations are basically of two kinds: those that calculate the thermal consequences of fluid flow and those that define its magnitude. Combining the two is particularly powerful. This chapter compiles paper and pencil tools the author has found useful and illustrates their use by applications to several kinds of hydrothermal ore deposits.

THERMAL CONSEQUENCES OF SUBSURFACE FLUID FLOW The interaction between temperature and subsurface fluid flow is governed by a conservation-of-energy equation that may be written

aT =A -

pmCm at

2

pC]· VT + K m V T

(5.1)

where P m an d Cm are the density and heat ca pacities of the media (rock plus pore fluid) in g/cm-' and caljg_ OC, respective ly, p and C are th e density an d heat capacity of the pore fluid in the same units, T is the temperature in °C of the pore fluid and immediately adjacent rock (assumed to be the same),] is the Darcy velocity, superficial velocity, or vo lume flux (all equivalent expressions for] ) of the pore fluid in crais, Km is the thermal conductivity of the media (rock and pore fluids) in cal/cm-s-r C, A is the rate of heat generation in caljcm 3-s, and V is the Lap lacian operator. The Darcy velocity, ] , deserves special comment. It is not really a velocity with dimensions of cm /s at all but a volume flux with dimensions of

Thermal Consequences of Subsurface Fluid Flow

J93

cm3/cm 2-s. It is the number of cubic centimeters of pore water that pass a l-cm 2 section perpendicular to the flow per second. The Darcy flux, as we shall call it henceforth, is the velocity the pore fluid would have if it occupied the entire space instead of the space represented by the pores through which water is moving in the porous medium. Because the moving water occupies only the flowing porosity, cPr, the true velocity of the pore water is J!cPr. It has been assumed in equation (S.l) that the pore fluid is incompressible (e.g., V·J = OJ, and that p C is constant. These are useful and nonrestrictive assumptions. Term by term, equation (S.l) states that the change of heat in a unit volume of the porous medium over a time, dt (first term on the left), equals the heat generated within that volume over dt (first term to the right of the equal sign), plus the heat advected into the volume by fluid flow, plus the net heat conducted into the volume (last term on right).

Migrating Thermal Fronts Heat advection can be investigated with equation (S.l) by switching to a coordinate system that moves with thermal anomalies. Subsurface temperature will, in general, vary so that temperature is a function of spatial position and time. Symbolically we write T = T(x, t}, wherex is the position vector with comp~nents of ~I along the x axis lying in the i direction, and so on, so that x = XI i + x2i + X3k. Because T depends on both x and t, the total derivative of T follows from the chain rule . The total change in T equals the partial change in T with respect to t, at a fixed location in space, multiplied by the incremental change in t, plus the partial change in T with respect to space, at a particular instant of time, multiplied by the change in spatial location with time, multiplied by the incremental change in time:

DT(x,t} = ( -aT)

at

Dt+ ( -sr ) xj

e

aXi

const,

t =const.

( -aXj ) at t =const.

Dt

Rearranging and letting (aXi/at}t =consr. = v, this equation may be rewritten:

sr =DT - -v·VT

-

at

Dt

(5.2)

To understand equation (S.2), imagine an observer standing beside a moving train of refrigerator cars. If the temperature of a particular car changes, DT/Dt :I- 0. If the train is moving and the cars do not all have the same temperature so that v- VT :I- 0, then the temperature of cars in front of a stationary observer will change. Suppose the train is traveling in the x direction at one car per minute, each car maintains constant temperature so that DT/Dt = 0, but the temperature in each adjacent car (in the pos-

194

Thermal ,\spects of Ore Formation

inve x direction) increases at 1°C per car. Equation (5.2) then becomes aT/at = - V x aT/ax, and the temperature in the cars at the observer's location (Eulerian or spatially fixed coordinate) will decrease at a rate of 1°C per minute. If there was the same difference in temperature, between cars but all cars were cooling at 1DC/ m in, that is, DT/Dt = - 1, the change observed at a stationary location would be - 2° C/ m in . If the observer rode on any particular car (e.g. , moved with the fluid in a moving Lagrangian coordinate system), only the change in temperature in the car to which the observer was attached would be observed, in this case - 1° C/ m in. There is nothing particularly sacred about the exact form of the transformation expressed in equation (5.2). A coordinate system that follows thermal anomalies, rather than the fluid motion, can be selected, for example, by modifying equation (5.2) slightly so that

v >!

pC PmC/II

DT

pC pmC/II

(5.3)

and aT

at = -D-t -

I : VT

(5.4)

That equations (5.3) and (5.4) describe the movement of a coordinate system in which temperature anomalies are stationary can be seen by substituting equation (5.4) into equation (5.1). The result is PmCIIl

DT Dt

.,

=A + Km V-T

(5.5)

Equation (5.5) states that if there is no heat generation (A = 0) and if thermal diffusion is ignored (K m = 0), then there is no change in temperature in the moving coordinate system we have chosen (e.g., DT/Dt = 0). In other words, in the absence of thermal diffusion and heating, a temperature anomaly will be swept by fluid flow through the porous medium at pC/PmC/II times th e Darcy flux of the fluid,]. If the observation grid moves at (pC/PmCIIl )] , any and all thermal anomalies will appear stationary. Since Pm == 2.7, C/II == 0.2, and o C == 1, the thermal anomalies will move at about 1/0.54 = 1.8 times the Darcy flux. For a thermal anomaly migrating at 1.8 times the Darcy flux, the water/rock ratio after passage of the anomaly will be 0.54 cm-' fluid/cm ' rock , or about 0.2 g fluid/g rock. If the flow porosity is 10 vol % of the sediment (or -0.1 % in the fractures of an igneous rock) and the true velocity of the fluid is thus 10 (or 1000) times the Darcy flux, the pore waters will move through the thermal anomaly at - 5 (or -500) times the rate at which the thermal anomaly migrates. The pore waters migrating from a hot area into cool country rock may

Thermal Consequences of Subsurface Fluid Flow

19S

deposit silica and other minerals. The deposition will be small because the water/rock ratio is small so long as the th ermal front is allowed to migrate. For example, taking into account both temperature and fluid density, the geologic solubility of quartz peaks at ab out 2500 ppm (Cathles, 198 3). At a water/rock mass ratio of 0.2 th e maximum silicification expected from a freely migrating thermal front is 0.5 wt % . More sophisticated computer simulations confirm this conclusion (Cathles, 1983). A silicification of 0.5 wr % is disappointingly small and we arrive at a first conclusion of importance to hydrothermal ore deposition: significant thermal deposition of silica (or any other mineral, because other minerals are generally a great deal less soluble than silica) can be ex pected only where thermal anomalies are immobilized. Immobilization can occur where the thermal front reaches the surface, where there is fluid boiling and the temperature is constrained to follow the liquid-vapor curve of water, or in areas wh ere cold fluids mix with hot ones. In such areas, significant amounts of silica and other minerals with prograde solubilities may be deposited as the result of temperature drop and other processes. But unless physical factors conspire to immobilize a thermal anomaly, only very minor amounts of material (silica or base metals) can be deposited as a result of temperature change by fluids moving through that anomaly. Note that we do not con sider contrasts in rock reactivity and skarn formation here. The approach outlined above is quite general. The th ermal front migrates at a rate equal to the product of the Darcy flux times the ratio of th e th ermal storage capacity of the fluid to that of the solid plus pore water combined. Chemical fronts migrate according to the same principle. For exa mple, warm water dissolves about 25 wt % NaCI. The salt dissolution front in a sandstone that has a density of 2.5 g/cm" , contains 20 wt % salt, and has 10 % porosity will move at about half the Darcy flux because the sandstone carries 0.525 g Na Cl/cm-' solid, whereas the fluid carries 0.25 g Na Cl/crn' fluid. Under isothermal conditions, an oxygen isotopic fro nt will migrate at about half the Darcy flux because water contains about 56 mol of oxygen per 1000 cm 3 of water, whereas rock contains about 100 mol of ox ygen per 1000 cm 3 of rock. Hydrogen isotopic anomalies will migrate at about 18 times the Darcy flux because water contains about 111 mol of hydrogen per 1000 cm 3 of water, whereas rock contains only about 6 mol of. hydrogen per 1000 cm-' of rock. In each case the movement of the chemical front could be followed by substituting an expression similar to (5.3) and (5.4) into a conservation equation similar to (5.1). In each case the motion of the front would adequately be described by the average velocity, but diffusion and dispersion would smear the front out with time. The smearing by diffusion and dispersion is characterized by a Peclet number; further smearing du e to finite reaction rates can be characterized by a Damkohler number. Examples and further discussion are available in the literature (Cathles, 1983; Lassey and Blatner, 1988; Blattner and Lassey, 1989; Banner and Hanson, 1990; Cathles and Shea, 1992).

196

Thermal Aspects of Ore Formation

In cases where the motion of an alteration front is tied to the motion of the thermal front, the magnitude of the initial alteration can be estimated. For example, knowing that the thermal front migrates with a water/rock ratio of 0.54 cm 3 fluid /crrr' rock , that the equilibrium water/rock isotopic fractionation changes by -10 0/ 00 between 350°C and 100°C, and that the ratio of oxygen concentrations in the water to that in the rock is 56/100 suggests that migration of a 350°C thermal front will enrich the rock downstream in 18 0 by about 3%0 (= 0 .54 x 0.56 x 10). The enrichment is actually greater than this because earlier enrichment enhances 180 in the fluid and leads to a greater increase in 18 0 in the rock as the anomaly migrates (Cathles, 1983). Nevertheless, simple therma l calculations determine the water/rock ra tio that is effective in many kinds of alteration and allow the magnitude of alteration to be estima ted .

Thermal Anomalies at the Surfa(e VERTICAL flOW A second very useful relationship can be deduced from equation (5.1). The steady-state temperature distribution near the surface can be determined, in the case where there are no thermal sources or sinks and the vertical outflow (J z > 0) of water is known, by so lving

a2 T az 2

sr

er

sc); 1 = Km az -{; az

=- -

(5.6)

For convenience we have defined a skin depth 0 = Km/pC]z. Substituting a trial solution T = e rz results in the algebraic equation, r 2 - (l/o)r = 0, which has the solutions r = 0 and r = 1/0. The solution to equation (5 .6) thus has the general form of T = A + Bez/ o. The boundary conditions T(z = 0) = 0 requires A = - B, and the condition tha t T(z = L) = To requ ires A = To/(1 - e- L/ o). Thus

T To

1 - ex p(z/o) ex p(- L/ o)

=1-

(5.7)

Plotting eq uation (5.7) shows tha t th e vertical out flow is small, th e th ermal gradient increases linearly with depth, and the hea t flow, ] , eq ua ls Km To/L. When the vertical fluid outflow is large and 0 small, the temperature increases rapidly with depth to To according to T =ToO - e Z/ O), and the heat flow, ], equals K m T% . Between a depth of -20 an d L the temperature is constant and equal to To. Remember in these equations that depth is positive upward, and negative as one proceeds into the subsurface. This problem of geothermal gradients was first solved by Bredehoeft and Papadopolos (1965) and plots and additional discussion can be found there.

Thermal Consequences of Subsurface Fluid Flow

J97

The concept of skin depth can be very useful. For exa m ple, trace-cl ement distributions suggest th at th e upper 20 m of the Amul et Rh yolite a t Noranda was originally a ba salt that has been silicified. If th e silicifica tio n reflects outflow and near-surfac e co ol ing , the n 0 ;: 20 m an d th e fluid ou tflow, taking Km = 6 x 10 - 3 cal /cm-s-PC, mu st have been about 94 g/cm 2-yr (= Km/pCo). For To ;: 350°C, heat flow over th e silicified area w as ToKm/o =(350) (6 x '10 - 3 )/(20 x 102 ) ;: 1000 I£cal/cm 2- s o r H FU, wh er e normal heat flow is about 1.5 HFU. To deposit 10 wt % silica over a 20-m depth range, o r about 520 g silica /crrr - , requires 2000 years of such discharge, assum ing the 350 °C so lutio ns precipitated 2500 ppm silica as th ey vented. In 2000 years at 1000 HFU, 6 x 10 17 cal (2.5 x 10 18 J) are vented per square kilometer. At 10 18 cal/km' (Norton and Cathles, 19 79) , th e heat required to silicify 1 km 2 of surface requires 0 .6 km' of ba saltic magma. There could easily be - 170 km 2 of silicification at Notanda (Uzma nn,

1993). An intrusive volume of 100 km 3 would be required to produce this possible silicification. The exposed area of the Flavrian tonn alite is - 150 km 2 , however, so even if the thickness of the Flavrian wer e only a few kilometers , there is plenty of heat a va ila ble from th e intrusive. This is not intended to resolve th e history of Noranda, but to illustrate the use of th e equations developed a bo ve. C lea rly a volume of silicified materi al requires a cert ain intrusive volume to supp ly enough heat. If hydrothermal silicifica tio n is extensive, a large intrusive heat so urce is requ ired. The size of th e intrusion can be estimated and could be of inter est in mineral exploration.

