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Earth and Planetary Science Letters 411 (2015) 298–309
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Earth and Planetary Science Letters
www.elsevier.com/locate/epsl
Stability and abundance of the trisulfur radical ion S−3
in hydrothermal fluids
Gleb S. Pokrovski a,∗, Jean Dubessy b
a Groupe Métallogénie Expérimentale, Géosciences Environnement Toulouse (GET), UMR 5563 of CNRS, University of Toulouse, 14 venue Edouard Belin, F-31400 Toulouse, Franceb CNRS, UMR 7359 GeoRessources, Université de Lorraine, BP 70239, F-54506 Vandœuvre-lès-Nancy, France
a r t i c l e i n f o a b s t r a c t
Article history:Received 6 April 2014Received in revised form 19 November 2014Accepted 21 November 2014Available online xxxxEditor: T. Elliott
Keywords:sulfurS−
3 ionRaman spectroscopythermodynamic propertiesgold depositsulfur isotopes
The interpretation of sulfur behavior in geological fluids and melts is based on a long-standing paradigm that sulfate, sulfide, and sulfur dioxide are the major sulfur compounds. This paradigm was recently challenged by the discovery of the trisulfur ion S−
3 in aqueous S-bearing fluids from laboratory experiments at elevated temperatures. However, the stability and abundance of this potentially important sulfur species remain insufficiently quantified at hydrothermal conditions. Here we used in situ Raman spectroscopy on model thiosulfate, sulfide, and sulfate aqueous solutions across a wide range of sulfur concentration (0.5–10.0 wt%), acidity (pH 3–8), temperature (200–500 ◦C), and pressure (15–1500 bar) to identify the different sulfur species and determine their concentrations. Results show that in the low-density (<0.2 g/cm3) vapor phase, H2S is the only detectable sulfur form. By contrast, in the denser liquid and supercritical fluid phase, together with sulfide and sulfate, the trisulfur radical ion S−
3 is a ubiquitous and thermodynamically stable species from 200 ◦C to at least 500 ◦C. In addition, the disulfur radical ion S−
2 is detected at 450–500 ◦C in most solutions, and polymeric molecular sulfur with a maximum abundance around 300 ◦C in S-rich solutions. These results, combined with revised literature data, allow the thermodynamic properties of S−
3 to be constrained, enabling quantitative predictions of its abundance over a wide temperature and pressure range of crustal fluids. These predictions suggest that S−
3 may account for up to 10% of total dissolved sulfur (Stot) at 300–500 ◦C in fluids from arc-related magmatic–hydrothermal systems, and more than 50% Stot at 600–700 ◦C in S-rich fluids produced via prograde metamorphism of pyrite-bearing rocks. The trisulfur ion may favor the mobility of sulfur itself and associated metals (Au, Cu, Pt, Mo) in geological fluids over a large range of depth and provide the source of these elements for orogenic Au and porphyry-epithermal Cu–Au–Mo deposits. Furthermore, the ubiquity of S−
3 in aqueous sulfate–sulfide systems offers new interpretations of the kinetics and mechanisms of sulfur redox reactions at elevated temperatures and associated mass-dependent and mass-independent fractionation of sulfur isotopes.
© 2014 Elsevier B.V. All rights reserved.
1. Introduction
All models of metal sulfide ore deposit formation and sulfur isotope fractionation require knowledge of sulfur speciation in geo-logical fluids at elevated temperature (T ) and pressure (P ). Because of the ubiquity of sulfate and sulfide minerals in hydrothermal–magmatic systems, by analogy, sulfur chemistry in aqueous fluids and silicate melts at depth has been believed to be controlled by sulfide (H2S, HS− and S2−) and sulfate (HSO−
4 and SO2−4 )
* Corresponding author. Tel.: +33 (0)5 61 33 26 18; fax: +33 (0)5 61 33 25 60.E-mail addresses: [email protected], [email protected]
(G.S. Pokrovski), [email protected] (J. Dubessy).
http://dx.doi.org/10.1016/j.epsl.2014.11.0350012-821X/© 2014 Elsevier B.V. All rights reserved.
(e.g., Boyle, 1969; Ohmoto and Lasaga, 1982; Barnes, 1997; Métrich et al., 2009; Mandeville, 2010). In addition, two intermediate-valence sulfur forms, sulfur dioxide (SO2) and native sulfur (S), produced by magma degassing and fluid or vapor cooling and condensation are commonly observed in volcanic gases and sub-limates.
This apparent simplicity of sulfur speciation at elevated T–Pcontrasts with the variety of sulfur redox states, from −2 to +7, and the corresponding species that exist in aqueous solution, non-aqueous solvents, glasses, and solid phases at ambient con-ditions (Fig. 1; e.g., Cotton et al., 1999; Steudel, 2003). Some of these S forms, such as thiosulfate (S2O2−
3 ), polysulfides (SnS2−), polymeric sulfur (S8), are found in geothermal springs (e.g., Kaasalainen and Stefánsson, 2011), fluid inclusions in minerals
G.S. Pokrovski, J. Dubessy / Earth and Planetary Science Letters 411 (2015) 298–309 299
Fig. 1. Inorganic chemical forms of sulfur known in aqueous solution and non-aqueous solvents at ambient conditions.
(e.g., Giuliani et al., 2003) and as reaction intermediates in biologi-cal sulfate–sulfide redox cycle (e.g., Miluska et al., 2012). Polysulfur radical ions (S−
2 , S−3 , S−
4 ) are important constituents of chemical engineering products (e.g., lithium-sulfur batteries, color pigments and glasses, zeolites; Chivers, 1974; Chivers and Elder, 2013; refer-ences therein). To the multitude of inorganic sulfur forms of Fig. 1may be added a plethora of organic thiol compounds (not shown) generated by microbial activity in near-surface environments (e.g., Amend and Shock, 2001; Schulte and Rogers, 2004). Although be-ing persistent at ambient temperatures because of slow rates of sulfur redox transformations (e.g., Ohmoto and Lasaga, 1982), all these S compounds are considered to be thermodynamically un-stable with respect to sulfide and sulfate. As a result, with the exception of the two latter forms for which robust thermodynamic data are available at elevated T (e.g., Murray and Cubicciotti, 1983;Williamson and Rimstidt, 1992; Johnson et al., 1992), the sta-bility of many intermediate-valence S species is insufficiently constrained in hydrothermal solutions owing to a lack of direct experimental or analytical data whose acquisition is challenging in the face of the dramatic changes in sulfur speciation, redox state, and solubility with T and P . For example, the solubil-ity of common sulfide minerals (pyrite, chalcopyrite, pyrrhotite) and native sulfur decreases by 3 to 7 orders of magnitude when the fluid cools from 500 to 100 ◦C (Dadze and Sorokin, 1993;Kouzmanov and Pokrovski, 2012). Sulfur dioxide (SO2), a major S form in magmatic vapors (Wallace, 2001), breaks down to sul-fate and sulfide upon aqueous fluid cooling below 400–500 ◦C depending on pH (Giggenbach, 1997). Thiosulfate (S2O2−
3 ), kineti-cally persistent at ambient conditions, decomposes to sulfate and sulfide upon aqueous solution heating above 150 ◦C (Ohmoto and Lasaga, 1982). Under atmospheric oxygen pressure at the Earth’s surface, all sulfur chemical forms ultimately oxidize to sulfate. These examples imply that many natural and laboratory products of sulfur reactions at high T–P brought to ambient conditions may not adequately reflect the true sulfur forms operating in melts and fluids at depth. In situ spectroscopic approaches are thus required to unambiguously assess S speciation at hydrothermal–magmatic conditions.
