The dissolution mechanism of sulphur in hydrous silicate melts. II: Solubility and speciation of sulphur in hydrous silicate melts as a function of fO2

Chemical Geology 322–323 (2012) 250–267

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The dissolution mechanism of sulphur in hydrous silicate melts. II: Solubility and speciation of sulphur in hydrous silicate melts as a function of fO2 Kevin Klimm a, b,⁎, Simon C. Kohn a, Roman E. Botcharnikov c a b c

School of Earth Sciences, University of Bristol, Wills Memorial Building, Bristol BS8 1RJ, UK Institut für Geowissenschaften, Goethe Universität Frankfurt, 60438 Frankfurt am Main, Germany Institut für Mineralogie, Leibniz Universität Hannover, Callinstrasse 3, 30167 Hannover, Germany

a r t i c l e

i n f o

Article history: Received 18 November 2011 Received in revised form 21 April 2012 Accepted 26 April 2012 Available online 5 May 2012 Editor: D.B. Dingwell Keywords: Sulphur solubility Sulphur speciation S K-edge XANES Raman spectroscopy Oxygen fugacity Silicate melt

a b s t r a c t Raman and X-ray absorption spectroscopy (XANES) measurements on a series of experimentally synthesised, sulphur (S)-bearing, hydrous silicate glasses were used to determine the S-speciation and S-oxidation state as a function of glass composition and oxygen fugacity (fO2) and to decipher the dissolution mechanism of S in silicate melts. Synthesised glasses include soda-lime (SLG), K2Si4O9 (KSG), albite and trondhjemite (TROND) compositions. A series of SLG and KSG glasses, doped with small quantities of Fe, was also studied in order to determine the effect of Fe/S on the S solubility. The experiments were performed in internally heated (IHPV) and cold seal (CSPV) pressure vessels at 200 MPa, 1000 and 850 °C and a range of fO2 from log fO2 = QFM − 2.35 to QFM + 4 (QFM is quartz–fayalite–magnetite oxygen buffer). The systematic correlation of features in Raman and XANES spectra allows the identification of at least four different S-species in the glasses depending on fO2 and Fe/S of the system. In XANES spectra of Fe-free glasses SH−, H2S and SO42 − are visible as peaks at 2466, 2471.8 and 2482 eV, respectively. In Raman spectra peaks at 2574 and 990 cm − 1 indicate the presence of H\S bonds and SO42 −, respectively, but SH− and H2S can not be distinguished using a Raman spectroscopy. In Fe-bearing glasses Fe\S bonding is identified at 2469 eV in the case of XANES and at 298, 372 and 420 cm − 1 in the case of Raman spectra. The intensities of peaks related to S\H bonding systematically decrease and the intensities of peaks related to Fe\S bonding systematically increase with increasing Fe/S in both the XANES and the Raman spectra indicating that in the presence of Fe, Fe\S bonding is preferred over S\H bonding. The total S solubility at sulphur saturation in the Fe-free melts is a function of the degree of melt polymerisation and it increases with increasing NBO/T (from 0.03 to 1.91 wt.% S). The S 2 − species are more soluble than the S 6 + species in contrast to previously studied Febearing “natural” compositions. The change from S 2 − to S6 + is observed at log fO2 = QFM − 1 to QFM + 1 which is ~ 1.5 log unit lower than the range of fO2 previously reported for Fe-rich compositions indicating that Fe influences not only the speciation but also the oxidation state of S in silicate melts at given redox conditions. The natural implications are that S 6 + in Fe-poor magmas can be stable at lower fO2 than previously predicted and, hence, S 6 + may act as an oxidising agent in the mantle wedge by successively oxidising Fe2 + to Fe 3 + via the reaction H2SO4 + 9FeO = FeS + 4Fe2O3 + H2O. For the silicate melt generated in the mantle wedge and containing about 10 wt.% total FeO, the change in the Fe 3 +/ΣFe ratio from 0.1 to 0.2 will correspond to an increase in the log fO2 from QFM − 0.5 to QFM + 1.5 and will require only 1000–3000 ppm S extracted from subducted slab. © 2012 Elsevier B.V. All rights reserved.

1. Introduction Sulphur (S) is the eighth most abundant element in the Earth (Henderson and Henderson, 2009) and plays an important role in processes occurring in all regions of the Earth from core to atmosphere. The storage and transport of sulphur in the Earth are

⁎ Corresponding author at: Institut für Goethe Universität Frankfurt, 60438 Frankfurt am Main, Germany. Tel.: + 49 6979840120; fax: + 49 6979840121. E-mail address: [email protected] (K. Klimm). 0009-2541/$ – see front matter © 2012 Elsevier B.V. All rights reserved. doi:10.1016/j.chemgeo.2012.04.028

strongly influenced by a wide range of oxidation states available, from −2 in the core and in the mantle sulphides to +4 in SO2 emissions from volcanoes and +6 in sulphates in the crust and atmospheric particulates. Sulphur is particularly important as a tracer of volcanic processes and sulphur emissions from volcanoes are extensively monitored as part of efforts to predict impending eruptions (e.g., Aiuppa et al., 2004). Volcanic SO2 is a local pollutant that causes acid rains, and sulphate aerosols formed from volcanic sulphur can reach the stratosphere and spread around the planet. The latter effect is implicated in episodic global cooling (Robock, 2000). The main sources of S in volcanic eruptions are magmatic

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and to understand in detail the transport, emission, and the entire budget of sulphur in magmatic systems. It is essential to quantify the solubility of sulphur in silicate melts, which is known to depend strongly on oxidation state and silicate melt composition (e.g., Jugo, 2009; Wallace and Edmonds, 2011; Webster and Botcharnikov, 2011). This dependence is controlled by structural characteristics of silicate melt and by speciation and redox state of S in the melts (e.g., Wilke et al., 2011). Solubility data, combined with information on the sulphur incorporation in the melt structure, are crucial for the development of a general model for the dissolution mechanism of S in silicate melts and for the interpretation of natural magmatic processes involving S. 1.1. Sulphur and iron redox state in magmas As an introduction to the oxidation behaviour of S in silicate melts, it is useful to review also the controls on oxidation states of Fe in silicate melts, because both S and Fe are very important elements involved in the redox reactions in magmatic systems. Iron and sulphur can be present in melts as Fe 2 + or Fe 3 + and S 2 − or S 6 +, respectively. The Fe 3 +/ΣFe (Fe 3 +/[Fe 3 ++Fe 2 +]) and S 6 +/ΣS (S 6 +/ [S 6 ++S 2 −]) ratios in silicate melts depend on the prevailing oxygen fugacity (fO2) and are determined by the following equilibrium reactions: 2FeO þ 1=2O2 ¼ Fe2 O3 2

S þ 2O2 ¼ SO4

2

ð1Þ ð2Þ

In the presence of water (H2O), the fO2 is affected by the dissociation of water, which is described by the following reaction: H2 O ¼ H2 þ 1=2O2

Fig. 1. Sulphur oxidation state (black lines) and iron oxidation state (grey lines) in silicate glasses expressed as the fraction of the oxidised cation as a function of fO2. The Fe3 +/ΣFe for basalt and andesite was calculated for T = 1200 °C, CAFS (stands for Ca–Al–Fe–Si glass) for T = 1000 °C using the model by Kress and Carmichael (1991). Selected S6 +/ΣS data are for EPMA (dashed lines) and XANES (solid line) derived data on natural glass compositions. Dotted curve is for sodium-silicate glass.

ð3Þ

The relationship between fO2 and Fe 3 +/ΣFe has been calibrated experimentally for various melt compositions (e.g., Kilinc et al., 1983; Kress and Carmichael, 1991) or spinel–pyroxene-(olivine) pairs (Wood and Virgo, 1989; Ballhaus et al., 1991) and is widely used as an oxygen geobarometer to determine the fO2 for natural rocks and melts. Basalt and andesite, typical primary melts in subduction zone volcanism, have Fe 3 +/ΣFe in the range of 0.1 to 0.2 at fO2 and temperatures representative for the mantle wedge with slightly higher values for basalt compared to andesite (Kress and Carmichael, 1991). Therefore, besides fO2, the melt composition has an influence on the oxidation state of Fe in the melt (Fig. 1). In the range of relevant fO2 for natural magmatic systems both, Fe 2 + and Fe 3 + are always present in significant proportions. The effect of water activity (aH2O) on Fe 3 +/Fe 2 + is indistinguishable from the effect of fO2 in H2-buffered basaltic systems (Botcharnikov et al., 2005) because aH2O and fO2 are both linked by the prevailing hydrogen fugacity (fH2) in the system. However, it was found that dissolved water may have an oxidizing effect on Fe in silicic and alkali-rich systems (e.g., Gaillard et al., 2001; Wilke et al., 2002; Schuessler et al., 2008). The change of S 6 +/ΣS with fO2 and the coexistence of both, S 2 − and S 6 +, in geologically relevant compositions has only been determined in three experimental studies so far using natural basaltic to dacitic compositions (Carroll and Rutherford, 1988; Jugo et al., 2010; Botcharnikov et al., 2011). In contrast to iron, sulphur changes its oxidation state in a narrow fO2 interval of ~ 2 log units (Fig. 1). The exact position of this change in terms of fO2 and the slope of the curve are currently debated as the precision of the determined S 6 +/ΣS is greatly affected by the spectroscopic technique applied (Jugo et al., 2005a; Jugo et al., 2010; Wilke et al., 2011; Klimm et al., in press). Recent improvements in XANES spectroscopy offer a higher precision on S 6 +/ΣS in silicate glasses whereas measurements

applying the routinely used λ SKα wavelength shift method using EPMA may suffer from beam damage (Rowe et al., 2007; Wilke et al., 2008; Klimm et al., in press) and the change from S2 − to S6 + generally shows a wider slope, shifted towards higher fO2 when using EPMA (Wallace and Carmichael, 1994; Matthews et al., 1999; Jugo et al., 2005a). In contrast to the case for Fe 3 +/ΣFe, a systematic relationship of S6 +/ΣS and melt composition has not been reported, yet. Although an investigation of S2 −/S6 + ratios in sodium silicate glasses indicates that in such “simple” systems the change of S2 − to S6 + may occur at fO2 that are ~2 log units below that of natural glasses (Nagashima and Katsura, 1973), it has been suggested that the results are erroneous because of the uncertainty in the fO2 (or partial pressure of oxygen, PO2) determination in their experiments (Jugo et al., 2005a). Another complication that arises from using the λ SKα wavelength shift method to determine S 6 +/ΣS in silicate glasses is that this method does not provide direct constraints on the S-species dissolved in the melt because it is based on the assumption that the shift of Δλ (SKα) is a linear function of the relative proportions of S 6 + and S 2 − compared to sulphate and sulphide reference compounds (Carroll and Rutherford, 1988; Klimm et al., in press). However, in order to quantify the solubility of sulphur in silicate melts and to formulate a structural model to describe specifically the dissolution mechanism of sulphur in a silicate melt and the relationship to fO2, the redox state, speciation and proportions of dissolved S-species must be known. Such information is essential to constrain the net fluxes of sulphur in subduction zones where the change from S 2 − to S 6 + is expected to occur at fO2 representative for the mantle wedge (Jugo et al., 2010) and where volcanic sulphur degassing may influence the climate on the Earth's surface (Robock, 2000).