FLOW FROM A SHALLOW·DIPPING AQUIFER A similar approach can be used to a na lyze th e steady-state temperature distribution due to outflow from a shallow-dipping aq uifer system . If th e aquifer segments are linear and th e aquifer is thin compared to its depth, th e temperature profile in each aquifer segm ent is determined by th e temper ature a t a single point in that segment. If temperature at a single point in th e segme nt is known, the temperature everywhere in the segment ca n be calculated. Segments may thus be joined together or daisy-ch ain ed with th e temperature at the end of each segment controlling th e temperature profile in the next segment. In this way, temperatures along the entire flow path ca n be determined by the flow rate and th e temperature at just one point along th e path. Since inflow temperatures will be close to ambient, temperatures can easi ly be calculated along any su bsu rface flow path in a sed imenta ry basin. The daisy-chain m ethod sketched a bove is fully described in Cathles (1987) and is used th er e to show that a Darcy flux of at least 15 rn/yr in a 30-m thick aquifer is required to produce a 35°C temper atu re an omaly a t l -km depth when the inflow to th e aquifer at 5-km depth is 300°C above th e ambient surface temperature of l5°C a nd the dip o f th e aquifer is 1% (0.5° ). The vertical flow rate in th e aq u ifer is 0.15 m/yr. The th ermal co nd uc tivity of the basin in this ca se is 3.5 x 10- 3 (a low value), and th e heat flow a rela-

198

Thermal Aspects of Ore Formation

tive ly high 2 .1 HFU, so that the no rm al (unpert urbe d) tem per atu re gradient in the basi n is 60°C/km. Under these circ umsta nces tem per atures of 110°C a rc reach ed at l-km depth. These temperatures a nd depth s are appro priate for th e fo rma tion of Mi ssissippi Valley-type lead-zin c deposits. T he 315 °C temper ature at 5- km depth is very hot for a sedimentary basi n. For coo ler basins grea te r ex pulsion rates would be required. Faster exp ulsio n rat es co uld incr ease aquifer tem per atures at I -km depth to the inflow tempera t ure at th e bottom of th e bas in (31 5° C in th e a bove exa m ple). If th e flow is to pographicall y driven , however, so th at th e fluids th at disch a rge at one mar gin rech arge at ano the r, th e max imum temperatu re th at can be att ained at l-km depth o n th e disch arge margin is o nly abo ut hal f th e temper ature in th e deepest parts o f the basin. This is becau se the fluids ente ring th e basin cool it substa ntially before th ey disch arge. T he th ermal gra dient ab ove th e aquifer in th is simple model is just the aq uife r tem perature divided by its depth. Becau se th ermal gra dient ma ps are ava ilable fo r most o f th e United Sta tes, method s like daisy-ch a ining allow ther mal gra dients to be inve rted a nd aquifer fl ow rates to be in ferred. N earsurface temperature gradie nts ca n be related very directl y to fluid flow an d ca n be used under certa in circumsta nces to mea sure th e rates of fluid fl ow in aq uifers.

RATES

OF SUBSURFACE FLUID FLOW

Th e other ph ysical side of hydrothermal ore depositi on invo lves th e rates and mechani sms of fluid movement. Sub surface fluid flow can be driven by a vari ety of pr ocesses. Flow may be dri ven by differ enc es in w ater-table elevation. Becau se th e wa ter tabl e tends to follow topography, thi s kind of flow is often ca lled topography-driven flow. Flow can be driven by den sit y differences produced by temperatu re o r sa linity ano malies . This kind of fl ow is ca lled thermal o r haline co nvection. A th ird kind o f flow is produced by th e co mpac tio n that occ urs wh en sediments accum ulate in a basin o r wh ere basin sed iments are overthrust. The ex pulsio n is similar to sq ueez ing wa ter o ut of a sponge and is referred to here as co mpactive ex pulsio n . Fin ally, fl ow can be produced in areas wh ere metamorphic reactions gen er ate fluids fro m m inerals. Amphi bolite grade devolatizati on reacti ons a nd th e pr oduction of hydrocarbon fl uids from th e maturation (slow cooking ) of kerogen a re o bvio us exa mples, whi ch a re referr ed to here as metamo rphic a nd matu rati on exp ulsio n. The flo w rates pr oduced th rou gh th ese mech anisms can be estima ted by simp le calculations. T he flow of water th at ca n be produced by eac h mecha nis m is conside red in turn in th e followin g secti o ns. The most important basic eq ua tio n go verni ng fluid fl ow is Darc y's Law. It expresses co nse rvation o f momentum in a dissipative porous med ium where inertial fo rces are un importan t. Darcy's Law ca n be written

-Rates af Subsurface Fluid Flaw

k

] = - - (VP - pg ) 1/

199

(5 .8)

where J is the Darcy flux in cm / s as before, k is th e perm ea bility o f th e po ro us medium in cm 2 , 1/ is the viscosity of th e por e fluid in g/crn- s o r po ise, P is the pressure in dyn/cm 2 , p is the density of th e po re fluid in g/ crrr", and g is the accelerat io n of gravity in cm/ s2 . The positive z axis points verti ca lly upward, so g = - gi.

Topography-Driven Fluid Flow Groundwater hydrology covers topography-driven fluid flow, and many good basic references are available (De Wiest , 1965; Bear, 1972; Freeze and Cherry, 1979). Groundwater hydrologists modify Darcy's Law to consider only constant-density water. Setting p = PO = constant in equation (5.8), -pg =pogi =V(pogz), so that equation (5.8) becomes

k

] =-

V(P + pgz)

(5.9)

1/

Suppose, as shown in Figure 5.1 , that th e elevation of th e perforat ed part of a well casing (where pore waters are allowed in and out of the well and ther efore where aquifer fluid pressure is measured) is measured in centimeters above sea level (or any other conven ient elevation datum) and ca lled th e

h z Sea Level Schematic diagram shawing the elevatian, z, af perforated ports af a well relative to a reference datum such as sea level, the elevatian af the water in the well relative ta the abservatian point (perfaratians), hp ' and the elevatian, h, relative ta sea level (h = hp + z) . FIGURE 5.1

200

Thermal Aspects of Ore Formolion

elevation head, z. Suppose also that the static height of the water in the well is measured relative to the same datum and called the hydraulic head, h. The pressure at z equals the static standing column of water in the well and can be expressed as: P = (h - z)POg. Substituting this expression for P into equation (5.9), and noting that PO and g are both constant and thus may be moved outside the gradient operator, yields the hydrogist's form of Darcy's Law:

J = - pogk Vh = - KVh

(5.10)

1/

where K is the hydraulic conductivity. In equation (5.10) K has units of cm/s. It is common for K to be given a wide variety of different units. For example, in Figure 5.2 K has units of m/yr, From equation (5.10) hydraulic conductivity, K, and permeability, k, are related: Kjcm/s] = pogk/1/ "" lO sk[cm 2 ] . A more usual unit of measure for k is the darcy, where 1 darcy =10- 8 cm 2 . Thus Kjrn/s] "" 10 - sk [darcies]. Hydraulic conductivity can be measured in a particularly simple way that is the basis for the percolation test required for septic tanks. Hydraulic conductivity is the maximum flux (sprinkling rate) a porous medium can accommodate without ponding. As illustrated in Figure 5.1, the hydraulic head, h, in equation (5.10) is the sum of the pressure head, h p , and the elevation head, z, so that h = h p + Z. If there is no ponding and the water saturation of the soil is almost, but not quite, 100%, the pressure head is zero (atmospheric) everywhere in the soil. The head is therefore exactly the elevation head. Consequently, Vh = Vz = 1, and, from equation (5.10), J z = - K. Thus hydraulic conductivity is directly measured by the maximum rate at which water can seep into the ground without ponding. Figure 5.2 shows that the hydraulic conductivity of naturally occurring material varies from 10 - 6 to 10 7m/yr. Gravels could allow water to seep into the ground at a rate that would allow a column of water 10,000 km high to disappear in one year, while clays could prevent even 0.1 Jtm of water from seeping into the ground in one year. There is a 13 order-of-magnitude range in the permeabilities of natural materials! Fortunately, yearly precipitation places an upper bound on topographically driven flow rates. Typically, one-third of the rainfall in an area infiltrates. Rainfalls of 10 m/yr and more occur in some rain forests but generally rainfall is less than 1 m/yr. Flow in a shallow-dipping aquifer is increased by the infiltration divided by the sine of the dip angle. For a 50 dip the infiltration is increased by an order of magnitude; for 0.5 0 (or 1%) dip the increase is two orders of magnitude. All considered, an approximate upper bound on Darcy flux for topographically driven flow is -30 m/yr, but usually the Darcy flux in aquifers is -0.1-1 m/yr, giving (for 10% porosity) true fluid velocities between 1 and 10m/yr. The true velocity of groundwater through the J Aquifer in the Great Artesian Basin of Australia, for example, is -1-5 m/yr (Halbermehl, 1984).

Rates of Subsunace Fluid Flow

K (m/y r)

k (d a r cy)

10*7 10*6 10*5 10*4 10*3 10*2

k (cm* 2)

10*5

10*-3

10*4

10*-4

10*3

10*-5

10*2

10*-6

10*1

10*-7

I

10*-8

!I

10*-1

10*-9

10*-2

10*-10

~ "0 ", _c

cB

'"

10*-3

10*-11

I

~

10*-4

10*-12

10*-2

10*-5

10*-13

10*-3

10*-6

10*-14

10*-4

10*-7

10*-15

10*-5

10*-8

10*-16

10*1

1

o

~

10*-1

I

"OCJ "0 C

201

10*-6

Hydraulic Conductivity FIGURE 5.2

Permeability

Hydraulic conductivifies and permeabilities of common sediment and rack types (odopted hom Freeze and

Cherry, 1979).

Convective Flow Convec tive flow is dri ven by differences in fluid densit y. As discu ssed in Chapter 10 of Barnes (1979), a great deal can be learned by exploring the con ditions under which no flow wi ll occur. From equation (5.9), J ca n be zero if the permeability is zero (no t a very interesting or even realistic case) or if 0 = VP - pg

Writ ing o ut th e vecto r co mponents of th is eq uation ex plicitly,

......

202

Thermal Aspects of Ore Formation

ap ax ap ay ap az

(5.11)

shows that there will be no convective flow if fluid density is a function of z only and the vertical pres sure gradient at any depth equals pg, Flow will be produced if there are any horizontal variations in p, however. This is because, from the third component of equation (5.11),

If p were a function of horizontal position, P would vary horizontally and the horizontal gradients of P would be nonzero, in contradiction to th e requirements that the first two components of equation (5.11) be zero. Thus a necessary condition for no flow in the subsurface is that density is a function of depth only; th at is, p = po(z) . Temp erature will typically increase linearly with depth and can be characterized by a geothermal gradient, GT[OCjcm] . With the coordinate system chosen such that the surface is at z = 0 and z is positive upward, G T might typically have a value of - 25° Cj km or - 25 x lO- SOCjcm. If the coefficient of thermal expansion of water is ex and the density of water at O°C is oi» , then the temperature, fluid density, and pressure under no-flow conditions (indicated by the subscript 0) can be expressed as To(z) =T s - GTZ

(5.12a)

po(z) = POO - POOexGTz

(5.12b)

PO(z) = -POOgz +

POOexGTz 2 2

(5.12 c)

where T s is the ambient average temperature at the surface. Note that, since G T is negative, and z is negative downward, equation (5.12c) indicates that th e pressure increases with depth at a decreasing rate. This is because increasing temperature causes increased thermal expansion and decreased fluid density. Po(z) is the cold water hydrostatic gradient. Fluid density may develop horizontal variations and cause convective fluid circulation in the sub surface in two ways . If the thermal gradient is too large or the subsurface is permeable over a substantial depth interval, a critical Ral eigh number (defined below) may be exceeded, in which case free-con-

Rotes of Subsurfoce Fluid Flow

203

vection cells will develop. This convecti on will produce and susta in horizon tal temperature gradients appropriate to drive th e convection. Th e free fluid circulation is more vigorous the more supercritical th e Ralei gh number and will continue until mineral precipitation or other processes reduce th e permeability and the Raleigh number falls below its critical value. Th ere is no free convection if the Raleigh number is less than its critical valu e. Convective pore-water circulation can also occur if externa l factors set up horizontal variations in fluid density. Th is is called for ced convection and is the most important kind of convection for hyd rothermal o re dep osits. Horizontal temperature gradients and forc ed convection are prod uced along th e near-vertical boundaries of an intrusive body and by horizon tal var iati ons in rock thermal conductivity. Convection is forc ed in the sense th at it will occur at the rate allowed by the permeability regardless of wh eth er a cr itical Raleigh number has been exceeded or not.

FREE CONVECTION The Raleigh number, R, is given by th e following ex pressio ns:

(5. 13)

where H is the thickness of the lithologic layer and the other parameters are as defined previously. The critical Raleigh number depends on whether free flow is allowed out the top surface of the layer. If free flow is allowed, Rc = 27 and if free flow is not allowed Rc = 411" 2 (Lapwood, 1948). For GT = 25°C/km and H = 5 km, the free-flow critical Rale igh number is exceeded if k > 0.3 millidarcy (0 .3 x 10- 11 cm 2 ). This is a reasonable permeability for the subsurface, as indicated by measurements in the Ga llapagos area (Fehn et aI., 1983), and therefore free convection could be an important geological process. It could cause significant subsurface redi stribution of silica. Salinity stratification (Frape and Fritz, 1987) probably eliminates free convection in most old terrains, however. The maximum vertical flow rate for th e no-outflow boundary condition is proportional to the square root of the difference between the Raleigh number and the critical Raleigh number (Donaldson, 1962; Combarnous and Bories, 1975):

i, = h K pcH

(R - Rd o.5

For K = 4 x 10 - 3 cal/crn-s-SC, H = 5 km , o C = 1, R = 50, and 1z = 0.012 m/yr. For R = 100,1 z = 0.028 m/yr.