Raman and UV–visible spectroscopy are the methods of choice for probing sulfur in aqueous solution. The majority of studies at elevated T have focused on the stable end-members – sulfide (e.g., Giggenbach, 1970; Ellis and Giggenbach, 1971; Suleimenov and Se-ward, 1997) and sulfate (e.g., Rudolph, 1996; Rudolph et al., 1997;Schmidt, 2009; Ni and Keppler, 2012). Less attention has been de-voted to systems containing intermediate-valence S species (e.g., polysulfides, thiosulfate, molecular sulfur, sulfur dioxide; Giggen-bach, 1971, 1974; Bondarenko and Gorbaty, 1997; Yuan et al., 2013). Pokrovski and Dubrovinsky (2011) performed Raman spec-troscopy in a diamond-anvil cell on aqueous solutions in which
sulfate and sulfide coexist; they found the trisulfur radical ion S−
3 to reversibly form in the range 250–450 ◦C and 5–50 kbar. Jacquemet et al. (2014) observed the formation of S−
3 in synthetic fluid inclusions from similar model systems in the same T interval but at lower P (<1 kbar). These findings suggest that, in addition to sulfide and sulfate, S−
3 might also be an important S form in natural hydrothermal fluids. Thus, knowledge of S−
3 concentrations is required for inclusion of this species in quantitative geochemi-cal models. Accurate thermodynamic properties of S−
3 are needed to be able to predict its amount in geological fluids.
We used quantitative in situ Raman spectroscopy on S-bearing aqueous solutions at P ≤ 1.5 kbar and T ≤ 500 ◦C, to determine the concentrations and thermodynamic properties of S−
3 . These new data combined with a critical revision of the literature en-able, for the first time, quantitative predictions of the abundance of S−
3 in geological fluids and evaluation of its potential geochemi-cal impact over the wide T–P range of the Earth’s crust.
2. Materials and methods
2.1. Sulfur-bearing experimental systems
Aqueous solutions containing 0.5 to 10.0 wt% of total sul-fur (Stot) were examined in this study in the range 200–500 ◦C and 15–1500 bar. The source of sulfur was potassium thiosulfate (K2S2O3 ± HCl in H2O or 1 : 1 H2O : D2O mixture) or potassium sulfate (K2SO4 or KHSO4 in H2O) plus H2S (Table 1). Potassium was preferred to sodium because of higher solubility of K2SO4 than Na2SO4 at elevated temperatures. The solutions were prepared by weight from analytical-grade chemicals (purity >99.9%) and deion-ized degassed H2O (±D2O) prior to loading in the spectroscopic cell. Hydrogen sulfide gas (purity >99.9%) was introduced via a pressure line in amounts controlled by P drop and/or optical ob-servations of the H2S vapor and liquid phase relationships in the cell below 100 ◦C (e.g., Appendix A).
Our experimental systems represent a good proxy for natu-ral S-rich fluids in arc-related magmatic–hydrothermal systems hosting porphyry Cu–Au–Mo deposits, which are characterized by acidic-to-neutral pH and the coexistence of sulfate and sulfide (Einaudi et al., 2003; Kouzmanov and Pokrovski, 2012). Further-more, they impose sulfur redox balance through stoichiometric breakdown of thiosulfate and enable oxygen fugacity ( fO2 ) and acidity (pH) buffering via the dominant reactions:
S2O2−3 + H2O = SO2−
4 + H2S (1)
H2S + 2O2 = SO2−4 + 2H+ (2)
SO2−4 + H+ = HSO−
4 (3)
The extent of these and accompanying reactions at equilibrium may accurately be calculated using available thermodynamic data (see Pokrovski et al., 2009; Jacquemet et al., 2014 for details). In addition to the free sulfate and hydrogen sulfate ions and molec-ular hydrogen sulfide, K+ ion pairs (KSO−
4 , KHSO04), and SO2 are
predicted to form with increasing T (>350–400 ◦C), polysulfide ions (SnS2−) and molten sulfur (as a discrete phase) contribute at moderate T (200–300 ◦C) in concentrated (>2 wt% S) solutions, and HS− at pH > 7. As shown by thermodynamic calculations of this and previous studies (Pokrovski and Dubrovinsky, 2011;Jacquemet et al., 2014), such species are relatively minor (<20%of total S) and do not significantly change the system proper-ties controlled by reactions (1) to (3). Because reactions (1) and (2) are sluggish below 150 ◦C (Ohmoto and Lasaga, 1982), most of our runs were conducted above 200 ◦C – the temperatures at which S2O2−
3 breakdown and H2S–SO4 exchange are fast enough to reach equilibrium within hours as shown by kinetic measure-ments and cooling–heating cycles to check for reversibility (see
300 G.S. Pokrovski, J. Dubessy / Earth and Planetary Science Letters 411 (2015) 298–309
Table 1Chemical composition, total density, T -range, Raman laser frequency, and phase equilibria in the experiments of this study.
# run Composition (mol/kg water)
Density (g/cm3)
T range (◦C)
Laser (nm)
Observed phases at indicated temperatures
#1_11 1.19m K2S2O3 0.36 150–300 514.5 L + V ± S; F + B + KS at >300 ◦C#2_11a 1.19m K2S2O3 0.65 300–500 514.5 L + V ± S at ≤350 ◦C; F (dominant) + B + KS at ≥400 ◦C#2_11 1.19m K2S2O3 0.65 200–500 457.9 L + V ± S at ≤350 ◦C; F (dominant) + B + KS at ≥400 ◦C#4_11 0.32m K2S2O3 0.54 500 457.9 SF#8_11 0.32m K2S2O3 0.70 200–400 457.9 L + V, homogenization to L at ∼350 ◦C, SF at 400 ◦C#1_12 0.68m K2S2O3, 0.22m HCl 0.74 200–400 457.9 L + V, homogenization to L at 340 ◦C, SF at 400 ◦C#5_12 0.68m K2S2O3, 0.22m HCl 0.60 200–400 457.9 L + V, homogenization to L at 380 ◦C; F + B + KS at 400 ◦C#6_12 0.68m K2S2O3, 0.22m HCl 0.74 200–500 457.9 L + V, homogenization to L at 340 ◦C; SF at 400–500 ◦C#7_12 0.68m K2S2O3, 0.22m HCl, 50 wt% D2O 0.83 200–500 457.9 L + V at ≤350 ◦C, F (dominant) + B (very minor) at ≥400 ◦C#8_12 1.17m KCl, 0.16m HCl 0.58 50–500 457.9 L + V, homogenization to L at 406 ◦C, SF at 450–500 ◦C#3_13 0.83m KHSO4, 0.41m KOH, 2.78m H2S 0.53 200–500 457.9 L + V + LH2S below 50 ◦C; L + V + S at 200–350 ◦C;
F (dominant) + B + KS at ≥374 ◦C#4_13 1.24m KHSO4, 1.69m H2S 0.62 200–500 457.9 L + V + S at ≤400 ◦C; F (dominant) + B (traces) + KS at >400 ◦C
m = molality (number of moles of solute per 1 kg of water). V = H2O–H2S vapor; L = aqueous liquid; LH2S = liquid H2S; S = molten sulfur; F = fluid (as opposed to brine above the water critical point, ≥374 ◦C), B = sulfate brine (above the water critical point); KS = potassium sulfate solid, SF = single-phase supercritical fluid. Run #8_12 is a blank experiment used for better constraining the spectral baseline and identifying peaks from the cell walls.
a Run #2_11 was conducted by alternating the two lasers.
Section 3.1 below). In contrast to redox reactions (1) and (2), proto-nation and ion pairing reactions (such as reaction (3)) are very fast processes reaching equilibrium within seconds to minutes (e.g., Martell and Hancock, 1996). As a result, reactions (1) to (3) impose robust chemical constraints (redox balance, fO2 , and pH) on the experimental systems independently of possible presence of low-to-moderate amounts of other S species, yielding fO2 close to that of the hematite–magnetite buffer and pH between 3 and 8 (de-pending of solution composition and temperature). Furthermore, the sulfate–sulfide redox balance allows in some cases more accu-rate constraints on these species concentrations than direct Raman spectroscopic analyses (see Appendix B).
