sn w heinrich

Economic Geology Vol. 85, 1990, pp. 457-481

The Chemistry of Hydrothermal Tin(-Tungsten) Ore Deposition CHRISTOPH A. HEINRICH

Bureau of Mineral Resources, Geology and Geophysics, Division of Petrology and Geochemistry, P.O. Box 378, Canberra, A.C.T. 2601, Australia

Abstract

A qualitative model for the formation of granite-related mesothermal cassiterite(-wolframite) deposits is derived from a review of geologic evidence and is then tested quantitatively with chemical mass transfer calculations based on published experimental and thermodynamic data.

Laboratory experiments and geologic, mineralogic, fluid inclusion, and stable isotope observations indicate that saline fluids of magmatic origin are involved in the formation of most tin ores that occur in veins, breccias, and replacement bodies of aluminosilicate or carbonate rocks. Transport of the ore fluid from a hot granitic source into a cooler depositional environment probably involves structural focusing and prevention of complete chemical reequilibration of the fluid with fresh quartzofeldspathic wall rocks. Under these conditions, a reduced acid fluid can transport high concentrations of Sn(II)-C1 complexes (hundreds of ppm metal) to a site of ore deposition at low temperature. Precipitation of cassiterite, Sn(IV)O2, requires oxidation and liberates acidity, which must be balanced by reduction and acid-consuming reactions involving other fluid and wall-rock components for cassiterite enrichment to proceed to economic concentrations. Several geologically likely deposition mechanisms have been tested, which differ in efficiency regarding the maximum tin ore grade that can be achieved in aluminosilicate host rocks. By contrast, wolframite can be precipitated by cooling of an Fe- W-bearing fluid without wall-rock reaction.

Single-step acid neutralization of magmatic fluids by feldspar hydrolysis to sheet silicates (phyllic alteration) probably produces subeconomic greisen mineralization, because the maximum tin ore grade is severely constrained by the high acid content of the fluid. Progressive fluid-rock reaction and multistage ore reworking at an advancing alteration front, described by a one-dimensional finite-element reactor model assuming local equilibrium, may be a more efficient and geologically realistic process to form tin-rich greisens and breccia pipes. Loss of H2 from a reduced tin-rich fluid by vapor separation, and simultaneous reaction with aluminosilicate rocks, is an alternative possibility for the formation of rich greisen-type deposits. Fluid mixing by injection of minor magmatic fluid into a cooler environment of convecting meteoric fluids could be a third, particularly efficient, tin-mineralizing mechanism in vein deposits without extensive wall-rock interaction. In this case and in the deposition of cassiterite by carbonate replacement, there are essentially no chemical limitations on tin ore grade, other than dilution of cassiterite by coprecipitating quartz and sulfides.

Introduction

A LARGE number of detailed geologic, mineralogic, fluid inclusion, and stable isotope studies of tin- (_ tungsten) ore deposits have been published in the last two decades (Kelly and Turneaure, 1970; Landis and Rye, 1974; Jackson et al., 1977, 1982; Charoy, 1979; Kelly and Rye, 1979; Collins, 1981; Patterson et al., 1981; Eadington, 1983; Andrew and Heinrich, 1984; Campbell et al., 1984; Eadington and Paterson, 1984; Higgins, 1985; Andrew et al., 1986; Ren and Walshe, 1986; Solomon et al., 1986; Witt, 1988; Schwartz and Askury, 1989; and others). These doc- ument a bewildering range of geochemical properties and structural settings but also point out a surprising number of common features, above all their intimate association with differentiated felsic magmatites. Tin deposits occur in volcanic and shallow subvolcanic settings (e.g., in Mexico, Huspeni et al., 1984; and

Bolivia, Sillitoe et al., 1975; Francis et al., 1981), but the majority of large tin mines host cassiterite in veins, breccias, and replacement bodies formed in moderately shallow plutonic environments (see reviews by Taylor, 1979; Heinrich and Eadington, 1986; Kwak, 1987; Plimer, 1987; Heinrich et al., 1989; Solomon et al., 1991).

Geochemical aspects of granite-related ore formation have recently been discussed by Burnham and Ohmoto (1980), Eugster (1985), Eugster and Wilson (1985), and Candela (1990a, b), with the main focus on chemical mechanisms responsible for the generation of metal-rich fluids in magmatic environments at high temperature. Less quantitative studies have focused on mechanisms of cassiterite precipitation and tin enrichment to form a primary hydrothermal tin or tungsten deposit. The present paper attempts to show, on the basis of published experimental thermodynamic data and mass balance considerations, that

0361-0128/90/1053/457-2553.00 457

458 CHRISTOPH A. HEINRICH

the chemistry of cassiterite precipitation severely constrains possibilities of forming a tin deposit with economic ore grade.

Based on a review of what are believed to be the

most crucial geologic observations common to many tin(___ tungsten) deposits, a number of geologically reasonable mechanisms of cassiterite and wolframite

deposition will be tested with chemical mass transfer calculations, using a set of thermodynamic data derived from recent experimental studies. The numerical experiments were designed with the basic idea that, among a range of equally likely geologic processes, those allowing the highest degrees of chemical enrichment of cassiterite in a given mass of rock (i.e., maximum tin ore grade) can be predicted as the most likely processes to form an economic ore deposit (as opposed to a subeconomic tin anomaly). For this modeling, the software package CSIRO-SGTE THERMODATA was applied (Turnbull and Wadsley, 1986), making use of various options that were orig- inally designed for the optimization of industrial chemical processes. The present study has been un- dertaken as groundwork for, and at the same time draws observations and ideas from, a number of on- going field studies of Australian tin and tungsten deposits (Andrew and Heinrich, 1984; Solomon et al., 1986; Higgins et al., 1987).

Results of mass transfer modeling are dependent on the geologic assumptions upon which they are based. Using the same experimental data and the same modeling capabilities, any number of computed re- suits may be obtained depending on what set of assumptions is chosen. Errors in the thermodynamic data used in the calculations will be reflected in the

accuracy of the model predictions. None of the calculations below reflect absolute constraints on the

formation of tin-tungsten ores, but they should help to point out the key chemical and physical parameters that may control ore-forming processes. These parameters should serve as useful guides for determining the essential ingredients of advanced genetic models for mineral exploration. The model calculations will also highlight the likely key problems for further geologic and experimental investigations into the formation of tin and tungsten deposits.

Geologic Processes of Sn-W Vein and Breccia Mineralization

Magmatic characteristics of source granitoids

Most tin deposits are associated with highly differentiated biotite granites free of hornblende or other calcium silicates (amphiboles, pyroxenes, sphene), whose absence may be a prerequisite for the enrichment of Sn as an incompatible element in the residual melt (Ishihara, 1977; Eugster, 1985). They are reduced and typically contain ilmenite rather than

magnetite as the main opaque Fe-Ti phase, with fo2 below Ni-NiO (NNO) and generally near the quartz- fayalite-magnetite (QFM) buffer (Ishihara, 1977; Burnham and Ohmoto, 1980; Shimazaki, 1980; Wones, 1981). This may reflect an origin of the granites from melting of crustal sources including pelitic metasediments which commonly contain organic matter. Many tin granites (for example, Jackson et al., 1982; Eugster, 1985; Schwartz and Askury, 1989) show characteristics of S-type granitoids in the clas- sification of Chappell and White (1974). Most Aus- tralian examples show mixed S and I characteristics but are invariably ilmenite series granitoids (Solomon et al., 1991).

Tin granites are 2 to 20 (rarely up to 100) times enriched in Sn compared with ordinary granites, due to magmatic fractionation from source magmas that are at most two- to threefold enriched compared with ordinary granites or average crustal abundance (Leh- mann, 1987; Solomon et al., 1991; Table 1). Tin

TABLE 1. Composition and Mineralogy of a Hypothetical Tin Granite

Average composition (wt %) based on Idealized composition

typical analyses in R1 used in Solomon et al. calculations of fluid-

(1991) rock reaction

SiO2 75.74 77.28 TiOz 0.06 A1203 12.87 12.77 Fe•O3 0.48 0.41 FeO 0.46 0.73

MnO 0.02

MgO 0.03 CaO 0.37

Na•O 3.54 3.62 K•O 4.75 4.86 P205 O. 11 H•O 0.27 0.33 CO2 0.14 SnO• 0.005 F 0.45

Others

(Rb, Li, SrZr) 0.41

The idealized rock composition R1 has been derived by adjustment to the stoichiometric formula Si•.868A12 so8-

+3 +2 Feo.o.5•Feo.•o.•Na•.•7K• o3.•Ho.•.5703o.•.56, which has a formula weight of 1 kg and is consistent with the following modal mineralogy:

Modal mineralogy of idealized rock R1

Mineral mole/kg rock Vol %

K feldspar KAISiaOs 0.828 21 Albite NaAISiaOs 1.170 30 Quartz SiO= 6.262 41 Muscovite KAIaSi30•o(OH)= 0.153 6 Biotite KFeaAISiaO•o(OOH) 0.051 2

HYDROTHERMAL Sn(- W) CHEMISTRY 459

granites are commonly enriched in F, Li, and/or B, with topaz, fluorite, lepidolite and tourmaline as magmatic or early subsolidus alteration minerals. Fluorine lowers the solidus temperature of granitic melts to eutectic temperatures as low as 650øC at 2 kbars, with residual melts enriched in albite and H20 relative to the F-free granite minimum (Manning, 1982; cf. Eadington, 1983; Weidner and Martin, 1987). These components also lower the melt viscos- ity at a given temperature, which may be instrumental in the physical separation between melt, crystals, and fluids (Dingwell, 1988).

Granitoids associated with tungsten-only deposits are not as well characterized as tin granites. They are generally I type, encompass a much wider compositional range from granodiorite to alkali feldspar granites, and contain magnetite rather than ilmenite (Newberry and Swanson, 1986; Kwak, 1987).

Alteration, ore mineralogy, and paragenesis Tin and tungsten granites show extensive and com-

monly pervasive metasomatic alteration (e.g., Shcherba, 1970; Landis and Rye, 1974; Pollard, 1983; Jackson and Helgeson, 1985b; Stemprok, 1987) which starts at high temperature where magmatic and postmagmatic processes are hardly distinguishable (cf. Burnham and Ohmoto, 1980). Early high-temperature stages include feldspar metasomatism (albitization or replacement by potassic feldspar), associated with miarolitic veins or as pervasive alteration zones near the roof of intrusions. This is followed by pervasive or vein-related (and hence clearly subsolidus) alteration to greisen consisting dominantly of muscovite + quartz ___ biotite or chlorite ___ topaz ___ fluorite. Ka- olinitc alteration is generally the latest and occurs at lower temperatures. The main alteration types and their relationship are similar to the alkalic, phyllic, and argillic alteration types associated with porphyry copper deposits (Landis and Rye, 1974; Eadington and Giblin, 1979; Pollard, 1983).

Cassiterite and wolframite mineralization in alu-

minosilicate host rocks (granites, felsic volcanics, quartzofeldspathic metasediments) is mostly associated with phyllic alteration. It occurs as low-grade disseminations in greisenized granite and in higher local concentration within veins and breccias. Alter-

ation in many rich vein deposits (particularly the ones hosted by semipelitic hornfelses) is confined to narrow selvages of complete replacement of feldspars by muscovite (Eadington, 1983; Solomon et al., 1986; Andrew and Heinrich, 1984). This, and vein textures indicating repeated overgrowths of hydrothermal minerals on the vein walls, suggests that chemical interaction between vein fluids and feldspathic rocks was commonly inhibited by low wall-rock permeability (Heinrich and Eadington, 1986). Carbonate host rocks are intensely altered to skarns, with a wide

range of silicate and sulfide assemblages (e.g., Collins, 1981; Patterson et al., 1981; Eadington and Kinealy, 1983; Holyland, 1987; Kwak, 1987).

Cassiterite in veins is generally precipitated to- gether with quartz and muscovite or Li micas, frequently associated with F minerals, biotite, chlorite, arsenopyrite, and/or pyrite. In many composite or crosscutting veins, this early oxide stage is followed by a sulfide stage comprising chalcopyrite, pyrrhotite, sphalerite, galena, stannite, and other sulfides. Chlo- rite and carbonates are the dominant gangue minerals associated with the sulfide stage. The mineralogy and the widespread occurrence of CH4 in fluid inclusions (below) suggests a relatively reducing depositional environment in most aluminosilicate-hosted tin de-

posits (Heinrich and Eadington, 1986).

Stable isotopes and fluid inclusions

Combined stable isotope and fluid inclusion studies provide the main evidence, besides the close geologic association, for a genetic relation of hydrothermal Sn ores with granites. Oxygen isotopes of quartz precipitated with cassiterite or wolframite are generally consistent with an isotopically heavy fluid that could have originated from a magmatic source (Landis and Rye, 1974; Kelly and Rye, 1979; Collins, 1981; Pat- terson et al., 1981; Andrew and Heinrich, 1984; Campbell et al., 1984; Higgins, 1985; Sun and Ead- ington, 1987). However, oxygen isotopes alone cannot generally differentiate a true magmatic fluid, pro- duced by exsolution of a crystallizing magma, from a fluid of meteoric origin that was equilibrated with a hot but solid granite under rock-dominated conditions (T > 400ø-500øC, fluid/rock ratio < 0.1).

Combination with hydrogen isotopic data can provide a more sensitive tracer for meteoric fluids, but only a small number of detailed studies of deposits at suitably high palcolatitude or palcotopography have been published (Andrew and Heinrich, 1984; Camp- bell et al., 1984, also quoting Landis and Rye, 1974, and Kelly and Rye, 1979; Sun and Eadington 1987). The data for the Mole Granite (New South Wales; Eadington, 1983; Sun and Eadington, 1987) clearly indicate direct input of saline magmatic fluids into a tin-depositing hydrothermal system. All other examples have in common a fluid •SO signature consistent with a high-temperature magmatic source, but a linear trend in •D extending to lighter than magmatic values. Campbell et al. (1984) interpreted this trend as a result of rock-dominated exchange of meteoric fluid with granite at 400 ø to 600øC. However, the ratio of meteoric fluid to rock required to explain the data of Campbell et al. (1984; 0.1-2 wt %) is substantially lower than the amount of magmatic fluid that is likely to be exsolved during crystallization of a water-saturated granite (5-10 wt %; Burnham, 1979). The interpretation of Campbell et al. (1984)


thus implies that comparatively large amounts of igneous fluid have been lost from the system at an early stage which is unrelated to the inferred later leaching of the granite and formation of the ore veins by a smaller amount of meteoric fluid. This seems improb- able, since the data of Campbell et al. (1984) could be explained equally well by minor admixture of meteoric water to a dominantly magmatic fluid, and reequilibration of the mixture with granite at 400 ø to 600øC (at higher total fluid/rock '• 0.15). This ex- planation is preferred for a similar •3•sO/•D trend observed at the Sundown tin prospect (SE-Queensland; Andrew and Heinrich, 1984). Other explanations for slightly reduced •D values in magmatic fluids are discussed by Higgins (1986). Irrespective of the ambi- guity regarding ultimate sources of H20, the stable isotope data from nearly all deposits indicate that tin- and tungsten-mineralizing fluids were equilibrated isotopically (and probably chemically as well) with a hot (y400øC) granitic source rock, prior to their transport into a cooler (•400øC) ore deposition site.

