# Multiple sulfur isotopes evidence deep intra-slab transport of sulfate-rich fluids

Isabelle Genot, Samuel Angiboust, Pierre Cartigny

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Isabelle Genot, Samuel Angiboust, Pierre Cartigny. Multiple sulfur isotopes evidence deep intra-slab transport of sulfate-rich fluids. Geochimica et Cosmochimica Acta, 2024, 377, pp.84-100. 10.1016/j.gca.2024.05.025. insu-04841359

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Distributed under a Creative Commons Attribution 4.0 International License
1 Multiple sulfur isotopes evidence deep intra-slab transport of sulfate-rich fluids

2 Isabelle Genot¹, Samuel Angiboust²,³, Pierre Cartigny¹

3

4 ¹ Université de Paris Cité, Institut de Physique du Globe de Paris, F-75005 Paris, France

5

6 ² Ecole Normale Supérieure of Lyon, Lyon University, LGL-TPE, 46 Allée d'Italie, F-69007

7 Lyon, France

8

9 ³ Institut Universitaire de France, Paris, France

10

11

12 Corresponding author: Isabelle Genot, genot@ipgp.fr

13

14 Abstract

15

16 Sub-arc mantle oxidation is often attributed to sulfates from seawater transported in aqueous
17 fluids derived from slab dehydration, but sulfur speciation and origin(s) in high-pressure fluids
18 remain weakly constrained. Here, we use the four sulfur stable isotopes (³²S, ³³S, ³⁴S, ³⁶S) to
19 decipher processes, e.g. mixing of two end-member reservoirs and open-system reaction
20 occurring during sulfur devolatilization, and deduce the nature and origin of sulfur species in
21 aqueous slab-derived fluids. We analyzed high pressure metamorphic rocks from the
22 Monviso massif Lower Shear Zone (LSZ; Western Alps), recording evidence of an interface-
23 parallel, channelized fluid flow near peak burial metamorphic conditions (~ 550°C and 2.7
24 GPa). We sampled variably metasomatized Fe-Ti metagabbros, hybrid (talc) schists and
25 serpentinites containing sulfides supposed to preserve the aqueous fluid composition, as well
26 as other alpine metagabbro, metabasalt and serpentinites from greenschist to eclogite-facies
27 devoid of the studied metasomatic overprint.
The millimeter to centimeter-sized sulfide crystals found in abundance within LSZ lithologies are mainly pyrites containing inclusions of the peak burial mineral assemblage (e.g. garnet, omphacite, rutile) indicating their crystallization at peak conditions or during incipient exhumation. The sulfides in the non-metasomatic metagabbros and metabasalts cover a narrow range in δ¹³S of 0.89 ± 0.63‰ and Δ¹³S of 0.017 ± 0.01 (n = 6) that could be inherited from hydrothermal sulfides formed on the seafloor. In contrast, the LSZ sulfides from the Fe-Ti metagabbros, serpentinites and hybrid schists (n = 24) are enriched in ³⁴S and ³³S (bulk δ¹³S from 4.91 to 21.32‰ and Δ¹³S from 0.02 to 0.06‰) compared to the sulfides from the non-metasomatic rocks. Mixing between reduced sulfur species cannot explain all the data without requiring an end-member composition as high as δ³⁴S ~ 22‰ and Δ³³S ~ 0.06‰, i.e. values rarely recorded by sulfides in the literature. Instead, we propose that this ³³S-³⁴S-enrichment results from open-system sulfate reduction to sulfides in aqueous fluids at 550°C, with an initial δ³⁴S ~ 12 ± 3‰ and Δ³³S ~ 0.025 ± 0.010‰. Because this signature is different from seawater sulfate, we suggest that sulfate was produced by hydrothermal sulfide oxidation during antigorite breakdown at 650°C, in agreement with Mg, Ni and Cr enrichments in the metasomatized HP mafic lithologies. This model is consistent with relatively oxidized ³⁴S-enriched arc lavas and can potentially explain recycled sulfur with positive δ³⁴S signature recorded by ocean island basalts.

46

(388 words)

48

Keywords: Sulfur isotopes, sulfur speciation, subduction zones, HP-fluids, sub-arc mantle oxidation

51

(Main text: 9186 words, 7 Figures, 3 Tables; Supplementary materials)

53
# 1. Introduction

Sulfur behavior in the sub-arc mantle is not well constrained, despite its potential role in mantle wedge oxidation (e.g. Kelley and Cottrell, 2009; Muth and Wallace, 2021; Taracsak et al., 2023), ore deposit formation (e.g. Simon and Ripley, 2011) and its climatic impact through SO₂ emission by arc volcanoes (e.g. Self et al., 1993). Melt inclusions in arc lavas generally contain higher S contents on average (around 300-3200 ppm with few values up to 6000 ppm in Central Mexico; Bouvier et al., 2008; Muth and Wallace, 2021; Wallace, 2005) than reduced mid-ocean ridge basalts (MORB, ~ 700-2000 ppm; Jenner and O'Neill, 2012). At given pressure and temperature conditions and FeO content, the high S contents of undegassed arc magmas suggest that the magmas are also elevated in oxygen fugacity (Wallace, 2005), which will stabilize S⁶⁺ in addition to S²⁻, the latter of which has limited solubility in basaltic silicate magmas (Jugo et al., 2010). This is consistent with both Fe³⁺/ΣFe of basaltic arc magmas and mantle xenoliths (Blatter and Carmichael, 1998; Brounce et al., 2014; Cottrell and Kelley, 2011; Kelley and Cottrell, 2009; McInnes et al., 2001) as well as the occurrence of oxidized sulfur (Bénard et al., 2018; Muth and Wallace, 2021). The sulfur isotopic data of arc lavas and arc melt inclusions show a wide and mostly positive δ³⁴S range from ~ -9.0 to 21.0‰, averaging at 4.7 ± 3.4‰ (n = 329) (Alt et al., 1993; Bénard et al., 2018; Bouvier et al., 2008; De Hoog et al., 2001; Gurenko, 2021; Gurenko et al., 2018; Ionov et al., 1992; Kawaguchi et al., 2022; Muth and Wallace, 2021; Rielli et al., 2018; Taracsák et al., 2023; Ueda and Sakai, 1983; Woodhead et al., 1987) compared to MORB with δ³⁴S ~ -1.0 ± 0.5‰ (Labidi et al., 2013; δ-values expressed here in the conventional notation, i.e. as [($^{34}$S/$^{32}$S<sub>sample</sub>)/($^{34}$S/$^{32}$S<sub>VCDT</sub>) -1] x 1000; Hulston and Thode, (1965), with V-CDT denoting for Vienna Canyon Diablo Troilite international standard). These positive signatures have been attributed to slab-derived fluids from sediments and altered oceanic crust (AOC), which reflect that of seawater sulfates (modern seawater δ³⁴S ~ 21‰; Ono et al., 2012), along with a causal link to their oxidized nature. Negative δ³⁴S have been proposed to reflect a higher
proportion of biotic sulfides derived from seafloor sediments (Figure 1A; Bouvier et al., 2008). 81

However, the ³⁴S-enrichment and/or S oxidation of arc magmas could also result from S 82 degassing from magmas at low pressures, sulfide segregation from arc magmas during melt 83 differentiation, and/or crustal thickening (e.g. Burgisser and Scaillet, 2007; Lee et al., 2018; 84 Marini et al., 1998; Marini et al., 2011; Tang et al., 2020). Although a significant S isotopic 85 fractionation at equilibrium during degassing was ruled out (Ueda and Sakai, 1984; Alt et al., 86 1993; Woodhead et al., 1987), kinetic fractionation during degassing could explain ³⁴S- 87 enriched degassed arc lavas up to ~ 21‰ (see Labidi et al., 2015). 88

Paleo-subduction zones can bring new insights on S cycling, specifically, ancient subducting 90 slabs provide a record of fluids released from the slab and bypass shallow processes 91 influencing arc magma signatures and the erupted rock record (Crossley et al., 2018; Evans 92 et al., 2014; Giacometti et al., 2014; Li et al., 2020; Li et al., 2021b; Su et al., 2019; Walters 93 et al., 2019). The exhumed (U)HP-LT rocks from the blueschist-eclogite facies transition from 94 which substantial amounts of fluids are released (Schmidt and Poli, 2003) contain sulfides 95 precipitated in veins and metasomatic rinds showing highly variable δ³⁴S from ~ -30 to 25‰ 96 (see compilation of Li et al., 2021a). Based on oxide mineral assemblages associated with 97 sulfides, reduced sulfur, with H₂S as major species, was suggested to be stable in aqueous 98 fluids (Crossley et al., 2018; Evans et al., 2014; Li et al., 2016), in agreement with 99 thermodynamic calculations (Li et al., 2020; Piccoli et al., 2019). Given the restricted isotopic 100 fractionation between H₂S and sulfide at high pressure (1000ln(αH₂S-S₂) < 2‰ at T > 500°C, 101 Eldridge et al., 2016), the fluid composition recorded by sulfides in HP veins could be 102 inherited from S isotopic composition of altered protoliths. Indeed, the sediments, AOC and 103 hydrated mantle carry sulfides (originally formed on the seafloor either through low 104 temperature biotic processes or high temperature abiotic processes) that cover a wide δ³⁴S 105 range from ~ -40 to +20‰ (Figs. 1A, B and C respectively; Alford et al., 2011; Alt, 1994; Alt 106 et al., 1989; Alt et al., 2012a; Alt et al., 2012b; Alt et al., 2013; Alt and Shanks III, 2003, Alt 107
and Shanks, 1998; Alt et al., 2007; Alt and Shanks, 2011; Delacour et al., 2008a; Delacour et al., 2008b; Ono et al., 2012; Rouxel et al., 2008; Schwarzenbach et al., 2012; Schwarzenbach et al., 2018a; Schwarzenbach et al., 2018b), potentially explaining the HP metasomatic sulfides isotopic signature. High-pressure metasomatic sulfides also record few $\delta^{34}\text{S}$ values higher than seawater sulfate (which varied through time from 12‰ during the Permian to 38‰ in the Cambrian; Kampschulte and Strauss (2004)), but these high values are rarely observed in the sulfides formed on the seafloor by alteration and/or low-pressure metasomatism. Rather, this signature was attributed to an additional process, possibly open system sulfate reduction in high-pressure fluids, or during seawater sulfate expulsion from pore water at low pressures (Crossley et al., 2018; Evans et al., 2014; Li et al., 2021b; Schwarzenbach et al., 2018; Su et al., 2019; Walters et al., 2019). This former possibility has important implications as it points to the sulfate stability in metamorphic fluids, in agreement with both recent thermodynamic calculations (Connolly and Galvez, 2018; Debret et al., 2017; Evans and Frost, 2020; Walters et al., 2020; Padrón-Navarta et al., 2023) and non-conventional isotopes such as $\delta^{56}\text{Fe}$ and $\delta^{66}\text{Zn}$ (Debret et al., 2015; Debret et al., 2021; Debret et al., 2016; Gerrits et al., 2019; Pons et al., 2016).

In this study, we investigate the sulfur speciation in HP aqueous fluids using the four sulfur stable isotopes ($^{32}\text{S}$, $^{33}\text{S}$, $^{34}\text{S}$ and $^{36}\text{S}$) systematics in a series of HP-rocks from the Monviso Lower Shear zone, which sample large-scale fluid-rock interactions, in a well-preserved oceanic crust section (Angiboust et al., 2012; Spandler et al., 2011). We infer the occurrence of slightly $^{34}\text{S}$-enriched sulfates in HP aqueous fluids and their reduction in an open system.

