The deposition of gold and other metals: Pressure-induced fluid immiscibility and associated stable isotope signatures

0016.7037/91/$3

Geochmuca et Counochimica Acro Vol. 55, pp. 2417-2434 Copyright 0 1991 Pergamon Press pk. Printed in U.S.A.

00 + .oO

The deposition of gold and other metals: Pressure-induced fluid immiscibility and associated stable isotope signatures TERESASUTER BOWERS* Department of Earth, Atmospheric, and Planetary Sciences, Massachusetts Institute of Technology, Cambridge, MA 02 139,USA (Received November 8. 1990; accepted in revi.sed./tirm June 11. 199 I )

Abstract-It is often proposed that the onset of fluid immiscibility may be a direct cause of ore deposition, with evidence cited from fluid inclusions associated with the ore minerals. A reaction-path model is used here to calculate fluid speciation, including pH, and mineral precipitation or dissolution as a function of volatile loss from a fluid undergoing phase separation. A modified Redlich-Kwong equation of state is used to model the H20-CO2 miscibility gap and calculate activity coefficients of Hz0 and CO*. This procedure is incorporated into the EQ3/6 chemical speciation and mass transfer computer codes to simulate the response of a mineral-fluid system to fluid immiscibility. Calculations for systems including Cu, Pb, Zn, Ag, and Au show that fluid immiscibility may induce metal deposition under a variety of conditions, but that simplistic assumptions regarding fluid response to volatile loss are not always correct because they may ignore the buffering capacity of a fluid-rock system. INTRODUCTION

Quebec (ROBERT and KELLY, 1987) and the HollingerMcIntyre mine, Ontario ( SPOONERet al., 1987). LYAKHOV and POPIVNYAK( 1978) present analyses of fluid inclusions in gold-quartz veins in the Soviet Union that exhibit variable amounts of COz and a wide range of homogenization temperatures, together with analyses of groups of inclusions with homogenization to liquid and gas phases at the same temperature. Such characteristics are not seen in the fluid inclusions present in the nonproductive zones. They suggest that such evidence is indicative of fluid immiscibility, and that the pressure drop and associated degassing of the gold-bearing solutions may have contributed to the breakdown of complexes responsible for transporting the gold, resulting in its deposition. Fluid inclusion evidence of immiscibility in the Finlandia vein is apparent during the mineralization stage where the gold and silver mineralization occurred (K.&MILL1 and OHMOTO, 1977). Several lines of evidence from this work, together with that of TSUI and HOLLAND( 1979), indicate that the ore precipitation commenced suddenly and rapidly, and ended equally suddenly, an observation which has been interpreted to implicate fluid phase separation as the primary mechanism of ore precipitation. ROBERT and KELLY ( 1987) also present fluid inclusion evidence for the Sigma Mine that is interpreted to represent trapping of two coexisting, yet immiscible, fluids and suggest again that the onset of gold deposition is contemporaneous with the onset of phase immiscibility. SPOONERet al. ( 1987), in one of the best documented studies of the relation of H20-CO2 phase separation to the deposition of gold-quartz veins suggest that the occurrence of fluid immiscibility may contribute to a number of common field observations, including the coprecipitation of gold with pyrite, other sulfides and carbonates, mineralization over large vertical extents, the oxidized features often associated with Archean gold-quartz deposits, and hydraulic fracturing and volumetric inflation. Stable isotopes have been used in addition to fluid inclusions to assess the likelihood and effects of fluid phase separation in hydrothermal systems. KAMILLI and OHMOTO

DEPOSITIONOF ORE MINERALS may be triggered by any of a number of physical or chemical changes to a fluid-rock system. In particular, hydrothermal gold has often been assumed to precipitate following a drop in temperature or pressure, an increase in oxygen fugacity, or a change in the chemical speciation of the transporting fluid. A drop in either pressure or temperature will effect a change in the chemical speciation, and if this results in a mixed volatile fluid intersecting a miscibility gap, profound changes in its aqueous speciation may occur. NADEN and SHEPHERD( 1989) have suggested that fluid phase separation of an HzO-C02-CH4-H2S fluid may be responsible for gold deposition in a variety of environments and structural levels in the Earth’s crust. For example, they suggest that Archean greenstone type gold deposits, with their attendant high-CO2 fluid inclusions, may result from phase separation of a predominantly HzO-CO2 fluid on intersecting the miscibility gap at high pressures ( l-3 kbar). Epithermal gold deposits with low-CO2 fluid inclusions may represent phase separation of H20-CO2 fluids occurring at low pressures, while black shale-hosted gold deposits may be related to immiscibility in the H20-CHI system (e.g., BOTTRELLet al., 1988). The strongest evidence for inducement of gold deposition by fluid phase separation comes from fluid inclusions that suggest the events were contemporaneous. Such evidence comes from a variety of localities and types of gold deposits, as implied by NADEN and SHEPHERD( 1989). Fluid phase immiscibility has been demonstrated for gold-quartz veins in Northern Buryatia and the Lena region in the Soviet Union (LYAKHOV and POPIVNYAK, 1978), silver-gold veins such as the Finlandia in the Colqui district of central Peru (K,+ MtLLt and OHMOTO, 1977), and Archean gold-quartz veins such as at the Sigma mine in the Abitibi greenstone belt in * Present address: Gradient Corporation, 44 Brattle Street, Cambridge, MA 02 138, USA.

2417

T. S. Bowers

2418

( 1977), in their study of the Finlandia vein system, observed a shift of a few permil in d I80 of vein quartz in the stage where fluid inclusions indicated that boiling had occurred. LARSON and TAYLOR( 1987) observed a gradient in 6 I80 in hydrothermal quartz from the Red Mountain system, San Juan Mountains, which they attributed to subsurface boiling of the fluid. TRUESDELLet al. ( 1977) present models of predicted 6 “0 fractionation that would accompany subsurface boiling and or dilution of geothermal waters, and apply their calculations to the Yellowstone geothermal system. A comprehensive study of the Ag-Pb-Zn district at Keno Hill in the Yukon Territory is presented in a recent study by LYNCH et al. ( 1990). They discuss fluid inclusion and oxygen isotope evidence of fluid phase separation leading to ore deposition. Unlike many of the systems described above, the fluid at Keno Hill was quite volatile rich, containing substantial quantities of CO2 and CH4. The authors estimate an initial ~~~~ of 0.25. They suggest that oxygen isotope measurements on quartz and siderite could be used as an exploration tool. Several standard expectations exist concerning phenomena caused by fluid phase separation and volatile loss. The reduced gases are more volatile, resulting in preferential loss of Hz, CH4, and H2S relative to SO?, SOS, and O2 ( DRUMMOND and OHMOTO, 1985). This suggests that phase separation, even in predominantly H20-CO2 fluids, will result in progressive oxidation of the fluid. Loss of CO2 and H2S may lead to an increase in pH by the mechanism proposed in the following reactions: H+ + HCOy = Hz0 + CO2 t

(I)

H+ + HS- = H2S t

(2)

where the reactions proceed to the right on volatile loss and H+ concentration decreases. Loss of CO2 is also related to carbonate deposition and can be expressed by Cat+ + 2HCO;

= CaC03 + Hz0 + CO2 t

(3)

which also proceeds to the right with progressive volatile loss. Loss of H$ has been suggested to result in gold deposition in systems where gold is transported primarily as a bisulfide complex. And yet, if this reaction is written as Au, + 2H2S = Au(HS);

+ H+ + %H2,

(4)

it can be seen that while H2S loss requires that the reaction proceed to the left with gold precipitation, Hz loss and increase in pH brought about by H2S and CO2 loss (see Eqs. 1 and 2 above) will both push the reaction to the right, suggesting increased solubility of gold. In fact, each of the expectations expressed above is simplistic and consistent with the assumption that the fluid is unaffected by any processes other than volatile loss through phase separation. Real systems are more complicated and may involve fluid-solid reactions occurring simultaneously with phase separation, including dissolution and precipitation of mineral phases. The chemical compositions and oxidation states of fluids are often buffered by interactions with wall rocks and precipitation of vein minerals. Significant loss of volatiles may overwhelm these buffers, but one may not expect a priori that a simple relationship exists between pro-

gressive volatile loss through phase separation as a function of changing temperature or pressure, and shifts in pH or oxidation state of a fluid-rock system. Such assessments require quantitative analysis of the particular physical and chemical conditions of the system. This communication attempts a step in that direction. PREVIOUS WORK A theoretical examination of fluid phase separation in low C02, low pressure hydrothermal systems and its effects on ore deposition has been conducted by DRUMMOND and OHMOTO( 1985 ) and COLE and DRUMMOND( 1986). These papers present the results of model calculations for fluid phase separation applied over a wide range of chemical and physical conditions, including open- and closed-system processes, isothermal and isenthalpic processes, and complete and partial redox equilibria between sulfur and carbon species in solution, at temperatures between I50 and 300°C. Partitioning of CO2 and HzS between liquid and vapor phases is accounted for by a Henry’s law constant ratio. The solutions are constrained initially to be at saturation with respect to several metal sulfide phases and either calcite or anhydrite. Among the conclusions of these studies are that ( 1) fluid phase separation results in precipitated assemblages with a higher ratio of ore minerals to silica than produced by other depositional mechanisms, (2) the most dramatic chemical changes to the solution are seen in the pH response to CO2 loss and in the oxygen fugacity response to H2S loss, and ( 3 ) open-system volatile loss results in greater changes to the chemistry of the system than that induced by closed-system processes. In addition, the paper by COLE and DRUMMOND ( 1986) reached several interesting conclusions about the relative transport and depositional properties of gold and silver in hydrothermal systems undergoing fluid phase separation. Conclusions by SEWARD( 1973) and others concerning the solubility of gold and the importance of sulfide over chloride as a gold-complexing agent were corroborated by this work. COLE and DRUMMOND( 1986) showed that Ag-Au ratios less than 1 in solutions are favored by temperatures below 250°C and low Cl, high H2S conditions, while high Ag-Au ratios are generally found in higher temperature solutions with high Cl and low HIS. Although the papers by DRUMMONDand OHMOTO( 1985) and COLE and DRUMMOND( 1986) have contributed significantly to the understanding of fluid immiscibility effects on mineral equilibria in hydrothermal systems, their model does have some limitations. Because of their employment of Henry’s law coefficients for liquid-vapor partitioning, the model is only applicable at xcq less than approximately 2 mol% and pressures somewhat above the water liquid-vapor coexistence curve. At 250°C they are apparently limited to pressures less than 150 bars and at higher temperatures to about 200 bars. Although epithermal deposits are often quite shallow, there are many gold deposits which have formed at significantly higher pressures. Examples include Archean goldquartz veins, which typically form at l-3 km depth (e.g., SPCXINER et al., I987), and the Carlin gold deposit in Nevada, which is estimated to have formed at pressures in excess of 1 to 2 kb ( PASTERISet al., 1986). Perhaps more important