1.2. Sulphur speciation and solubility in melts Several attempts have been made to describe the sulphur solubility mechanism in silicate melts. The solubility of S from the S-bearing gas mixtures in binary melt composition was first studied by Fincham and Richardson (1954). In their experiments, decreasing

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S solubility with increasing silica content and increasing S solubility with increasing mole fraction of metal oxides lead to the conclusion that that the oxygens participating in the S2 − and S6 + equilibria were non-bridging oxygens. Burnham (1979) described the S solubility mechanism analogously to his models for H2O and CO2 solubilities as H2S and SO2 incorporation to account for S 2 − and S 6 +, respectively following the reactions:

SO2 ðvÞ þ O

 ðmÞ

2 ðmÞ

¼ SH

2 ðmÞ

¼ SO3

H2 SðvÞ þ O

 ðmÞ

þ OH

2 ðmÞ

ð4Þ ð5Þ

where (v) and (m) refer to the vapour and melt phases. However, this model was considered to be unlikely because: a) this mechanism requires that an SH− ion breaks into the melt structure substituting a bridging oxygen to form Si\S bonds which are highly energetic and especially unlikely in the presence of more chalcophile elements such as Fe2 + or Ni (Nockolds, 1966) and b) in the model the metal cation does not affect the solubility mechanism of S which is contrary to the findings of Fincham and Richardson (1954). Because of the strong tendency of Fe–S ion pairing, Holloway (1980) suggested a more complex mechanism for S 2 − dissolution: H2 SðvÞ þ FeOðmÞ ¼ FeSðmÞ þ 2H2 OðvÞ

ð6Þ

This brief summary already illustrates the importance of identifying the S-species, S-bonding and the local environment in silicate melts. The total amount of S dissolved in the melt also depends on the sulphur fugacity (fS2) and is proportional to (fS2/fO2) ½ (O'Neill and Mavrogenes, 2002). In addition, the role of H2O on the dissolution of sulphur in silicate melts is poorly constrained (e.g., Moretti and Ottonello, 2003; Liu et al., 2007; Moune et al., 2009; Webster and Botcharnikov, 2011). Although the change from S 2 − to S 6 + is primarily a function of fO2 which is linked to H2O through Eq. (3) (see above), the amount of S dissolved will be governed by the Sspecies present in the melt. Recent experimental and spectroscopic investigations on S in silicate glasses emphasised that in the presence of H2O, sulphide may be dissolved as SH −, H2S or FeS-species in silicate melts (Clemente et al., 2004; Klimm and Botcharnikov, 2010; Stelling et al., 2011) or even as S3− in water rich fluids (Pokrovski and Dubrovinsky, 2011) whereas in dry systems additional species such as S2−, S3− or S4− (McKeown et al., 2001; Tsujimura et al., 2004) have been suggested. Here we present Part II of a combined experimental and multispectroscopic approach investigating the speciation and solubility of S as a function of fO2 in “simple” water-bearing silicate melt compositions. In Part I of this paper (Klimm et al., in press) we performed a detailed study of the methods for determining the oxidation state of S in glasses (λ SKα wavelength shift, X-ray absorption, Raman and 33S NMR spectroscopy). Here in Part II we apply XANES and Raman spectroscopy (which were found in Part I of the paper to be the most effective) to study the systematics of S6 +/ΣS as a function of fO2, Fe concentration and melt polymerisation. 2. Experimental methods 2.1. Starting materials and glass synthesis The experimental details are summarised in Table 1. As starting materials we used soda-lime glass (SLG), K2Si4O9 (KSG), NaAlSi3O8 (Ab) and trondhjemite glass (TROND). SLG consists of 70 wt.% SiO2, 7.5 wt.% CaO and 22.5 wt.% Na2O. TROND is identical in composition (except that it is Fe-free) to the trondhjemitic residual liquid coexisting with eclogite residue under subduction-zone relevant P-T-fO2 conditions of 2.5 GPa, 850 °C and QFM + 0.7 (Klimm et al., 2008a). It

contains 72.2 wt.% SiO2, 0.3 wt.% TiO2, 16.0 wt.% Al2O3, 0.9 wt.% CaO, 9.4 wt.% Na2O and 1.2 wt.% K2O. The glasses had been synthesised by mixing synthetic oxides (SiO2, TiO2, Al2O3) and carbonates (CaCO3, Na2CO3, K2CO3). The mixtures were decarbonated by stepwise heating from 700 to 1000 °C with a heating rate of 100 °C/h in a muffle furnace. The now carbonate-free mixtures were melted twice in Pt-crucibles at 1500 °C for 2 h in a tube furnace (with grinding between the two melting steps) and quenched to crystal free glasses. Homogeneity of the glasses was confirmed by EPMA analysis. Mixtures of glass, elemental S and Fe2O3 (in the case of Fe-doped runs only) were then homogenised in an agate mortar. These mixed starting materials were then welded together with H2O in Au capsules. Most of the experiments were performed in an internally heated pressure vessel (IHPV) equipped with a Shaw membrane and a rapid quench device at the Leibniz University of Hannover (Berndt et al., 2002). One run per composition was performed in cold seal pressure vessels (CSPV) also equipped with a rapid quench device (Carroll and Blank, 1997). Run conditions were 200 MPa and 1000 °C for IHPV and 200 MPa and 850 °C for CSPV. The fO2 during the experiment was either intrinsic (controlled by the pressure vessel material) or controlled by the use of an Ar–H2 gas pressure medium in the IHPV. The maximum fO2 during the experiments at intrinsic conditions was log fO2 = QFM + 1.5 in CSPV (Fabbrizio et al., 2006) and log fO2 = QFM + 4 in IHPV when only Ar was used as pressure medium (Berndt et al., 2002) with QFM being the fO2 for the quartz–fayalite–magnetite buffer assemblage (Chou, 1978). In all other experiments in the IHPV the fO2 was varied by adjusting the fH2 using an Ar–H2 mixture and monitored with a Shaw membrane (Berndt et al., 2002). Within the capsule the fO2 is fixed by the external fH2 through the equilibrium reaction of water formation (H2 + ½ O2 = H2O) due to diffusion of H2 through the capsule wall. Therefore, the fO2 in the capsule during the experiments at water saturated conditions can be calculated as fO2 = (fH2O/(Kw × PH2)) 2 with fH2O after Burnham et al. (1969), Kw from Robie et al. (1978) and the experimental PH2 measured with the Shaw membrane. For water-undersaturated conditions (aH2O b 1) the reduced aH2O in the system (and therefore fH2O) results in a decrease of fO2, because of the fixed fH2. The fO2 in the capsule at water-undersaturated conditions is calculated as log fO2 = log fO2(aH2O = 1) + 2log aH2O (see also Botcharnikov et al., 2005). After the high pressure and temperature runs each capsule was reweighed to ensure that no vapour loss had occurred during the experiment and then carefully opened while being viewed using a binocular microscope in order to determine the presence of excess sulphur phases, such as sulphate or sulphide blebs. Glass chips together with excess sulphur phases (if present) were embedded in araldite epoxy and oil-polished to avoid the dissolution of sulphur blebs in H2O. All runs were performed at water-undersaturated conditions (aH2Ob 1) to avoid the preferred partitioning of sulphur into a coexisting hydrous fluid phase (e.g., Lesne et al., 2011; Scaillet et al., 1998; Keppler, 1999; Webster and Botcharnikov, 2011). In order to determine the aH2O during the experiment, the concentration of H2O in the melt during the experiment (cH2Omelt) and the concentration of H2O at water saturation (cH2Osat) for each composition must be known. Assuming almost ideal behaviour of H2O at studied conditions, the aH2O can then be calculated as aH2O =cH2Omelt/cH2Osat (e.g., Botcharnikov et al., 2005). The cH2Osat for all compositions was determined by equilibrating 10 and 15 wt.% of H2O together with glass powder in Au capsules at identical pressure of 200 MPa and temperature of 1000 °C. Details of these experiments are summarised in Table 2. The runs were performed in a different IHPV without a Shaw-membrane resulting in slightly lower intrinsic fO2 (log fO2 ~ QFM+ 3.2; Schuessler et al., 2008). After the runs the capsules were pierced using a steel needle and then placed into a drying furnace at 110 °C for 3–5 min. The occurrence of an excess fluid after piercing and the weight loss of the capsule after drying confirmed the fluid-

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Table 1 Summary of experimental details of sulphur bearing runs at 1000 °C (and 850 °C) and 200 MPa. Run no.a

Duration [d]

PH2b [Bar]

Sloaded [wt.%]

Sglass [wt.%]

Fe2O3 [wt.%]

Fe/Sglass Molar

cH2Oinitial [wt.%]

fO2c ΔQFM H2Osat

S6 +/ΣSd XANES

S6 +/ΣSe Raman

cH2Omeltf [wt.%]

aH2Og

fO2h ΔQFM

SLG RS47 RS58 RS19 RS39 RS60 RS61 RS62 RS63 RS64 RS29 RS59 RS55 RS75

2 1 2 3 5 5 5 5 5 5 1 1 4

– – – 8.1 7.6 7.6 7.6 7.6 7.6 7.9 12.9 12.9 13.8

3.00 0.54 1.15 0.08 0.38 0.40 0.34 0.41 0.21 2.00 3.00 0.40 0.36

0.69 0.54 1.15 0.08 0.38 0.40 0.34 0.41 0.21 1.91 1.71 0.40 0.36

– – – – 0.66 1.19 2.24 0.43 – – – – –

– – – – 0.69 1.19 2.66 0.50 – – – – –

11.0 4.0 5.6 5.2 5.0 4.9 5.0 5.1 5.0 5.2 5.1 4.0 0.8

+ 4.00 + 4.00 + 1.50 + 0.72 + 0.77 + 0.77 + 0.77 + 0.77 + 0.77 + 0.74 + 0.31 + 0.31 + 0.25

