Earth and Planetary Science Letters 411 (2015) 298–309
Contents lists available at ScienceDirect
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, France
b
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
Article history:
Received 6 April 2014
Received in revised form 19 November 2014
Accepted 21 November 2014
Available online xxxx
Editor: T. Elliott
Keywords:
sulfur
S−
3 ion
Raman spectroscopy
thermodynamic properties
gold deposit
sulfur isotopes
a b s t r a c t
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 lowdensity (<0.2 g/cm3 ) vapor phase, H2 S 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 massindependent 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 geological 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
2−
by sulfide (H2 S, HS− and S2− ) and sulfate (HSO−
4 and SO4 )
*
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.2014.11.035
0012-821X/© 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 intermediatevalence 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 sublimates.
This apparent simplicity of sulfur speciation at elevated T–P
contrasts 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 conditions (Fig. 1; e.g., Cotton et al., 1999; Steudel, 2003). Some of
these S forms, such as thiosulfate (S2 O23− ), polysulfides (Sn S2− ),
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
Fig. 1. Inorganic chemical forms of sulfur known in aqueous solution and nonaqueous solvents at ambient conditions.
(e.g., Giuliani et al., 2003) and as reaction intermediates in biological sulfate–sulfide redox cycle (e.g., Miluska et al., 2012). Polysulfur
−
−
radical ions (S−
2 , S3 , S4 ) are important constituents of chemical
engineering products (e.g., lithium-sulfur batteries, color pigments
and glasses, zeolites; Chivers, 1974; Chivers and Elder, 2013; references therein). To the multitude of inorganic sulfur forms of Fig. 1
may 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 being 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 unstable 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 stability 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 solubility 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 sulfate and sulfide upon aqueous fluid cooling below 400–500 ◦ C
depending on pH (Giggenbach, 1997). Thiosulfate (S2 O23− ), kinetically 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 Seward, 1997) and sulfate (e.g., Rudolph, 1996; Rudolph et al., 1997;
Schmidt, 2009; Ni and Keppler, 2012). Less attention has been devoted to systems containing intermediate-valence S species (e.g.,
polysulfides, thiosulfate, molecular sulfur, sulfur dioxide; Giggenbach, 1971, 1974; Bondarenko and Gorbaty, 1997; Yuan et al.,
2013). Pokrovski and Dubrovinsky (2011) performed Raman spectroscopy in a diamond-anvil cell on aqueous solutions in which
299
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 geochemical 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 enable, for the first time, quantitative predictions of the abundance
of S−
3 in geological fluids and evaluation of its potential geochemical 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 sulfur (Stot ) were examined in this study in the range 200–500 ◦ C
and 15–1500 bar. The source of sulfur was potassium thiosulfate
(K2 S2 O3 ± HCl in H2 O or 1 : 1 H2 O : D2 O mixture) or potassium
sulfate (K2 SO4 or KHSO4 in H2 O) plus H2 S (Table 1). Potassium
was preferred to sodium because of higher solubility of K2 SO4 than
Na2 SO4 at elevated temperatures. The solutions were prepared by
weight from analytical-grade chemicals (purity >99.9%) and deionized degassed H2 O (±D2 O) 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 observations of the H2 S vapor and liquid phase relationships in the
cell below 100 ◦ C (e.g., Appendix A).
Our experimental systems represent a good proxy for natural 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). Furthermore, they impose sulfur redox balance through stoichiometric
breakdown of thiosulfate and enable oxygen fugacity ( f O2 ) and
acidity (pH) buffering via the dominant reactions:
S2 O23− + H2 O = SO24− + H2 S
H2 S + 2O2 =
SO24−
+
SO24−
+
+ 2H
−
+ H = HSO4
(1)
(2)
(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 molec0
ular hydrogen sulfide, K+ ion pairs (KSO−
4 , KHSO4 ), and SO2 are
predicted to form with increasing T (>350–400 ◦ C), polysulfide
ions (Sn S2− ) 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 properties 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 S2 O23− breakdown and H2 S–SO4 exchange are fast enough
to reach equilibrium within hours as shown by kinetic measurements 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 1
Chemical 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
#2_11a
#2_11
#4_11
#8_11
#1_12
#5_12
#6_12
#7_12
#8_12
#3_13
1.19m K2 S2 O3
1.19m K2 S2 O3
1.19m K2 S2 O3
0.32m K2 S2 O3
0.32m K2 S2 O3
0.68m K2 S2 O3 , 0.22m HCl
0.68m K2 S2 O3 , 0.22m HCl
0.68m K2 S2 O3 , 0.22m HCl
0.68m K2 S2 O3 , 0.22m HCl, 50 wt% D2 O
1.17m KCl, 0.16m HCl
0.83m KHSO4 , 0.41m KOH, 2.78m H2 S
0.36
0.65
0.65
0.54
0.70
0.74
0.60
0.74
0.83
0.58
0.53
150–300
300–500
200–500
500
200–400
200–400
200–400
200–500
200–500
50–500
200–500
514.5
514.5
457.9
457.9
457.9
457.9
457.9
457.9
457.9
457.9
457.9
#4_13
1.24m KHSO4 , 1.69m H2 S
0.62
200–500
457.9
L + V ± S; F + B + KS at >300 ◦ C
L + V ± S at ≤350 ◦ C; F (dominant) + B + KS at ≥400 ◦ C
L + V ± S at ≤350 ◦ C; F (dominant) + B + KS at ≥400 ◦ C
SF
L + V, homogenization to L at ∼350 ◦ C, SF at 400 ◦ C
L + V, homogenization to L at 340 ◦ C, SF at 400 ◦ C
L + V, homogenization to L at 380 ◦ C; F + B + KS at 400 ◦ C
L + V, homogenization to L at 340 ◦ C; SF at 400–500 ◦ C
L + V at ≤350 ◦ C, F (dominant) + B (very minor) at ≥400 ◦ C
L + V, homogenization to L at 406 ◦ C, SF at 450–500 ◦ C
L + V + LH2 S below 50 ◦ C; L + V + S at 200–350 ◦ C;
F (dominant) + B + KS at ≥374 ◦ C
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 = H2 O–H2 S vapor; L = aqueous liquid; LH2 S = liquid H2 S; 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), protonation 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, f O2 , and pH) on the
experimental systems independently of possible presence of lowto-moderate amounts of other S species, yielding f O2 close to that
of the hematite–magnetite buffer and pH between 3 and 8 (depending of solution composition and temperature). Furthermore,
the sulfate–sulfide redox balance allows in some cases more accurate 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 preparation 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 solution and a vapor phase (±H2 S 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 expansion (<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