2419

Deposition of metals by fluid immiscibility

is the restricted CO2 region of the model, because fluids in deposits forming at higher pressures such as many of the Archean goldquartz veins may have initial COZ contents as high as 30 mol% (ROBERT and KELLY, 1987). Reaction path calculations for boiling fluids have been presented by REED and SPYCHER( 1985) and SPYCHERand REED ( 1986). Their model is based on a dilute Broadlandstype water that boils and cools over a temperature interval from 278 to lOO”C, precipitating minerals as a function of reaction progress. The calculations show an increase in pH with gas loss and dropping temperature that results in precipitation of sulfides and carbonates. Silicates precipitate as well, driven by the drop in temperature. Throughout the calculations the vapor phase is maintained in equilibrium with the liquid and minerals and is not allowed to leave the system. This means that components of the vapor can back-react with the fluid, and mass transfer may be either from liquid to vapor or vapor to liquid. Several scenarios for gold deposition are presented. Although this model uses equilibrium constants for computing the composition of the gas phase rather than Henry’s law coefficients, the assumption of ideality in the gas phase requires that the accessible range of xco, be small, similar to that in the Drummond and Ohmoto model, and pressure is restricted to approximately that corresponding to the water saturation curve. Neither the Drummond and Ohmoto model nor the Reed and Spycher model include the ability to calculate stable isotope fractionation. FLUID PHASE EQUILIBRIA AND ISOTOPIC FR4CIIONATION Review of Fluid Phase Relations The binary system HzO-CO2 has a miscibility gap at high pressures and temperatures that has been delimited experimentally by MALININ ( 1959), TODHEIDE and FRANCK ( 1963 ), and TAKENOUCHIand KENNEDY ( 1964). The critical composition curve in this system extends from the critical point of HZ0 at 374°C and 22 1 bars to a temperature minimum of approximately 26O’C at about 2000 bars. Temperature then increases slightly with increasing pressure to the limit of available experimental data, suggesting that the miscibility gap may be larger at higher pressures. Perhaps the most familiar expression of the HzO-CO2 miscibility gap is that shown on a temperature-composition cross section such

0 0

I 0.2

/

,

I 0.8

1

o.4 xco,".6

FIG. 1.The miscibility gap for the binary system HzO-CO2at 500 bars (solid curve) and 1500 bars (dashed curve), as a function of temperature.

0

0.8

0.2 o.4 9

1

o.6

RG. 2. The miscibilitygap for the binary systemHZO-CO2at 200°C (solid curves) and 250°C (dashed curves), as a function of pressure.

as Fig. 1. At 1500 bars it can be seen that an HzO-rich liquid coexists with a COz-rich vapor in the center of the diagram below a temperature of approximately 265°C. At 500 bars the critical temperature is higher and the region of two coexisting fluids is larger. Phase relations can also be expressed as pressure versus composition, as shown in Fig. 2. The solid curves correspond to the miscibility gap at 200°C while the dashed curves represent 25O’C. Again, the two-fluid phase region is in the center of the diagram. The addition of ionic species such as NaCl serves to increase the size of the miscibility gap ( TAKENOUCHIand KENNEDY, 1965; BOWERSand HELGESON,1983). Although phase equilibria relations for binary mixtures are often well known, the same is not true of ternary mixtures. No comprehensive study of fluid phase relations in the system H20-C02-H$ is available, but some information for specific compositions has been determined ( HUANG et al., 1985). For the purposes of the calculations presented here it is necessary to assess the proportion of H2S that will partition into the vapor phase that arises from intersection of the fluid with the H20-CO2 miscibility gap. DRUMMOND and OHMOTO ( 1985 ) present volatility ratios for H2S, CO*, and other gases as a function of temperature at pressures corresponding to the water saturation curve. They define volatility ratio as the molality of the species in the vapor phase divided by its molality in the coexisting liquid phase. No information concerning volatility ratios for H2S is available at higher pressures; however, for COZ the work of TODHEIDEand FRANCK( 1963) and TAKENOUCHIand KENNEDY ( 1964) can be recast as volatility ratios as a function of both temperature and pressure. The assumption can then be made that the dependence of volatility ratio on pressure is the same for CO2 and H2S, and in this way the amount of H2S to be partitioned into the CO&h vapor phase can be calculated. At 250°C the ratio of the volatility ratio of H$ to that of CO2 is 0.14 (DRUMMONDand OHMOTO, 1985). This reflects the fact that H2S is less strongly partitioned into a vapor phase than is COZ. Under reducing conditions CH4 becomes a dominant fluid component rather than CO*. Methane and water exhibit a miscibility gap that is quite similar to that seen in the H20CO2 system, although with a higher critical temperature, somewhat in excess of 350°C between 500 and 2000 bars ( WELSCH, 1973). Calculations presented below are restricted to conditions such that CH4 is not a major fluid component.

2420

T. S. Bowers

Isotopic Fractionation Equilibrium fractionation factors for ‘sO/‘6O, D/H, and 34S/32Shave been experimentally or theoretically determined as a function of temperature for a variety of mineral pairs and mineral-solution systems. Fractionation factors between two phases A and B are commonly expressed by cy,defined as 1000 + SA CYA-B = 1000 i- Ijg where &, and SSare the isotopic signatures of the two phases with respect to a common standard in units of permil. When (Yis close to 1 the following approximate relationship holds: r)~ - 68 = lo3 ln &&_B.

(6)

Summaries of fractionation factors are available in FRIEDMAN and O’NEIL ( 1977), BOWERSand TAYLOR ( 1985), KYSER (1987), and BOWERS(1989). The application of these fractionation factors to systems of minerals coexisting with mixed volatile fluids requires some assumptions. RUMBLE( 1982) discusses the control of isotope fractionation by fluid composition. He presents an equation to relate the 6 I80 signature of a mineral (M) to the 6 I80 of a mixed volatile fluid (FL):

where xI is the atomic fraction of oxygen in volatile species f relative to the total amount of oxygen contained in the mixed volatile fluid. For a binary fluid mixture of Hz0 and C02, Eqn. (7) will reduce to

=

xH20

lo3 In

%-Ha0 xH,O

+ +

2xCo2to3 In ffM_C02 232%

($1

where xHZoand xcoZ are now simple mole fractions and sum to 1. Note that while fractionation factors for carbonates in isotopic equilibrium with CO* have been determined, those for silicates in equilibrium with CO:! are not known. If the proportion of CO2 in the fluid is small, Eqn. (8) will reduce to 6180~ - 6’*0~~ = lo3 In &-~~o.

(9)

The value of 6 IgOm always depends on XC%, so there will be a change in 6180,,, in equilibrium with the fluid if there is a change in the proportions of Hz0 and COZ in the fluid. Implicit in Eqn. (9) is that the difference between 6 “0 of the mineral phase and 6 I80 of the fluid remains constant at a given temperature, and does not change with changing concentration of CO2 in the fluid. Equation (9) has been used for all oxygen isotope fractionation calculations in this study. Such calculations have been carried out for a fluid containing 15 mol% CO,, where the alternative use of Eqns. (8) or (9) will result in different calculated values for 6 180h,. Equation (9) was used primarily because of the paucity of information on fm~tionation between minerals and CO,. Note that alternative assumptions concerning the distribution of oxygen isotopes between dissolved species in solution and the relationship of oxygen in pr~pi~ting minerals to such solution

species can lead to formulations other than either Eqn. (8) or (9 ) . it must be emphasized that both Eqns. (8 ) and (9 ) involve simplistic assumptions, and that the validity of any of the alternative assumptions suggested above can only be ascertained by further experimental work. Isotopic Fractionation across the HzO-CO2 Miscibili~ Gap The fmctionations of stable isotopes between coexisting fluid phases on either side of the H20-CO2 miscibility gap have not been measured experimentally and are calculated in this study based on a set of assumptions. The fractionation of I80 between HZ0 liquid and Hz0 vapor and between HZ0 Iiquid and CO, vapor is known, and several determinations are summarized by FRIEDMANand O’NEIL ( 1977). The calculations summarized here employ the curve for H20 liquidHZ0 vapor from B~TTINGA and CRAIG f 1968) and the curve for Hz0 liquid-CO2 vapor from TRUESDELL( 1974). All calculations begin with one fluid phase, a liquid, present with an assigned S’80. With dropping pressure or temperature, a liquid of specified composition intersects the H20-CO2 miscibility gap and a vapor phase of composition corresponding to the vapor-rich limb of the mi~ibility gap forms. Relative proportions of the liquid and vapor phases can be calculated from the lever rule and depend on the size of each pressure or temperature step, that is, how far the solvus has been overstepped. It is then assumed that the CO2 portion of the vapor phase is in isotopic equilibrium with the liquid, and d “OFe can be described by the fractionation factor for HZ0 liquid-CO* vapor. The Hz0 portion of the vapor phase is also assumed to be in isotopic equilibrium with the liquid, and 6 ‘80Z$+?can be described by the fractionation factor for H20 liquid-H20 vapor. A mass balance on the isotopes is performed and the bulk 6 “0 of the vapor phase is determined from g”Ovapor = 2xco$‘80~~

+ &,H206’80~Z~.

