# Effect of sulfur speciation on chemical and physical properties of very reduced mercurian melts

Brendan A. Anzures, Stephen W. Parman, Ralph E. Milliken, Olivier Namur, Camille Cartier, Sicheng Wang

► To cite this version:

Brendan A. Anzures, Stephen W. Parman, Ralph E. Milliken, Olivier Namur, Camille Cartier, et al.. Effect of sulfur speciation on chemical and physical properties of very reduced mercurian melts. Geochimica et Cosmochimica Acta, 2020, 286, pp.1-18. 10.1016/j.gca.2020.07.024. insu-03712896

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1 **Effect of Sulfur Speciation on Chemical and Physical Properties of Very**

2 **Reduced Mercurian Melts**

3

4 **Brendan A. Anzures¹, Stephen W. Parman¹, Ralph E. Milliken¹, Olivier Namur², Camille**

5 **Cartier³, Sicheng Wang⁴**

6 ¹Brown University, Department of Earth and Planetary Sciences, USA

7 ²KU Leuven, Department of Earth and Environmental Sciences, Belgium

8 ³CRPG/CNRS, University of Lorraine, France.

9 ⁴Centre for Mined Land Rehabilitation, Sustainable Minerals Institute, The University of

10 Queensland, Brisbane 4072, Australia

11

12 Corresponding author: first and last name (brendan_anzures@brown.edu)

13
14 Abstract

The NASA MESSENGER mission revealed that lavas on Mercury are enriched in sulfur (1.5-4 wt.%) compared with other terrestrial planets (<0.1 wt.%), a result of high S solubility under its very low oxygen fugacity (estimated fO₂ between IW-3 and IW-7). Due to decreasing O availability at these low fO₂ conditions, and an abundance of S²⁻, the latter acts as an important anion. This changes the partitioning behaviour of many elements (e.g. Fe, Mg, and Ca) and modifies the physical properties of silicate melts. To further understand S solubility and speciation in reduced magmas, we have analysed 11 high pressure experiments run at 1 GPa in a piston cylinder at temperatures of 1250 to 1475 °C and fO₂ between IW-2.5 to IW-7.5. S K-Edge XANES is used to determine coordination chemistry and oxidation state of S species in highly reduced quenched silicate melts. As fO₂ decreases from IW-2 to IW-7, S speciation goes through two major changes. At ~IW-2, FeS, FeCr₂S₄, Na₂S, and MnS species are destabilized, CaS (with minor Na₂S) becomes the dominant S species. At ~ IW-4, Na₂S is destabilized, MgS becomes the dominant S species, with lesser amounts of CaS. The changes in S speciation at low fO₂ affect the activities of SiO₂, MgO and CaO in the melt, stabilizing enstatite at the expense of forsterite, and destabilizing plagioclase and clinopyroxene. These shifts cause the initial layering of Mercury's solidified magma ocean to be enstatite-rich and plagioclase poor. Our results on S speciation at low fO₂ are also applicable to the petrologic evolution of enstatite chondrite parent bodies and perhaps early Earth.
## 1. Introduction

Sulfur is one of the most abundant and versatile volatile elements in magmas. Magmatic volatiles such as H, C, S, Cl, N, and F play crucial roles in the production, transport, and eruption of magmas under terrestrial and lunar conditions. However, the behavior of sulfur is highly complex, partly due to its heterovalent nature (oxidation state ranging from 2- to 6+) (Behrens & Webster, 2011). Under different redox conditions, sulfur can bond with both more electropositive and more electronegative elements (Fleet, 2005), having significant effects on the partitioning behaviors of these elements between silicate melts (S²⁻ and S⁶⁺), liquid metals and sulfides (S⁰ and S²⁻), gases (H₂S, S₂, SO₂, and SO₃), and solids (S²⁻, S¹⁻, S⁰, S²⁺, S⁴⁺, and S⁶⁺).

Thus, sulfur is important in understanding many geologic problems including the eruptive style of volcanoes (e.g. Head et al., 2012; Shinohara, 2008; Sparks, 1997; Sutton et al., 2001) subduction zone recycling (e.g. Hilton et al., 2002; De Hoog et al., 2001; Métrich et al., 1999; Wallace, 2005), the formation of magmatic-hydrothermal ore deposits (e.g. Hamlyn et al., 1985; Hattori & Keith, 2001; Jenner et al., 2010; Righter et al., 2008), and atmospheric fluxes of volcanic gases and related climate effects (e.g. Andres & Kasgnoc, 1998; Hammer et al., 1980; Lamb, 1969; McCormick et al., 1995; Robock, 2000; P. J. Wallace, 2001). Therefore, detailed understanding of the concentration of sulfur as it moves from magmas to the surface is vital.

Sulfur in silicate melts is dissolved as S²⁻ and S⁶⁺ species (Fincham & Richardson, 1954) with previous detections of S⁴⁺ (Métrich, 2002) now recognized as beam damage artifacts from the electron microprobe or synchrotron x-rays (Wilke et al., 2008) during measurement.

Experiments have shown that sulfur solubility at saturation is controlled by the nature of the saturating phase and the composition of the silicate melt (M. R. Carroll & Rutherford, 1985, 1987; Luhr, 1990), but also intensive conditions such as pressure, temperature, and oxygen
fugacity ($fO_2$). The transition of the primary dissolved sulfur species from $S^{2-}$ to $S^{6+}$ occurs over a narrow range of oxygen fugacity (FMQ (Fayalite-Magnetite-Quartz buffer) to about FMQ+2) witnessed in sulfur solubility and $S^{6+}/\Sigma S$ sulfur speciation curves (Jugo et al., 2010). Many terrestrial magmas are oxidized with their maximum dissolved sulfur concentration controlled by saturation of either anhydrite (M. R. Carroll & Rutherford, 1987; Luhr, 1990) or immiscible molten sulfate (Jugo et al., 2005). On the other hand, most mid ocean ridge basalts (MORB) have $fO_2$ below FMQ meaning sulfur concentration in reduced silicate melts is controlled by sulfide saturation (Fincham & Richardson, 1954; O'Neill & Mavrogenes, 2002; P. Wallace & Carmichael, 1992). Sulfide species have a stronger affinity for network-modifying cations such as Fe, Mg, and Ca than network forming cations Si and Al such that as $S^{2-}$ dissolves in the silicate melt, distinct sulfide complexes should be formed (Fleet, 2005). This may act to break up the silicate melt structure similar to the formation of phosphates or carbonates (Hess, 1995).

Mercury evolved as a geochemical and mechanical endmember among the terrestrial planets, largely due to its low $fO_2$. The Earth, Moon, and Mars have similar mantle $fO_2$ between IW-2 and IW+2 (IW = iron-wüstite equilibrium) (Wadhwa, 2008) as seen in **Figure 1**. Sulfur solubility in magmas is low (<0.1 wt.%) at these conditions and is present as an important but minor element (Namur, Charlier, et al., 2016). In contrast, the Mercurian surface exhibits high S concentrations (1.5-4 wt.%) (Nittler et al., 2011). If these values reflect magma compositions, then a mantle $fO_2$ between IW-3 and IW-7 is implied from silicate-metal-sulfide equilibria (McCubbin et al., 2012; Zolotov et al., 2013). Based on models of S solubility in silicate melts, Namur, Charlier, et al. (2016) estimated an average $fO_2$ of IW-5.5 for Mercury's mantle. The only other planetary materials thought to have formed at such low $fO_2$ conditions are enstatite chondrites and aubrites (Berthet et al., 2009; Casanova et al., 1993; Fogel et al., 1989; McCoy et
al., 1999; Sears et al., 1982; Watters & Prinz, 1980). Additionally, some studies have suggested a highly reduced early Earth (Wohlers & Wood, 2015). Studies of low fO₂ systems also are also applicable to the petrologic evolution of exoplanets orbiting C-rich suns.

High S solubility in silicate melts has been observed in numerous experimental studies at highly reducing conditions (Berthet et al., 2009; Cartier et al., 2014; Chabot et al., 2014; Holzheid & Grove, 2002; Vander Kaaden et al., 2015; Malavergne et al., 2007, 2014; McCoy et al., 1999; Namur, Charlier, et al., 2016) (Figure 2). Due to decreasing O concentrations at these low fO₂ conditions, S²⁻ fills in for O²⁻ as a second significant anion, changing the partitioning behavior of many elements (e.g. Fe, Mg, and Ca and other elements that become more chalcophile) between silicate and sulfide as well as modifying the structure and physical properties of the silicate melt. Enstatite chondrite-derived thermodynamic models (Ebel et al., 2011), thermodynamic models based on Mercury's surface composition (Zolotov et al., 2013), and raman spectra (Namur et al., 2016a) suggest S switches from bonding with Fe to bonding with Mg and Ca in silicate melts at these reducing conditions. The formation of MgS and CaS lowers the activity of MgO and CaO in the melts and decreases the stability of minerals rich in these oxides (e.g. olvine and plagioclase). Cations bonded with S will be available to precipitate sulfides once saturation of the sulfide species is attained but will not be available for silicate mineral reactions. Sulfide speciation is particularly important under reducing conditions due to high sulfur solubility in the silicate melt (Figure 2). As an added consequence, sulfur at sulfide saturation may not act as a volatile element because the system would precipitate immiscible sulfide melts or crystals rather than gaseous species.