( P sat = PH2 O + PH2 S ) along the vapor–liquid coexistence curve up to
the homogenization temperature (T h ∼ 340–380 ◦ C, see Table 1);
above T h 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 excitation (∼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 phenomena, 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 geometry 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 acquisition 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 spectrometer was calibrated using the Raman stretching vibrations of a Si
wafer (520.7 cm−1 at 20 ◦ C), and oxygen (1555 cm−1 ) and nitrogen (2331 cm−1 ) gas from the air (Dubessy et al., 2012). In
addition, the narrow intense Raman bands of SO24− and H2 S 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 H2 O stretching band that serves as an internal standard. In addition, external standard solutions of HSO−
4,
SO24− and H2 S were measured and processed identically to the experimental 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 T
are controlled both by the ρ tot values and solute concentrations. In
all experiments, an H2 S–H2 O vapor and an aqueous liquid coexist
up to 340–350 ◦ C. Above that T , high-density moderately concentrated thiosulfate systems (ρtot ≥ 0.7 g/cm3 , ≤0.7m K2 S2 O3 , 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 polymers, 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 conditions, 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 develops, 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 patterns of the minor phases are detailed in Appendix A; those of
the dominant aqueous liquid and supercritical fluid phase are discussed 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. H2 S, SO24−
−
0
(±KSO−
4 ), and HSO4 (±KHSO4 ) are ubiquitous species in the aqueous liquid and fluid phase as indicated by their most intense
Raman peaks at 2590–2580, 980–970, and 1050–1055 cm−1 , respectively (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 (Appendix B), reflect the solubility of molten sulfur, partitioning of H2 S
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 balance 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 identified 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−
3
of the laser radiation, which results in enhancement of the symmetric 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 spectrum 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 concentration. 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 direct 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−
3 , H2 S, sulfate, and hydrogen sulfide as the major species;
Fig. 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 T
using the analogy with those of sulfate, hydrogen sulfate and H2 S;
the established coefficients were finally used to determine S−
3 concentrations 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 another 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 concentrated thiosulfate solutions (0.7m K2 S2 O3 –HCl); in more concentrated 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, Sn0 (aq), other than S08 (aq), forming in
equilibrium with molten sulfur (see Appendix B for additional arguments). Aqueous S08 itself was only detected in aqueous solution in one S-rich experiment (#4_13, KHSO4 –H2 S) at 350–500 ◦ C
(Fig. 4b) by narrow bands at ∼144 cm, 214, and 470 cm−1 , similar to the S8 ring molecules in molten sulfur (Fig. A2). The large
width of the Sn0 bands (full width at half maximum, FWHM, of
the 400 cm−1 band ≈60 cm−1 ) compared to those of the S8
molecule (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–H2 O(–NaOH)
systems using hydrothermal batch reactors based on analyses of
different S forms in the sampled fluid, which reported the presence 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−
3 ), accompanied by weaker bands at 1160 ± 10 and 1740 ±
20 cm−1 (Figs. 3, 4). The 580 and 1160 cm−1 bands may belong to
−1 ;
the following species: S−
3 (S–S asymmetric stretch, ν3 ∼ 580 cm
Chivers and Elder, 2013; references therein), SO2 (S–O symmetric 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). However, the following arguments suggest that all the three peaks arise
−
−
from S−
2 rather than S3 and SO2 : 1) the ν3 mode of S3 was not
observed below 450 ◦ C; 2) the SO2 concentrations at 500 ◦ C predicted 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 spectrum 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 similar to that of S−
3 , Appendix B), but is expected to increase above
−
500 ◦ C. Although S−
2 is encountered, together with S3 , 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.
0
In all experimental systems the formation of S−
3 (as well as Sn
−
−
and S2 ) is rapid and fully reversible. On heating, the S3 intensity
stabilizes 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
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 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 (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 quantitative reversibility with temperature (see text).
cell below 250 ◦ C results in the almost instantaneous disappearance 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 formation and abundance are independent of the initial sulfur state
(S–S bonds in S2 O23− or contrasting S redox forms in SO24− and
H2 S). All these findings suggest that S−
3 is not a short-lived transient 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 S3 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, alkali 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
303
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 solution in some earlier studies (Giggenbach, 1971; Uyama et al., 1985;
Bondarenko and Gorbaty, 1997) were erroneously attributed to S−
2
or other species (see Appendix C). Note that hydrothermal batchreactor experiments in S-rich high-T aqueous solutions (Dadze and
Sorokin, 1993; Pokrovski et al., 2008) did not detect S−
3 likely because 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 unambiguously 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 tendency for S-bearing radicals, among which S−
3 is the most energetically favorable (e.g., Steudel and Steudel, 2013). Furthermore,
additional factors at the molecular level may contribute to its enhanced 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 H2 O
molecule, may allow energetically favorable hydration structures
in the H2 O 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 silicate cages of zeolites (Reinen and Lindner, 1999) or by forming
−
+ −
Na+ S−
3 or K S3 ion pairs. Thus, the effect of fluid salinity on S3
stability 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 hydration energy of ions in this density range (Pokrovski et al., 2013).
