# Protracted hydrothermal alteration recorded at the microscale in the Chenaillet ophicarbonates (Western Alps): Insights from in situ δ18O thermometry in serpentine, carbonate and magnetite

Maria Rosa Scicchitano, Romain Lafay, John W. Valley, Noriko T. Kita, William O. Nachlas

► To cite this version:

Maria Rosa Scicchitano, Romain Lafay, John W. Valley, Noriko T. Kita, William O. Nachlas. Protracted hydrothermal alteration recorded at the microscale in the Chenaillet ophicarbonates (Western Alps): Insights from in situ δ18O thermometry in serpentine, carbonate and magnetite. Geochimica et Cosmochimica Acta, 2022, 318, pp.144-164. 10.1016/j.gca.2021.11.025. insu-03661249

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Distributed under a Creative Commons Attribution 4.0 International License
1 Protracted hydrothermal alteration recorded at the microscale in

2 the Chenaillet ophicarbonates (Western Alps): Insights from *in*

3 *situ* δ¹⁸O thermometry in serpentine, carbonate and magnetite

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5 Maria Rosa Scicchitano<sup>a, b, *</sup>, Romain Lafay<sup>c</sup>, John W. Valley<sup>b</sup>, Noriko T. Kita<sup>b</sup>,

6 William O. Nachlas<sup>b</sup>

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8 <sup>a</sup> Deutsches GeoForschungsZentrum GFZ, D14473 Potsdam, Germany (maria.rosa.scicchitano@gfz-

9 potsdam.de)

10 <sup>b</sup> Department of Geoscience, University of Wisconsin, Madison 53706, WI, USA

11 (valley@geology.wisc.edu, noriko@geology.wisc.edu, nachlas@wisc.edu)

12 <sup>c</sup> Géosciences Montpellier, Université Montpellier-II, 34095 Montpellier, France

13 (romain.lafay@yahoo.fr)

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15 * Corresponding author. E-mail: maria.rosa.scicchitano@gfz-potsdam.de

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17 Keywords: ophicarbonates, Alps, oxygen isotopes, geothermometry, SIMS

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27 Abstract

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29 The study of serpentinites and ophicarbonates from ophiolitic terrains provides a
30 three-dimensional perspective on the hydration and carbonation processes affecting
31 modern oceanic lithosphere. The Chenaillet ophiolite (western Alps) is interpreted as
32 a fragment of an oceanic core complex that resembles a modern slow spreading
33 center, and it was weakly affected by Alpine metamorphism. Ophicarbonates from the
34 Chenaillet ophiolite were targeted in this study for *in situ* analysis by Secondary Ion
35 Mass Spectrometry (SIMS) of oxygen and carbon isotopes in serpentine, calcite,
36 dolomite and magnetite. The high spatial resolution of SIMS allowed us to target
37 different serpentine, carbonate and magnetite generations intergrown at scales ≤ 50
38 µm, and reveal systematic zoning in δ¹⁸O with a range of 5.8‰ in serpentine (from
39 3.0 to 8.8‰, V-SMOW), 21.2‰ in carbonate (9.4 to 30.6‰), and 5.6‰ in magnetite
40 (−5.0 to −10.6‰). Coupled analysis of oxygen isotopes in seven different touching-
41 pairs of co-crystallized serpentine+carbonate and serpentine+magnetite provides
42 independent constraints on both the temperatures and δ¹⁸O(water) values during
43 serpentinization and carbonation responsible for the formation of the Chenaillet
44 ophicarbonates. The new stable isotope data and thermometric estimates can be
45 directly linked to textural and petrographic observations. These new results identify at
46 least four different stages of hydrothermal alteration in the Chenaillet ophicarbonates:
47 (1) peridotite hydration during seafloor exhumation at temperatures down to 200-130
48 °C and water δ¹⁸O values varying from 5 to 2‰, as documented by
49 serpentine+magnetite in mesh textures; (2) carbonation during exhumation near the
50 seafloor at temperatures as low as 10 °C assuming water δ¹⁸O values of −1‰, as
51 documented by the highest oxygen isotope ratios in texturally older calcite; (3)
serpentinization and carbonation at temperatures up to 240 °C and water δ¹⁸O values
of 2-3‰, as documented by serpentine+magnetite in veins crosscutting mesh textures
(T = 192±66 °C, δ¹⁸O(water) = 2±1‰, 2 standard deviation), serpentine+magnetite (T =
182±32 °C, δ¹⁸O(water) = 2±1‰) and serpentine+dolomite (T = 243±79 °C, δ¹⁸O(water) =
3±2‰) in recrystallized hourglass domains within serpentinite clasts,
serpentine+dolomite (T = 229±50 °C, δ¹⁸O(water) = 3±1‰) and serpentine+calcite (T =
208±40 °C, δ¹⁸O(water) = 2±1‰) within the fine-grained calcite matrix surrounding
serpentinite clasts; (4) late stage carbonation at temperatures down to 70-40 °C
assuming water δ¹⁸O values of 3 to -1‰, as documented by the highest oxygen
isotope ratios in a large calcite vein crosscutting both serpentinite clasts and fine-
grained carbonate matrix. We suggest that the textural and isotopic observations are
consistent with a protracted serpentinization and carbonation of the lithospheric
mantle that started during progressive exhumation to the seafloor and continued due
to interaction with hot and isotopically shifted seawater, which circulated at depth in
the oceanic crust and was then discharged near the seafloor, similar to modern mid-
ocean ridge venting systems.