Sulfur isotopes of sulfides associated with cassiterite deposits are generally consistent with a reduced igneous sulfur source (e.g., Landis and Rye, 1974; Kelly and Rye, 1979; Collins, 1981; Patterson et al., 1981). Some sulfide-rich deposits contain a second sulfur component contributed by sedimentary wall rocks (e.g., Andrew and Heinrich, 1984).

Fluid inclusions in most tin deposits are saline (5- 50 equiv wt % NaC1) and commonly contain CO2 (Taylor, 1979; Heinrich and Eadington, 1986; Kwak, 1987). Limited data are available on the salt components, but FeCI• is probably a major component besides NaC1, KC1, and minor CaCI•, MgCI•, and ZnCI• (Eadington, 1983; Bottrell and Yardley, 1988). Substantial fractions of CH4 have been observed in CO2-rich inclusions, suggesting fairly reduced fluids in these instances (e.g., Patterson et al., 1981; Ramboz et al., 1985; Hoffmann et al., 1988).

Reported homogenization temperatures of fluid inclusions associated with cassiterite mineralization

vary from about 220 ø to over 500øC, but main ore- stage filling temperatures are mostly between 280 ø and 400øC (Haapala and Kinnunen, 1982; Kwak, 1987). Where fluid inclusion data are available in combination with paragenetic observations, they commonly indicate a sequence of mineral deposition with falling temperature (e.g., Kelly and Turneaure, 1970; Kelly and Rye, 1979; Collins, 1981; Patterson et al., 1981; Andrew and Heinrich, 1984; Solomon et al., 1986). The early oxide stage of many deposits typically precipitated at 300 ø to 400øC, i.e., lower than the temperatures inferred for the last rock-dominated oxygen isotope exchange of the ore fluids with the magmatic source rock. A further temperature decrease is recorded by the late-stage sulfide precipi-

tation at 200 ø to 300øC (Heinrich and Eadington, 1986, table 1).

Correlations between filling temperature and total salinity are common, suggesting mixing of a magmatic saline fluid with cooler, lower salinity meteoric waters at the site of ore deposition (Kelly and Rye, 1979; Eadington, 1983; Davis and Williams-Jones, 1985; Witt, 1988). Meteoric fluids increasingly predominate toward the paragenetically late stages of mineralization (sulfides, carbonates; e.g., Kelly and Rye, 1979; Eadington, 1983; Witt, 1988).

Fluid inclusion evidence for boiling, or separation of a CO•- (CH4-)-rich gas phase, is recorded from many tin and tungsten deposits (see Heinrich and Eadington, 1986, table 1; Solomon et al., in press). Some tin systems record coexistence of two fluids, a high-density brine or salt melt and a medium-density vapor, already from an early magmatic stage (e.g., Eadington 1983; Solomon et al., 1986). Where si- multaneously trapped liquid- and vapor-rich inclusions have been studied quantitatively, they suggest fluid pressures of 200 to 600 bars for the ore deposition stage (Solomon et al., 1986; Eadington and Pat- erson, 1984; Davis and Williams-Jones, 1985; Hig- gins, 1985; Ramboz et al., 1985; Ren and Walshe, 1986). These relatively low pressures are interpreted as near-hydrostatic in studies where independent estimates of overburden are available (e.g., Landis and Rye, 1974), or where lithostatic pressures were estimated from preore fluid inclusions (e.g., Kelly and Rye, 1979; Eadington and Paterson, 1984) or from contact-metamorphic mineral assemblages (Solomon et al., 1986; Hoffmann et al., 1988).

Interpretation: Hydrodynamic, thermal, and chemical evolution

The formation of mesothermal tin(-tungsten) deposits is probably dominated by fluid processes at the transition, in time and space, from a magmatic lithostatic source regime to a cooler hydrostatic depositional environment (Fig. 1). The chemical and isotopic constitution of the ore-forming fluid(s), including ligand and ore metal aquisition, is probably dominated by a direct contribution of magmatic fluids.

Source regime--hot geopressured stage: Many tin granites contain miarolitic cavities suggesting fluid saturation at near-solidus conditions. Experimental data suggest that the aqueous fluid exsolved by a crystallizing hydrous granodiorite at P > I kbar is moderately saline and single phase (Burnham, 1979), but elevated C1/H in the melt, lower pressure, and higher temperatures typical for porphyry-copper stocks favor two fluids, a low-salinity vapor and a dense brine, to separate directly from the melt (e.g., Eastoe, 1978, 1982; Bodnar et al., 1985). High-density brine inclusions are observed in some tin systems

HYDROTHERMAL Sn(-W) CHEMISTRY 461

• F 7'0 F }t xx 300C

,,' I

D

C

B ,

* "1 A

+ + + + + + + + + +l 16-3/276

Rock/Process

Quartzofeldspathic country rock (granite, volcanics, semipelitic hornfels)

Cassiterite ore vein or

breccia pipe, restricted phyllic alteration

Fluid Characteristics

Mostly meteoric: Low- salinity cool aqueous solution + CO2/HCO3-

Mostly magmatic, or mixed

magmatic+meteoric: Medium-salinity aqueous solution + low-density vapour

Fluid Pressure

Near-hydrostatic

Near-hydrostatic after initial vein formation

Site of initial vein Magmatic Near-lithostatic at formation (e.g. high hydraulic failure curvature in contact)

Pervasive or vein - Mostly magmatic: Lithostatic related phyllic Medium salinity changing to alteration (greisen) _+ aqueous solution hydrostatic disseminated Sn

High-T postmagmatic Magmatic: Residual alteration: dispersed salt melt + buoyant feldspathisation etc. medium-salinity

vapour

Crystallizing source Magmatic: Medium- granite salinity fluid, super-

critical or coexisting with salt melt

FIG. l. Schematic cross section through a shallow plutonic tin-mineralizing system. Tin(-tungsten) ores are interpreted to form by venting of magmatic fluid from a lithostatic igneous source regime into an overlying hydrostatic ore-depositional environment, where the magmatic fluid cools and interacts with wall rocks and/or external fluids.

?Lithostatic

Lithostatic

(e.g., Eadington, 1983; Solomon et al., 1986; Hoff- mann et al., 1988; Witt, 1988) but are not as common as in porphyry copper deposits, possibly due to greater emplacement depth of most tin granites and/ or somewhat lower solidus temperatures due to additional melt components such as F (Manning, 1982; Weidner and Martin, 1987). On the other hand, the presence of CO2 extends the vapor-brine solvus at high P and T (Holloway, 1976; Henley and McNabb, 1978; Gehrig, 1980) and probably favored separation of two fluids from the final melt in some tin-tungsten systems (e.g., Aberfoyle; Solomon et al., 1986).

Burnham (1979) has pointed out that crystallization reactions of the type:

hydrous granitoid melt --* granite

4- aqueous fluid(s), (1)

involve a positive volume change up to several tens of percent (depending on H20 content and emplacement depth) which leads to a fluid overpressure within the pluton. Failure of early crystallized outer shells of the intrusion (carapace) and of the intruded country rocks by hydraulic fracturing occurs when the fluid pressure exceeds the lithostatic load pressure (Burn- ham, 1979) or more exactly the sum of lowest prin-

cipal stress and rock strength. Initial hydraulic failure will be localized at points of stress concentration, for example, the point of highest curvature in granite cu- polas (Solomon et al., 1986; Fig. 1) and this probably controls the localization and geometric style of many tin-tungsten deposits.

Ore depositional regime--cooler hydrostatic stage: Initial fracturing leads to fluid loss, decrease in fluid pressure, and a tendency for veins to collapse under the load of the rock column. The hydrodynamic evolution after initial rupture depends on the rate of fluid loss to the surface (or an overlying hydrostatic regime), relative to the supply rate of magmatic fluid from the granite into the vein, breccia, or replacement body. Well-documented examples of vein and breccia deposits record a change with time from high-temperature lithostatic conditions prior to ore formation, to near-hydrostatic fluid pressures attending the main stage of tin or tungsten precipitation (Landis and Rye, 1974; Kelly and Rye, 1979; Jackson et al., 1982; Ead- ington, 1983; Eadington and Paterson, 1984; Ren and Walshe, 1986; Solomon et al., 1986). These deposits have probably not formed by diffusive closed-system ore transport within a stagnant ore fluid (which requires veins to be kept open by lithostatic fluid pressure for periods of 10,000-100,000 yr; e.g., Kigai,


1978), but rather by focused fluid advection in a dynamic open environment.

There is widespread evidence for decreasing temperatures with time during the evolution of many tin deposits. These observations might be explained by general cooling of the deposit and its igneous fluid source as a whole, due to dissipation of heat from an igneous intrusion through conduction and convection of meteoric water (Cathies, 1977). However, the interpretation of stable isotope data from many tin- tungsten deposits implies cooling of the fluid itself between source and deposition site. Since the site of ore formation is likely to be a site of focused fluid flow through a small volume of rock (below), the heat transferred by the fluid is probably too great to be dissipated by local heating of the wall rock (Barton and Toulmin, 1961). The two most effective cooling mechanisms consistent with observations on tin-tungsten deposits are probably, adiabatic expansion (re- versible, or irreversible throttling, with or without vapor separation); and heat transfer between hot magmatic and cooler meteoric fluids, which may or may not include physical fluid mixing (Barton and Toulmin, 1961). Eadington (1983) and Sun and Ead- ington (1987) have presented convincing evidence that the tin deposits around the Mole Granite were located at an interface between a regime of high-temperature igneous brines, and a lower temperature environment of low-salinity meteoric fluids (at a local scale of hundreds of meters or less). The outer meteoric fluid system was probably under hydrostatic pressures and convecting, thus providing a mechanism of efficient heat dissipation on a larger scale of kilo- meters (Eadington, 1983; Hoffmann et al., 1988; see also Eastoe, 1978, 1982; Henley and McNabb, 1978; and Fig. 1).

Composition of Magmatic Tin(-Tungsten) Ore Fluids

The foregoing review suggests that the composition of tin ore fluids, prior to their arrival at the ore deposition site, are primarily controlled by chemical processes in the magmatic environment. Even where some infiltration of meteoric water into the source

area does occur, the major element solute load of tin and tungsten ore solutions is probably dominated by element partitioning between final melt and exsolving magmatic fluids and by high-temperature subsolidus reequilibration with magmatic minerals. Based on this assumption, the major element composition of a simplified Fe-Na-K-Si-C1-O-H-bearing model ore fluid (fluid F1, Table 2) has been estimated, using high- temperature equilibria and limits from experimental data by Hemley (1959), Chou and Eugster (1977), Lagache and Weisbrod (1977), Chou (1978), Burn- ham (1979), Holland and Malinin (1979) and Gunter

TABLE 2. Composition of Hypothetical Tin Ore-Forming Fluids (concentrations in mole/kg H•O)

Range used in Component Fluid Fll additional calculations

NaCI 0.62 0.32-1.86 KCI 0.21 0.10-0.63

FeCI• 0.07 0.03-0.42 HCI 0.02 0.002-0.06

H• 0.3 0.03-3.0 CO• 0.075-1.0 • H•S 0.06-0.84 SnCI• 0.004 (500 ppm) 0.0004-0.012 H•WO4 0.00055 (100 ppm) SiO2 0.04

I Hypothetical fluid composition in equilibrium with K feldspar + albite + muscovite + quartz, oxidation potential buffered by quartz + fayalite + magnetite, at 600øC and 1 to 2 kbars; used in calculation of Figures 3 to 8 and 11

2 Additionally used in fluids similar to F1, for calculation of Figure 9

and Eugster (1980). A total salinity of 1 m Cltotal, a temperature of 600øC, a lithostatic fluid pressure of 1 to 2 kbars, and at a redox level near QFM was assumed to derive the composition of fluid Fl. A wider range of total salinities and solute ratios as well as a number of additional components (Table 2, right) were used in various numerical experiments discussed below. Metal ratios in the same range as these experimental estimates have been observed by leachate analyses of magmatic brine inclusions from tin granites (Eadington, 1983; Bottrell and Yardley, 1988).

Taylor et al. (1984) and Taylor (1988) have measured the partitioning of Sn between granitic melts and supercritical fluids of low to intermediate chloride concentration, at controlled oxidation potentials. Tin partitioning in favor of the fluid phase increases with increasing fluid salinity and hydrogen fugacity, indicating the presence of complex Sn(II)-C1 species in the fluid and of Sn(IV) in the coexisting silicate melt (e.g., Snfluid/Snmelt = 10-30 at QFM and 1 m Cltotal in the fluid; Taylor, 1988). Considering that average tin granites contain a few tens of ppm Sn (Table 1), a concentration of 500 ppm Sn in the model ore fluid was assumed (Table 1). This is probably at the upper limit for any tin ore fluid of this salinity arriving at the site of ore deposition (below). The less extensive data available for tungsten (Manning and Henderson, 1984) suggest that similarly high concentrations of tungsten could be partitioned from granitic melts into exsolving saline fluids. An arbitrary concentration of 100 ppm tungsten (5.5 X 10 -4 m as He, WO4) has been selected to illustrate some aspects of wolframite deposition.

A fluid with the bulk composition F1 (Table 2) is supercritical in its magmatic source region at 600øC


and 1 kbar, but after cooling below 400øC and a transition to hydrostatic conditions at similar depth it rep- resents a subcritical aqueous liquid slightly above its boiling pressure ('--300 bars; Bodnar et al., 1985). Sufficient thermodynamic data are now available to calculate, with some confidence, many multicompo- nent fluid-mineral equilibria at conditions approxi- mating the depositional regime of tin deposits, using experimental measurements on various subsystems at the saturation vapor pressure up to 350øC (Table 3). For lack of quantitative experimental data and geologic information on the intervening cooling and de- pressurization path of magmatic ore-forming fluids, the mass transfer calculations below will primarily explore the chemical consequences of direct transfer of a hypothetical tin ore fluid from its magmatic source into a cooler low-pressure depositional environment.

Selected Thermodynamic Data for the System Sn-W-Fe-Na-K-A1-Si-CI-S-C-O-H

The free energy minimization code CHEMIX of the CSIRO-SGTE THERMODATA software package (Turnbull and Wadsley, 1986) was used for all calculations presented in this paper. This package requires representation of thermodynamic data for chemical species in the form of Gibbs free energies of formation from elements, AGp, and provides rou- tines for fitting free energies as a function of temper-

ature, T(øK), to experimentally determined equilibrium constants of balanced reactions. A data base of

free energy functions for selected aqueous species of the form:

AG[(T) = A + BT + CT In T, (2)

was derived from published experimental data within the restricted temperature range of 200 ø to 350øC and at the saturation vapor pressure of H•.O (Table 3; Appendix). Equation (2) is used merely for the pur- pose of representing equilibrium constant data within quoted experimental uncertainties. It does not imply (although it is formally consistent with) an assumption of constant changes in standard enthalpy, entropy, and heat capacity of formation of the species from elements, which is clearly not valid across this temperature range (Helgeson et al., 1981). To minimize errors in the geologic results due to internal incon- sistency, the free energies of species were derived from (or at least checked against) published equilibrium constants of balanced reactions of importance to the following calculations (e.g., feldspar-fluid alkali exchange and hydrolysis equilibria).