1 1 0.98 0.95 – – – – 0.57 0.30 0.20 0.12 0

– 1 0.96 – – – – – 0.57 0.37 0.26 0.22 0

4.26 2.79 3.07 5.03 – – – – 4.73 3.85 3.75 3.89 0.77

0.39 0.26 0.28 0.46 0.46 0.45 0.46 0.47 0.44 0.36 0.35 0.36 0.07

+ 3.19 + 2.82 + 0.41 + 0.05 + 0.10 + 0.08 + 0.10 + 0.12 + 0.05 − 0.16 − 0.61 − 0.58 − 2.01

KSG RS44 RS41 RS18 RS36 RS65 RS66 RS67 RS68 RS69 RS53 RS26 RS76

2 2 2 3 5 5 5 5 5 1 5 2

– – – 7.8 7.5 7.5 7.5 7.5 7.5 7.4 7.9 13.8

3.00 0.37 1.80 0.07 0.21 0.46 0.45 0.43 0.45 0.44 2.00 0.41

0.35 0.37 0.66 0.07 0.21 0.46 0.45 0.43 0.45 0.44 0.96 0.41

– – – – – 0.44 0.76 1.18 2.48 – – –

– – – – – 0.38 0.68 1.10 2.20 – – –

11.2 5.6 5.6 5.1 5.0 5.0 5.0 5.0 4.9 4.0 5.2 0.7

+ 4.00 + 4.00 + 1.50 + 0.75 + 0.78 + 0.78 + 0.78 + 0.78 + 0.78 + 0.80 + 0.74 + 0.25

1 1 0.71 0.63 0.29 – – – – 0.23 0.26 0

1 – 0.73 – 0.32 – – – – 0.48 – 0

4.46 4.77 2.73 5.00 4.86 – – – – 3.77 4.03 0.70

0.34 0.36 0.21 0.38 0.37 0.38 0.38 0.38 0.37 0.28 0.30 0.05

+ 3.06 + 3.11 + 0.13 − 0.10 − 0.09 − 0.07 − 0.07 − 0.07 − 0.08 − 0.29 − 0.29 − 2.30

TROND RS12 RS31 RS24 RS50 RS78

1 2 5 1 2

– – 7.9 12.3 13.8

0.10 0.10 2.00 0.05 1.10

0.06 0.06 0.07 0.05 0.31

– – – – –

– – – – –

5.1 4.9 5.2 3.8 0.6

+ 4.00 + 4.00 + 0.74 + 0.35 + 0.25

1 1 0.55 0.28 0

1 – 0.35 0.40 –

4.88 4.68 2.73 3.77 0.64

0.82 0.78 0.46 0.63 0.11

+ 3.82 + 3.79 + 0.06 − 0.05 − 1.69

Ab RS30 RS42 RS23 RS10

2 2 5 1

– – 7.9 20.8

0.11 3.00 2.00 0.05

0.11 0.31 0.03 0.03

– – – –

– – – –

5.4 11.3 5.3 5.0

+ 4.00 + 4.00 + 0.74 − 0.11

1 1 0.56 0.31

1 – 0.54 0.53

5.15 4.56 2.79 4.97

0.96 0.85 0.52 0.93

+ 3.97 + 3.86 + 0.17 −0.17

a

Bold numbers indicate runs saturated with a sulphide/sulphate phase. Numbers in italic indicate Fe-doped runs. Measured with a Shaw membrane during the experiment (setup described in detail in Berndt et al., 2002. c ΔQFM H2Osat = log fO2(experiment) − log fO2 (QFM; Chou, 1978). For water saturated conditions (aH2O = 1) in the CSPV the fO2 is ~ QFM + 1.5 (Fabbrizio et al., 2006). For the water saturated conditions in the IHPV the fO2 is either intrinsic (QFM + 4) or calculated by using the equation fO2 = (fH2O/(Kw × PH2))2 with fH2O from (Burnham et al., 1969), Kw from (Robie et al., 1978) and the experimental PH2 run, assuming that the reaction of water formation controls the redox conditions in the experiments (H2 + 1/2 O2 = H2O; at fixed fH2). d Determined by a linear combination of reduced and oxidised reference XANES spectra (Klimm et al., in press). e Calculated by ratioing the intensities of Raman bands related to S6 + and S2 − (Klimm et al., in press). f Corrected water content in the melt assuming that H2O is consumed to generate SO42 − (S6 +) following the reaction: S + 4 H2O = SO42 − + 2 H+ + 3 H2. The amount of SO42 − formed by this reaction is deduced from the Sloaded and the S6 +/ΣS determined by XANES. g aH2O = cH2Omelt/cH2Osat, where cH2Osat is the experimentally determined water content at water saturation at 200 MPa and 1000 °C (see table 2). For the Fe-bearing runs aH2O was calculated as aH2O = cH2Oinitial/cH2Osat. h Because the aH2O in all runs is aH2O b 1 the fO2 in all charges is calculated as log fO2 = log fO2(aH2O = 1) + 2log aH2O (Botcharnikov et al., 2005). b

Table 2 Summary of experimental details of water saturation determination at 1000 °C and 200 MPa. Comp.

Duration [d]

cH2Oinitial [wt.%]

fO2a ΔQFM H2Osat

cH2Omeltb [wt.%]

SLG KSG TROND Ab

5 5 5 5

15.0 15.0 10.0 10.0

+ 3.20 + 3.20 + 3.20 + 3.20

10.82 ± 0.09 13.25 ± 0.07 5.97 ± 0 07 5.34 ± 0 09

a ΔQFM H2Osat = log fO2(experiment) − log fO2 (QFM; Chou, 1978). For the water saturated conditions in the IHPV the fO2 is intrinsic (QFM + 3.2). Note that the IHPV used here is a different one from those equipped with the Shaw membrane, which was used in all other experiments. b Determined by Karl-Fischer Titration (KFT; Behrens et al., 1996).

saturated conditions during the experiment. The water content of the glasses at water-saturated conditions was then determined by KarlFischer Titration (KFT; Behrens et al., 1996). For SLG and KSG we obtained surprisingly high water solubilities of 10.82 and 13.25 wt.% H2O at 200 MPa. These high levels are confirmed by the fact that first set of experiments doped with 10 wt.% H2O did not release any water upon piercing the capsules after experiments, indicating water-undersaturated conditions. In addition, in order to exclude any contamination of our glasses by water-bearing phases that may have precipitated from the fluid upon quenching and ultimately result in an overestimation of the water content of the glasses, additional runs with higher cH2Oinitial were performed. If hydrous quench phases had precipitated upon quench, then the cH2Omelt determined by KFT will be a function of the cH2Oinitial. However, we did not observe any such

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relationship, assuring that the levels of water saturation for KSG and SLG are correct. In the experiments at water-undersaturated conditions the cH2Omelt should be the amount of water loaded in the capsule. The deficit to 100% of the microprobe totals can be used to estimate the H2O content in the glass (Devine et al., 1995). However, because we used elemental S as the starting S species, the oxidation of elemental S to SO42 − consumes H2O to form SO42 −, as demonstrated for the oxidation of S 2 − (Stelling et al., 2011) and the reduction of SO42 − (Backnaes et al., 2007). Therefore, the cH2Omelt is not simply the amount of water loaded in the capsule and was corrected for water loss due to S oxidation assuming that the oxidation of S to SO42 − consumes 4 mol of H2O per mole of SO42 −. The amount of SO42 − in the glass can be deduced from the amount of S loaded in the capsule and the S6 +/ΣS ratio determined by XANES. Runs RS42 and RS47 provide evidence that this assumption of water loss due to S oxidation is correct; although both runs were loaded at water levels above that of water saturation (cH2Oinital = 11.3 and cH2Osat = 5.3 for RS42; cH2Oinital = 11.0 and cH2Osat = 10.8 wt.% for RS 47, respectively) no excess fluid phase was observed upon opening the capsules or by weight loss after drying. The deficit to 100% of the microprobe totals of the glasses (available as Supplementary data at http://www.sciencedirect.com/ science/journal/00092541) are in agreement with the calculated water contents considering that the accuracy of this method is only within 1–2 wt.% H2O when not calibrated with reference standards (e.g., Klimm et al., 2008b). Exceptions here are the KSG glasses that suffered from potassium loss during EPMA reflected in lower and almost constant totals of ~ 90%. 2.2. Analytical methods Compositions of glasses were determined by EPMA using a CAMECA SX-100 at the University of Bristol. To minimize the loss of alkalis during the measurements a low beam current of 2 nA at 20 kV with a defocused beam of 20 μm was used. Na and K were measured first with 5 s total counting time. However, even under these conditions we could not avoid K loss from the KSG glasses in the order of a few wt.%. S concentrations were measured with 20 kV, 30 nA and 30 s total counting time. The XANES spectra were collected at the Swiss Light Source (SLS), Villigen, Switzerland at the LUCIA beamline. Conditions and procedures for the S K-edge XANES measurements are presented in detail in Part I of this paper (Klimm et al., in press). To avoid any beam-damage during analysis, the samples were moved during spectra acquisition with a rate of ~1 s per step (Wilke et al., 2008). In addition, spectra were collected at the Fe K-edge in the range of 7050 to 7300 eV to determine the valence state of Fe in Fedoped glasses. Close to the Fe K-edge (7105 to 7120 eV) small step sizes of 0.1 eV and long acquisition times of 4 s were used. In other regions the step size was 0.2 eV (7120 to 7145 eV) or 1 eV (7050 to 7105 eV and 7145 to 7300 eV), with 1 s acquisition time per step. Raman spectra were collected with a Renishaw micro-Raman spectrometer (RM-1000) using the 532 nm line of a Nd: YAG laser at the Goethe University of Frankfurt. Details on the Raman setup and analytical conditions are also reported in Part 1 of this paper (Klimm et al., in press). 3. Results 3.1. Run products All experimentally produced samples consisted of homogenous crystal-free glass with the only exception of RS78 that contained some additional plagioclase crystals (~ 5%). The compositions of all glasses are provided as Supplementary data (available at http:// www.sciencedirect.com/science/journal/00092541) S-saturated runs also contained sulphide or sulphate phases. Typical BSE-images of S