(5.14)

Rc

= 411" 2,

-204

Thermal Aspects of Ore Formation

FORCEDCONVECTION NEAR AN INTRUSION OR SALT DOME Forced convection is of particular interest to hydrothermal ore deposits because it can be much more vigorous than free convection and therefore can sustain high temperatures and steep temperature gradients near the surface. Estimates of the magnitude of forced convective flow are made most easily if flow is considered to result from a perturbation pressure, PI, an anomalous fluid density, PI, and an anomalous temperature, TI, defined: P(x, t) = Po(z) + PI (x, r)

(5. 15a)

p(x, t) = po(z) + P I (x, t)

(5. 15b)

=To(z) + TI (x , r)

(5.15e)

TI (x, t)

Since from equation (5.11) no fluid flow requires VPO - POg = 0, substitution of equation (5.15b) into (5.8) yields a new form of Darcy's Law: (5.16)

It can be seen explicitly from equation (5.16) that flow in the subsurface is driven by pressure and fluid density perturbations relative to the normal or zero-order (no flow) values of these variables that are given in equations (5.12a,b,c). Convective inflow near an intrusive body will not perturb subsurface fluid pressures very much, because the inflow will be over a large area and slow. If there is no barrier between areas of cold water inflow and hot fluid upflow near an intrusive, cold water hydrostatic pressures will impose themselves on the upflow zone (Elder, 1966). They will squeeze and narrow it until the rate of upflow compensates for the lower density of the hot fluids and the pressure gradient in the upwelling areas matches the cold water hydrostatic gradient. Mathematically, this means that in equation (5.16) the upflow zone \l PI = O. The rate of upflow is thus (5.17)

Since g = - gi, and PI = - POOexT I , the rate of thermally forced convective flow at the margin of an intrusive is

J=

kpoOgexTI

i

(5.18)

TJ 11

2

For ex = 10 3 JOC, g = 10 3 cm/s 2 , k = 1 millidarcy = 10 - cm , TJ =0.002 g/cms, and TI =300°C, J =0.47 m/yr. The upflow rate near an intrusion margin

Rates of Subsurface Fluid

F10Vl

205

is thu s about

J = 0.4 7k[md] ml yr If the permeability w er e 10 millidarcies (rnd ), for exa mple, th e vert ica l Darcy flux would be 4.7 m/yr; if T] were 150°C rather th an 300°C th e flow rate would be halved. Rates of haline convection ca n be estima ted in th e sa me way. The perturbation in densi ty is du e to sa linity rather th an temperature. If sa linity, C, is measured as the mass fracti on salt, PI '" pooC, and cases of haline co nvection near the margin of a salt dome can immediately be esti ma ted from equation (5.17). For the same parameters as before and I -mi llida rcy permeability, the downflow near a sa lt dome th at ca uses adjace nt fluids to be 22 wt % more saline than a d jace nt regional fluids would be -0.34 m/yr, The flow rate is a little less than th e th erma l upflow rate in th e pr evious exa mple because the perturbation of density cau sed by increas ing temperature by 300°C is 0.3 g/crrr", wher eas the increase in density due to salt add ition is -0.22 g/cm-'.

FORCED CONVECTION IN SEDIMENTARY STRATA Another example of forced co nvectio n th at is indirectl y of co nside ra ble economic interest is the forced convection that occurs wh en lithol ogic st ra ta of differing thermal co nd uct ivity are folded. If sedi ment ary strata with differing thermal conductivities are deformed such th at th eir co ntact is no lon ger horizontal, horizontal temperature gr adients are esta blished and th ese gra d ients driv e forced convection. The magnitude of th e forced convect ion ca n be est ima ted in a fashio n identical to that used in the preceding section. Co nside r a sa nd layer enclosed by shale (Figure 5.3). If a sa nd layer, o f thickness ~h and th ermal co nd uctivity Ks , is dipping at a sma ll a ngle w ith slope S, th e region al heat flow, h -I , is uniform , and the thermal conductivity of th e sh ale is K, th e difference in temperature horizontally across a sand layer, ~ T, as illustrated in Figure 5. 3 is

K

Shale

Shale FIGURE 5.3

Temperatures at the some elevation on either side of a shale-hosted sand layer of thickness tJ.h.

206

Thermal Aspects of Ore Formation

t:..T = (D ) H + t:..h } H ) _ (D ) H + t:..h } H ) KKK k, = t:..h } H (K s _ 1)

Ks

(5.19)

K

If th e th erm al co nd uc tivity o f th e sa nd is greate r than th e shale that encloses it, th e temper ature a t th e top of th e sa nd lay er will be t:.. T/2 greater th an at th e sa me dep th in the middle of th e lay er, a nd th e temper ature at the bottom o f th e lay er will be - t:.. T/2 less than th at in th e middle of th e layer a t th e sa me depth . W ater at th e top o f th e lay er will circ ula te upward , fluid at th e base downward , a nd fluid in th e center o f th e lay er will be static. Th e zero -o rde r (st agnant) hydrost at ca n th er efore be conveniently defined by the temperatures and fluid den sities in th e middle of th e lay er. The Darc y flux al ong th e top of th e sa nd layer,lII " then follows immediately from eq ua tio n (5. 18):

k ( POOCi } 11 + = ~

t:..T) 2"

(goS)

(5.20)

wh er e t:.. T is given by eq ua tion (5.19) . N ote th at only th e component of gO pa ra llel to the sid es of th e sa nd lay er is effect ive in driving flow and th er efore gO has been modified to goS, wher e S is the slo pe of th e la yer. If th e sa nd laye r is sinusoida lly folded w ith a m plit ude a and w a velength A, an d S is ta ken to be th e ma xi m um slo pe of th e lay er at x = A/2 so th at S = 27ra/A , equa t ion (5.20 ) yields

(5.21)

This is identical to th e exact solution for a simple folded layer surro unded by shale (Davis er a l., 1985). The vertical fluid velocity is 111' S = 111' (21ra/A). T aking} H = 1.2 x 10- 6 calfcm 2-s, t:..h = 100 m, K = 3 x 10- 3 cal/cm-s-Xl, K, = 5 x 10- 3 cal/cm-s-r C, A = 10 km, a = 100 m, k = 10 darcies = 10- 7cm 2, and th e other paramet ers are as in the previous calculations, } II' = 0.8 m/ yr a nd } z = 0.05 m/ yr, values close to those estimate d by Wood and H ew ett (198 2) . Thus substa nt ia l forced circ ula tio n can occur in thick, permeable, nonhorizonral, sha le-hosted sa nd bodies just as a consequence of contrasting th erm al co nd uc tivity. Becau se co nve ctio n ca n persist for very lon g periods o f time, substa nt ial red istr ibution of silica may result (Wood a nd H ewett, 1982 ). For exa m ple, foll owing Wood a nd H ewett a nd taking a typ ical tem perature dependen ce of silica sol ubi lity of 5 ppm Si02;oC, a vertical upflow o f 5 cm/yr will dep osit 2.3x 10- 9 cm 3 Si02/yr[= (5x lO-6 g Si02/g H2 0 )(PH20/PSi02 )G TVz], a nd 10

Rates of Subsurface Fluid Flow

207

vol % silica can be deposited in 43 million years [= (0.1 cm 3 Si02/cm3 )/(2.5 x 10- 9 cm 3 Si02/cm3-hr)].

FORCED CONVECTION IN A RADIOGENIC INTRUSION Finally, radiogenic heat can steadily maintain horizontal temperature gradients and drive forced convection. Fehn et al. (1983) showed that if the permeability was very low so that there was no convection, an intrusion having the size and radioactivity of the Conway, New Hampshire Granite would steadily maintain its center at 10-km depth at temperatures -140°C hotter than those at the same depth far from the intrusion. If the granite and its surroundings had O.l-millidarcy permeability, the Darcy flux through the intrusion would be a few centimeters per year, and a mass of water equal to that of the Conway Granite could be circulated through it in less than 10 million years. Fehn et al. suggested that such radiogenic thermal circulation could account for vein uranium deposits associated with radiogenic intrusions. Whenever a radiogenic granite is fractured, circulation may remobilize uranium into the fracture system. Long-lived radiogenic convection may also have produced the china clays in Cornwall (U. Fehn, personal communication, 1983).

Expulsive Flow Fluids can be expelled from a basin by compaction or as the result of reactions that convert solids to lower density fluids. The expulsion rate can be estimated by equations expressing conservation of solid and fluid mass. Conservation of solid mass for uniform sediments requires that the same solid mass pass all depth horizons as the basin subsides. If vs(z) indicates the subsidence rate (relative to the basin surface), this means that vs(z)[l - ¢(z)] = constant Application to both the surface and depth z shows that the subsidence at depth z is a function of the surface sedimentation rate S, the deposition porosity ¢o, and the porosity at z, ¢(z):

1 - cPo vs(z) = S - - 1 - ¢(z)

(5.22)

COMPACTIVE EXPULSION Conservation of pore-fluid mass requires that the compactive efflux of pore fluids at a depth z, ] c(z), equal the difference between the water carried in at z and out at the base of the basin, z = b:

208

Thermol Aspects of Ore Formotion

If the basement is considered impermeable and there is no horizontal flow (as is assumed here) , the expulsive Darcy flux will be vertically upward. The direction of flow (sign of Jc) is provided by geological considerations. From equation (5.22),

Jdz) = (¢(Z)

- ¢(b)

~ =:~~~)

vz

(5.23)

A "typical" compaction curve (Hanor, 1979) is ¢(z) = 0.7 exp(0.57z[km])

(5.24)

where z is measured in kilometers and is negative downward into the basin from z = 0 at the basin surface. Taking b = - 8 krn and z = - 3, we find ¢(z = 0) = 0.70 = ¢o ¢(z = - 3 krn) = 0.13 ¢(z

= - 8 krn) = 0.007

Jc(z =- 3 krn) =O.12vz vz(z = - 3 krn)

=0.345

Jc(z =- 3 krn) =0.045 For 5

=0.1

(5.25)

to 1 krn/Ma, ] c(z =- 3 krn)

=4 x

10 - 5 t04 x 10 - 6 rn/yr

Cornpacrive fluid expulsion rates from sedimentary basins are about four orders of magnitude smaller than the topographic and convective Darcy flux rates we have previously been considering. AOUATHERMAL EXPULSION

Pore fluids are also expelled from a basin as the pore fluids are buried, heated, and thermally expand. For uniform burial rates the rate of heating is approximately constant and can be characterized by Hr. For example, in a basin accumulating sediments at 1 krn/Ma, H, would be about 25 °C/Ma if the geothermal gradient in the basin, GT, were -25°C/km. If the coefficient of thermal expansion of the pore fluid is ex, the increase in pore fluid volume at any depth is ¢(z)exH r, and the total rate of expansion between z and b, and therefore the total aqua thermal expulsion, ] at, is

] adz) =

J~ ¢(z)exHr dz

Rates of Subsurface Fluid Flow

209

If we approximate H , = vz(z = - 3 km )GT, integration together with eq uation (5.24) yields

Jadz = - 3

km) =

¢(z = - 3 km) - ¢(b)

0.57

exGTVz(Z =- 3 km)

For ex = 10 - 3 ;oC and a thermal grad ient GT = 25°C jkm, and b before, eq ua tio n (5.26 ) gives

(5.26)

=8

km as

Jad z = - 3 km ) = 0.005vz Aquathermal expulsion is about 4 % of compacti ve expulsion.

MATURATION EXPULSION The vertical flux of fluids from chemical reactions in a sedimentary basin is easily estimated if those reactions begin below z and go to completion above the basement, z = b. If th e grade of the reacting material is Cj grams of i per gram of solids, Ps is the density of the solids, and ~ Vi is the change in volume upon complete reactio n of i, th e upper bound on th e fluid flux, Ji, due to reaction i is

l:

< ~ V CjPsVz

(5.2 7)

The efflux will be less than indicated by equation (5.2 7) because the reaction goes to completion below z, where the subsidence rate is less than V z. Kerogen matures progressively to gas in the deep portions of sedimentary basins (see Chapter 12). The deeper methane-generating rea ctions that gen erally occur below a 3-km depth are dom inated by the conversion of earlier generated oil to methane. For typ e II (marine ) kero gen about 0.4 g of oil with density 0.8 g/cm-' are generated per gram of o rigin al kerogen , and th e oil decomposes to 0.24 g CH4 with dens ity - 0.16 g/ crn-' and 0.16 g of residu a with density 2 gjcm 3 (J. Hunt, personal communication , 19 89 ). The density of methane giv en here was calculated for 87 °C and 290 bars usin g th e Behar (Behar et aI., 1985) equation of state. These pressures and temperatures are typical of conditions at a 3-km depth in hydrostatically pres sured sed iment s above the overpressured zones that typically lie below this depth in actively infilling basins. The volume change of the oil decomposition reaction is 1.1 cm 3 j g kerogen. Thus if CKPs = 0.1, from equatio n (5.25)

JCH4 = 0 .1 1v z

(5.28)

Comparing eq ua tio n (5 .2 8) with (5.25) shows that th e efflux o f pore fluids due to th e maturation of reasonably organic-rich sediments in a basin is comparable to the effl ux that could be caused by compaction! The burial of organic-rich sediments into the gas generation window can be as effective in generating overpressures and ex pelling pore fluids as compaction, a

Thermal Aspects of Ore Formation

21 0

possibility of considerable interest to the formation of Mississippi Valleytype (MVT) deposits. It was suggested as a fluid expulsion mechanism for the classic midcontinent of the United States by Rich (1927) and has interesting implications for coupling fluid expulsion and the formation of MVT lead-zinc deposits to sea level regressions (Eisenlohr et aI., 1994). METAMORPHIC EXPULSION

Amphibolite grade metamorphic reactions liberate water bound in hydrous minerals. Likely rates of metamorphic fluid expulsion can be estimated in a fashion similar to that discussed above if Vz is considered the migration rate of a thermal front away from an intrusive body rather than the subsidence rate at z. This is because, in equation (5.24), Vz is really an estimate of the rate at which the thermal gradient moves through sediments in the basin. For example, far enough up in the sedimentary sequence, temperatures are cool enough that the reaction in question has not occurred. Deep enough into the basin , temperatures are hot enough that the reaction is fully completed. Because the geothermal gradient in a basin is not affected greatly by the sedimentation rates considered here, the subsidence velocity, Vz, is the rate at which sediments move across the maturation boundary. We need a way to estimate how rapidly the intrusive will heat its surroundings. A conductive thermal front migrates a distance z = 2~ away from the hot surface of an intrusion in time t (Carslaw and Jaeger, 1959). Taking the derivative of this expression, we see that the rate of migration of the temperature isotherms is proportional to the thermal diffusivity of the rock and inversely proportional to time

(5.29)

For

K

1

= 0.01 cm-/s,

vz[km/Ma] =

5.6 ,,It[Ma]