Estimated P-T conditions of tin deposition are in many instances somewhat above 350øC and close to the critical point of the pure solvent, H•O, where thermodynamic properties of solutes at infinite dilution vary steeply with minor changes in pressure and temperature (Helgeson et al., 1981). However, tin

TABLE 3. Sources of Thermodynamic Data

Species Reference and remarks

All elements, CO2(g), CH4(g), SO2(g), HCI(g), HaO(g), HaO(/)

Minerals KAISiaOs, KAIaSiaO•0(OH)•, NaAISiaOs, SiOs, FeS2, FeS, FeaOa, Fe304, SnOa, FeWO4, FeCOa

FeaSiO4, NiO Simple aqueous ions H +, Na +, K +, Fe +a, OH- KC1 ø, A12SiaOs(OH)4, SiOn,

NaA1SiaO•o(OH)• KFeaAISiaO•o(OH)a KFeaAISiaO•o(OOH)

FesAlaSiaO•o(OH)s, Fe7SiaO•o(OH)s

HC1 ø

NaOH ø, KOH ø COg, CH•, H2S ø HCOg, HS-, HSO•, SO• •, Hg SnCI +, SnCI•, SnCI5, SnCI• •, SnCIa(OH)•,

SnCI(OH)•, SnCI•(OH)• 2, Sn(OH)g FeCl +, FeCl•

HaWO], HWO•, WO• a

CPDMRLDATA, Turnbull and WadsIcy (1986)

CPDMRLDATA, Turnbull and WadsIcy (1986)

CPDNPLDAT, Turnbull and WadsIcy (1986) Cobble et al. (1982) SOLMIN, M. H. Reed, pers. commun. (1986)

Hewitt and Wones (1984, p. 216) Oxyannite component in biotite; free energy chosen so that fo2 near

Ni/NiO in intermediate annite-oxyannite solid solutions Fe chlorite solid solution, Walshe (1986); enthalpy estimate based on

James et al. (1976) Ruaya and Seward (1987) Lukashov et al. (1975) Drummond (1981) Cobble et al. (1982) Pabalan (1986)

Heinrich and Seward (1990; above 200øC, isocoulombic extrapolation using data from Ruaya and Seward, 1986)

Wesolowski (1984), Jaireth et al. (1990), and S. Jaireth, pets. commun.

Compounds in the system Sn-W-Fe-Na-K-A1-Si-Ni-CI-S-C-O-H considered in mass transfer calculations; superscripts for charge denote aqueous species; sequence in the list corresponds to the order of sequential fitting in the buildup of the data base


ore fluids contain at least 1 m total chloride, which extends the liquid behavior of these solutions to about 430øC or above (Bodnar et al., 1985) and probably reduces the effect of pressure on fluid-mineral equilibria in the 350 ø to 400øC range. Empirical free energy vs. temperature relations of equation (2) have therefore been tentatively extrapolated to 400øC, without explicit consideration of pressure. Empirical observations show that supercritical aqueous equilibria vary primarily as a function of solvent density (e.g., Marshall and Mesmer, 1981), and extrapolated equilibrium constants are therefore believed to approximate the true values at pressures somewhere between the liquid-vapor curve of HeO-NaC1 (270 bars at 400øC) and the isochore of pure water extending from 350øC/Psaturated to 400øC/•'•500 bars (Vn2o = 1.74 cm3/g). Although crude, this assumption is consistent with a recent prediction of Pb speciation in supercritical fluids using more advanced and theoretically based prediction methods (Sverjensky, 1987b). The free energy of HeO itself was extrapolated along the 1.74 cm3/g isochore using data from Burnham et al. (1969). Free energy data for minerals and gases were taken from the file CPDMRLDAT by Turnbull and Wadsley (1986) and from other sources given in Table 3.

Activity coefficients for aqueous ions were calculated from the b-dot extension of the Debye-Hiickel equation of Helgeson (1969), using ft values from Truesdell and Jones (1974). Electrostatic properties of ricO were extrapolated isochorically to 400øC/500 bars (Helgeson and Kirkham, 1974), consistent with the free energy function of HeO. Activity coefficients of neutral aqueous species were set to 1, except for COg where data from Drummond (1981) were used. Fugacity coefficients in the low-density gas phase were ignored in boiling calculations. Biotite and chlorite were treated as ideal binary solid solutions, and all other minerals were considered as pure end mem- bers. In all, uncertainties in calculated metal solubilities are estimated to be on the order of 0.5 log units up to 350øC, but at least one log unit in extrapolations to 400øC.

Hydrothermal Geochemistry of Tin

Tin in the earth's crust occurs mainly in two valence states. Sn(IV) forms the main ore mineral, cassiterite, some aqueous complexes, and probably predominates in granitic melts (Taylor, 1988). Sn(II) complexes are probably the most important species in deep hydrothermal fluids over a wide range of P-T-X conditions (Eadington and Giblin, 1979; Taylor et al., 1984; Jackson and Helgeson, 1985a; Eugster and Wilson, 1985; Pabalan, 1986; Taylor, 1988).

Cassiterite solubility in supercritical H20-HC1- NaC1 fluids was measured by Eugster and Wilson

(1985), who concluded that Sn(II)-C1 complexes predominate at reduced as well as moderately oxidized conditions. Pabalan (1986) measured cassiterite solubility in aqueous chloride solutions between 200 ø and 350øC. His data confirm that Sn(II)-C1 complexes predominate at 350øC throughout the redox stability field of magnetite. Sn(IV)-C1-OH complexes become relatively more important with decreasing temperature and increasingfo2 but do not allow high solubilities of cassiterite except in very acid solutions. This qualitatively confirms earlier extrapolations of low- temperature data (Eadington and Giblin, 1979; Pat- terson et al., 1981; Jackson and Helgeson, 1985a; Heinrich and Eadington, 1986) which predicted somewhat higher solubilities. Sn-F complexes have not yet been studied with thermodynamically controlled high-temperature experiments, but extrapolations of low-temperature data suggest that they are generally less important than C1 complexes, because of the strong predominance of C1 over F even in fluorite- or topaz-saturated magmatic fluids (Eadington and Giblin, 1979; Patterson et al., 1981; Jackson and Helgeson, 1985a, b). Tin solubility is probably dominated by equilibria of the type:

Sn(II)Cl$ -x + 2HeO = Sn(IV)Oe (aqueous) (cassiterite)

+ 2H + + xCl- + H0, (3) and cassiterite precipitation from a reduced acid solution is favored by oxidation (He consumption) or acid neutralization (Pabalan, 1986).

Figure 2 (solid contours) indicates the solubility of cassiterite in a fluid with 2 m total chloride as NaC1

+ KC1, buffered with regard to acidity by the assemblage potassic feldspar + albite + muscovite + quartz (KAbMQ). Cassiterite solubility as Sn(IV) species is very low, and unusually reducing conditions (fo2 < Ni/NiO) or temperatures in excess of 400øC are required to reach Sn(II) solubilities of 10 ppm or more. By contrast, 100 to 1,000 times higher solubilities are supported in feldspar-absent (muscovite + quartz + kaolinitc) assemblages at given T and redox conditions (Fig. 2; broken contours). The experiments of Eugster and Wilson (1985) suggest even lower solubilities in the feldspar-excess field, below concentrations that are realistically required to form an economic tin deposit below 400øC under chemical conditions buffered by the mineralogy of typical tin granites. This indicates that conditions of incomplete chemical buffering of the fluids by feldspathic host rocks are required to transport economic amounts of tin to an ore depositional environment at less than 400øC (see Eugster and Wilson, 1985, p. 99). Incom- plete wall-rock buffering will occur if the flow of hydrothermal fluid through a confined channel (such as


A 0

-4

magnetite -hematite

Sn02 cassiterite

250

2 m C/to t K-feldspar

Albite

Muscovite

Quartz

I•-

300 350 400

Temperature (øC) -- -- 2 m Clto t

(K/K + Na) = O. 25

+ Kaolinite

Muscovite

Quartz

FIG. 2. Temperature vs. redox diagram showing contours for the solubility of cassiterite (SnO2) in a 2-m chloride fluid, calculated from data referenced in Table 3 including experiments by Pabalan (1986). Note that high tin solubilities in fluids buffered with regard to acidity by muscovite + feldspars + quartz are restricted to high temperatures and/or very reducing conditions (solid lines). High solubilities extend to much lower temperatures in more acid fluids in equilibrium with muscovite and quartz alone (limited by the dashed contours for the buffer assemblage muscovite + kaolinitc + quartz).

a vein) is rapid relative to the kinetics controlling the interaction with fresh wall rocks (for example, the diffusion of aqueous components through an exhaustively altered vein selvage; Heinrich, 1986; Heinrich and Eadington, 1986; Holyland, 1987).

Reaction of Magmatic Fluids with Feldspathic Rocks

Fluid-buffered vs. rock-buffered cooling of magmatic fluid

Two limiting cases for the chemical evolution of a single-phase magmatic fluid on its cooling and de- pressurization path from the source to the ore deposit are considered first (Fig. 3; Table 4): cooling of the igneous fluid in equilibrium contact with excess granitic source rock: i.e., rock-buffered evolution; and cooling in isolation from reactive host rocks, such that fluid speciation is controlled entirely by equilibria among aqueous species (and possibly congruent precipitation of minerals): i.e., fluid-buffered evolution.

These two extremes (in contrast to any intermediate cases of partial rock reaction) lead to an equilibrium state of the fluid at any P and T of a depositional site that is independent of the intervening P-T

path. Neither process may be realized in nature, but the rock-buffered evolution is approached by in situ cooling of an igneous fluid in contact with its source granite, which will typically involve a 10- to 100-fold excess of rock over fluid (by weight; Burnham, 1979). The fluid-buffered evolution can be approached by quenching of a fluid or any other physical process of fluid focusing, such that the amount of fluid relative to reactive wall rock is sufficiently high to dominate the equilibrium mineralogy of the rock, without significant modification of the fluid composition by rock interaction. Figure 3 compares the predicted chemical evolution of a fluid in the system Sn-W-Fe-Na-K-A1- Si-C1-O-H for these two cases.

For the rock-buffered evolution of the fluid from

magmatic conditions to 250øC (Fig. 3a, calculated for 1-kg F1 in equilibrium with 30-kg model granite R1; Tables 1 and 2), the initial high concentrations of acidity (HC1 ø, H +) in the fluid and its content of the dominant reducing species (H• ø) drop off steeply with falling temperature. Acidity is consumed by hydrolysis of feldspars to hydrous silicates (muscovite)'

3KA1Si3Os + 2HC1 ø

: KAI3Si30]o(OH)•, + 6SiOn, + 2KC] ø (4)

and

3NaA1Si3Os 2HC1 ø + KC1 ø (feldspars)

: KA13Si30]o(OH)•, + 6SiOn, + 3NaC1 ø. (5) (muscovite) (quartz)

H• ø is used in the reduction of ferric silicate components, e.g., in biotite:

KFe•(II)Fe(III)AiSi30,o(OOH) + 0.SH• ø (oxy-annite)

= KFe3(II)A1Si30]0(OH)•,, (6) (annRe)

or by more complex Fe(III)-Fe(II) reactions involving Ti minerals (Wones, 1981). Beduetion and acid neutralization both promote reaction (3), and under rock- buffered conditions 99 percent of the initial Sn content of fluid F1 will be preeipitated at temperatures above 400øC. After cooling, low concentrations of eassiterite will be dispersed over a large quantity of slightly altered granite.

In the absence of rock-buffering, the high initial concentrations of HC1 ø + H + and of H0 ø, are main- tained throughout cooling, and the acidity and redox level of the fluid are only constrained by equilibria among aqueous species, i.e.:

HCl ø = H + + C1- (7)


(a) Granite-buffered fluid cooling

HO....-" -3

-4

'" ::::• ..... 'i .............

(b) Fluid-buffered cooling

....................................... . .... ,/ H2

H2 ///' .... ß 1000 / $n(tot) ......

? ........ '...•w(,,,) •(t,,t) """ ß ........................... -1

,,o0,) 'o /,//

200 2•0 300 350 400 200 250 300 350 400 Inp•

( Falling Tem•m *C Ig•uS fluid

•c. 3. Predicted chemical evolution of a magmatic fluid (model fluid F1, Table 2) during cooling in equilibrium with an excess of its granitic source rock (a, rock-buffered evolution based on model granite R1, Table 1), compared with the evolution of the same fluid cooling in isolation from feldspathic wall rock (b, fluid-buffered evolution). The latter may be realized by relatively rapid fluid flow in a confined channel (e.g., vein) with limited exposure of fresh wall rock.

and

+ = .H2o. (8) Fluid F1 is close to eassiterite saturation at 400øC, but with falling temperature under fluid-buffered

conditions it becomes increasingly undersaturated with SnO2, because the precipitation reaction (3) is counteracted by increasing hydrogen ion concentration and a constant high Hø• concentration.

The behavior of tin indicated by Figure 3a and b is exhibited by any reduced, acid, and chloride-rich fluid and any alkali feldspar-bearing rock. It indicates that neither simple quenching of a tin-rich ore fluid, nor in situ cooling of a magmatic fluid in contact with an excess of quartzofeldspathic rock, is likely to form an economic tin deposit at the typical temperature range of 300 ø to 400øC recorded by fluid inclusion studies. Figure 3b, however, emphasizes that high concentrations of tin can be transported by a saline fluid to any low-temperature site, provided that the fluid is prevented from chemical interaction with feldspathic rock during parts of its cooling history (Heinrich, 1986).