saturated runs are shown in Fig. 2. The composition of the “excess” sulphide and sulphate phases depends on the starting bulk composition. For instance KSG contained potassium-bearing phases such as K2S and/or K2SO4 (Fig. 2a and c), whereas TROND contained mixtures of Na and Ca sulphides and/or sulphates (Fig. 2b). The sulphides and/or sulphates were not evenly distributed within the glasses but occurred concentrated in blebs with sizes ranging from 10 to 200 μm in diameter (Fig. 2a, b, d) or at the edge of the glasses in contact to the Au-capsule (Fig. 2c). In the latter case the sulphides (when present) also contained dissolved Au from the capsule wall. The shape and size of the sulphide and sulphate blebs demonstrate that sulphide/sulphate liquid was present at high T in agreement with the observations of similar features in experiments with S-saturated basalts (Jugo et al., 2005b) and are not a product of insufficiently rapid quenching. The importance of avoiding any contact of the sample with H2O during polishing of the samples can be seen Fig. 2c which shows the sample with empty blebs not filled with any S-phase. This sample had been polished in the presence of H2O resulting in the dissolution of the S-phases from the blebs in water. Another piece of the glass from the same charge but polished with oil in the absence of H2O still contained blebs filled with Na-sulphide/sulphate. The glass compositions of sulphur-undersaturated samples are identical to the starting bulk composition, but in samples containing excess sulphide/sulphate phases, the silicate glasses are slightly depleted in alkalis consumed by the formation of the excess sulphide or sulphate phases. The S concentrations of the glasses are given in Table 1 and the S contents of the glasses at sulphur saturation are presented in Fig. 3. The highest sulphur concentrations are observed in the SLG glass with 1.71 to 1.91 ± 0.03 wt.% S and the KSG glass with 0.66 to 0.96 ± 0.02 wt.% S at log fO2 = QFM + 0.5 to QFM − 1. At similar fO2 range the maximum S concentration of the two other compositions is 0.07 and 0.03 ± 0.01 wt.% S for TROND and Albite, respectively, i.e. 1 to 2 log units lower than SLG and KSG. All glass compositions except Albite show a decrease of S solubility with increasing fO2. In the range log fO2 = QFM + 3 to QFM+ 4, S concentrations in the glass in the presence of a sulphate phase are 0.69 ± 0.02 wt.% S for SLG, 0.35 ± 0.01 wt.% S for KSG and 0.06 ± 0.01 wt.% S for TROND. The S content in the Albite glass at log fO2 = QFM + 3.8 is 0.31± 0.01 wt.% S and therefore much higher to that at reducing conditions. However, for the other compositions the general observation that the S solubility decreases with increasing fO2 is further confirmed by one experiment with TROND (RS78) at reducing fO2 that was also saturated with a sulphide phase. The S concentration in that glass was 0.31 wt.% S at log fO2 = QFM − 1.66. 3.2. XANES and Raman spectroscopy on hydrous glasses 3.2.1. S K-edge XANES of Fe-free hydrous glasses XANES spectra for all compositions synthesised at varying fO2 are presented in Figs. 4 to 7. Because the edge position is a function of the sulphur oxidation state (e.g., Fleet, 2005) peaks in XANES spectra can be assigned to different S-species present in the run products when compared to spectra of sulphur reference compounds. More details on the interpretation and assignment of features in S K-edge XANES spectra are given in Part I (Klimm et al., in press). The most oxidised glasses (fO2 = QFM ≥ 2.83) only contain S 6 + as indicated by the presence of one characteristic peak at 2482 eV. In contrast, the most reduced glasses (fO2 ≤ QFM − 1.66) show two characteristic peaks at 2471.8 and 2476 eV related to S 2 − dissolved in the silicate glass. All other glasses synthesised at intermediate fO2 show peaks related to both, S 6 + and S 2 −, with a systematic increase of the peak at 2482 eV and decrease of the peaks at 2471.8 and 2476 eV and thus, an increase of the relative proportion of S 6 +/S 2 − with increasing fO2. An additional feature of the spectra at reducing fO2 is a characteristic peak present at 2466 eV in glasses containing S 2 −. This peak has been observed previously in a few studies (Wilke et

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Fig. 2. BSE images of sulphur saturated glasses: a) KSG glass containing blebs filled with K2SO4; b) TROND glass containing blebs filled with a mixture of CaSO4 and Na2SO4; c) KSG glass and a mixture of K2SO4, K2S and Au at the edge of the glass; d) albite glass containing empty blebs from where sulphide/sulphate phases have been dissolved during polishing. Run numbers, fO2 and levels of initial S contents are provided in the black boxes.

al., 2008; Métrich et al., 2009; Stelling et al., 2011) but has not yet been definitively assigned to any specific S-species dissolved in the glass. It is most prominent at the most reducing conditions when S 6 + is absent and it roughly decreases with increasing fO2 (exceptions: RS64 and RS23 see explanation in Section 4.4). Thus, this feature is related to reduced S-species in the glass. The absence of any significant signal at 2478 eV, i.e., at the energy characteristic for the presence of S4 +, indicates that the glass was not damaged during the XANES measurement (Klimm et al., in press; Wilke et al., 2008). However, some of the KSG glasses (Fig. 5) show a very small signal at 2478 eV indicating that even when the sample was moved under the beam, some of the KSG glasses experienced a small degree of beam damage. This is in agreement with the observation, that KSG is the most vulnerable composition in terms of being affected by interaction with X-ray beam exposure (Klimm et al., in press; Wilke et al., 2008). The amount of S4 + generated during beam exposure is very small (as indicated by the small signal at 2478 eV) and does not materially affect our results. The spectra of fully oxidised glasses of each bulk composition are almost identical (top spectra in Figs. 4–7). All spectra show a strong peak at 2482 eV (related to S 6 +) and a relatively small shoulder after the edge at ~ 2500 eV. In contrast, spectra obtained from fully reduced glasses (bottom spectra in Figs. 4–6) show slight differences in the intensities of the peaks at 2466, 2471.8 and 2476 eV depending on the bulk composition. For instance, the features at 2466 and 2476 eV are most prominent in SLG, whereas the peak at 2471.8 eV is most intense in KSG. The most intense feature in TROND is in between 2471.8 and 2476 eV.

3.2.2. S K-edge XANES of Fe-bearing hydrous glasses The XANES spectra of SLG and KSG glasses doped with Fe and synthesized under similar P-T-fO2 conditions are shown in Fig. 8. All

spectra are characterized by a strong peak at 2482 eV indicating the presence of S 6 + in the glasses. In addition, peaks at 2466, 2471.8 and 2476 eV indicate the presence of reduced S-species such as S 2 − in the glass in agreement with the spectra observed in the Fe-free system. There are very clear systematic changes in the peak intensities depending on the amount of Fe added to the glass: i) adding Fe to the glass results in the appearance of an additional feature at 2469 eV that was not observed in the Fe-free glasses; ii) the intensity of this peak systematically increases with the amount of Fe added; iii) the peaks at 2466 and 2471.8 eV systematically decrease with increasing Fe/S in both compositions. The feature at 2466 eV is absent for Fe/S > 2 whereas a small peak is still visible at Fe/S ~ 1; iv) the drop of the intensity at 2471.8 eV from Fe-free to Fe-bearing conditions is more pronounced in KSG than in SLG; and v) there is no systematic relationship between the normalized peak intensity at 2482 eV and Fe/S, therefore S 6 +/ΣS is constant for the Fe/S ratios examined here (Fig. 9). The only exception is the KSG glass with Fe/S = 0.38 which, for some unknown reason, shows a significantly increased intensity at 2482 eV compared to all other KSG glasses.

3.2.3. Fe K-edge XANES Fe K-edge XANES spectra for the pre-edge region at 7110 to 7120 eV of the Fe-doped SLG and KSG glasses are shown in Fig. 10 together with olivine and hematite as Fe 2 + and Fe 3 + reference compounds. The energy position of the pre-edge feature between 7104 and 7118 eV is a function of the Fe oxidation state (Berry et al., 2003; Wilke et al., 2004; Botcharnikov et al., 2005). Compounds containing Fe 2 + such as olivine are characterized by a pre-edge feature at ~ 7113 eV whereas Fe 3 + as in hematite shows a pre-edge feature at higher energies of ~ 7116 eV. The energy position of the pre-edge feature of the SLG and KSG glasses containing Fe/S from 0 to 2.66 (Table 1) is all almost identical at ~ 7114.5 eV.

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Fig. 4. Normalised XANES spectra at the S K-edge for SLG glasses. Grey lines indicate the characteristic energies for the different S-oxidation states: S6 + at 2482 eV; S2 − at 2476 and 2471.8 eV; and the SH− species at 2466 eV.

Fig. 3. Sulphur concentrations in the melt as a function of: (a) fO2 and (b) NBO/T calculated on a hydrous basis. Grey lines in a) represent S concentrations for basalt and trachyandesite following the S-solubility model by Jugo (2009). Solid line in a) indicate the level of S contents for pure S2 − and S6 +, respectively, and the dotted extensions reflect the range of fO2 where S2 − and S6 + coexist; Filled circles in b) indicate runs that contain a mixture of S6 + and S2 − and open circles runs that only contain S6 +. Solid and dashed curves in b) are linear fits for reducing and oxidising conditions, respectively.

The calculated Fe 3 +/ΣFe of these silicate glasses by fitting two Gaussian functions for the pre-edge peak following the method by Wilke et al. (2004) is Fe 3 +/ΣFe ~ 0.5 without any significant variations amongst the two compositions and the level of Fe-doping.

3.2.4. Raman spectroscopy of Fe-bearing hydrous glasses Raman spectra and their band assignment of the S-bearing glasses from Table 1 have already been shown and discussed in more detail in Klimm and Botcharnikov (2010) and in Part I of this contribution (Klimm et al., in press). Here we will focus on the results of Raman spectroscopy on Fe-bearing glasses having various Fe/S ratios. Raman spectra of SLG and KSG runs are shown in Fig. 11. In addition to the bands related to vibrations of the silicate glass structure, all spectra show two peaks, one at 990 cm− 1 and the other at

2574 cm− 1. These bands are related to SO42 − and S–H vibrations, respectively, in the glass structure and they are consistent with the observation of both sulphate and sulphide in these glasses by XANES. Compared to the Fe-free glasses, all Fe-bearing glasses show three additional bands at 298, 372 and 420 cm− 1. All three bands are attributed to the presence of Fe–S complexes in the glass, but the details of the assignments are unclear (see Section 4.4). With increasing Fe/S, the band at 2574 cm− 1, related to S–H, decreases and the bands at 298, 372 and 420 cm− 1, related to Fe–S, systematically increase. The intensities of the bands are related to the abundance of the species in the glass (Klimm et al., in press). Thus, the relative quantities of S–H (as SH− and H2S) and Fe\S bonds present in the glass change from only S–H present in the case of Fe-free glasses to Fe–S dominant at Fe/S > 2. It has to be noticed that even at Fe/S= 2.66 a small peak at 2574 cm− 1 is still visible in SLG, indicating that some S\H bonds are still present in the glass. The band at 990 cm− 1, which is related to S6 +, does not show any significant or systematic variation with Fe/S implying that the amount of S 6 + is constant in all the glasses, and that for these levels of Fe addition there is no effect on the position of the sulphur oxidation state curve (Fig. 1).