=5.6-177 km/Ma

for

t

= 10 6 to 10 3 yr

This calculation presumes that metamorphism is driven by igneous intrusion or delamination of the lithosphere. Evidence of fluid flow and highT, low-P metamorphism during deformation suggests this geologic concept (Etheridge er aI., 1987; Loosveld and Etheridge, 1990; Sandiford and Powell, 1991). Metamorphism by overthrusting produces slower rates of fluid expulsion, as will be briefly reviewed below. It can be seen from equation (5.29) that the migration rate of a thermal front, although very rapid immediately following intrusion, drops to 600°C are as permeable as their host is clearly wrong. There is certainly a dependence of permeability on temperature. At some temperature rocks will become plastic enough that the fractures through which flow takes place will heal and the rocks will become impermeable. Observations of typical venting temperatures in forced thermal convection systems and computer simulations that incorporate the thermodynamic properties of water indicate that the permeability cutoff occurs at T'" 350°C (Cathles, 1983). All convective systems seem to have maximum fluid temperatures of -350°C. Black smoker venting on the seafloor is at this temperature, as are the deep, preboiling fluids in continental geothermal systems. Maximum temperatures in volcanogenic massive sulfide deposits and hydrostatically pressured vein deposits are inferred from mineralogical, isotopic, and fluid inclusion data to be in this range. The frequency with which fluids of 300-350°C are encountered in active and fossil hydrothermal systems is a well documented (Polster and Barnes, 1993) observation that must be explained. It has been suggested that the fluid temperature is related to the triple point of water. Bischoff and Rosenbauer (1985), for example, argued that convection downward toward a hot surface will turn upward when low densities near the critical point are reached. This scenario is possible for a carefully contrived case of uniform heating from an absolutely flat hot pad (free convection), but hydrodynamic calculations show that the Bischoff-Rosenbauer mechanism wiII not work in the more common case of forced convection, where, for example, the exposed margin of an intrusive heats adjacent pore waters and causes them to convect upward. Calculations show that in this situation the thermodynamic properties of water make the upward transport of heat more rather than less efficient and cause the fluids that exit the surface to

216

Thermal Aspects of Ore Formation

reflect more rather than less closely the temperature at which they are able to interact with the rock (Cathles, 1983). In this case, the thermodynamic properties of water are completely ineffective in controlling water temperature. The only control on the preboiling discharge temperature during forced thermal convection appears therefore to be the temperature at which most of the circulating water flows through the rock. If hydrothermal fluids are to be ::;350°C, permeability must be a function of temperature such that it is much smaller above 300-350°C. This dependence of permeability on temperature does not mean that water cannot interact with rock at higher temperatures. Some interaction can and does take place as discussed in the next section. The dependence of permeability on temperature simply means that most of the pore water circulates through rock at T:5: 350°C. The permeability reduction appears to be related to pressure-driven rates of dissolution, where opposite sides of a fracture make contact. When temperatures of about 350°C are reached, the rate of dissolution at these locations is rapid enough to heal the fracture between rupture events. All other factors being equal, tectonically active areas may typically have slightly higher hydrothermal system temperatures than tectonically inactive sites. The difference in flow temperature is not expected to be large, however, because the rate of the dissolution reactions doubles for every -10°C increase in temperature. Therefore a site 10 times more tectonically active would be expected to operate at only -35°C higher temperature. The idea that there is a strong temperature effect on permeability that is effective above -350°C has many implications. For example, this hypothesis suggests that hydrothermal systems that operate at significantly different temperatures, such as hotter porphyry copper and cooler MVT lead-zinc systems, involve a nonconvective flow mechanism. It suggests that thermodynamic parameters at T::;350°C will suffice to characterize much of the alteration. It suggests that 350°C venting will operate as long as partially molten crystalline mush remains for a thermal cracking front to attack, and will cease as soon as this is not the case. Temperatures in a hydrothermal system will drop very rapidly as soon as there are no more magmas to buffer the heat supply (see next section). Black smoker venting thus requires the presence of magma. The 350°C cutoff suggests hydrostatic conditions may persist in the Earth's crust to about the 350°C isotherm, and that therefore topographyor convection-driven flow may occur to depths of -15 km but not at greater depths were fractures will be closed and impermeable. Because the rocks are so impermeable below 15 krn depth and above 350°C, pore fluids in these environments will tend to be overpressured. This basic permeability and pressure division may explain Archean lode gold deposits. These deposits are thought to form by fluid decompression at the amphibolite-greenschist boundary when fluids decompress from lithosratic to hydrostatic pressures. A temperature-dependent (350°C) pressure and permeability division is also compatible with the general assumption in metamorphic petrology that meta-

Fluid Fluxes, Temperature, and Ore Deposition

217

morphic reactions at amphibolite or higher grad e occur at lithostatic fluid pressures. MIDOCEAN RIDGE HYDROTH ERM AL SYS TE MS Four observe d characteristics of midocean ridge hydrothermal systems allow four principal unknowns of the hydrothermal system to be determined by the algebra ic so lutio n of four equations . The four observations are as follows: 1. Black smo ker fields commonly discha rge at a total rate Qw == 75 kgjs (Ca nn a nd Strens, 19 89 ). 2. The discharge temp erat ur e, T yc nt , is - 350°C. 3. The resid ence time, tfz, of the fluids in the flow zone, from when they are heated to 150°C and a nhydri te precipitates to whe n they vent, is "'''' ~iz

,

I

0- M'ne .hall

• - Sam pl. locality

o I

1000

2000

I

I

seere In 'M'

FIGURE 6.7 Map of the Tonopah Mining District, Nevada, showing the 0180 contours at 0, - 2, - 4, - 5, and - 6

(solid lines) for wholfHock samples (black dots) of the altered volcanic rocks in the mine workings (adapted from a figure in Taylor, 19740; based an dolo of Taylor, 1973). Nate the rough correspondence between the central zone of lowest 0180 values and the apex of the elevation contours (after Nolan, 1935) of both the upper su~ace (dashed line = 5925 h) of the productive ore zane in the vein systems at Tonopah and the elevation of the apex of the bottom su~ace of this productive ore zane (dotted line =5525 h). Vorious mine shafts ore indicated by open squares,

3.1 % NaCI equivalent) that have unusually high 00 = - 69; this might be explained by mixing with magmatic waters in the deeper zones of the Comstock Lode, compatible with the depth and relatively high calculated temperature and salinity (Figure 6.6 ). Criss and Champion (1991) followed up the study of Taylor (1973), carrying out an extensive 180 /16 0 study of more than 250 samples of volcanic and intrusive rocks from outcrops, mine workings, and drill cores from the Comstock Lode mining district (Figure 6.8). They observed a bimodal distribution of whole-rock 0 18 0 values: a generally distal group with 0 18 0 '" +5 to +9 made up of fresh to weakly altered samples, and a central group with 0 180'" - 4 to +3 made up of samples that have been intensely mineralogically

258

Oxygen and Hydrogen Isotope Relafionships in Hydrothermal Mineral Deposits

39"20'

•

o'---'1km 39·15'

'---------.a...,-:-::':~-.....:....~!L!.;!...-------.L__:__------J

119"35'

a (hatched), +3, and +6 for intermediate composifion volcanic rocks of the Comstock lode Mining DislTict (adapted from a map in Criss ond Champion, 1991). A75-1:.101 zone of low.180 rocks is centered on a small, 2·km1 stock at MI. Davidson (diagonal ruled pottern) that represents on ancient volcanic center. The famous Comstock lode ore deposits occur neor and along the prominent, NNE'lTending Comstock and Silver City faults (heovy block lines).

FIGURE 6.8 Sample leteficns of wholeiock «5 180 values (dots) showing «5 180 contours at

a ltered. The gradients in whole-rock 0 18 a are very syst ematic on a region al sca le, but the isotopic variations at a given outcrop are sm all , indicating pervasive water-rock interaction guided by ubiquitous fractu re-controll ed permeability (on all scales from microfractures to major faults). A contour map of wh ole-rock 0 18 0 values of surface sa m ples reveals a 75-km 2 zon e of 18 0 depletion, centered on the 2-km 2 granodiorite sto ck at Mr. David son and suggesting that the hydrothermal syste m was principally driven by this intrusion (Figure 6.8). The fine structure of th is 18 a a nomaly is dominated by two elongate, NNE-trending zones with negative 0 18 0 values st raddling the mineralized Comstock-Silver City-Occidental fault system. Movements on these faults occurred repeatedly, both during and after th e main epi sode of hydrothermal activity, and it is clear th at these large fracture syste ms were the major conduits for the hydrothermal fluid s that pervas ively alte red th e country rocks (Figure 6.9). Criss and Champion (1991) sugges t th at th e size discrepancy between the altered zone a nd that of the centra l intrusion could be explained if this intrusive center was a major stra to volcano, a feature of appropriate areal extent in which successive pulses of magma utili ze th e sa me central vent, thereby providing th e required excess

Epithermal Ore Deposits in Volcanic Terrones

259

West 6800

Comb ination Shaft

6400

g 6000

,.

c 5600 ""',"""··0 ..

.9 r.i

>

-......

5200

III

Ul

Sutro Tun nel

4800

14000

4400

-2,.

4000

East

3600

6800 6400 6000 tTl

ro

-_~

x PYROPHYLLITE • SERICITE o CLAY(HYPOGENE) o CLAY(SUPERGENEl () CLAY - SERICITE • BIOTITE ... CHLORITE- SERICITE -lS0 '-----L..>II~--'-----'-----L.L...---.1..------'------'-----'

-15

-5

0

5

10

15

20

8'so (%0) 18 FIGURE 6.10 Plot of &0 versus & 0 for most of the presently available analyses of OH·bearing minerals from porphyry copperole deposits and from a few other types of deposits, such as Tonopah, Nevada; Creede, Colorado; Climax, Colorado; Pasta Bueno, Peru; and Butte, Montana (Sheppard et ol, 1969; 1971; Taylor, 1973; Hall et ol, 1974; Landis and Rye, 1974; Sheppard and Taylor, 1974; Sheppard and Gustafson, 1976; Bethke and Rye, 1979; and Bowman et 01., 1987). The kaolinite line of Savin and Epstein (1970a) is shown far reference. The dotted line is drawn merely to emphasize the separation of supergene and hypogene clays. Abbreviations: CP =Cerro de Pasco; CC =Capper Creek; B = Bethlehem; AI = Morenci; ES = EI Salvador; MP = Mineral Pork; LR = Lost River, Alaska; BC = Bond Creek, Alaska; G = Gilman; SG = SI. George; NB = New Boston; and at Butte AA = Advanced Argillic alteration; PZ= Peripheral Zone; IZ = Intermediate Zone; CZ = Central Zone; EOM = Early Oark Micaceous alteration. The stippled pattern indicates the range of isotopic values inbiotiles from the patassic alteration zones of most porphyry copper deposits, EI Salvador biotites also lie within the stippled zone, but for clarity are not plotted (four samples, &180 =+4.0 to +5.0, &0 =- 73 to - 83; Sheppard and Gustafson, 1976). Final~, note that although they are not plotted on the diagram, essentially all of the biotite data paints from Bingham by Bowman et ol (1987) confirm the plotted results af Sheppard et ol (1971 ) and would a~o lie within the stippled fieldonthis diagram (nine samples, &18 0 = +4.0 to +4.7, &0 = - 76 to - 86). Two of the Bowmanet ol (1987) Bingham samples have just slightly lower &0 values of - 90 and - 92, perhaps as a result of overprinting by the prapylitic alteration event at Bingham; these prapylitic biotites coexist with chlorite and have &18 0 =+3.0 to +4.8 and &0 = - 84 to - 97 (e~ht samples).

present at most of these deposits. In addition, th e very low 5 180 valu es in sericites and clays from some of th e northerly or high-elevation deposits (Butte, Climax) independently argue tha t these effects are not due to secondary exchange, as also does the similarity in 5D va lues of sericites of vastly different grain size from the Climax molybdenum deposit (Sheppard er a I., 1971 ). Hall et al. (1974) also demonstrated very nicely that th e sericites at

Porphyry Copper cnd Molybdenum Deposits

265

Climax preserved their oD values even though nearby potassium feldspars were drastically depleted in 18a by exchange with heated groundwaters. In striking contrast to the sericites and hypogene clays, the oD and 0 180 values of the hydrothermal biotites from the potassic alteration zones of porphyry copper deposits form a very tight grouping similar to "normal" igneous biotites and show no relationship to latitude (Figure 6.10). Biotites from Panguna, Bougainville (Ford and Green, 1977), and from El Salvador, Chile (Sheppard and Gustafson, 1976), also display this characteristic isotopic composition, as do essentially all of the samples analyzed from Bingham, Utah, by Bowman et al. (1987). The only biotites from porphyry Cu deposits known to fall outside this narrow range are some that may be inferred to have been overprinted by later-stage fluids, such as the finegrained EDM biotites at Butte, Montana, those at Copper Canyon, Nevada (oD = - 100 to -139; Batchelder, 1977), and those in the propylitic zone at Bingham (Bowman et al., 1987). The EDM biotites almost certainly exchanged with 10w-oD, main-stage ore fluids at Butte, because they are from a zone where there was intense overprinting by this main-stage meteorichydrothermal mineralization (Sheppard and Taylor, 1974). Bowman et al. (1987) followed up the studies of Sheppard et al. (1971) at Bingham, Utah, and showed that Na Cl-rich fluids in equilibrium with the potassic core of the Bingham stock had a narrow range of calculated 0 180 (+5.1 to +6.9) and oD (-51 to - 70 ), just like normal magmatic waters. However, fluids in the outer transitional and propylitic zones are progressively enriched in deuterium and depleted in 180 (0 180 as low as +2.8 and oD as high as - 40), leading Bowman et al. (1987) to conclude that the external, non magmatic hydrothermal system at Bingham was dominated by relatively D-rich formation waters, and that these were inward-directed toward the Bingham stock during sericitic and argillic alteration. Low-D, meteoric groundwaters were only important during late-stage, fracture-controlled deposition of clay minerals and sericite, and during alteration of the nearby Last Chance stock (where the biotites have 0 180 = +2.2 to +4.1 and oD = -103 to - 112). Recently, Dilles et al. (1992) have described a virtually complete 6-km vertical cross section through a tilted and deeply eroded porphyry copper ore body, the Ann-Mason deposit near Yerington, Nevada. They provided details of the fluid-flow geometry and evidence for multiple types of fluids with different isotopic compositions, including (1) a high-temperature, high-salinity magmatic water that was reponsible for potassic alteration and Cu mineralization and two types of "external" water; (2) a synchronous, deep, convecting, saline, sodic-calcic fluid; and (3) a late-stage, lower-temperature, more dilute fluid prevalent at shallow levels, which caused the sericite alteration. The latter has the isotopic characteristics of seawater or low-latitude meteoric H20. Such waters are compatible with the known paleogeography at the time that this Jurassic porphyry deposit was emplaced (Solomon and Taylor, 1991). However, the isotopic contrasts between these various kinds of aqueous fluids in the Yerington area are not very large, making it difficult to sep-