The predicted behavior of tungsten is different. Wolframite (ferberite) deposition, like that of cassiterite, probably liberates acid (H + or HClø):

H•WO• + FeCIø• = FeWO4 + 2HCI ø, (9) (ferberite)

but this is outweighed by a steep increase of the calculated equilibrium constants of reactions of this type with falling temperature. Wolframite precipitation proceeds over a relatively narrow cooling interval through 350 ø to 300øC, with or without rock interaction (Jaireth et al., 1990). The 100-ppm W solubility limit is indeed slightly extended to lower temperature in the presence of excess granite because of the de-

TABLE 4. Processes ofTin(-Tungsten) Mineralization Tested with Mass Transfer Modeling

Geologic process (model features) Results and conclusions for tin(-tungsten) mineralization

Exsolution of magmatic fluid from granitic melt Rock-buffered cooling of magmatic fluid with excess

granite; no fluid focusing Fluid-buffered cooling and decompression; ___ muscovite

_ quartz but isolated from feldspars

Single-step reaction of magmatic fluid with feldspath, ic rocks at T < 400øC (titration)

Progressive fluid-rock reaction (staged equilibration of fluids along one-dimensional flow path)

Vapor separation at T < 400øC (single-step boiling)

Simultaneous vapor separation and reaction with quartz + muscovite + chlorite or biotite

Presence of additional CO2 or arsenic in cooling magmatic fluid

Mixing of magmatic fluid with epithermal fluid

100-1,000 ppm Sn partitioned into reduced chloride fluid Tin precipitated above 400øC as low-grade dissemination

No cassiterite saturation; effective transport of high Sn concentrations to low temperatures in increasingly acid and reducing fluid; quantitative deposition of wolframite from W, Fe-rich fluid

Cassiterite precipitation by acid neutralization (feldspar --* muscovite + quartz), acid content of fluid (including metal precipitation) restricts Sn enrichment to subeconomic ore grade; titration model and F/R concept geologially unrealistic description of fluid-rock reaction in advective system

Chromatographic enrichment of cassiterite at feldspar hydrolysis front to economic greisen ore (>1 wt % Sn)

Partial precipitation of cassiterite due to oxidation by H• loss; relatively inefficient tin mineralization process

Economic greisen mineralization by combined oxidation (H• loss) and acidity control by Fe silicate reactions

Redox equilibria between H• and CH] (or As(III) and arsenopyrite) may limit Sn solubility; potential mechanisms for fluid-buffered cassiterite precipitation (insufficient data for testing)

Deposition of cassiterite by cooling, decomplexation, oxidation (CO•), and acid neutralization (HCO•); efficient Sn mineralization, no chemical limit on grade


piction of the Fe concentration of the fluid by the formation of Fe silicates (biotite, chlorite) from feldspars and muscovite (Fig. 4). This suggests that the formation of an economic tungsten deposit primarily depends on (1) high-temperature magmatic processes leading to generation of high W concentrations in the fluid (fluid-melt partioning or subsolidus leaching?; Campbell et al., 1984), (2) physical processes of cool-

ing a sufficient quantity of fluid within a restricted ore volume, and/or (3) reaction of relatively Ca-Fe- poor fluids with calcareous (scheelite skarns; Weso- lowski, 1984) or iron-rich wall rocks (Jaireth et al., 1990).

Observations on the contrasting distributions of tin and tungsten at the Panasqueira deposit (Polya, 1988) may be easily explained by the different chemical be-

lOO

• 80 o

60

o 40

.½- 20

0

O -7 ,.J

2

A B C D A

240øC

qtz

ksp •

B C A B

320øC

qtz

1 ,

400øC

qtz

mu$ Io

(a)

(b)

i i i i

'•(c) .....

w

i i i

I 10 100 1000

i i

.1 I 10 100 1000 0 .1 1 10 100 10•

(d)

Ratio Fluid F1 / Granite R1 (bywt.)

FIG. 4. Calculated results of single-step isothermal equilibration of a magmatic fluid (F1, Table 2) with a typical tin granite (R1, Table 1) at three temperatures. The mineralogic composition of the resulting altered rock (row a), the chemical composition of the residual solution (rows b and c), and the tin enrichment in the final rock (row d) are shown as a function of the bulk fluid to rock ratio (by weight; note that these calculations are time and space invariant). Labels A, B, C, D refer to titration steps discussed in the text. Abbreviations (used throughout): alb = albite, bio = biotite, chl -- chlorite, kao -- kaolinite, ksp -- potassium feldspar, mus -- muscovite, qtz -- quartz, tot -- total.


havior of the two elements predicted above. Wol- framitc is the main economic ore mineral at Panas- queira because it is concentrated within thick veins (focused fluid flow with restricted wall-rock contact). Cassiterite occurs only in minor quantities within the veins (mostly with muscovite at vein walls; Kelly and Rye, 1979, p. 175), whereas a much larger quantity of tin (five times the tonnage of mineable tungsten!) is disseminated in a wide halo of uneconomic veinlets

in the semipelitic schists (dispersed fluid flow with intimate rock contact).

Single-step fluid-rock reaction The calculations for the two extreme cases of pos-

sible fluid evolution (Fig. 3) suggest that economic cassiterite mineralization in aluminosilicate host rocks

could form by a process in which a tin-rich magmatic fluid undergoes part of its early evolution above 400øC in effective isolation from excess feldspar and rock redox buffers (rapid quenching and focused fluid flow), and is later exposed to renewed rock interaction at some lower temperature (e.g., in a zone of brecciation), leading to localized cassiterite precipitation by acid neutralization and oxidation (Eadington and Giblin, 1979; Jackson and Helgeson, 1985b).

The conceptually and computationally simplest model for such a process includes a step of fluid-buffered cooling and decompression of a single batch of magmatic fluid, followed by a step of irreversible reaction with an amount of fresh granite at constant temperature (Fig. 4). Mass transfer calculations of this type were introduced by Helgeson et al. (1969) and were first applied to ore-forming systems by Helgeson (1970). If it is assumed that the reaction reaches a state of local equilibrium between fluid and rock at the deposition site, then the reaction progress in these models can be simply expressed in terms of the ratio of input fluid to reactive rock. Titration diagrams of this type have been presented by Reed (1983), Sverq jensky (1987a), and others. It should be noted that these plots have no spatial dimension or direction and are therefore at best qualitatively related to alteration geometry such as mineral zonation.

Calculated results of equilibrating fluid F1 with variable amounts of model granite R1 at three temperatures are shown in the three columns of Figure 4, with the weight ratio of fluid/rock (F/R) plotted on the horizontal axes. Relative amounts of minerals in

the resulting rock, and species concentrations in the equilibrated fluid, are plotted on the vertical axes. The equilibrium states along the left (low F/R) and the right (high F/R) extremity of each diagram cor- respond, respectively, to points on the rock-buffered and fluid-buffered cooling paths shown in Figure 3a and b. Amounts of silicates in Figure 4 (row a) are presented in terms of oxygen units (moles of minerals

multiplied by the number of O atoms in their formula unit) and normalized to 100 percent. This approxi- mates the volume proportions of the phases (modal mineralogy) in the altered rock, because the molar volume of oxides and silicates is approximately proportional to the number of oxygen atoms contained in one mole (Thompson et al., 1982; 12 +_ 1 cm3/O - unit).

With an increasing fluid/rock ratio, the mineralogy of the altered rock changes from feldspars + minor sheet silicates (with partial replacement of albite by potash feldspar), to quartz + muscovite _ kaolinitc and variable amounts of ferrous sheet silicates. The

predicted mineralogic sequence with increasing degree of alteration also varies with temperature (see Fig. 4 for abbreviations):

bio-fsp-mus -• bio-mus-qtz -• mus-qtz at high T, bio-fsp-mus -• chl-mus-qtz at medium T, bio-fsp-mus -• chl-mus-qtz -• (feldspathic) (phyllic alteration,

greisen)

kao-mus-qtz at low T. (argillic)

This is in broad agreement with geologic observation of biotite-bearing phyllic assemblages in relatively high-temperature greisens (commonly early in paragenetic sequence), the widespread occurrence of quartz-muscovite-chlorite assemblages at somewhat lower temperature (frequently associated with a later sulfide stage) and the restriction of kaolinitc to low temperatures (generally late) in many Sn-W systems. The nature and amount of the predicted iron silicate phase is sensitive to small variations in the iron concentration of the input fluid. Doubling Fetotal stabilizes chlorite relative to biotite up to 400øC, and halving Fetotal restricts the stability of ferrous sheet silicates relative to pure quartz-muscovite greisen. Significant iron addition is recorded by mass balance studies of tin greisens replacing granite (e.g., Taylor, 1979, p. 176; Charoy, 1979; Adam, 1983; Stemprok, 1987; Schwartz and Askury, 1989), but part of this may be incorporated in other minerals including muscovite (FeSiAI_2 substitution in phengitic solid solution), zinnwaldite, tourmaline, and iron sulfides (pyrite, pyrrhotite; see below).

The composition of the fluid resulting from interaction at an increasing fluid/rock ratio displays three to four steps (Fig. 4, rows c and d and labels A to D). Point A marks the exhaustion of the redox exchange capacity of the rock, which in R1 is given by its content of Fe(III) hosted by biotite (eq. 6). For most quartzofeldspathic rocks (and particularly for Fe(III)-

HYDROTHERMAL Sn (- W) CHEMISTRY 4 6 9

and Fetotal-poor granitoids typically related to tin deposits; Table 1) reacting with reduced (H2-rich) tin ore fluids, the redox exchange capacity of the rock will be exhausted at relatively low fluid to rock ratios (0.03-1.0), irrespective of the actual Fe(III) mineral.

Substantially higher fluid/rock ratios will generally be required to consume the acid exchange capacity of most feldspathic rocks. This defines the second step in the variation of fluid composition in Figure 4 (point B, rows b and c), which occurs where feldspars are completely consumed by reactions (4) and (5). Given the modal composition of a feldspathic protolith, the F/R at which feldspar exhaustion occurs (point B) is proportional to the total acid content of the input fluid. This consists not only of H + and HC1 ø but also of contributions from all dissolved metals that are precipitated as a result of acid neutralization by reactions producing H + or HC1 ø. This includes tin precipitation according to reaction (3), but the most important contribution is from iron, which forms biotite at the expense of feldspars according to reactions of the type:

3FeCIø• + KA1SiaOs + 4H20

(K feldspar)

= KFeaA1SiaO•0(OH)2 + 6HC1 ø.

(biotite)

(10)

In the present case (F1 with 0.07 m FeC12), and probably in many magmatic fluids reacting with aluminosilicate rocks, the contribution of reaction (10) (or similar reactions involving plagioclase and other Fe minerals) to the total acid budget substantially exceeds the contributions due to the original HC1 ø input, and thereby, determines the fluid to rock ratio (B) where feldspar is exhausted. Note that the arguments regarding acid budget (as H + + HC1 ø) derived from equations (3) to (5) and (9) through 12 (below) are an expression of mass and charge balance only. They do not depend on the actual coordination number of the metal chloride complexes present in solution, which varies as a function of temperature and total chloride ligand concentration. For example, equation (10) can equally be written in terms of3FeC1 + on the left and 3H + + 3HC1 ø on the right side of the reaction.

At F/R above the point of feldspar exhaustion, acidity is controlled by equilibria between aqueous iron and sheet silicates, i.e., in biotite greisens by:

9FeClø• + KAlaSiaO•0(OH)• + 6SiO• + œKC1 ø (muscovite) (quartz)

+ 12H•O = 3KFeaA1SiaO10(OH)• + 20HC1 ø, (11) (biotite)

and in chlorite-muscovite assemblages by:

15FeCl• ø + 2KAlaSiaO10(OH)2 + 3SiO• + 24H•O

(muscovite) (quartz)

= 3FesAl•SiaO10(OH)s + 2KCI ø + 28HC1 ø. (12) (chlorite)

A fourth step associated with further hydrolysis of muscovite to kaolinitc is restricted to highly fluid- dominated conditions and temperatures below 280øC (point D, Fig. 4).

Tin solubility (Fig. 4, row c) is very low at a low fluid/rock ratio (wholly rock-buffered conditions left of point A), but somewhat higher concentrations are supported where the redox exchange capacity of the feldspathic rock is exhausted so that H• ø in the fluid dominates the redox state of the system (A to B; 80 ppm Sn at m H• ø = 0.3, T -- 400øC; 10 ppm Sn at 3œ0øC). Once feldspar is exhausted (F/R above point B) and acidity as well as redox level becomes increasingly fluid-dominated, the tin concentration remains at the initial 500 ppm without cassiterite saturation at any temperature.

While feldspar remains in the equilibrium assemblage, aqueous tin is almost quantitatively extracted as cassiterite. The tin ore grade in the rock (Fig. 4, row d) therefore increases nearly proportionally with the fluid/rock ratio, to reach a maximum near point B where feldspars are just consumed. The magnitude of this maximum, i.e., the highest ore grade that can be achieved by the present model process of acid neutralization by feldspar hydrolysis, is therefore in- versely related to the total acid content of the fluid. Given that the latter is dominated by aqueous iron species (eq. 10), a reduction of the Fe concentration of the incoming fluid would increase the maximum possible cassiterite enrichment by allowing a higher F/R for point B. However, the solubilities of cassiterite and of the Fe silicates are coupled at given acidity and chloride (and H•) concentrations. A series of similar calculations using variations in input fluid compositions (Table 2, right) indicated that maximum tin enrichment cannot be raised substantially above the value of 0.35 wt percent shown in Figure 4d. An alternative assumption, of feldspar-buffered cooling to 400øC followed by fluid-buffered cooling to some lower temperature and renewed feldspar interaction, results in even lower grades of tin enrichment.

A maximum tin enrichment of 0.35 wt percent is probably insufficient to produce an economic tin ore- body, considering that bulk mineable ores are typically inhomogeneous and are likely to contain a predominant proportion of rock affected by fluid to rock ratios that were higher or lower than the optimum value. Bulk equilibration of a magmatic fluid with a


feldspathic rock (at least in the simple chemical system Vluid considered so far) therefore appears as a rather un- batches likely and probably inadequate model for the formation of an economic tin deposit. 3

Progressive fluid-rock reaction at an alteration front

The titration calculations depicted in Figure 4 point out general mass balance and thermodynamic relations and are useful for determining the dominant chemical reactions in a hypothetical fluid-rock reaction process. However, the concepts of fluid/rock ratio and single-step equilibration of one batch of rock with one large batch of fluid clearly do not allow a geologically realistic description of hydrothermal ore formation by advective mass transfer. Fluid-rock interaction in a dynamic (flow-through) system is progressive, in the sense that it involves repeated over- printing of a rock volume by sequential smaller batches of fluid, each of which has itself been chemically modified by previous steps of fluid-rock reaction (Korzhinskii, 1970). Reaction progress along a flow line of hydrothermal fluid passing through reactive rock has been described by simultaneous calculation of fluid flow and chemical reaction kinetics (e.g., Lichtner, 1988; and Baumgartner and Rumble, 1988, for stable isotope exchange). Similar models have been applied to specific situations of hydrothermal ore formation, to obtain chemical predictions that include time and space as explicit variables (e.g., Merino et al., 1986; Holyland, 1987; Ague and Brimhall, 1989). Time-specific information on in situ permeability and reaction rate constants are difficult to estimate for

high-temperature magmatic systems. However, even without such data, a one-dimensional advection model assuming local equilibrium between successive small batches of fluid and segments of a column of reactive rock along a fluid flow line (as indicated by Fig. 5) can be used, which affords a geologically more realistic description of hydrothermal alteration than single-step titration calculations.

Figure 6 shows part of the results of a one-dimensional local-equilibrium flow path calculation, based on identical input fluid and rock compositions as used in the calculation of Figure 4 (center column, 320øC). A total of 0.2 kg of rock R1 is subdivided into ten segments along a linear array. 1 kg of fluid F1 is partitioned into ten batches, which sequentially equili- brate with each segment of the rock column. Figure 6 shows the states of each rock segment after the passage of one, six, and eight of the ten fluid batches, and the concentration of some components in the fluid after each equilibrium contact.

Reaction of the first batch of fluid with rock segment 1 (Fig. 6a) is identical to single-step equilibration

Rock column, 6 segments

Progressing alteration front

Fluid

batch

FIG. 5. Cartoon showing the principle of a finite element, local equilibrium model for progressive fluid-rock interaction along a one-dimensional fluid flow path through a permeable rock (see Fig. 6, below). This model probably affords a geologically more realistic description of an advancing alteration front (or progressive alteration in a confined channel such as a breccia pipe) than the single-step equilibration (titration) depicted in Figure 4.