2.2. Spectroscopic cell
We used a recently developed cell (Caumon et al., 2013;Dargent et al., 2013) similar to that described by Chou et al. (2008)and Wang et al. (2011) that enables in situ measurements up to ∼500 ◦C and ∼2 kbar using micro beam laser Raman spectroscopy. The cell consists of round cross-section silica-fused capillary tubing of 323 μm external and 100 μm internal diameters; its prepara-tion and loading protocols are detailed in Chou et al. (2008) and Caumon et al. (2013). The capillary is sealed with a micro-torch at both ends; at ambient temperature it contains an aqueous so-lution and a vapor phase (±H2S liquid, see Appendix A). The total system density (ρ tot, equivalent to the degree of filling of the cell) is estimated by measurement of the volume of each phase on a micrometric stage. The cell is brought to the desired temperature on a heating stage (®CAP-500 Linkam), which insures a very good thermal stability (±0.1 ◦C) and negligible T gradients (<1 ◦C over the cell length of 10–15 mm). Because of its negligible thermal ex-pansion (<1% of volume up to 500 ◦C; Dubessy et al., 2009), the cell is considered to be isochoric. Thus, the internal pressure in the cell corresponds to the saturated vapor pressure of the system (Psat = PH2O +PH2S) along the vapor–liquid coexistence curve up to the homogenization temperature (Th ∼ 340–380 ◦C, see Table 1); above Th the pressure evolves in the single-phase field along an isochore imposed by ρ tot and composition (see Jacquemet et al., 2014).
2.3. Raman spectroscopy measurements
Spectra were obtained at the GeoRessources Laboratory (Nancy, France) with a LabRam HR spectrometer (®Jobin Yvon Horiba), using either 514.5 nm (green) or 457.9 nm (violet) Ar+ laser ex-citation (∼1–2 μm spot size on the sample). The use of multiple
excitation wavelengths allows for more robust identification and quantification of sulfur species that exhibit Raman resonance phe-nomena, such as S−
3 , polysulfides or polymeric sulfur (e.g., Clark and Franks, 1975; Clark and Cobbold, 1978) whose Raman signal is selectively enhanced at certain laser wavelengths, whereas that of non-resonant species such as sulfate and sulfide is little affected. The acquisition was performed on the liquid, vapor, supercritical, and solid/molten phases (Section 3.1) in a backscattering geome-try using an Olympus ×20 objective, a 1800 lines/mm grating, an entrance slit of 200 μm, and a confocal hole of 500 μm (spectral resolution ∼3–5 cm−1). The spectra were recorded over the range 100–4500 cm−1 split in 7 spectral windows with 5–60 s acquisi-tion time per window and 2–10 acquisitions (depending on signal intensity). The typical laser power on the liquid/fluid and vapor phases was 0.6 and 6.0 mW, respectively, which was enough to yield exploitable Raman signal while not overheating the sample that might lead to phase changes or laser-induced photochemical reactions, as carefully checked in each experiment. The spectrome-ter was calibrated using the Raman stretching vibrations of a Si wafer (520.7 cm−1 at 20 ◦C), and oxygen (1555 cm−1) and ni-trogen (2331 cm−1) gas from the air (Dubessy et al., 2012). In addition, the narrow intense Raman bands of SO2−
4 and H2S in our samples provide a complementary check of the energy position during a Raman session. Raman spectra were baseline subtracted and fitted using pseudo-Voigt functions to determine each Raman peak position and integrated intensity (i.e., peak area), which was normalized to that of the H2O stretching band that serves as an in-ternal standard. In addition, external standard solutions of HSO−
4 , SO2−
4 and H2S were measured and processed identically to the ex-perimental samples to establish calibration relationships allowing quantitative analyses of each of these sulfur forms (Appendix B).
3. Results
3.1. Phase composition and equilibria in the experimental systems
All experimental systems display a very similar phase behavior between 200 and 500 ◦C as shown by optical observations (Table 1; Appendix A). The phase relative amounts and their changes with Tare controlled both by the ρ tot values and solute concentrations. In all experiments, an H2S–H2O vapor and an aqueous liquid coexist up to 340–350 ◦C. Above that T , high-density moderately concen-trated thiosulfate systems (ρtot ≥ 0.7 g/cm3, ≤0.7m K2S2O3, where m is the molality) homogenize to liquid, while less dense and/or more concentrated thiosulfate and sulfate–sulfide systems contain a dominant aqueous phase and a small fraction of sulfate brine.
G.S. Pokrovski, J. Dubessy / Earth and Planetary Science Letters 411 (2015) 298–309 301
Molten sulfur, consisting of S8 ring molecules and other Sn poly-mers, is also present as a discrete phase in the form of balls or bullets (Appendix A) in the concentrated runs typically between 200 and 400 ◦C; its amount being larger at more acidic condi-tions, consistent with thermodynamic predictions of S solubility. In low-density S-rich runs, a potassium sulfate solid precipitates at ≥400 ◦C. In most experiments above 300–350 ◦C (where S−
3 devel-ops, see below), the contributions of the vapor, molten sulfur or sulfate salt/brine to the total sulfur balance are minor compared to the aqueous liquid or supercritical fluid. Phase changes occur within minutes (between vapor, liquid, and sulfate solid) to hours (formation of sulfur melt) at a given T , and are fully reversible in heating–cooling cycles, suggesting equilibrium. The Raman pat-terns of the minor phases are detailed in Appendix A; those of the dominant aqueous liquid and supercritical fluid phase are dis-cussed below.
3.2. Aqueous liquid and supercritical fluid phase
The aqueous liquid and fluid phase in all thiosulfate and sulfate–sulfide systems at elevated temperatures displays very similar Raman patterns and the corresponding sulfur species (Figs. 2–4).
Sulfate, hydrogen sulfate, and hydrogen sulfide. H2S, SO2−4
(±KSO−4 ), and HSO−
4 (±KHSO04) are ubiquitous species in the aque-
ous liquid and fluid phase as indicated by their most intense Raman peaks at 2590–2580, 980–970, and 1050–1055 cm−1, re-spectively (the range of values corresponds to a typical frequency change from 200 to 500 ◦C). Their concentrations, as determined from the integrated peak intensities and calibration solutions (Ap-pendix B), reflect the solubility of molten sulfur, partitioning of H2S into the vapor phase, formation of potassium sulfate salt/brine, and the presence of other S species, depending of the run and T (e.g., Fig. B.3). These concentrations are in agreement, within errors, with thermodynamic predictions and the sulfur redox bal-ance equations (1) and (2), and confirm that these three S forms account for the major part of dissolved sulfur.
S−3 ion. The most remarkable feature of all thiosulfate (T ≥
200 ◦C) and sulfate–sulfide (T ≥ 300 ◦C) runs is the systematic presence of the trisulfur radical ion S−
3 , which is clearly identi-fied by its symmetric S–S bending vibration (δ) at ∼235 cm−1, symmetric S–S stretching vibration (ν1) at 530–535 cm−1, and their higher-order overtones (2ν1 ≈ 1070, 3ν1 ≈ 1600, 4ν1 ≈ 2140, 5ν1 ≈ 2670 cm−1) and combination bands (ν1 − δ ≈ 295, ν1 + δ ≈770, 2ν1 − δ ≈ 830, 2ν1 + δ ≈ 1305 cm−1). This Raman pattern is due to a resonance phenomenon induced by the absorption by S−
3of the laser radiation, which results in enhancement of the sym-metric modes and their overtones (see Clark and Franks, 1975;Chivers and Drummond, 1972). The blue S−
3 ion absorbs light over a broad range of wavelengths (450–750 nm) with a maximum at 590–620 nm (Chivers and Elder, 2013) at which the resonance is largest. Thus, the ν1 peak intensity with the 514.5 nm laser whose frequency is closer to that maximum is a factor of ∼15 higher (Fig. 2) than with the 457.9 nm laser (Figs. 3, 4). The third Raman active vibrational mode of S−
3 , the S–S asymmetric stretch ν3 at ∼580 cm−1 (Chivers and Elder, 2013; references therein) was not detected because of its too low intensity in the resonance spec-trum dominated by symmetric vibrations (ν1). The ν1 intensity grows from 250 to 500 ◦C, accompanied by an enhancement of the fluid blue color (Fig. 3) consistent with an increase in S−
3 concen-tration. The Raman and color patterns of S−
3 are unique and cannot be mixed up with those of any other species. In the absence of di-rect standards, S−
3 concentration was first estimated using S redox and mass balance at 500 ◦C in experiment #6_12 (Fig. 3) showing only S− , H2S, sulfate, and hydrogen sulfide as the major species;
3Fig. 2. Raman spectra, at 514.5 nm excitation, of the liquid and supercritical fluid phase in a thiosulfate experiment at the indicated composition and temperatures. Vertical dashed lines denote the vibration modes and major Raman peak positions of the labeled species. The spectra are normalized to 25 s acquisition time and offset vertically for clarity. Panel (b) is a zoom of the low-frequency spectral part (100–1200 cm−1) outlined by a rectangular contour in panel (a).