266

Oxygen and Hydrogen Isotope Relationships in Hydrothermal Mineral Deposits

aratc out the different effects of temperature and water/rock ratio from the sources of the H20. The importance of the study of Dilles et al. (1992) lies in demonstrating that the external fluids responsible for sodic-ealcic alteration circulated very deeply (to 3- to 6-km paleodepth) laterally alongside and below an ore zone that was simultaneously being formed from magmatic waters. Summing up, the hydrothermal biotires in all porphyry copper deposits that have so far been studied appear to have formed from H20 with an isotopic composition that coincides with the hypothetical field of primary magmatic waters, whereas the clays and sericites generally appear to have formed from heated waters of external derivation (Figure 6.10). In most cases this external H20 has a meteoric source; however, the amounts of meteoric H20 are much less than in the epithermal deposits associated with the higher permeability volcanic-intrusive terranes discussed above. This implies that significant hydrothermal convective systems were also established around the epizonal porphyry copper stocks, but that the potassium-silicate alteration assemblages in the cores of the stocks are in general not related to these external systems. A complicating factor, however, is that the coexisting quartz and potassium feldspar are consistently out of isotopic equilibrium as a result of postdepositional ISO exchange between the feldspars and the various aqueous solutions that affected these deposits. Sheppard et al. (1969, 1971) and Taylor (1974a) developed the following model to explain the isotopic relationships in porphyry copper deposits. Magmatic-hydrothermal solutions under lithostatic pressure are formed during the late stages of crystallization in the upper, interior portions of a porphyry stock. Outside the stock, a convective-hydrothermal circulation system is established under hydrostatic pressures that are probably only one-third as high as the pressures inside the stock. These external waters might be either meteoric groundwaters or saline formation waters, or both. The internal and external hydrothermal systems are simultaneously present early in the history of the stock, although the expanding envelope of magmatic H20 and the higher lithostatic H20 pressures of the internal system prevent the external hydrostatic system from invading the stock (Taylor, 1987). However, the lower-temperature external system will persist even after the magmatichydrothermal system fades away. Thus, with time, as the "heat engine" of the stock cools off, the external hydrothermal system collapses inward onto the higher-temperature hydrothermally altered rocks that formed from the internal system; the meteoric-hydrothermal argillic and sericite-pyrite alteration zones with their characteristic isotopic compositions are then locally overprinted upon the magmatic-hydrothermal, potassium feldspar-biotite alteration zones or upon the fresh pluton. The major locus of chalcopyrite deposition in the porphyry copper deposits tends to be along the boundary between the tWO different types of hydrothermal systems, suggesting that the mixing interface may have been important in ore deposition (see Lowell and Guil bert, 1970).

Supergene Alteration 01 Porphyry Copper Deposits

267

In some cas es, the primary ores may be dissolved and redeposited by the later-stage meteoric-hydrothermal fluids, wh ich typically have low pH . For example, at Butte, Montana, the repeated opening and reopen ing of tectonic fractures in a fresh, unaltered qu artz mon zonite pluton provided th e permea bility necessary for a very large externa l meteoric system to circulate downward and to leach out primary o re at depths of 2 km o r more. These ho t (-300°C) meteoric fluids th en tra nspo rted the metals upwa rd into the fabu lously rich Butte Main-Stage Cu-Zn-Pb veins, producing low -e 180 , 10w-oO sericite alteration envelo pes, the whole system being dr iven by a largely hidden porphyry copper intrusio n at depth (Sheppa rd and T aylor, 1974; Brimhall, 1980). This porphyry Cu model envi sag es th at a lthough the H 20 in th ese systems has at least two major sources, the ma jor source of the sulfur, copper, and other heavy metals is basically th e magmatic system repre sented by th e porphyry stock itself, w ith a variab le but lar gely unknown contribution of materials leached from th e country rocks. Th e size of the "e xterna l" hydrothermal system varies considerably in known porphyry copper depo sits, from "wet" types (e.g., Butte, Bingham, Morenci) having high pyrite-to -chalcopyrite ratios and surro unde d by eno rmo us halos of pyrit e-sericite-quartz, to " dry" deposits (e.g., Bethlehem ) w ith relat ively min or sericite- pyrite zo nes (Lowell and Guilbert, 1970). These vari at ions in size of the externa l system are to be expected because the permeabilities and avail abil ity of H 20 in th e country rocks will be quite variable. Also, as sho wn by Oilles et al. (1992) and Bowman et al. (1987), more than on e type of wat er may be involved in the external system .

SUPERGENE ALTERATION OF PORPHYRY COPPER DEPOSITS Given favorable circumstances, stable isotope techniques can clearly distinguish between supergen e and hypogene mineral assembl ages. Clay-rich so ils in the United States all plot very close to a curve lab eled " kao linite line " in Figure 6.10 (Lawrence and Taylor, 1971 ); this line represents the locu s of isotop ic data points obtained on pure kaolinites from weathering zon es (Savin and Epstei n, 1970a). These data imp ly that kao linites and montmori llonites formed during weathering are in approximat e isotopic equ ilibrium with their coexisting meteoric waters , and that the fraction ation factors at ambient surface temperatures are such that th ese clays are about 27% 0 enr iched in IS O and about 30% 0 depleted in deuterium relative to th e water from whi ch th ey formed (see Figures 6.1 a nd 6.2 ). We should th er efore ex pect supergene clay mineral s to plot in the vicinity of the " ka o linite line" if they fo rm in eq uilibrium with meteoric H20 at Earth-surface temperatures. Sheppard et al. (1969 ) showed thi s to be the case, if allowance is made for the fact that temperatures of supergene deposi-

268

Oxygen and Hydrogen Isotope Relationships in Hydrothermal Mineral Deposits

rion may range up to 50-60°C because of the large amount of heat produced by oxidation of pyrite during the production of the acid supergene solutions. Based on geologic relations, the probably supergene clays in Figure 6.10 are distinguished by a different symbol from the hypogene clays, and all of the former either plot on the "kaolinite line" or slightly to the left of it; the sh ift to the left is readily explained by slightly higher temperatures of forma tion than are involved in surface weathering (e.g., by the exothermic reaction of pyrite oxidation). On a oD-oISO diagram, such as Figure 6.10, there is a clear-cut gap between the shallow supergene clay minerals and the deep hypogene clays. For example, at Santa Rita, which is the most extensively studied of these deposits, the deeper clay samples (below 5000-ft elevation) have olSO =+6.4 to +14.9 and oD =- 62 to - 71 , whereas the shallow clay samples have 0 18 0 = +14.5 to +18.9 and oD = - 71 to - 88.

MIS SIS SIP P I VALL EY-T YPE LE AD-Z I NC-F LU0 RITE DEPOSITS Fluid inclusion, D/H, and 1S0/160 studies of vein minerals and wall rocks from several Mississippi Valley ore deposits have demonstrated rather conclusively that these lead-zinc-fluorite vein deposits in Paleozoic sedimentary rocks were formed from circulating saline formation waters at approximately 70-1S0°C (Hall and Friedman, 1963, 1969; Roedder et aI., 1963; Pinckney and Rye, 1972; Roedder, 1972, 1977; Richardson et aI., 1988; see reviews by Heyl et aI., 1974; Sverjensky, 1981 , 1986; and Ohle, 1980). Several recent isotopic studies confirm these conclusions. The salinities, chemical compositions, and the range of oD in the fluid inclusions are all remarkably similar to those observed in modern oilfield brines in the nearby Illinois basin (Figure 6.4), as discussed in detail in Chapter 4. Sediment-hosted massive sulfide deposits also may be deposited from relatively low-temperature formation waters or basinal brines. A possible example, shown in Figure 6.16 below, is the Silvermines Zn-Pb-barite deposit in Ireland (Samson and Russell, 1987), which formed at about 140-220°C by mixing of at least two types of NaCI-rich fluids. Ore deposition may have occurred when the higher-temperature fluid (oD = - 2 3 to - 3 0, 0 18 0 = +5 to +9, and 8-12 wt % NaCI eq.) mixed with a lower-temperature, more saline fluid (oD =- 40 to - 50 0 1S0 =+2 to +7,18-22 wt % NaCI eg.; see Figure 6.16 below).

VOLCANOGENIC MASSIVE SULFIDE DEPOSITS Most recent genetic models of massive sulfide ore deposits envisage the op eration of submarine hydrothermal systems, in which ocean water was the major constituent. The metal-rich Red Sea brine may be such a potential ore

Volconogenic Mossive Sulfide Deposits

269

fluid (see Cra ig, 1966). Also, a ll isotopic st udies of hyd rother mal a lteration in ophiol ite co mplexes co ncl ude that seawater wa s by far the dom ina nt component of th ese hydrotherm al so lutions (Spooner et a l., 1974; Magari tz and Taylor, 19 76 ; Gregory a nd Taylor, 1981; Stakes an d Taylor, 1992). A good exa mple o f thi s phen omenon is provided by the pyr ite-cha lcopyrite deposits in Cretaceous pill ow lav as of th e Troodos op hio lite on Cyprus (H eato n and Sheppard, 1977; Schiffman et al., 1987; Schiffma n an d Smit h, 1988). Th e striking similarities between hydrothermal vent frag ments in these massive sulfi de deposits a nd fragm ents of pr esent-day " black smokers" suggest th at the Tro od os ore bodies wer e formed by submarine hyd rothermal systems analogous to th e mod ern on es. The Solea gra be n is th e o ldest and best-exp osed of three asymmetric struc tural grabens ide nti fied within th e sheeted dike comp lex exposed a long the northern margin of th e Troodos o phiolite. Som e o f the largest ma ssive sulfide deposits on Cyprus lie within thi s tectonic grabe n (Schiffma n et a l., 1987) . Skouriotissa (6 x 10 9 kg of 2% Cu prepro duction reserves) is locat ed alo ng the graben axis and Apl iki and M avrovouni (15 x 10 9 kg of 4 % Cu) lie 3 km to the west (Figure 6.11 ). Ther e is a close connec tion between areas of ISO depletion a nd areas of epidos ite alterati on of th e sheeted dike co mplex (Figure 6.11 ). The epidos ites represent zo nes of strong metasomatic hydrothe rmal changes in th e bulk che mistry of th ese mafic roc ks, lar gely in resp on se to replacement of the origin al minerals by epido te and quartz (Schiffma n et al., 1987). The ba saltic ma gm as th at fo rmed th e sheeted dike co mplex initially had ISO ", +5.6 to +6.0 (e.g. , see Grego ry and T aylor, 1981), impl ying that t he ISO -r ich zones in Figur e 6.11 formed by seawa ter a ltera tio n at - 50-250°C, whereas the areas of 18 0 depletion repr esent zo nes of perv asive excha nge by large amounts of hot (>25 0°C) seawa ter. In the ophiolite str atigraphy, th e low_I SO and ep idot e-rich zo nes within the Solea gra be n project directly underneath tw o major o re bod ies in the ove rlying pillow lavas and dik es, notably th e largest such deposit in th is area (Mavrovouni, Figure 6.11). These zones of ISO depletion and epidos ite alteratio n a ppea r to postdate both ma gm atic a nd st ruc tura l sprea ding in the Solea graben, perhaps as a result o f ren ew ed magmatism in thi s region immediately before a ridge jump approximately 8 km eas t of th e Solea ax is (Schiffma n et a I., 1987). Th e extre me depl etions of C u, Zn, and ISO in th e epidos ites sugges t th at high-temper ature hyd rothermal a lteratio n of the sheeted dik e complex co ntributed ore constituents to overlying (now ero ded ) massive sulfide o re bodies (Schiffman et a l., 19 87). If the o re deposits and epidosi tes in th e Solea graben formed in a n off-axis setting, ther e must have been a related magm at ic event capable of driving hydrothermal circulation. Th e outc ro p pattern s o f upwelling zo nes within the dik e co mplex thu s may reflect th e distribution of these o ff-ax is, hypabyssal intrusio ns. Although th e inferred off-axis intrusions are largely un exposed , Schiffman et a l. (1987 ) have identifi ed o ne sma ll gabbro st ock northwest o f Pedhoulas (Figure 6.11) th at is relatively un altered

< 270

Oxygen ond Hydrogen Isotope Relotionships in Hydrothermol Minerol Deposits

I

o

•

2

L......l...-J km

.....