(near point B of feldspar exhaustion and maximum tin enrichment of 0.35 wt % shown in Fig. 4). This ex- tracts about 60 percent of the tin contained in fluid batch 1. The remaining 40 percent (which would be lost in the single-step process, Fig. 4) is deposited almost quantitatively in rock segment 2. After this, the metal and acid content of the fluid is exhausted

and the fluid composition and mineralogy of the sub- sequent rock segments passed by fluid batch 1 remain unchanged.

At a more advanced stage after passage of the sixth fluid batch (Fig. 6b) a larger number of rock segments are exhaustively altered to chlorite q- muscovite q- quartz, and tin becomes progressively enriched near the alteration front of feldspar consumption. This is achieved by dissolution of cassiterite at the back of the ore zone (where the incoming acid fluid is undersaturated with cassiterite), a transient increase in the fiuid's tin concentration, followed by quantitative deposition where the fluid encounters unreacted feldspar-bearing rock. After passage of several fluid batches through the system, the grade of cassiterite enrichment in one rock segment reaches a maximum or plateau value after which the ore zone still widens but no longer increases in ore grade (Fig. 6c).

The maximum tin enrichment achieved by this process is still limited by the mass balance between advected acidity and the acid exchange capacity of the protolith, but it is three to five times higher than the maximum ore grade resulting from single-step equilibration, at given compositions of input fluid and host rock. The predicted higher chemical efficiency and the greater physical reality of the advection model suggest that this chromatographic enrichment process could be a widespread mechanism for the formation of economic greisen-hosted tin deposits. A similar process may also occur in extended but laterally confined breccia pipes draining a granite pluton. Ore reworking may be realized by local reequilibration of


(a) 1/10 of 0.2 kg fluid passed lOO

(b) 6/10 of 0.2 kg fluid passed (c) 8/10 of 0.2 kg fluid passed

[] qtz

[] mu$

[] ksp

[] alb

[] bio

ß chl

•_ t)

O) -2

=o -3

Sn(tot) _•

H+(tot)

''T

.c: Cassiterite 0.5

0.0

I 2 3 4 5 6 7 8 9 10 1 2 3 4 5 6 7 8 9 10 I 2 3 4 5 6 7 8 9 10

Distance (rock segment) • Distance (rock segment) • Distance (rock segment)

FIG. 6. Progressive alteration and cassiterite mineralization of a granite by a magmatic fluid of constant input composition, as predicted by the flow-line model sketched in Figure 5 (calculated with the STAGE option of program CHEMIX; Turnbull and Wadsley, 1986). Ten batches of fluid of identical starting composition (F1, Table 2) successively flow from left to right through a column of granitic rock (R1, Table 1) subdivided into ten segments. The three columns of diagrams show the state of all rock segments and coexisting fluid compositions after passage of the first, sixth, and eighth fluid batch. An advancing alteration front of feldspars reacting to muscovite q- quartz + chlorite develops, which is associated with progressive cassiterite enrichment by dissolution and reprecipitation. This results in tin ore grades that are considerably higher than the maximum ore grade achieved by a single-step fluid- rock reaction (Fig. 4).

the fluids with previously altered rocks in the lower portions of the pipe or in intensely greisenized apical portions of the intrusion (Fig. 1). Possible examples include the chlorite- or tourmaline-rich breccia col-

umns at Ardlethan (Eadington and Paterson, 1984; Ren and Walshe, 1986), in the Herberton district (Taylor, 1979), or at Zaiplaats (Andrew et al., 1986).

Vapor Separation

Two-phase fluid separation or boiling can influence mineral solubilities due to preferential fractionation of components between the liquid and the vapor phase (e.g., Reed and Spycher, 1984; Drummond and Ohmoto, 1985; Spycher and Reed, 1986). Ore mineral deposition may occur due to ligand loss (e.g., S in the case of gold bisulfide complexes; Seward,

1974), shifts in acidity due to preferential loss of weak acids (e.g., CO2; Giggenbach, 1981; Henley and Ellis, 1983), or changes in redox level (loss of C-S-N-O-H gases; Giggenbach, 1980). Ignoring possible gaseous species of tin (which has been detected in volcanic gases; Symonds et al., 1987) effects of boiling on cassiterite solubility in an Sn-Fe-Na-K-Si-C1-O-H fluid are primarily due to oxidation resulting from preferential loss of the main reducing species, Hg. Figure 7 shows the calculated compositional evolution and tin solubility in model fluid F1, assuming single-step closed-system isothermal vapor loss at 380øC, and Figure 8 summarizes similar calculations for other temperatures. Two hypothetical processes are considered.

If the magmatic fluid is quenched without any rock interaction prior to boiling at 380øC (open symbols


Fluid composition •(tot)

o

• H2(-q) Fa(lol)

• -1

E H*(IOl)

o • .2 .J

OO04

0OO0

Minerals A • • a

, Ce*,titaritl 4 M% Sn

.,,,,, ,•in rock c,,,i,•;ii; .....................................

20 40 60

Water Loss to Vapor (percent H20)

10o

G.

FIG. 7. Isothermal boiling offiuid F1 at 380øC, with and without simultaneous fluid/rock reaction. Broken lines and open symbols -- evolution of residual liquid as a function of single-step closed-system vapor loss (in % of initial H20); full lines and symbols -- vapor loss from a fluid in equilibrium contact with muscovite + quartz + an Fe mineral (chlorite or biotite), equivalent to a high F1/R1 ratio • 100. The higher efficiency of tin extraction of combined boiling and fluid-rock reaction is due to acidity control by equilibria between aqueous iron and Fe silicate.

equilibrium contact with a greisen assemblage of muscovite + quartz + chlorite or biotite during cooling and boiling, its acidity is controlled by equilibria (11) and (12) between aqueous iron and ferrous silicate. This acidity control (represented in Fig. 4 by the region at high F/R between B and C) is insufficient by itself to precipitate cassiterite from a reduced tin- rich fluid, but in conjunction with the oxidizing effect of fluid boiling provides a potentially significant precipitation mechanism for cassiterite. Compared with simple boiling, the precipitation mechanism of combined vapor separation and wall-rock reaction is effective over a wider temperature range, and the efficiency of tin extraction at any temperature is en- hanced (Figs. 7 and 8b). The process leading to point A in Figure 7 (point of chlorite exhaustion, and maximum tin extraction after 40% H20 loss) produces a rock consisting of muscovite + quartz + cassiterite with a tin content of 4 wt percent. This ore grade is more than an order of magnitude greater than the maximum value (0.35%) that can be obtained by single-step feldspar titration of an identical ore fluid with the same feldspathic protolith (Fig. 4).

Additional Fluid Components

The calculations presented so far are based on a highly simplified chemical system. Nevertheless, given the assumptions and the thermodynamic data for all compounds in the simple system, the results with regard to cassiterite mineralization will not be altered significantly by any additional fluid species unless these are present in comparable or larger concentrations than those of the dominant acid and redox

species of the simple model system. If the high concentrations of HC1 ø and H• assumed so far (Table 2) are indeed required for effective tin transport, then the possibilities for competing components are quite restricted.

in Figs. 7 and 8a), it will reach cassiterite saturation after 8 percent H20 loss, as a consequence of the oxidizing effect of preferential loss of H2. On further boiling, up to one-third of the tin in solution is precipitated, but then the tin solubility increases again as a consequence of rising concentrations of total ligand, Cltotal, and acidity, Htotal, due to the loss of solvent. Vapor loss without concomitant rock reaction is a moderately efficient cassiterite precipitation mechanism at 400øC where the fluid is already close to saturation prior to boiling, but with decreasing temperature it becomes less efficient because of increasing H + activity (eq. 7; Fig. 8a). Simple fluid boiling therefore does not appear to be an effective tin- mineralizing process, all the more as it will generally be associated with cooling (adiabatic boiling).

Alternatively, if the same starting fluid remains in

(a)

- / 16o% Fluid bolling without / rock Interaction

ø'øø3 1 ?*C I ""1

:• o.ooo I r • , . , 3•, •o 0 20 40 60

Water Loss to Vapor

(b)

Fluid bolling + rack reaction 400 1 100 ' 36O

340 6O

)

r/ (qtz+m"+bl•hl) / 2O

(percent H20)

FIG. 8. Influence of temperature on the two boiling processes compared in Figure 7. H2 loss due to fluid boiling alone (a) is only effective in precipitating cassiterite if the fluid is already close to saturation (400øC in this case). In the more realistic case (b) of fluids reequilibrating with a greisen assemblage containing an Fe mineral (chlorite, biotite, or Fe sulfide), the oxidation effect associated with vapor loss contributes to a more effective cassiterite- precipitating mechanism over a wider temperature range.


Aluminum and boron

In all calculations and reactions presented above, aluminum was assumed to be conserved among rock- forming silicates. This is not entirely consistent with evidence for at least some AI mobility during greisen- type alteration (e.g., Taylor, 1979; Stemprok, 1987). At near-solidus temperatures, polymetallic A1-Na-K hydroxy complexes are probably present in solution (Anderson and Burnham, 1983), and species of this type could potentially neutralize part of the high acidity of magmatic fluids during cooling and decompression. However, experiments by Woodland and Walther (1986) suggest that their concentration in the model ore fluid F1 is significantly smaller than its estimated HC1 ø content. The omission of aqueous aluminum from the present calculations, for lack of thermodynamic data and knowledge of its complex stoichiometry and concentration at high T, should therefore not significantly alter the results discussed so far. Moreover, the observation of ubiquitous phyllic (acid) alteration associated with tin deposits provides strong geologic evidence that the high initial HC1 content of the ore fluids outweighs their total content of alkaline species such as A1 hydroxy complexes.

Boron could influence the acid balance of fluid-

rock interaction in tin deposits, if it is present in high concentrations comparable to HC1 and FeCb.. If tourmaline is relatively insoluble up to high temperature, its precipitation could deplete the fluid in iron close to the source and thereby favor higher tin enrichment by phyllic alteration at lower temperatures (greisen ores). Again, there are not enough experimental data to test this quantitatively.

Carbon species

Redox effects: Many cassiterite-related fluid inclusions contain CO•. as the main gaseous component, besides comparable or smaller concentrations of CH 4 . At temperatures near the granite solidus, the C-O-H components of a supercritical fluid are dominated by H•.O, CO•., and H•.. With falling temperature, CO•. and H•. react to CH 4 according to:

CO• ø + 4Hø• = CH4 ø + 2H•.O, (13)

(Ohmoto and Kerrick, 1977) down to about 250øC, below which temperature the reaction becomes ki- netically inhibited (Giggenbach, 1980). This equilibrium probably exerts the main redox control in relatively reduced magmatic fluids, if they contain a total CO•. concentration comparable to or larger than that of H•. (Burnham and Ohmoto, 1980). Figure 9 shows the equilibrium evolution (without rock interaction) of a fluid similar to F1 (Table 2) with additional CO• concentrations of 0.075 moles/kg H•.O (just bal- ancing the 0.3 m H• ø according to eq. 13; Fig. 9a) or

1

'• -2 E

-4

(a)

mC02 = 1/4 mH2

• CH4 •

H2 'C• Sn(tot)

Temperature øC Temperature øC

F•C. 9. Schematic diagram indicating the likely variation of cassiterite solubility in COg-bearing fluids at high temperature, for fluid F1 with addition of (a) 0.075 ra COg and (b) 1 ra COg. Redox conditions in the absence of mineral buffers are dominated by C-O-H equilibria; the consumption of H2, by reaction with COg to CH4, with falling temperature may lead to cassiterite precipitation without any fluid-rock interaction.

1 mole/kg (COø• in excess over Hø•; equivalent to Xcog "• 0.02; Fig. 9b). The extrapolation above 400øC is schematic, but qualitatively shows the shift of equilibrium (13) to the right with falling temperature. The consumption of Hø• effectively oxidizes the fluid, and as a consequence, a major fraction of its initial tin content is precipitated as cassiterite, at temperatures well above 400øC, by reactions of the type:

4SnCIø• + COø• + 6H•.O = 4SnO•. + CH• + 8HC1 ø, (14)

i.e., an HE-balanced combination of equations (3) and (13). As with the case of H• ø loss by vapor separation (Figs. 7 and 8), tin solubility as a function of temperature passes through a minimum. In presence of excess COø•, cassiterite precipitation is associated with the crossover from Hi as the dominant reduced species at high T, to CH4 ø at lower T (Fig. 9b, arrow). The temperature of this crossover is quite sensitive to uncertainties in the thermodynamic data and activity coefficients for the C-O-H species. For the calculation of Figure 9, data for aqueous species derived from Henry constant measurements below 300 ø to 400øC were extrapolated. This is certainly not valid up to 500øC where the relations will be strongly pressure dependent (in a low-density gas with H•.O > CO•. > H•. the crossover is shifted to lower temperature). Even though the high-temperature solubility of cassiterite according to equilibrium (14) cannot be fully evalu- ated, the available data suggest that it may significantly limit hydrothermal transport of tin in CO•.- bearing magmatic fluids. On the other hand, reaction (14) might offer a possible mechanism of cassiterite precipitation by simple cooling of a single-phase fluid, without any wall-rock interaction, which is not possible in the C-free system discussed in the previous sections.

Acid-base effects: In the lower temperature, lower salinity, and less acid fluids typical of geothermal sys-


tems and epithermal precious metal deposits, carbonic acid and bicarbonate ion constitute the dominant acid-

base pair in the fluid (Giggenbach, 1980, 1981; Hen- ley and Ellis, 1983). They control solution acidity by the equilibrium:

COg + H20 = H + + HCOg (15)

or by the carbonate-silicate equilibria of the type:

CO• + (anorthite)

= (calcite) + (sheet silicates) + H + (16)

(Giggenbach, 1981). In those fluids, vapor separation leads to loss of the dominant acid (CO• ø) in preference over the base (HCOS), which strongly affects solution acidity and thereby mineral solubility (Drummond and Ohmoto, 1985; Spycher and Reed, 1986). How- ever, the importance of CO2 for the acid balance of the typically more saline, hotter, and more acid magmatic fluids associated with tin deposits is diminished for a combination of reasons (Fig. 10).

The dominant acidity buffer within a fluid can be considered as that acid-base pair of which both species are present in concentrations that outweigh the concentrations of any other acid-base pair. In low-temperature, near-neutral, and/or low-salinity CO•- bearing solutions this is generally the case for CO• ø and HCOS, but with increasing temperature, the dissociation constant of equation (15) becomes smaller (i.e., CO• ø becomes a very weak acid) so that the concentration of HCO5 becomes very low. Conversely, the dissociation constant of HC1 ø according to equation (7) is high at low T, but it decreases with increasing temperature so that HC1 ø and C1- outweigh

450 I •' + ksp I • '. HCIø/Cl' dominant + alb | '. • ' - acid/base pair + mus

[ c02O/HCO '' ' 3. e E 300 q dominant . I • t- / acid/base pair In fluid 0.1 . me02

25o / -2.0 -•'.0 010 110

log mCl(tot) (mol/kg fluid)

FIG. 10. Diagram of temperature versus total chloride molality, delimiting the two regions where COø•-HCO• (eq. 15) and HC1 ø- C1- (eq. 7) are the predominant acid-base pairs in fluids with 1 m CO•,o,• (ca. 2 mole % = solid line; 0.1 m and 3 m CO• indicated by dashed lines), buffered with regard to acidity by muscovite, quartz, and feldspars (KAbMQ). Upon CO•-1oss by vapor separation, fluids originating from the CO•-HCO• region may become more alkaline, whereas typical tin-transporting fluids originating from the HCIø-CI - region will not experience significant changes in acidity and hence cassiterite solubility.