## 2. Definition of $\delta'^{34}\text{S}$-$\Delta'^{33}\text{S}$-$\Delta'^{36}\text{S}$ systematics

Given that multiple sulfur isotope geochemistry is rarely applied to metamorphic rocks, here, we provide some basic background. At equilibrium, an isotopic exchange reaction is similar to any cationic exchange reaction, where the fractionation factor $^{34/32}\alpha = (^{34}\text{S}/^{32}\text{S})_A/(^{34}\text{S}/^{32}\text{S})_B$
describes the partitioning of two isotopes between two phases at a given temperature. In this study, we use the logarithmic delta notation to express the isotopic compositions of a phase, as $\delta'^{34}\text{S} = \ln[(^{34}\text{S}/^{32}\text{S}_{\text{sample}})/(^{34}\text{S}/^{32}\text{S}_{\text{VCDT}})] \times 1000$ (Hulston and Thode, 1965), because the fractionation laws relating the different isotope compositions are linear and can be predicted from isotope masses between two phases (Bigeleisen and Mayer, 1947; Urey, 1947; Young et al., 2002):

$$141 \quad ^{33/34}\lambda_{A-B} = \ln(^{33/32}\alpha_{A-B}) / \ln(^{34/32}\alpha_{A-B}) \sim (1/m_{32} - 1/m_{33}) / (1/m_{32} - 1/m_{34}) \sim 0.515$$

$$142 \quad ^{36/34}\lambda_{A-B} = \ln(^{36/32}\alpha_{A-B}) / \ln(^{34/32}\alpha_{A-B}) \sim (1/m_{32} - 1/m_{36}) / (1/m_{32} - 1/m_{34}) \sim 1.89$$

with m being the mass of each isotope, $^{3x/32}\alpha_{A-B}$ being the isotope fractionation factor between two phases (x = 3 or 6) and $^{3x}\lambda_{A-B}$, (or $^{3x}\theta_{A-B}$ at equilibrium) being the “mass exponent”. We emphasize that the difference between the conventional and logarithmic notations is here as high as 0.22‰, averaging at 0.07 ± 0.06‰ for $\delta'^{34}\text{S}$ and < 0.003‰ for $\Delta'^{33}\text{S}$.

At equilibrium, $^{3x}\lambda_{A-B}$ describes the relative fractionation between $^{3x}\text{S}/^{32}\text{S}$ and $^{34}\text{S}/^{32}\text{S}$ (Fig. S1). Therefore, a slope of ~0.515 in $\delta'^{33}\text{S}-\delta'^{34}\text{S}$ and ~1.89 in $\delta'^{36}\text{S}-\delta'^{34}\text{S}$ between two components would be consistent with mass-dependent fractionation at equilibrium.

Deviations of $\lambda$ values from the equilibrium reference are quantified with the $\Delta$ parameter expressed in logarithmic notations as (Ono et al., 2003; Fig. S1):

$$153 \quad \Delta'^{33}\text{S} = \delta'^{33}\text{S} - \delta'^{34}\text{S} \times 0.515$$

$$154 \quad \Delta'^{36}\text{S} = \delta'^{36}\text{S} - \delta'^{34}\text{S} \times 1.89$$

Small non-zero $\Delta'^{33}\text{S}$ values (typically between -0.2 and +0.2‰) can result from mass-dependent processes such as Rayleigh distillation or other mass conservation effects such as mixing and/or processes out of equilibrium (Farquhar et al., 2007a; Ono et al., 2006). As
detailed further in the discussion, the mass exponent will allow us to decipher these processes.

### 3. Geological context and samples

The Monviso metaaphiolite is part of the Liguro-Piemontese domain in the Western Alps (Fig.

2A) and represents an eclogitized portion of oceanic crust formed during the opening of the slow-spreading Tethyan basin in the Jurassic (163 ± 2 Myr; Rubatto and Hermann, 2003).

The Monviso massif is located to the East of the Schistes Lustrés, a paleo-accretionary domain, which corresponds to a fossil accretionary wedge formed of Mesozoic metasediments and embedded mafic and ultramafic blocks metamorphosed in the blueschist facies (Agard et al., 2002; Schwartz et al., 2000; Herviou et al., 2022). To the East of the Monviso massif, the Dora Maira massif represents a stack of remnants from the thinned European continental margin subducted to eclogite facies in the Eocene (Bonnet et al., 2022; and references therein). The Monviso ophiolite is divided in two metamorphic units (Figs. 2B and C): the Monviso and Lago Superiore units (LSU) separated by the Upper Shear Zone (USZ) (Angiboust et al., 2012). The LSU recrystallized ca. 45 Myr ago (Rubatto and Hermann, 2003) in the lawsonite-eclogite facies, at ~550°C/2.7 GPa (Angiboust et al., 2012; Groppo and Castelli, 2010) and partially during the exhumation in the epidote-blueschist and greenschist facies (Angiboust and Glodny, 2020; Gilio et al., 2020; Locatelli et al., 2019; Schwartz et al., 2000). The LSU preserves a unique, km-thick, continuous section of altered oceanic crust and lithospheric mantle (e.g. Angiboust et al., 2011). The Intermediate Shear Zone (ISZ) and Lower Shear Zone (LSZ) separate the metabasalt portion from the Fe-Ti metagabbros and the Mg-Al metagabbros from the Iherzolitic serpentinite sole respectively (Fig. 2C). The petrology and geochemistry of these lithologies were extensively studied (Angiboust et al., 2012; Angiboust et al., 2014; Groppo and Castelli, 2010; Hattori and Guillot, 2007; Locatelli et al., 2018; Nadeau et al., 1993; Philippot and Kienast, 1989; Philippot and Selverstone, 1991; Spandler et al., 2011). Based on these studies, we focus
here on the mylonitized Fe-Ti metagabbros found as eclogite breccias in the LSZ embedded in the serpentinite matrix, because they preserve different types of veins and matrices sampling multiple fluid-rock interaction events, occurring at peak burial conditions (Angiboust et al., 2011; 2014; Locatelli et al., 2019; Hoover et al., 2022). The veins that have the same mineral assemblage and composition as the host rock, i.e. omphacite, garnet and rutile, are a record of both local and external fluid influx from the metagabbros and serpentinites. The metasomatic rinds of the eclogite breccias contain hydrous minerals (lawsonite, phengite, chlorite, talc) and disseminated sulfides (Figs 3A and B) showing marked enrichments in Cr, Ni, Mg and high δ¹¹B as well as the large ion lithophile elements (LILE), reflecting extensive external fluid circulation from serpentinite dehydration at sub-arc depth and sediments/AOC interaction (Angiboust et al., 2014; Locatelli et al., 2019; Spandler et al., 2011). The hybrid schists, which form around the eclogite blocks at the contact with serpentinite, are dominated by chlorite and talc with (up to) cm-scale sulfide crystals (Figs. 3C, D and E). Hybrid schists carry a large amount of water in chlorite and talc and could thus represent an important reservoir of fluids released at sub-arc depth or beyond (Spandler et al., 2008).

The LSZ is generally seen as allowing a focused channelized fluid flow, where growing sulfides would record the sulfur isotopic signature and speciation of fluids released during serpentinite dehydration at sub-arc depth. To better address this issue, we sampled mafic (metabasalt and metagabbro), ultramafic rocks (serpentinites and hybrid schists) and metasediments from different metamorphic grades and deformation/metasomatism stages described in Table 1, that show geochemical, structural and age similarities as evidence of their relationship in the Liguro-Piemontese domain (Bertrand et al., 1987; Lombardo et al., 2002; Piccardo et al., 2002; Tricart and Lemoine, 1991). Stage T1 corresponds to the absence of deformation from greenschist to eclogite facies, before the mylonitization in the eclogite facies that corresponds to stage T2. Stage T3 represents the external fluids metasomatic episode in the eclogite facies (Fig. 2D). Comparing samples from stage T3 to samples from stages T1 and T2 will bring insights into fluid origins and sulfur speciation.

213
214 4. Analytical methods

215

216 4.1. Major and trace element analyses

217

218 Bulk-rock chemical analyses were carried out on the metasediments, mafic and ultramafic
219 rocks from each of the three stages. The major (including Sc) and trace element contents
220 were obtained by alkali fusion with LiBO₂ then by acid digestion and measured by ICP-OES
221 and ICP-MS respectively, following the protocol implemented by Carignan et al. (2001), at
222 the SARM, at CRPG (France). Uncertainties and detection limits are presented in Table S1.

223

224 In order to identify the type of sulfides present (e.g. pyrite, chalcopyrite, pyrrhotite) and
225 characterize any intragrain and/or intergrain heterogeneities, we collected Fe, S, Co, Ni and
226 Cu chemical maps of sulfides and performed spot analyses with a Cameca SX Five electron
227 microprobe at the Camparis facility (Sorbonne Université). The Co, Ni, Cu, Fe and S
228 contents of sulfides were mapped using an acceleration voltage of 20 keV, a beam current of
229 200 nA, a pixel size of 3-5 µm and a dwell counting time of 80 ms. The counting time for spot
230 analyses was set to 50 s. Each analysis corresponds to four measurements with
231 reproducibility for S, Fe, Co, Ni and Cu of 0.25 wt%, 0.39 wt%, 0.07 wt%, 0.06 wt% and 0.08
232 wt% (1σ) respectively. Mean detection limits for S, Fe, Co, Ni and Cu are 0.11 wt%, 0.19
233 wt%, 0.10 wt%, 0.20 wt%, 0.32 wt% and mean standard deviations are 1.0 wt%, 2.5 wt%, 0.1
234 wt%, 0.2 wt%, 0.3 wt% respectively (Table S2).

235

236 The major element contents of mineral inclusions in the samples VI1907, VI1018, VI1924
237 and VI1912 were measured using an acceleration voltage of 15 keV, a diameter size of 5 µm
238 and a beam current of 10 nA on 3-4 different spots. The standards used for the calibration
239 were orthoclase (Al, K), Cr₂O₃ (Cr), diopside (Mg, Si), Fe₂O₃ (Fe), MnTiO₃ (Mn, Ti), anorthite
240 (Ca), albite (Na), pyrite (S), CaF₂ (F). We also collected chemical maps and performed spot
241 analyses on the thin section VI1912 (core to rim of omphacite) and MV with a scanning electron microscope (SEM) Zeiss EVO MA10 using internal calibration standards at IPGP.

242
243
244 4.2. Bulk sulfur isotope analyses

245
246 Bulk sulfur was extracted chemically from ~2 to 8 grams of bulk rock powder when the
247 sulfide grains were too small to be picked by hand. In table 2, an asterisk indicates samples
248 from which the sulfides were picked by hand. Under a continuous nitrogen flux, 5 mL of
249 degassed HCl 6N was added to the powder to dissolve and reduce monosulfide phases
250 ("acid volatile sulfur" or AVS such as FeS, NiS) into H₂S which was subsequently trapped in
251 a silver nitrate solution to precipitate Ag₂S (Rickard and Morse, 2005). Then 20 mL of a
252 chrome reducing solution (CRS; see Canfield et al., 1986) was used to reduce disulfide
253 phases (e.g. FeS₂, CuFeS₂) into H₂S and form Ag₂S. The extraction of sulfates necessitated
254 20 mL of a strongly reducing solution of HCl, H₂PO₃ and HI (Thode et al., 1961). The
255 released H₂S was then trapped in 0.1 M of NaOH, into which the subsequent additions of
256 HNO₃ and AgNO₃ yielded a clear Ag₂S precipitate (Geng et al., 2018).

257 The Ag₂S precipitate was weighed to 0.001 mg precision that allowed us to determine the S
258 content from the mass of sample powder dissolved (for bulk extraction only), and to
259 characterize the sulfide nature as mono- or di- sulfide based on the obtained fraction. Then,
260 the Ag₂S was wrapped in Al-foil packets and placed in nickel bombs under vacuum for
261 fluorination with F₂ at 250°C overnight. The produced SF₆ was purified by cryogenic
262 techniques and gas chromatography following Farquhar et al. (2007b). The 127, 128, 129,
263 131 mass/charge ratio for ³²SF₅⁺, ³³SF₅⁺, ³⁴SF₅⁺ and ³⁶SF₅⁺ respectively, was then analyzed
264 using a dual inlet ThermoFinnigan MAT 253 mass spectrometer. The resultant δ¹³S are
265 expressed versus V-CDT after calibration of our in-house SF₆ tank with the known IAEA S1
266 $\delta^{34}$S of -0.30‰ (Robinson, 1993). The $\Delta^{33}$S and $\Delta^{36}$S values are scaled with V-CDT mean
267 value of $\Delta^{33}$S$_{CDT}$ = -0.029 ± 0.005 and $\Delta^{36}$S$_{CDT}$ = -0.129 ± 0.087‰ (Defouilloy et al., 2016).