Following our own observations and the entropy principles postulating 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 S3
above 500–600 ◦ C.
4.2. Thermodynamic properties of S−
3
The equilibrium concentrations of S−
3 measured in this study allow 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 molal Gibbs free energy G0T , P was varied to match its Raman-derived
concentrations at each T–P-composition point (Table 2). The minor S species in solution (Sn , Sn2− , 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 G0T , 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 thermodynamic equilibrium constant of the reaction:
2H2 S(aq) + SO24− (aq) + H+ (aq)
= S−
3 (aq) + 0.75O2 (gas) + 2.5H2 O(liq),
log10 K 4
(4)
The values of log10 K 4 are reported in Table 2 and compared
with literature data in Fig. 5. The data from this study are identical within errors at a given T in different runs over a large range
of S concentration; this consistency further strengthens the validity of our derivations. Our log K 4 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 2
−
Concentrations of S−
3 in the Raman experiments as a function of T and P , and the corresponding Gibbs free energies of S3 and equilibrium constants (log10 K ) for reaction (4):
2H2 S(aq) + SO24− (aq) + H+ (aq) = S−
3 (aq) + 0.75O2 (gas) + 2.5H2 O(liq), derived in this study.
T
(◦ C)
P
(bar)a
m(S−
3)
(mol/kg H2 O)
G0T , P (S−
3)
log10 K 4
(kcal/molb )
#1&2_11: 1.19m K2 S2 O3
200
250
300
350
400
450
500
30
59
90
182
500
600
750
0.0014 ± 0.0006
0.011 ± 0.007
0.028 ± 0.010
0.070 ± 0.014
0.098 ± 0.020
0.112 ± 0.017
0.114 ± 0.029
7.2 ± 1.0
3.1 ± 0.7
1.8 ± 1.0
−1.4 ± 1.4
−3.2 ± 1.7
−5.3 ± 1.2
−7.9 ± 1.7
−24.4 ± 0.9
−19.5 ± 0.6
−16.2 ± 0.7
−12.0 ± 1.0
−10.4 ± 1.1
−7.7 ± 0.7
−5.3 ± 0.9
#4&8_11: 0.32m K2 S2 O3
300
400
500
93
650
750
0.0028 ± 0.0010
0.0033 ± 0.0009
0.0031 ± 0.0006
1.4 ± 0.7
−3.0 ± 1.0
−7.8 ± 1.4
−16.1 ± 0.5
−10.8 ± 0.6
−5.3 ± 0.8
#1&5&6_12: 0.68m K2 S2 O3 + 0.22m HCl
250
60
300
100
350
200
400
500
450
950
500
1400
0.005 ± 0.003
0.032 ± 0.020
0.036 ± 0.012
0.061 ± 0.012
0.070 ± 0.010
0.065 ± 0.010
3.8 ± 1.0
1.2 ± 0.7
0 .0 ± 0.5
−4.1 ± 0.5
−7.4 ± 0.7
−10.0 ± 1.2
−19.8 ± 0.8
−16.0 ± 0.5
−12.7 ± 0.3
−10.1 ± 0.3
−8.1 ± 0.4
−6.1 ± 0.7
#3_13: 0.83m KHSO4 + 0.41m KOH + 2.78m H2 S
300
133
350
201
375
300
400
350
450
500
500
700
0.058 ± 0.006
0.13 ± 0.05
0.09 ± 0.06
0.095 ± 0.060
0.181 ± 0.060
0.204 ± 0.058
3.7 ± 1.0
0.5 ± 2.4
0.5 ± 2.4
−0.7 ± 1.9
−4.3 ± 2.6
−7.2 ± 2.9
−16.9 ± 0.7
−12.9 ± 1.7
−12.3 ± 1.6
−10.6 ± 1.2
−7.2 ± 1.6
−5.1 ± 1.6
#4_13: 1.24m KHSO4 + 1.69m H2 S
300
109
350
185
400
350
450
500
500
700
0.015 ± 0.008
0.049 ± 0.018
0.049 ± 0.018
0.069 ± 0.030
0.191 ± 0.051
2.4 ± 1.9
−0.2 ± 1.2
−2.2 ± 1.2
−4.3 ± 1.9
−8.4 ± 1.7
−16.4 ± 1.5
−12.6 ± 0.8
−10.1 ± 0.8
−7.2 ± 1.2
−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–H2 O system (Driesner and Heinrich, 2007) for PH2 O plus PH2 S from Raman
measurements in the vapor phase (below T h ). Uncertainties of the P estimation are within ±20 bar below T h , ±50 bar above T h , 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 reciprocal 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 K 4 = −(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 modeling (Tossell, 2012); however, the uncertainties associated with
that work are unknown and the reported absolute K 4 values would
imply S−
3 to be at least ∼10 times more abundant at the conditions 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 reaction (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 Dubrovinsky (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
Table 3
Standard molal thermodynamic properties at 25 ◦ C and 1 bar and parameters of the revised HKF equation of state for S−
3 retrieved in this study.