HCO• in hot, saline, and/or acid fluids even in the presence of high CO• concentrations. Figure 10 shows the approximate boundary in temperature versus salinity space, between the fields of dominance of the CO•ø-HCO• and HCIø-C1 - acid-base pairs at a given total CO• concentration. The boundary was calculated for an Na-K-C1-C-O-H fluid in equilibrium with K feldspar q- albite + muscovite q- quartz, to summarize results from calculations using more complex fluids. Upon boiling (after isolation from the minerals) the acid balance of fluids originating from the HClø-C1 - field are not significantly affected by CO• loss to the vapor, because the dominant acid, HCI ø, does not strongly partition into the vapor phase. In contrast, fluids in the CO•ø-HCO• field may experience a significant acidity shift as a consequence of preferential loss of their dominant acid, CO•, unless other acid/ base pairs are present in higher concentration (e.g., HSO•-SO•2; cf. Drummond and Ohmoto, 1985, figs. 5-8).

Magmatic fluids capable of retaining high tin concentrations down to temperatures below 400øC have to be either highly saline or one to two orders of magnitude more acid than the KAbMQ buffer (Fig. 4). This further reduces the field of dominance of CO•-HCO5 as an acidity buffer within the fluid, equivalent to a shift of the curves in Figure 10 to the left by 1 to 2 log concentration units. This indicates that CO• loss by fluid boiling is not an important pro- moter of cassiterite deposition, but that it could aid precipitation of wolframite from a low-salinity CO•- rich magmatic fluid (Higgins, 1985).

Sulfur

Most tin deposits contain at least minor amounts of iron (and/or base metal) sulfides. Traces of pyrite or pyrrhotite also occur in most of their parent granitoids, but it is uncertain whether they represent sat- urating solidus minerals. The initial sulfur content of magmatic tin ore fluids is therefore difficult to estimate and probably varies considerably among individual systems. Under the reducing conditions that favor partitioning of tin into the magmatic fluid, H2S ø is expected to be the predominant sulfur species in solution.

The addition of H•S has a relatively minor effect on the results of calculations similar to the ones presented in Figures 3 to 9, except that the iron silicates (chlorite, biotite) are replaced by pyrrhotite or pyrite under certain conditions. Fluid-rock titration calcu-

lations similar to Figure 3 with fluids containing H•S of twice the concentration of FeC12 precipitate iron and sulfur almost quantitatively as pyrite at very low fluid/rock ratio, where acidity as well as redox conditions are rock buffered:

FeClz ø + 2H2S ø: FeSz + 2HC1 ø + H• ø. (17)

HYDROTHERMAL Sn (- W) CHEMISTRY 475

Pyrrhotite deposition at 400øC is restricted to intermediate fluid/rock ratios where acidity is feldspar buffered but the fluid remains reduced due to its high H2 concentration:

FeCI• + H2S ø = FeS + 2HC1 ø. (18)

At temperatures below 350øC, pyrrhotite replaces biotite and chlorite throughout the range of F/R where redox levels are low (compare Heinrich and Eadington, 1986, Fig. 5a and b).

The precipitation of one mole of iron as sulfide liberates 2 moles of acidity (eq. 17 and 18), just like its precipitation as ferrous silicate (eq. 10). The acidity balance of alteration processes and its consequences for cassiterite ore grade discussed in previous sections are therefore not significantly altered by additional H•S, even if it is present in high concentrations.

Preliminary calculations considering calcium, fluorine, arsenic, and base metals (Cu, Pb, Zn) indicate that none of these additional components is funda- mentally likely to alter the general relations discussed above. However, they provide additional constraints on fluid evolution in individual ore systems where combined paragenetic information and fluid inclusion data are available. They may also impose alternative acidity controls, e.g., fluorite deposition in Ca,F-rich systems (Jaireth et al., 1990), or redox controls. Heinrich and Eadington (1986) suggested that coprecipitation of arsenopyrite, FeAs(-I)S, from As(III)- rich solutions could dominate the oxidation potential and hence cassiterite deposition in certain systems, but insufficient experimental data are available to test this further.

Fluid Mixing

Mixing between magmatic and external fluids is well documented in a number of tin-depositing systems (Kelly and Rye, 1979; Davis and Williams-Jones, 1985; Sun and Eadington, 1987; Witt, 1988) and could be important in many more. Little is known about the chemical composition of the nonmagmatic fluid, except that it is cooler and of low salinity (correlations between fluid inclusion temperature and salinity being the main evidence of fluid mixing) and that its water is largely of meteoric origin. Fluids in continental geothermal systems, associated with intermediate to acid volcanics and inferred intrusions

at depth, such as the Taupo zone (New Zealand) or Yellowstone, probably represent the nearest recent analogues from which supplementary indications on the composition of external fluids in Sn-W ore systems may be inferred.

Mixing of granite-derived tin ore fluids with oxidized sulfate waters would cause a dramatic reduction

of cassiterite solubility. This might be a significant mechanism of cassiterite deposition in subvolcanic

(acid sulfate waters) or possibly in marine-exhalative (seawater sulfate) tin ore-forming environments.

In the mesothermal regime of typical tin veins and breccias, mixing of magmatic fluid with moderately reduced fluids comparable to those in deeper parts of active geothermal fields (and epithermal ore deposits) is more likely. There is strong evidence that the chemical composition of deep low-salinity fluids from the Taupo geothermal fields in New Zealand indeed originates from minor admixture of magmatic fluid to a large convecting body of surface-derived waters (Giggenbach, 1981; Henley and Ellis, 1983) and a similar process is consistent with observations from fossil tin systems (e.g., Eadington, 1983).

Compared with the magmatic ore fluid, most geothermal fluids contain lower concentrations of nearly all species considered in the previous calculations (Henley and Ellis, 1983, table I). Carbon species represent the only component occurring in most geothermal fluids at concentrations comparable to those in the tin ore fluid F1 (Table 2). Carbon speciation in the deep Taupo fluids is buffered by carbonate- silicate equilibria (eq. 16; Giggenbach, 1981), and carbonate saturation is also typical for the paragenetically late, meteoric water-dominated, stages in mesothermal cassiterite deposits (e.g., Kelly and Tur- neaure, 1970; Kelly and Rye, 1979; Andrew and Heinrich, 1984). Low-salinity geothermal fluids are typically near-neutral or slightly alkaline, so that the concentration of HCO5 is significant (Fig. 10). The deep fluid of the Rotokawa system in the Taupo region (Krupp and Seward, 1987, table 2, RK4a) is estimated to contain between 3 X 10 -4 m and 2 X 10 -3 m HCO5 at depth (calculated, respectively, from tabu- lated pH = 5.75 at 320øC, or assuming equilibrium with K feldspar + albite + muscovite + quartz). The higher bicarbonate value was adopted for the calculation of Figure 11, which predicts the results of mixing the magmatic CO2-free model fluid F1 (600øC) with variable proportions of a simplified Rotokawa fluid (in moles/kg H20 at T = 320øC): CO• -- 0.280, NaHCO3 = 0.002, CH4 = 0.009, NaC1 = 0.019, and KC1 = 0.003. For comparison, the effect of simple dilution of the magmatic fluid by pure water of 320øC is indicated by stippled lines in Figure 11. The temperature of the mixture was assumed to vary linearly with the mixing ratio (Fig. 11a), i.e., differences in heat capacity of the two fluids and any heat exchange with rock were ignored (cf. Henley and McNabb, 1978). The calculations are restricted to temperatures below 400øC, i.e., to mixing ratios of magmatic/meteoric fluid of less than 0.3, which is realistic in light of the observations from geothermal fields.

The calculations indicate that fluid dilution alone

(stippled; Fig. 11) may lead to substantial precipitation of cassiterite, due to lowered acidity and total ligand concentration (decomplexation of Sn-C1 spe-


ø0600{ 500 t 300 I ß , ß , . , ß ,

(b)

._x E cl(tot)

0.4 ••OC 02 0.2 (tot)

(c)

(no C02)

(d)

o.oool I ' ,

.,,.

._x ,:: E

.... o

• HC03'

0 0000 . , . , ."'T'"

100 --;

extraction

', 90 (

0 20 40

Cassiterite

efficiency

(e)

, , ß

60 80 100

Percent magmatic fluid in mixture

FIG. 11. Predicted results of mixing a hot saline tin-rich magmatic fluid (F1, Table 2) with a cooler low-salinity CO2 bicarbonate solution, such as the deep meteoric fluid of the Rotokawa geothermal field (Krupp and Seward, 1987). The stippled lines indicate the effect of dilution of the magmatic fluid with pure water of the same temperature (320øC) for comparison. Partial cassiterite precipitation occurs by acid dilution and decomplexation of stannous chloride species (stippled), but the oxidizing and acid- neutralizing effect of COøe and HCO•, respectively, greatly en- hance the tin extraction efficiency, particularly at low ratios of magmatic to meteoric fluid (full lines).

cies). However, up to 20 percent of the initial 500 ppm Sn will remain in solution and be dispersed in the absence of carbon species. By contrast, tin is almost quantitatively deposited upon mixing with the feldspar-saturated geothermal fluid containing carbon species in the concentrations above. This is partly due to the oxidizing effect of CO2 (eq. 14; Fig. 11c) and partly due to the acid neutralization of HC1 ø by HCO• (Fig. 11d; low mixing ratios of magmatic to geothermal fluid). The model calculations suggest that minor admixture of granitic tin ore fluid into the base of a convecting epithermal environment should be a very favorable mechanism for the formation of a mesothermal cassiterite ore deposit.

Summary and Conclusions

A review of geologic evidence on mesothermal cassiterite veins, breccias, and greisen-type replacement bodies indicates that these ores are formed by saline fluids, which are initially equilibrated with a high-temperature (>400øC) granitic source, and which then react irreversibly with rocks _ fluids in a cooler depositional environment. Thermodynamic mass transfer calculations based on published experimental data show that high tin concentrations (hundreds or possibly thousands of ppm) can be transported as Sn(II)-C1 complexes to any low-temperature (<400øC) deposition site, provided that chemical buffering of the cooling fluids by feldspathic wall rocks is prevented by relatively rapid fluid flow through a confined conduit such as a vein. Several alternative mechanisms to saturate cassiterite, Sn(IV)O2, at temperatures below 400øC are consistent with geologic evidence, but most of these are subject to severe thermodynamic and mass balance constraints on the degree of tin enrichment, and only a few processes are likely to accumulate economic primary ore grades in aluminosilicate host rocks (Table 4).

Acid neutralization by feldspathic host rocks through hydrolysis to muscovite + quartz _ biotite _ chlorite _ Fe sulfides (phyllic or greisen alteration) is an effective tin extraction mechanism, but the maximum ore grade achieved by this is limited by the acid exchange capacity of the rock (essentially given by its feldspar content) relative to the total acidity advected by the fluid. The latter consists not only of HC1, but includes all metals contributing acidity by their precipitation reactions, notably iron deposited as ferrous sheet silicates or sulfides. Acid neutraliza-

tion by feldspars may be insufficient to produce an economic tin deposit, except by multistage enrichment at an advancing (chromatographic) alteration front or in a confined breccia pipe.

Fluid-boiling or vapor loss probably has little effect on the acid balance of a typical tin ore fluid even if it is COe-rich, because HClø/C1 - is the predominant acid-base pair, rather than COeø-HCO• as in most

HYDROTHERMAL Sn(- W) CHEMISTRY 4 7 7

cooler lower salinity geothermal-epithermal fluids. Vapor separation may nevertheless lead to effective cassiterite mineralization through the oxidizing effect of H2 loss, provided that boiling is associated with simultaneous fluid/rock equilibration with a greisen assemblage (e.g., muscovite + quartz + Fe chlorite or biotite). More than ten times higher ore grade can be achieved from a given fluid and protolith, compared with single-step feldspar hydrolysis.

Fluid mixing of small proportions of magmatic tin fluid with cooler low-salinity meteoric fluid is an ex- tremely efficient mechanism of tin extraction, es- pecially if the meteoric water has near-neutral acidity and contains CO2-HCO•, like many active geothermal fluids. Coprecipitation of gangue minerals (e.g., quartz in a fissure vein) places the only chemical limitation on tin ore grade in this case, in contrast to any alternative mechanism of cassiterite deposition in aluminosilicate host rocks. Dilution by quartz + sulfides is also the main chemical limitation on tin en-

richment in carbonate replacement deposits such as Renison Bell, where cassiterite is precipitated by acid neutralization associated with carbonate dissolution

(Patterson et al., 1981; Holyland, 1987; not further modeled here).

The principal uncertainties and shortcomings in the geochemical predictions presented above are probably the following. First, the assumed quenching and decompression of ore fluids by direct transfer from a magmatic source to a low P-T depositional environment clearly is an oversimplification of the real geologic process. The simplified treatment is largely dic- tated by the sparsity of experimental data available to quantify the speciation of saline fluids in the 400 ø to 600øC and 300 to 1,000 bars region as a function of temperature and pressure (both of which will be equally important in this transitional P-T region), the virtual absence of experiments on metal partitioning between coexisting medium-salinity (_CO2-rich) vapor and high-density salt brines, and also by the lack of detailed geologic evidence for the P-T evolution of magmatic fluids prior to their arrival at the deposition site.

Second, it is concluded that iron species exert the prime control on the acid balance of feldspar alteration reactions and tin enrichment in aluminosilicate-

hosted cassiterite ores, and yet the aqueous Fe species and particularly the ferrous sheet silicates are among the thermodynamically least characterized species in the calculations; these ferrous compounds clearly re- quire additional experimental study.

Finally, improved models and more specific applications to individual deposits should test the effects of additional components. For this, further experimental data on the aqueous speciation of aluminum, arsenic, and fluorine (including tests on Sn-F complex stability) and on the effect of CO2 on cassiterite solubility would be of particular interest.

Acknowledgments

This study has been stimulated by discussions with Mike Solomon, who has provided important ideas and has greatly contributed with his broad geological knowledge of tin and tungsten deposits (but who is not to blame for any undue generalizations). I would also like to thank Dick Henley, Neville Higgins, Sub- hash Jaireth, Jeff Taylor, Vic Wall, and John Walshe for stimulating discussions, helpful comments, and access to unpublished information. Particular thanks are due to Alan Turnbull for his advice with the use of the CSIRO-SGTE THERMODATA software. The

manuscript has greatly benefited from thoughtful and detailed Economic Geology reviews. Published with permission of the Director, Bureau of Mineral Re- sources, Geology and Geophysics.