268 The accuracy and precision of the measurements are based on the reproducibility of IAEA
269 S1 standard analyses at each session (Table S3), averaging for $\delta^{34}$S at -0.40 ± 0.07‰, for
270 $\Delta^{33}$S at 0.085 ± 0.007‰ and for $\Delta^{36}$S at -0.914 ± 0.134‰ (n = 13; 1σ), within error of those
271 measured at IPGP and other laboratories (e.g. Farquhar et al., 2007b; Labidi et al., 2012;
272 Ono et al., 2006). Most analyses were duplicated with a maximum chemical and analytical
273 reproducibility of 1.11‰, 0.015‰ and 0.582‰ for $\delta^{34}$S, $\Delta^{33}$S and $\Delta^{36}$S respectively
274 (standard reproducibility was indicated instead when samples were not duplicated).

275
276 4.3. In-situ sulfur isotope analyses

277
278 We also performed *in situ* sulfur isotope analyses at CRPG (Nancy) using a secondary ion
279 mass spectrometry Cameca IMS 1270. The thin sections were coated with pure gold and
280 inserted in the sample chamber and pumped to pressures below 5.10<sup>-9</sup> mbar. A Cs<sup>+</sup> primary
281 ion beam of 2.5 nA was focused to 10 µm diameter and electrons were bombarded to
282 compensate positive charges accumulated on the surface. The background was measured
283 while pre-sputtering a cleaned surface of 20x20 µm<sup>2</sup> during 30-120s. The two stable isotopes
284 <sup>32</sup>S and <sup>34</sup>S were measured on Faraday cup detectors with a mass resolving power of ~ 2500,
285 at mass/Δmass of 10% peak height, to separate S ions from isobaric interferences. The counts
286 on <sup>32</sup>S were ~ 8.10<sup>8</sup>/s and the isotope measurements performed over 40 cycles.

287 Two reference materials were measured at the beginning and end of the session to correct
288 for time drift and instrumental mass fractionation (IMF) with a reproducibility of 0.05‰ (n =
16, 1σ) for Sierra pyrite (δ³⁴S = 2.17 ± 0.28‰ (2σ); LaFlamme et al., 2016) and 0.10‰ (n = 6, 289 1σ) for Spain pyrite (δ³⁴S = -1.61 ± 0.20 ‰ (2σ); Kitayama, 2018) (Table S4). The IMF value 290 obtained from the standards was -2.05‰, which would be somewhat inconsistent with the *in situ* δ³⁴S values obtained for our samples (also expressed in logarithmic notation), since the 291 *in situ* values are similar to the bulk values (i.e. a correction of +2.05‰ would shift the mean 292 to higher values, inconsistent with the bulk signature; Table S4). Alternatively, the pyrites 293 with Ni/M < 0.003 could be associated with higher IMF than the pyrites with Ni/M > 0.003 294 (Kitayama, 2018). Assuming an IMF of 0‰ for Ni/M = 0.003 and using the pyrite Spain, with 295 Ni/M = 0.001281 associated with an IMF of -2.05‰, we derived a relationship scaling IMF 296 with Ni/M ratios. The resulting average δ³⁴S values are within 1‰ similar to the bulk value, 297 except for VI1914, which does not contain Ni (Table S4). 298

300

301 5. Results

302

303 5.1. Mineral assemblages and their chemical composition

304

305 As the lithologies and petrology of the Monviso metaophiolite have been extensively 306 described (Angiboust et al., 2011; Locatelli et al., 2018; Philippot and Kienast, 1989; 307 Schwartz et al., 2000; Spandler et al., 2011), we briefly summarize the mineral assemblages 308 with special emphasis on the sulfides *sensu lato* (Table 1). Here, we describe the different 309 types of sulfides encountered within the six samples (MV, VI1018, VI1102-d, VI1912, VI1914 310 and VI1924) in section 5.4.

311 The sulfides in the undeformed (T1) and mylonitized samples (T2), i.e. away from the LSZ, 312 characterize the evolution from alteration on the seafloor to eclogite facies. The undeformed 313 pillow lava from the Pic du Pelvat (Schistes Lustrés domain, Queyras region, France; PV on 314 Fig. 2A) contains cm-scale sulfides associated with epidote and carbonate veins, which 315 precipitated at LT/HP greenschist facies conditions (319°C, ~1.4 GPa; Herviou et al., 2022).
The sulfides would thus have formed either during hydrothermal alteration on the seafloor or at the onset of slab subduction. The metagabbro from the Queyras region (QM on Fig. 2A) has mm- to cm-scale disseminated sulfides associated with relict clinopyroxene surrounded by glaucophane and plagioclase replaced by lawsonite, fingerprinting blueschist facies metamorphism.

The mylonitized Fe-Ti metagabbros (VI1913, VI1914, VI1915, VI1916; stage T2) contain omphacite, garnet, rutile and quartz, the peak burial mineral assemblage at eclogite-facies conditions, with mm-scale deformed sulfides along the foliation, evidencing their formation before or during the mylonitization. A static eclogite facies metagabbro (VI1102-d; stage T1) also contains < mm-scale sulfides in equilibrium with the peak mineral assemblage, suggesting that the sulfides in the mylonitized metagabbros formed before the deformation.

The serpentinites from the sole (VI1901, VI1902, VI1026; stage T2) show antigorite and magnetite as the main phases, but sulfides < mm-scale are sparse and were only found in the VI1026 sample. Two metasediments, SL1801 and VI1003, from the Schistes Lustrés Unit and Monviso Unit respectively, have disseminated sulfides < mm-scale.

Among the metasomatized samples (VI1904, VI1906, VI1907, VI1912, MV, VI1018; stage T3) collected on the LSZ, the Fe-Ti metagabbros show mm to cm-scale sulfides both in the garnet-omphacite matrix associated with rutile (Fig. 4A) and in veins with omphacite, chlorite ($X_{Mg}$ = 0.7-0.9), talc ($X_{Mg}$ = 0.9), phengite ($Si^{4+}$ = 3.2-3.8 (atoms per formula unit, or a.p.f.u.) and lawsonite replaced by epidote and paragonite (Figure 4B-F; mineral compositions in Table S5). The sulfides also contain inclusions of garnet, rutile, omphacite, Mg-chlorite and phengite with $Si^{4+}$ up to 3.8 a.p.f.u. In some samples, the occurrence of epidote, glaucophane and phengite with low Si content (~ 3.2-3.4 a.p.f.u.) and $Fe^{3+}$-enriched rim in omphacite and zoisite (Figs. 4C, E, F) result from both the epidote-blueschist and greenschist exhumation overprint (Angiboust and Glodny, 2020; Gilio et al., 2020; Locatelli et al., 2019).
The hybrid schists (LSZ05, VI1908, VI1920 to VI1927; stage T3) contain cm-scale sulfides associated with serpentine, talc and chlorite ($X_{Mg}$ = 0.92 in VI1924). The pyrite in the VI1924 sample also contains inclusions of tremolite (Ca-amphibole), talc (Fig 6I) and diopside, which may reflect lower pressure conditions. But their stability together with the occurrence of rare omphacitic clinopyroxenes (see Angiboust et al., 2014; Gilio et al., 2020) may indicate that the hybrid schists and hence the pyrite formation occurred at near peak burial conditions or upon early exhumation, still in the eclogite-facies. Low-pressure exhumation overprint is also discussed below. The serpentinites in the LSZ (VI1910-17a-28) contain < mm-scale sulfides associated mostly with magnetite, antigorite, lawsonite and rare carbonates (VI1928).

## 5.2. Bulk major and trace element compositions

The major and trace element compositions of the Monviso lithologies have also been widely described (Angiboust et al., 2014; Locatelli et al., 2019; Spandler et al., 2011). We complete the dataset with 10 new samples, reported in Table S1.

The trace element signatures measured in this study are represented together with those previously published (Angiboust et al., 2014; Locatelli et al., 2019; Spandler et al., 2011) of the sole serpentinites, the Fe-Ti and Mg-Al metagabbros and the marine sediment average composition GLOSS II (Plank, 2014); Fig. S2). As outlined earlier by these authors, the metasomatized eclogite breccias (T3) are enriched in LILE (Ba up to 14.6 ppm; Cs = 0.27-0.6 ppm; Rb = 1.82-4.94 ppm), Pb (up to 3.11 ppm), and Ni (249-554 ppm), Cr (273-805 ppm), MgO (8.88-14.10 wt%) and depleted in FeO (8.20-12.64 wt%), V (170-286 ppm) and Sc (19.40-34.97 ppm) compared to the non-metasomatized Fe-Ti metagabbros (T2), which have, for VI1914, LILE and Pb contents below the detection limit, Ni = 28.9 ppm, Cr = 9.6 ppm, MgO = 5.79 wt%, FeO = 19.62 wt%, V = 1065 ppm and Sc = 54.35 ppm. The hybrid schists show the same pattern as the LSZ05 and LSZ18 reported in Angiboust et al. (2014) showing ultramafic compositions with MgO = 25.32-29.08 wt%, Ni = 698-1301 ppm and Cr = 266-1750 ppm, tending toward the serpentinite compositions with MgO = 34.85-36.31 wt%,
372 Ni = 1716-2132 ppm and Cr = 2724-2839 ppm. The hybrid schist VI1920 also shows a
373 particular pattern with enrichment in LREE and MREE except in Pb and Sr.

374

375 5.3. Bulk sulfur content and multi-isotope compositions

376

377 The sulfur isotope compositions of sulfides are illustrated in Figure 5 and reported in Tables
378 2 and 3. No AVS phase such as pyrrhotite was present. The sulfides are mainly disulfides,
379 such as pyrites as confirmed by the microprobe analyses for some samples (Table 1; Table
380 S2). The non-metasomatized samples (T1 and T2, n = 6), i.e., the greenschist pillow lava,
381 the blueschist metagabbro as well as the Fe-Ti metagabbro contain no detectable amount of
382 sulfate and have between 969 and 7017 ppm of reduced sulfur (although these contents are
383 likely biased to higher values because only samples with visible sulfides were selected) and
384 a rather homogeneous isotopic signature averaging at 0.89 ± 0.63‰, 0.017 ± 0.010 and -
385 0.004 ± 0.016‰ for δ¹³⁴S, Δ¹³³S and Δ¹³⁶S respectively. The sole serpentinites (T2) VI1901
386 and VI1902 contain no sulfide but preserved 12-26 ppm of oxidized sulfur with δ¹³⁴S of -4.04
387 and -2.49‰, Δ¹³³S of 0.041 and 0.035‰ and Δ¹³⁶S = -0.301 and -0.245‰ respectively while
388 the VI1026 has no sulfate but 39 ppm of reduced sulfur with δ¹³⁴S of 7.16‰, Δ¹³³S of 0.006‰
389 and Δ¹³⁶S of 0.198‰. The metasediments from the Monviso and Schistes Lustrés units
390 (VI1003 and SL1801) have only reduced sulfur varying between 70 and 1230 ppm with δ¹³⁴S
391 from -20.54 to -5.12‰, Δ¹³³S from -0.015 to 0.005‰ and Δ¹³⁶S from 0.032 to 0.403‰ (n = 2).
392 The metasomatized Fe-Ti metagabbros and the hybrid schists (T3) of the LSZ contain no
393 detectible sulfates and their reduced sulfur content range between 114 and 3474 ppm. Their
394 isotopic compositions are all ³³S-³⁴S-enriched compared to the mylonitized Fe-Ti
395 metagabbros and the sole serpentinite (T2), from 4.90 to 21.10‰ in δ¹³⁴S and from 0.023 to
396 0.062‰ in Δ¹³S (n = 24). Their Δ¹³S values are in the same range as their protolith, from -
397 0.177 to 0.188‰ and we observed no clear relation between the protolith and the
398 metasomatized samples (Fig. 5D). The metasomatized serpentinites (T3) have on average
399 higher sulfide content (up to 64 ppm) with a δ¹³S from 6.85 to 15.85‰, Δ¹³S from 0.038 to
400 0.044‰ and Δ¹³S from -0.055 to 0.262‰ (n = 3) and have sulfate S contents of about 14-72
401 ppm with δ¹³S not higher than their respective sulfides, from -3.67 to 4.86‰, with Δ¹³S from
402 0.031 to 0.034‰ and Δ¹³S from -0.316 to -0.216‰. The Lower Shear Zone metasediments
403 VI1918 and VI1919 contain sulfide S between 590 and 2180 ppm, respectively, with δ¹³S =
404 0.93 and -12.16‰, Δ¹³S = 0.003 and 0.046‰, Δ¹³S = 0.081 and 0.321‰, respectively.
405 It is worth noting that the measured Δ¹³S values for samples in this study are significantly
406 different from 0‰. However, we note that the analytical reproducibility of these
407 measurements is low, as high as 0.58‰ in 1σ (see section 4.2 and Table 3). Although our
408 models predict a relationship in the Δ¹³S-Δ¹³S systematic (see black curve, Figure 5d), the
409 large error bars on the Δ¹³S measurement obscure a clear assessment of the quality of the
410 fit of our preferred model with the data at this time.