MgsAhSiJOlo(OH)s + 10 H+ (sea) (rock) (rock) (aq) (chl) (aq) (7.1)

The H+ production drives the system toward small values of 10g(aK+jaw) and eventually to quartz and pyrophyllite precipitation at point Q (Figure 7.3) (Reed, 1983). The subsequent evolution of the aqueous phase is dominated by acid neutralization, yielding muscovite then microcline (point M), as well as the normal suite of greenschist minerals including epidote, albite, and actinolite. Although this diagram and others like it perfectly represent mineral stability relationships in subsystems of the whole, they are not suited for predicting the process of reaction because they do not provide for complex mass action effects involving simultaneous equilibria, such as acidification of the fluid by chlorite precipitation. No fully satisfactory alternatives to the conventional diagrams have been developed, but some of their shortcomings are overcome by the stacked, reaction path diagrams used here to display the results of water-rock reactions. In these diagrams, the abscissa represents a measure of composition in the fluid-rock system in terms of wjr, which is an explicit expression of composition in a water-rock system computed by dividing the mass of rock in the system into the mass of solvent water in the initial aqueous phase. Increasing rock mass corresponds to decreasing the wjr ratio when reading the graphs from left to right. Most of the diagrams displayed in this chapter apply to closed-system rock titrations, wherein no material leaves the chemical system after it has been added. As such, each position along the abscissa represents a unique and complete expression of the equilibrium condition for that particular compositional mixture of rock and initial fluid. Because such diagrams are based on whole-system calculations of simultaneous equilibria, every chemical variable having a bearing on the fluid-rock reaction can be represented. To simplify the reading of the diagrams, the number of variables displayed is limited to those necessary to communicate the fundamental relationships among minerals and aqueous species. Although each such set of diagrams is based on reaction of a single starting fluid with a single rock, the generality of the diagrams is broader than it appears because the composition of the fluid changes continuously with changing w jr; each different wjr defines reaction of the rock with a different fluid than those at other wjr values. For example, in the following examination of magmatic acid condensate reaction with andesite, at large wjr the reaction of andesite with an acidic, oxidizing fluid yielding one min.eral assemblage is explored, but at small w Ir, a reaction of andesite with a neutral-pH, moderately saline, reducing fluid yielding a different mineral assemblage is explored. The small wjr portion of the calculation can readily be regarded as a neutral-pH reaction without ever considering any relation-

316

Hydrothermol Alteration ond Its Relotionship to Ore Fluid Composition

ship to the aci d ic origins of the fluid in th is particu la r calc ulation. It is clea r th at the nume rica l val ue o f w/r on ly has mea ning w he n th e in it ial w ater co mposit ion is speci fied; w/r is never th eless useful as a rela ti ve varia ble, independe nt of its pa rticu la r numerical va lue.

W ATE R/R 0 (K RAT lOA SAD ES( RIP T I V E V A R I A BLE It is a matter of numerical practica lity th at ca lculatio ns of wa ter-rock reaction a re executed by titrating rock into wa ter; co nseq ue ntly, it is customary to gra ph results sho w ing th e am ount o f titr ated rock inc reasi ng fro m left to right, co rresponding to w/r cha ng ing from large to sma ll, as in Figure 7.4 . This a rra nge ment o f th e graph co nveys no specia l geochemica l significa nce. Th e geoc he mica l sett ing d ictates wh eth er it is more useful to co nside r th e evo lut ion of th e geoche mica l syste m from sma ll w/r to large, o r vice versa. Co rrespo nd ingly, th e gra phs ca n be read fro m left to righ t o r from right to left . Ap plication of th e closed-system w/ r to geoche mica l sys te ms is strictly valid only wh ere the aq ueo us phase is sta tic. This limits somew ha t th e a pp licabi lity of ca lculation resu lts to interpreting spa tia l a nd temporal re lations in th e natural setting. The sta tic w/r is, however, a power ful a nd sim p le va ria ble a pplica ble in a broa d range of systems. The use o f w/r a nd th e ca lculatio ns it represents is a co mpromise between co nsidera tion o f sim ple co nve ntiona l equilibrium relati on sh ips in limited subsyste ms (as represented in activity diagra ms) a nd th e idea l of a fully integrated tr eat ment th a t incl ud es infiltrati on , diffusion , and kin et ics in sp ecific spa tia lly a nd temp orally d efin ed syste ms. React ion ca lculatio ns ex pressed in terms o f th e sim ple wa te r/rock rati o ca n be appl ied in an approximate wa y to inte rpreting spat ial a nd tempora l relati on sh ips in real systems, as illustrated in Figure 7 .2, wh ich shows a fracture bordered by a di ffusion-gener ated, zo ned alte ra tio n hal o, including a veinwa rd zo ne that pinches o ut in th e direct ion of fluid flow. The vei nwa rd zon e pinch es out as th e a bility of th e fluid to effect th e veinward a lte rat ion is diminished by reaction with th e wall rock (e.g., acid ic fluid is neu tral ized) . Through tim e, th e diffusion halo and its individ ua l zo nes adva nce outward from th e vein (Sales and M eyer, 1948) and th e pinchout o f th e veinward zo ne adv ances in th e flow di recti on . The seque nce o f a lte ration min eral asse m blages from th e vei n outwa rd a nd th e seq ue nce o f th e pinching veinwa rd assemblages along th e fluid infiltration di recti on cor respond to th e sequence o f miner als th at form with decr ea sing w/r in a water-rock reaction calculation (e.g., Figure 7 .4 ). Rock that is pr oxim al to th e in filtra tion source or to th e fluid-filled fra cture "sees" more fluid th an d ist al rock , a nd fluids that arrive at distal points have changed com position by reaction with rock along th e way, eithe r directl y or by excha nging solutes through th e diffusion halo with fluids reacting with fresh ro ck. Thus proximal zones correspond to larg er w/r ratio th an do distal zon es. Any indi-

Reaction of Acidic Fluid with Wall Rocks of Basic to

Fe~ic (ompos i~on

3 J7

vidual parcel o f rock in th e a lte ra tio n halo evo lves fro m sma ll to large w/r, corresponding to th e seq uence o f asse mblages from th e o utside of the halo inward , and an y given fluid parcel tr aversing the fra cture evolves fro m lar ge to small w/r. At an y fixed po int along th e flow path, the tem por al evol utio n of the fluid and adjacent wall ro ck is fro m small to large w/ r, as fres h pr imary fluid (Figure 7.1) co nti nues to ente r the syste m (e.g., Figu res 7.9 and 7.10). On th e othe r hand , if fresh fluid does not co nt inue to ente r th e system, co rrespo nd ing in Figure 7.2 to stopping th e fluid influx while fresh ro ck remain s beyond th e alte ration halo, th e co ntinued reaction of fluid with fresh rock would ca use th e a lrea dy -fo rmed large w/r assemb lages to be repl aced by a smaller w/r assemblage . Clea r exa mp les o f lar ge-t o-small w/ r evolution systems include sandsto ne diagenesis o r seawa te r react ing with basaltic hyaloclastites (Reed, 19 83). Another setting where thi s view clearly appl ies is in the bulk-system evolution of geo ther ma l syste ms in genera l: ove r time, mor e and more origi na lly fresh ro ck reacts with th e fluids, th e alteration asse mb lages evo lve toward the rock-dominated , propyliti c asse mblage, and the fluids evolve tow ard a neutral-pH , reduced , ro ck-imp rinte d form in equilibrium with a pr opylitic assembl age . In co ntras t, in portion s of th e system wh ere th e supply of fresh fluid is dominant, th e syste m evo lves from sma ll to large w/ r. If all of th e rock at a fixed loc ati on in th e rock mass has alrea dy evolved to eq uilibriu m with a sta tic po re fluid, th en th e fluid is repeat edl y replaced by un reacted fluid; for example, when fresh magmatic co nde nsa te infiltr at es th e base of a breccia body, th e min eral asse mblage will evolve from small to lar ge w/r. If the supply o f infiltrat ing fluid is sufficient and sustai ned, th e alt eration assemblage would evo lve to a fluid-d ominated on e. For exa mple, rock attack by magmatic acid co nde nsa tes yield qu artz or quartz-alun ite (Figure 7.4 a) and seawate r yield s chl orite (Reed , 19 83).

REAOCTION 0 F A CI DI C FLU I D WIT H W ALL ROC KS 0 F BASIC TO FELSIC COMPOSITION Many fundamental alteratio n processes essentia l to alte ration of all typ es can be explo red in a reaction o f an acidic (pH 0.8 , 300°C) fluid with ro cks ranging in composition from basalt to dacite. Such reaction s are of interest primarily for a direct explo ra tion of th e nature of alteration reactions, bu t they also have implications fo r the origins of epithermal or e fluids and the role of wall rocks in controlling met al concentrations in fluids as they tr averse from a hypogene source to th e ore body. Both as pects are explo red below. The acidic fluid (Table 7.1) is a met al-ri ch volcanic gas condensa te from a high -temperature fumarole on Augustin e volcano (Dome-3 sample, Symonds er aI., 1990) diluted with ten parts water to simul at e th e mixing of ma gmatic gas condensate with ov erlying groundwater (Reed, 1992a). Although recent reevaluation of thi s gas composition (R. B. Symonds, personal comrnunica-

-

CD

W

°3

0 .79

0 .403 0 .2 17 0 . 160 0 .00885 0 .000545 O.l 77 x 10- 5 0 .30 2 x 10- 5 0 .315 x 10- 6 0 .105 x 10- 4 0 .57 1 X 10- 4 0.920 X 10- 4 0 11 2 x 10- 6 0 .420 x 10- 5 0 .453 x 10- 6 0 .158 x 10- 5 0 .65 7 x 10- 10 0657 x 10- 10 0 . 118 x 10 - 7 0. 184 x 10- 8 0 .250 x 10- 6 0 .650 x 10- 5

13643 19955 9 272 280 31.3 00457 0 .1 16 0 0073 3 0 .561 2.14 0 .0586 0 .263 0 .0275 0 .312 0 .677 x 10 - 5 0 .16 8 x 10 - 4 0 .0038 7 0 .00024 ) 0783 0 .04 12 0.783

A 1: 10 (Magmatic: Meteoric) ppm m

- -

Model Fluid Composifions

0.79

0 .996 x 10- 6 0 .4 18 0 . 159 0 .00708 0000545 0 .177 x 10- 5 0 .302 x 10- 5 0 .315 x 10- 6 0 . 105 x 10 - 4 057 1 x 10- 4 0 .9 19 x 10- 4 0 . 112 x 10 - 5 0 .420 x 10- 5 0 .453 x 10- 6 0 .158 x 10- 5 0 .657 x 10- 10 0 .657 x 10- 10 0 . 118 x 10 - 7 0 . 184 x 10- 8 0 .250 x 10- 6 0 .650 x 10- 5

m 0.033 6

ppm 37292 9 228 223 31 2 0 .04 55 0 11 5 0 .00729 0 .558 2 13 2.01 0 .00584 0 26 1 00274 0 .31 1 0 .67 4 x 10- 5 0 .167 x 10- 4 0 00 385 0 .000240 0 .04 10 0.779

(low·HCI Fluid)

B

4 .03

0 .68 2 0 .234 x 10- 4 0 .0 101 000 160 0 .00770 0 .500 x 10- 4 0 .0250 0 .350 x 10- 4 0 .000240 0 .050 0 .580 0 .500 x 10- 5 0 .000300 0 .000105 0600 x 10- 4 0900 x 10- 6 0 430 x 10- 8

m

27083 - 0 .70 2 257468 52.05 474 0 .0263 37 .9 0 .0 195 0 .0 124 2502 16399 0285 0.0440 0 .890 0.00746 0 .009 27

I 104 - 0 . 105)( 10- 4 6 . 10 000228 0 .0114 141 x 10- 5 000 1368 0 . 116 x 10- 5 0 .321 x 10- 6 0 .092 5 1.031 2 10- 5 10- 5 10- 5 10- 6 10- 7

232 13 2. 15 59 2 50 .9 444 1.30 96 2 0 .82 12.86 1877 12801 0 .264 1883 6 .38 11.93 0 .0932 0 .00 110

5 . 18

0 .630 x 0 . 100 x 0 .62 1 x 0 . 100 x 0 .500 x

ppm

m

0 (Greenstone)

ppm

(Dolomite)

C

A =Magmatic condensate mixed with pure water; used for fluid- rock reactions shown In Figures 7 4, 7.5, 7 .6, and 7 10 B = HypolheticallowHCI magmatic condensate mixed With pure water, used for fluid- rock reactions shown in F'gures 7 .6 and 7.7 C = Hypothetical lhnd for rea ction With dolomite to yield Zo-Pb skarn (Figure 7 I 1I. D = Hypothetical fluid for reoction with amphibolite to yield greenstone gold ore IFigure 7, 121.

pH 1300 °CI

No' Mn2• Zn2• Cu' Pb2• Ag+ Au ' Hg2. Ba 2 • SblOHb H3As

K'

~f Fe •

Co2•

AI)'

S024 HCOJ HSSi02

(1-

Component

TABLE 7.1

Reaction of Acidic Fluid with Wall Rocks of Basic fa Felsic Composition

TABLE 7.2

Bulk Rock Iomposiliors?

Basalt Si02 AI203 Fe203 FeO MnO MgO CoO Na 20 K20 BoO NaCI H2O CO2

319

50 .1 2 wt % 17 .4 3 3.09 6 .5 2 0 .1 7 6 .74 10.15 3 .0 3 0 .80 0 .0478 0 .033

Andesite 59 .07 w t % 17 .4 3 2.56 4 .05 0 .12 3.35 6 .77 3.9 3 1.48 0 .0692 0 .033

Dacite

Amphibolite

66.88 wt % 15.57 1.09 2.74 0 . 11 0 .60 2.16 5 .55 4 .09 0 .102 0 .115

5 2.8 wt % 15 .4 0 .76 5 .4

S CU20 PbO ZnO Sb203 As203

40 0 ppm 105 6 .7 92 0 .1 2 2.64

200 pp m 48 8 .2 97 0 .24 3.3

100 ppm 33 39 97 0 .48 4.6

Au Ag Hg

4 ppb 100 10

4 ppb 80 10

2 ppb 50 10

7 .59 9 .77 3.08 0 .7 0 .033 2.40 .55 25 ppm 6 .7 75

4 ppb 1 ppm

8 Au(HS)i + 6 H+ + 4 H20

(aq)

(aq)

(7.15)

(aq)

The overall dilution-driven, gold dissolution reaction is the sum of reaction (7.15) and the buffer reaction (7.7). The wall rock-mediated dilution effect on gold is probably the fundamental cause of the distal zonal position of gold in many magmatic-hydrothermal systems. It is also likely that such epithermal gold concentrations form where early-precipitated gold is remobilized by evolved, diluted fluids. The combined effects of pH increase, dilution of ligand chloride, and increased total aqueous sulfide cause base metal concentrations to decrease out of proportion to simple dilution, as is apparent from their steep ~ega­ tive slopes in Figure 7.8b (compare to Hg 2+). For example, early in the dilution, alteration galena precipitates in accordance with the following reaction, which is driven to the right by dilution: PbCl J + H2S =:> PbS + 3 Cl" + 2 H+ (aq) (aq) (gn) (aq) (aq)

(7.16)