Thermodynamic propertya
S−
3 ion
Gibbs energy f G , kcal/mol
Enthalpy, f H0 , kcal/mol
Entropy, S0 , cal/(mol K)
Heat capacity, C p0 , cal/(mol K)
Volume, V 0 , cm3 /mol
13.16 ± 3.50b
10.84 ± 2.50b
28.6 ± 8.0b
62.3 ± 12.8b
37.7 ± 3.0b
HKF coefficientsc
a1 × 10, cal/(mol bar)
a2 × 10−2 , cal/mol
a3 , cal K/(mol bar)
a4 × 10−4 , cal K/mol
c 1 , cal/(mol K)
c 2 × 10−4 , cal K/mol
ω × 10−5 , cal/mol
2.5 ± 4.0
19.9 ± 11.0
9.2 ± 5.0
−3.6 ± 2.0
50.2 ± 5.0
9.6 ± 2.6
0.8 ± 0.2
0
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 G01 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 C p0 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−
3
from this study, Pokrovski and Dubrovinsky (2011) and revised data of
Giggenbach (1971) at a function of T and P , using the HKF model equations 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 predictive capacity of the revised HKF model (Oelkers et al., 2009;
Sverjensky et al., 2014), these data enable quantification of S−
3 concentrations 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 conditions and fluid compositions using the HCh computer code. Fig. 6
shows 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 pressure 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 f O2 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 abundance 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 decompression upon its rise in a porphyry–epithermal system. Total salt
(10 wt% NaCl + KCl) and S (2 wt%) contents and sulfur speciation (H2 S : 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;
305
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 (H2 S–SO2 -sulfate; f O2 ∼ 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 H2 S (see Appendix C). Note that the situation shown in Fig. 7a represents the optimal conditions for S−
3
in 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 acidity (pH < 3) produced by SO2 disproportionation to H2 S and sulfuric 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 H2 S, 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 separation in porphyry systems is expected to lead to (partial) breakdown of S−
3 , which is not stable in the low-density vapor phase
(see Section 4.1 and Fig. 6d). Upon the H2 S–SO2 vapor ascent
and cooling, S−
3 may re-form in the condensed S-rich liquid carrying 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 system.
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 liberating sulfur and accompanied by chlorite breakdown releasing water (Tomkins, 2010), generates S-rich fluids that may carry large
amounts of S−
3 . This is illustrated in Fig. 7b showing the solubility of the pyrite–pyrrhotite–magnetite assemblage along a typical
geothermal gradient of regional Archean metamorphism and modern hot subduction zones (Peacock, 1990). While S−
3 is negligible
below 500 ◦ C (<0.001 wt% S) because of the low aqueous Stot concentration (<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 margins 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 f O2 buffered by the HM assemblage; (b) oxygen fugacity at pH ∼ 5; (c) fluid acidity
at f O2 between HM (acidic pH) and NNO (basic pH); (d) pressure at pH ∼ 5 and f O2 of HM (at ≥0.3 kbar – the pressure above which the HKF model is valid at 450 ◦ C).
+ and Na+ ion pairs); the vertical dashed lines in (b) denote
Curves show concentrations of each indicated species (sulfate stands for the sum of SO24− , HSO−
4 , and their K
the f O2 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 S3 , 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 geological fluids may have far-reaching implications for the formation
of economic ore deposits of Cu, Au, Mo, and Pt. This is because S−
3 is expected to form stable complexes with these metals in aqueous solution, similar to traditional polysulfides (Sn S2− ;
Berndt et al., 1994; Rickard and Luther, 2006; Liu et al., 2013)
and related sulfur-nitrogen ions (e.g., S3 N− ; Bojes et al., 1981,
1982). Thus S−
3 may compete with hydrogen sulfide (H2 S 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
−
+ and Cu+ in aqueof S−
3 , comparable to that of HS , for Au
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, potentially, 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 porphyry 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 complexes with different metals. Furthermore, the presence of significant 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 H2 S 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 responsible for S isotope signatures. In particular, the formation of S−
3
in sulfate–sulfide systems, as evidenced in this study, should be
taken into account when interpreting kinetics of S isotope massdependent 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 resulting 34 S/32 S 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 measure S−
3 . In addition, our work and Pokrovski and Dubrovinsky’s
(2011) study in concentrated solutions show thiosulfate decomposition rates at 200–300 ◦ C to be higher than those predicted using
Ohmoto and Lasaga’s (1982) model for dilute solutions that suggests thiosulfate intermediates as rate-controlling species (H2 S2 O3 ,
2−
HS2 O−
3 , and S2 O3 , 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 pHdependent 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) experiments 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
of MIF in younger sediments (e.g., Farquhar and Wing, 2003;
Bekker et al., 2004). This model is used for interpreting any significant MIF (33 S > 0.2h), 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 33 S
(Farquhar and Wing, 2003; Johnston, 2011). The particular properties 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 17 O MIF
anomalies due to symmetry-driven differences in allowed quantum energy levels of their different isotopomers (e.g., 16 O16 O16 O
vs 16 O17 O18 O; 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 (33 S from
−2 to + 3h) 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 preservation of the Archean crust is difficult but the conditions are fa33
vorable for S−
3 formation. Finally, reports of significant S values
(from −1.1 to 13.0h; 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.
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 (H2 S/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 aqueous 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 metamorphism 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 SO24− , 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 experimental 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 isotope 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 massindependent sulfur isotope fractionation (MIF), discovered in pyrite
and barite from Archean sedimentary rocks (33 S = δ 33 S − 0.515 ×
δ 34 S = −4 to +14h; 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, prevented SO2 photolytic reactions and caused the disappearance
6. Conclusions and perspectives
The key points of this study are the following:
This work confirms the previous findings of S−
3 in aqueous solution and provides a consistent set of its thermodynamic properties
allowing, for the first time, quantitative predictions of S−
3 abundance 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 realized 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 volcanic 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 isotope signatures. Predictions of significant concentrations of S−
3 in
natural and laboratory aqueous systems should motivate future experimental 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 aqueous systems suggests that this ion may also form in hydrous silicate melts. Because S−
3 breaks down to other sulfur species on
quench, it may only be unambiguously analyzed by in situ spectroscopic 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−
2
ion, 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 antiquity 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 National de la Recherche Scientifique (grants CESSUR-ORPY and PNPS3MIF from the Institut des Sciences de l’Univers), University of
Toulouse (grant CO2 MET), 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. Caumon, 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 online at http://dx.doi.org/10.1016/j.epsl.2014.11.035. The supplementary material contains three pdf files that are cited in the main text
as Appendix A (Phase composition and Raman spectra of the vapor, solid and melt phases), Appendix B (Raman spectra analysis),
and Appendix C (Thermodynamic analysis and data sources).