411
412 5.4. Cobalt and nickel distribution and in-situ δ³⁴S in sulfides
413
414 In order to better trace the different sources of fluids and processes (e.g. mixing, open
415 system crystallization, redox changes) occurring during fluid percolation, we performed
416 chemical maps of Co and Ni of six sulfides from the different lithologies and measured in situ
417 the δ³⁴S value in five of the most characteristic grains (Table S4). We also measured the Fe,
S and Cu contents, which indicate that all measured sulfides are pyrite (with Cu content below detection limit), consistent with the absence of AVS phases (Table S2).

The pyrite in the mylonitized Fe-Ti metagabbro VI1914 (T2) has Ni below detection limit but patchy Co content enrichment up to 0.63 wt% and homogeneous $\delta^{34}\text{S}$ values similar to the bulk composition (Fig. 6A; $\delta^{34}\text{S}_{\text{bulk}} = 1.23 \pm 0.16\text{‰}$). We observed the same Co pattern in the static metagabbro VI1102-d (T1; Fig. 6B; Co averaging at $0.53 \pm 0.14$ wt%, Ni averaging at $0.01 \pm 0.06$ wt%, a sample where $\delta^{34}\text{S}$ was not measured).

The pyrite MV in the metasomatized Fe-Ti metagabbro (T3) shows a core enriched in Ni with concentrations as high as 0.84 wt%, and $\delta^{34}\text{S}$ up to 21.69‰ with no detectable Co, mantled with zones that are depleted both in Ni (as low as 0.14 wt%) and $\delta^{34}\text{S}$ (~ 11‰). After an episode of dissolution, a rim enriched in Co (up to ~ 0.8 wt%) and depleted in Ni (~ 0-0.11 wt% Ni), with $\delta^{34}\text{S}$ ~ 8‰ precipitated (Figs. 6E,F).

The metasomatic pyrite VI1912 (T3) reveals patchy zonation in Co, with concentrations as high as 0.4 wt% Co and ~ 0.18-0.65 wt% Ni in its core, mantled with homogeneous $\delta^{34}\text{S}$ values ~ 8‰, similar to the bulk value. Similar to the MV pyrite, after an episode of dissolution, a rim enriched in Co (up to 0.34 wt%), but depleted in Ni, with higher $\delta^{34}\text{S}$ of 13.78‰ precipitated (Figs. 6C,D). It is worth noting that this Co-enriched rim is in textural equilibrium with phengite with low Si content (~ 3.2-3.4 a.p.f.u.; Figs. 4C,D,E and Fig. S3). The core of the metasomatic pyrite VI1018 (T3) is richer in Co (~ 0.43-0.5 wt%) than in Ni (up to 0.14 wt%) with the same Co-rich rim up to 0.6 wt% as in VI1912 and $\delta^{34}\text{S}$ varying between ~ 8-13‰ (Fig. 6G).

The pyrite in the hybrid schist VI1924 (T3) preserved a core enriched in Co ~ 0.55-0.65 wt% and Ni ~ 0.23-0.29 wt% with homogeneous $\delta^{34}\text{S}$ values ~ 11‰. As for the other pyrites, a rim
enriched in Co (up to 0.83 wt%) and δ¹³S (up to 15.76‰) precipitated after an episode of dissolution, which also contains tremolite and talc inclusions (Figs. 6H,I). A single spot depleted in Co has a high δ¹³S of 18.87‰.

In summary, Co and Ni variations among pyrites are not correlated (Fig. S4) but a marked intergrain variability exists between Ni/Co and δ¹³S. Cobalt and Ni in sulfides represent 76% and 26% of the bulk rock in eclogites, the remaining being incorporated in e.g. omphacite, garnet and chlorite while sulfides represent the main host of S (Li et al., 2013; Li et al., 2021b; Walters et al., 2021). Thus, Ni and Co budgets could also be controlled by dissolution/precipitation of silicate minerals that would contribute to the fluid Ni and Co budget.

6. Discussion

## 6.1. Isotopic signature of undeformed (T1) and mylonitized samples (T2)

In order to determine the isotope characteristics of the fluids and therefore their origin, we first constrain the sulfur isotopic composition of the pyrites precipitated before the metasomatic episode at the peak burial conditions. All of the undeformed and mylonitized samples (T1 and T2), i.e. the pillow lava from the Pic du Pelvat, the metagabbro from the Queyras and the Fe-Ti metagabbro of the LSU, contain disulfides spanning a narrow range of δ¹³S = 0.89 ± 0.63‰, Δ¹³S = 0.021 ± 0.010‰ and Δ¹³S = -0.004 ± 0.160‰. This range is consistent with the AOC average δ¹³S of 0.9‰ (Alt, 1995) and within the uncertainty of reported Δ¹³S-Δ¹³S values (Fig. 1), suggesting that little isotopic fractionation occurred if S was devolatilized during subduction. Furthermore, the pyrites VI1914 (T2) and VI1102d (T1) have Ni contents below detection limit and Co ~ 0.6 wt%. This is consistent with the
466 hydrothermal pyrites formed on the seafloor, which contain little Ni (up to 0.048 wt%) but can
467 have Co up to 0.944 wt% (Keith et al., 2016), resulting from a higher solubility for Co than Ni
468 at low temperature in hydrothermal fluids (Brugger et al., 2016; Liu et al., 2011). The
469 available evidence is therefore consistent with the sulfides from the undeformed (T1) and
470 mylonitized (T2) samples having precipitated on the seafloor.
471

472 The sole serpentinites (T2) contain no or few sulfides (S ~ 40 ppm) and show δ¹³S of 7.16 ± 0.07‰, Δ¹³S of 0.006 ± 0.007‰ and Δ¹³S of 0.198 ± 0.134‰, within the admittedly large range of seafloor and subducted serpentinites S content and isotopic signature (Alt et al., 2012a; Alt et al., 2012b; Alt et al., 2013; Debret et al., 2014; Hattori and Guillot, 2007; Li et al., 2020). Sulfur might have devolatilized from the serpentinites as inferred by Debret et al. (2014), but with only one data point, we cannot assess this issue or whether S isotopic fractionation occurred.
473
474
475
476
477
478

479
480 The sulfur isotopic compositions of the metasediments SL1801 and VI1003 (T2) show δ¹³S of -5.12 ± 0.07‰ and -20.54 ± 0.21‰, Δ¹³S of 0.005 ± 0.007‰ and -0.015 ± 0.003‰ and Δ¹³S of 0.032 ± 0.134‰ and 0.403 ± 0.165‰ respectively, which is within the range of sulfides formed through bacterial sulfate reduction and that of subducted sediments, with δ¹³S varying from -50 to 30‰ and averaging at -20‰ since the Phanerozoic (Canfield and Farquhar, 2009; Goldhaber and Kaplan, 1980; Lew, 1981; Li et al., 2020; Walters et al., 2019); Fig. 1). As for the serpentinites, given the large range of sediments δ·¹³S, we cannot assess whether these values were modified during T1-T2 stages.
481
482
483
484
485
486
487
488
In summary, the sulfides (mainly pyrites) from the T1 and T2 stages seem to have precipitated on the seafloor, or at least kept the signature of the subducted protolith. Thus, we will refer to those isotopic ranges to compare with the metasomatic sulfides from the LSZ.

## 6.2. Isotopic signature and speciation of sulfur in intraslab fluids

The occurrence of disulfides in veins, mainly pyrites, with omphacite, garnet and rutile as well as hydrous minerals such as Mg-chlorite, phengite with high Si content (~ 3.7 a.p.f.u.) or lawsonite (pseudomorphed) testifies that the sulfide precipitation occurred at peak metamorphic conditions under fluid-saturated conditions. The exhumation overprint is characterized by pyrite with thin Co-enriched rims having $\delta^{34}\text{S} \sim 13-16\text{‰}$, associated with the occurrence of lower-pressure minerals (e.g. tremolite, phengite with low Si content), but seems to be volumetrically of restricted magnitude. Indeed, for instance, the bulk $\delta^{34}\text{S}$ of the VI1912 pyrite is similar to the pyrite core $\delta^{34}\text{S}$, consistent with the thinness of the Co-rich rims.

All the metasomatized sulfides in the eclogite veins and the hybrid schists (T3) have variable yet higher $\Delta^{33}\text{S}$ (up to 0.062‰) and $\delta^{34}\text{S}$ (up to 21.10‰) compared to their respective protoliths, including the mylonitized Fe-Ti metagabbro and the sole serpentinite from the Monviso unit whereas the $\Delta^{36}\text{S}$ values are in the same range of the non-metasomatized samples resulting from high analytical uncertainty (T1 and T2; Fig. 5A). The *in situ* analyses also reflect this variability with $\delta^{34}\text{S}$ from ~ 8‰ in the VI1912 and VI1018 samples to ~ 22‰ in the MV pyrite core (Fig. 6). This results in $^{33}\lambda$ (the exponent that relates variation in $^{33}\text{S}/^{32}\text{S}$ and $^{34}\text{S}/^{32}\text{S}$ abundances) of 0.5169 ± 0.0015 and $^{36}\lambda$ of 1.903 ± 0.025 for the metasomatized
512 Fe-Ti metagabbros, and $^{33}\lambda$ of 0.5171 ± 0.0018 and $^{36}\lambda$ of 1.88 ± 0.03 (2σ) for the hybrid schists, which are within error of the equilibrium values $^{33}\lambda$ of 0.515 and $^{36}\lambda$ = 1.89. In the following sections, we discuss processes such as mixing and open-system reaction that could have led to this isotopic enrichment compared to the host rock and we estimate intraslab fluid composition.

517

518 6.2.1. Kinetic isotope effects

519

520 Determining the isotopic composition of fluids requires that sulfides precipitated in equilibrium with the fluids, i.e. without kinetic isotope fractionation. Kinetic isotope fractionation, especially through diffusion, leads to heavy isotope enrichment in the residual phase (e.g. Mg evaporation experiments in Young et al. (2002)) and could explain the $^{34}$S-enrichment of the metasomatic sulfides. Walters et al. (2019) previously described $^{34}$S-enriched metasomatic sulfides and argued for a sluggish S self-diffusion (i.e. a kinetic isotope effect) during sulfide precipitation as the mechanism responsible for fluid (and hence sulfide) $^{34}$S-enrichment through a Rayleigh distillation model. However, sluggish S self-diffusion cannot explain the sulfide $^{34}$S-enrichment in our metasomatic sulfides. Indeed, the mass exponent $^{33}\lambda$ should tend toward a theoretical value of 0.504 (gray field on Fig. 5A), in response to the different isotope diffusivities (see Young et al., 2002). Thus, to the best of our knowledge, our observations are not consistent with available data on kinetic isotope effects involving multiple S-isotopes and are therefore not considered further to account for the metasomatic sulfides $^{33}$S-$^{34}$S enrichment at high pressure. Overall our multiple S-isotope data are in support of isotope equilibrium.

535

536 6.2.2. Mixing in reduced environment

537
Strong chemical potential gradients at the interface between two different lithologies facilitate large element mass transfer (Breeding et al., 2004; Hermann et al., 2006; John et al., 2004; Spandler and Pirard, 2013) that could lead to a mixing between different sulfur pools during fluid migration. In our isotope space, mixing would yield convex curves on a $\delta^{34}\text{S}-\Delta^{33}\text{S}$ diagram (Farquhar et al., 2007a; Ono et al., 2006). If a mixing between $^{33}\text{S}-^{34}\text{S}$-depleted and $^{33}\text{S}-^{34}\text{S}$-enriched components results in the enrichment in $\Delta^{33}\text{S}$ and $\delta^{34}\text{S}$ and slightly higher $^{33}\lambda$ than 0.515, the depleted component should have a $\delta^{34}\text{S} < 1.00\%$ and $\Delta^{33}\text{S} > 0.010\%$ to encompass the whole data range (Fig. 5B). Both the metagabbros (T1-T2) and the metasediments meet this condition. We thus graphically defined a metasediment end-member ($\delta^{34}\text{S} = -12\%$ and $\Delta^{33}\text{S} = 0.046\%$) and an AOC end-member ($\delta^{34}\text{S} = 1\%$ and $\Delta^{33}\text{S}$ from 0.02 to 0.028\%) (Fig. 5B).