The stoichiometric relationships in the reaction make it clear that increased pH drives galena precipitation as does increased activity of H2S. Dilution

342

Hydrothermal Alteration and Its Relationship to Ore Fluid Composition

of aqueous CI- also drives this reaction because diluted chloride activity is cubed in the ma ss-action relationship wh ereas diluted PbCl 3 activity goes as the first power. The increase in aqueous H2S (Figure 7.8c) results from dis solution of pyrite, driven by transfer of iron from pyrite to actinolite in response to increasing pH: 5 FeS2 + 13 H20 + 8 Si02 + 2 Ca 2+ (py) (aq) (qz) (aq) => Ca2FeSSiS022(OHh + 6.5 H+ + 1.25 SO~- + 8.75 H2S (act) (aq) (aq) (aq) (7.1 7)

At extreme dilution, the total concentrations of aqueous Cu " , PbL + , Zn 2+, and Ag" increase relative to the dilution trend (slopes of concentration curv es are shallower than that of Hg 2+, Figure 7.8b) because the complexing ligand changes from CI- to HS- in response to dilution of chloride and increasing bisulfide concentration. BARITE IN AlTERATION AND GANGUE

Reaction of the acidic magmatic fluid with dacite, andesite, or basalt yields alteration barite (e.g., Figure 7.4d) as barium from the rock reacts with sulfate in the fluid. Barite in altered wall rock, as opposed to gangue, is an indication of the action of sulfate-bearing alteration fluids . Equilibrium with alteration barite dictates the mirror-image relationship between ,t he aqueous Ba2+ and SO~- of Figure 7.4c,e; where SO~- concentration is large, Ba2+ concentration is small, and vice versa. Because barite is relatively insoluble, a significant concentration of sulfate depresses aqueous Ba 2+ concentration to small values. Aqueous sulfate is reduced by rock reaction [reaction (7.9)], ultimately causing the barite to dissolve, releasing barium to the aqueous phase (Figure 7.4e). The aqueous Ba2+ grows to extreme concentrations (e.g., > 5000 ppm, Table 7.4, columns F and G) at small w/r in the same fluids that contain substantial concentrations of Ag with lesser Au. The same buildup of aqueous Ba2+ would be expected to occur in any fluid-rock system involving reaction with a reducing rock (e.g., seawater-basalt systems), except C02rich systems, where witherite precipitation limits aqueous Ba 2+. When a Ba-bearing, reduced fluid mixes in a shallow environment with acid-sulfate waters descending from a zone of surficial oxidation of sulfide gas condensate, gangue barite would precipitate with gold, the latter by acidification (Reed and Spycher, 1985; Spycher and Reed, 1989). Such an assemblage with high-grade gold formed in late, shallow veins at Summitville (Stoffregen, 1987), suggesting that late, reduced, gold-rich fluids evolved in this system. In general, gangue barite forms where reduced waters

Wol~Rock

Reoction ConrlOls on Ore Fluid Composition

343

mix with sulfate-beari ng waters. T hus a bundant ga ngue ba rite marks th e arrival of quite reduced fluid s at a site of mixing with sulfate-bea ring fluids. Th e sulfate-bearing fluid ma y be seawa ter, for exa mple, yielding th e upper barite zo ne in Kuroko-type ma ssive sulfi de deposit s or similar vein deposits in marine volcaniclastic s suc h as at Johnson River, Alaska (Steefel, 1987). EPITH ERMAl Al TERA TlO N AN 0 0RE: CO NClUSI 0 N

Wallrock neutraliz ation and redu ction of an init ially acidic, oxidized, ma gmatically derived fluid mak e it possible th at a single magmatic-hydrothermal system co uld evo lve th rou gh a series of chem ical states yielding a range of epithe rma l fluid typ es cha racteristic of acid-sulfate Cu-Au (ev-en-Au) to base metal-silver-gold veins. Upon diluti on , fluids from the latter system may evo lve to th ose th at form adularia-silica- Au- Ag ores. In th e abs ence of rep eated pul ses of ma gm at ic volatil es (an d excluding surficially produced acidic water s) (Reed and Spycher, 1985), we would expect a single hydrothermal system and its product o res to evolve from the acid-sulfate to neutralpH typ e as the initially acidic, oxidizing fluids are neutralized and reduced by wall-rock reacti on (M argolis et a!., 1991; Reed , 1994). Such evo lution is found at El Indi o (Jannas et al., 1990) and is indicat ed at Summitville by the late barite-gold veins (see above) and lat e coatings of covellite by ch alcopyrite, but in gene ra l, more th an o ne epitherma l style may not be ex pressed in a singl e dep os it becau se th e details of the district-scale hydrolo gic regim e favor on e o r ano ther of th e enviro nments.

Andesite Alteration by an Infiltrating Fluid Another view o f th e int erd epend ence betw een th e alteration assem blage and fluid compositi on in a flow ing system is obtained by calculating a single-volume fluid -flush process wh er ein the pore volume of a body of rock is rep eatedly filled with fluid , equilibrat ed with the rock, then flushed o ut and refilled with fresh fluid (Figure 7.9). This calcul ation explo res th e alteratio n history of a single volume of rock along a fluid infiltr ation path , assuming local eq uilibrium between fluid and rock , using the andesite (Tabl e 7.2) and initial fluid A (Tabl e 7.1) applied in the andesite reaction, above. In the gr aphical display of th e calculation (Figure 7.10), the a bscissa shows th e logarithm of th e throughput w/r. As the reacti on proceeds, th e porosity of the rock increases from its initial a rbit rary valu e of 25 % (Figure 7.1Oc) in response to the changing alter ation assemblage (Figure 7.10a); the volume of each successive aliquot of input fluid is adjusted to fit the ava ilabl e pore space in each reaction segment. A comparison of this calculation typ e to the conventional rock -into-water calculations (e.g., Figure 7.4) illustrates the point stressed above that the conventional w/r diagrams can be applied in either sense, incr easing or decrea s-

344

Hydrothermal Alteration and Its Relationship to Ore Fluid Composition

VOLUME UNIT

0 1nll

, 0

=25% 43%

en d =

I

~ "'''§*('&''§*~ ''§>",'N'*''>¥f&K~~ffi%'''~~~'''''2S'%-m''''S *2,' J

Fluid In pH 0 .8 (always)

.

~

J .._____r

~

"'-..

I~

/

r

Fluid Out pI I 5.7 (early) pI10 .8 (end)

.:;e. III

Q)

_

--....... ,- _ _ -qz

..-' /" fij'\

~

o

"0 E C>

o

-1

...J

-1

-2

-3 -4 -5

-o~

.~ III o o a. ~

-6 h;:::;':==~=-==-=~

~n: --: ~ '-' -.~ ---:---.-:~ >~j -1.0

-0 .5

0.0

0.5

1.0

1.5

Log w/r FIGURE 7.10 Infiltration reaction of andesite (Table 7·2) with diluted acidic, magmatic gas condensate (Table 7·1, column A) at 300°e. Reaction scheme is iIIuslJOted in Figure 7.9. The abscissa shows the logorithm of the throughput I'I/r, increasing fram left to right, which in this calculation also corresponds to increasing time. Mineral abbreviations ore shown in table 7·3. (0) log moles of major minerals present per kilogram of initial rock. (b) pH and log total molality of principal component species in the aqueous phose. Each concentration plateau corresponds to a distinct mineral assemblage and is defined by many calculation points. (c) Porosity (%) of the rock volume. (d) log moles of minor alteration minerals present per kilogram of initial rack. (e) pH and log totol molality of are metels in the aqueous phose. (I) log of molalities of individual species that ore significant in redox and odd-bose equilibria.

are fixed by th e solubilities of th e respect ive min erals. Otherwise, aq ueo us concentrations a re thos e of th e input fluid (e.g., th e base metal s at large wjr, Figure 7.10e ). The calculation clearly illustrat es th e rol e of wall rock in modulating th e metal concentrations in hyp ogene fluid emerging from a hypothetic al deep magmatic ga s condensate environment such as that suggested a bove for th e andesite reaction setting. Becau se th e propylitic assemblage buffers pH in th e neutral range, st abilizing alteration sphalerite, galena, and chalcopyrite (Figure 7.10d) , ba se met al concentrations are small in the early fluid (Figure 7.10e). Addition of acidic fluid suc cessively destroys silicate buffers in th e pH buffer series ex plained above, causing pH to step downward and solubility-

aq

346

Hydrothermal Alterotian and Its Relationship to Ore Fluid Composition

- -ca - "' I

0)

-

.:.: I II

-3

Q)

.~====4----SI_==:::-_---------- - - - - - - - - - 1 ....... I cp • \

(5

E o

0)

I

-4

\

!

I '- '- 'gn'- 'l '

--I

O)

o

--I

-3

I

!

-,

-e-

__.

bar

@

zn2+ J' r: - \ \0 _ Bl!..2+.r. _ - J c: + .~ + N ._._._ ._ . / . ~ \ zn2+ ~ .... aH + ---J h - r--_:J". .

:~.

Iron

+ pyrrhotite Pyrite

+

sulfur

Fe

FIGURE 8.3 Phose with permissionl.

80

relo~ons

s

inthe Fe-S syslemobove 400 0 ( at low pressure (from Vaughan and (raig, 1978; used

matrix, but during slow cooling such text ures are often tr an sformed into a nhedra l masses with no evidence of their original compl exity. On the sulfur-rich side of pyrrhotite, the min imum temper ature where a sulfide liquid exists is 988 °C; above this temperat ur e, two melts fo rm a br oad field of liquid immiscibility in wh ich the sulfide liquid coexists with a liquid that is very sulfur-rich. Pyrite , th e Earth 's most abundant sulfide, has at low pressures a maximum thermal stability of 742 °C ± 1°C. The upper th ermal stability of pyrite rises -14°C per kilobar of confining pressure (Kullerud and Yod er, 1959). The common occurrence of pyrite with pyrrhotite and the dependence of th e composition of that pyrrhotite on temperature led R. J. Arnold (1962) to propose the pyrrhotite-pyrite geothermometer. He determined compositions on the pyrrhotite solvus and calibrated an Xray diffraction spacing curve to find the composition of the pyrrhotite by the position of the (102) reflection. The Xray spacing curve was subsequently improved by Toulmin and Barton (1964) and Yund and Hall (1969) resulting in the equation for hex agonal pyrrhotites: Atomic % iron

=45.212 + 72.86(d(l02l -

2.0400) + 311.5(d( 102) - 2.0400)2 (8.3)

The Iron-Sulfur System

375

0.98 0.96 0.94 0.92 0.90 N FeS FIGURE 8.4 Variation of the XRD measured spacing of the (102) peak, d002l , of hexagonal pyrrhotite as a function of composition (from Vaughan and Craig, 1978, adopted from Toulmin and Borton, 1964; used with permission). The uncertainty in the dimension is about 0.005 A.

This method remains the most convenient accurate means of determining the composition of hexagonal pyrrhotites (see Figure 8.4). Unfortunately, the utility of the pyrrhotite-pyrite geothermometer is severely limited in natural systems where rapid reequilibration between pyrrhotite and pyrite occurs while samples cool from their initial crystallization temperatures. During this cooling, the pyrrhotite composition slides down the solvus to a more iron-rich composition. ~fu rth er complication in many ore deposits is recrystallization to monoclinic Ry-.rrhotite_either during , cooling below- about256" ~~ a lo~-tem erature oxidation product. Consequently, the pyrrhotite-pyrite geothermometer can be used best for experimental studies in which rapid quenching is easily achieved, the few natural occurrences where rapid cooling has taken place, or where the pyrrhotite exsolved during cooling is texturally distinctive and its concentration can be determined. Thermochemical studies of the iron-sulfur system were initially conducted in order to understand various aspects of smelting and refining processes (e.g., Richardson and Jeffes, 1952; Rosenqvist, 1954); however, studies that are more complete and important from a mineralogical and petrological

376

Sulfide Ore Mineral Stabilities, Morphologies, and Inlergrowth Textures

viewpoint are th ose o f Toulmin and Barton (196 4 ) and Rau (1 976). The form er related pyrrhot ite co mposition to th e activity o f S2 gas an d presented th e now fami liar activiry'-L'T diagram (Figur e 8.5 a ) th at is used com monly to correla te sulfide sta bilities. They also esta blished th e co nventio n o f ex pressing pyrrh otite co mposi tions in terms of th e mole fracti on o f FeS, XFcS, in the syste m FeS-52 , moles FeS XFcS =-....,.----=-::-- --,;----::;-moles Fe5 + moles 52

(8.4)

Thus th e mol e fracti on of Fe5 is equivalent to twice th e a to m ic perc ent of iron in pyrrh ot ite (e.g., an XFcS of 0.98 is the same as 49 mol % iron and a n XFcS of 0.94 is eq ual to 47 mol % iron). Using the e1 ectrum tarnish method as a calibration technique, Barton and Toulmin (1964 ) determ ined th e variation of the activity of 52 gas in equilibrium with pyrrhotite ov er a wid e temperature range. An expression relating the activity of su lfur to pyrrhotite co m pos itio n is log as!

=(70.30 -

85.83X)(1000/XFcS - 1) + 39.30(1 - 0.9981XFcS) - 11.91

(8.5) a nd th e result is seen in Figure 8.5a in which the pyrrhotite field is contoured in terms of XFcS. Barker and Parks (1986) have reevaluated th e th ermodynamics of pyrrhotite and concluded that " Rau's (1976 ) data for pyrrhotite . .. and the pyrite/pyrrhotite buffer ar e more precise th an , but consiste nt with, th e results of Toulmin and Barton (1964)." Gronvold a nd Stolen (19 92) also explo red the variety of mod els used to int erpret experimental m easurem ents of sulfur fugacities a ~o~e the pyrrhotite solid solution and found th at a sta tisti cal model based on a regular solution model is capable of reproducing most of the experimental results. ' The most important and useful curve in Figure 8.5 a shows the activity of S2 gas as buffered by coexisting pyrite plus hexagonal pyrrhotite. Barton and Skinner (1979) have referred to this curve as the " ma in line" because of its importance in defining and maintaining the activity of sulfur over the assemblages in so many ore deposits as they cool and during any subsequent metamorphism. Although, at temperatures below 598 K (325°C), no measurements of the asz above pyrite-pyrrhotite have been published due to experimental problems in making reliable measurements of extremely low sulfur pressures, caculations have been performed (A. Lennie, personal communication) using thermodynamic data published by Gronvold and Stolen (1992). Above about 300°C, during any temperature change, th ere is relatively rapid reequilibration of pyrrhotite compositions with coexisting pyrite. This eq u ilibratio n changes not only the proportions of the phases but also th e composition of the pyrrhotite as it slides along th e solvus (Figure 8.3).