(10)

Fractionation calculations of 6 “0 between coexisting liquid and vapor phases on either side of the HzO-CO2 miscibility gap as a function of temperature at a constant pressure of 500 and 1500 bars (Fig. 1) , and as a function of pressure at a constant temperature of 200 and 250°C (Fig. 2) are presented in Figs. 3-6. These calculations are for an open system, where the vapor is removed at each step. In each calculation the initial 6”O of the fluid is +6%0. The calculation shown in Figs. 3a and 4a corresponds to a fluid composition with initial xco, set at 0. t25; this results in an intersection with the 500 bar solvus at approximately 300°C. Initial conditions for the calculation shown in Figs. 3b and 4b are xc% = 0.15 with intersection of the 1500bar miscibility gap at approximately 250°C. Figure 3 shows S”O for the liquid and vapor phases as a function of temperature, while Fig. 4 shows 6180 from the same calculations plotted as a function of XC* in the liquid phase. Figures 5 and 6 show comparable calculations at constant temperature. The initial fluid composition is XCO,= 0.075 for Figs. 5a and 6a with the miscibility gap at 2OO”C, and XC- = 0.15 for Figs. 5b and 6b with the miscibility gap at 250°C. In both examples the miscibility gap is intersected at approximately 1500 bars. In each of Figs. 3-6 it can be seen that S’*O of the fluid

2421

Deposition of metals by fluid immiscibility

40

0

-I VAPOR

3o

/

/

+,_-”

?T

20 -

300

/

--

250

:'

200 150 100 TEMPERATURE,'?

50

0

2000

1500

1000 PRESSURE,BARS

500

0

FIG. 3. Calculated 6 ‘8O of coexisting H1O-rich liquid and COzrich vapor phases as a function of temperature (which varies with xc% as shown in Fig. 1) for (a) 500 bars and (b) 1500 bars.

FIG. 5. Calculated 6 ‘*O of coexisting HzO-rich liquid and CO>rich vapor phases as a function of pressure (which varies with xco, as shown in Fig. 2) for (a) 200°C and (b) 250°C.

becomes lighter as a function of progressive vapor phase separation. This is in contrast to boiling of pure water where the lighter isotope partitions into the vapor and the liquid becomes progressively heavier. Although water vapor is isotopically lighter than liquid water, CO2 gas is isotopically heavier

than liquid water. The direction and magnitude of change in the isotopic composition of a mixed volatile fluid from which a vapor phase is separating will depend on the fluid composition, while the magnitude of change will also depend on changing temperature and pressure conditions. This

I

50

I

25

I

a

40 -

I --_

20 -

i

I

1 a

v;p,x

- - -

0'5 'D m 10 -

VAPOR

5

t 50

I

/

,

LKXJID

-

0

< I

I

I

1

b 40 -

I I'

3o % 0 20 -

VAPOR

iJ

___----

10 -

LlQUlD LCUUI

0

I

0.15

0.12

I

I

0.09 0.06 XCO, inthe LIOUID

-t

I

0.03

0

FIG. 4. Calculated 6180 of coexisting HZO-rich liquid and C02rich vapor phases as a function of xc% in the liquid phase (which varies with temperature as shown in Fig. 1) for (a) 500 bars and (b) 1500 bars.

I

0.2

0.15

I

I

0.1 0.05 XCO, in theLIOUID

I 0

FIG. 6. Calculated 6’*0 of coexisting H*O-rich liquid and COzrich vapor phases as a function of xc% in the liquid phase (which varies with pressure as shown in Fig. 2) for (a) 200°C and(b) 250°C.

T. S. Bowers

2422

reverse oxygen isotope shift has also been remarked upon by TRUESDELL et al. ( 1977) and LYNCH et al. ( 1990). When a fluid with xco, = 0.125 and 6 I80 = +6%0 intersects the 500 bar miscibility gap (Fig. 3a) the vapor phase that initially separates will have 6 I80 approximately equal to + 14%0.After cooling to lOO“C, 6”O of the liquid will equal +3.0%0 and 6 “0 of the vapor +32. Figure 3b shows that for intersection of a fluid with xco2 = 0.15 with the miscibility gap at 1500 bars the effect on the liquid is more dramatic, with 6”O at 100°C = +2.0%0, while the vapor varies from + 17 to +31%0. Figure 4 shows results of the same calculations plotted as a function of xco, in the liquid (corresponding to the liquid side of the miscibility gap). 6 ‘*O of the vapor phase changes most rapidly at low xco2 where xco, changes slowly as a function of temperature. Much of the change in isotopic signatures of the liquid and vapor phases shown in Figs. 3 and 4 is a result of the strong dependence of isotopic fractionation on temperature. Figures 5 and 6 show the analogous calculations at constant temperature where the miscibility gap is expressed as a function of pressure. Here all change in d ‘*O values of both liquid and vapor may be attributed to phase separation because temperature is not a variable. Both liquid and vapor 6”O values decrease with dropping pressure: at 200°C liquid from +6 to +4.2k and vapor from +23 to +21%0 (Fig. 5a), and at 250°C liquid from +6 to +2.2%0 and vapor from +18 to +14%0. The greater change at 25O’C may be attributed to the higher initial content of CO1 in the fluid: 0.15 compared to 0.075. Figure 6 is comparable to Fig. 4 and shows the results plotted as a function of xco, on the liquid side of the miscibility gap. These are a few sample calculations to show the effects of temperature, pressure, and fluid composition on fractionation of 6”O across the HzO-CO2 miscibility gap. Under some conditions 6 “0 of the liquid phase changes by as much as 4%0, and this change may be observed at constant (Fig. 5b) as well as at changing (Fig. 3b) temperature. Although CH4 and Hz are not considered in this study because the calculated models are constrained by oxygen fugacities consistent with the predominance of COz, some comment about the effects of H20-CH4 immiscibility on hydrogen isotope fractionation is warranted. At 250°C CH4 gas is approximately 7OL lighter than liquid water. Hz gas is approximately 6000/wlighter than liquid water. It is apparent from these values that an H20-CH., system undergoing phase separation will result in a liquid phase with a 6D signature that becomes progressively heavier, perhaps by several permil. THE MODEL

The basic computer model used in the following calculations is the EQ3/6 computer software package developed by WOLERY(1978, 1979, 1983). EQ3 computes a distribution of species in the aqueous solution and generates concentrations and activities of ions and complexes; it then calculates which, if any, minerals are saturated with respect to the solution. EQ6 calculates chemical equilibrium and mass transfer in aqueous solution-mineral systems. The codes were extended by BOWERSand TAYLOR( 1985 ) and BOWERS( 1989) to include fractionation of oxygen, hydrogen, and sulfur iso-

topes between the fluid and minerals. Isotopic exchange is assumed to take place only through the processes of solution and deposition; diffusional mechanisms have not been included. Conservation of mass of the isotopes is carried out mathematically in EQ6 in much the same way as is conservation of the elements. The EQ3 / 6 computer programs were further modified for the calculations presented here to incorporate a modified Redlich-Kwong (MRK) equation of state for HzO-CO2 fluids (BOWERSand HELGESON, 1983) to describe activity coefficients for Hz0 and COz in the liquid phase, and fluid unmixing with vapor loss consistent with the H20-CO2 miscibility gap. Although BOWERSand HELGESON( 1983) report a lower temperature limit of 350°C for the use of their MRK equation of state for H20-C02-NaCl fluids, the equation is valid to temperatures as low as 1OO“Con the HrO-CO2 binary (see their Fig. 22). The MRK equations generate fugacity coefficients or fugacities for Hz0 and CO1 that are consistent with a gas standard state ( 1 bar and any temperature), while liquid water and CO* in the aqueous phase are handled in EQ3/6 with a liquid standard state (any pressure and any temperature). For this reason a conversion is necessary from the fugacity coefficients calculated from the MRK equations to an activity of Hz0 and an activity coefficient for CO* consistent with EQ3 / 6 usage. This can be done by writing a reaction between liquid water and steam: HIOiiquid = H~Osteam

(11)

for which a log K is available in the EQ3 / 6 database or can be calculated from SUPCRT ( HELGESONet al., 1978), where log K = log&am - log awater

(12)

where f is fugacity which is calculated from the MRK equations, and a represents the activity required in the EQ3/6 programs. Activity of water is therefore calculated from

1%aH20= kf&m

- log K.

(13)

Likewise, for CO*,

(15) where f is again the fugacity calculated from the MRK equations and a is activity, which is equal to y * m where y is the activity coefficient and m is molality. EQ3 /6 requires a value for the activity coefficient: log

YCO~ =

log fc~~,,,

-

log

K

-

log

m.