References
Amend, J.P., Shock, E.L., 2001. Energetics of overall metabolic reactions of thermophilic and hiperthermophilic Archaea and Bacteria. FEMS Microbiol. Rev. 25,
175–243.
Babikov, D., Kendrick, B.K., Walker, R.B., Schinke, R., Pack, R.T., 2003. Quantum origin of an anomalous isotope effect in ozone formation. Chem. Phys. Lett. 372,
686–691.
Barnes, H.L., 1997. Geochemistry of Hydrothermal Ore Deposits. Wiley, New York.
Bekker, A., Holland, H.D., Wang, P.-L., Rumble III, D., Stein, H.J., Hannah, J.L., Coetzee, L.L., Beukes, N.J., 2004. Dating the rise of atmospheric oxygen. Nature 427,
117–120.
Berndt, M.E., Buttram, T., Earley III, D., Seyfried, W.E., 1994. The stability of gold
polysulfide complexes in aqueous sulfide solutions: 100 to 150 ◦ C and 100 bars.
Geochim. Cosmochim. Acta 58, 587–594.
Bojes, J., Chivers, T., Codding, P.W., 1981. Preparation and X-ray crystal structures of copper(I) complexes of the S3 N− ion. J. Chem. Soc., Chem. Commun.,
1171–1173.
Bojes, J., Chivers, T., Laidlaw, W.G., Trsic, M., 1982. The thio analogue of peroxynitrite, [SSNS]− : preparation, electronic structure, resonance Raman spectrum,
and formation of complexes with nickel and cobalt. J. Am. Chem. Soc. 104,
4837–4841.
Bondarenko, G.V., Gorbaty, Y.E., 1997. In situ Raman spectroscopic study of sulfursaturated water at 1000 bar between 200 and 500 ◦ C. Geochim. Cosmochim.
Acta 61, 1413–1420.
Boyle, R.W., 1969. Hydrothermal transport and deposition of gold. Econ. Geol. 64,
112–115.
Brimhall, G.H., Crerar, D.A., 1987. Ore fluids: magmatic to supergene. Rev. Miner. 17,
235–321.
Buchachenko, A.L., 2001. Magnetic isotope effect: nuclear spin control of chemical
reactions. J. Phys. Chem. A 105, 9995–10011.
Cabral, R.A., et al., 2013. Anomalous sulphur isotopes in plume lavas reveal deep
mantle storage of Archaean crust. Nature 496, 490–493.
Caumon, M.-C., Dubessy, J., Robert, P., Tarantola, A., 2013. Fused silica capillary
capsules (FSCCs) as reference synthetic aqueous fluid inclusions to determine
chlorinity by Raman spectroscopy. Eur. J. Mineral. 25, 755–763.
Chivers, T., 1974. Ubiquitous trisulfur radical ion S−
3 . Nature 252, 32–33.
Chivers, T., Drummond, I., 1972. Characterization of the trisulfur radical ion S−
3
in blue solutions of alkali polysulfides in hexamethylphosphoramide. Inorg.
Chem. 11, 2525–2527.
Chivers, T., Elder, P.J.W., 2013. Ubiquitous trisulfur radical ion: fundamentals and applications in materials science, electrochemistry, analytical chemistry and geochemistry. Chem. Soc. Rev. 42, 5996–6005.
Chivers, T., Lau, C., 1982. Raman spectroscopic identification of the S4 N− and S−
3
ions in blue solutions of sulfur in liquid ammonia. Inorg. Chem. 21, 453–455.
Chivers, T., Gilmour, E., Kydd, R.A., 1978. Vibrational spectra of sulphur-doped borate
glasses. J. Mater. Sci. 13, 1585–1588.
Chou, I.-M., Song, Y., Burruss, R.C., 2008. A new method for synthesizing fluid inclusions in fused silica capillaries containing organic and inorganic material.
Geochim. Cosmochim. Acta 72, 5217–5231.
Chu, X., Ohmoto, H., Cole, D.R., 2004. Kinetics of sulfur isotope exchange between
aqueous sulfide and thiosulfate involving intra- and intermolecular reactions at
hydrothermal conditions. Chem. Geol. 211, 217–235.
Clark, R.J.H., Cobbold, D.G., 1978. Characterization of sulfur radical anions in solutions of alkali polysulfides in dimethylformamide and dexamethylphosphoramide and in the solid state in ultramarine blue, green, and red. Inorg.
Chem. 17, 3169–3174.
Clark, R.J.H., Franks, M.L., 1975. The resonance Raman spectrum of ultramarine blue.
Chem. Phys. Lett. 34, 69–72.
Cotton, F.A., Wilkinson, G., Murillo, C.A., Bochmann, M., 1999. Advanced Inorganic
Chemistry, 6th edition. Wiley, Chichester. 1355 pp.
Dadze, T.P., Sorokin, V.I., 1993. Experimental determination of the concentrations of
0
H2 S, HSO−
4 , SO2aq , H2 S2 O3 , Saq , and Stot in the aqueous phase in the S–H2 O
system at elevated temperatures. Geochem. Int. 30, 38–53.
Dargent, M., Dubessy, J., Truche, L., Bazarkina, E.F., Nguyen-Trung, C., Robert, P.,
2013. Experimental study of uranyl(VI) chloride complex formation in acidic LiCl
aqueous solutions under hydrothermal conditions (T = 21 ◦ C–350 ◦ C, Psat) using Raman spectroscopy. Eur. J. Mineral. 25, 765–775.