the established calibration coefficient (analogous to the Raman cross-section of the S−
3 ν1 peak) was then extrapolated to lower Tusing the analogy with those of sulfate, hydrogen sulfate and H2S; the established coefficients were finally used to determine S−
3 con-centrations from the measured ν1 integrated peak intensity in all other experiments (see Appendix B for details and uncertainties). The resulting S−
3 concentrations are reported in Table 2 and shown in Fig. B.3 (for selected runs). They increase by a factor of 10 to 70 (depending of the run) from 250 to 500 ◦C, accounting for up to ∼20% of Stot in most concentrated runs at the highest T (#3_13 and #4_13).
Polymeric sulfur. Another Raman feature of all experimental solutions at equilibrium with molten sulfur is a broad peak at ∼400 cm−1 with a shoulder at 450 cm−1, accompanied by an-other poorly resolved band at 800–850 cm−1 (Figs. 3, 4). All these features are better visible in violet-laser spectra (Fig. 3b) because of the much lower intensity of the overlapping ν1(S−
3 )
band compared to green-laser spectra (Fig. 2b). Both features have their intensity maximum at 300 ◦C, decrease with further T rise, and become undetectable above 450 ◦C in moderately concen-trated thiosulfate solutions (0.7m K2S2O3–HCl); in more concen-trated sulfate–sulfide solutions they persist up to 500 ◦C (Fig. B.7). The wavenumber range of the 400–450 cm−1 and 800–850 cm−1
features corresponds to S–S stretching vibrations and their 2nd or-
302 G.S. Pokrovski, J. Dubessy / Earth and Planetary Science Letters 411 (2015) 298–309
Fig. 3. Raman spectra, at 457.9 nm excitation, of the liquid and supercritical fluid phase in a thiosulfate experiment at the indicated composition and temperatures. Vertical dashed lines denote the vibration modes and major Raman peak positions of the labeled species. The spectra are normalized to 25 s acquisition time and offset vertically for clarity. Panel (b) is a zoom of the low-frequency part of the spectrum (200–1300 cm−1) outlined by a rectangular contour in panel (a). The photos of the capillary cell show color changes in the fluid with increasing temperature (refer to the web version for colors). The black curves labeled “blank” correspond to Raman spectra of an aqueous KCl–HCl solution recorded at 200 and 500 ◦C with the same acquisition parameters.
der overtones, respectively (Steudel, 2003). Similar features were observed in the spectra of molten sulfur (Fig. A.2) suggesting that those in aqueous solution may also arise from some zero-valent sulfur polymeric molecules, S0
n(aq), other than S08(aq), forming in
equilibrium with molten sulfur (see Appendix B for additional ar-guments). Aqueous S0
8 itself was only detected in aqueous solu-tion in one S-rich experiment (#4_13, KHSO4–H2S) at 350–500 ◦C (Fig. 4b) by narrow bands at ∼144 cm, 214, and 470 cm−1, sim-ilar to the S8 ring molecules in molten sulfur (Fig. A2). The large width of the S0
n bands (full width at half maximum, FWHM, of the 400 cm−1 band ≈60 cm−1) compared to those of the S8molecule (FWHM of the 470 cm−1 band ≈10–15 cm−1) suggests the presence of multiple chain-like molecules of low symmetry (Meyer, 1976; Steudel, 2003). The direct in situ observation of aqueous S polymers at elevated temperatures in this study is a new finding; it corroborates the experiments in S–H2O(–NaOH) systems using hydrothermal batch reactors based on analyses of different S forms in the sampled fluid, which reported the pres-ence of zero-valent sulfur in solution (Dadze and Sorokin, 1993;Pokrovski et al., 2008). Maximal possible concentrations of these species do not exceed 0.15m S (which corresponds to 5% of Stot) in our most concentrated experiments as may roughly be estimated using S mass balance (Appendix B).
S−2 ion. Another systematic feature, observed at 450 and 500 ◦C
in violet-laser spectra of thiosulfate and sulfate–sulfide solutions, is a small peak at 580 ± 5 cm−1, on the high-frequency side of ν1(S−), accompanied by weaker bands at 1160 ± 10 and 1740 ±
320 cm−1 (Figs. 3, 4). The 580 and 1160 cm−1 bands may belong to the following species: S−
3 (S–S asymmetric stretch, ν3 ∼ 580 cm−1; Chivers and Elder, 2013; references therein), SO2 (S–O symmet-ric stretch ∼1150 cm−1; Risberg et al., 2007; Ni and Keppler, 2012), and the disulfur radical ion S−
2 (S–S symmetric stretch ∼580–590 cm−1; Chivers and Lau, 1982; Ledé et al., 2007). How-ever, the following arguments suggest that all the three peaks arise from S−
2 rather than S−3 and SO2: 1) the ν3 mode of S−
3 was not observed below 450 ◦C; 2) the SO2 concentrations at 500 ◦C pre-dicted by thermodynamics vary over 2 orders of magnitude (e.g., from 0.002m in #1_11 to 0.3m in #4_13) in experiments that show similar 1160 cm−1 band intensities; 3) both 1160 and 1740 cm−1
features match well the 2nd and 3rd order harmonics of the 580 cm−1 stretch of S−
2 , consistent with its known resonance spec-trum under the violet laser (Ledé et al., 2007). The S−
2 amount does not exceed 1% of Stot in our experiments (<0.04m S, as estimated from its 580 cm−1 band assuming a Raman cross-section simi-lar to that of S−
3 , Appendix B), but is expected to increase above 500 ◦C. Although S−
2 is encountered, together with S−3 , in S-doped
borate glasses and minerals ultramarines (e.g., Chivers et al., 1978;Ledé et al., 2007), our study is the first report of S−
2 in aqueous solution.
In all experimental systems the formation of S−3 (as well as S0
n
and S−2 ) is rapid and fully reversible. On heating, the S−
3 intensitystabilizes within a few min at 350–500 ◦C following the T rise and remains unchanged over up to 20 h at constant T . Cooling the
G.S. Pokrovski, J. Dubessy / Earth and Planetary Science Letters 411 (2015) 298–309 303
Fig. 4. Raman spectra, at 457.9 nm excitation, of the liquid and supercritical fluid phase in a sulfate–sulfide experiment at the indicated composition and temper-atures. Vertical dashed lines denote the vibration modes and major Raman peak positions of the labeled species. The spectra are normalized to 25 s acquisition time and offset vertically for clarity. Panel (b) is a zoom of the low-frequency part of the spectrum (100–1200 cm−1) outlined by a rectangular contour in panel (a). Spectra at 350 ◦C were obtained both on heating (350 ◦C up) and cooling (350 ◦C down) to check for reversibility. All other experiments (not shown) also demonstrate quanti-tative reversibility with temperature (see text).
cell below 250 ◦C results in the almost instantaneous disappear-ance of S−
3 and quantitative return to the sulfate–sulfide–sulfur spectrum. Furthermore, the remarkable similarity in the behavior of S−
3 and the other species in two very different types of sulfur systems (thiosulfate vs sulfate–sulfide) demonstrates that their for-mation and abundance are independent of the initial sulfur state (S–S bonds in S2O2−
3 or contrasting S redox forms in SO2−4 and
H2S). All these findings suggest that S−3 is not a short-lived tran-
sient or metastable complex forming in a certain step of redox exchange between sulfide, sulfate and sulfur, but it is instead a product of true thermodynamic equilibrium.