4. Discussion Complete understanding of the complexities of the dissolution mechanisms of S in a silicate melt requires knowledge of the nature of the dissolved sulphide and sulphate species as well as a quantitative description of the effect of fO2 and melt compositions on the sulphide–sulphate equilibrium. In the following sections we will

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Fig. 5. Normalised XANES spectra at the S K-edge for KSG glasses. Grey lines indicate the characteristic energies for the different S-oxidation states: S6 + at 2482 eV; S2 − at 2476 and 2471.8 eV; and the SH− species at 2466 eV.

Fig. 6. Normalised XANES spectra at the S K-edge for TROND glasses. Grey lines indicate the characteristic energies for the different S-oxidation states: S6 + at 2482 eV; S2 − at 2476 and 2471.8 eV; and the SH− species at 2466 eV.

combine and discuss the information deduced from the different spectroscopic techniques in terms of the S-speciation in the glass.

the similarity in wavenumber compared to sulphate reference compounds suggests that S 6 + is present as SO42 − also confirming the observation from XANES and NMR. It thus can be inferred that the S in the sulphate oxy-anion has a structural environment in these glasses that is very similar to that of crystalline reference compounds as suggested by Wilke et al. (2011). This structural similarity is further constrained by structural data obtained from analysis of the first EXAFS data published for S 6 + in borosilicate waste glass (Brendebach et al., 2009), that show that the sulphate oxy-anion is present in the glass as isolated groups surrounded by cations and not associated with bridging oxygens. The variation in the exact wavenumber related to S6 + in silicate glasses of different compositions is ~10 cm− 1 which is much less compared to the variation in sulphate reference compounds where, depending on the composition and structure of the compound, especially on the type of the charge balancing cation, the peak of ν1(SO4) stretching vibration is observed in the range of 975 to 1025 cm− 1 (Beny et al., 1982; Burke, 2001; McKeown et al., 2001; Tsujimura et al., 2004; Lenoir et al., 2009; White, 2009). This implies, that the effect of cations on the vibrational mode of the S\O bond is much less pronounced in a silicate glass than in a crystalline compound.

4.1. The local environment of sulphate Fully oxidised glasses that only contain S 6 + are characterised in XANES spectra by a very intense peak at around 2482 eV, and a small shoulder on the high-energy side is observed throughout (Figs. 4 to 7). The XANES spectra of fully oxidised “synthetic” Fe-free glasses (Paris et al., 2001; Fleet et al., 2005; Backnaes et al., 2007; Métrich et al., 2009; Stelling et al., 2011) are identical to fully oxidised spectra of “natural” Fe-bearing glasses (Paris et al., 2001; Métrich et al., 2002, 2009; Jugo et al., 2010). Overall, these spectra are very similar to those of the crystalline sulphate model compounds (e.g., Fleet, 2005; Klimm et al., in press; Métrich et al., 2009). Even the region at higher energies (2490 to 2510 eV) shows a quite strong similarity to sulphate model compounds and almost no variability depending on glass composition. The same can be noticed from 33S MAS NMR. Spectra obtained from S6 +-bearing glasses and sulphate reference compounds are similar in the chemical shift and the peak shapes (Couch et al., 2004; Klimm et al., in press). In the Raman spectra of silicate glasses S6 + can be identified by a peak at around 1000 cm− 1 corresponding to the ν1(SO4) stretching vibration. The exact wavenumber is depending on the glass composition and the S6 + band position varies from 990 cm− 1 in “synthetic” silicate and borosilicate glasses (McKeown et al., 2001; Tsujimura et al., 2004; Lenoir et al., 2009; Klimm and Botcharnikov, 2010) to 1000 cm− 1 in natural basaltic to rhyolitic glasses (Klimm and Botcharnikov, 2010). The occurrence of this band is a direct evidence of S\O bonding in the silicate glass structure and

4.2. The local environment of sulphide The local structural environment of sulphide is much more complex and variable than that of sulphate. The key features to interpret our data are: i) the asymmetric peak at 2574 cm − 1 in Raman spectra; ii) the set of three peaks at 298, 372 and 420 cm − 1 in Raman spectra; iii) the set of three small peaks at 2466, 2469 and 2471.8 eV in the

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Fig. 7. Normalised XANES spectra at the S K-edge for Albite glasses. Grey lines indicate the characteristic energies for the different S-oxidation states: S6 + at 2482 eV; S2 − at 2476 and 2471.8 eV; and the SH− species at 2466 eV.

XANES spectra; and iv) the main, broad sulphide peak in the XANES spectra at 2476 eV.

4.2.1. Evidence from the Fe-free samples There is little doubt that the Raman peak at 2574 cm − 1 is the result of the S–H stretching vibration and hence, it should originate from some combination of SH − and H2S species (Klimm and Botcharnikov, 2010; Klimm et al., in press). We assign the XANES peak at 2471.8 eV to dissolved H2S because it is close to the experimentally observed position for H2S (Prange et al., 2002). This assignment is supported by a correlation between the intensities of the 2574 cm − 1 Raman peak and the 2471.8 eV XANES peak (Fig. 12). Furthermore there is an equally good correlation between the 2574 cm− 1 Raman peak and the 2466 eV XANES peak. This implies that the 2466 eV XANES peak results from an S–H species that is not H2S. It was noted earlier that the 2466 eV XANES peak occurs in all the four glass compositions studied here and therefore cannot be the result of species such as NaSH, CaSH, AlSH etc. The only common component of all the glasses is Si, therefore we assign the 2466 eV XANES peak to Si–S–H. Although the 2466 eV XANES peak seems to be the most well developed in hydrous, Fe-free glasses (Klimm et al., in press; Stelling et al., 2011), small 2466 eV peaks were reported by Métrich et al. (2009) in Fe-free glass synthesised at 1 bar pressure (and thus presumed to be dry). It is possible that these glasses somehow retained traces of H even at 1 bar pressure, or alternatively the 2466 eV peak could also be produced by Si–S− as well as Si–S–H. Theoretical studies are probably required to resolve this issue. The coexistence of Si–S–H and H2S proposed here is directly analogous to the coexistence of associated and dissociated species of

Fig. 8. Normalised XANES spectra at the S K-edge for SLG and KSG glasses with different Fe/S ratios. Grey lines indicate the characteristic energies for the different S-oxidation states: S6 + at 2482 eV; S2 − at 2476 and 2471.8 eV; and the SH− species at 2466 eV. The inset shows an enlargement of the energy regions from 2464 to 2747 eV that are most significantly affected by the amount of Fe added to the glass.

Fig. 9. Normalised XANES intensities at 2482 eV vs. the molar Fe/S in the glasses. Note the (almost) constant intensity at 2482 eV for various Fe/S.

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Fig. 10. Normalised XANES spectra at the Fe K-edge for KSG and SLG glasses with Fe/S ranging from Fe/S = 0 to 2.66. Also shown are the reference spectra of olivine and hematite that only contain Fe2 + and Fe3 +, respectively.

water (e.g., Kohn, 2000) and CO2 (e.g., Morizet et al., 2001) as part of the dissolution mechanism of volatile species in silicate melts. Furthermore, it can be seen in Fig. 4 (which shows XANES spectra normalised to the edge step, i.e. total sulphur concentration) that samples RS29 and RS59 have anomalously low Si–S–H peaks compared with the other samples in the SLG series (same was observed for RS26 in the KSG series). Since these are the two samples with the highest dissolved sulphur (and sulphide) concentrations, it appears that high sulphide concentrations reduce the concentration of SH − relative to H2S. This phenomenon is explored in Fig. 13 which shows the absolute height of the SH − and H2S XANES peaks for SLG glasses (in arbitrary units) as a function of the total sulphide concentration of the sample. The peak heights (peak intensities) were obtained by fitting the un-normalised peaks at 2466 and 2471.8 eV (plus 2469 eV in the case of Fe-bearing samples; see below) with Gaussian functions after subtraction of a polynomial background to remove the edge step and the overlap of the 2471.8 eV peak with the feature at 2476 eV. The trends in Fig. 13 are very reminiscent of those for OH − and H2O in hydrous silicate glasses especially in the case of the SLG (Fig. 13a). The concentration of SH − rises steeply for low concentrations of sulphide, then the slope shallows. For H2S the opposite is true. By analogy with the H2O/OH − system (e.g., Nowak and Behrens, 1995; Shen and Keppler, 1995; Behrens and Nowak, 2003; Behrens et al., 2009) one can imagine a temperature dependent equilibrium 2−

O

−

−

þ H2 S ¼ SH þ OH

ð7Þ

which could possibly be affected by retrograde reaction upon quenching. 4.2.2. Evidence from the Fe-bearing samples The other informative changes in the spectra that can be used to elucidate the sulphide dissolution mechanisms are those that accompany modification of the glass compositions by addition of Fe. As noted previously, the Raman 2574 cm − 1 peak (SH − + H2S) is reduced by addition of Fe. Fig. 14a shows the relationship between the 2574 + 2601 cm− 1 peaks and the sum of the 2466 and 2471.8 eV

Fig. 11. Normalised and baseline subtracted Raman spectra of Fe doped SLG (RS60 to RS63) and KSG (RS66 to RS69) glasses and Fe-free reference glasses (RS64 and RS65) synthesised at identical fO2. Grey lines indicate S\H bonds at 2574 cm− 1, SO42 − at 990 cm− 1 and the most prominent feature related to Fe–S at 372 cm− 1.