In this study, we use S K-edge X-ray Absorption Near Edge Structure (XANES) spectroscopy to quantify the coordination chemistry of sulfur in quenched silicate melts under
very reducing conditions (IW-2.5 to IW-7.2). We show quantitative S speciation trends in very reduced S-rich and Fe-poor silicate melts for MgS, CaS, FeS, FeCr₂S₄, Na₂S, and MnS as a function of S solubility in silicate melt and of oxygen fugacity. The oxide activities of cations that bond with S change, thus affecting silicate phase equilibria (Namur, Charlier, et al., 2016; Namur, Collinet, et al., 2016). We discuss the effect changes in activity have on previous oxygen fugacity models and how S speciation may itself be a potential oxybarometer at very reducing conditions. We also discuss how changes in cation activities as a function of sulfur dissolution (or fO₂) affect the initial mineralogical layering of Mercury's mantle, post magma ocean crystallization. This layering sets the stage for the subsequent thermal and magmatic evolution of the planet. We finally evaluate how changes in silicate melt structure may affect physical properties such as viscosity, density, and melting temperature (Fogel, 2005; Holzheid & Grove, 2002; Namur, Charlier, et al., 2016).

# 2. Materials and Methods

## 2.1 Starting Material and Bulk Composition

The primitive Mercurian mantle composition used in this study was determined by comparing the chemical fractionation of Mercurian lavas to chondritic compositions representing the possible building blocks of the planets, a method employed by Dreibus et al. (1998) for Mars and by Drake & Righter (2002) for Earth (Figure 3). CH3 chondrite Allan Hills 85085 (Weisberg et al., 1988) was chosen as a Mercurian analog due to its similarities in major element ratios (Mg/Si, Al/Si, Fe/Si, and Ni/Si) to an estimated bulk composition derived from Mercurian surface measurements, high metal content consistent with Mercury's large core (Ebel & Stewart, 2018), and relatively reducing conditions (estimated IW-5.5 (Petaev et al., 2003). The bulk composition of ALH 85085 was used to determine the 13 elemental concentrations of this
primitive Mercurian mantle though iron and sodium were modified (Table 1). Iron (4.0 to 0.54 wt.% FeO) and sodium (0.4 to 0.08 wt.% Na₂O) were modified to reflect a primitive mantle after core formation and to be consistent with the high sodium abundances (2.75 wt.% to 7.00 wt.% Na₂O) detected on the Mercurian surface (Nittler et al., 2011, 2018; Peplowski et al., 2011, 2014). While many experimental studies have used Indarch as a Mercury bulk composition, this composition appears to be too rich in Si to produce Mercury's lavas (i.e. it does not fall on the line defined by Mercurian lavas in Figure 3). Removal of Si into the mostly Fe core (Berthet et al., 2009; Charlier & Namur, 2019) could increase Mg/Si and Al/Si ratios of an Indarch composition to be consistent with the lavas. While Mercury may have an enstatite chondrite bulk composition, it seems like the mantle does not; using an enstatite chondrite for core-mantle partitioning experiments is fine, but not for mantle melting. Such a mantle composition would be similar to the ALH85085 used here. We recognize that this meteorite is not an enstatite chondrite and likely formed in the outer solar system. Nevertheless, its composition appears to be a good model for Mercury's mantle. Melting experiments on it have produced melt compositions nearly identical to Northern Smooth Plains lavas. Compositions of our primitive Mercurian mantle composition that are used as the starting compositions are compared with previous estimates of the Mercurian mantle, terrestrial mantle, and Mercurian surface lavas are included in Table 1.
142

<table><caption>**Table 1.** Estimates of Mercury's Mantle Composition (wt.%)</caption><thead><tr><th></th><th>Primitive<br>Mercury<br>Mantle*</th><th>Primitive<br>Mercury<br>Mantle (80%<br>crystallized)*</th><th>Pre-<br>melting<br>Mantle<br>(NSP)<sup>a*</sup></th><th>Pre-melting<br>Mantle (IcP-<br>HCT)<sup>a*</sup></th><th>Terrestrial<br>Mantle<sup>b</sup></th><th>CH<br>ALH85085<sup>c</sup></th><th>EH4<br>Indarch<sup>d</sup></th><th>NVP<sup>a</sup></th><th>IcP-HCT<sup>a</sup></th></tr></thead><tbody><tr><td>SiO<sub>2</sub></td><td>51.32</td><td>52.8</td><td>53.67</td><td>51.98</td><td>49.5</td><td>48.9</td><td>65.3</td><td>58.7</td><td>52.7</td></tr><tr><td>TiO<sub>2</sub></td><td>0.21</td><td>1.17</td><td>0.24</td><td>0.21</td><td>0.2</td><td>0.2</td><td>0.1</td><td>0.04</td><td>0.4</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>3.78</td><td>17.33</td><td>4.57</td><td>4.24</td><td>3.6</td><td>3.6</td><td>2.1</td><td>13.8</td><td>8.79</td></tr><tr><td>Cr<sub>2</sub>O<sub>3</sub></td><td>0.41</td><td>0.09</td><td>**</td><td>**</td><td>0.4</td><td>0.8</td><td>0.1</td><td>**</td><td>**</td></tr><tr><td>FeO</td><td>0.54</td><td>0.02</td><td>0.02**</td><td>0.03**</td><td>7.1</td><td>4</td><td>0.4</td><td>0.04**</td><td>0.04**</td></tr><tr><td>MnO</td><td>0.09</td><td>0.2</td><td>**</td><td>**</td><td>0.1</td><td>0.4</td><td>0.3</td><td>**</td><td>**</td></tr><tr><td>MgO</td><td>40.72</td><td>14.92</td><td>36.89</td><td>37.64</td><td>35.7</td><td>38.8</td><td>26.9</td><td>13.9</td><td>27.8</td></tr><tr><td>CaO</td><td>2.73</td><td>10.91</td><td>2.26</td><td>3.84</td><td>2.8</td><td>2.6</td><td>1.4</td><td>5.81</td><td>7.27</td></tr><tr><td>Na<sub>2</sub>O</td><td>0.08</td><td>2.22</td><td>1.97</td><td>1.29</td><td>0.3</td><td>0.4</td><td>1</td><td>7</td><td>2.75</td></tr><tr><td>K<sub>2</sub>O</td><td>0.02</td><td>0.44</td><td>0.05</td><td>0.04</td><td>0</td><td>0.1</td><td>0.1</td><td>0.2</td><td>0.08</td></tr><tr><td>P<sub>2</sub>O<sub>5</sub></td><td>0</td><td>0</td><td>**</td><td>**</td><td>0</td><td>0.2</td><td>0</td><td>**</td><td>**</td></tr><tr><td>NiO</td><td>0.11</td><td>0.01</td><td>**</td><td>**</td><td>0.2</td><td>3.5</td><td>2.4</td><td>**</td><td>**</td></tr><tr><td>Total</td><td>100.01</td><td>100.11</td><td>99.67</td><td>99.27</td><td>99.9</td><td>103.5</td><td>100.1</td><td>99.49</td><td>99.83</td></tr></tbody></table>

143

<sup>a</sup>(Nittler et al., 2018), <sup>b</sup>(Hart & Zindler, 1986), <sup>c</sup>(Weisberg et al., 1988), <sup>d</sup>(Berthet et al., 2009)

144

*Primitive Mercury Mantle is modelled from CH ALH85085, while Primitive Mercury Mantle (80% crystallized) is the experimental residual liquid composition in this work. Pre-melting Mantle compositions are modelled by (Nittler et al., 2018).

145

146

**Cr, Mn, P, and Ni not constrained by MESSENGER affecting NVP and IcP-HCT compositions as well as Pre-melting Mantle estimated compositions. These FeO contents do not include reduced Fe (1-2 wt.% Fe on Mercury's surface measured by MESSENGER (Nittler et al., 2011, 2018; Weider et al., 2014)

147

148
Following previous experimental studies (Berthet et al., 2009; Namur, Charlier, et al., 2016), $fO_2$ was controlled (but not buffered) by changing the ratio of $Si^0$ metal to $SiO_2$ in the starting composition. All of the $Si^0$ is oxidized during the experiment, lowering the $fO_2$, though not buffering it. Based on previous experiments, 95 mol% $SiO_2$ and 5 mol% Si metal produces an $fO_2$ of about IW-3.

Silicate starting compositions were produced from high-purity reagent grade powder oxides and carbonates. Two separate powder mixes were made, one in which all Si is added as $SiO_2$ and one in which all Si is added as $Si^0$. Mixing these two allows variation of $Si^0/SiO_2$ in the mix while maintaining the same molar concentrations of all cations. The $Si^0$ is added after decarbonation and conditioning. The oxide/carbonate powders were ground mixed under ethanol in an agate mortar and pestle. This mix was decarbonated and conditioned in a horizontal gas flow furnace (900 °C for 4 hours) with a CO/H$_2$ gas ratio representing an $fO_2$ near IW-1. $Si^0$ powder was then added to the $SiO_2$-free mix to match the molar concentration of $SiO_2$ in the starting composition in order to not change the bulk composition. Finally, elemental sulfur powder was added to ensure sulfide saturation (depending on $fO_2$, P, and T).