4. Discussion
4.1. Ubiquitous S−3 ion
The S−3 ion is one of the most well studied S forms since 1970s
in a variety of non-aqueous materials such as S-bearing organic and inorganic solvents, alkali halide melts, borosilicate glasses, al-kali metal–sulfur batteries, ultramarine pigments, and zeolite-type minerals in which this ion is responsible for their blue color (see Chivers and Elder, 2013 for review). Although being unstable in
the presence of water at ambient conditions, the blue S−3 ion is
known to also form on heating above 100 ◦C in aqueous solution of polysulfides and sulfur (e.g., Chivers, 1974). However, its blue color and UV–visible and Raman spectral patterns in aqueous solu-tion in some earlier studies (Giggenbach, 1971; Uyama et al., 1985;Bondarenko and Gorbaty, 1997) were erroneously attributed to S−
2or other species (see Appendix C). Note that hydrothermal batch-reactor experiments in S-rich high-T aqueous solutions (Dadze and Sorokin, 1993; Pokrovski et al., 2008) did not detect S−
3 likely be-cause of its rapid breakdown to sulfate, sulfide and sulfur during fluid sampling or quench. The formation of S−
3 with increasing temperature in aqueous solution of thiosulfate and sulfur was un-ambiguously demonstrated by recent Raman spectroscopy work (Pokrovski and Dubrovinsky, 2011; Jacquemet et al., 2014).
The increasing stability of the trisulfur ion with T revealed in this and previous studies is consistent with the general ten-dency for S-bearing radicals, among which S−
3 is the most ener-getically favorable (e.g., Steudel and Steudel, 2013). Furthermore, additional factors at the molecular level may contribute to its en-hanced stability in aqueous solution compared to other radical and non-radical S polymers. For example, the S−
3 geometry (S–S–S angle ≈105–115◦; Tossell, 2012), similar to that of the H2O molecule, may allow energetically favorable hydration structures in the H2O hydrogen-bond network of liquid-like fluids (Pokrovski and Dubrovinsky, 2011). In addition, alkali cations (Na+ , K+) may also act by solvating S−
3 similarly to its coordination by Na+ in sil-icate cages of zeolites (Reinen and Lindner, 1999) or by forming Na+S−
3 or K+S−3 ion pairs. Thus, the effect of fluid salinity on S−
3stability requires further investigation. The rise of pressure or fluid density leads to an increase in S−
3 abundance, particularly in the transition region between vapor-like (ρ ≤ 0.3 g/cm3) to liquid-like (ρ ≥ 0.5 g/cm3) fluids, consistent with the sharp increase in hy-dration energy of ions in this density range (Pokrovski et al., 2013). Following our own observations and the entropy principles pos-tulating increasing stability of smaller and more compact species with increasing T (e.g., Brimhall and Crerar, 1987), it is expected that the disulfur radical ion S−
2 might form at the expense of S−3
above 500–600 ◦C.
4.2. Thermodynamic properties of S−3
The equilibrium concentrations of S−3 measured in this study al-
low derivation of its thermodynamic stability constants from 200 to 500 ◦C and from P sat to 1.5 kbar. This was achieved using the HCh computer code enabling chemical equilibrium calculations in fluid–mineral–gas systems using the chemical composition of the system and thermodynamic properties of its constituents (Shvarov, 2008). The S−
3 ion was included in the code and its apparent mo-lal Gibbs free energy G0
T ,P was varied to match its Raman-derived concentrations at each T–P-composition point (Table 2). The mi-nor S species in solution (Sn , S2−
n , SO2) and the presence of vapor phase and sulfate salt or brine in the system were also taken into account, but their effect on the final G0
T ,P (S−3 ) values was found
to be negligible. The obtained values (Table 2), along with those of other constituents (Table C.1), allow calculation of the thermody-namic equilibrium constant of the reaction:
2H2S(aq) + SO2−4 (aq) + H+(aq)
= S−3 (aq) + 0.75O2(gas) + 2.5H2O(liq), log10 K4 (4)
The values of log10 K4 are reported in Table 2 and compared with literature data in Fig. 5. The data from this study are identi-cal within errors at a given T in different runs over a large range of S concentration; this consistency further strengthens the valid-ity of our derivations. Our log K4 values systematically increase with increasing T from 200 to 500 ◦C, but are independent of P
304 G.S. Pokrovski, J. Dubessy / Earth and Planetary Science Letters 411 (2015) 298–309
Table 2Concentrations of S−
3 in the Raman experiments as a function of T and P , and the corresponding Gibbs free energies of S−3 and equilibrium constants (log10 K ) for reaction (4):
2H2S(aq) + SO2−4 (aq) + H+(aq) = S−
3 (aq) + 0.75O2(gas) + 2.5H2O(liq), derived in this study.
T(◦C)
P(bar)a
m(S−3 )
(mol/kg H2O)G0
T ,P (S−3 )
(kcal/molb)
log10 K4
#1&2_11: 1.19m K2S2O3
200 30 0.0014±0.0006 7.2±1.0 −24.4 ±0.9250 59 0.011±0.007 3.1±0.7 −19.5 ±0.6300 90 0.028±0.010 1.8±1.0 −16.2 ±0.7350 182 0.070±0.014 −1.4±1.4 −12.0 ±1.0400 500 0.098±0.020 −3.2±1.7 −10.4±1.1450 600 0.112±0.017 −5.3±1.2 −7.7±0.7500 750 0.114±0.029 −7.9±1.7 −5.3±0.9
#4&8_11: 0.32m K2S2O3
300 93 0.0028±0.0010 1.4±0.7 −16.1 ±0.5400 650 0.0033±0.0009 −3.0±1.0 −10.8±0.6500 750 0.0031±0.0006 −7.8±1.4 −5.3±0.8
#1&5&6_12: 0.68m K2S2O3 + 0.22m HCl250 60 0.005±0.003 3.8±1.0 −19.8 ±0.8300 100 0.032±0.020 1.2±0.7 −16.0 ±0.5350 200 0.036±0.012 0.0±0.5 −12.7 ±0.3400 500 0.061±0.012 −4.1±0.5 −10.1±0.3450 950 0.070±0.010 −7.4±0.7 −8.1±0.4500 1400 0.065±0.010 −10.0±1.2 −6.1±0.7
#3_13: 0.83m KHSO4 + 0.41m KOH + 2.78m H2S300 133 0.058±0.006 3.7±1.0 −16.9 ±0.7350 201 0.13±0.05 0.5±2.4 −12.9 ±1.7375 300 0.09±0.06 0.5±2.4 −12.3 ±1.6400 350 0.095±0.060 −0.7±1.9 −10.6±1.2450 500 0.181±0.060 −4.3±2.6 −7.2±1.6500 700 0.204±0.058 −7.2±2.9 −5.1±1.6
#4_13: 1.24m KHSO4 + 1.69m H2S300 109 0.015±0.008 2.4±1.9 −16.4 ±1.5350 185 0.049±0.018 −0.2±1.2 −12.6 ±0.8400 350 0.049±0.018 −2.2±1.2 −10.1±0.8450 500 0.069±0.030 −4.3±1.9 −7.2±1.2500 700 0.191±0.051 −8.4±1.7 −4.8±0.9
Uncertainties are evaluated for each data point using error propagation analysis and are reported at 2σ level; those on m(S−3 ) stem from the Raman measurements (see
Appendix B for detailed evaluation), those on G and K values include, in addition, uncertainties associated with the concentrations of other fluid constituents, activity coefficients, and thermodynamic properties of the constituents of reaction (4) (see Appendix C for details).
a P = total pressure in the system, calculated using the data at the same ρ tot from the NaCl–H2O system (Driesner and Heinrich, 2007) for PH2O plus PH2S from Raman measurements in the vapor phase (below Th). Uncertainties of the P estimation are within ±20 bar below Th, ±50 bar above Th, and ±100 bar at 400–500 ◦C for highly concentrated systems of ρtot ≤ 0.65 g/cm3 in which homogenization was not reached (see Table 1).
b Apparent Gibbs free energy of formation from the elements at the subscripted P and T as defined in Shock and Helgeson (1988).