(16)

Modifications to EQ3/6 require a call to a Redlich-Kwong subroutine which returns values of fugacities for Hz0 and CO2 for water-rich compositions up to those corresponding to the water-rich limb of the miscibility gap. Further subroutines and procedures that use the activity of either water or carbon dioxide then substitute values calculated from the MRK equations. The calculations presented here are all made for a fluid evolving in pressure at a constant temperature of 250°C. The pressure at which the HzO-COz gap is intersected is dependent on initial fluid composition (Fig. 2). The mole fraction of

Deposition of metals by fluid immiscibility CO2 corresponding to the water-rich and vapor-rich limbs of the miscibility gap were fit to a polynomial in pressure. Calculations are begun at a pressure such that the fluid is in the one-phase region, and pressure decreases as a function of reaction progress. At each pressure step the code calculates the x,-o2 position of the miscibility gap and compares it to the xco, content of the fluid. When the calculation indicates that the water-rich limb of the miscibility gap has been intersected or overstepped, a packet of CO-rich vapor of composition consistent with the vapor-rich limb of the miscibility gap at that pressure is subtracted from the system. This portion of the calculation is consistent with an open system; that is, when this packet of vapor leaves the system it is not held in equilibrium with the fluid over future steps and it is not available for further reaction. Thus, there is a net change to the mass balance of the elements 0, H, C, and S, as well as to the isotopes of 0, H, and S. EQ3/6 requires a separate database to describe the pressure-temperature path along which calculations are performed. The database distributed with EQ3/6 ranged from 0 to 300°C at pressures corresponding to the liquid-vapor curve for water. The version of EQ3 /6 used here accesses a pressure database generated at 250°C and covering a pressure range from that corresponding to the saturation curve for pure water to 3 kbars. Although the database is isothermal, it is approximately equivalent to an adiabat for pure water. This database was generated in the same manner as are constant pressure databases used by BOWERSand TAYLOR( 1985 ) and JANECKYand SEYFRIED( 1984). Log K values for hydrolysis reactions of the minerals were calculated as a function of pressure at 25O’C using the computer program SUPCRT ( HELGESONet al., 1978 ) . Log K values for reactions between aqueous complexes and ions were generated from the saturation pressure database by applying the isodielectric correction suggested by HELGESON( 1969). Log K values for all reactions between aqueous species at 250°C at a pressure equal to that of the saturation curve for pure water, together with the source of the data, are given in Appendix 1. Supporting thermodynamic data are found in HELGESONet al. (1978), ROBIE et al. (1978), WOLERY (1978), HELGESON ( 1969), and in the databases provided with the EQ3 /6 code. Databases containing isotopic fractionation factors for oxygen, hydrogen, and sulfur are those used by BOWERS( 1989). A large number of variables must be specified at the outset of a calculation. These include temperature, Con content, sulfur and oxygen fugacity, pH, concentrations of cations and anions in solution, oxygen, hydrogen, and sulfur isotopes of the fluid, and whether the sulfide and sulfate are chemically and isotopically in equilibrium or disequilibrium. Temperature is currently constrained by the database to be 25O’C: however, further database development would remove this constraint. The user is free to select each of the other variables. Obviously, large numbers of possible combinations of these variables could be investigated, only a few of which are presented in this preliminary study. In each of the calculations presented below, xco2 equals 0.15 ( xco, + xu20 = 1); however, the code is not restricted to this value. Examples of COz contents in this range have been cited in the Keno Hill AgPb-Zn district in the Yukon (xcoz = 0.25; LYNCHet al., 1990) and in Archean goldquartz veins at Sigma Mine, Quebec

2423

( xc02 = 0.15-0.30; ROBERT and KELLY, 1987). BODNAR ( 1986) reports a typical range of xco, from 0.10 to 0.20 in silver-bearing skam deposits. The water-rich limb of the H@COz miscibility gap is intersected at xco, = 0.15 and 25O’C at approximately 1500 bars. Two sets of calculations are presented: those where initial input includes the elements H, 0, S, C, Cl, Na, K, Ca, Mg, Al, Si, Fe, Zn, Pb, Ag, Cu, and Au; and those where the input elements are restricted to H, 0, S, C, Cl, Na, K, Ca, Si, Fe, and Au. The latter set approximates Al-poor fluids that deposit quartz rather than alumina-silicates, and precludes Cu, Pb, Zn, and Ag in an attempt to better define conditions favoring gold deposition. Oxygen fugacity either corresponds to the coexistence of magnetite and hematite (mt-hm; a log value of approximately -35.3) or has a log value of -33, which is in the hematite stability field. Sulfur fugacity is set equal to a log value of -9.6 (approximately equal to the coexistence of enargite and tennantite), -6.8 (coexistence of chalcocite and covellite), or - 11. The coexistence of pyrite and pyrrhotite is at an approximate log sulfur fugacity value of - 14. Calculations run at this value are not included here because the onset of fluid immiscibility had little apparent effect on mineral precipitation patterns. Some of the values of sulfur and oxygen fugacity used in this study border on those shown in HEALDet al. ( 1987) in their Fig. 5 for likely environments of formation of major types of epithermal mineral deposits at 250°C. Ionic strength is varied from 0.5 to 0.1 m. The concentration of K is set at 10% of the concentration of Na, and in several calculations the concentrations of Na + K sum to that of Cl. The value of pH is specified by a charge balance over all ions in solution, and variations in pH are obtained by allowing the concentrations of Na + K to diverge from Cl. The concentrations of the remaining major cations are constrained by specifying them to be in equilibrium with a suite of stable minerals. These include quartz (Si), muscovite, Ca-beidellite or Kfeldspar (Al), calcite or anhydrite (Ca) , dolomite (Mg) , pyrite or siderite (Fe), chalcopyrite or bomite (Cu), galena (Pb), sphalerite (Zn), argentite (Ag), and gold (Au). The constraints on solution composition for all of the calculated models are set at 1500 bars and 250°C with pH calculated by charge balance from the selected variables. Initial constraints of all models are listed in Table 1. Initial isotopic ratios of the fluid are set, arbitrarily, at 6 “0 = -2%~ and 634S = O%O.Sulfate and sulfide are assumed to be both compositionally and isotopically in equilibrium and to remain in equilibrium throughout the calculation. This assumption is supported by laboratory experiments on seawater interaction with magnetite and fayalite by SHANKSet al. ( 198 1) who show that sulfate reduction is rapid at temperatures in excess of 250°C in acidic solutions and that isotopic equilibrium is attained between coexisting sulfide and sulfate in solution at 250 and 350°C. Contradicting this is work by OHMOTO and LASAGA( 1982) suggesting that the temperature threshold for attainment of sulfide-sulfate equilibrium in natural systems is higher, possibly 350°C. Sulfide and sulfate should probably not be considered to reach equilibrium at temperatures below 250°C. The calculations are initiated at 2000 bars. This is accomplished by a redistribution of species from 1500 to 2000 bars witb no change in bulk composition of the fluid. Nevertheless,

T. S. Bowers

2424 Table 1. Initial Constra~ts Log fo, Logfs, -9.6 Model A -35.3 (en-tn) (mt-hm)

for Model Calculations Cl pH* Na + K 0.5 4.63 0.5

at 1500 bars and 250°C Minerals Chalcopyrite Quartz Galena Muscovite Sphalerite Calcite Argentite Dolomite Pyrite Gold

Mode1 B

-35.3 (mt-hm)

-6.8 (cc-w)

0.5

0.5

4.63

Quartz Muscovite Calcite Dolomite Pyrite

Bornite Gabna Sphalerite Argentite Gold

Model C

-35.3 (mt-hm)

-9.6

0.1

0.1

4.54

(en-tn)

Quartz Ca-beideilite Calcite Dolomite Pyrite

Chakopyrite Gdena Sphalerite Argentite Gold

-35.3 (mt-hm)

-9.6 (en-tn)

0.5

0.05

5.93

Quartz K-Feldspar Calcite Ddomite Pyrite

Chalcopyrite Galens Sphalerite Argentite Gold

Model D

Model E

-33

-9.6 (en-tn)

0.5

0.5

4.46

Quartz Muscovite Anhydrite Doiomite Pyrite

Chalcopyrite Galena Sphalerite Argentite Gold

Model F

-33

-9.6 (e&n)

0.1

0.02

4.98

Quartz Anhydrite

Pyrite Gold

Model G

-33

-63 (cc-w)

0.1

0.02

4.35

Quartz Anhydrite

Pyrite Gold

Model H

-33

-11

0.1

0.02

5.18

Quartz Calcite

Siderite Gold

-9.6 0.02 5.19 Quartz 0.1 (en-tn Calcite * Initial pB is constrained by the other selected variables and calculated from them by a charge balance Model .I

-33.5

pH calculated from charge balance at 1500 and at 2000 bars may differ; and, as a result, the following diagrams show an initial pH value at 2000 bars different from that listed in Table 1 at 1500 bars. The redistribution of species at 2000 bars occasionally resulted in a supersaturation of one or more mineral phases, usually gold or pyrite. When this occurred the concentration of the relevant cation (Au or Fe) was reduced until the solution was just at saturation at 2000 bars. RESULTS

Results of nine model ~cula~ons, labeled A-H and J, are presented. Model A was calculated for the full suite of elements (H, 0, S, C, Cl, Na, K, Ca, Mg, Al, Sk Fe, Zn, Pb, Ag, Cu, Au) with a log oxygen fugacity of -35.3 (approximately the magnetite-hematite buffer), a log sulfur fugacity of -9.6 (approximately the enargite-tennantite buffer), and an ionic strength of 0.5 m. The resulting pH calculated from charge balance at 1500 bars was 4.63. Results of the calculations are presented in Figs. 7- 10. Figure 7a shows log moles of precipitates that are calculated to form as a function of reaction progress (equivalent to decreasing pressure). The

Pyrite Gold

calculation is initiated at 2000 bars and the H20-COz miscibility gap is intersected at approximately 1500 bars. This is also the pressure at which several minerals are constrained to be at saturation. As the miscibility gap is intersected and volatiles begin to leave the system, precipitation of dolomite, galena, argentite, sphalerite, muscovite, and calcite occur, followed by chakopyrite at approximate 1470 bars. Each of these minerals continues to precipitate as pressure drops and volatiles leave the system. In this diagram, as well as all following diagrams of calculated precipitates versus pressure, the sulfides are depicted with solid curves, while the nonsulfides are shown with dashed curves. Figure 7b shows the cumulative percent of total moles of H20, C02, and H2S that are removed as vapor with decreasing pressure. When the calculation terminates at approximately 200 bars, over 80% of the COr has left the system, along with 22% of the total sulfur and 6% of the H20. Figure 8a shows the calculated pH and oxygen fugacity as a function of pressure. The pH initially increases when pressure is between 2000 and 1500 bars, a trend that can be attributed to a redistribution of aqueous species over this pressure interval. When the miscibility gap is intersected and volatile loss occurs together

2425

Deposition of metals by fluid immiscibility

r--

a

MODEL A

60

iY

I

-

? 0

$

60 -

5

40 5 a 20 y

1500

2000

500

1000 PRESSURE, BARS

0 2m:

FIG. 7. Reaction path model at 250°C calculated as a function of pressure for initial conditions of log oxygen fugacity = -35.3, log sulfur fugacity = -9.4, and an ionic strength of 0.5 m with concentrations of Na + K = Cl. Results of this model are shown in Figs. 7- 10. (a) Calculated log moles of precipitates. Solid curves represent sulfide minerals and dashed curves are non-sulfide minerals. (b) The cumulative amount of HzO, CO*, and HrS lost from the system as a result of vapor phase separation. 2000 5.25

I

I

’

I

I

al - -35.2

5.00 -.