Then, as neither the Fe-Ti metagabbros nor the metasediments can account for the Ni, Cr, MgO and the $\delta^{34}\text{S}-\Delta^{33}\text{S}$ enrichments recorded in the metasomatized samples (Fig. S2), we emphasize that the enriched component is the consequence of fluid infiltration released by serpentine breakdown, with $\delta^{34}\text{S}$ from 15.00 to 21.00\% and $\Delta^{33}\text{S} = 0.050-0.060\%$ (determined graphically on Fig. 5B). Following from recent thermodynamic models and isotopic studies (Li et al., 2020; Piccoli et al., 2019), we first assume that aqueous sulfur released during antigorite breakdown is reduced. Given the limited isotope fractionation between sulfide and aqueous sulfide ($^{34}\alpha_{\text{H}_2\text{S-FeSH}+} = 0.9993$ from Eldridge et al. (2021) or $^{34}\alpha_{\text{H}_2\text{S-pyrite}} = 1.0006$ from Ohmoto and Lasaga (1982) at 550°C), it would imply that sulfides in the hydrated mantle were as high as 22.00\% for $\delta^{34}\text{S}$ and 0.060\% for $\Delta^{33}\text{S}$. This model is however unlikely for at least two reasons: (1) the solubility of reduced sulfur ($\text{H}_2\text{S}$, HS⁻) is low, below 70 ppm during the antigorite dehydration (Piccoli et al., 2019) and could not account
for significant S-mobility; (2) the high δ¹³S-Δ¹³S signature of the second end-member is rarely recorded by sulfides and is mostly restricted to sulfates (Fig. 1).

To match such values, pristine (or unreduced) seawater sulfates could be a viable enriched end-member. The altered oceanic crust sampled at the Monviso formed ~ 165 Ma ago and subducted before 45 Ma (Rubatto and Hermann, 2003), with δ³S seawater sulfate value varying between 15.1 ± 1.4‰ and 20.7 ± 2.2‰ (Kampschulte and Strauss, 2004) and Δ³S of 0.043 ± 0.016‰ during the last 142 Ma (Masterson et al., 2016). The seawater sulfate enters the lattice of serpentine either during serpentinization on the seafloor or before subduction where the plate bends (Debret et al., 2017). Anhydrite preservation after pore water expulsion at low temperature and beyond blueschist-eclogite facies is still debated (see Walters et al., 2019) but could potentially release ³S-³S-enriched fluids if dissolved at 450-650°C (Tomkins and Evans, 2015). The seawater sulfates would thus be released in fluids during antigorite dehydration and anhydrite dissolution. The high solubility of sulfates in aqueous fluids (D(S)rock-fluid ~ 0.25, estimated with chlorite-harzburgite from Cerro del Almirez; Alt et al., 2012a) allows large S transfer through the slab when antigorite dehydration occurs, up to peak conditions ~ 550°C/2.7 GPa. The sulfide would thus record the pristine seawater sulfur isotopic composition if sulfates are reduced quantitatively, a model that is hard to reconcile with the evidence for open system evolution occurring during fluid transfer. Additionally, this mixing end-member is difficult to consider because the sole serpentinites from the Monviso preserved sulfates with δ¹³S from -4.04 to -2.49‰ (Table 3), inconsistent with the pristine seawater sulfates isotopic signature mentioned above. It is also worth mentioning that the mafic rocks from T1 and T2 stages contain no detectable sulfates. The ³S-³S-enriched component thus requires a more complex model than a quantitative (i.e. total) seawater sulfate reduction.

### 6.2.3 Open-system reactions
587 Open-system reactions (Rayleigh distillation) can produce the ³⁴S-³³S-enriched metasomatic sulfides (see section 2), where the instantaneous product is in equilibrium with the residue leaving the system. If we assume a temperature of 550°C, corresponding to the peak temperature, we can compute ³⁴α and ³³λ for pair species according to Eldridge et al. (2016), Eldridge et al. (2021) and Otake et al. (2008), and we can model the resulting S-isotope compositions of both product and residual reactant.

588 First, we consider a reducing fluid from which reduced sulfur species such as H₂S react with aqueous iron sulfide FeSaq to precipitate pyrites through the “hydrogen sulfide mechanism” (see Eldridge et al., 2021 and references therein). Given the restricted isotopic fractionation, i.e. ³⁴αH₂Saq-FeSH⁺ = 0.9993, ³³λ = 0.5156 and ³⁶λ = 1.8911 (Eldridge et al., 2021), the isotopic range of the metasomatic sulfides can be covered only if the initial H₂S source is heterogeneous in δ¹³S, Δ¹³S and Δ¹³S (Figs. 5C and D; red line). Although the sulfides from the hydrated mantle and AOC overlap with these isotopic ranges, fluids with such heterogeneous isotopic compositions should rather correspond to local fluids that remain unmixed. This is inconsistent with the large abundance of hydrous minerals and the Cr, Ni, MgO and δ¹¹B enrichment in the LSZ, which result from large fluid fluxes from serpentine dehydration and point to (and are modeled using) a homogeneous composition (Spandler et al., 2011; Angiboust et al., 2014; Locatelli et al., 2019). It is worth noting that Ohmoto and Rye (1979) experimentally obtained ³⁴αH₂S-pyrite = 1.0006 at 550°C, which should produce ³⁴S-enriched residual H₂S in fluids and hence, ³⁴S-depleted instantaneous pyrites, in opposite direction to our observed trend. Therefore, an open-system precipitation of sulfide from H₂S cannot explain the ³³S-³⁴S-enriched metasomatic sulfides.

600 Radical S₃⁻ ions have been observed in hydrothermal fluids and high temperature experiments. Pokrovski and Dubessy (2015) and Pokrovski and Dubrovinsky (2011) showed that radical S₃⁻ ions are stable above 250°C when the total S content exceeds 0.1-0.5 wt% and could be involved in the pyrite precipitation at peak conditions. However, the
experimental fractionation factor (Kokh et al., 2020) and fractionation factor $ ^{34}\alpha_{S3-/H2S} $ of
0.9988 and $ ^{33}\lambda $ of 0.5158 at 550°C calculated by Eldridge et al. (2021) cannot explain our
data either.

We now investigate whether the $ ^{33}S-^{34}S $-enriched metasomatic sulfides could result from
sulfate reduction in an open system. A Rayleigh distillation model was previously proposed to
explain the $ ^{34}S $ variability of ~ 30‰ in the metasomatic sulfides, i.e. a variability that could
not result from the preservation of the protolith signature heterogeneity (Walters et al., 2019).

Starting here from a single initial composition of $ \delta^{34}S = 12 \pm 3\text{‰} $, $ \Delta^{33}S = 0.025 \pm 0.010\text{‰} $ and $ \Delta^{36}S = 0.085 \pm 0.200\text{‰} $ for sulfates (represented by the blue circle and brown field for source variability on Figs. 5C,D and Fig. 7) and a fractionation factor $ ^{34}\alpha_{SO4-H2S} = 0.9903 $ at 550°C, sulfate reduction in aqueous fluids into $ H_2S $ and precipitated as pyrite would yield $ ^{33}\lambda $ as measured in the sulfides that is consistent with the observation and would cover the whole sulfur isotopic range (Fig. 5C,D and Fig. 7). This initial sulfate $ \delta^{34}S $ signature is also in agreement with the $ \delta^{34}S $ estimates of ~ 14‰ based on a mass balance between serpentinite and chlorite-harzburgite from Cerro del Almirez (Alt et al., 2012). Consequently, a (single) fluid homogeneous in $ \delta^{34}S $ could produce the intragrain and intergrain $ \delta^{34}S $ variability. The sulfate will be reduced to variable extents depending on the electron donor availability, such as $ Fe^{2+} $ liberated following garnet dissolution (see garnet atolls on Fig. S3 and in Faryad et al. (2010), Angiboust et al. (2011) and Angiboust et al. (2014)). For instance, the $ \delta^{34}S $ decrease from core to rim of the pyrite MV can result from the increase in the residual fraction of sulfate in successive fluid pulses, through e.g. a decrease of $ Fe^{2+} $ availability (Fig. 6G). However, evidence for oxidation in response to sulfate reduction should be further investigated, as the $ Fe^{3+} $-enriched rim of omphacite and zoisite in the VI1912 sample could relate to exhumation (Figs. 4E and F). Alternatively, this intragrain variation from core to rim could result from an increase in the rock's $ fO_2 $ that could be buffered by oxidizing fluids,
640 limiting the sulfate-reducing iron-oxidizing reaction, and hence leading to an increase in the
641 fraction of residual sulfate and a decrease in δ²³S at any given time.

642 A fluid isotopic signature of δ¹³S = 12 ± 3‰ and Δ¹³S = 0.025 ± 0.010‰ does not
643 correspond to seawater sulfates, nor to the sulfates measured in the sole serpentinites (see
644 section 6.2.2.). However, the fluids released during antigorite breakdown could oxidize the
645 hydrothermal sulfides in serpentinites into sulfates (Fig. 7) and produce this isotopic
646 signature. This mechanism was proposed to explain the negative correlation between δ⁵⁶Fe
647 and Fe³⁺/ΣFe during the serpentinite prograde metamorphism (Debret et al., 2016; 2021),
648 where sulfide oxidation iron-reduction from magnetite and serpentine reaction leads to
649 Fe³⁺/ΣFe decrease and loss of the lighter isotopes to the fluid under the form of Fe(II)-SO₄
650 complex, following this equation (Debret et al., 2016):

651
652
653
4 Fe₂O₃ (serp-magnetite) + FeS = 8 FeO + Fe-SO₄(aq) (2)

654 Ferric iron incorporated in serpentine and magnetite during peridotite alteration on the
655 seafloor, then released and reduced to Fe²⁺ at high PT during serpentinite dehydration could
656 thus be a viable mechanism to produce sulfates in aqueous fluids. Thermodynamic
657 calculations provide other evidence for the sulfate occurrence in HP aqueous fluids (e.g.
658 Debret and Sverjensky, 2017; Connolly and Galvez, 2018). Recently, thermodynamic
659 calculations of the position of the SSO buffer, i.e. the H₂S-SO₂ equal activity line, at different
660 PT conditions in serpentinites showed that for T < 550°C, sulfates are not stable in
661 equilibrium with magnetite only, Hematite-Magnetite buffer (HM) has to be reached; whereas
662 at T > 550°C, sulfates could be stable in equilibrium with magnetite only, as the SSO buffer is
663 reached before the HM buffer in the oxidized serpentinites (Evans and Frost, 2021). In the
664 Monviso specifically, the serpentinites close to the LSZ contain awaruite and pentlandite,
665 while the sole serpentinites contain magnetite and haezlewoodite, indicating a redox gradient
666 from ΔFMQ in the latter to ΔFMQ ~ -4 in the former (Caurant et al., 2023). Thus, it is likely
that sulfide oxidation in the sole serpentinites during antigorite breakdown could have formed
667
sulfates, certainly under Fe(II)-SO₄ complex. Then, the sulfates could have been
668
subsequently reduced when circulating along the LSZ between the serpentinites and the
669
metagabbros. It is worth noting that sulfate observation in the form of anhydrite and/or
670
gypsum in the Monviso eclogite fluid inclusions are evidence of their stability in the
671
metagabbros (Philippot and Selverstone, 1990), implying that sulfates in aqueous fluids are
672
not totally buffered by the reducing LSZ serpentinites.
673

The antigorite breakdown requires temperature conditions of 600-700°C (Ulmer and Trommsdorff, 1995), implying that it should occur deeper in the slab (the peak conditions at the Monviso locality being only at ~ 550°C and ~ 2.7 GPa). Considering serpentinites at T = 650°C with a S content of 40 ppm as sulfides as observed in the Monviso sole serpentinite VI1026, D(S)rock-fluid= 0.25 and ³⁴αH₂S-SO₄ = 1.0078, the partially oxidized sulfides should have an initial signature averaging at δ¹³S = 7 ± 3‰, Δ¹³S = 0.020 ± 0.010‰ and Δ¹³S ~ 0.100 ± 0.200‰ (green square on Fig. 5C and Fig. 7, represented by dashed and solid curves for the produced sulfates and the residual pyrites in the serpentinites respectively) to obtain the above sulfate-bearing fluids signature after releasing 12 wt% H₂O from antigorite. This signature is consistent with hydrothermal sulfide signature from the literature (Alt et al., 2012a; Alt et al., 2012b; Alt et al., 2013; Alt and Shanks III, 2003; Alt and Shanks, 1998; Delacour et al., 2008a; Delacour et al., 2008b; Ono et al., 2012; Schwarzenbach et al., 2012; Schwarzenbach et al., 2018a; Schwarzenbach et al., 2018b; see Fig. 1) and within uncertainty of the Monviso sole serpentinite VI1026 δ¹³S -Δ¹³S value (Fig. 5C).