## 2.2 Experimental Methods

Experiments were run in a ¾" piston-cylinder (PC) apparatus using inner pieces of crushable MgO, straight-walled graphite furnaces, and outer sleeves of Pyrex and sodium chloride as the pressure cell displayed in **Figure 4**. Runs were completed using graphite capsules inside molybdenum capsules for a sample size of 2mm height and 2mm diameter. Graphite inner capsules were used because graphite minimally reacts with silicate liquids under reducing conditions (Ardia et al., 2013; Vander Kaaden & McCubbin, 2016). Unfortunately, graphite is also porous and could allow rapid diffusion of oxygen into the sample especially with such a
large $fO_2$ difference between the sample and air (over 12 orders of magnitude). Terrestrial mantle melting experiments typically use graphite inner capsules with platinum outer capsules where the platinum metal acts as a barrier for oxygen and volatiles. Molybdenum outer capsules were used to minimize oxygen diffusion in and volatile loss out of the sample and act to smooth out the thermal gradient. The graphite prevents the melt from touching the molybdenum; post-experiment analysis of the molybdenum capsules showed molybdenum and a thin layer of molybdenum oxide, but no molybdenum sulfide.

Experiments were run for 12 to 48 hours to attain equilibrium at 1250 to 1475 °C, typically between the liquidus and solidus to produce a silicate melt, silicate crystal, and sulfide assemblage. The homogeneity of each phase and euhedral pyroxene and olivine crystals suggest experiments obtained equilibrium. Pressure was applied using the cold piston-in technique to 1 GPa to simulate Mercurian upper mantle conditions. Run temperature was monitored and controlled to within 1 °C using a W3Re/W25Re thermocouple, and no pressure correction was applied to thermocouple EMF. A summary of experimental conditions is included in **Table 2**.

After quenching, each capsule was sliced in half with a diamond wafer blade and one half was mounted in epoxy. Samples were dry-polished using hexagonal boron nitride as a lubricant to avoid loss of any sulfides (Murthey et al. 2003) or volatiles such as S and Na in the silicate glass (Vander Kaaden & McCubbin, 2016); Vander Kaaden & McCubbin (2016) found reduced silicate glasses polished using water and acetone had 0.2-1.2 wt.% lower volatile abundances ($Na_2O$, S, and Cl) than when dry polished using hexagonal boron nitride. Additionally, Na loss is only significant in 1 atm vertical furnaces but not in piston cylinders or multi anvils. Polishing steps included 300, 600, and 1200 grit sandpaper, 1-micron diamond powder, and 0.3 and 0.05-
micron alumina polishing powder cleaning the sample in between steps with a Kimwipe and
compressed air.

## 2.3 Analytical Methods

Phase identification and quantitative chemical analyses were completed through electron
probe microanalysis (EPMA) using the CAMECA SX100 electron microprobe at Brown
University. Silicate glass phases were analyzed using a 15 keV accelerating voltage, a 10 nA
beam current, and a 5 µm spot size. Based on repeated analyses, relative errors (100 x 1σ/mean
value) for silicate glass are less than 12% for the major oxides (SiO₂, MgO, Al₂O₃, CaO, and
Na₂O). Relative error on sulfur concentration in silicate glass is less than 10% using pyrite as the
S standard. Even though sulfur in reduced silicate melt is more akin to troilite than pyrite, pyrite
is a common S standard for silicate melt analyses and has no detectible difference in sine theta
peak position (0.614681 for pyrite and 0.614686 for troilite). Additionally, no glass standard
with so much sulfur exists to independently verify the accuracy of the electron probe at such high
sulfur concentrations in glass. Synthesized sulfides were identified using a 15 keV accelerating
voltage, a 10 nA beam current, and a 1 µm spot size because the sulfides were typically very
small (<1µm). Based on repeated analyses, purity is >99% with relative errors less than 0.5%
(Anzures et al., 2020).

Sulfur speciation was analyzed using an undulator-based, hard x-ray microprobe at
beamline 13IDE of GSECARS, Argonne National Laboratory. To prevent interaction with air
during X-ray exposure, measurements were performed inside a helium environment. X-ray
fluorescence (XRF) maps with 2 µm/pixel resolution were created using the program Larch
(Newville, 2013). These XRF maps were used to guide XANES analyses of the quenched silicate
glass and sulfides. The spot size for XANES spectroscopy was 1-2 microns, with a nominal
penetration depth of 2-3 microns for S. XANES spectra were collected from 2442 to 2542 eV
with a step size of 0.25 eV near the S K-edge and 2 eV in the other energy regions (1 s
acquisition time per step). Spectra were normalized using an edge step normalization algorithm
fitting the energy region before and after the edge using a linear regression (Ravel et al., 2005) to
produce 451 different processing combinations as detailed in (Anzures et al., 2020). Standard
error for each sample spectra was then calculated by propagating the 1 σ standard deviation of
the top 10 normalizations for multiple measurements of the same sample (2 to 5 clean spectra per
sample).

Beam damage was avoided during XANES analysis by keeping the acquisition time low, analyzing multiple spots on the sample before averaging, and monitoring for the potential presence of a S<sup>4+</sup> species known to occur by X-ray irradiation (Wilke et al., 2008). While XANES analyses were completed after EPMA, surface oxidation and possible beam damage were removed by re-polishing. Recent S K-edge XANES analyses at 13IDE did not observe the potential presence (or increase with time) of S<sup>4+</sup> with 2.5 seconds per energy point (Head et al., 2018), with these analyses having an even lower collection time of 1 second per energy point. Although S<sup>4+</sup> has been observed in some terrestrial volcanic glasses (Métrich et al., 2009) and minerals (Konecke et al., 2017), no clear evidence was observed in the sulfide standards (Anzures et al., 2020) or quenched silicate glass (a few samples were tested with 5-10 repeat analyses on the same spot).

## 2.4 Calculations of Oxygen Fugacity

The intrinsic fO₂ of the experiments were calculated using Namur et al. (2016a)'s empirically derived sulfur solubility-oxygen fugacity model because although all samples were sulfide-saturated, most were not Fe metal-bearing. Normally, fO₂ is calculated using Fe³⁺/Fe²⁺
under oxidizing conditions (Kress & Carmichael, 1988; Wilke et al., 2004) and equilibrium between Fe-rich metal and FeO in the silicate melt under reducing conditions (Zolotov et al., 2013). Under very reducing conditions however, the FeO content of the silicate melt was less than 0.12 wt.% in all samples, which is close to the 0.1 wt.% detection limit of EPMA. At such low concentrations, the activity coefficient of FeO is poorly constrained (Wykes et al., 2015). The absence of a metal phase in some experiments also hampers using iron activity to calculate $fO_2$. Namur's method to calculate $fO_2$ from S content at sulfide saturation (SCSS) was recently developed for Mercurian volcanic compositions and was calibrated using Fe-FeO and Si-SiO$_2$ equilibrium (Namur, Charlier, et al., 2016). However, the low Fe concentrations in silicate melt in some experiments made it so that only $fO_2$ values calculated with Si-SiO$_2$ were used to regress the S solubility equation. The expression for S solubility is

$$ \ln[S \text{ wt. \%}]_{\text{scss}} = a + \frac{b}{T} + \frac{cP}{T} + d \log fO_2 + \sum e_i \frac{X_i}{X_{SiO_2}} $$

where S is the sulfur content at sulfide saturation (SCSS), T is the temperature (K), P is the pressure (bar), $fO_2$ is the oxygen fugacity (bar), $X_i$ are the mole fractions of oxides in the melt recalculated to a total of 100% on a sulfur-free basis, and a, b, c, d, and e are derived coefficient constants (a=7.25, b=-2.54x10$^4$, c=0.04, d=-0.551, eTiO$_2$=-62, eMgO=0.426, and eNa$_2$O=1.56) (Namur, Charlier, et al., 2016). The estimated uncertainty on $fO_2$ using this technique is 0.5 log units (Namur, Charlier, et al., 2016).

## 2.5 Modelling Methods

Quantitative S speciation data for the quenched silicate glasses were calculated by linearly unmixing the quenched silicate glass spectra into their component pure sulfide spectra. Sulfide standard spectra are taken from Anzures et al. (2020). While linear unmixing using representative endmember spectra has been used successfully to distinguish between $S^{2-}$ and $S^{6+}$
and quantitatively determine S<sup>6+</sup>/ΣS in quenched silicate glasses (Head et al., 2018; Jugo et al., 2010; Klimm, Kohn, O'Dell, et al., 2012), this study quantitatively determines sulfide speciation between different cations (e.g. FeS, MgS, and CaS) in reduced experimental quenched silicate melts. This is possible because the K-edge position of FeS, MgS and CaS are distinct from each other.