- -35.4

4.75 I,

-35.0

4.50

-

\ 4.25

\

\

-35.6

Q”

- -35.6

2

\

1500

1000 PRESSURE,

500

500

0

BARS

FIG. 9. Initial conditions are as described in the caption of Fig. 7. Log concentration of metal complexes in the fluid phase: (a) Fe and Cu complexes; (b) Ag and Pb complexes; (c) Zn and Au complexes. Only complexes with a log concentration greater than - 10 have been plotted.

A

-36.0

2000

1000 PRESSURE,

8

\

4.00

1500

0

BARS

FIG. 8. Initial conditions are as described in the caption of Fig. 7. (a) Calculated pH (solid curve) and log fugacity of oxygen (dashed curve). Note the slight change in slope in the pH curve at 1500 bars where initial fluid phase separation occurs. (b) Log concentration of carbonate, sulfide, and sulfate species in solution.

with mineral precipitation, a discontinuity is observed in the pH curve. It continues to increase with dropping pressure, although with a slightly lower slope, which gradually steepens again as pressure continues to drop. Oxygen fugacity decreases slightly over the course of the calculation and no response can be seen in its value to the initiation of fluid phase immiscibility. Figure 8b shows speciation of several carbon- and sulfur-bearing complexes in solution. (Carbonate and sulfate complexes of Na, Ca, Mg, and other cations are included in the calculation but not shown in this figure.) The major species show no change in concentration between 2000 and 1500 bars, although some of the minor species do change over this pressure interval as a result of species redistribution. No net change in the mass balance of carbon or sulfur occurs between 2000 and 1500 bars as the system remains closed over this pressure interval. At the point that the miscibility gap is intersected at 1500 bars most of the curves show a discontinuity as CO*, H& and Hz0 begin to leave the system. These trends are not all in the negative direction, however, as the concentration of HS- increases with progressive volatile loss. This

T. S. Bowers

2426

-10 2

/ ’

I

I

I

I

1

I

I I

b

consequence of increased sulfur concentration in a fluid that is constrained to be at equilibrium with respect to specified metal-sulfide phases is that less of the metal will be transported. This results in the deposition of less galena and argentite. Model C, shown in Fig. 11b, has a lower ionic strength than Model A (0.1 m compared to 0.5 m) . One consequence of this is that pyrite is among the precipitates in Model C, although not in Models A or B. Ca-beidellite is stable over muscovite in the early part of the calculation. Somewhat lower amounts of each of the minerals precipitate as compared to Model A; this may be attributed to the transport of fewer cations in solution. Model D (Fig. 11c) has an initial concentration of Na + K equal to 0.5 m, the same as in Model A, but the concentration of Cl has been reduced by a factor

0

3 t

-2-

5

-4-

E LL

-6m

3

-&

!2

-10

t

-6

/ 2000

1

1500

I

1000 PRESSURE, BARS

I I

500

0

0

I

/

M30EL B

I

M#~~VITE

a -

_~z~~~z~y~~~ /:-

_

WLCM,TE

_

_ _ _ _ SPHALEAITE

ARGENTlTE GALENA SORNITE g

FIG. 10. Initial conditions are as described in the caption of Fig. 7. (a) d ‘*O of coexisting minerals and fluid, also d “0 of the separating vapor phase. Note that the overall change for the pressure interval plotted is approximately 4%0for the fluid as well as the minerals. (b) h3?S of coexisting minerals and fluid. The values of the fluid and minerals approach each other as the fluid becomes more reduced with decreasing pressure (Fig. 8). The total change in 634Sof the

-

-12

,

8 i 3

I

minerals is on the order of 2L, while change to the fluid is negligible.

is linked to the increase in pH and the decrease in oxygen fugacity; note the drop in concentration of the sulfate species. Figure 9a-c shows concentrations of all major aqueous species of Fe, Cu, Ag, Pb, Zn, and Au. (Those with log concentrations less than -10 are not included in the diagrams, although present in the calculations.) The concentrations of several of the species decrease from 1500 bars with dropping pressure as sulfides are precipitated. At 200 bars approximately 94% of the Fe, 95% of the Pb, 92% of the Zn, 36% of the Ag, and 5% of the Cu of that originally carried in the fluid has been deposited in the form of sulfide minerals. Isotopic signatures of the minerals and fluid are shown in Fig. 10, oxygen in Fig. lOa, and sulfur in Fig. lob. b “0 of both fluid and minerals decreases by approximately 4%0over the pressure interval of the calculation. Little change is seen in 6 34Sof the fluid, however 634S of the precipitating sulfide minerals increases by about 2% over the pressure interval. The effects of alternative constraints on sulfur fugacity, ionic strength, and pH can be seen in Models B, C, and D shown in Fig. 11. This figure shows log moles of precipitates as a function of pressure and can be compared to Fig. 7a for Model A. These models differ from Model A by alteration of only one initial constraint in each case. Model B is constrained by a higher sulfur fugacity, a log value of -6.8 which corresponds to the chalcocite-covellite buffer. The order and amounts of minerals that form in Models A and B are quite similar, but Model B has less argentite and galena, and bomite forms in place of chalcopyrite, later joined by chalcocite. A

2 t

-2-

;

-4-

B B

-6

ul u

-0-

CHALCOPYRITE

/

d

i-l0

-12

I 2000

1500

/ 1000 PRESSURE, BARS

500

0

FIG. Il. Reaction path calculations at 250°C as a function of pressure similar to that shown in Fig. 7 for varied initial conditions. Solid curves represent sulfide minerals; dashed curves are non-sulfide minerals. (a) Effect of increased sulfur over that shown in Fig. 7. Log oxygen fugacity = -35.3, log sulfur fugacity = -6.8, ionic strength = 0.5 m with concentrations of Na + K = Cl. (b) Effect of decreased ionic strength from that shown in Fig. 7. Log oxygen fugacity = -35.3, log sulfur fugacity = -9.6, ionic strength = 0.1 m with concentrations of Na + K = Cl. CHALCOPY = chalcopyrite. (c) Effect of higher initial pH over that shown in Fig. 7. Log oxygen fugacity = -35.3, log sulfur fugacity = -9.6, Na + K = 0.5 m, Cl = 0.05 m. The lower concentration ofCl requires a higher pH, everything else being equal, to balance the charge.

2427

Deposition of metals by fluid immiscibility

of 10, to 0.05 m. This requires a higher pH to balance the charge, 5.93 in Model D compared to 4.63 in Model A. The set of initial constraints employed in Model D produces a different mineral paragenesis, with galena, pyrite, argentite, and chalcopyrite all beginning to precipitate before the miscibility gap is intersected. As fluid phase separation occurs, additional precipitates include dolomite, K-feldspar, calcite, and sphalerite. These are followed by muscovite replacement of K-feldspar at approximately 1380 bars and formation of d&pore at about 500 bars. An increase in the log oxygen fugacity from -35.3 to -33 has considerable effect on the reaction path. Model E is calculated with initial constraints identical to those for Model A with the exception of the higher oxygen fugacity, and the results are shown in Figs. 12-14. Figure 12a shows the log moles of precipitates calculated to form as a function of decreasing pressure, and this diagram can be compared to Fig. 7a for Model A. Several minerals begin precipitating before the fluid intersects the HzO-CO2 miscibility gap, including pyrite, argentite, galena, sphalerite, chalcopyrite, and dolomite. Muscovite begins precipitating as fluid phase separation begins, followed by anhydrite. Pyrite becomes unstable and redissolves. At 1130 bars gold begins to precipitate and remains stable through 550 bars. Argentite becomes unstable and redissolves shortly after gold begins to form. In the middle of the gold interval chalcopyrite becomes unstable and hematite forms in its place. Note that gold precipitation persists

5.0 4.8

’

,

-

-5 - -10

4.6 I

a

-

4.4

-

4.2

-

-15 /

- -20

QN

8J

- -25

4.0

I ,, 1

1

2000

1500

I 1000 PRESSURE,

BARS

- -30 ,

I

500

0

-35

FIG. 12. Reaction path calculation at 250°C as a function of pressure showing the effect of increased oxygen fugacity over that shown in Fig. 7. Initial conditions are log oxygen fugacity = -33, log sulfur fugacity = -9.6, ionic strength = 0.5 m, concentrations of Na + K = Cl. Results of this model are shown in Figs. 12- 14. (a) Calculated log moles of precipitates. Solid curves are sulfide minerals; dashed curves correspond to non-sulfide minerals. Note the presence of gold. ARGEN = argentite, CP = chalcopyrite. (b) pH (solid curve) and oxygen fugacity (dashed curve).