It is worth noting that the partial oxidation of sulfides during prograde fluid loss leads to a
689
restricted isotopic fractionation at 650°C < 3‰ for δ¹³S and negligible Δ¹³S-Δ¹³S variation
690
between the residual sulfides and the produced sulfates, either by pulses or continuous fluid
691
flow (dashed and solid curves on Figs. 5C and 7; Hoover et al., 2022). In other words, this
692
693 mechanism produces sulfate-bearing fluids enriched in <sup>34</sup>S and residual sulfides depleted in
694 <sup>34</sup>S by less than 3‰, consistent with the limited isotopic fractionations between sulfide and
695 fluids observed by Walters et al. (2019), but without requiring sluggish S self-diffusion. This
696 small isotopic fractionation is then hardly detectable considering the sulfur isotopic variability
697 of the slab reservoirs. In this case, the residual sulfides in the serpentinites would then have
698 δ<sup>34</sup>S = 4 ± 3‰, Δ<sup>33</sup>S = 0.020 and Δ<sup>36</sup>S = 0.085‰, considering a dehydration by pulses
699 (white square, Figs. 5C and 7).

700 To summarize, we propose here that the partial oxidation of sulfides in the serpentinites
701 during antigorite dehydration at 650°C led to the formation of sulfate-bearing fluids with initial
702 δ<sup>34</sup>S = 12 ± 3‰, Δ<sup>33</sup>S = 0.025 ± 0.010‰ and Δ<sup>36</sup>S = 0.085 ± 0.200‰. These sulfate-
703 bearing fluids were partially reduced to form mostly pyrite in an open system at 550°C,
704 resulting in the precipitation of <sup>33</sup>S-<sup>34</sup>S-enriched metasomatic sulfides in the metasomatized
705 Fe-Ti metagabbros and the hybrid schists (T3) compared to their protoliths, i.e. the
706 mylonitized eclogites and the sole serpentinites (T1 and T2; Fig. 7). The information carried
707 by δ<sup>34</sup>S-Δ<sup>33</sup>S-Δ<sup>36</sup>S systematics thus provide a new and independent evidence for the
708 occurrence of sulfates in aqueous fluids, which could be ultimately transported to the sub-arc
709 mantle. This model does not rule out the possibility of a mixing between a fluid with partially
710 reduced sulfates (δ<sup>34</sup>S = 15-21‰) and sulfur from sediments or mafic rocks (Figs. 5B and
711 D).

712

713 6.3. Implications for sulfur cycling in sub-arc and deep mantle

714

715 The δ<sup>34</sup>S-Δ<sup>33</sup>S values preserved by the metasomatized sulfides provide a new evidence for
716 sulfate stability in HP aqueous fluids at least down to 550°C, the peak temperature of the
717 Lago Superiore eclogites, consistent with evidence of oxidant aqueous fluids in the sub-arc
718 mantle inferred by xenoliths, melt inclusions and volcanic gases (Bénard et al., 2018; Blatter
and Carmichael, 1998; de Moor et al., 2022; McInnes et al., 2001; McInnes et al., 1999; Muth and Wallace, 2021; Widom et al., 2003). Compared with melt inclusions and glasses from arc lavas having δ³⁴S from ~ -9 to 21‰ (Woodhead et al., 1987; Alt et al., 1993; De Hoog et al., 2001; Bouvier et al., 2008; Muth and Wallace, 2021), our model can explain most ³⁴S-enriched signatures without requiring isotopic fractionation during S degassing or other shallow processes. Thus, a mixing between a depleted mantle source (δ³⁴S ~ -1 ± 0.5‰; Labidi et al., 2013) and sulfates from slab-derived fluids could produce the arc lavas δ³⁴S values from -1 to 12 ± 3‰, while a higher proportion of ³⁴S-depleted sediments in the source could explain δ³⁴S values lower than -1‰ (Bouvier et al., 2008). The basalts with δ³⁴S up to ~ 21‰ would then require a source having at least a δ³⁴S ~ 18‰, considering an isotopic fractionation lower than 3‰ during partial oxidation. Given that sulfides with this signature are not abundant in the literature (Fig. 1), δ³⁴S up to 21‰ could result from additional processes, such as (1) kinetic degassing occurring during the eruption (Aubaud et al., 2004; Labidi et al., 2015) or (2) open system sulfate reduction in the aqueous fluids or hydrous melt being transported in a sub-arc mantle still reduced such as at the onset of subduction (Brounce et al., 2021). Indeed, evidence for both oxidizing aqueous fluids (e.g. Blatter and Carmichael, 1998; McInnes, 1999; 2001; Widom et al., 2003) associated with ³⁴S-enriched values in mantle xenoliths and melt inclusions (Rielli et al., 2018; Bénard et al., 2018, Muth and Wallace, 2021) and mantle xenoliths with fO₂ similar to the depleted mantle have been observed (Song et al., 2009; Wang et al., 2008). Providing that sulfates are not totally reduced by circulating through the reducing LSZ serpentinites and mafic crust, as observed in arc xenoliths (see Blatter and Carmichael, 1998; McInnes et al., 2001), the transport of sulfates in a sub-arc mantle already oxidized or still reduced will not result in the same isotopic composition and S budget. If the mantle is oxidized, i.e. on or above the SSO (H₂S-SO₄ equal activity line) buffer, the sulfate-bearing fluid can preserve its signature of δ³⁴S = 12 ± 3‰. However, if the mantle is reduced, as observed in fore-arc basalts at the onset of subduction (Brounce et al., 2021), we speculate that sulfates will be continuously reduced
(following a Rayleigh distillation process) when reacting with Fe²⁺, resulting in even higher δ³⁴S and Δ³³S of residual sulfates, until the SSO buffer is reached (Mungall, 2002). The isotopic variability could thus result from both the degree of sulfate reduction and the source signature. The way sulfur is transported, by diapir, focused or pervasive flow and hence the water (or melt)/rock ratio (Hall and Kincaid, 2001; Pirard and Hermann, 2015) is also important because the extent of equilibration will be different and remains to be investigated.

The residual sulfides in the downgoing slab will be recycled in the deep mantle. The ocean island basalts probing the deep mantle reservoirs have δ³⁴S values in the range of “depleted” MORB δ³⁴S signature (~ -1‰) or slightly enriched up to ~ 3‰, implying a recycled component with positive δ³⁴S values (Dottin III et al., 2020b; Labidi et al., 2015; Labidi et al., 2022). We specify that, here we do not consider few samples having lower δ³⁴S values from ~ -17 to -2‰ than MORB because the data were obtained on lavas, not glasses, which could be degassed. We also discard the sulfides measured using ion-probe, which show a significant instrumental variability (up to 1.67‰ in 2σ) (Cabral et al., 2013; Delavault et al., 2016; Dottin III et al., 2020a), certainly veiling the source signature. Considering the isotopic variability of the subducted protoliths (Fig. 1), the limited isotopic fractionation < 3‰ produced during the partial sulfide oxidation model still allows the recycling of a component with a positive δ³⁴S signature, in agreement with the ³⁴S-enriched OIB glasses compared to MORB.

## 7. Conclusion

The Monviso metaophiolite constitutes a unique and coherent portion of subducted lithosphere that allows studying fluid circulation through the slab at eclogite-facies conditions. The analyses and comparison of the non-metasomatized and metasomatized samples have highlighted ³³S-³⁴S-enrichment of metasomatic sulfides in veins and metasomatic rinds
compared to their host-rock. Such $ ^{33}S-^{34}S $-enriched isotopic compositions are inconsistent
with abiotic or biotic sulfides formed on the seafloor. The isotopic enrichment rather reflects
sulfate reduction in an open-system, implying the occurrence of sulfates in HP fluids with an
initial $ ^{34}S $ of ~ 12‰. This signature does not correspond to pristine seawater sulfates and
rather points to the oxidation of hydrothermal sulfides from hydrated mantle. Therefore, we
suggest that sulfates were produced during the oxidation of sulfides triggered by antigorite
dehydration liberating $ Fe^{3+} $ at greater depth in the slab. These sulfate-bearing fluids traveled
along intra-slab shear zones parallel to the subduction interface until their subsequent
reduction (into pyrite) when they met the Monviso metaaphiolite at peak burial conditions.

Last, sulfate-bearing fluids with $ ^{34}S $ of 12‰ can account for the occurrence of oxidized and
$ ^{34}S $-enriched arc lavas, except for values up to 21‰ that may require additional processes.

The residual sulfides in the slab would also carry a signature in agreement with S isotope
signature recorded in OIB, reflecting S isotopic fractionation (< 3‰) during the oxidation.

786

787 8. Acknowledgments

788

We thank the editor Esther Schwarzenbach, Jesse Walters and two anonymous reviewers
for their constructive comments. We are also grateful to Maryjo Brounce for helping in
clarifying our manuscript. We also thank Stefan Borensztajn for assistance during SEM
analyses, Nicolas Rividi and Michel Fialin for their assistance during EPMA analyses,
Guillaume Landais, Gérard Bardoux and Amaury Bouyon for their help during isotopic
analyses at IPGP, Guillaume Elise for sample polishing as well as Lithologie Bourgogne for
thin section preparation, and Emilie Thomassot and Baptiste Debret for discussions. We are
very grateful for the help of Nordine Bouden who performed ion-probe analyses during the
pandemic at LG-SIMS Nancy (CRPG - CNRS / Univ. Lorraine). This study was supported by
Grants ANR-10-BLAN-0603 and ANR-15-CE31-005 of Agence Nationale de la Recherche,
France.

800
801 Credit Author Statement:

802
803 Isabelle Genot: Conceptualization, Methodology, Validation, Investigation, Writing – Original draft, Writing – Review and Editing, Visualization. Samuel Angiboust: Conceptualization, 804 Methodology, Validation, Investigation, Writing – Original draft, Writing – Review and Editing, 805 Visualization, Resources. Pierre Cartigny: Conceptualization, Methodology, Validation, 806 Investigation, Writing – Original draft, Writing – Review and Editing, Visualization, 807 Resources. 808

809
810 9. Appendix A: Supplementary materials

811
812 Figure S1 shows the mass fractionation line in a δ¹³S-δ¹³S plot. Figure S2 presents the bulk 813 rock trace element patterns of HP metasediments, mafic and ultramafic rocks. Figure S3 814 shows chemical maps of the eclogite breccia VI1912. Figure S4 shows the Ni vs Co contents 815 in pyrites.

816 Table S1 reports the bulk major and trace element contents in HP metasediments, mafic and 817 ultramafic rocks. Table S2 shows Cu, Fe, S, Co and Ni contents in pyrites. Table S3 presents 818 the mean S isotopic compositions of international standard IAEA S1. Table S4 reports the in- 819 situ δ³S of pyrites using ion-probe. And Table S5 shows the composition of minerals in 820 eclogite breccias and hybrid schists.

821
822 10. Data availability

823
824 Data are available through Mendeley data repository at:
825 https://doi.org/10.17632/3fs2386t2s.1

826
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1263
1264
Figure captions

1265
1266

**Figure 1:** Sulfur isotopic composition compilation of sulfides and sulfates in the sediments, altered oceanic crust and hydrated mantle. (A) δ³⁴S variability of Phanerozoic sediments sulfides. (B) δ³⁴S variability of sulfides and sulfates in the altered oceanic crust. (C) δ³⁴S variability of sulfides and sulfates in seafloor serpentinites. (D) δ³⁴S-Δ³³S and (E) Δ³³S-Δ³⁶S variability of sulfides in AOC and seafloor serpentinites. Data are from Alford et al. (2011); Alt (1994); Alt et al. (1989); Alt et al. (2012a); Alt et al. (2012b); Alt et al. (2013); Alt and Shanks (2003), Alt and Shanks (1998); Alt et al. (2007); Alt and Shanks (2011); Delacour et al. (2008a); Delacour et al. (2008b); Ono et al. (2012); Rouxel et al. (2008); Schwarzenbach et al. (2012); Schwarzenbach et al. (2018a); Schwarzenbach et al. (2018b). Δ³³S and Δ³⁶S values of Ono et al. (2012) have been recalculated in the conventional notation.