The S species endmembers used in the unmixing starting model– FeS, MgS, CaS, FeCr<sub>2</sub>S<sub>4</sub>, Na<sub>2</sub>S, MnS, NiS- were guided by sulfides present in enstatite chondrites that formed under similarly reducing conditions (Keil, 1968; Zhang et al., 1995), thermodynamic predictions (Ebel et al., 2011), and previous Raman measurements (Namur et al., 2016a). Troilite (FeS), niningerite (MgS endmember of (FeMgMnS)), and oldhamite (CaS) are all common sulfide minerals in enstatite chondrites. While FeS has been confirmed to be the dominant sulfide species in reduced terrestrial magmas (O'Neill & Mavrogenes, 2002), (Mg,Ca)S has been thought to be the dominant sulfide in extremely reduced melts (Cartier & Wood, 2019; Fogel, 2005; Namur, Charlier, et al., 2016; Stockstill-Cahill et al., 2012). Alabandite (MnS endmember of (Mn,Fe)S) has been observed in EL chondrites instead of niningerite that is found in EH chondrites. Daubreelite (FeCr<sub>2</sub>S<sub>4</sub>) is another accessory sulfide mineral common in enstatite chondrites (Keil, 1968). Na<sub>2</sub>S acts as a proxy for alkali sulfides caswellsilverite (NaCrS<sub>2</sub>) and djerfisherite (K<sub>6</sub>Na(Fe<sup>2+</sup>,Cu,Ni)<sub>25</sub>S<sub>26</sub>Cl) that are rare phases in enstatite chondrites (Keil, 1968; Zhang et al., 1995). NiS was also included in the endmember suite because it can substitute into pyrrhotite, the nonstoichiometric form of troilite. These sulfide endmembers are distinguishable by their different shape XANES spectra and S K-edges (2469.0-2472.2 eV) (Anzures et al., 2020) as shown in **Figure 5**. The unmixing fitting window (Figure 5b) was chosen to be 2463 to 2473.8 eV to encompass this spectral variability and emphasize the sulfide K-edge where sulfide
286 endmember predictive power is maximized. This focus on the K-edge region allows the use of
287 pure, solid end-member phases (e.g. MgS, CaS) to calibrate for sulfide bonds in silicate glasses,
288 as the K-edge position is not sensitive to matrix effects. Use of higher energy parts of the
289 XANES spectrum will require more matrix-matched standards that better represent the local
290 coordination environment within 5 Å of S that is probed by S K-edge XANES spectra. No
291 weighting function is used due to XANES spectra mixing linearly.

292 Quenched silicate glass spectra were linearly unmixed using all possible endmember
293 combinations and then compared using RMSE and adjusted r². We used a chemical constraint
294 where proportions were limited by their concentration in the silicate glass. If the calculated
295 proportion was larger than the maximum proportion attributed to the cation (e.g. Fe, Cr, Mn), the
296 proportion was fixed at that maximum and the unmixing was re-run. This is particularly
297 important for distinguishing FeS and Na₂S. These have similar K-edge positions, but at low fO₂,
298 concentrations of Fe in the melt are below 0.1 wt.% (Table 3). Therefore, the K-edge can be
299 confidently ascribed to Na₂S. No other initial constraints were used in order to obtain the best
300 spectral match, but instead calculated proportions were used to remove endmembers that were
301 not detected. In other words, negative proportions or proportions smaller than their uncertainty
302 were set to 0%. RMSE was used to rank results for each endmember combination using a set
303 number of endmembers. Knowing that RMSE will always decrease with additional endmembers
304 (Draper & Smith, 1998), the final unmixing combination was chosen based on the minimum
305 RMSE and highest adjusted r², which is a variant of r² that adjusts for the number of terms in the
306 model. When combinations with different numbers of endmembers are compared, there is an
307 optimal number of endmembers where adding components does not significantly improve the fit,
308 at which point additional endmembers are not meaningful (Draper & Smith, 1998) (Figure A1).
309 Sulfide speciation proportions presented below represent the optimal number of endmembers for each fit. Final sulfide speciation proportions were renormalized to sum to 100%. 310

311 Sulfide proportion uncertainties were calculated as the 2σ standard error of the linear unmixing model. The estimated covariance matrix is given by: 312

313 $$ \Sigma = \text{MSE}(X' \cdot X)^{-1} $$

314 where MSE is the mean squared error of the measured and modelled spectra, and X is the matrix 315 of observations in this case the endmembers. The 95% standard error is then calculated as the 316 diagonal of the square root of the coefficient covariance matrix. This accounts for the uncertainty 317 given an imperfect fit, which is the largest source of uncertainty. Uncertainty on the proportions 318 from propagating measured spectra uncertainty was tested by fitting 100 different randomly 319 distributed measured spectra within 1σ of intensities and only resulted in an additional error of 320 ~0.1%, which is encompassed by the 1-20% uncertainty calculated above. Typical uncertainties 321 in linear unmixing of XANES spectra are ~10% due to imperfect fitting, local coordination 322 differences between the measured quenched glass and endmember crystalline samples, and other 323 gaps in pitfalls of interpreting XANES spectra. 324

## 3. Results

### 3.1 Experimental phase assemblages

326 All experiments have a multi-component phase assemblage with greater than 20% silicate 327 glass, minor amounts of sulfide ± metal, and sometimes silicate crystals (orthopyroxene ± olivine) 328 (Supplementary Tables). The dominant sulfide identified through energy-dispersive x-ray 329 spectroscopy (EDS), when present, is FeCr₂S₄ (daubréelite), with minor amounts of Mg and Ca. 330 The dominant metal phase identified through energy-dispersive x-ray spectroscopy (EDS), when 331 present, is an FeSi alloy, with minor amounts of Ni and S. Since most of the sulfide and metal
blebs were very small (<1 µm), sulfide and metal compositions are not included. Both orthopyroxene and olivine are nearly pure (>99%) Mg-endmembers of enstatite and forsterite. Samples displayed equilibrium textures with large euhedral silicate crystals and homogenous silicate glass as revealed by XRF mapping (Figure 6).

336

<table><caption>Table 2. Table of experimental run conditions</caption><thead><tr><th>Sample</th><th>T (°C)</th><th>Si/SiO<sub>2</sub> (in wt.%)</th><th>Time (hrs)</th><th>ΔIW</th><th>Phase assemblage</th></tr></thead><tbody><tr><td>MER17<sup>a</sup></td><td>1475</td><td>0.02</td><td>24</td><td>-3</td><td>Liq, for, sulf</td></tr><tr><td>MERC2a<sup>a</sup></td><td>1420</td><td>0.05</td><td>23</td><td>-3.4</td><td>Liq, ens, for, sulf</td></tr><tr><td>PM6<sup>a</sup></td><td>1420</td><td>0.05</td><td>12</td><td>-5.5</td><td>Liq, ens, sulf, met</td></tr><tr><td>SW1<sup>a</sup></td><td>1475</td><td>0.05</td><td>48</td><td>-2.9</td><td>Liq, ens, for, sulf</td></tr><tr><td>MER15<sup>a</sup></td><td>1475</td><td>0.05</td><td>24</td><td>-7.5</td><td>Liq, ens, sulf, met</td></tr><tr><td>MER18<sup>a</sup></td><td>1475</td><td>0.1</td><td>24</td><td>-6.2</td><td>Liq, sulf, met</td></tr><tr><td>MER22<sup>a</sup></td><td>1475</td><td>0.15</td><td>24</td><td>-6.9</td><td>Liq, ens, for, sulf</td></tr><tr><td>MC3_4<sup>b</sup></td><td>1250</td><td>0.04</td><td>24</td><td>-4.5</td><td>Liq, ens, sulf</td></tr><tr><td>MC3_2<sup>b</sup></td><td>1300</td><td>0.04</td><td>24</td><td>-4.9</td><td>Liq, ens, sulf</td></tr><tr><td>MC2_2<sup>b</sup></td><td>1300</td><td>0.05</td><td>24</td><td>-2.5</td><td>Liq, sulf</td></tr><tr><td>MC1_1<sup>b</sup></td><td>1350</td><td>0.18</td><td>24</td><td>-5.7</td><td>Liq, sulf</td></tr></tbody></table>

Liq = silicate melt; ens = enstatite; for = forsterite; sulf = liquid sulfide; met = metallic alloy

<sup>a</sup>Primitive Mercurian Mantle, <sup>b</sup>Primitive Mercurian Mantle (80% crystallized)