Driesner, T., Heinrich, C.A., 2007. The system H2 O–NaCl. Part I: Correlation formulae for phase relations in temperature–pressure-composition space from 0
to 1000 ◦ C, 0 to 5000 bar, and 0 to 1 XNaCl . Geochim. Cosmochim. Acta 71,
4880–4901.
Dubessy, J., Ding, J., Robert, P., Leisen, M., Giraud, A., 2009. Capillary Techniques
Combined with Vibrational Spectroscopies Applied to Fluid Inclusions and Phase
Equilibria, ECROFI, Granada, Spain.
Dubessy, J., Caumon, M.-C., Rull, F., Sharma, S., 2012. Instrumentation in Raman
spectroscopy: elementary theory and practice. In: Dubessy, J., Caumon, M.-C.,
Rull, F. (Eds.), Raman Spectroscopy Applied to Earth Sciences and Cultural Heritage. In: EMU Notes in Mineralogy, vol. 12. European Mineralogical Union and
the Mineralogical Society of Great Britain & Ireland, London, pp. 83–165.
Einaudi, M.T., Hedenquist, J.W., Inan, E.E., 2003. Sulfidation state of fluids in active
and extinct hydrothermal systems: transitions from porphyry to epithermal environments. Soc. Econ. Geol. Spec. Publ. 10, 285–313.
Ellis, A.J., Giggenbach, W., 1971. Hydrogen sulphide ionization and sulphur hydrolysis in high-temperature solution. Geochim. Cosmochim. Acta 35, 247–260.
Farquhar, J., Wing, B.A., 2003. Multiple sulfur isotopes and the evolution of the atmosphere. Earth Planet. Sci. Lett. 213, 1–13.
Farquhar, J., Bao, H., Thiemens, M., 2000. Atmospheric influence of Earth’s earliest
sulfur cycle. Science 289, 756–758.
Farquhar, J., Savarino, J., Airieau, S., Thiemens, M.H., 2001. Observation of
wavelength-sensitive mass-independent sulfur isotope effects during SO2 photolysis: implications for the early atmosphere. J. Geophys. Res., Planets 106,
32829–32839.
Farquhar, J., Wing, B.A., McKeegan, K.D., Harris, J.W., Cartigny, P., Thiemens, M.H.,
2002. Mass-independent sulphur of inclusions in diamond and sulphur recycling
on early Earth. Science 298, 2369–2372.
Gao, Y., Marcus, R.A., 2001. Strange and unconventional isotope effects in ozone formation. Science 293, 259–263.
Giggenbach, W., 1970. Optical spectra of highly alkaline sulphide solutions and the
second dissociation constant of hydrogen sulfide. Inorg. Chem. 10, 1333–1338.
Giggenbach, W., 1971. The blue solutions of sulfur in water at elevated temperatures. Inorg. Chem. 10, 1306–1308.
Giggenbach, W., 1974. Equilibria involving polysulfide ions in aqueous sulfide solutions up to 240 ◦ C. Inorg. Chem. 13, 1724–1730.
Giggenbach, W.F., 1997. The origin and evolution of fluids in magmatic–
hydrothermal systems. In: Barnes, H.L. (Ed.), Geochemistry of Hydrothermal Ore
Deposits. John Wiley & Sons, New York, pp. 737–796.
Giuliani, G., Dubessy, J., Banks, D., Vinh, H.Q., Lhomme, T., Pironon, J., Garnier, V.,
Trinh, V.T., Long, P.V., Ohnenstetter, D., Schwarz, D., 2003. CO2 –H2 S–COS–S8 –
AlO(OH)-bearing fluid inclusions in ruby from marble-hosted deposits in Luc
Yen area, North Vietnam. Chem. Geol. 194, 167–185.
Hayashi, K.I., Ohmoto, H., 1991. Solubility of gold in NaCl- and H2 S-bearing aqueous
solutions at 250–350 ◦ C. Geochim. Cosmochim. Acta 55, 2111–2126.
Hedenquist, J.W., Lowenstern, J.B., 1994. The role of magmas in the formation of
hydrothermal ore deposits. Nature 370, 519–527.
Heinrich, C.A., 2005. The physical and chemical evolution of low-salinity magmatic
fluids at the porphyry to epithermal transition: a thermodynamic study. Miner.
Depos. 39, 864–889.
Helgeson, H.C., Kirkham, D.H., Flowers, G.C., 1981. Theoretical prediction of the thermodynamic behaviour 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 600 ◦ C and 5 kb. Am.
J. Sci. 291, 1249–1516.
Jacquemet, N., Guillaume, D., Zwick, A., Pokrovski, G.S., 2014. In situ Raman spectroscopy identification of the S−
3 ion in S-rich hydrothermal fluids from synthetic fluid inclusions. Am. Mineral. 99, 1109–1118.
Johnson, J.W., Oelkers, E.H., Helgeson, H.C., 1992. SUPCRT92: a software package
for calculating the standard molal thermodynamic properties of minerals, gases,
G.S. Pokrovski, J. Dubessy / Earth and Planetary Science Letters 411 (2015) 298–309
aqueous species, and reactions from 1 to 5000 bar and 0 to 1000 ◦ C. Comput.
Geosci. 18, 899–947. http://geopig.asu.edu/?q=tools.
Johnston, D.T., 2011. Multiple sulfur isotopes and the evolution of Earth’s surface
sulfur cycle. Earth-Sci. Rev. 106, 161–183.
Kaasalainen, H., Stefánsson, A., 2011. Sulfur speciation in natural hydrothermal waters, Iceland. Geochim. Cosmochim. Acta 75, 2777–2791.
Kouzmanov, K., Pokrovski, G.S., 2012. Hydrothermal controls on metal distribution
in porphyry Cu(–Au–Mo) systems. Soc. Econ. Geol. Spec. Publ. 16, 573–618.