February 16, November 28, 1989 REFERENCES

Adam, J., 1983, The Elsemore greisen, northern N.S.W.: Unpub. B. So. thesis, Armidale, Univ. New England, 85 p.

Ague, J. J., and Brimhall, G. H., 1989, Geochemical modeling of steady state fluid flow and chemical reaction during supergene enrichment of porphyry copper deposits: ECON. GEOL., v. 84, p. 506-528.

Anderson, G. M., and Burnham, C. W., 1983, Feldspar solubility and the transport of aluminuin under metamorphic conditions: Ain. Jour. Sci., v. 283-A, p. 283-297.

Andrew, A. S., and Heinrich, C. A., 1984, Isotopic and fluid inclusion evidence for sources of inineralizing fluids at the Sun- down tin prospect, S. E. Queensland labs.l: Geol. Soc. Australia Abs., v. 12, p. 36-38.

Andrew, A. S., Pollard, P. J., Taylor, R. G., and Groves, D. I., 1986, Closed-system behaviour for the tin-tungsten system at Zaaiplaats, South Africa: Genesis of tin-tungsten deposits and their associated granitoids labs.l: [Australia] Bur. Mineral Re- sources Rec. 1986/10, p. 5-6.

Barton, P. B., and Touhnin, P., 1961, Some mechanisms for cooling hydrothermal fluids: U.S. Geol. Survey Prof. Paper 424-D, p. 348-352.

Baumgartner, L. P., and Rumble, D., III, 1988, Transport of stable isotopes: I: Development of a kinetic continuum theory for stable isotope transport: Contr. Mineralogy Petrology, v. 98, p. 417-430.

Bodnar, R. J., Burnham, C. W., and Sterner, S. M., 1985, Synthetic fluid inclusions in natural quartz. III. Determination of phase equilibrium properties in the system H20-NaC1 to 1000øC and 1500 bars: Geochim. et Cosmochim. Acta, v. 49, p. 1861-1873.

Bottrell, S. H., and Yardley, B. W. D., 1988, The composition of a primary granite-derived ore fluid from S. W. England, determined by fluid inclusion analysis: Geochim. et Cosmochim. Acta, v. 52, p. 585-588.

Burnham, C. W., 1979, Magmas and hydrothermal fluids, in Barnes, H. L., ed., Geochemistry of hydrothermal ore deposits: New York, Wiley Intersci., p. 71-136.

Burnham, C. W., and Ohmoto, H., 1980, Late-stage processes of felsic magmatism: Soc. Mining Geologists Japan Spec. Issue 8, p. 1-11.

Burnham, C. W., Holloway, J. R., and Davis, N. F., 1969, Ther- modynamic properties of water to 1000øC and 10,000 bars: Geol. Soc. America Spec. Paper 132, 96 p.

Campbell, A. R., Rye, D. M., and Petersen, U., 1984, A hydrogen and oxygen isotope study of the San Cristobal mine, Peru: hn- plications of the role of water to rock ratio for the genesis of wolframite deposits: ECON. GEOL., v. 79, p. 1818-1832.

Candela, P. A., 1990a, Magmatic ore-forming fluids: Thermody-

4 7 8 CHRISTOPH A. HEINRICH

namic and mass transfer calculations of metal concentrations:

Rev. Econ. Geology, v. 4. (in press). -- 1990b, Felsic magmas, volatiles and metallogenesis: Rev.

Econ. Geology, v. 4 (in press). Cathies, L. M., 1977, An analysis of the cooling of intrusives by

ground-water convection which includes boiling: ECON. GEOL., v. 72, p. 804-826.

Chappell, B. W., and White, A. J. R., 1974, Two contrasting granite types: Pacific Geology, v. 8, p. 173-174.

Charoy, B., 1979, Greisenisation, min•ralisation et fluides associ•s a Cligga Head (Cornwall, sud-ouest de l'Angleterre): Bull. Mi- n•ralogie, v. 102, p. 633-641.

Chou, I. M., 1978, Calibration of oxygen buffers at elevated P and T using the hydrogen fugacity sensor: Am. Mineralogist, v. 63, p. 690-703.

Chou, I. M., and Eugster, H. P., 1977, Solubility of magnetite in supercritical chloride solutions: Am. Jour. Sci., v. 277, p. 1296- 1314.

Cobble, J. W., Murray, R. C. J., Turner, P. J., and Chen, K., 1982, High-temperature thermodynamic data for species in aqueous solution: Palo Alto, Electric Power Research Inst. Rept. NP- 2400, 97 p.

Collins, P. L. F., 1981, The geology and genesis of the Cleveland tin deposit, Western Tasmania: Fluid inclusion and stable isotope studies: ECON. GEOL., v. 76, p. 365-392.

Davis, W. J., and Williams-Jones, A. E., 1985, A fluid inclusion study of the porphyry-greisen, tungsten--molybdenum deposit at Mount Pleasant, New Brunswick: Mineralium Deposita, v. 20, p. 94-101.

Dingwell, D. B., 1988, The structures and properties of fluorine- rich magmas: A review of experimental studies: Canadian Inst. Mining Metallurgy Spec. Vol. 39, p. 1-12.

Drummond, S. E., 1981, Boiling and mixing of hydrothermal fluids: Chemical effects on mineral precipitation: Unpub. Ph.D. thesis, Pennsylvania State Univ., 380 p.

Drummond, S. E., and Ohmoto, H., 1985, Chemical evolution and mineral deposition in boiling hydrothermal systems: ECON. GEOL., v. 80, p. 126-147.

Ladington, P. J., 1983, A fluid inclusion investigation of ore formation in a tin-mineralized granite, New England, New South Wales: ECON. GEOL., v. 78, p. 1204-1221.

Ladington, P. J., and Giblin, A., 1979, Alteration minerals and the precipitation of tin in granitic rocks: Sydney, CSIRO Inst. Earth Resources Tech. Commun., 37 p.

Ladington, P. J., and Kinealy, K., 1983, Some aspects of the hydrothermal reactions of tin during skarn fomration: Geol. Soc. Australia Jour., v. 30, p. 461-471.

Ladington, P. J., and Paterson, R. G., 1984, Microdeformation and fluid inclusions and their significance in mineralized breccia columns in the Ardlethan tin mine, NSW labs. I: Geol. Soc. Aus- tralia Abs., v. 12, p. 154-155.

Eastoe, C. J., 1978, A fluid inclusion study of the Panguna porphyry copper deposit, Bougainville, Papua New Guinea: ECON. GEOL., v. 73, p. 721-748.

-- 1982, Physics and chemistry of the hydrothermal system at the Panguna porphyry copper deposit, Bougainville, Papua New Guinea: ECON. GEOL., v. 77, p. 127-153.

Eugster, H. P., 1985, Granites and hydrothermal ore deposits: A geochemical framework: Mineralog. Mag., v. 49, p. 7-23.

Eugster, H. P., and Wilson, G. A., 1985, Transport and deposition of ore-forming elements in hydrothermal systems associated with granites, in Halls, C. et al., eds., High heat production (HPP) granites, hydrothermal circulation and ore genesis: St. Austell, England, Inst. Mining Metallurgy, p. 87-98.

Francis, P. W., Baker, M. C. W., and Halls, C., 1981, The Kari Kari caldera, Bolivia, and the Cerro Rico stock: Jour. Volca- nology Geotherm. Research, v. 10, p. 113-124.

Gehrig, M., 1980, Phasengleichgewichte und PVT-Daten tren/irer Mischungen aus Wasser, Kohlendioxid und Natriumchlorid bis

3 kbar und 550øC: Freiburg. Hochschulsammlung Naturw., Chemic 1,109 p.

Giggenbach, W. F., 1980, Geothermal gas equilibria: Geochim. et Cosmochim. Acta, v. 44, p. 2021-2032.

-- 1981, Geothermal mineral equilibria: Geochim. et Cos- mochim. Acta, v. 48, p. 2693-2711.

Gunter, W. D., and Eugster, H. P., 1980, Mica-feldspar equilibria in supercritical alkali chloride solutions: Contr. Mineralogy Pe- trology, v. 75, p. 235-250.

Haapala, I., and Kinnunen, K., 1982, Fluid inclusion evidence on the genesis of tin deposits, in Evans, A.M., ed., Metallization associated with acid magmatism: Chichester, Wiley, p. 101- 111.

Heinrich, C. A., 1986, Some thermodynamic and mass balance considerations on the hydrothermal precipitation ofcassiterite: Genesis of tin tungsten deposits and their associated granitoids [abs.]: [Australia] Bur. Mineral Resources Rec. 1986/10 p. 26- 27.

Heinrich, C. A., and Ladington, P. J., 1986, Thermodynamic predictions of the hydrothermal chemistry of arsenic, and their significance for the paragenetic sequence of some cassiterite- arsenopyrite-base metal sulfide deposits: ECON. GEOL., v. 81, p. 511-529.

Heinrich, C. A., and Seward, T. M., 1990, A UV-spectrophoto- metric study of aqueous iron(II) cloride complexation at 25 to 200øC: Geochim. et Cosmochim. Acta, v. 54, no. 8.

Heinrich, C. A., Henley, R. W., and Seward, T. M., 1989, Hy- drothermal systems: Adelaide, Australian Mineral Found., 74 p.

Helgeson, H. C., 1969, Thermodynamics of hydrothermal systems at elevated temperatures and pressures: Am. Jour. Sci., v. 267, p. 729-804.

-- 1970, A chemical and thermodynamic model of ore deposition in hydrothermal systems: Mineralog. Soc. America Spec. Paper, v. 3, p. 155-186.

Helgeson, H. C., and Kirkham, D. H., 1974, Theoretical prediction of the thermodynamic behavior of aqueous electrolytes at high pressures and temperatures: II. Debye-Hiickel parameters for activity coefficients and relative partial molal properties: Am. Jour. Sci., v. 274, p. 1199-1261.

Helgeson, H. C., Garrels, R. M., and Mackenzie, F. T., 1969, Evaluation of irreversible reactions in geochemical processes involving minerals and aqueous solutions--II. Applications: Geochim. et Cosmochim. Acta, v. 33, p. 455-481.

Helgeson, H. C., Kirkham, D. H., and Flowers, G. C., 1981, Theoretical prediction of the thermodynamic behavior of aqueous electrolytes at high pressures and temperatures: IV. Calculation of activity coefficients, osmotic coefficients, and apparent molal and standard and relative partial molal properties to 600øC and 5kb: Am. Jour. Sci., v. 281, p. 1249-1516.

Hemley, J. J., 1959, Some mineralogical equilibria in the system K20-AI2Os-SiO2-H20: Am. Jour. Sci., v. 257, p. 241-270.

Henley, R. W., and Ellis, A. J., 1983, Geothermal systems, ancient and modern: A geochemical review: Earth-Sci. Rev., v. 19, p. 1-50.

Henley, R. W., and McNabb, A., 1978, Magmatic vapor plumes and ground-water interaction in porphyry copper emplacement: ECON. GEOL., v. 73, p. 1-20.

Hewitt, D. A., and Wones, D. R., 1984, Experimental phase relations of the micas: Rev. Mineralogy, v. 13, p. 201-256.

Higgins, N. C., 1985, Wolframite deposition in a hydrothermal vein system: The Grey River tungsten prospect, Newfoundland, Canada: ECON. GEOL., v. 80, p. 1297-1327.

-- 1990, Moderately depleted oxygen isotope composition of waters associated with tin- and tungsten-bearing quartz veins: An evaluation of models, in Herbert, H. K., and Ho, S. E., eds., Stable isotopes and fluid processes in mineralization: Perth, Univ. Western Australia Geol. Dept. Univ. Ext. Pub. 23, p. 204- 214.


Higgins, N. C., Forsythe, D. L., Sun, S.-S., and A. ndrew, A. S., 1987, Fluid and metal sources in the Mt Carbine tungsten deposit, North Queensland, Australia labs.l: Pacific Rim Congress '87 Gold Coast, Australia, Ext. Abs., p. 173-177.

Hoffmann, C. F., Henley, R. W., Summons, R. E., Solomon, M., and Higgins, N. C., 1988, Biogenic hydrocarbons in fluid inclusions from the Aberfoyle tin-tungsten deposit, Tasmania: Chem. Geology, v. 70, p. 287-299.

Holland, H. D., and Malinin, S. D., 1979, The solubility and occurrence of nonore minerals, in Barnes, H. L., ed., The geochemistry of hydrothermal ore deposits: New York, Wiley In- tersci., p. 461-508.

Holloway, J. R., 1976, Fluids in the evolution ofgranitic magmas. Consequences of finite CO2 solubility: Geol. Soc. America Bull., v. 87, p. 1513-1518.

Holyland, P. W., 1987, Dynamic modelling at the Renison tin mine labs.l: Pacific Rim Congress '87, Gold Coast, Australia, Ext. Abs., p. 189-193.

Huspeni, J. R., Kesler, S. E., Ruiz, J., Tuta, Z., Sutter, J. F., and Jones, L. M., 1984, Petrology and geochemistry of rhyolites associated with tin mineralization in northern Mexico: ECON. GEOL., v. 79, p. 87-105.

Ishihara, S., 1977, The ilmenite-series and magnetite-series granitic rocks: Mining Geology, v. 27, p. 293-305.

Jackson, K. J., and Helgeson, H. C., 1985a, Chemical and thermodynamic constraints on the hydrothermal transport and deposition of tin: I. Calculation of the solubility of cassiterite at high pressures and temperatures: Geochim. et Cosmochim. Acta, v. 49, p. 1-22.

-- 1985b, Chemical and thermodynamic constraints on the hydrothermal deposition of tin: II. Interpretation of phase relations in the Southeast Asian tin belt: ECON. GEOL., v. 80, p. 1365-1378.

Jackson, N.J., Moore, J. M., and Rankin, A. H., 1977, Fluid inclusions and mineralization at Cligga Head, Cornwall, England: Geol. Soc. London Jour., v. 134, p. 343-349.

Jackson, N. J., Halliday, A. N., Sheppard, S. M. F., and Mitchell, J. G., 1982, Hydrothermal activity in the St. Just mining district, Cornwall, England, in Evans, A.M., ed., Metallization associated with acid magmatism: Chichester, Wiley, p. 137-179,

Jaireth, S., Heinrich, C. A., and Solomon, M., 1990, Chemical controls on hydrothermal tungsten transport in some magmatic systems and the precipitation of ferberite and scheelite labs.l: Geol. Soc. Australia Abs., v. 25, p. 269-270.

James, R. S., Turnock, A. C., and Fawcett, J. J., 1976, The stability and phase relations of iron chlorite below 8.5kb pH20: Contr. Mineralogy Petrology, v. 56, p. 1-25.

Kelly, W. C., and Rye, R. O., 1979, Geologic, fluid inclusion and stable isotope studies of the tin-tungsten deposits of Panasqueira, Portugal: ECON. GEOL., v. 74, p. 1721-1822.

Kelly, W. C., and Turneaure, F. S., 1970, Mineralogy, paragenesis and geometry of the tin and tungsten deposits of the eastern Andes, Bolivia: ECON. GEOL., v. 65, p. 609-680.