Fig. 5. Decimal logarithm of the equilibrium constant of reaction (4) versus the re-ciprocal of absolute temperature (in Kelvin). Symbols show the data derived in this study and from the literature at different T–P conditions, solid line represents a weighted least-square fit of experimental data points in the 25–500 ◦C range from this study, Pokrovski and Dubrovinsky (2011) and Giggenbach (1971) using a three parameter equation: log10 K4 = −(123.3 ± 20.5) − (10 585 ± 1380)/T + (45.57 ±6.50) × log10(T ) where T is temperature in Kelvin, (Eq. (C.4), Table C.3, Appendix C), dotted line indicates its extrapolation to 700 ◦C. Errors on individual data points from the three experimental studies cited above are less than the symbol size.
at <1.5 kbar; they are similar within errors to those reported by Pokrovski and Dubrovinsky (2011) at similar temperatures but higher pressures (5 ≤ P ≤ 15 kbar). The temperature trend of our data is qualitatively supported by recent quantum-chemistry mod-eling (Tossell, 2012); however, the uncertainties associated with that work are unknown and the reported absolute K4 values would imply S−
3 to be at least ∼10 times more abundant at the condi-tions of our study than the Raman analyses demonstrate. Finally, our data are in quantitative agreement with those of Giggenbach(1971) by UV–visible spectroscopy of aqueous polysulfide–sulfide solutions from 60 to 260 ◦C and P sat as recalculated to reac-tion (4) (see Appendix C for details). Giggenbach’s revised values are identical within errors to those measured in our study at 200 and 250 ◦C and follow a similar temperature trend down to 25 ◦C (Fig. 5). The coherence between our and Giggenbach’s (1971)studies, employing very different methods and chemical systems, further supports the validity of the S−
3 stability constants and their consistency with the thermodynamic data of other S species adopted in our study.
The stability constants of S−3 derived in this study together
with those from Giggenbach (1971) and Pokrovski and Dubrovin-sky (2011) allow generation, for the first time, of a self-consistent set of S−
3 thermodynamic properties in the framework of the HKF (Helgeson et al., 1981) equation of state. Details of this analysis and associated uncertainties are discussed in Appendix C, and the
G.S. Pokrovski, J. Dubessy / Earth and Planetary Science Letters 411 (2015) 298–309 305
Table 3Standard molal thermodynamic properties at 25 ◦C and 1 bar and parame-ters of the revised HKF equation of state for S−
3 retrieved in this study.
Thermodynamic propertya S−3 ion
Gibbs energy � f G0, kcal/mol 13.16 ± 3.50b
Enthalpy, � f H0, kcal/mol 10.84 ± 2.50b
Entropy, S0, cal/(mol K) 28.6 ± 8.0b
Heat capacity, C0p , cal/(mol K) 62.3 ± 12.8b
Volume, V 0, cm3/mol 37.7 ± 3.0b
HKF coefficientsc
a1 × 10, cal/(mol bar) 2.5 ± 4.0a2 × 10−2, cal/mol 19.9 ± 11.0a3, cal K/(mol bar) 9.2 ± 5.0a4 × 10−4, cal K/mol −3.6 ± 2.0c1, cal/(mol K) 50.2 ± 5.0c2 × 10−4, cal K/mol 9.6 ± 2.6ω × 10−5, cal/mol 0.8 ± 0.2
a Gibbs free energy and enthalpy correspond to those of formation from the elements at standard T–P conditions (298.15 K and 1 bar). The reference states for the elements (for which � f G0
1 bar,298 K and � f H01 bar,298 K = 0) in
the system S–O–H are S (orthorhombic), O2 (ideal gas), and H2 (ideal gas).b Calculated from the reaction (4) properties in Table C.3 and using the
� f G0, � f H0, S0 and C0p at 298.15 K and 1 bar of the reaction constituents
from the sources cited in Table C.1.c Derived from fitting the apparent Gibbs free energy values of S−
3from this study, Pokrovski and Dubrovinsky (2011) and revised data of Giggenbach (1971) at a function of T and P , using the HKF model equa-tions and correlations among parameters as implemented in the OptimB program (Shvarov, in press); see Appendix C for discussion of uncertainties.
resulting HKF parameters of S−3 are reported in Table 3. They allow
calculation of the Gibbs free energy values of S−3 within better than
±2 kcal/mol over the T–P range covered by experimental data (25–500 ◦C, <5 kbar), and predictions up to ∼700 ◦C and 15 kbar within ±10 kcal/mol. Combined with the available thermodynamic properties of the major sulfur species and minerals and the pre-dictive capacity of the revised HKF model (Oelkers et al., 2009;Sverjensky et al., 2014), these data enable quantification of S−
3 con-centrations in geological fluids over the Earth’s crust conditions.
5. Geological applications
5.1. Abundance of S−3 in geological fluids
Equilibrium concentrations of S−3 and other major S species
in the fluid phase were modeled in a wide range of T–P condi-tions and fluid compositions using the HCh computer code. Fig. 6shows the effect of total S content, redox potential, acidity, and pressure on the S−
3 abundance at a typical temperature of 450 ◦C, but the results are similar in a wide T range, from 300 to 600 ◦C. The following physical–chemical parameters were found to be fa-vorable for S−
3 : 1) temperatures above 250 ◦C in a wide pres-sure range of liquid-like fluids (ρ > 0.4–0.5g/cm3); 2) elevated Stot concentrations (>0.5 wt%); 3) moderately acidic-to-neutral pH (4 < pH < 6, depending on T and P ); 4) redox conditions of the sulfide–sulfate (±SO2) coexistence, which are within the range of fO2 between the hematite–magnetite (HM, at T ≤ 500 ◦C) and nickel–nickel oxide (NNO, T ≥ 600 ◦C) buffers. Such conditions occur in two crustal settings: magmatic–hydrothermal porphyry–epithermal Cu(–Au–Mo) systems associated with volcanic arcs, and metamorphic belts hosting orogenic Au deposits. The S−
3 abun-dance in typical fluids from these environments is discussed below.
Fig. 7a shows the distribution of sulfur species in an aqueous fluid degassing from magma and undergoing cooling and decom-pression upon its rise in a porphyry–epithermal system. Total salt (10 wt% NaCl + KCl) and S (2 wt%) contents and sulfur specia-tion (H2S : SO2 ≈ 1) in the initial magmatic fluid are typical of those found in fluid inclusions and magmatic gases in back-arc settings (e.g., Hedenquist and Lowenstern, 1994; Wallace, 2001;
Heinrich, 2005; Seo et al., 2009; Kouzmanov and Pokrovski, 2012). The fluid acidity and redox potential are assumed to be controlled, respectively, by fluid equilibrium with alkali aluminosilicate rocks (pH ∼ 5) and equilibria among the major dissolved sulfur species in the fluid itself (H2S–SO2-sulfate; fO2 ∼ HM assemblage). The fluid is assumed to evolve in a Fe-poor environment and have an elevated initial sulfur/metals (Fe, Cu) ratio, which is one of the case scenarios of porphyry fluid evolution generating S-rich and Fe-poor fluids forming epithermal deposits above porphyry plutons (e.g., Heinrich, 2005; Richards, 2011). At such conditions, S−
3 accounts for 5 to 10% of Stot between 300 and 500 ◦C (which corresponds to 0.1–0.2 wt% S). Considering the uncertainties of S−
3 thermodynamic properties (see above), its maximal abundance may reach 20% of Stot at such conditions. Our extrapolations to T above 500 ◦C, though less certain, also suggest potentially high abundance of S−
3 , from 10% (Fig. 7a) to as high as 40% of Stot
(Fig. C.1a), depending on the choice of thermodynamic data for major S species such as H2S (see Appendix C). Note that the situ-ation shown in Fig. 7a represents the optimal conditions for S−
3in porphyry settings. Other scenarios of fluid evolution are less favorable. For example, if the rock buffering capacity is too low and/or fluid flow is too fast for efficient neutralization of the acid-ity (pH < 3) produced by SO2 disproportionation to H2S and sul-furic acid in the cooling fluid (Pokrovski et al., 2014; references therein), such acidic fluids will contain little S−
3 (<100 ppm S). If the amount of major metals (Fe ± Cu) released from the magma is comparable with or in excess to that of H2S, a large part of S will be removed by pyrite precipitation upon fluid cooling (e.g., FeS2 solubility is <1000 ppm Stot below 500 ◦C; Kouzmanov and Pokrovski, 2012); such S-depleted fluids will carry S−
3 amounts ∼100 times smaller than those in Fig. 7a. Vapor-brine separa-tion in porphyry systems is expected to lead to (partial) break-down of S−
3 , which is not stable in the low-density vapor phase (see Section 4.1 and Fig. 6d). Upon the H2S–SO2 vapor ascent and cooling, S−
3 may re-form in the condensed S-rich liquid car-rying sulfur to epithermal deposits forming around 300 ◦C. Thus, S−
3 concentrations may exhibit a large variability, depending of the details of magma evolution and volatile release, S vs Fe budget, and fluid T–P path and hydrodynamics in a given porphyry sys-tem.