1277
1278
**Figure 2:** (A) Global geological map of the Liguro-Piemontese domain in the Western Alps. SL: Schistes Lustrés; DM: Dora Maira massif. Samples VI1102-d and VI1018 are outside the map (see GPS coordinates in Table 1). (B) Geological map of the Lago Superiore Unit and location of samples. LSZ: Lower Shear Zone, BSZ: Basal Shear Zone, ISZ: Intermediate Shear Zone, USZ: Upper Shear Zone. Modified after Rubatto & Angiboust (2015). (C) Cross-section view (dashed line in (B)) of the Lago Superiore Unit. (D) P-T evolution of LSZ samples with T1: undeformed blueschist (BS) and eclogites (Ecl), T2: mylonitized eclogites, T3: metasomatized eclogites. Modified after Angiboust et al. (2014).

1286
1287
**Figure 3:** Hand-specimen and field views of sulfides in garnetite, i.e. metasomatized Fe-Ti metagabbro (A), (B) and hybrid schists (C), (D), (E). Lws ps.: Lawsonite pseudomorph of Omp : omphacite ; Serp : serpentinite ; Py : pyrite.

1290
**Figure 4:** (A) Photomicrograph of sulfides disseminated in both garnet(Gt)-omphacite(omp) matrix, with rutile and chlorite, and omphacite vein in sample VI1906. Chemical analysis was not performed to determine the type of sulfide. (B) SEM image of pyrites disseminated in omphacite vein, with apatite and zircon, in sample MV. (C) SEM image of pyrites in metasomatic rind between omphacite vein and Gt-omp matrix of sample VI1912. (D) Photomicrograph of pyrites and omphacite veins in VI1912 (crossed polar). (E) SEM image zoom on metasomatic pyrite in sample VI1912 with inclusions of chlorite (Chl), garnet (Gt) and omphacite, in equilibrium with chlorite and omphacite rim enriched in Fe³⁺ (omp2). (F) SEM image of lawsonite replaced by paragonite (pg), phengite (phg), zoisite (zo1 and zo2), Fe³⁺-enriched rim) in sample VI1912.

1301

**Figure 5:** (A) δ¹³⁴S and Δ¹³³S values of metamorphic (T1 and T2 stages) and metasomatic (T3) sulfides. Dashed lines represent the composition for mass exponents ³³λ of 0.5153 and 0.5189 and delineate our metasomatic sulfides composition. The gray field represents isotopic compositions that should result from processes out of equilibrium (i.e. kinetic isotope effects). Error bars represent the analytical (chemistry and analysis) reproducibility in 1σ for each sample. Standard reproducibility is indicated when samples have not been duplicated. (B) Mixing model, represented by convex curves (see Ono et al., 2012), between depleted components, i.e. sediments (δ¹³⁴S = -12‰, Δ¹³³S = 0.046‰, similar to sample VI1919) and metagabbros (δ¹³⁴S = 1‰, Δ¹³³S from 0.020 to 0.028‰) and the enriched component that should have δ¹³⁴S and Δ¹³³S values from 15 to 21‰ and 0.062 and 0.049‰ respectively to encompass all the samples. (C) Sulfate reduction model in an open system from an initial sulfate-bearing fluid having δ¹³⁴S = 12 ± 3‰ and Δ¹³³S = 0.025 ± 0.01‰ (represented by the blue circle and brown field for source isotopic variability), with an isotopic fractionation ³⁴αₕ₀₄- H₂S = 0.9903 at T = 550°C. The instantaneous pyrite composition resulting from reaction of H₂S with FeSₐq is also represented by the red line. The solid gray curve represents the composition of residual sulfides in serpentinites following their oxidation into sulfates by the
1318 release of 12 wt% H₂O (white square) during the antigorite breakdown at T = 650°C with
1319 ³⁴αH₂S-SO₄ = 1.0078 and using a demixing calculation between the residual sulfides and the
1320 produced sulfate-bearing fluids (dashed gray curve). Initial hydrothermal sulfides should have
1321 δ¹³⁴S = 7‰ and Δ¹³³S = 0.02‰ (green square) to produce the above sulfate-bearing fluids
1322 composition. F is the residual fraction of sulfates. (D) Δ¹³³S-Δ¹³⁶S values of the metamorphic
1323 and metasomatic sulfides analyzed in this study, in comparison with those from the literature
1324 recalculated in the logarithmic notation (see Fig. 1 for symbols) and open-system model for
1325 sulfate and H₂S reduction. The source isotopic variability of 0.01 and 0.2‰ for Δ¹³³S and
1326 Δ¹³⁶S is represented by the brown field.
1327

1328 **Figure 6:** In-situ δ¹³⁴S points (in ‰, white dots) and Co, Ni X-ray maps of sulfides in sample
1329 VI1914 (A), VI1102-d (B), VI1912 (C and D), MV (E and F), VI1018 (G), VI1924 (H and I). (J):
1330 Simplified explanation of intragrain δ¹³⁴S variation resulting from variable reduction potential,
1331 through electron donor availability such as Fe²⁺.
1332

1333 **Figure 7:** Scheme of processes leading to the formation of ³³S-³⁴S-enriched metasomatic
1334 pyrites in the LSZ. 1) Sulfide oxidation occurs during antigorite breakdown at 650°C, allowed
1335 by magnetite and antigorite buffering, from which Fe³⁺ is reduced and produces little
1336 fractionated Fe(II)SO₄-bearing fluids (³⁴αH₂S-SO₄ = 1.0078). 2) These sulfate-bearing fluids are
1337 then reduced in an open-system at 550°C in the LSZ, following the dissolution of Fe²⁺-
1338 bearing minerals such as garnet, precipitating pyrites and leading to continuous ³³S-³⁴S-
1339 enrichment of residual sulfates (³⁴αSO₄-H₂S = 0.9903), and hence instantaneous produced
1340 pyrites (modified after Angiboust et al., 2014).
1341
1342
1343
1344
1345
1346
<figure><img src="image_2.png" /><figcaption>Figure 1</figcaption></figure>
<figure><img src="image_3.png" /><figcaption>Figure 2</figcaption></figure>

1371

1372

1373

1374

1375

1376

1377

1378

1379

1380

1381

1382

1383

1384

1385

1386

1387
<figure><img src="image_4.png" /><figcaption>Figure 3</figcaption></figure>

1388

1389

1390

1391

1392

1393

1394

1395

1396

1397

1398

1399

1400

1401

1402

1403

1404

1405

1406

1407

1408

1409

1410
<figure><img src="image_5.png" /><figcaption>Figure 4</figcaption></figure>

1411

1412

1413

1414
<figure><img src="image_6.png" /><figcaption>Figure 5</figcaption></figure>

1415

1416

1417

1418

1419

1420

1421

1422

1423

1424

1425

1426

1427

1428

1429
<figure><img src="image_7.png" /><figcaption>Figure 6</figcaption></figure>

1430

1431

1432

1433

1434

1435

1436

1437

1438

1439

1440

1441

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1443

1444

1445

1446

1447

1448
<figure><img src="image_8.png" /><figcaption>Figure 7</figcaption></figure>

1449
1450

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1453
1454
1455
1456
<table><thead><tr><th>Sample</th><th>Stages</th><th>Rock type</th><th>Matrix mineral assemblage</th><th>Locality</th><th>Latitude</th><th>Longitude</th></tr></thead><tbody><tr><td>Pelvat (PV)</td><td>T1</td><td>Greenschist-facies metabasalt</td><td>Ep, Chl, Tremolite, Calcite, disulfides</td><td>Pic du Pelvat, Ubaye Valley, France</td><td>N 44° 34' 51"</td><td>E 06° 55' 07"</td></tr><tr><td>Queyras (QM)</td><td>T1</td><td>Blueschist-facies metagabbro</td><td>Gln, Lawsonite, relict cpx, jadeite, disulfides</td><td>Château-Queyras, France</td><td>block</td><td>block</td></tr><tr><td>VI1102-d</td><td>T1</td><td>Undeformed Fe-Ti metagabbro</td><td>Omp, Gt, Rt and ilmenite, pyrite</td><td>Lago di Luca (South Monviso)</td><td>N 44° 37' 07"</td><td>E 07° 08' 40"</td></tr><tr><td>SL1801</td><td>T1</td><td>Metapelite</td><td>Phg, Chl, Qz, disulfides</td><td>Forest-Vetan (South Monviso)</td><td>N 44° 36' 18"</td><td>E 07° 09' 27"</td></tr><tr><td>VI1003</td><td>T1</td><td>Metapelite</td><td>Phg, Chl, Qz, Gt, disulfides</td><td>500 m W. Rifugio Sella, Monviso</td><td>N 44° 39' 29"</td><td>E 07° 06' 28"</td></tr><tr><td>VI1913-16</td><td>T2</td><td>Mylonitized Fe-Ti metagabbro</td><td>Gt, Omp, Rt, Qz, disulfides (VI1914: pyrite)</td><td>Traversette path, Pian del Re</td><td>N 44° 41' 55"</td><td>E 07° 05' 44"</td></tr><tr><td>VI01-02</td><td>T2</td><td>Sole serpentinite</td><td>Ant, Mgt, sulfates</td><td>Pian del Re</td><td>N 44° 42' 00"</td><td>E 07° 06' 04"</td></tr><tr><td>VI1026</td><td>T2</td><td>Sole serpentinite</td><td>Ant, Mgt, disulfides</td><td>Monte Grané</td><td>N 44° 41' 24"</td><td>E 07° 07' 17"</td></tr><tr><td>VI1904</td><td>T3</td><td>Eclogite breccia, Fe-Ti metagabbro</td><td>Vein1: Omp, disulfides; Retro: vein 2 gln, Ep</td><td>Ghincia Pastour, Monviso</td><td>N 44° 40' 53"</td><td>E 07° 06' 44"</td></tr><tr><td>VI1906</td><td>T3</td><td>Eclogite breccia, Fe-Ti metagabbro</td><td>Matrix: Gt, Omp, Rt, Chl, disulfides; Vein: Omp, disulfides</td><td>Ghincia Pastour</td><td>N 44° 40' 53"</td><td>E 07° 06' 44"</td></tr><tr><td>VI1907</td><td>T3</td><td>Eclogite breccia, Fe-Ti metagabbro</td><td>Vein: dissolved Gt and Omp, Chl, Tlc, Apatite, disulfides</td><td>Ghincia Pastour</td><td>N 44° 40' 53"</td><td>E 07° 06' 44"</td></tr><tr><td>VI1912</td><td>T3</td><td>Eclogite breccia, Fe-Ti metagabbro</td><td>Matrix: Gt (atoll), Omp, Rt, pyrite; Vein1: Omp; Vein2: Omp, Chl, Phg</td><td>Ghincia Pastour</td><td>N 44° 40' 53"</td><td>E 07° 06' 44"</td></tr><tr><td>VI1018</td><td>T3</td><td>Tlc-bearing omphacitite</td><td>Omp, Tlc, Chl, pyrite</td><td>Punta Murel</td><td>N 44° 38' 48"</td><td>E 07° 08' 01"</td></tr><tr><td>MV2</td><td>T3</td><td>Eclogite breccia, Fe-Ti metagabbro</td><td>Vein: Omp, Chl, Psd lawsonite, pyrites</td><td>Ghincia Pastour</td><td>N 44° 40' 53"</td><td>E 07° 06' 44"</td></tr><tr><td>MV3</td><td>T3</td><td>Eclogite breccia, Fe-Ti metagabbro</td><td>Vein: Gt, Omp, pyrites</td><td>Ghincia Pastour</td><td>N 44° 40' 53"</td><td>E 07° 06' 44"</td></tr><tr><td>VI1908</td><td>T3</td><td>Hybrid schist</td><td>Chl, Tlc, disulfides</td><td>Ghincia Pastour</td><td>N 44° 40' 53"</td><td>E 07° 06' 44"</td></tr><tr><td>VI1920-24</td><td>T3</td><td>Hybrid schist</td><td>Chl, Ant, Tremolite, Ep, Tlc, pyrites</td><td>Lago Chiaretto</td><td>N 44° 41' 20"</td><td>E 07° 05' 52"</td></tr><tr><td>VI1925-27</td><td>T3</td><td>Hybrid schist</td><td>Chl, Tlc, Tremolite, Pyrites</td><td>Lago Chiaretto</td><td>N 44° 41' 20"</td><td>E 07° 05' 52"</td></tr></tbody></table>
<table><tr><td>LSZ05</td><td>T3</td><td>Hybrid schist</td><td>Cpx, Magnesite, Tlc, Chl, actinolite, Na-Ca clinopyroxene, Mgt, disulfides</td><td>Punta Forcione</td><td>N 44° 40' 37"</td><td>E 07° 07' 34"</td></tr><tr><td>VI1910</td><td>T3</td><td>Serpentinite</td><td>Ant, Mgt, sulfides, sulfates</td><td>Ghincia Pastour</td><td>N 44° 40' 53"</td><td>E 07° 06' 44"</td></tr><tr><td>VI1911</td><td>T3</td><td>Serpentinite</td><td>Ant, Mgt, sulfates</td><td>Ghincia Pastour</td><td>N 44° 40' 53"</td><td>E 07° 06' 44"</td></tr><tr><td>VI1917a</td><td>T3</td><td>Serpentinite</td><td>Ant, Mgt, Plag; Vein: cpx, titanoclinohumite, disulfides</td><td>Pian del Re</td><td>N 44° 42' 09"</td><td>E 07° 05' 47"</td></tr><tr><td>VI1928</td><td>T3</td><td>Serpentinite</td><td>Ant, Mgt, lawsonite, dolomite, disulfides, sulfates</td><td>Lago Fiorenza</td><td>N 44° 41' 53"</td><td>E 07° 05' 34"</td></tr><tr><td>VI1918</td><td>T3</td><td>Quartzite</td><td>Qz, Chl, Phg, Lws (psd.), pyrites</td><td>Lago Fiorenza</td><td>N 44° 41' 31"</td><td>E 07° 05' 50"</td></tr><tr><td>VI1919</td><td>T3</td><td>Metapelite</td><td>Chl, Phg, Qz, Lws (psd.), disulfides</td><td>Lago Fiorenza</td><td>N 44° 41' 35"</td><td>E 07° 05' 40"</td></tr></table>