340
341

<table><caption>**Table 3.** Average compositions (wt.%) for experimental silicate melts</caption><thead><tr><th rowspan="2"></th><th colspan="2">MER17</th><th colspan="2">MERC2a</th><th colspan="2">PM6</th><th colspan="2">SW1</th><th colspan="2">MER15</th><th colspan="2">MER18</th><th colspan="2">MER22</th><th colspan="2">MC3_4</th><th colspan="2">MC3_2</th><th colspan="2">MC2_2</th><th colspan="2">MC1_1</th></tr><tr><th>Na=9</th><th>1 σ</th><th>Na=9</th><th>1 σ</th><th>N=10</th><th>1 σ</th><th>Na=10</th><th>1 σ</th><th>Na=5</th><th>1 σ</th><th>Na=5</th><th>1 σ</th><th>Na=7</th><th>1 σ</th><th>Na=10</th><th>1 σ</th><th>Na=10</th><th>1 σ</th><th>Na=10</th><th>1 σ</th><th>Na=10</th><th>1 σ</th></tr></thead><tbody><tr><td>Si</td><td>27.62</td><td>0.27</td><td>30.04</td><td>0.08</td><td>24.53</td><td>0.30</td><td>24.78</td><td>0.08</td><td>25.03</td><td>0.32</td><td>24.91</td><td>0.19</td><td>24.70</td><td>0.14</td><td>25.16</td><td>0.10</td><td>25.13</td><td>0.20</td><td>24.57</td><td>0.09</td><td>26.86</td><td>0.61</td></tr><tr><td>Ti</td><td>0.47</td><td>0.03</td><td>0.23</td><td>0.02</td><td>0.20</td><td>0.02</td><td>0.48</td><td>0.03</td><td>0.26</td><td>0.05</td><td>0.17</td><td>0.02</td><td>0.24</td><td>0.01</td><td>0.82</td><td>0.03</td><td>0.82</td><td>0.04</td><td>0.81</td><td>0.03</td><td>0.09</td><td>0.02</td></tr><tr><td>Al</td><td>7.12</td><td>0.29</td><td>6.50</td><td>0.07</td><td>8.38</td><td>0.10</td><td>7.89</td><td>0.07</td><td>5.65</td><td>0.13</td><td>4.14</td><td>0.19</td><td>5.49</td><td>0.21</td><td>9.02</td><td>0.04</td><td>9.28</td><td>0.20</td><td>8.62</td><td>0.07</td><td>9.26</td><td>0.17</td></tr><tr><td>Cr</td><td>0.04</td><td>0.02</td><td>0.02</td><td>0.01</td><td>0.07</td><td>0.01</td><td>0.05</td><td>0.01</td><td>0.01</td><td>0.01</td><td>0.01</td><td>0.01</td><td>0.03</td><td>0.01</td><td>0.04</td><td>0.01</td><td>0.04</td><td>0.01</td><td>0.06</td><td>0.01</td><td>bdl</td><td>bdl</td></tr><tr><td>Fe</td><td>0.09</td><td>0.04</td><td>0.06</td><td>0.02</td><td>0.04</td><td>0.02</td><td>0.05</td><td>0.01</td><td>bdl</td><td>bdl</td><td>0.06</td><td>0.01</td><td>0.02</td><td>0.01</td><td>0.04</td><td>0.02</td><td>0.02</td><td>0.02</td><td>0.04</td><td>0.02</td><td>0.01</td><td>0.01</td></tr><tr><td>Mn</td><td>0.08</td><td>0.03</td><td>0.03</td><td>0.01</td><td>0.15</td><td>0.01</td><td>0.03</td><td>0.02</td><td>0.02</td><td>0.01</td><td>0.04</td><td>0.01</td><td>0.06</td><td>0.03</td><td>0.10</td><td>0.02</td><td>0.10</td><td>0.02</td><td>0.11</td><td>0.01</td><td>0.04</td><td>0.01</td></tr><tr><td>Mg</td><td>6.95</td><td>0.80</td><td>5.40</td><td>0.15</td><td>9.94</td><td>0.12</td><td>9.74</td><td>0.04</td><td>14.80</td><td>0.44</td><td>18.28</td><td>0.85</td><td>14.82</td><td>0.57</td><td>7.63</td><td>0.05</td><td>7.39</td><td>0.16</td><td>8.40</td><td>0.09</td><td>9.19</td><td>0.09</td></tr><tr><td>Ca</td><td>7.56</td><td>0.27</td><td>5.90</td><td>0.07</td><td>7.17</td><td>0.09</td><td>7.40</td><td>0.07</td><td>5.90</td><td>0.13</td><td>4.59</td><td>0.30</td><td>5.41</td><td>0.26</td><td>7.82</td><td>0.06</td><td>8.13</td><td>0.10</td><td>7.44</td><td>0.07</td><td>7.54</td><td>0.21</td></tr><tr><td>Na</td><td>1.80</td><td>0.15</td><td>3.43</td><td>0.09</td><td>1.21</td><td>0.06</td><td>2.15</td><td>0.08</td><td>0.31</td><td>0.05</td><td>0.44</td><td>0.04</td><td>0.76</td><td>0.09</td><td>1.68</td><td>0.07</td><td>1.51</td><td>0.08</td><td>1.50</td><td>0.08</td><td>0.36</td><td>0.03</td></tr><tr><td>K</td><td>0.07</td><td>0.01</td><td>0.17</td><td>0.01</td><td>0.58</td><td>0.01</td><td>0.07</td><td>0.01</td><td>0.02</td><td>0.01</td><td>0.02</td><td>0.01</td><td>0.02</td><td>0.01</td><td>0.31</td><td>0.01</td><td>0.30</td><td>0.01</td><td>0.29</td><td>0.01</td><td>0.15</td><td>0.01</td></tr><tr><td>Ni</td><td>0.01</td><td>0.01</td><td>0.01</td><td>0.01</td><td>bdl</td><td>bdl</td><td>0.01</td><td>0.01</td><td>0.01</td><td>0.01</td><td>0.01</td><td>bdl</td><td>bdl</td><td>0.01</td><td>0.01</td><td>0.01</td><td>0.01</td><td>0.01</td><td>0.01</td><td>0.01</td><td>0.02</td><td>0.02</td></tr><tr><td>S</td><td>1.08</td><td>0.08</td><td>1.31</td><td>0.06</td><td>3.31</td><td>0.16</td><td>1.01</td><td>0.03</td><td>7.56</td><td>0.72</td><td>4.68</td><td>0.12</td><td>6.68</td><td>0.32</td><td>2.01</td><td>0.05</td><td>2.67</td><td>0.08</td><td>0.88</td><td>0.06</td><td>3.90</td><td>0.17</td></tr><tr><td>O*</td><td>45.88</td><td>1.26</td><td>46.69</td><td>0.30</td><td>43.92</td><td>0.51</td><td>45.24</td><td>0.25</td><td>42.17</td><td>0.52</td><td>43.90</td><td>1.05</td><td>42.07</td><td>0.72</td><td>45.08</td><td>0.26</td><td>44.85</td><td>0.58</td><td>44.91</td><td>0.29</td><td>46.19</td><td>0.94</td></tr><tr><td>Total</td><td>98.77</td><td>3.26</td><td>99.79</td><td>0.89</td><td>99.50</td><td>1.41</td><td>98.90</td><td>0.69</td><td>101.74</td><td>2.39</td><td>101.25</td><td>2.80</td><td>100.3</td><td>2.37</td><td>99.72</td><td>0.73</td><td>100.25</td><td>1.50</td><td>97.64</td><td>0.85</td><td>103.61</td><td>2.30</td></tr></tbody></table>

342

Na= number of analyses; bdl = below detection limit

343

*Oxygen calculated by stoichiometry
344 **3.2 Sulfur speciation in silicate melts**

Qualitative analysis of the XANES spectra of experimental quenched glasses reveal
several key trends as fO₂ decreases including: 1) a diminishing low energy shoulder at 2470 eV,
2) flattening of the near-edge structure from 2473 to 2477 eV, and 3) a positive K-edge shift
(Figure 7). Although it is well-established that the sharp S²⁻ peak present in crystalline sulfides is
absent in broadened silicate glass spectra due to additional scattering and slightly different
coordination in multicomponent glasses (Fleet, 2005; Jugo et al., 2010; Métrich, 2002; Paris et
al., 2001), the energy location of the initial K-edge determined by s-p orbital energy differences
is more robust. Thus, these spectral changes correspond to a decrease in the lower energy
sulfides (FeS, FeCr₂S₄, and Na₂S) responsible for the shoulder at 2470 eV and an increase in
higher energy sulfides (CaS and MgS) as seen in Figure 7a and 7b. As fO₂ decreases, the
spectra become more similar to pure MgS that has a flat near-edge structure, whereas CaS and
MnS that both have positive sloped near-edge structures are only obvious under less reducing
conditions. MnS has an intermediate K-edge at 2470.6 eV that can be observed in a small
additional shoulder in a few samples.

These qualitative spectral shifts were quantified into sulfide species proportions using
linear unmixing as shown in **Figure 8 and S2**. Modelled spectra have component sums between
98% and 122% with RMSE below 0.05 as well as ordinary and adjusted r² values greater than
0.98 (supplementary tables). After normalizing the percentages to 100%, normalized spectra
have RMSE below 0.11 as well as ordinary and adjusted r² values greater than 0.95. Since fitting
was focused on the K-edge in the XANES region of the spectrum, deviations are larger at higher
energies in the near-edge structure and XAFS region. Less reduced silicate glasses (IW-2 to IW-
4) had more complicated XANES spectra, and thus higher RMSE and residual deviations as seen
in **Figure 8a**. However, we have high confidence that CaS is the dominant sulfide species in less reduced silicate glasses because CaS is the primary end member in all fits (**Table 4** and **Figure A1**).

<table><caption>**Table 4.** Fit statistics for different endmember combination of SW1 (IW-2.9, 1.01 wt.% S).</caption><thead><tr><th></th><th>Endmember<br>combination</th><th>RMSE of fitted<br>region</th><th>R<sup>2</sup><br>(ordinary)</th><th>R<sup>2</sup><br>(adjusted)</th></tr></thead><tbody><tr><td>One endmember</td><td>MgS</td><td>0.20</td><td>0.77</td><td>0.77</td></tr><tr><td>One endmember</td><td>CaS</td><td>0.14</td><td>0.87</td><td>0.87</td></tr><tr><td>Two<br>endmembers</td><td>CaS, Na<sub>2</sub>S</td><td>0.05</td><td>0.98</td><td>0.98</td></tr><tr><td>Three<br>endmembers</td><td>CaS, Na<sub>2</sub>S, FeS</td><td>0.05</td><td>0.98</td><td>0.98</td></tr></tbody></table>

At high $fO_2$, from IW to ~IW-2, S solubilities are below 0.5 wt% and the S speciation is predominantly FeS (Lodders & Fegley, 1997; Saal et al., 2008; P. Wallace & Carmichael, 1992). Around IW-2 to IW-4.5 (S=0.5-2.0 wt%), FeS and $FeCr_2S_4$ disappears from the spectra and CaS (40-65%) becomes the dominant S species, with substantial amounts of $Na_2S$ (~35%) and minor MnS (~7%) (**Figure 9**). At ~IW-4.5, MgS becomes the dominant species, rising to ~85% and remaining at that level to the lowest $fO_2$ we examined (IW-7.2), even while S concentration increases from 2.5 wt% to 7.5 wt%. CaS decreases from 65% to 15%, and $Na_2S$ disappears from the spectra (**Figure 9**). At IW-4.9 and below, MgS and CaS proportions do not significantly change from ($Mg_{0.85}Ca_{0.15}S$), which is remarkably similar to an estimated sulfide composition at the surface of Mercury ($Mg_{0.74}Ca_{0.26}S$ (Nittler et al., 2011; Stockstill-Cahill et al., 2012; Weider et al., 2012)) and sulfide composition in EH Indarch melting experiments ($Mg_{0.79}Ca_{0.21}S$ (McCoy et al., 1999)).