2000

1500

1000 PRESSURE, BARS

500

0

FIG. 13. Initial conditions are as described in the caption of Fig. 12. (a) Log concentration of carbonate, sulfide, and sulfate species in solution. (b) Calculated 634Sof fluid and sulfur-bearing minerals. As the fluid oxidizes and becomes sulfate-rich. the anhydrite and liquid curves merge.

to a lower pressure than any of the sulfide minerals. It forms later in the paragenesis than any other ore minerals, and it maintains a stability field alone after the other ore minerals cease to precipitate. Figure 12b shows the calculated pH and oxygen fugacity as a function of decreasing pressure. The pH trend is very similar to that shown for Model A in Fig. 8a in that it increases between 2000 and 1500 bars, shows a discontinuity in slope at 1500 bars where fluid phase separation begins, and continues to increase with dropping pressure and subsequent volatile loss. In contrast, oxygen fugacity has a much different behavior in this system from that shown in Model A, initially constant and then increasing rapidly at approximately 520 bars. Note that the increase in oxygen fugacity coincides with the cessation of gold precipitation. At high oxygen fugacities the mineral assemblage contains only dolomite, anhydrite, muscovite, and hematite. Figures 13a and 14a-c show the speciation of carbon, sulfur, and metals in the aqueous solution. At the point that the system undergoes rapid oxidation, the concentration of the sulfide species, the Fe++ species, and the Cu’ species drop to nearzero concentrations. Figure 13b shows calculated d34Svalues of the fluid and S-bearing precipitates. The diagram varies considerably from its counterpart for Model A shown in Fig. lob, and the differences are attributable to the predominance of sulfate species over sulfide species in solution (compare Figs. 13a and 8b). In Model E the sulfide minerals have a low 634Ssignature, ranging from -20 to - 25%0.Each mineral varies in 634S by approximately 1%0 over its precipitation interval. 6 34Sof the fluid increases by approximately 4%0over the course of the calculation, and its curve merges with that

T. S. Bowers

2428

-30

I

1 4.50

1500

I

,

500

0

I

I 2000

1000 PRESSURE,

-35

BARS

FIG. 15. Reaction path calculation at 250°C as a function of pressure for the system Si-Al-Ca-Fe-Au-HzS-HrO-COz . Initial conditions are log oxygen fugacity = -33, log sulfur fugacity = -9.6, concentrations of Na + K = 0.1 m, Cl = 0.02 m. (a) Calculated log moles of precipitates where sulfide minerals are shown with solid curves and non-sulfide minerals with dashed curves. (b) Corresponding calculated pH (solid curve) and oxygen fugacity (dashed curve).

2000

1500

1000 PRESSURE,

500

0

BARS

FIG. 14. Initial conditions are as described in the caption of Fig. 12. Log concentrations of metal complexes in the fluid phase: (a) Fe and Cu complexes; (b) Ag and Pb complexes; (c) Zn and Au complexes. Metal complexes with log concentration less than -10 are not plotted.

0

3 5 t

-2

4

-4

I

1

,

1

I

tVWELG

a

% ; -6 0 rn y -0

of anhydrite when the fluid becomes sufficiently oxidized such

8 o-1o

that essentially no sulfide remains. The remaining reaction path calculations were generated for a simpler system of elements including only H, 0, S, C, Cl, Na, K, Ca, Si, Fe, and Au. Without aluminum in the system, quartz may precipitate where muscovite might have been stable in the previous calculations. These systems approximate not only fluids with no Al, but fluids which are poor in Al where quartz would be the favored silicate precipitate. In an effort to understand conditions of gold precipitation in systems undergoing fluid phase separation, Cu, Pb, Zn, and Ag were also precluded from the system for the following calculations. Models F, G, and H, shown in Figs. 15, 16, and 17, respectively, were initiated with log oxygen fugacity = -33, Na + K = 0.1 m, and Cl = 0.02 m. Sulfur fugacity was varied to assess its effect on gold precipitation. Its log value was set at -9.6 in Model F, -6.8 in Model G, and -11 in Model H. Figure 15 shows that pyrite precipitates early in the calculation of Model F and gold precipitates over a narrow pres-

3

-12

,

,

b

-20.0

- -22.5

5.0 I

a 4.5

4.0

II 2000

-35.0 1500

1000 PRESSURE,

BARS

500

0

FIG. 16. Reaction path calculation at 250°C showing the effect of increased sulfur concentration over that shown in Fig. 15. Initial conditions are the same as those for Fig. 15 except that log sulfur fugacity = -6.8. (a) Log moles of precipitates. (b) pH (solid curve) and oxygen fugacity (dashed curve).

2429

Deposition of metals by fluid immiscibility

In

y

-6

zi

-10

6.25 6.00 1

4.75 1 2000

I 1500

I

I

1000 500 PRESSURE, BARS

I -35 0

FIG. 17. Reaction path calculation at 250°C showing the effect of decreased sulfur concentration from that shown in Fig. 15. Initial conditions are the same as those for Fig. 15 except that log sulfur fugacity = - 11. (a) Log moles of precipitates. Note that gold is present as a very narrow spike at the boundary between siderite and hematite and reaches a maximum log value of approximately -10. (b) pH (solid curve) and oxygen fugacity (dashed curve).

sure interval between approximately 1260 and 1180 bars, where it coexists with quartz and anhydrite. Initial precipitation of hematite occurs simultaneously with the cessation of gold precipitation. pH increases continually (Fig. 15b) throughout the calculation with no apparent response to the onset of fluid phase immiscibility. Oxygen fugacity increases abruptly at 1180 bars, which coincides with the beginning of hematite stability and termination of gold deposition. With a higher initial sulfur fugacity in Model G (Fig. 16) gold initially forms at lower pressure, where its precipitation persists over a 350 bar interval. Its precipitation again terminates with the appearance of hematite and shortly after the point of rapid increase in oxygen fugacity. Like Model F, gold coexists with quartz and anhydrite in Model G. Model H demonstrates the case for lower sulfur fugacity in the fluid (Fig. 17 ) Here gold forms over a very narrow pressure interval between 1385 and 1345 bars and coexists with quartz, calcite, and siderite. It ceases to precipitate as hematite replaces siderite and oxygen fugacity increases rapidly. Gold is often reported coexisting with or in close proximity to quartz, calcite, and pyrite; however, many of the models shown above included anhydrite instead of calcite and siderite or hematite instead of pyrite. Anhydrite, siderite, and hematite also commonly occur with gold and silver; see SPOONER et al. ( 1987) and LYNCH et al. ( 1990). Note that alunite may form in place of anhydrite under certain conditions. The stability of these Ca and Fe phases depends on the fugacities of COz, sulfur, and oxygen. A stability diagram is calculated at

250°C and 1500 bars for the initial conditions of these models and is shown in Fig. 18. The fugacity of CO2 is fixed at a value calculated from the MRK equation of state for xco2 = 0.15, which corresponds to the edge of the miscibility gap at this temperature and pressure. The crossing stability lines for anhydrite versus calcite and siderite versus pyrite form four fields where the possible combinations of these minerals are stable. The dotted lines on the diagrams show the stability relation at 250°C 350 bars, and xco, = 0.05, pressure-temperature-composition conditions that also correspond to the HzO-rich limb of the miscibility gap. Examination of this diagram shows that under the pressure-temperature-xco2 conditions of these calculations, calcite plus pyrite will not coexist stably at log oxygen fugacity = -33 as used in the calculation of Models F, G, and H above. However, gold did not form in any of the models where log oxygen fugacity was set at -35.3, corresponding to the magnetite-hematite boundary. Model J, shown in Figs. 19 and 20, has an initial oxygen fugacity log value of -33.5 and sulfur fugacity log value of -9.6. This places the fluid near the upper comer of, but within, the calcite plus pyrite stability field on Fig. 18. Figure 19a shows that gold does precipitate in Model J simultaneously with pyrite, quartz, and calcite. Gold is stable over a pressure interval from 750 to 6 10 bars, although pyrite is replaced by hematite at 7 10 bars. Gold ceases to precipitate shortly after the oxygen fugacity increases sharply (Fig. 19b), similar to its behavior in Models F, G, and H. Speciation of carbon, sulfur, and gold in solution is shown in Fig. 20a and b, where it can be seen that concentrations of the sulfide species drop as the oxygen fugacity increases. The total concentration of gold in solution must increase as native gold dissolves; however, gold is carried in solution as chloride complexes and as Au+ because the Au-bisulfide complexes are no longer stable as the system oxidizes and sulfate species become predominant. DISCUSSION A comparison of the 6 I80 of the fluid phase shown in Fig. 1Oa with that in Fig. 5b shows a similar drop of approximately

-33.5-

bNHYDRlTE

PYiTE

-35.0 i -14

-12

-10

-6 LOG fS,

-6

-4

FIG. 18. Log fugacity of oxygen versus log fugacity of sulfur showing stability fields for anhydrite versus calcite and siderite versus pyrite calculated at 25OT and 1500 bars for xc& = 0.15 (solid lines) and at 250°C and 250 bars for x co2 = 0.05 (dotted lines). Both of these pressure-temperature-composition points correspond approximately to the water-rich limb of the H20-CO2 miscibility gap.

T. S. Bowers

2430

MODELJ

a

-10 -15 -20 -25

QN 8

-I

-30 -35 2000

1500

500 1000 PRESSURE, BARS

0

FIG. 19. Reaction path calculation at 250°C showing the effect of decreased oxygen fugacity from that shown in Fig. 15. Initial conditions are the same as those for Fig. 15except that log oxygen fugacity = -33.5. (a) Log moles of precipitates. Note the coexistence of gold, pyrite, quartz, and calcite. (b) Calculated pH (solid curve) and oxygen

fugacity (dashed curve).