XANES peaks. The near-linear correlation implies that either: i) all the sulphide in the Fe-free glasses is in the form of SH− or H2S; or ii) that the ratio (SH− + H2S + “FeS”)/total sulphide is constant despite the changing Fe/S ratio in the samples. In this context it is worth noticing that the feature at 2476 eV, which is also related to S2 −, is present in all spectra (Figs. 4 to 8), but has not been assigned to any S-species, yet. It has also been identified in other glass compositions including Fe-bearing natural glasses (Métrich et al., 2002, 2009; Jugo et al., 2010; Wilke et al., 2011). The similarity in energy of this feature with XANES spectra obtained for CaS and MgS reference compounds leads to the conclusion that CaS and MgS species may also be present in silicate glasses (Métrich et al., 2009; Wilke et al., 2011). However, when comparing all spectra available, the bulk chemistry does not affect the energy position of this peak. Also, addition of Fe does not systematically affect the intensity at 2476 eV (Fig. 8). This suggests that this feature may reflect the overall presence of S 2 − but does not provide a direct information on the S-complexing. Therefore, in the absence of any other evidence, we will assume the first scenario, i.e. that all sulphide in the Fe-free glasses is in the form of SH− or H2S. The absence of any other sulphide species is also supported by the

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Fig. 12. Normalised XANES intensities of Fe-free SLG glasses at 2466 (filled diamonds) and 2471.8 eV (open diamonds) vs. the sum of normalised Raman intensities at 2574+ 2601 cm− 1. Dashed and solid lines are linear fits.

agreement of S6 +/ΣS ratios determined by XANES and Raman spectroscopy (Klimm et al., in press). As Fe is added, the concentrations of SH− and H2S are reduced and the concentration of S, complexed by Fe (“FeS”), increases. This model is supported by the change in the intensity of the 2469 eV XANES peak as function of the intensity of the “FeS”-related Raman peaks (Fig. 14b), particularly for the SLG glasses. The scatter in the KSG glasses is probably caused either by the sloping background and therefore the difficulty in uniquely fitting the Raman spectra, or by a difference in the details of the FeS speciation and different Raman scattering cross sections for the three different Raman peaks. In the XANES spectra the 2469 eV peak grows at the expense of the SH − and H2S peaks as Fe is added (Fig. 15) consistent with the observations in the Raman spectra. The effect of Fe on the SH −/ (SH − + H2S) can also be investigated. Fig. 16 shows the changes in peak intensities for SH −/(SH − + H2S) as a function of Fe/S for the SLG and KSG glasses. There is a very clear trend in the relative abundances of H2S and SH − in SLG, with H2S becoming relatively more abundant as Fe/S increases. The equivalent plot for KSG glasses is less clear. It could be interpreted as a similar trend for Fe/S > 0, but shifted to slightly lower SH −/(SH − + H2S) compared to that in SLG. Alternatively the data are consistent with a near zero-dependence of SH −/(SH − + H2S) on Fe/S. In addition it should be remembered that SH −/(SH − + H2S) could be modified upon quenching (Eq. (7)), possibly to different extents for different melt compositions with different viscosities. 4.3. Sulphur solubility at sulphur saturation As shown before, with increasing fO2 the S-species in the Fe-free glasses change from SH− and H2S to SO42 −. This change in Scomplexation also determines the quantities of S that can be incorporated in the glass structure. In Fig. 3b the total amount of sulphur of S-saturated glasses is plotted as a function of the nonbridging oxygens per tetrahedron (NBO/T) to evaluate a structural effect on S-solubility. At fully oxidised conditions, the SO42 − solubility shows a positive correlation with NBO/T for TROND, KSG and SLG glasses. The Albite glass (RS42), which is the most polymerised glass, has higher SO42 − contents compared to the more depolymerised

Fig. 13. XANES peak intensities at 2471.8 (filled diamonds) and 2466 (open diamonds) vs. the total S2 − content calculated from the S concentration in the glass and the S6 +/ ΣS determined by XANES (Table 1) for Fe-free SLG (a) and KSG (b) glasses. Error bars reflect the error of peak intensity determination by Gaussian fitting. Solid and dotted curves are fits of the data showing the relative change of H2S and SH− abundances in the glasses.

TROND and SLG glasses and does not follow this trend. This indicates that the SO42 − incorporation in Albite is more complex than that simply controlled by the abundance of non-bridging oxygens. At more reducing conditions, when the glasses are saturated with a mixture of sulphate and sulphide phases, there is a much stronger correlation of the S-content in the melt and NBO/T (Fig. 3b). This suggests that the incorporation of the reduced S-species (SH− and H2S) requires nonbridging oxygens in agreement with Eq. (7) and the model of (Burnham, 1979). It has to be noted that we do not have experiments that only contain SH− and H2S for SLG, KSG and Albite. However, when comparing the S-saturation level of TROND glass (Fig. 3b), that is higher at fully reduced conditions compared to the range of fO2 where S2 − and S6 + coexist, we can predict an increased gradient for the correlation with NBO/T at lower fO2 for SLG and KSG.

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Fig. 14. a) Sum of normalised XANES intensities at 2466 + 2471.8 eV vs. the sum of normalised Raman intensities at 2574 + 2601 cm− 1 for Fe-bearing KSG and SLG glasses. b) Normalised XANES intensities at 2469 eV vs. the sum of normalised Raman intensities at 298 + 372 + 420 cm− 1 for Fe-bearing KSG and SLG glasses.

The relatively high sulphide contents and the decreasing S-solubility when switching from S2 − to S 6 + is contrary to most observations and models for natural Fe-bearing systems (see Baker and Moretti, 2011, for a summary). For instance the model by Jugo (2009) predicts much lower S 2 − contents for basalt and andesitic compositions and an increase of S-solubility from reducing to oxidising conditions (Fig. 3a). It has to be noted that the concentration of S at sulphate saturation is comparable to basalt and andesite. Our observation that S6 + incorporation is enhanced in depolymerised structures is in agreement with the models of Jugo (2009) and other models (Baker and Moretti, 2011, and references therein) predicting higher sulphate contents in relatively depolymerised basalt compared to slightly more polymerised andesite and rhyolite compositions. This suggests that the SO42 − incorporation in Fe-bearing glasses is similar to those of Fe-free glasses. The observation that basalt incorporates more S6 + than the more

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Fig. 15. XANES peak intensities at 2469 eV (filled diamonds) and at 2466 + 2471.8 eV (open diamonds) vs. the molar Fe/S in the glasses for SLG (a) and KSG (b). Error bars reflect the error of peak intensity determination by Gaussian fitting.

depolymerised SLG composition requires an additional explanation such as that the type and amounts of cations in the melt (Mg, Ca, Na etc.) may also play a role. However, the S-contents of our depolymerised glasses at reducing conditions are an order of magnitude higher than that found for basalts, andesites and rhyolites. The only experimental evidence for higher sulphide solubility is the study of Scaillet and Macdonald (2006) reporting two times higher concentrations of sulphide relative to sulphate in peralkaline rhyolites. High sulphide contents in silicate melts is often correlated with the total Fe content (e.g., Carroll and Rutherford, 1987; Luhr, 1990; Wallace and Carmichael, 1992; O'Neill and Mavrogenes, 2002) and in empirical models the sulphide capacity at sulphide saturation (SCSS) of a silicate melt is largely determined by an Fe-controlled compositional factor (e.g., Mavrogenes and O'Neill, 1999; O'Neill and Mavrogenes, 2002; Liu et al., 2007; Baker and Moretti, 2011). High sulphide concentrations in simple systems such as diopside, albite and haplogranite compositions have been reported before (Mysen and Popp, 1980; Keppler,

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press). Because of the lack of a fully reduced Albite glass, the S 6 +/ ΣS of Albite was determined by using the fully reduced TROND spectrum as the compositions of these glasses are comparable. The obtained relative proportions of S 6 + and S 2 − expressed as S 6 +/ΣS are given in Table 1. S 6 +/S 2 − can be derived from S 6 +/ΣS by applying the following conversion:        6þ 2− 6þ 6þ ¼ S =ΣS = 1− S =ΣS S =S

ð8Þ

As already discussed in detail in previous studies (Wallace and Carmichael, 1994; Matthews et al., 1999; Métrich et al., 2009; Jugo et al., 2010; Wilke et al., 2011), the oxidation from S 2 − to S 6 + according to Eq. (2) has the following equilibrium constant: K ¼ aSO4

2−

  2 2− = aS  f O2

ð9Þ

with a for activity and f for fugacity. Assuming that the activities of SO42 − and S 2 − can be described by their concentrations expressed as c we obtain:   2− 2− ¼ 2 log f O2 þ log K log cSO4 =cS Fig. 16. Calculated SH−/(SH− + H2S) vs. the molar Fe/S of Fe-bearing SLG and KSG glasses. SH−/(SH− + H2S) was obtained using XANES peak intensities at 2466 and 2471.8 eV. Error bars are the propagated error of peak intensity determination by Gaussian fitting.

1999) indicating that Fe is not instrumental in achieving high Sconcentrations in a silicate melt. The water content of a melt also correlates positively with the S-content as suggested by Moune et al. (2009) but the effect is much less pronounced than that for Fe. Clemente et al. (2004) observed high sulphide contents in hydrous rhyolitic glasses at very low Fe/S ratios correlating with increasing fH2S in these experiments. Based on this observation they suggested that H2S may be present in these glasses but did not provide spectroscopic evidence for such an S-species. In this work we provide evidence that such sulphide species indeed may account for the high S-concentrations in the rhyolitic glasses of Clemente et al. (2004) and when adding Fe, FeS becomes the dominant S-species resulting in a drop in S2 − solubility. We have to emphasise that our determination of sulphur saturation is far from being complete as we do not provide systematic data for the S2 − saturation, activity–composition relationships of the melt components and the sulphur fugacity in our experiments. However, the high SH− and H2S solubilities are striking and together with the observation by Clemente et al. (2004) provide important implications on the behaviour of sulphur in natural Fe-poor silicate melts. 4.4. S 6 +/ΣS as a function of fO2 The linear combination of XANES spectra from glasses containing only S 6 + or S 2 −, respectively, can be used to quantify the relative proportions of S 6 + and S 2 − and the S 6 +/ΣS ratio of the glasses (Jugo et al., 2010; Klimm et al., in press; Wilke et al., 2011). Because this type of fit is also sensitive to regions in the spectra that are not sensitive to the sulphur oxidation state (Wilke et al., 2011), it is probably more precise focussing on features in the spectra that are particular sensitive to either S 2 − or S 6 + as suggested by Jugo et al. (2010). However, in our XANES spectra we can identify more than one energy region that is highly sensitive to S 2 −. For instance fully reduced KSG glass (RS76) shows highest intensities at ~ 2576 eV comparable to that in the natural basaltic glasses from Jugo et al. (2010), but glasses containing both, S 6 + and S 2 − (RS26 and RS53), show the most prominent feature related to S 2 − at 2471.8 eV (Fig. 5). For that reason we calculated S 6 +/ΣS by simple linear combination of fully reduced and oxidised glasses (Klimm et al., in