These trends are robust over the compositional range (Si = 24.53-30.04 wt.%, Fe = 0-0.09 wt.%, Mg = 5.4-18.28 wt.%, Ca = 4.59-8.13 wt.%, Na = 0.31-3.43 wt.%), and they show a slight
correlation with Fe, Mg, Ca, and Na, while the Cr trend is unclear due to low concentrations.

FeS, MgS, and CaS trends are consistent with previous Raman measurements (Namur et al., 2016a). The presence of $Cr^{3+}$ and $Na^+$ are consistent with precipitation of solid sulfides stable under very reducing conditions $FeCr_2S_4$ (daubréelite ($FeCr_2S_4$); caswellsilverite ($NaCrS_2$) and djerfisherite ($K_6Na(Fe^{2+},Cu,Ni)_{25}S_{26}Cl$)) as well as predictions of Na as a charge-balancing cation (Métrich et al., 2009).

## 4. Discussion

### 4.1 Potential effects of local coordination environment

It is possible to distinguish between sulfur species bonded to divalent cations (e.g. $Fe^{2+}$, $Mg^{2+}$, and $Ca^{2+}$) even though the sulfur ion is nominally present as $S^{2-}$. In contrast, distinguishing sulfur species bonded in sulfates is quite difficult, even though the sulfur ion is nominally present as $S^{6+}$. As an example, The S K-edge energies vary more in sulfides (from 2469.0 to 2472.2 eV) compared to those for sulfates (2480.5-2481.3 eV) (Anzures et al., 2020; Fleet, 2005). Note that determining gross oxidation states (i.e. the S K-edge difference between $S^{2-}$ and $S^{6+}$) is not difficult and possible using EPMA (M. Carroll & Rutherford, 1988; Metrich & Clocchiatti, 1996). Sulfide S K-edge energies are known to increase with increasingly ionic bonding from covalent/metallic FeS to more ionic MgS and CaS as covalent sharing of electrons decreases the energy differences in the innermost electron shells of bonded ions (Anzures et al., 2020; Fleet, 2005). Sulfides also show differences in near-edge peaks indicative of a wider range in overall S K-edge XANES structure than sulfates due to differences in the *p*-d hybridization of the *1s* to metal *3d* states (Fleet, 2005). The smaller variation in sulfates may be explained by the sulfur ion being surrounded by 4 oxygens to form an anionic complex shielding the sulfur ion from cation effects. The fact that all sulfates exhibit 100% ionic character of the S-O bond as calculated
409 using Pauling (1960) reinforces the idea that the local bonding environment of the sulfate ion is 410 dominated by the greater electronegativity of the oxygen. Presumably, this stereochemical 411 shielding by oxygen also hinders using X-ray absorption spectroscopy to differentiate between 412 different cation species for elements such as C and P in compounds like carbonates and 413 phosphates but could be used to differentiate between carbides and phosphides.

414 One of the challenges in determining sulfur speciation in silicate melt are the inherent 415 differences in bonding of S in a crystal (our endmembers), quenched glass (our samples), and 416 melt (our interest). These differences in the local coordination environment of S (within 5 Å of 417 the S atoms) between the quenched silicate glass, crystalline standards, and silicate melt at run 418 conditions may have resulted in systematic errors in the S speciation not accounted for in 419 calculated uncertainties. The fitting window choice to focus on the K-edge avoids most of higher 420 energy near-edge and extended fine structure features that are more affected by the local 421 coordination environment scattering (Als-Nielsen & McMorrow, 2011; S. R. Sutton et al., 2020); 422 S K-edge XANES spectra of complex silicate glasses have been observed to diminish as well as 423 shift and broaden slightly to higher energies compared with crystalline sulfides (Fleet, 2005; 424 Head et al., 2018; Jugo et al., 2010). In the future, matrix-matched glass endmembers (Anzures 425 et al., 2020) and XANES analysis at run conditions (Berry et al., 2003) may improve fitting of 426 unknown silicate glass spectra. Nevertheless, the observed systematic changes in the S K-edge 427 XANES spectra (Figure 7b) strongly support that changes in XANES features are due to 428 different S species in the quenched glass and are representative of S speciation at run conditions.

429 Since sulfide speciation was not determined in the melt at run conditions, it is possible 430 that there could be some quench modification of valence state due to electron exchange (Berry et 431 al., 2003; Head et al., 2018). This electron exchange has been used to explain the coexistence of
S²⁻ and S⁶⁺ in reduced melt inclusions quenched below IW-4 where S should only be dissolved as S²⁻ (Métrich et al., 2009). This quench modification could occur between S²⁻ and a transition element like iron, chromium, or titanium, which would explain the small S⁶⁺ peaks at ~2482 eV that persisted in a few samples even after re-polishing (Figure 7b). These sulfur bearing reactions are unlikely however, because they require large structural changes in the local coordination environment between S²⁻ and SO₄²⁻ during quench and are likely proportional to the relative abundances of the transition element electron acceptors (Head et al., 2018; Métrich et al., 2009). Quenched glasses have low transition element contents at <0.1 wt.% Fe, <0.08 wt.% Cr, and <0.83 wt.% Ti. Additionally, this peak at ~2482 eV is unlikely to be related to H₂ (Klimm, Kohn, & Botcharnikov, 2012) due to extremely low hydrogen solubility at very reducing conditions (Zolotov, 2011).

## 4.2 Effect of S on Fe, Mg, Ca, and Si activities and mineral stability

The sulfide speciation trends as a function of oxygen fugacity have important ramifications for Mercurian magmas and other reduced rocky bodies. The bonding of S with Fe, Mg, Ca, Na, Mn, and Cr reduces the concentrations of the cations available to form oxides affecting the silicate phase equilibria. The effective concentrations of MgO, CaO, and Na₂O are reduced by up to 30%, 35%, and 39% respectively, while FeO, MnO, and Cr₂O₃ are reduced by effectively 100% (Figure 10). Manganese and chromium sulfide species detections are minor (<10%), and consistent with the presence of similar sulfides in enstatite chondrites and aubrites (Keil, 1968; Zhang et al., 1995).

Under less reducing conditions (above IW-1 but below sulfate stability ~IW+2), it is well known that S dissolves in terrestrial, lunar, and Martian magmas as mostly FeS (Lodders & Fegley, 1997; Saal et al., 2008; P. Wallace & Carmichael, 1992). Applying the S speciation
trends to Mercurian melt chemistry, any iron ($X_{Fe}<0.11$ wt.%) in the silicate melt exists as FeS rather than FeO at these reducing conditions (below IW-2) consistent with the VNIR data of Mercury's surface showing no FeO related absorptions (Izenberg et al., 2014; Riner et al., 2010).
Thus, models of $fO_2$ based on FeO in the melt may be in error, as most Fe will be present as FeS.
However, the Fe content measured by MESSENGER on the Mercurian surface of 1-2 wt.% (Nittler et al., 2011) is much higher than would be dissolved in the silicate melt at these reducing conditions. With low stability of Fe in the silicate melt, Fe would exsolve as FeSi, FeS, and/or $FeCr_2S_4$. If these iron sulfides are entrained and physically carried to the surface within the silicate melt, this could account for up to 1.15 wt.% or 4.60 wt.% S if pure FeS and $FeCrS_4$ respectively. These surface sulfides would dramatically reduce the assumed elevated S solubility in Mercurian lavas (1-4 wt.% S measured on the surface) (Nittler et al., 2011) and could favor a Mercurian mantle source on the oxidized side of the current $fO_2$ bounds (IW-3 to IW-7) (McCubbin et al., 2012; Zolotov et al., 2013). However, it is currently very difficult to place constraints on the relative abundances of $Fe^0$ versus FeS without knowing Mercury's bulk composition as well as how much of the surface lavas are silicate crystals rather than quenched silicate glass without higher resolution chemical and mineralogical data (Namur & Charlier, 2017).

The S speciation data shows that MgS is the dominant S species at $fO_2$ below IW-4, as suggested by Namur et al (2016a) from Raman spectroscopy. This is consistent with predictions for extremely reduced (below IW-4) and S-rich silicate melts relevant for Mercury, enstatite chondrites, and aubrites that the dominant melt species is $(Mg_{75}Ca_{25})S$ (Fogel, 2005; Namur, Charlier, et al., 2016; Stockstill-Cahill et al., 2012). The substantial bonding of Mg with S at highly reducing conditions decreases MgO concentrations in the melt by an average of 22.5%
relative. This is consistent with the destabilization of Mg-rich olivine in favor of enstatite. If Mercurian lavas have an intrinsic fO₂ below IW-4 as suggested by Namur et al. (2016a), (Mg,Ca)S should be the dominant sulfide in Mercury's crust.