The Iron-Sulfur Sysrem

473

373

273

573 673

377

T/K

-10 N

~

Ol

-20

.3 -30

3 .6

3.2

2.8 10 3fT,

2.4

2.0

1.6

K

FIGURE 8.S (0) The composition of pyrrhotite in the Fe--S system as 0 function of temperature and the activity of S2 (adapted from Toulmin and Borton, 1964). The bold, solid line is the pyrite-pyrrhotite buffer curve, and the light solid contours are pyrrhotite compositions in terms of XfeS, the mole fraction of FeS in the FeS-S2 system. (b) The pyrite-pyrrhotite and iron-rlOilite buffer curves at lower temperatures (adapted from A. R. Lennie, personal communication). Filled squares represent experimental data from Schneeberg (1973) ond open circles are dora colculored from Gronvold and Stolen (1992).

Pyrrhotite initi~.1!Y...J9rmed with pyrite or metamorphosed in equilibrium with pyrite, relJases sulfur during slow cooling. The released sulfur-forms ~ pyrite, thus increasing the pyrite/pyrrhotite ratio. During any prograde metamorphic episode, the reverse process would occur; that is, pyrrhotite would take up sulfur from the pyrite that is consumed. Craig and Vokes (1986, 1993) have considered these effects and pointed out that the pro-

378

Sulfide Ore Mineral Slobililies, Morphologies, and Intergrowlh Textures

portions of the pyrite consumed o r p roduced as a result of the changing pyrrhotite composition is dependent on the initia l ratio of pyrite to pyrrhotite and on the change in temperature. Thus ores with high pyrrhotite/pyrite ratios have greater sulfu r exchange, and greater change in the amounts of pyrite present, than do ores that contain low pyrrhotite to pyrite ratios . Figure 8.6 illustrates the amounts of py rite that would have formed on cooling over three different temperature intervals as a function of the amount o f pyrite fou nd in t he ore after coo ling. Such relations are based on assumed equi libri um throughout a deposit, and that the ex change of sulfur is isoche mica l during cooling or du ring the metamorphic cycle and must, of course, be used with considerable caution. Even if these conditions are maintai ned down to 300°C, the proportions of pyrrhotite to py rite in vario us hand sa mples w ill show large varia tions . D uri ng cooling from initia l crysta llization co nditions or coo ling du ring ret ro grad e metam orphism, the sulfur lost by pyrrho tite may form pyrite eit he r by cxso lurion or by overgrowth of pyrite o n preexisting pyrite crysta ls. W he re both processes a re active, th e resulting pyrites are indistinguis hab le in casual, and sometimes ca refu l, o bservation. The tendency of pyri te to grow as euhed ra l crys ta ls (usually cubes) has commonly been ca lled the "force of crys ta lliza tion " and is so effective that the tiny laths or lenses seen during rapid exsol urion in experimental studies do no t long persist. Evidence that pyrite growth does take p lace d uri ng coo ling is commonly observed in the forms of euhedral crystals, overgrowth

1/1 1/1

60

I'll

E

50 0 a. >a. 40 I

I'll

c:

=

30

-...

20

c: C)

>- 10

a.

::!! 0

0

o

10

20

30

40

50

60

% Pyrite due to sulfur loss by pyrrhotite FIGURE 8.6 Ttuee curves iIIusllOting the amount of pyrite formed os a result of sulfur released from pyrrhotite during cooling from 6000 ( to 300'( (open squares), from 500' ( to 3000 ( (filled squares), and from 400 0 ( to 3000 ( (mostly filled squares) (from (ro~ and Vokes, J993).

The Iron-Sulfur System

379

patterns in wh ich it is possibl e to distinguish inn er from o uter zo nes o f cry stals by inclusions, ch emical differen ces, ha rdn ess va riation s, and so o n, and som etimes in isotopic variations. Cra ig et al. (199 1) sho wed a n increase in 0345 from core to rim in a large pyrite cr yst al th at a ppa rent ly grew during cooling. Figures 8. 7a a nd 8. 7b sho w two co ntinuo us overgrowth pattern s seen in pyrite a nd o thers hav e been sho wn by Brook er et a l. (1987). The ch an ge in the number o r nature o f inclusion s, as sho wn in Figure 8.8, no doubt reflects changes in th e (fo r mat io n co ndi tions between th e tim e th e core formed and th~ o~tsid'e rIm "fbr med . In Figure 8.7a, th e lon g dim ension s of th e incl usions di splay a moderately well-d eveloped co ncent ric pattern. Th is may result either from simultaneous growth o utwa rd in whi ch th e pyrite overgrows minerals that first la y with th eir lon g axes adjace nt to th e pyrite crystal surface, o r from rotation o f th e gro wing crystal such th at it incorporates inclusions lik e a rolling "snow ba ll." The pattern displ ayed by a rotating and growing pyrite crys ta l w ould dep end o n th e manner in which it is cut and examined, as show n in Figure 8.9. A section cut parallel to th e axis of rotation would yield a co ncentric pattern (Figures 8.7a and 8.9) wh ereas a sectio n cut perpendicular to th e axis o f ro ta tion would yield a helical pattern (Figures 8.7b and 8. 9). Ca rste ns (1941 , 1944) published th e first reports of such pat-

. ,-

. . "~,:

,~

. .,

~

-(' ''':'. ' \

..

.. ..". t~

.

"' ,

I ~ .' .. :

,

... . . -. ... '

\

I

~-

..

.

,

.

'",

FIGURE 8.7 (0) Concentric growth pattern in a pyrite crystal, as revealed by inclusions. from Ducktown. Tennessee (centimeter scale). (b) Spiral inclusion pattern ina 1Q-cm pyrite cube. As described by Craig et 01. (1991). the inclusion

pattern implies a 3600 rolotion of the pyrite relative !o th~ host pyrrhotite during growth. ('f"

• f:

• •

('~) c...

380

Sulfide Ore Mineral Stobilities, Morphologies, and Inlergrowth Textures

FIGURE 8.7

(Continued)

terns. Brooker et al. (1987), Craig et al. (1991), and Craig and Vokes (1993 ) have also demonstrated that such patterns are, at least locally, reasonably common. The growth of euhedral pyrite crystals in hydrothermal systems is commonplace and produces remarkably beautiful crystals in some localities. Logrono (Spain) and Elba (Italy) are especially well known for museum quality specimens up to 5 ern across. The largest pyrite crystals appear to form, however, under metamorphic conditions. Templernan-Kluir (1970) and McCla y and Ellis (1983) have shown that, in general, increasing grade of metamorphism results in overall increase in the size of pyrite crystals. Typically, this relationship shows that pyrite crystals formed in the micrometer size range at lower temperature increase to the millimeter size range at amphibolite grade metamorphism. The localities with the largest documented pyrite crystals are the Gresli deposit of southwestern Norway, where crystals up to 15 ern are visible (Craig and Vokes, 1993) and the Ducktown,

The

lron-5u~ur

5ysrem

381

FIGURE 8.8 Pyrite crystols from the Moshon DislTict, Anhui Province, Chino, showing differences in the numbers of inclusions os growth proceeded. (Field of view is 1.2 mm.l

Tennessee, dep osit wh ere crystals up to 30 ern are observed (Brooker et aI., 1987). In both dep osits, th e pyrite crystals occur in a matrix of hexagonal pyrrhotite and underwent amph ibolite grade metamorphism (550°C, 6.1 kbar at Ducktown; Brooker et aI., 1987). Below about 300°C, it is likely that reequ ilibrati on between the phases is retarded by slow reaction rates and that the pyrrhotite would react to lower temperatures, independently of the pyrite, as described below.

Low-Temperature Phase Relations In spite of numerous studies, the phase relations in the iron-sulfur system at temp eratures below 350°C are incompletely understood. Earlier work on the system is reviewed in detail by Power and Fine (1976) and Vaughan and Craig (1978). The most recent comprehensive study undertaken using hydrothermal recrystallization methods between 350°C and 115°C has been that of Kissin and Scott (1982). A temperature-composition diagram for the system between FeS and FeS2 from O°C to 350°C, modified after Kissin and Scott (1982), is shown in Figure 8.l0a. Phase relations in the compositional regions Fe-FeS and FeS2-S remain straightforward, essentially as shown in Figure 8.3 , but the central part of the system is exceedingly complex. This complexity is caused primarily by the crystal chemistry of the pyrrhotites. The name " pyrrho tite " is given to a group of mineral s with compositions

382

Suijide Ore Mineral Srobilities, Morphologies, ond Inlergrowth Textures

Section B

Rotation Axis

Cut perpendicular to Rotation Axis

Cut parallel to Rotation Axis

B

FIGURE 8.9

Diogrammotic presenrotion of the spiral (lower left) ond concentric (lower right) inclusion potterns observed

a0 cubic crysrol thot hos grown during rorotion is cut perpendicular or parollel, respectively, to the rorotian axis.

between FeS and Fe7Sg. All have structures based on the nickel arsenidetype in which metals occur in octahedral coordination and anions in trigonal prismatic coordination. Layers of metals and anions occur parallel to the basal plane. An important feature of this structure is an ability to omit metal atoms, leaving holes or vacancies. In the case of the pyrrhotites, approximately one-eighth of the iron atoms can be omitted, giving rise to an omission solid solution. At higher temperatures (~ 350°C), the vacancies are ran domly distributed throughout the structure (disordered), and solid solution is complete. At low temperatures, ordering of vacancies occurs, resulting in the development of various superstructures based on the nickel arsenide structure. The best known of these superstructures is that of monoclinic pyrrhotite (Fe7Sg). In this case, a superstructure is found in which the vacancies occur in alternate layers of iron atoms parallel to the basal plane and in alternate rows in those layers (Figure 8.10b). This is known as the 4C structure, because the superstructure involves a cell that is four times the c dimension of the parent nickel arsenide-type cell. The monoclinic pyrrhotite structure, first sug-

The Iron-Sulfur System

,

,

,

,

,

383

I

,

HEXAGONAL PYRRHOTITE (ICI + PYRITE

-

300

"HEXAGONAL·

PYRR~OTlTE

(MC) + PYRITE

("HEXAGONAL" PYRRHOTITE (NAI + MONOCLINIC PYRRHOTITE)

.

_("HEXAGONAL" PYRRHOTITE (NC) + MONOCLINIC PYRRHOTITE)

MONOCLINIC PYRRHOTITE

+ PYRITE

-

-

I

SMYTHITE

'I

-

+ PYRITE

I

,

, w

.... ....x>:::E

'"

44

.T W ....

" 2

W a:

, 40

42

Atomic % Fe

,

, 38

36

,

, 34

w ....

iE >-

Q.

(a)

FIGURE 8.10 (0) Phose relations in the cenflol portion of the Fe-S system below 350°C, based on Kissin and Scott (1982) and references therein. (b) Diagrams to show the ordered distribution of vacancies in three pyrrhotites (j) the 4C sflucture of Fe,S8, (iil the 5C sflucture of Fe9SIO, and (iii) the 6C sflucture of Fell S12. On~ cation layers (parallel to basal planes) ore shown. (c) Schematic, isothermal free energy-{omposition diagram for the iron sulfide minerals. Here the gain in free energy involved in forming the vorious sulfides from the elements in standord states (6 Gil in kilojoules per mole (of Fe) is plotted against composition. Stable phases fall on the dashed line (adopted from Vaughan and Lennie, 1991). (d) Revised phase diagram for the region FeS to Fe,SS (adopted from Griinvold and Stolen, 1992). Phase designations ore used in this diagram: alpha-iron (Fe); Iei , xS (10; intermediate temperature FeS (2A, 0; unknown superstructures for Fe] _xS (NA, NO; floilite FeS (3A, 20; Fell SI2 (60; FelOSll (110; Fe9SlO (50; noncommensurate phose FeO.89S (Ml; Fe,S8 (40; pyrite FeS2 (py). Solid lines ore macroscopic phase boundories; lines broken by dots ore suggested extensions of phase boundories. Dashed lines ore tentative phose boundories for Fe9SIO and Fe,SS; dotted lines ore possible existence ranges for 6C and 11 C. Crosses separated by two dashes show the antiferro- to poromognetic flansitions. Oats separated by two dashes indicate the spin-flip flansition.

gested by Bertaut (1956), has been confirmed by detailed Xray, spectroscopic, and electron microscope studies (e.g., Tokonami et aI., 1972). Monoclinic pyrrhotite is difficult to distinguish from other pyrrhotites in hand samples or microscopically but is readily identified because it is magn etic and because Xray diffraction reveals two peaks in the vicinity of the single sharp 102 reflection of hexag~nal pyrrhotite (Figure 8.11).

384

Sulfide Ore Mineral Stobilifies, Morphologies, ond Intergrowth Textures

5c

Fe9SlO (b)

6Gi 4c

2c

FeS

50

49

48

47

46

Alo mic % Fe

(e)

FIGURE8.1 0 ((ontinued)

Although the structure of monoclinic pyrrhotite is well established, those compositions lying between FeS and Fe7SS remain uncertain. Numerous superstructure types have been reported for such intermediate pyrrhotites, including examples with nonintegral multiples of the parent cell c and a dimensions (so-called NC and NA types). Many reported superstructures

The

Iron-Su~ur

System

385

Atom ic Percent Fe

49

50

700

46

o IC

+

600

IC + Py

Solid Solut ion

Q)

u,

~

47

48

-+--+--+--+--+--+--+--+--+-~==~~==~~~

:\

eli

I

:;

--tf"- - - ---.,I

o

iii lii

0.500 E