Ledé, B., Demortier, A., Gobeltz-Hautecoeur, N., Lelieur, J.-P., Picquenard, E., Duhayon,
C., 2007. Observation of the ν3 Raman band of S−
3 inserted into sodalite cages.
J. Raman Spectrosc. 38, 1461–1468.
Liu, X., Sprik, M., Cheng, J., 2013. Hydration, acidity and metal complexing of
polysulfide species: a first principles molecular dynamics study. Chem. Phys.
Lett. 563, 9–14.
Loucks, R.R., Mavrogenes, J.A., 1999. Gold solubility in supercritical hydrothermal
brines measured in synthetic fluid inclusions. Science 284, 2159–2163.
Mandeville, C.W., 2010. Sulfur: a ubiquitous and useful tracer in Earth and planetary
sciences. Elements 6, 75–80.
Martell, A.E., Hancock, R.D., 1996. Metal Complexes in Aqueous Solutions. Plenum,
New York.
Masterson, A.L., Farquhar, J., Wing, B.A., 2011. Sulfur mass-independent fractionation
patterns in the broadband UV photolysis of sulfur dioxide: pressure and third
body effects. Earth Planet. Sci. Lett. 306, 253–260.
Mei, Y., Sherman, D.M., Liu, W., Brugger, J., 2013. Complexation of gold in S−
3 -rich
hydrothermal fluids: evidence from ab-initio molecular dynamics simulations.
Chem. Geol. 347, 34–42.
Métrich, N., Berry, A.J., O’Neill, St C, H., Susini, J., 2009. The oxidation state of sulfur
in synthetic and natural glasses determined by X-ray absorption spectroscopy.
Geochim. Cosmochim. Acta 73, 2382–2399.
Meyer, B., 1976. Elemental sulfur. Chem. Rev. 76, 367–388.
Miluska, J., et al., 2012. Zero-valent sulphur is a key intermediate in marine methane
oxidation. Nature 491, 541–546.
Murray, R.C., Cubicciotti, D., 1983. Thermodynamics of aqueous sulphur species to
300 ◦ C and potential-pH diagrams. J. Electrochem. Soc. 130, 866–869.
Ni, H., Keppler, H., 2012. In-situ Raman spectroscopic study of sulfur speciation in
oxidized magmatic–hydrothermal fluids. Am. Mineral. 97, 1348–1353.
Oduro, H., Harms, B., Sintim, H.O., Kaufman, A.J., Cody, G., Farquhar, J., 2011. Evidence of magnetic isotope effects during thermochemical sulfate reduction.
Proc. Natl. Acad. Sci. USA 108, 17635–17638.
Oelkers, E.H., Bénézeth, P., Pokrovski, G.S., 2009. Thermodynamic databases for
water–rock interaction. Rev. Mineral. Geochem. 70, 1–46.
Ohmoto, H., Lasaga, A.C., 1982. Kinetics of reactions between aqueous sulfates and
sulfides in hydrothermal systems. Geochim. Cosmochim. Acta 46, 1727–1745.
Peacock, S.M., 1990. Fluid processes in subduction zones. Science 248, 329–337.
Philippot, P., van Zuilen, M., Rollion-Bard, C., 2012. Variations in atmosphere sulphur
chemistry on early Earth linked to volcanic activity. Nat. Geosci. 5, 668–674.
Pokrovski, G.S., Dubrovinsky, L.S., 2011. The S−
3 ion is stable in geological fluids at
elevated temperatures and pressures. Science 331, 1052–1054.
Pokrovski, G.S., Borisova, A.Y., Harrichoury, J.-C., 2008. The effect of sulfur on
vapor-liquid fractionation of metals in hydrothermal systems. Earth Planet. Sci.
Lett. 266, 345–362.
Pokrovski, G.S., Tagirov, B.R., Schott, J., Hazemann, J.-L., Proux, O., 2009. A new view
on gold speciation in sulfur-bearing hydrothermal fluids from in-situ X-ray absorption spectroscopy and quantum-chemical modeling. Geochim. Cosmochim.
Acta 73, 2506–2527.
Pokrovski, G.S., Borisova, A.Y., Bychkov, A.Y., 2013. Speciation and transport of metals
and metalloids in geological vapors. Rev. Mineral. Geochem. 76, 165–218.
Pokrovski, G.S., Akinfiev, N.N., Borisova, A.Y., Zotov, A.V., Kouzmanov, K., 2014. Gold
speciation and transport in geological fluids: insights from experiments and
physical-chemical modeling. In: Garofalo, P., Ripley, E. (Eds.), Gold-Transporting
Fluids in the Earth’s Crust. In: Geological Society of London Special Publications,
vol. 402, pp. 9–70.
Reinen, D., Lindner, G.-G., 1999. The nature of the chalcogen colour centres in
ultramarine-type solids. Chem. Soc. Rev. 28, 75–84.
Rickard, D., Luther III, G.W., 2006. Metal sulfide complexes and clusters. Rev. Mineral. Geochem. 61, 421–504.
Richards, J.P., 2011. Magmatic to hydrothermal metal fluxes in convergent and collided margins. Ore Geol. Rev. 40, 1–26.
Risberg, E.D., Eriksson, L., Mink, J., Pettersson, L.G.M., Skripkin, M.Y., Sandström, M.,
2007. Sulfur x-ray absorption and vibrational spectroscopic study of sulphur
dioxide, and sulfonate solutions and of the substituted sulfonate ions X3 CSO−
3
(X = H, Cl. F.). Inorg. Chem. 46, 8332–8348.
309
Rudolph, W., 1996. Structure and dissociation of the hydrogen sulfate ion in aqueous
solution over a broad temperature range: a Raman study. Z. Phys. Chem. 194,
73–95.