**Table 1:** Mineralogical description of collected samples with their location. Cpx : clinopyroxene ; Gt : garnet ; Omp : omphacite ; Rt : rutile ; Qz : quartz ; Ant : antigorite ; Mgt : magnetite ; Chl : chlorite ; Tlc : talc ; Gln : glaucophane ; Phg : phengite. Pyrite is indicated when the sulfide was identified by chemical analysis (see Table S2). Instead, disulfide is indicated when only chemical extraction allowed us to discriminate between Acid Volatile Sulfide (AVS) phase, i.e. monosulfide or Chromium Reducible Sulfide (CRS) phase, i.e. disulfide.
<table><caption>Table 2</caption><thead><tr><th></th><th>Sulfur ppm</th><th>δ<sup>34</sup>S V-CDT ‰</th><th>1σ</th><th>Δ<sup>33</sup>S V-CDT ‰</th><th>1σ</th><th>Δ<sup>36</sup>S V-CDT ‰</th><th>1σ</th></tr></thead><tbody><tr><td colspan="8"><strong>Metabasalt-metagabbro (T1)</strong></td></tr><tr><td>Pelvat PV*</td><td></td><td>1.74</td><td>0.07</td><td>0.030</td><td>0.007</td><td>-0.189</td><td>0.134</td></tr><tr><td>Queyras<br>QM*</td><td></td><td>0.56</td><td>0.07</td><td>0.011</td><td>0.007</td><td>-0.075</td><td>0.134</td></tr><tr><td colspan="8"><strong>Metasediments (T1-T2)</strong></td></tr><tr><td>SL1801</td><td>1231</td><td>-20.54</td><td>0.21</td><td>-0.015</td><td>0.003</td><td>0.403</td><td>0.165</td></tr><tr><td>VI1003</td><td>69</td><td>-5.12</td><td>0.07</td><td>0.005</td><td>0.007</td><td>0.032</td><td>0.134</td></tr><tr><td colspan="8"><strong>Mylonitized Fe-Ti metagabbro (T2)</strong></td></tr><tr><td>VI1913</td><td>969</td><td>0.29</td><td>0.04</td><td>0.026</td><td>0.007</td><td>0.194</td><td>0.064</td></tr><tr><td>VI1914</td><td>7017</td><td>1.23</td><td>0.16</td><td>0.020</td><td>0.006</td><td>0.171</td><td>0.016</td></tr><tr><td>VI1915</td><td>4485</td><td>0.20</td><td>0.07</td><td>0.026</td><td>0.007</td><td>-0.109</td><td>0.134</td></tr><tr><td>VI1916</td><td>4866</td><td>1.33</td><td>0.07</td><td>0.011</td><td>0.007</td><td>-0.018</td><td>0.134</td></tr><tr><td colspan="8"><strong>Metasomatized Fe-Ti metagabbros (T3)</strong></td></tr><tr><td>VI1904</td><td>280</td><td>4.9</td><td>0.75</td><td>0.026</td><td>0.001</td><td>-0.168</td><td>0.003</td></tr><tr><td>VI1906</td><td>2972</td><td>7.06</td><td>1.11</td><td>0.024</td><td>0.004</td><td>0.026</td><td>0.044</td></tr><tr><td>VI1907*</td><td></td><td>13.2</td><td>0.31</td><td>0.032</td><td>0.006</td><td>0.117</td><td>0.105</td></tr><tr><td>VI1912</td><td>3474</td><td>8.59</td><td>0.2</td><td>0.037</td><td>0.003</td><td>-0.11</td><td>0.064</td></tr><tr><td>MV2-1*</td><td></td><td>10.07</td><td>0.1</td><td>0.026</td><td>0.003</td><td>-0.124</td><td>0.019</td></tr><tr><td>MV2-2*</td><td></td><td>9.9</td><td>0.13</td><td>0.03</td><td>0.004</td><td>-0.083</td><td>0.047</td></tr><tr><td>MV2-3*</td><td></td><td>7.76</td><td>0.17</td><td>0.023</td><td>0.002</td><td>-0.003</td><td>0.018</td></tr><tr><td>MV2-4*</td><td></td><td>14.36</td><td>0.01</td><td>0.05</td><td>0.015</td><td>-0.131</td><td>0.007</td></tr><tr><td>MV3-1*</td><td></td><td>11.49</td><td>0.07</td><td>0.041</td><td>0.007</td><td>-0.085</td><td>0.134</td></tr><tr><td>MV3-2*</td><td></td><td>11.72</td><td>0.07</td><td>0.038</td><td>0.002</td><td>-0.177</td><td>0.003</td></tr><tr><td>MV3-3*</td><td></td><td>11.28</td><td>0.07</td><td>0.039</td><td>0.007</td><td>-0.054</td><td>0.134</td></tr><tr><td colspan="8"><strong>Hybrid schists (T3)</strong></td></tr><tr><td>VI1908</td><td>179</td><td>10.29</td><td>0.71</td><td>0.026</td><td>0.003</td><td>-0.068</td><td>0.252</td></tr><tr><td>VI1920</td><td>311</td><td>10.82</td><td>0.36</td><td>0.04</td><td>0.003</td><td>0.107</td><td>0.207</td></tr><tr><td>VI1921*</td><td></td><td>11.56</td><td>0.11</td><td>0.044</td><td>0.011</td><td>0.078</td><td>0.178</td></tr><tr><td>VI1922</td><td>114</td><td>8.57</td><td>0.16</td><td>0.040</td><td>0.002</td><td>0.188</td><td>0.224</td></tr><tr><td>VI1923</td><td>384</td><td>7.79</td><td>0.61</td><td>0.026</td><td>0.01</td><td>0.052</td><td>0.246</td></tr><tr><td>VI1924</td><td>1866</td><td>14.46</td><td>0.41</td><td>0.062</td><td>0.012</td><td>-0.134</td><td>0.145</td></tr><tr><td>VI1925*</td><td></td><td>9.48</td><td>0.63</td><td>0.038</td><td>0.003</td><td>0.005</td><td>0.347</td></tr><tr><td>VI1926*</td><td></td><td>21.1</td><td>0.47</td><td>0.049</td><td>0.007</td><td>-0.066</td><td>0.015</td></tr><tr><td>VI1927*</td><td></td><td>12.03</td><td>0.51</td><td>0.049</td><td>0.006</td><td>-0.16</td><td>0.027</td></tr><tr><td>LSZ05*</td><td></td><td>8.67</td><td>0.07</td><td>0.025</td><td>0.007</td><td>0.31</td><td>0.134</td></tr></tbody></table>

Metasediments (T3)
<table><tr><td>VI1918</td><td>2180</td><td>0.93</td><td>0.07</td><td>0.003</td><td>0.007</td><td>0.081</td><td>0.134</td></tr><tr><td>VI1919</td><td>591</td><td>-12.16</td><td>0.5</td><td>0.046</td><td>0.001</td><td>0.321</td><td>0.093</td></tr></table>

**Table 2:** Sulfur content and isotopic composition of sulfides (expressed in logarithmic notation) in metasediments, mafic rocks and hybrid schists. Asterisk means that sulfides were picked, instead of bulk chemistry. Uncertainties represent the analytical (chemistry and analysis) reproducibility. Uncertainty of samples not duplicated corresponds to the standard IAEA S1 reproducibility in italic (see Table S3).
<table><caption>Table 3</caption><thead><tr><th></th><th>Sulfide<br>ppm</th><th>δ<sup>34</sup>S V-CDT ‰</th><th>1σ</th><th>Δ<sup>33</sup>S V-CDT ‰</th><th>1σ</th><th>Δ<sup>36</sup>S V-CDT ‰</th><th>1σ</th><th>Sulfate<br>ppm</th><th>δ<sup>34</sup>S V-CDT ‰</th><th>1σ</th><th>Δ<sup>33</sup>S V-CDT ‰</th><th>1σ</th><th>Δ<sup>36</sup>S V-CDT ‰</th><th>1σ</th></tr></thead><tbody><tr><td colspan="15" style="text-align: left; font-weight: bold;">Sole serpentinites (T2)</td></tr><tr><td>VI1901</td><td>nd</td><td></td><td></td><td></td><td></td><td></td><td></td><td>12</td><td>-4.04</td><td>0.07</td><td>0.041</td><td>0.007</td><td>-0.301</td><td>0.134</td></tr><tr><td>VI1902</td><td>nd</td><td></td><td></td><td></td><td></td><td></td><td></td><td>26</td><td>-2.49</td><td>0.07</td><td>0.035</td><td>0.007</td><td>-0.245</td><td>0.134</td></tr><tr><td>VI1026</td><td>39</td><td>7.16</td><td>0.07</td><td>0.006</td><td>0.007</td><td>0.198</td><td>0.134</td><td>nd</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td colspan="15" style="text-align: left; font-weight: bold;">LSZ serpentinites (T3)</td></tr><tr><td>VI1910</td><td>64</td><td>6.85</td><td>0.51</td><td>0.044</td><td>0.001</td><td>0.262</td><td>0.582</td><td>59</td><td>0.18</td><td>0.07</td><td>0.033</td><td>0.007</td><td>-0.248</td><td>0.134</td></tr><tr><td>VI1911</td><td>nd</td><td></td><td></td><td></td><td></td><td></td><td></td><td>14</td><td>-3.67</td><td>0.07</td><td>0.031</td><td>0.007</td><td>-0.316</td><td>0.134</td></tr><tr><td>VI1917a</td><td>64</td><td>15.85</td><td>0.12</td><td>0.038</td><td>0.004</td><td>-0.023</td><td>0.018</td><td>nd</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>VI1928</td><td>20</td><td>14.21</td><td>0.05</td><td>0.044</td><td>0.007</td><td>-0.055</td><td>0.003</td><td>72</td><td>4.86</td><td>0.07</td><td>0.034</td><td>0.007</td><td>-0.216</td><td>0.134</td></tr></tbody></table>

**Table 3:** Sulfur content and sulfur isotopic composition of sulfides and sulfates in serpentinites. Uncertainties represent the analytical (chemistry and analysis) reproducibility. Uncertainty of samples not duplicated corresponds to the standard IAEA S1 reproducibility in italic (see Table S3)