ð10Þ

Eq. (10) is in the form of y = 2x + b and will therefore define a straight line with a slope of 2 and an intercept that is a function of the equilibrium constant. Jugo et al. (2010) confirmed this theoretical approach in a XANES spectroscopic study on experimentally synthesised S-bearing basaltic glasses and suggested that previous determination of S 6 +/ΣS in natural glass compositions that do not fall on a trend with a slope of 2 (Carroll and Rutherford, 1988; Wallace and Carmichael, 1992; Nilsson and Peach, 1993; Wallace and Carmichael, 1994; Matthews et al., 1999; de Hoog et al., 2004; Jugo et al., 2005a) may have suffered either from beam damage during S 6 +/ΣS determination using EPMA or from modification of the S oxidation state during cooling of the glasses (Métrich et al., 2009). Our calculated S 6 +/S2 − ratios for the glasses that contain both, S6 + and S2 −, are shown in Fig. 17a together with the data of Jugo et al. (2010) and Botcharnikov et al. (2011) as a function of fO2. In the compositions studied here, S6 + and S2 − coexist in a range of fO2 from log fO2 = QFM− 1 to QFM + 1. Linear fitting of SLG and KSG data follow the theoretical consideration derived from Eq. (9) with a slope of 2.07 and 1.92 for SLG and KSG, respectively. The scattering of our data (r 2 = 0.63 and 0.71) is most likely related to the way we have determined the fO2 of the experiments, which is related to uncertainties in the calculations of the cH2O and aH2O. However, it is remarkable that the calculated slopes of the S 6 +/S 2 − equilibrium is very close to theoretical value of 2 indicating that our fO2 determination is correct. Because we only have two glasses for each of the other compositions (TROND and Albite) where S6 + and S2 − coexist, we have refrained from fitting those results. Qualitatively both, the TROND and Albite S6 +/S2 − are very close to SLG and KSG at constant fO2 with a similar slope of ~2. The calculated equilibrium of S6 +/S2 − at a constant fO2 is nearly identical for SLG and KSG implying that the compositional effect on S-speciation is minimal. So, amongst the compositions studied and within the resolution of our experiments, it is hard to determine any structural effect (e.g. degree of polymerisation) from playing a major role in the S6 +/ΣS–fO2 relationship as it, for instance, has been shown for the equilibrium of Fe 3 +/Fe2 + in silicate glasses (Kress and Carmichael, 1991). If there is any compositional effect on the S 6 +/S2 − equilibrium, it must be small (i.e. less than 0.5 log units of fO2 for any given S6 +/ΣS). We can also exclude an effect of fS2 on the S6 +/S 2 − in the glasses because both S-saturated and S-undersaturated experiments yield similar S 6 +/ ΣS ratios (Table 1) at comparable fO2. Our results are in close agreement with the determined S6 +/ΣS ratios in sodium-silicate glasses

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S≫ 2. In particular, it has to be noted that TROND and Albite glasses show the same shift towards lower fO2 when compared to Fe-bearing dacite glasses (Carroll and Rutherford, 1988) which are compositionally and structurally similar. As we have shown and discussed above, the Fe/S of the glass controls the dissolved S 2 −-species in the glass. Therefore, we also conclude that Fe has a fundamental influence on the S6 +/S2 − equilibrium. If the Fe/S is in the range where all SH− and H2S are consumed, additional Fe will incrementally reduce SO42 − to S2 − and therefore, decrease the S6 +/ΣS according to Eq. (17). 4.5. The overall solubility mechanism Taking all the data and previous discussion into account, we can propose a series of reactions that describe the interaction of elemental sulphur with hydrous silicate melts and the subsequent speciation reactions within the melt. In our experiments we have elemental sulphur and water in the starting materials, and in the glasses we produce, sulphur is in the form of either sulphide or sulphate. We can write the following equations based solely on the stoichiometry of the overall reactions: S2ðsÞ þ 8H2 OðmÞ ¼ 2H2 SO4ðmÞ þ 6H2ðvÞ ðoxidation of sulfurÞ

ð11Þ

S2ðsÞ þ 2H2 OðmÞ ¼ 2H2 SðmÞ þ O2ðvÞ ðreduction of sulfurÞ

ð12Þ

These reactions are not meant to imply that S2 is a major species in a fluid coexisting with a hydrous melt, they are simply intended to show the consequences of oxidation or reduction when adding elemental sulphur to a hydrous melt and are particularly appropriate for fluid undersaturated conditions (as is the case for most of our experiments). In the more general case of excess hydrous fluid we also have to consider the speciation of sulphur in the vapour phase (Scaillet et al., 1998): S2ðsÞ þ4H2 OðvÞ ¼ 2SO2ðvÞ þ4H2ðvÞ ðpartial oxidation of sulfur in the vapourphaseÞ

ð13aÞ 2SO2ðvÞ þ4H2 OðmÞ ¼ 2H2 SO4ðmÞ þ2H2ðvÞ ðfurther oxidation of sulfur upon dissolutionÞ

ð13bÞ Fig. 17. Sulphur oxidation state as a function of fO2 expressed as ΔQFM for SLG (filled circles), KSG (grey circles), TROND (open squares) and albite (open diamonds) glasses compared to basaltic and andesitic glasses determined by Jugo et al. (2010; crosses) and Botcharnikov et al. (2011; stars). a) Log (S6 +/S2 −) vs. ΔQFM. Solid and grey lines are linear regressions for SLG and KSG values, respectively. Dotted curve is the linear regression for basalt. All fits have a slope of ~ 2. b) S6 +/ΣS vs. ΔQFM solid and dashed lines are the linear regressions shown in a). The change from S2 − to S6 + in the Fe-free glasses studied here occurs at ~ 1.5 log units lower fO2 compared to basalt.

synthesised experimentally at 1 atm (Fig. 1, Nagashima and Katsura, 1973), a study that has previously been excluded from models aiming to calculate the S 6 +/ΣS ratio in natural silicate liquids (Matthews et al., 1999; Jugo et al., 2005a). When comparing our results on “synthetic” glasses to those obtained for “natural” glasses from Jugo et al. (2010) and Botcharnikov et al. (2011) it is apparent that our data are shifted to ~1.5 log units lower fO2 for a constant S6 +/S2 −. Fig. 17b shows all of our data and those for basalt from Jugo et al. (2010) and for andesite from Botcharnikov et al. (2011) together with the fits obtained for the equilibrium of S6 +/S2 − according to Eq. (9) for SLG and basalt from Jugo et al. (2010). In the “synthetic” glasses, sulphur is dissolved as S2 − at fO2 b QFM− 1 and as S6 + at fO2 >QFM+ 1. In the range of QFM− 1 to QFM +1, sulphur is dissolved as a mixture of S6 + and S2 − species. The difference between the “synthetic” and “natural” glasses is the high Fe of the latter with Fe/

S2ðsÞ þ2H2 OðvÞ ¼2H2 SðvÞ þO2ðvÞ ðreduction of sulphur in the vapour phaseÞ

ð13cÞ In Eqs. (11) and (13b) sulphate is expressed as H2SO4 for simplicity, but in fact sulphate appears to be dissolved as an isolated SO42 − group charge balanced by cations. The type of cation depends on the overall glass chemistry and for a Ca-bearing system the sulphate complexing reaction can be written as: H2 SO4ðmÞ þ CaOðmÞ ¼ CaSO4ðmÞ þ H2 OðmÞ

ð14Þ

As shown earlier, we have good evidence for an H2S speciation reaction analogous to that for H2O (Eq. (7)). Our data are consistent with all dissolved sulphide being present as H2S or Si–SH in ironfree melts confirming the Burnham (1979) model for the sulphide solubility mechanism (Eq. (4)). This result disagrees with some previous suggestions that, depending on the bulk glass chemistry, alkali or alkaline earth metal sulphide groups such as NaS, CaS, MgS etc. are present in the glass (Métrich et al., 2009; Wilke et al., 2011). The equilibrium position of Eq. (7) depends on total sulphide concentration (Fig. 13) and glass composition (compare SLG and KSG compositions where Fe/S = 0 in Fig. 16). The sulphide speciation reaction changes in the presence of iron. In our samples with low concentrations of Fe (Fe/S up to approximately

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2.6), the following reactions apply. It is important to note that these reactions do not involve any oxidation or reduction. H2 SðmÞ þ FeOðmÞ ¼ FeSðmÞ þ H2 OðmÞ

ð15Þ

Si  SHðmÞ þ FeOðmÞ ¼ FeSðmÞ þ Si  OHðmÞ

ð16Þ

The position of equilibrium of both of these reactions is apparently far to the right under the conditions of glass synthesis, but there does not appear to be a quantitative scavenging of S 2 − by Fe, because there is some residual S–H species even where none is required by stoichiometry. The relative importance of these two reactions appears to depend on glass composition; as shown in Fig. 16, SH −/(H2S + SH −) decreases with increasing Fe/S for SLG and KSG compositions, implying that reaction (16) is favoured relative to reaction (15). KSG has slightly lower SH −/(H2S + SH −) compared to SLG. In ironbearing glasses that have much more Fe than the Fe-bearing samples from this study, the 2469 eV XANES peak is absent and another feature at ~2470 eV is visible especially in glasses that are saturated with a Fe–S phase (Métrich et al., 2009; Jugo et al., 2010). It is therefore implied that in natural glasses the S 2 − has a different coordination environment from that observed here. This change is likely to be related to an increased number of Fe 2 + ions coordinating the S 2 −. Once reactions (15) and (16) have gone to completion (or if the equilibrium is far on the RHS), addition of further FeO requires a new reaction, if further FeS(m) is to be created. This, therefore, occurs by reduction of sulphate and consequent oxidation of Fe (of course only if the fO2 is in the range that SO42 − is present in the Fe-poor melt; see Fig. 17): H2 SO4ðmÞ þ 9FeOðmÞ ¼ FeSðmÞ þ 4Fe2 O3ðmÞ þ H2 OðmÞ