4%~between 1500 and 200 bars at 250°C. The calculations presented in Figs. 3-6 are for partitioning of isotopes between liquid and vapor phases on either side of the miscibility gap without consideration of possible reaction with coexisting minerals. The model calculations presented in Figs. 7-l 7 and 19-20 are for fluid-dominated systems. Although mineral precipitation occurs from the fluid, no interaction with wall rock has been included. This is a realistic scenario for fluid conduits that allow little access of the fluid to the wall rock, or for systems with rapid fluid flow or in which fluid immiscibility and mineral deposition occur over short time intervals. Under such a scenario, the distribution of oxygen isotopes among the mineral phases has little influence on the isotopic signature of the fluid produced by volatile loss when the H20CO2 miscibility gap is intersected. (This observation minimizes the impact of the choice of assumptions leading to calculation of 6 “0 of minerals from Eqn. 8 or 9 or possible alternatives.) This analysis suggests that the calculations shown in Figs. 3-6 for isotopic fractionation under varying pressure and temperature conditions have general applicability to more complex systems involving mineral deposition from fluid-dominated systems. For example, the curves in Fig. 10a demonstrate a 4% variation in precipitating calcite as well as in the fluid over the course of the calculation at 250°C. The liquid curve in Fig. 5a suggests that calcite precipitating along this pressure-temperature path at 200°C would show a 1%Ovariation in its 6 “0 signature. Because the pressure-temperature path was the same for calculations of Models A-J, the calculated 6 “0 values for fluid and minerals were all quite similar to those shown in Fig. 1Oa for

Model A, and oxygen isotope calculations for the subsequent models have not been shown. Of Models A-J, distinction can be made between those with a fluid initially dominated by sulfate and those where the fluid is dominated by sulfide. All of the models that were initiated with a log oxygen fugacity value of -35.3 corresponding to the magnetite-hematite boundary (Models AD) have fluids dominated by sulfide species. In these models the oxygen fugacity drops slightly with dropping pressure and progressive volatile loss, and gold is never deposited because it remains stably complexed with bisulfide. Models E-J begin with higher oxygen fugacities and have fluids dominated by sulfate species. These models show deposition of gold and substantial oxidation at some point in the reaction path. This oxidation is not gradual or linearly related to loss of volatile species such as suggested by the chemical reactions described in the introduction to this paper. Rather, it is an abrupt oxidation that occurs at different pressures depending on the initial constraints on the fluid chemistry of the system. A lower concentration of sulfur in the initial fluid results in this oxidation occurring earlier in the reaction progress at higher pressures (compare Figs. 17b, 15b, and 16b). A comparison of Figs. 15b, 19b, and 8a indicates that decreasing the oxygen fugacity of the starting fluid decreases the pressure at which rapid oxidation occurs, and in fact it does not occur at all in the sulfide-dominated fluids. The mechanism that triggers this abrupt increase in oxygen fugacity is related to the sulfide content of the initial fluid and the amount of H2S loss to vapor as a consequence of phase separation. Sudden oxidation occurs when the two are approximately equal. For sulfidedominated fluids (Models A-D) this does not occur; the

1

-6

1

b -0

2

nqtis);

E-l0

HAI,(

w -12 AUTOS=

ti ’

&.A”+

I-14 -16

/I 2000

1500

1000 500 PRESSURE,BARS

0

FIG. 20. Concentration of species in solution corresponding to the calculations shown in Fig. 19. (a) Carbonate, sulfide, and sulfate species. (b) Gold complexes. Species with a concentration below - 16 are not plotted.

2431

Deposition of metals by fluid immiscibility amount of H2S loss to the vapor phase is substantially less than the initial sulfur content of the fluids. The sulfate-dominated fluids (Models E-J) have much less sulfide initially present, and because HIS is preferentially lost to a vapor phase over sulfate species ( DRUMMONDand OHMOTO,1985 ) , it is possible for all sulfide to be removed from the fluid. Other factors are involved, such as deposition of sulfides and sulfates, and the initial buffering capacity of the fluid-rock system, which acts to control oxygen fugacity and mediate the equilibrium between sulfide and sulfate. For Models F, G, H, and J, oxidation begins at a pressure where the amount of H$ removed to vapor is roughly equivalent to the initial sulfide content of the fluid. For Model E, oxidation does not occur until nearly two times the initial sulfide content of the fluid has been removed to vapor, an indication of sulfate reduction that occurs during phase separation in order to replace sulfide loss and maintain equilibrium. Note than an assumption of sulfide-sulfate disequilibrium would give very different results, at least in the case of Model E. Although gold is carried in solution primarily as bisulfide complexes, it appears that the sulfate-rich fluids have a greater potential for gold deposition. This is related to the progressive destabilization of gold-bisulfide complexes as sulfide is oxidized to sulfate. Note however (Figs. 12, 15, 16, and 17) that gold does not form simultaneously with hematite under extreme oxidized conditions. Gold may coexist with hematite (Fig. 12); it may continue to precipitate as the rapid oxidation of the fluid takes place (Fig. 16); but in either case such a scenario signals the approaching limit of the region of gold stability. In most calculations (Figs. 15 and 17 as examples) the cessation of gold precipitation occurred almost simultaneously with hematite precipitation and rapid oxidation of the fluid. SPOONERet al. ( 1987) have discussed the common development of oxidized systems associated with Archean gold-quartz vein deposits. This is in qualitative agreement with calculations here, showing that only the systems which undergo oxidation deposit gold. The sulfur isotope signatures of sulfide-bearing minerals may provide the key as to whether a fluid is initially sulfide- or sulfate-dominated and therefore has the potential to deposit gold as a result of fluid immiscibility. A comparison of Fig. lob for Model A and Fig. 13b for Model E shows significant differences in both the value of 634Sof precipitating sulfides and the range over which this value might change with progressive volatile loss. SPOONER et al. ( 1987) cite several examples, including the McIntyre mine, the Renabie deposit, and the Canadian Arrow deposit, of Archean gold-quartz veins bearing anhydrite, barite, and hematite with negative 634S signatures in the sulfides (compare to Figs. 12a and 13b). At the Renabie mine, CALLAN and SPOONER( 1987) record pyrite 634Sas low as -7%, with coexisting anhydrite between 9 and 13%. Model E may serve as a general model for such systems. Because sulfur isotope fractionation is very sensitive to several fluid compositional variables as well as the assumption of sulfide-sulfate equilibrium, it will probably be difficult to evaluate the significance of sulfur isotopic compositions of sulfides in natural systems on the basis of the general models discussed here. Rather, site-specific models must be generated. The models presented here show variations in the pressure interval over which metals are deposited, ranging from a spike

of gold over a very narrow pressure interval ( Fig. 17 ) to quite broad depositional ranges in some models for sphalerite, galena, chalcopyrite, and argentite (Fig. 7 ). Both of these extremes may be useful models of metal deposition under specific conditions. A narrow spike of an ore metal over a very limited pressure range may be representative of localized and concentrated deposition. The models showing metal deposition over a broad pressure range have two possible interpretations. Reduction of pressure on a rock from lithostatic to hydrostatic through fracturing would serve to compress the predicted mineral precipitation over a large pressure interval into a narrow vertical extent, serving to localize ore metal deposition. An alternative possibility is that SPOONER et al. (1987) and SIBSON et al. (1988) have reported the occurrence of incremental gold deposition over large vertical extents ( l-3 km). This type of mineralization is often localized along vertical faults and may represent an area of sudden fluid pressure release ( SIBSONet al., 1988)) initiating phase separation and subsequent ore deposition. CONCLUSIONS The computer codes employed in this study could be used for several options and types of models not included above. For example, the initial CO2 content of the fluid can be altered, sulfide and sulfate can be held at disequilibrium, databases for various pressure-temperature paths can be constructed, and interaction of the fluid with wall rock can be considered through the addition of minerals to the equilibrium system as reactants using the standard EQ6 approach. The large number of independent and yet interrelated variables which must be set to specify input for the models presented here precludes a thorough presentation or even investigation of all combinations of potential geochemical significance. The next step in such work is to tailor models to specific field areas where ore deposition is suspected to have occurred simultaneously with, or have been induced by, fluid phase immiscibility, Further work on database and program development is also necessary. For example, the fractionation of oxygen isotopes across the H20-CO2 miscibility gap and between minerals and mixed volatile fluids needs to be determined experimentally. When available, such information can readily be substituted into the current databases, No provision has been made in this study for the effect of CO* on the dielectric constant of water and its subsequent effect on fluid speciation and mineral stabilities as a function of pressure and temperature ( WALTHERand SCHOTT, 1988 ). Nevertheless, the preliminary work presented here provides a first step towards quantifying the conditions under which metal deposition is induced by the onset of fluid phase immiscibility and the response recorded in oxygen and sulfur isotopes. Many of the observations made in the course of this modeling study have been described in field areas. For example, LYNCH et al. ( 1990) describe the relationship of Ag-Pb-Zn deposits in the Keno Hill district, Yukon, to zones of fluid phase immiscibility where deposition of calcite and siderite accompany the pH increase resulting from CO2 loss. Deposition of siderite as a late-stage boiling product occurs with a decreasing 6 ‘*O fluid trend caused by exsolving C02. They

T. S. Bowers

2432

use quartz 6 “0 values to identify areas of fluid phase immiscibility and note that these areas coincide with ore veins with the highest ratio of Ag to Pb. The work of LYNCH et al. ( 1990) reinforces the value of stable isotope fractionation accompanying fluid phase immiscibility as a potential exploration tool. Analogous models to those presented in this study could be calculated for CH,-bearing fluids. The large D/H fractionation between liquid water and CH4 and H2 gas will be manifested through phase immiscibility and progressive volatile loss as increasingly heavy 6D signatures in precipitating minerals. A zoning of several permil could be recorded in Hbearing vein minerals under some conditions, serving as a strong indicator of fluid history. And finally, this study has shown that one cannot expect that CO2 or other volatile loss will be accompanied by linear or simply predicted increases in pH and oxygen fugacity of the fluid. The changes in these variables are a complex function of fluid speciation, buffering capacity of the solution, precipitating minerals, and possible reactions with the wall rock. Acknowledgments-This work was supported by NSF grant RII8800108 in the form of a visiting professorship for TSB at Harvard University. I would like to thank H. C. Helgeson, R. R. Jannas, U. Petersen, and H. D. Holland for their thoughtful comments during the course of this work. Constructive reviews of the manuscript together with helpful suggestions for improvement were provided by Lukas Baumgartner, Dave Janecky, Dimitri Sverjensky, Anthony Williams-Jones, and Scott Wood. Editor&l handling: S. A. Wood