The low solubility of pyrite in aqueous fluids is a key lim-itation for S−
3 formation at moderate temperatures (<400 ◦C) in Fe-dominated systems. In prograde metamorphic settings of greenschist–amphibolite facies above 500 ◦C, the increasing pyrite solubility, followed by pyrite-to-pyrrhotite transformation liberat-ing sulfur and accompanied by chlorite breakdown releasing wa-ter (Tomkins, 2010), generates S-rich fluids that may carry large amounts of S−
3 . This is illustrated in Fig. 7b showing the solubil-ity of the pyrite–pyrrhotite–magnetite assemblage along a typical geothermal gradient of regional Archean metamorphism and mod-ern hot subduction zones (Peacock, 1990). While S−
3 is negligible below 500 ◦C (<0.001 wt% S) because of the low aqueous Stot con-centration (<0.1 wt%), it grows at higher T when approaching the stability limit of pyrite (670 ◦C in our case) and attains absolute concentrations of several wt% S corresponding to half of Stot in the fluid above 600 ◦C. More reducing environments produced via metamorphism of carbonaceous shales (e.g., Tomkins, 2010) would lead to lesser abundances of S−
3 (e.g., Fig. 6b). Although error mar-gins of our predictions are large (see Appendix C), we hypothesize that S−
3 may represent a significant contribution to sulfur budget of metamorphic fluids in oxidizing environments and enlarge the T–P window of sulfur liberation.
306 G.S. Pokrovski, J. Dubessy / Earth and Planetary Science Letters 411 (2015) 298–309
Fig. 6. Distribution of S−3 and other sulfur species (expressed as ppm of S) calculated in model hydrothermal fluids of 10 wt% salinity (NaCl + KCl) at 450 ◦C and 1 kbar
(except (d)) and 2 wt% Stot (except (a)) as a function of (a) total S content at pH ∼ 5 and fO2 buffered by the HM assemblage; (b) oxygen fugacity at pH ∼ 5; (c) fluid acidity at fO2 between HM (acidic pH) and NNO (basic pH); (d) pressure at pH ∼ 5 and fO2 of HM (at ≥0.3 kbar – the pressure above which the HKF model is valid at 450 ◦C). Curves show concentrations of each indicated species (sulfate stands for the sum of SO2−
4 , HSO−4 , and their K+ and Na+ ion pairs); the vertical dashed lines in (b) denote
the fO2 values of the major mineral buffers at these T–P conditions (QFM = quartz–fayalite–magnetite, NNO = nickel–nickel oxide, PPM = pyrite–pyrrhotite–magnetite, and HM = hematite–magnetite). Thermodynamic properties of S−
3 , and other species plus minerals are from Table 3 and Table C.1 (Appendix C), respectively. Concentrations of S−
2 and Sn are at least 10 times less than those of S−3 , those of S8 are less than 50 ppm S (not shown).
5.2. Geochemical significance of S−3
Metal ore deposit formation. The presence of S−3 in geologi-
cal fluids may have far-reaching implications for the formation of economic ore deposits of Cu, Au, Mo, and Pt. This is be-cause S−
3 is expected to form stable complexes with these met-als in aqueous solution, similar to traditional polysulfides (SnS2−; Berndt et al., 1994; Rickard and Luther, 2006; Liu et al., 2013) and related sulfur-nitrogen ions (e.g., S3N−; Bojes et al., 1981, 1982). Thus S−
3 may compete with hydrogen sulfide (H2S and HS−), which has always been believed to be the principal ligand for gold in hydrothermal fluids (e.g., Boyle, 1969; Seward, 1973;Pokrovski et al., 2014; references therein). The strong affinity of S−
3 , comparable to that of HS−, for Au+ and Cu+ in aque-ous solution was surmised by molecular modeling (Tossell, 2012;Mei et al., 2013). Even through S−
3 is rarely a major S species, its absolute concentrations exceed by orders of magnitude those of HS− in the acidic-to-neutral pH range of most hydrothermal fluids (Figs. 6, 7). Thus S−
3 might be a major carrier of Au and, po-tentially, other S-loving metals in S-rich fluids. Stable complexes with S−
3 greatly enhancing metal solubility may thus favor Cu, Mo and Au extraction from magma and their transport to por-phyry and high-sulfidation epithermal systems which cannot be fully accounted for, in some cases, using their known species with Cl− and HS− ligands (e.g., Kouzmanov and Pokrovski, 2012), and Au mobilization from pyrite during metamorphism and the metal transfer to orogenic gold deposits. Inclusion of S−
3 in quantitative models of ore formation awaits experimental data on its com-plexes with different metals. Furthermore, the presence of signif-icant amounts of S−
3 , as suggested by its stability constants from our work, in S-rich aqueous solutions used in some laboratory studies of Au and Mo solubility (e.g., Hayashi and Ohmoto, 1991;Loucks and Mavrogenes, 1999; Pokrovski et al., 2009; Zajacz et al.,
2010; Zhang et al., 2012), requires a revision of metal speciation models that assume H2S and HS− to be the only sulfur ligands.