ð17Þ

Free oxygen, O2 does not appear in reaction (17), yet sulphur is oxidised on the LHS and reduced on the RHS. Thus for a fixed oxygen fugacity addition of FeO beyond that required in reactions (15) and (16) converts sulphate to sulphide. Therefore, reaction (17) qualitatively explains why the S oxidation state curve (Fig. 17) for Fe-rich melts is displaced to the right by more than 1 log unit in fO2. This reaction gives the surprising prediction that there is no unique sulphur oxidation curve as a function of fO2. The more FeO is added to the melt, the more reduced the sulphur becomes. On closer consideration this suggestion is not ruled out by the existing database. Many of the previously published sulphur oxidation curves have been already shown to be inaccurate because of analytical problems (Jugo et al., 2010; Wilke et al., 2011), back-reaction on cooling (Métrich et al., 2009) etc. It also has to be noted that the fO2 determination of natural melt inclusions is inferred from the Fe 3 +/ΣFe of the glass (e.g., Nilsson and Peach, 1993; Wallace and Carmichael, 1994; Matthews et al., 1999; de Hoog et al., 2004) according to an empirical model (Kilinc et al., 1983; Kress and Carmichael, 1991) that has never been calibrated for a sulphurbearing system. It is quite possible that the remaining data give an apparently consistent single curve because of the limited range of Fe/S ratio studied and the uncertainties on both parameters (fO2). Further experimental investigations are required to clarify this important observation. 4.6. Speciation in the melt at high temperature Upon cooling in each experiment, the high temperature melt was transformed into a glass. A fundamental question is therefore whether the determined S-species in the glass and the oxidation states of S and Fe are modified from the speciation at high temperature or not. The glass transition temperature (Tg) is the temperature at which speciation reactions are frozen in (Dingwell and Webb, 1990). Tg is

usually ~2/3 of the melting temperature (Tm), therefore if a speciation reaction is T-dependent, then the position of equilibrium recorded in the glass is not the same as that in situ in the super-liquidus melt. The H2O + O2 − = 2OH− conversion, moves to the right with increasing T (Behrens and Nowak, 2003; Zhang et al., 1997), and therefore measurements of OH−/H2O in glasses do not reflect the actual speciation at run temperature, and are shifted towards higher H2O abundances. If the SH−/H2S-reaction has a similar enthalpy to the OH−/H2O reaction, then the SH−/(SH− + H2S) ratio may be similarly affected upon cooling and the H2S abundances would be lower in the melt at high temperature than those observed in our glasses. However, the thermodynamics of H2S-speciation in melts is unknown and it is not possible to predict a priori the direction of changes in speciation upon quenching. Nonetheless, it is likely that SH− is a stable species in reduced, low-Fe, hydrous silicate melts, in contrast to the proposal of Stelling et al. (2011) that SH− is just a quench related species by the reaction of H2O and S2 −. Métrich et al. (2009) suggested that cooling may also affect the S6 +/ ΣS and Fe3 +/ΣFe due to an electron exchange reaction between S2 − and Fe 3 + producing S6 + and Fe2 +. In our experiments we do not find evidence for such processes. The S6 +/S 2 − ratios of our glasses follow the theoretical dependence on fO2, even if they have been synthesized in different pressure vessels with different cooling rates (IHPV and CSPV) and adding Fe does not affect the S6 +/ΣS ratio or the relative abundances of Fe3 + and Fe2 +. 4.7. Implications for natural magmatic systems The observations that at low Fe/S the S 2 − solubility can be high and that the Fe/S influences the equilibrium of S 6 +/ΣS and fO2 provide some interesting implications for natural felsic melts, where the Fe/S ratio is low and where a change of Fe/S during the magmatic evolution of the system occurs. Such felsic melts occur in large volumes in subduction zone settings and are either the result of fractionation during magmatic evolution in the upper crust as in the case of rhyolites or represent first melting products just above the solidus of the subducted oceanic crust deep in the mantle. In the case of rhyolites at upper crustal conditions, our findings suggest that magma may contain larger quantities of S, dissolved as SH − and H2S at logfO2 b QFM, if the Fe/S is around 2 or less (as predicted from our S-speciation curve in Fig. 17b), compared to Fe richer compositions. Such a scenario is in perfect agreement with the experiments by Clemente et al. (2004) that also determined the highest S-concentration of ~1000 ppm S in rhyolitic compositions at logfO2 b QFM and Fe/S ~ 2. However, considering that for subduction zone magmas the determined fO2 is generally logfO2 > QFM (e.g., Carmichael, 1991), SH − and H2S will not be stable even at Fe/S b 2 and thus, these S-species will not play a role on the total net S-load of rhyolites in subduction zones. In the case of melting or fluid transfer from the subducted oceanic crust, the relationship of SH −/H2S, SO42 − and fO2 may be more important as the fO2 is generally lower compared to arc magmas. Dehydration or melting of the subducted oceanic lithosphere occurs at fO2 determined for MORB or upper mantle at logfO2 b QFM (although some upper parts of the slab may be more oxidised through hydrothermal alteration or contribution of sediment) and the liberated liquids reflect the fO2 of their source region (Carmichael, 1991). Close to the basalt solidus such melts are trondjemitic in composition and carry only small amounts of FeO (in some cases less than 0.5 wt.%; Prouteau et al., 2001; Kessel et al., 2005; Klimm et al., 2008a). As we have shown before and in agreement with Clemente et al. (2004), such liquids are able to carry larger quantities of S than predicted from Fe-rich systems, presumably in the form of SH − and H2S. Due to the strong partitioning of sulphur into fluids (Scaillet et al., 1998; Keppler, 1999; Webster and Botcharnikov, 2011) the S concentration may be even higher considering that the

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dehydration or melting of the slab may occur in the supercritical field, where fluids and melts become fully miscible (e.g., Bureau and Keppler, 1999; Klimm et al., 2008a). The resulting Fe/S ratio of such supercritical slab liquids is ≪ 2 and a significant proportion of sulphur is present as SO42 − according to the S-speciation curve for Fe-free (Fe-poor) systems (Fig. 17b). If the initial oxidation state of sulphur in the slab is less oxidised than that of the released fluids, the slab becomes more reduced and continued subduction therefore removes a reduced component from the arc system. The slab liquids migrate through the mantle and carry sulphate to the mantle wedge, the primary source region of arc magmas, and react with the mantle- and sediment-derived melts to form basaltic to andesitic melts (Elliott et al., 1997). The FeO content in the produced melt continuously increases with progressive melting of the mantle wedge and Fe/S ratio increases to ≫2. This shifts S from the stability of sulphate to that of sulphide and, according to reaction 17, Fe 2 + is successively oxidised to Fe 3 + resulting in an increase of the Fe 3 +/ ΣFe ratio in the mantle wedge. Thus, the fO2 prevailing in the mantle wedge, calculated using the Fe 3 +/ΣFe ratio (Ballhaus et al., 1991; Kress and Carmichael, 1991; Wood and Virgo, 1989), increases in response to the reaction of Fe with sulphate from a slab liquid. The scale of such a process can be estimated assuming, for example, an andesitic melt containing FeOtotal = 10 wt.%. The change in the Fe 3 +/ΣFe ratio of this melt from 0.1 to 0.2 corresponds to an increase of the fO2 of 2 log units (from QFM − 0.5 to QFM + 1.5; Fig. 1) and requires only 3000 ppm SO42 − (=1000 ppm of S) according to reaction 17. Considering that some of the S in the Fe-poor slab liquid may still be present as S 2 − according to the S 6 +/ΣS ratio of 0.6 to 0.2 at fO2 = QFM to QFM − 0.5 (Fig. 17b), the total amount of S required is in the range of 1000 to 3000 ppm S, which matches the range of S concentrations determined for primary arc magmas (Portnyagin et al., 2007). The extent of increasing the Fe 3 +/ΣFe ratio by adding oxygen due to the reduction of S 6 + to S 2 − is limited to maximum values of fO2 = QFM + 2 when all S is dissolved as S 6 + in Fe-bearing systems. Thus, the whole range of mantle wedge fO2 from log fO2 = QFM to QFM + 2 (Arculus, 1994; Brandon and Draper, 1996; Parkinson and Arculus, 1999) can be explained by ingress of SO42 − from subducted slab. A further increase of fO2 in arc magmas to values up to QFM + 4 is most likely related to volatile degassing during magma ascent (Burgisser and Scaillet, 2007). 5. Conclusions The combination of Raman and X-ray absorption spectroscopy on experimentally equilibrated sulphur-bearing, hydrous silicate glasses under controlled fO2 provides some unique new insights on the dissolution mechanism of S in silicate melts: i) in Fe-free melts S is dissolved as SH −, H2S and/or SO42 − depending on the prevailing fO2; ii) adding Fe results in the formation of Fe–S complexes at the expense of SH − and H2S, which are still observed up to Fe/S ~ 2.6 depending on the bulk composition; iii) the S 6 +/S 2 − equilibrium in Fe-free systems is shifted by ~ 1.5 log units to lower fO2 compared to Fe-rich systems; iv) the total sulphur solubility depends on the degree of polymerisation of the melt and increases with increasing NBO/T. This correlation is much more pronounced for SH− and H2S than for SO42 −; and v) in Fe-free systems S2 − is more soluble than S6 +. The oxidation state of S therefore depends not only on fO2, as previously assumed, but also on composition, in particular Fe. This may have consequences for the S-cycle in subduction zones as Fe-poor slabderived liquids are able to carry S6 + and the S6 + may act as an oxidising agent in the mantle wedge by successively oxidising Fe2 + to Fe3 +. Acknowledgements This work was supported by the NERC grant NE/C510967/1 and the British German Academic Research Collaboration Programme (ARC)

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through a grant number 1280. Access to SLS was supported by EU (proposal ID: 20060889 and 20070428). The help and advice of M. Janousch, L.A. O'Dell, S. Kearns, and K. Schollenbruch using the LUCIA beamline at SLS, the EPMA at University of Bristol and the Raman-microscope at Goethe-University of Frankfurt is greatly appreciated. The work benefitted greatly from stimulating discussions with J.D. Blundy, A. Berry, M.J. Walter, H. Behrens, M. Wilke and A.B. Woodland and the thoughtful review by B. Scaillet, and we thank them all for their contributions. Appendix A. Supplementary data Supplementary data to this article can be found online at http:// dx.doi.org/10.1016/j.chemgeo.2012.04.028. References Aiuppa, A., Federico, C., Giudice, G., Gurrieri, S., Paonita, A., Valenza, M., 2004. Plume chemistry provides insights into mechanisms of sulfur and halogen degassing in basaltic volcanoes. Earth and Planetary Science Letters 222 (2), 469–483. Arculus, R.J., 1994. 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