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145. BOURCIERW. L. and BARNESH. L. ( 1987) Ore solution chemistryVII. Stabilities of chloride and bisulfide comolexes of zinc to 350°C. &on. Geol. 82, 1839-1863. BOWERST. S. ( 1989) Stable isotope signatures of water-rock interaction in mid-ocean ridge hydrothermal systems: Sulfur, oxygen, and hydrogen. J. Geophys. Res. 94, 5775-5786. BOWERST. S. and HELGESONH. C. ( 1983) Calculation of the thermodynamic and geochemical consequences of nonideal mixing in the system H20-C02-NaCI on phase relations in geologic systems: Equation of state for H,O-CO,-NaCl fluids at high pressures and temperatures. Geochim. Cosmochim. Acta 47, 1247-1275. BOWERST. S. and TAYLORH. P., JR. ( 1985) An integrated chemical and stable-isotopemodel of the origin of mid-ocean ridgehot spring systems. J. Geophys. Rex 90, 12,583-12,606. CALLANN. J. and SPOONERE. T. C. (1987) Archean Au-quartz vein mineralization hosted in a tonalite-trondhjemite terrane, Renabie mine area, Wawa, North Ontario, Canada. Ontario Geol. Surv. Misc. Paper 136, l-10. COLE D. R. and DRUMMONDS. E. ( 1986) The effect of transport and boiling on Ag/Au ratios in hydrothermal solutions: A preliminary assessment and possible implications for the formation of epithermal precious-metal ore deposits. J. Geochem. Expl. 25,4579. DRUMMONDS. E. and OHMOTOH. ( 1985) Chemical evolution and

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2433

Deposition of metals by tluid immiscibility

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Seawater sulfate reduction and sulfur isotope fractionation in basaltic systems: Interaction of seawater with fayalite and magnetite at 200-350°C. Geochim. Cosmochim. Acta 45,19X’- 1995. SHENBERGER S. M. and BARNESH. L. ( 1989) Solubility of gold in aqueous sulfide solutions from 150 to 350°C. Geochim. Cosmochim. Acta 53,269-218. SIBSONR. H., ROBERTF., and POUL~ENK. H. (1988) High-an8le reverse faults, fluid-pressure cycling, and mesothermal goldquartz deposits. Geology 16,55 l-555. SPWNER E. T. C., BRAYC. J., WOOD P. C., BURROWSD. R., and CALLANN. J. ( 1987) Au-Quartx vein and ~-Au-~-M~hy~te mine~tion, Hank-McIn~ mines, Timmins, Ontario: d 13C values (McIntyre), fluid inclusion gas chemistry, pressure (depth) estimation, and H@-Ca phase separation as a precipitation and dilation mechanism. OntarioGeol. Surv. Misc. Paper 136,35-56. SPYCHERN. F. and REEDM. H. ( 1986) Boiinn of aeothermal waters: Precipitation of base and pre&ous’metals~sp&iation of arsenic and antimony, and tlte role of gas phase metal transport. In Proceedings ofthe Workshopon Ge~hemicaI fueling (eds. K. J. JACKSONand W. L. BOURCIER),pp. 58-65. Fallen Leaf Lake, CA. TAKENOUCHIS. and KENNEDY G. C. ( 1964) The binary system H&-COz at high temperatures and pressures. Amer. J. Sci. 262, 1055-1074. TAKENOUCHI S. and KENNEDYG. C. ( 1965) The solubility of carbon dioxide in NaCl solutions at high temperatures and pressures. Amer. J. sci. 263,445-454. TODHEIDEK. and FRANCK E. U. ( 1963) Das Zweiphasengebiet und die kritische Kurve im Kohlendioxid-Wasser bis zu Drucken von 3500 bar. Z. Phys. Chem. N. F. 37,387-401. TREMAINEP. R. and LEBLANCJ. C. ( 1980) The solubility of mag netite and the hydrolysis and oxidation of Fe+* in water to 3OO’C. J, S&r. Chem. 9,4i 5-442.

Appendix Log&

are from the EQ3/6 database

1. LogKs at 250°C and Saturation

unless otherwise noted.

Note that the validity and consistency

discussed in this study, and many values wilI llkeIy be superseded Cu++ + 0.5 Ha0 = cu+ + H+ f 0.2502 -6.3329 CuOH+ + H+ = Cu++ + Hz0 -1.02 cucq

= cu+

CuClF = cut

Pressure of these data are not

by future work. AgClT = Ag+ + 3ClAgC1,3 = A&

-5.3

+ 4Cl-

-6.0

+2c1-

-6.5

AgSO,

= Ag+ + SO;

+ 3c1-

-7.2

AuC&= = Au+ + 2CP

-6.2 -19.3484

-2.69

+ cl-

-4.8

b

Au(HS);

CUCI, = cu++

+ 2c1-

-4.9

c

Au2(HS)$=

&lCl,

= cut+

+ 3Cl-

-4.6

c

HAu(HS)a

CuC@ = cut+

+ 4c1-

-3.9

Fe+++ + 0.5Hs.O + Fe++ + Ht + o.2502

-0.2263

cuso4

f

-1.37

Fe(OH)+

7.12

cuc1+

= cu++

=

cd+

SOT

= Au+ + 2HS-

+ H+ = ZAu+ + 3HS= H+ + Au+ + 2HS-

+ H+ = Fe++ + Hz0

-31.76 -24.26

PbCl+ = Pb++

f Cl-

-2.4

Fe(OH)2 + 2H+ = Fe++ -+ 2HzO

12.99

PbCla = Pb++

+ 2Cl-

-3.1

Fe(OH),

22.1

+ 3H+ = Fe++ + 3Hz0

PbCI,- = Pb++

+ 3Cl-

-3.6

Fe(OH)s -l- ZHt = Fe++ + 2.5H20 + 0.2502

PbCly

= Pb++

+ 4Cl-

-4.0

Pe(OH);

PbCOs

+ H+ = Pb++ + HCO,

0.124

-3.33 -3.06

+ 3H+ = Fe++ + 3.5HsO + 0.25&

ZnCl+ = Zn++ + Cl-

-4.4

FeCl2 = Fe++ + 2ClFeCl; = Fe++ + 4Cl-

Z&la

11.7438

-5.6

Feel++

+ Cl-

-4.67

ZnClg = Zn++ + 3C1-

-6.7

FeCl$ = Fe+++ + 2Cl-

-1.45

Z&l:

-6.0

FeCb = Fe+++

+ 3CI-

-7.1

-16.4

FeCll = Fe+++

+ 4Cl-

-6.6

Zn(HS)a = Zn++ + 2HS-

-13.8

FeSO4 = Fe++ + SO;

-4.75

Zn(HS); Zn(HS);

-15.3

FeSO$ = Fe+++

-7.91

-15.0

FeCO3 f H+ = Fe++ + HCO,

ZnSO4 = Zn++ + SO;

-5.34

FeCOt

A&l = Ag+ $ Cl-

-3.0 -4.7

FeHCOg = Fe++ + HCO;

-1.1992

NaOH = Na+ + OH-

-0.01

+ Zn++ + 2Cl= Zn++ + 4Cl-

Zn(OH)(HS)

AgCF

= Zn++ + OH-

= Zn++ f 3HS= Zn++ + 4HS-

= A&

+ ZCl-

+ HS-

= Fe+++

3.5438

+ SO;

+ H+ = Fe+++

+ HCO;

12.8268 1.7057

T. S. Bowers

2434

APPENDIX1. (continued) f

NaCl = Na+ + Cl-

-0.50

MgSOl

= Mg++

+ SOT

NaSO;

= Na+ + SO;

-1.71

M&O3

= Mg++

+ COY

NaCOs

= Na+ + COT

2.4474

MgHCO;

= Mg++

1.6125

Al(OH)++

KC1 = K+ + Cl-

0.3

Al(OH),

KSO;

-2.5

AlSO:

NaHC03

= Nat

+ HCO,

= K+ + SOT

Ca(OH)+

= Cat+

+ OH-

-3.4

HsSiOr h

-5.4226

+ HCO,

= Al+++

+ OH-

= Al+++ = Al+++

-5.18

+ 40H-

-13.1 -33.75

+ SOT

+ H+ = SiOz(,,)

-2.3035

-5.53 + 2H20

HCl = H+ + Cl-

9.6976

g

cac1+

= ca++

+ Cl-

-0.7

g

C&l2

= Cat+

+ 2Cl-

-2.3025

HS-

-4.31

HSO;

= H+ + SO:

-5.41

= H+ + HS-

-7.56

CaSO4 = Cat+ CaCOs

Mg(OH)+

= Cat+ = Mg++

MgCl+ = Mg++

+ 202 = H+ + SO;

63.8499

5.564

H2S,,

+ HCO,

-2.756

SE + H+ = HS-

+ OH-

-5.06

CO;

=HCO,

11.286

-2.28

C02(aq) + H20 = H+ + HCO,

-7.763

+ H+ = Cat+

CaHCOt f

+ SOT

-0.68

+ HCO,

+ Cl-

+H+

12.14

a) Bourcier and Barnes (1987), b) Shenberger and Barnes (1989), c) S eward (1973), d) Tremaine and LeBlanc (1980), e) B. Bourcier, pers. comm. and Helgeson (1985).

(1987), f) D. Sverjensky, pers. comm.

(1986), g) W’lli I ams-Jones

and Seward (1989), h) Jackson