At intermediate fO₂ conditions (IW-2 to IW-4), CaS is the major S species, and the concentration of CaO is decreased by 8 to 15%. At very reducing conditions (<IW-4), CaS is a minor component, and CaO concentrations decreases by up to 35%. This functions to decrease the CaO concentration by up to 35% and an average of 13 wt.%, which may destabilize Ca-bearing silicates like feldspar and clinopyroxene as shown in the reaction below.

$$CaO + Al_2O_3 + 2SiO_2 \leftarrow CaAl_2Si_2O_8 \text{ (anorthite)}$$

If Mercurian lavas have an intrinsic mantle source fO₂ between IW-2 and IW-4 as suggested by McCubbin et al. (2017) (IW-3.2 to IW-4.3), CaS should be the dominant sulfide in Mercury's crust, with substantial Na₂S.

Up to 16% Na-bearing sulfide species is detected in the intermediate fO₂ range, causing a decrease in Na₂O concentration of up to 39% and an average of 28% when present. If Mercury lavas have fO₂ between IW-2 and IW-4, they should precipitate substantial amounts of Na-bearing sulfide. This may be related to the high-Na lavas detected on the surface of Mercury (Nittler et al., 2011). Melts rich in Na₂S should precipitate caswellsilverite (NaCrS₂) and djerfisherite (K₆Na(Fe²⁺,Cu,Ni)₂₅S₂₆Cl) found in enstatite chondrites.

Another interesting yet poorly understood geologic feature on Mercury are hollows—shallow, flat-floored, steep-sided, rimless depressions generally found in or around much smaller impact craters (Thomas et al., 2014). Hollows-forming-material appear to form through sublimation of a solid layer of rock in response to the intense solar wind Mercury experiences (Blewett et al., 2011, 2016). Hollows-forming-material must be moderately volatile on the
surface, although the specific material (maybe sulfide or graphite (Blewett et al., 2016)) is still
unclear. These alkali-sulfides (caswellsilverite (NaCrS₂) and djerfisherite
(K₆Na(Fe²⁺,Cu,Ni)₂₅S₂₆Cl)) may be candidate hollows-forming material since they have
significantly lower melting temperatures than MgS (2000°C at 1atm) and CaS (2525°C at 1atm).
Alkali-sulfides may consolidate and float to the surface during flash-melting and low-pressure
cooling within shallow impact melt sheets.

Sulfide complexing also likely increases the activity of SiO₂ as sulfide complexes
"remove" network-modifying cations from oxides that would otherwise be available for silicate
phase equilibria. This would drive the reaction from forsterite to enstatite following the reaction
below.

$$511 \qquad Mg_2SiO_4 \text{ (forsterite)} + SiO_2 \rightarrow 2MgSiO_3 \text{ (enstatite)}$$

This is consistent with the expansion of the enstatite stability field at the expense of forsterite
found in (Namur, Charlier, et al., 2016; Namur, Collinet, et al., 2016). The decrease in MgO
concentration may also drive the reaction toward enstatite as the more Mg-poor phase.

## 4.3 Constraining the oxygen fugacity of Mercury

The oxygen fugacity of Mercury has been estimated in a variety of ways. The high S and
low Fe contents from XRS and GRS (Evans et al., 2012; Nittler et al., 2011) imply fO₂ of IW-3
to IW-7 (McCubbin et al., 2012; Namur, Charlier, et al., 2016; Zolotov et al., 2013), which is
consistent with the absence of Fe³⁺ charge transfer absorptions that indicate low-Fe silicate
minerals from VNIR (Izenberg et al., 2014; Riner et al., 2010). Sulfur speciation may also be a
useful oxybarometer. S speciation can be described by the following equilibrium reactions with
components after the polymer model of silicate melts (Hess, 1995) :

$$523 \qquad 2SiOMgOSi + 1/2S_2 = MgSMg + 2SiOSi + O_2$$
524

$$K_{eq} = \frac{a_{MgSMg} a_{SiOSi}^2 f_{O2}}{a_{SiOMgOSi}^2 f_{S2}}$$

525

$$2\text{SiOCaOSi} + 1/2\text{S}_2 = \text{CaSCa} + 2\text{SiOSi} + \text{O}_2$$

526

$$K_{eq} = \frac{a_{CaSCa} a_{SiOSi}^2 f_{O2}}{a_{SiOcaOSi}^2 f_{S2}}$$

527

where the equilibrium between MgO, MgS, CaO, CaS, and SiO₂ are a function of both fO₂ and

528

fS₂. This is a potential oxybarometer at very reducing conditions. However, additional

529

measurements and comparison with calculations with equilibrium of other S species such as CaS

530

are necessary to determine what S species and reaction controls fO₂.

531

Interestingly, the conservative estimate of fO₂ of Mercury's mantle between IW-3 to IW-

532

7 (Zolotov et al., 2013) includes both the CaS (IW-2 to IW-4) and MgS (<IW-4) sulfur

533

speciation regimes. The fO₂ region where CaS dominates and MgS is absent, would be

534

consistent with the observed strong correlation between Ca and S, and the weaker correlation of

535

Mg and S on the surface of Mercury (Nittler & Weider, 2019; Weider et al., 2012). An fO₂ range

536

of IW-2 and IW-4 would also be consistent with a less reduced (IW-3.2 to IW-4.3) mantle

537

predicted to account for the low O/Si ratio measured on Mercury (McCubbin et al., 2012, 2017).

538

Thus, the fO₂ of Mercury's mantle may be between IW-2 and IW-4 rather than IW-3 to IW-7

539

(Zolotov et al., 2013). However, a less reduced Mercurian mantle may not be able to dissolve

540

enough S (Namur, Charlier, et al., 2016) to match the surface measurements of 1-4 wt.% S

541

(Nittler et al., 2011). Unfortunately, the large spatial resolution of XRS data (50-1000km used

542

for mapping) (Nittler et al., 2011; Weider et al., 2015) complicates attempts to fingerprint S

543

speciation from orbit; we must assume chemical homogeneity within the measurement footprint

544

and that S chemical correlations are due primarily to differences in sulfide content and not to

545

different silicate crystal fraction (Mg-, Ca-, and Na-bearing olivine, pyroxene, and feldspar).
546

## 4.4 Implications for Mercury's magma ocean

547

While Mercury is similar in size to the Moon, differences in $fO_2$ caused different magma ocean evolutions. With a mantle $fO_2$ between IW and IW-2 (Kevin Righter et al., 2016; Wadhwa, 2008), the Moon had an anorthosite flotation crust due to the very FeO-rich nature of the residual melt at the timing of plagioclase saturation (Elkins-Tanton et al., 2011; Jolliff et al., 2000). However, the low-FeO content of Mercurian melts at more reducing conditions (IW-3 to IW-7 (Zolotov et al., 2013)) results in only graphite (Vander Kaaden & McCubbin, 2015) or buoyant sulfides (Boukaré et al., 2019) as a potential minerals comprising a flotation crust. S speciation shows that Fe solubility at very reducing conditions (<0.15 wt.% Fe at IW-4 to IW-7) is controlled by FeS such that any Fe that dissolves in the silicate melt is bonded with S.

548

549

550

551

552

553

554

555

556

Most of the compositional fractionation of the Mercurian magma ocean will be caused by the crystallization of silicates (olivine, orthopyroxene, clinopyroxene, feldspar). By changing the stabilities of these silicate components in the melt, it will have a strong influence on the compositional evolution of the Mercurian magma ocean. This is similar to how H$_2$O strongly affects magma fractionation on Earth, despite it having low concentrations in most crystallizing phases (olivine, pigeonite, augite, and plagioclase) (Bell & Rossman, 1992; Hirschmann et al., 2005; Hirth & Kohlstedt, 1996; Saal et al., 2002).

557

558

559

560

561

562

563

Changes in sulfur speciation influence the initial silicate layering through silicate phase equilibria, setting the stage for the subsequent thermal and magmatic evolution of Mercury. Plagioclase crystallization would be delayed with the reduction of CaO by up to 35%, driving the reaction away from anorthite, the Ca and Al endmember of plagioclase, and becoming more sodic in composition as seen in enstatite chondrites, aubrites, and aubrite basalt vitrophyres (Fogel, 2005). Consequently, the primitive Mercurian mantle might be not be as depleted in

564

565

566

567

568
aluminum as the Moon because of the lesser amount of plagioclase. The initial mineralogical
layering of Mercury should also be enriched in orthopyroxene at the expense of olivine because
of S reducing MgO and CaO and increasing SiO₂ activities, the same process that would stabilize
orthopyroxene in enstatite chondrites and aubrites. Reduced CaO would also delay
clinopyroxene crystallization.

## 4.5 Effect on physical properties

Changing S speciation should also affect the physical properties of the melts. Sulfide
complexing increases the activity of SiO₂, which should increase polymerization and viscosity
(Fogel, 2005; Namur et al., 2016). However, S may preferentially react with bridging oxygen
atoms in the melt, which would produce the opposite effect (Holzheid & Grove, 2002). Si-S
bonds are not predicted to be abundant because stable silicon sulfide precipitates have not been
observed in experiments at very reducing conditions implying such phases are unstable
(MacHacek et al., 2010), the S XANES spectra are fit well without an additional Si-S
endmember (although silicon sulfide XANES spectra have not been collected), and the
mineralogical structure of SiO₂ and SiS₂ are dramatically different (Zwijnenburg et al., 2007).
The effect of sulfide complexes and perhaps other redox changes at low fO₂ in the silicate melt
is not well understood, especially at high S solubility; viscosity measurements may show a
negligible effect of S (Morizet et al., 2015; Vetere et al., 2017). Melt density should also be
affected by overall S concentration and what S species form in the melt: (density Na₂S < CaS,
MgS < FeS). Changing viscosity and density will have substantial influences on both magma
eruption dynamics, as well as the evolution of a Mercurian magma ocean.