Rudolph, W., Brooker, M.H., Tremain, P.R., 1997. Raman spectroscopic investigation
of aqueous FeSO4 in neutral and acidic solutions from 25 ◦ C to 303 ◦ C: innerand outer-sphere complexes. J. Solution Chem. 26, 757–777.
Rumble, D., 2005. A mineralogical and geochemical record of atmospheric photochemistry. Am. Mineral. 90, 918–930.
Schmidt, C., 2009. Raman spectroscopic study of a H2 O + Na2 SO4 solution at
21–600 ◦ C and 0.1 MPa to 1.1 GPa: relative differential v1 –SO24− Raman scattering cross sections and evidence of the liquid–liquid transition. Geochim. Cosmochim. Acta 73, 425–437.
Schulte, M.D., Rogers, K.L., 2004. Thiols in hydrothermal solution: standard partial
molal properties and their role in the organic geochemistry of hydrothermal
environments. Geochim. Cosmochim. Acta 68, 1087–1097.
Seo, J.H., Guillong, M., Heinrich, C.A., 2009. The role of sulfur in the formation
of magmatic–hydrothermal copper–gold deposits. Earth Planet. Sci. Lett. 282,
323–328.
Seward, T.M., 1973. Thio complexes of gold and the transport of gold in hydrothermal ore solutions. Geochim. Cosmochim. Acta 37, 379–399.
Shock, E.L., Helgeson, H.C., 1988. Calculation of the thermodynamic and transport
properties of aqueous species at high pressures and temperatures: correlation
algorithms for ionic species and equation of state predictions to 5 kb and
1000 ◦ C. Geochim. Cosmochim. Acta 52, 2009–2036.
Shvarov, Y.V., 2008. HCh: new potentialities for the thermodynamic simulation of
geochemical systems offered by windows. Geochem. Int. 46 (8), 834–839.
Shvarov, Y.V., in press. A suite of programs, OptimA, OptimB, OptimC, and
OptimS compatible with the Unitherm database, for deriving the thermodynamic properties of aqueous species from solubility, potentiometry
and spectroscopy measurements. Appl. Geochem. http://dx.doi.org/10.1016/
j.apgeochem.2014.11.021. Programs are available at: http://www.geol.msu.ru/
deps/geochems/soft/index_e.html.
Steudel, R., 2003. Elemental Sulfur and Sulfur-Rich Compounds I, II. Springer, Berlin.
Steudel, R., Steudel, Y., 2013. Polysulfide chemistry in sodium–sulfur batteries and
related systems – a computational study by G3X(PM2) and PCM calculations.
Chem. Eur. J. 19, 3162–3176.
Suleimenov, O.M., Seward, T.M., 1997. A spectrophotometric study of hydrogen sulphide ionization in aqueous solutions to 350 ◦ C. Geochim. Cosmochim. Acta 61,
5187–5198.
Sverjensky, D.A., Harrison, B., Azzolini, D., 2014. Water in the deep Earth: the dielectric constant and the solubilities of quartz and corundum to 60 kb and 1200 ◦ C.
Geochim. Cosmochim. Acta 129, 125–145.
Tomkins, A.G., 2010. Windows of metamorphic sulfur liberation in the crust: implications for gold deposit genesis. Geochim. Cosmochim. Acta 74, 3246–3259.
Tossell, J.A., 2012. Calculation of the properties of the S−
3 radical anion and its complexes with Cu+ in aqueous solution. Geochim. Cosmochim. Acta 95, 79–92.
Truche, L., Bazarkina, E.F., Barré, G., Thomassot, E., Berger, G., Dubessy, J., Robert, P.,
2014. The role of S−
3 ion in thermochemical sulphate reduction: geological and
geochemical implications. Earth Planet. Sci. Lett. 396, 190–200.
Uyama, F., Chiba, H., Kusakabe, M., Sakai, H., 1985. Sulfur isotope exchange reaction
in the aqueous system: thiosulfate–sulfide–sulfate at hydrothermal temperature.
Geochem. J. 19, 301–315.
Wallace, P.J., 2001. Volcanic SO2 emissions and the abundance and distribution of
exsolved gas in magma bodies. J. Volcanol. Geotherm. Res. 108, 85–106.
Wang, X., Chou, I.-M., Hu, W., Burruss, R.C., Sun, Q., Song, Y., 2011. Raman spectroscopic measurements of CO2 density: experimental calibration with highpressure optical cell (HPOC) and fused capillary capsule (FSCC) with application
to fluid inclusions observations. Geochim. Cosmochim. Acta 75, 4080–4093.
Watanabe, Y., Farquhar, J., Ohmoto, H., 2009. Anomalous fractionations of sulfur isotopes during thermochemical sulfate reduction. Science 324, 370–373.
Williamson, M.A., Rimstidt, J.D., 1992. Correlation between structure and thermodynamic properties of aqueous sulfur species. Geochim. Cosmochim. Acta 56,
3867–3880.
Yuan, S., Chou, I.-M., Burruss, R.C., Wang, X., Li, J., 2013. Disproportionation and
thermochemical sulfate reduction reactions in S–H2 O–CH4 and S–D2 O–CH4 systems from 200 to 340 ◦ C at elevated pressures. Geochim. Cosmochim. Acta 118,
263–275.
Zajacz, Z., Seo, J.H., Candela, P.A., Piccoli, P.M., Heinrich, C.A., 2010. Alkali metals control the release of gold from volatile-rich magmas. Earth Planet. Sci. Lett. 297,
50–56.
Zhang, L., Audétat, A., Dolejš, D., 2012. Solubility of molybdenite (MoS2 ) in aqueous fluids at 600–800 ◦ C, 200 MPa: a synthetic fluid inclusions study. Geochim.
Cosmochim. Acta 77, 175–185.