Kigai, I. N., 1978, Two hydrodynamic types of ore-forming systems, in Stemprok, M., Burnol, L., and Tischendorf, G. eds., Metallization associated with acid magmatism: Prague, Czechoslovakia Acad. Sci., p. 343-348.

Korzhinskii, D. S., 1970, The theory of metasomatic zoning: Ox- ford, Clarendon Press, 162 p.

Krupp, R. E., and Seward, T. M., 1987, The Rotokawa geothermal system, New Zealand: An active epithermal gold-depositing environment: ECON. GEOL., v. 82, p. 1109-1129.

Kwak, T. A. P., 1987, W-Sn skarn deposits and related metamorphic skarns and granitoids: Amsterdam, Elsevier, 451 p.

Lagache, M., and Weisbrod, A., 1977, The system: Two alkali feldspars-KC1-NaCI-H•O at moderate to high temperatures and low pressures: Contr. Mineralogy Petrology, v. 62, p. 77-101.

Landis, G. P., and Rye, R. O., 1974, Geologic, fluid inclusion, and

stable isotope studies of the Pasto Bueno tungsten-base metal ore deposit, northern Peru: ECON. GEOL., v. 69, p. 1025-1059.

Lehmann, B., 1987, Tin granites, geochemical heritage, magmatic differentiation: Geol. Rundschau, v. 76, p. 177-185.

Lichtner, P. C., 1988, The quasi-stationary state approximation to coupled mass transport and fluid-rock interaction in a porous medium: Geochim. et Cosmochim. Acta, v. 52, p. 143-165.

Lukashov, Y. M., Komissarov, K. B., Gobulev, B. P., Smirnov, S. N., and Svistunov, E. P., 1975, An experimental investigation of the electroytic properties of uni-univalent electrolytes at high parameters of state: Teploenergetika, v. 22, p. 78-81.

Manning, D. A. C., 1982, An experimental study of the effects of fluorine on the crystallization ofgranitic melts, in Evans, A.M., ed., Metallization associated with acid magmatism: Chichester, Wiley, p. 191-204.

Manning, D. A. C., and Henderson, P., 1984, The behaviour of tungsten in granitic melt-vapour systems: Contr. Mineralogy Petrology, v. 86, p. 286-293.

Marshall, W. L., and Mesmer, R. E., 1981, On the treatment of pressure effects on ionization constants in aqueous solutions: Jour. Solution Chemistry v. 10, p. 121-127.

Merino, E., Moore, C., Ortoleva, P., and Ripley, E., 1986, Mineral zoning in sediment-hosted copper-iron sulfide deposits--a quantitative kinetic approach, in Friedrich, G. H., ed., Geology and metallogeny /•f copper deposits: Heidelberg, Springer- Verlag, p. 559-571.

Newberry, R. J., and Swanson, S. E., 1986, Scheelite skarn granitoids: An evaluation of the roles of magmatic source and process: Ore Geology Rev., v. 1, p. 57-81.

Ohmoto, H., and Kerrick, D. M., 1977, Devolatilization equilibria in graphitic systems: Am. Jour. Sci., v. 277, p. 1013-1044.

Pabalan, R. T., 1986, Solubility of cassiterite (SnO•) in NaC1 solutions from 200øC-350øC, with geologic applications: Unpub. Ph.D. thesis, Pennsylvania State Univ., 140 p.

Patterson, D. J., Ohmoto, H., and Solomon, M., 1981, Geologic setting and genesis of the cassiterite-sulfide mineralization at Renison Bell, western Tasmania: ECON. GEOL., v. 76, p. 393- 438.

Plimer, I. R., 1987, Fundamental parameters for the formation of granite-related tin deposits: Geol. Rundschau, v. 71, p. 23- 40.

Pollard, P. J., 1983, Magmatic and postmagmatic processes in the formation of rocks associated with rare-element deposits: Inst. Mining Metallurgy Trans., v. 92, sec. B p. B1-B9.

Polya, D. A., 1988, Efficiency of hydrothermal ore formation and the Panasqueira W-Cu(Ag)-Sn vein deposit: Nature, v. 333, p. 838-841.

Ramboz, C., Schnapper, D., and Dubessy, J., 1985, The P-V-T- X-fo2 evolution of H•O-CO•-CH4-bearing fluid in a wolframite vein: Reconstruction from fluid inclusion studies: Geochim. et

Cosmochim. Acta, v. 49, p. 205-219. Reed, M. H., 1983, Seawater-basalt reaction and the origin of

greenstones and related ore deposits: ECON. GEOL., v. 78, p. 466-485.

Reed, M. H., and Spycher, N., 1984, Calculation ofpH and mineral equilibria in hydrothermal waters with application to geother- mometry and studies of boiling and dilution: Geochim. et Cos- mochim. Acta, v. 48, p. 1479-1492.

Ren, S. K., and Walshe, J. L., 1986, Geology, brecciation and paragenesis of the Ardlethan tin field: Genesis of tin tungsten deposits and their associated granitoids labs.l: [Australia] Bur. Mineral Resources Rec. 1986/10, p. 63-64.

Ruaya, J. R., and Seward, T. M., 1986, The stability of chloro- zinc(II) complexes in hydrothermal solutions up to 350øC: Geochim. et Cosmochim. Acta, v. 50, p. 651-661.

-- 1987, The ion-pair constant and other thermodynamic properties of HC1 up to 350øC: Geochim. et Cosmochim. Acta, v. 51, p. 121-130.

Schwartz, M. O., and Askury, A. K., 1989, Geologic, geochemical,

4 8 0 CHRISTOPH A. HEINRICH

and fluid inclusion studies of the tin granites from the Bujang Melaka pluton, Kinta Valley, Malaysia: ECON. GEOL., v. 84, p. 751-779.

Seward, T. M., 1974, Thio complexes of gold in hydrothermal ore solutions: Geochim. et Cosmochim. Acta, v. 37, p. 379- 399.

Shcherba, G. N., 1970, Greisens: Internat. Geology Rev., v. 12, p. 114-150 and 239-255.

Shimazaki, H., 1980, Characteristics of skarn deposits and related acid magmatism in Japan: ECON. GEOL., v. 75, p. 173-183.

Sillitoe, R. H., Halls, C., and Grant, J. N., 1975, Porphyry tin deposits in Bolivia: ECON. GEOL., v. 70, p. 913-927.

Solomon, M. S., Collins, P. L. F., Etheridge, M. A., Halley, S., Hellsten, K. J., Higgins, N. C., and Wall, V. J., 1986, Formation of the Aberfoyle and Lutwyche cassiterite-wolframite veins, Tasmania, Australia: Terra Cognita, v. 6, p. 531-532.

Solomon, M. S., Etheridge, M. A., and Groves, D. I., 1991, The geological setting and nature of Australia's mineral deposits: Oxford, Oxford Univ. Press, in press.

Spycher, N. F., and Reed, M. H., 1986, Boiling of geothermal waters: Precipitation of base and precious metals, speciation of arsenic and antimony, and the role of gas phase metal transport, in Jackson, K. J., and Bourcier, W. L., eds., Proceedings of the workshop on geochemical modelling, CONF-8609134: Cali- fornia, Lawrence Livermore Natl. Lab., p. 58-65.

Stemprok, M., 1987, Greisenization (a review): Geol. Rundschau, v. 76, p. 169-175.

Sun, S., and Eadington, P. J., 1987, Oxygen isotope evidence for the mixing of magmatic and meteoric waters during tin mineralization in the Mole Granite, New South Wales, Australia: ECON. GEOL., v. 82, p. 43-52.

Sverjensky, D. A., 1987a, The role of migrating oil flied brines in the formation of sediment-hosted Cu-rich deposits: ECON. GEOL., v. 82, p. 1130-1141.

-- 1987b, Calculation of the thermodynamic properties of aqueous species and the solubilities of minerals in supercritical electrolyte solutions: Rev. Mineralogy, v. 17, p. 177-209.

Symonds, R. B., Rose, W. I., Reed, M. H., Lichte, F. E., and Fin- negan, D. L., 1987, Volatilization, transport and sublimation

of metallic and non-metallic elements in high-temperature gases at Merapi volcano, Indonesia: Geochim. et Cosmochim. Acta, v. 51, p. 2083-2101.

Taylor, J. R., 1988, The partitioning of tin from granitic magmas: Unpub. Ph.D. thesis, Melbourne, Monash Univ., 223 p.

Taylor, J. R., Wall, V. J., and Bloom, M. S., 1984, The mobilisation of tin from granitoid magmas labs. I: Geol. Soc. Australia Abs., v. 12, p. 514.

Taylor, R. G., 1979, Geology of tin deposits: Devel. Econ. Geology, v. 11,543p.

Thompson, J. B., Laird, J., and Thompson, A. B., 1982, Reactions in amphibolite, greenschist and blueschist: Jour. Petrology, v. œ3, p. 1-27.

Truesdell, A. H., and Jones, B. F., 1974, WATEQ, a computer program for calculating chemical equilibria of natural waters: U.S. Geol. Survey Jour. Research, v. 2, p. 233-248.

Turnbull, A. G., and Wadsley, M. W., 1986, The CSIRO-SGTE THERMODATA system, version V: Melbourne, CSIRO Div. Mineral Chemistry Commun., v. 1-7, 413 p.

Walshe, J. L., 1986, A six-component chlorite solid solution model and the conditions of chlorite formation in hydrothermal and geothermal systems: ECON. GEOL., v. 81, p. 681-703.

Weidner, J. R., and Martin, R. F., 1987, Phase equilibria of a fluorine-rich leucogranite from the St. Austell pluton, Cornwall: Geochim. et Cosmochim. Acta, v. 51, p. 1591-1597.

Wesolowski, D., 1984, Geochemistry of tungsten in scheelite deposits: The skarn ores at King Island, Tasmania: Unpub. Ph.D. thesis, Pennsylvania State Univ., 431 p.

Witt, W. K., 1988, Evolution of high-temperature hydrothermal fluids associated with greisenization and feldspathic alteration of a tin-mineralized granite, northeast Queensland: ECON. GEOL., v. 83, p. 310-334.

Wones, D. R., 1981, Mafic silicates as indicators of intensive variables in granitic magmas: Mining Geology, v. 31, p. 191-212.

Woodland, A., and Walther, J., 1986, Experimental determination of the solubility of Na and K feldspar q- mica q- quartz assemblages in supercritical H•O labs.]: Am. Geophys. Union Trans., v. 67/16, p. 388.

APPENDIX

Summary of Thermodynamic Data

Equilibrium constants of formation from elements, log Kf, at selected temperatures and the equilibrium vapor pressure of H9•O. Extrapolations to 400øC are tentative. References and comments on derivation see Table 3 and text.

log Kr

Temperature (øC) 200 250 300 350 (400)

H20(g) 24.3 21.7 19.6 17.8 (16.3) HCI(g) 10.7 9.7 8.9 8.2 (7.6) CO2(g) 43.6 39.4 36.0 33.1 (30.7) CH4(g ) 3.9 3.1 2.3 1.7 (1.2) SO•lg ) 33.2 30.1 27.4 25.2 (23.3) H20 23.2 20.3 17.8 15.8 (14.0) H• -2.8 -2.6 -2.4 -2.1 (-1.O) H + 0.0 0.0 0.0 0.0 (0.0) OH- 11.7 9.0 6.6 4.4 (2.5) CI- 11.4 0.3 7.5 5.0 (4..5) Na + 30.4 27.8 25.6 23.8 (22.2) K + 33.4 30.7 28.4 26.5 (24.9) NaCI ø 41.4 37.3 33.9 31.1 (28.6) KC1 ø 43.9 39.6 36.0 33.0 (30.5) NaOH ø 40.4 36.9 33.9 31.5 (29.4)


APPENDIX (Cont.)

log Kt

Temperature (øC) 200 250 300 350 (400)

KOH ø 43.6 39.8 36.6 34.0 (31.7) Fe +2 10.1 9.1 8.3 7.6 (7.1) FeCI + 22.6 20.61 19.41 18.5 • (18.1) FeCI• 32.8 29.11 26.71 25.01 (24.1) SnCI + 18.8 17.3 16.4 15.8 (15.4) SnCI• 31.9 28.3 25.8 23.9 (22.5) SnCI• 43.5 38.3 34.0 30.6 (27.7) SnCI• 2 53.7 46.3 40.0 34.4 (29.5) SnCI•(OH)• 79.6 69.3 61.3 55.1 (50.2) SnCI(OH)• 100.9 87.9 77.2 68.2 (60.5) SnCI•(OH)• • 113.7 97.8 84.8 73.8 (64.4) Sn(OH)• 107.5 92.9 80.9 70.8 (62.2) H•WO• 86.8 77.2 69.6 63.7 (59.0) HWO• 82.6 73.1 65.7 59.8 (55.2) WO• • 77.2 67.0 58.8 52.1 (46.6) SiO• 88.6 79.2 71.5 64.9 (59.4) CO• 41.6 37.5 34.2 31.5 (29.2) CH• 1.1 O.6 O.2 -0.1 (-O.2) HCO• 57.5 49.9 43.6 38.2 (33.6) H2S ø 2.8 2.5 2.3 2.2 (2.0) HS- -4.4 -5.1 -5.8 -6.6 (-7.3) HSO• 73.9 64.1 55.9 49.0 (43.0) KAISiaO8 397.7 355.9 321.4 292.4 (267.8) NaA1SiaO8 393.7 352.3 318.1 289.4 (265.0) KAI•SiaO1o(OH)• 589.5 526.7 474.9 431.4 (394.4) NaAl•SiaOl o(OH)• 584.2 521.8 470.3 427.2 (390.4) A12Si•Os(OH)4 395.0 351.8 316.2 286.3 (260.8) SiO2 91.0 81.4 73.5 66.8 (61.2) FeS• 16.8 14.9 13.3 12.0 (10.8) FeS 11.2 10.2 9.3 8.6 (8.0) Fe203 76.9 68.2 61.0 55.1 (50.0) Fe•O4 105.4 93.7 84.0 75.9 (69.1) SnO• 53.5 47.3 42.2 37.9 (34.3) FeWO4 101.3 89.9 80.6 72.7 (66.0) Fe•SiO4 174.7 155.9 140.4 127.3 (116.2) FeCO3 68.9 61.1 54.6 49.2 (44.7) KFe3A1Si301o(OH)• 508.8 454.3 409.3 371.5 (339.4) KFe3A1Si301o(OOH) 508.0 453.7 408.9 371.3 (339.3) FesAl•Si30•o(OH)• 679.7 604.8 543.0 491.2 (447.2) Fe?SiaOlo(OH)s 588.9 522.9 468.6 422.9 (384.2) NiO 21.7 19.1 17.1 15.3 (13.9)

Fitted to log K (Fe +• + C1-) = FeCI +) = 1.1 (200øC), 2.2 (250øC), 3.6 (300øC), .5.0 (350øC) and log K (Fe +• + 2C1- = FeCI•) -0.1 (200øC), 1.4 (250øC), 3.4 (300øC), 5.6 (350øC) estimated from isocoulombic extrapolation of data by Heinrich and Seward

(1990).

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