Mass-dependent sulfur isotope fractionation (MDF). The finding of S−
3 in aqueous fluids may affect S isotope fractionation models, which are based on a fundamental assumption that sulfate and sulfide (±SO2 in gas phase) are the major S-bearing forms respon-sible for S isotope signatures. In particular, the formation of S−
3in sulfate–sulfide systems, as evidenced in this study, should be taken into account when interpreting kinetics of S isotope mass-dependent fractionation (MDF) between sulfate and sulfide. The actual MDF models consider the formation of either thiosulfate or polysulfide species as reaction intermediates to account for the rates of isotope exchange between sulfate and sulfide and the re-sulting 34S/32S fractionation as a function of pH and T (Ohmoto and Lasaga, 1982; Uyama et al., 1985; Chu et al., 2004). We do not detect thiosulfate ions in our experimental sulfate–sulfide systems above 250 ◦C, and we do not have clear spectroscopic evidence for polysulfides (although their spectral pattern might be hidden by polymeric molecular sulfur, see Appendix B). By contrast, in all thiosulfate and sulfate–sulfide solutions we systematically mea-sure S−
3 . In addition, our work and Pokrovski and Dubrovinsky’s (2011) study in concentrated solutions show thiosulfate decompo-sition rates at 200–300 ◦C to be higher than those predicted using Ohmoto and Lasaga’s (1982) model for dilute solutions that sug-gests thiosulfate intermediates as rate-controlling species (H2S2O3, HS2O−
3 , and S2O2−3 , depending on pH). These rates are known to
be fastest at strongly acidic pH (<2), decrease with increasing pH to ∼4, stay constant between pH 4 and 6, and then further decrease at more basic pH. This rate pattern reflects well the pH-dependent abundance of Sn (maximum concentrations at pH < 4) and S−
3 (4 < pH < 6; Fig. 6c) found in our study. The systematic presence of S−
3 in thermochemical sulfate reduction (TSR) exper-iments in various solutions at T ≤ 300 ◦C was also qualitatively
G.S. Pokrovski, J. Dubessy / Earth and Planetary Science Letters 411 (2015) 298–309 307
Fig. 7. Concentrations of S−3 and other sulfur species (in wt% S) in natural fluids as
a function of temperature. (a) An aqueous iron-poor fluid degassed from magma at 700 ◦C and containing 2 wt% Stot (H2S/SO2 molal ratio = 1) and 10 wt% NaCl equivalent. The fluid is assumed to cool down and decompress in the liquid state (from 2000 bar at 700 ◦C to 100 bar at 200 ◦C) in a porphyry-epithermal setting, with no loss of sulfur and in equilibrium with alkali aluminosilicate rocks (quartz–muscovite–(K)feldspar assemblage (QMK), pH ≈ 5 at all temperatures). (b) An aque-ous fluid of sea-water salinity (3 wt% NaCl) in equilibrium with the PPM and QMK mineral assemblages along a typical geothermal T–P gradient of prograde meta-morphism in subduction zones. The solid blue curve in each panel represents the predicted S−
3 concentration while the dashed blue curves outline the error margins associated with its estimation. Sulfate stands for the sum of SO2−
4 , HSO−4 , and their
K+ and Na+ ion pairs. Thermodynamic properties of S−3 , and other species plus
minerals are from Table 3 and C.1 (Appendix C), respectively. Concentrations of S−2
and Sn in (a) are tentative; those of S8 (not shown) are less than 0.001 wt% S. The gray shaded zone in both panels indicates the extrapolated region beyond the ex-perimental data range above 500 ◦C. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)
confirmed by in situ Raman analyses (Truche et al., 2014). Thus, S−
3 together with Sn might be a key player in sulfur redox and iso-tope exchange processes in hydrothermal systems.
Mass-independent sulfur isotope fractionation (MIF). In addition to MDF, the findings of S−
3 may provide new insights into mass-independent sulfur isotope fractionation (MIF), discovered in pyrite and barite from Archean sedimentary rocks (�33S = δ33S −0.515 ×δ34S = −4 to +14�; see Farquhar et al., 2000; Johnston, 2011;Philippot et al., 2012; for details and definitions). These MIF anomalies are believed to be due to SO2 photolysis by UV light from the Sun in an oxygen-poor atmosphere, producing sulfuric acid and elemental sulfur, which were incorporated in sulfate and sulfide minerals preserving the anomaly (e.g., Farquhar et al., 2001;Masterson et al., 2011). The rapid rise of atmospheric oxygen at ∼ 2400 Ma, shielding the Earth from the UV radiation, pre-vented SO2 photolytic reactions and caused the disappearance
of MIF in younger sediments (e.g., Farquhar and Wing, 2003;Bekker et al., 2004). This model is used for interpreting any signif-icant MIF (�33S > 0.2�), found in terrestrial samples of different age, depth, and temperature of formation, as a contribution from the Archean crust. All other inorganic or biological reactions in aqueous and mineral systems involving common S species (e.g., sulfate, sulfite, thiosulfate, sulfide) produce close to zero �33S(Farquhar and Wing, 2003; Johnston, 2011). The particular prop-erties of S−
3 make it different from the other S species as to potential MIF generation. First, its radical nature allows for so called magnetic isotope effects known for other radical species (e.g., Buchachenko, 2001). Second, S−
3 is a structural and electronic analog of ozone (O3) and its radicals, which exhibit large 17O MIF anomalies due to symmetry-driven differences in allowed quan-tum energy levels of their different isotopomers (e.g., 16O16O16O vs 16O17O18O; Rumble, 2005); this general quantum effect is also applicable to other triatomic molecules (Gao and Marcus, 2001;Babikov et al., 2003). Third, several findings of MIF (�33S from −2 to + 3�) in sulfide inclusions in diamond (Farquhar et al., 2002) and young ocean–island basalts (Cabral et al., 2013) might suggest a contribution from S−
3 , because these samples originate from high T–P magmatic or metamorphic settings in which preser-vation of the Archean crust is difficult but the conditions are fa-vorable for S−
3 formation. Finally, reports of significant �33S values (from −1.1 to 13.0�; Watanabe et al., 2009; Oduro et al., 2011) in TSR experiments in aqueous solution between 150 and 300 ◦C suggest the implication of S−
3 , alternative to thiol-disulfide radicals hypothesized in those studies. Thus, if reactions of S−
3 formation or breakdown in hydrothermal fluids generate MIF patterns, they may be used for better tracing S fluxes and redox conditions of Archean fluids and sulfur and metal sources in younger rocks and ore deposits.
6. Conclusions and perspectives
The key points of this study are the following:This work confirms the previous findings of S−
3 in aqueous solu-tion and provides a consistent set of its thermodynamic properties allowing, for the first time, quantitative predictions of S−
3 abun-dance in geological fluids across a wide T–P-depth range.
Temperatures above 250 ◦C, high dissolved S concentrations (>5000 ppm), acidic-to-neutral pH (4–6), and redox conditions enabling coexistence of sulfate/sulfur dioxide and sulfide are the main factors that favor S−
3 formation. These factors are real-ized during some stages of magmatic fluid evolution in porphyry Cu(–Au–Mo) systems leading to excess of sulfur over metals above 250 ◦C, and during prograde metamorphism of sedimentary or vol-canic rocks above 500 ◦C causing pyrite breakdown to pyrrhotite and generation of S-rich fluids forming orogenic Au deposits.
The findings of S−3 shift a long-standing paradigm that sulfide
and sulfate are the primary S species responsible for chalcophilic metal (Cu, Au, Mo, Pt,) transport in geological fluids and sulfur iso-tope signatures. Predictions of significant concentrations of S−
3 in natural and laboratory aqueous systems should motivate future ex-perimental studies to quantify the effect of S−
3 on metal transport and both sulfur MDF and MIF phenomena.
The enhanced stability of S−3 at elevated temperatures in aque-
ous systems suggests that this ion may also form in hydrous sil-icate melts. Because S−
3 breaks down to other sulfur species on quench, it may only be unambiguously analyzed by in situ spec-troscopic methods.
This study has also revealed the formation in S-rich fluids of other surprising S species. Aqueous Sn forms at T around 300 ◦C in the presence of molten sulfur and dominates over S8 – the only aqueous molecular sulfur form known to date. The radical S−
2ion, detected at 450–500 ◦C, may further be favored at magmatic
308 G.S. Pokrovski, J. Dubessy / Earth and Planetary Science Letters 411 (2015) 298–309
temperatures. These findings show that old sulfur known from an-tiquity has yet to reveal all its surprises.
Acknowledgements
This work was funded by the French National Research Agency (grant SOUMET ANR-2011-Blanc SIMI 5-6 009), the Centre Na-tional de la Recherche Scientifique (grants CESSUR-ORPY and PNP-S3MIF from the Institut des Sciences de l’Univers), University of Toulouse (grant CO2MET), and Institute Carnot (grant ISIFoR). We thank T. Chivers and M. Kokh for fascinating discussions about the S−
3 ion, T. Pokrovski for her assistance in writing, M.-C. Cau-mon, P. Robert, and A. Randi for their help with the experiments, A. Zwick and W. Rudolph for advice on Raman spectroscopy, and C. Cavaré-Hester for drawing. The article benefited from comments of the editor T. Elliott and two anonymous reviewers.
Appendix. Supplementary material
Supplementary material related to this article can be found on-line at http://dx.doi.org/10.1016/j.epsl.2014.11.035. The supplemen-tary material contains three pdf files that are cited in the main text as Appendix A (Phase composition and Raman spectra of the va-por, solid and melt phases), Appendix B (Raman spectra analysis), and Appendix C (Thermodynamic analysis and data sources).
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