## 5. Conclusions
In very reduced basaltic melts, FeS is destabilized in the melt structure. CaS and Na₂S are the primary S species from IW-2 to IW-4. MgS becomes the dominant species at oxygen fugacities below IW-4. The shifts of S speciation have substantial effects on the activities of oxide components of the melt and act to increase enstatite stability and decrease plagioclase stability. Enstatite should be the primary mafic silicate mineral in the Mercurian mantle.
Depending of fO₂, (Mg,Ca)S or CaS may be the dominant sulfide in Mercury's volcanic crust.
Sulfur speciation also affects physical properties (viscosity, density, melting temperature), with implications for magma ocean dynamics and volcanism.

Future experiments are needed to constrain the initial upper mantle minerals and proportions, especially what phases are major hosts of Ca and Al in the mantle. Experimental determination of the Ca-bearing phase, sulfide or silicate, is significant because they are the main reservoirs for a range of minor and trace elements, and most importantly, for the main heat-producing elements U, Th, and K (Boujibar et al., 2019; Boukaré et al., 2019). On the Moon Ca-bearing plagioclase concentrates the heat-producing elements while in enstatite chondrites CaS (oldhamite) does. The increased stability of (Mg,Ca)S in the melt structure suggests it may be the dominant sulfide and heat-producing element reservoir in Mercury's crust. More experiments are needed in the CaS dominated regime between IW-2 to IW-4, which would place further constraints on the fO₂ of Mercury's mantle.

## Acknowledgements

The authors thank Antonio Lanzirotti and Matthew Newville who aided in collection of the XANES spectra. The authors also thank Joseph T. Cukjati for helpful discussion on local coordination environment of sulfides and sulfates. We thank the three anonymous reviewers, associate editor Raj Dasgupta, and executive editor Jeffrey G. Catalano for helping improve this
614 manuscript. We acknowledge the experimental and analytical facilities at Brown University and
615 GeoSoilEnviroCars (Sector 13), Advanced Photon Source, Argonne National Laboratory. This
616 research used resources of the Advanced Photon Source, a U.S. Department of Energy (DOE)
617 Office of Science User Facility operated for the DOE Office of Science by Argonne National
618 Laboratory under Contract No. DE-AC02-06CH11357. We acknowledge the funding of NASA
619 grant No. NNX15AH63G and NASA Earth and Space Sciences fellowship No.
620 80NSSC18K1245.

621
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<figure><img src="image_2.png" /><figcaption>**Figure 1.** Terrestrial planets and meteorites cover a range of fO₂ relative to IW (Fe-FeO buffer) (Righter et al. (2006), Wadhwa (2008)). The fO₂ for typical magmas on Earth is denoted by a gray circle with black outline. Low fO₂ is observed for Mercury (McCubbin et al. 2012, Zolotov et al. 2013, Namur et al., 2016a), enstatite chondrites (e.g. Berthet et al. (2009)), aubrites (e.g. McCoy et al. (1997)), and early Earth (Javoy et al. (2010)).</figcaption></figure>
<figure><img src="image_3.png" /><figcaption>**Figure 2.** S content at sulfide saturation (SCSS) in silicate melts as a function of oxygen fugacity. The fO₂ is given relative to the iron-wustite (IW) oxygen buffer. As oxygen fugacity decreases, S solubility increases dramatically in silicate melts. The range of S concentrations for the Mercurian surface (blue field) from MESSENGER XRS (Nittler et al. (2011)) and the implied range of possible fO₂ for Mercurian lavas (green field). The spread in the experimental data represents differences in P, T, and composition among the studies (McCoy et al. (1999), Holzheid & Grove (2002), Malavergne et al. (2007), Berthet et al. (2009), Malavergne et al. (2014), Cartier et al. (2014), Namur et al. (2016)).</figcaption></figure>
16

<figure><img src="image_4.png" /><figcaption>**Figure 3.** Plot of Al/Si vs. Mg/Si showing the chemical fractionation lines for Earth (Drake & Righter (2002)), Mars (Dreibus et al. (1998)), and Mercury to determine the initial bulk composition of a terrestrial planet. Mercury's olivine+pyroxene fractionation line is a fit through the different geochemical terrains on the surface (Weider et al. (2015)) excluding the anomalous low Mg NVP region that has a different slope. In Al/Si vs. Mg/Si space, chondrites form a linear trend, which represents the chemically primitive building blocks of the solar system (SR Taylor (2001)). The bulk composition of the planet is then determined as the intersection between the fractionation line (higher Mg/Si ratio identifies more primitive magmatic compositions) and the chondrite line. For Earth, this primitive upper mantle composition agrees well with estimates based on primitive ultramafic rocks and the slope of the line through ultramafic magmas indicates primarily olivine-fractionation (Jagoutz et al., 1979; Ringwood, 1979; Palme and Nickel, 1986). The lower slope of the ultramafic lavas on Mercury suggest mostly orthopyroxene fractionation, with minor olivine (intersection with y-axis gives some indication of proportions).</figcaption></figure>

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CH ALH85085 is used in this study as a primitive Mercurian bulk composition because of its
similar Al/Si and Mg/Si ratios to the predicted bulk chemistry, high metal content and density
consistent with Mercury's internal structure with a large core, and low fO₂ (IW-5 from Petaev et
al., 2003).
Cell Assembly (3/4" PC)

<figure><img src="image_5.png" /><figcaption>**Figure 4.** ¾" Piston Cylinder cell assembly. 2mm tall x 2mm diameter sample powder is enclosed inside a graphite capsule inside a molybdenum capsule 7.2mm tall x 6.4 mm diameter.</figcaption></figure>

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<figure><img src="image_6.png" /><figcaption>**Figure 5.** a) Spectral variability as a function of energy for sulfide minerals (MgS, CaS, FeS, FeCr₂S₄, Na₂S, and MnS), sulfate minerals (CaSO₄, MgSO₄, FeSO₄, Fe₂(SO₄)₃, and reduced</figcaption></figure>
silicate glass (reduced CH composition). The green fitting window from 2463 to 2473.8 eV emphasizes the sulfide K-edges (2469.0-2472.2 eV) where sulfide endmember predictive power is maximized. The variability in silicate glass spectra is broader than for crystalline sulfides, which is consistent with multiple scattering of other components in silicate glass (Fleet, 2005). b)

Sulfur K-edge XANES spectra of endmember sulfides compared to one intermediate silicate glass (MC3_2: IW-4.9, 2.67 wt.% S) within the fitting window. A spectral shoulder at ~2469 eV is consistent with Fe-, Na-, and Cr-bearing sulfides, while the dominant feature is a spectral slope at ~2472 eV similar to MgS.
<figure><img src="image_7.png" /><figcaption>**Figure 6.** X-ray Fluorescence RGB (R=Si, G=S, B=Fe) image of an experimental charge showing equilibrium texture. Euhedral, homogeneous silicate crystals of forsterite and enstatite are in red, separate sulfide phase of daubréelite ($FeCr_2S_4$) in blue, and silicate glass in yellow-green. The spot size for XANES spectroscopy analyses is 1-2 microns with a penetration depth of 2-3 microns for S.</figcaption></figure>
<figure><img src="image_8.png" /><figcaption>**Figure 7.** a) Pure endmember standards used in this paper (Anzures et al., 2020) and b)</figcaption></figure>

experimental glasses. The silicate glass spectra display a systematic change in shape with the sulfur content (proxy for fO₂) with the spectral shoulder at ~2470 eV disappearing and the near-edge structure flattening. These spectral changes are due to changes in sulfur speciation. Inset brown box shows the complete spectra from 2440 to 2540 eV with the black lines showing expanded region. Vertical colored lines show the K-edge locations of the MgS and FeS species (thin black lines in figure 7a). The spectra were collected in total fluorescence yield.

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<figure><img src="image_9.png" /><figcaption>**Figure 8.** Linear unmixing results for representative silicate glasses that span a range in fO₂ showing the full measured spectra, modelled spectra, and residual within the green fitting window.</figcaption></figure>

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<figure><img src="image_10.png" /><figcaption>**Figure 9.** Sulfide speciation in experimental glasses as a function of S content and fO₂ from linear unmixing of XANES spectra. The proportions of FeS, FeCr₂S₄, MnS, and Na₂S species decrease and MgS increases with more reducing conditions. CaS reaches a peak at IW-3 and then becomes the minor species. The regions where FeS, CaS & Na₂S, and MgS & CaS are the dominant S species are highlighted in yellow, blue, and orange, respectively. Oxygen fugacity for the melts was calculated using the fO₂-sulfur solubility model of Namur et al. (2016). Experiments run at 1GPa and 1250-1475 °C. Error bars indicate standard error during unmixing.</figcaption></figure>
<figure><img src="image_11.png" /><figcaption>**Figure 10.** Relative percent decrease in oxide concentration due to cations bonding with S as a function of oxygen fugacity. Oxide concentrations for cations that bond with S decrease in very reduced silicate melts. Oxide concentrations of FeO, Cr₂O₃, and MnO will be reduced by 100%. On average, oxide concentrations of CaO will decrease by 13%. When MgS and Na₂S are detected, oxide concentrations will decrease by 22% and 28% respectively.</figcaption></figure>