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# 1. INTRODUCTION

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Serpentinization and carbonation of ultramafic rocks occur over a wide range of environments and physical-chemical conditions. The majority of modern serpentinites form at slow and ultraslow spreading ridges during mantle unroofing and exhumation at the seafloor at temperatures varying from 500 °C to <100 °C (e.g., Mével, 2003; Barnes et al., 2009; Boudier et al., 2010; Bach et al., 2011; Rouméjon et al., 2018).

Carbonation reactions resulting in the formation of veins and ophicarbonates (i.e., carbonated ultramafic breccias resulting from repeated tectonic and/or hydraulic fracturing of serpentinites) are usually associated with late stages of low-temperature alteration near the seafloor (Bonatti et al., 1974; Mével, 2003; Früh-Green et al., 2004; Manatschal and Müntener, 2009). Serpentinization of the subcontinental mantle is also documented at ocean-continent transition (OCT) boundaries (e.g., Iberia-Newfoundland rifted margins) where seawater infiltration is favored by crustal thinning and tectonic activity (e.g., Boillot et al., 1989; Skelton and Valley, 2000; Manatschal and Müntener, 2009; Schwarzenbach et al., 2013; Alt et al., 2018; Picazo et al., 2019). In addition to passive margins, serpentinization and carbonation of peridotite also occur at active margins and on continental settings. Ultramafic rocks exposed in ophiolitic terrains can undergo further hydration and carbonation due to low-temperature (<50 °C) interaction with meteoric water (e.g., Barnes and O'Neil, 1969; Barnes et al., 1978; Neal and Stanger, 1985; Streit et al., 2012; Barnes et al., 2013; Falk et al., 2016).

The study of ophiolitic terrains is fundamental to unravel the magmatic, tectonic and hydrothermal evolution of ocean basins from rift-drift and seafloor spreading stages to subduction initiation and final closure (Dilek and Furnes, 2014). In fact,
ophiolites represent an analogue for present-day oceanic crust and preserve the only evidence for ocean basins older than ~170 Ma (Dilek and Furnes, 2014; Pearce, 2014). Ophiolites provide better access than the one-dimensional drill cores from modern oceanic crust, but they often record intense alteration during their passage from rift to subduction/collision zone to mountain belt (Pierce, 2014; Staudigel et al., 2014). Constraining P-T conditions and fluid sources of hydration and carbonation reactions in ultramafic rocks from ophiolitic terrains is also important for better characterizing the conditions that favor dissolution and precipitation of carbonate minerals in these lithologies, with implications for better estimating the potential for permanent storage of CO₂ in ultramafic rocks (Kelemen and Matter, 2008; Kelemen et al., 2011). It is also crucial for tracing the mobilization and fluxes of water, carbon and fluid-mobile elements among Earth's different reservoirs (e.g., Scambelluri et al., 2004; Alt et al., 2012; Collins et al., 2015) and the production of fluids enriched in hydrogen, methane and other hydrocarbons both at near-surface conditions (Kelley et al., 2005; Miller et al., 2016) and in subduction zones (Vitale Brovarone et al., 2017).

Oxygen isotopes have been extensively used as geothermometers and tracers to determine the temperatures and fluid sources during peridotite hydration and carbonation in various tectonic settings (e.g., Wenner and Taylor, 1971, 1973; Agrinier and Cannat, 1997; Früh-Green et al., 1990, 1996; Barnes et al., 2009; Lafay et al., 2017; Noël et al., 2018; Rouméjon et al., 2018; Scicchitano et al., 2018a,b, 2020). This method requires knowledge of mineral-fluid or mineral-mineral oxygen isotope fractionation factors and assumption that isotope equilibrium between co-precipitated minerals and fluids was achieved and preserved. Measured δ¹⁸O values in co-precipitated minerals can, then, be combined with corresponding calibrations of oxygen isotope fractionations to independently calculate both the equilibrium
temperature and the oxygen isotope composition of the interacting fluid. Several studies (e.g., Wenner and Taylor, 1971; O'Hanley and Offler, 1992; Früh-Green et al., 1996; Alt and Shanks, 2006; Barnes et al., 2009) have estimated serpentinization temperatures and fluid $\delta^{18}$O values using the equilibrium oxygen isotope fractionation between coexisting serpentine and magnetite previously separated from the rock.

Geothermometry in carbonated serpentinites has recently been achieved using carbonate clumped isotope techniques that allow calculating equilibrium temperatures independently of the isotopic composition of the fluids from which carbonate formed (e.g., Streit et al., 2012; Falk and Kelemen, 2015; de Obeso and Kelemen, 2018).

Studies based on mineral separates and clumped isotope techniques represent a fundamental achievement in our understanding of the temperatures at which serpentinization and carbonation of ultramafic rocks occur. Nevertheless, they likely represent average estimates and overlook the larger complexity of these reactions recorded at the microscale by the occurrence of multiple serpentine and carbonate generations. In fact, recent *in situ* oxygen isotope studies by Secondary Ion Mass Spectrometry (SIMS) in serpentinites and carbonated serpentinites from different tectonic settings have shown a significant variation in the oxygen isotope composition of serpentine and carbonate minerals recorded at the microscale (Noël et al., 2018; Rouméjon et al., 2018; Scicchitano et al., 2018a, 2020). The high spatial resolution of SIMS techniques allows the analysis of oxygen isotopes to be measured along detailed traverses through co-precipitated host rock and vein minerals that are in textural equilibrium at scales $\le 50$ µm (Turnier et al., 2020). Consequently, it is a powerful tool for obtaining independent estimates of serpentinization temperatures and fluid $\delta^{18}$O values framed in a petrological context in rocks that record multiple stages of fluid-rock interactions (e.g., Scicchitano et al., 2018b, 2020) as well as for
152 testing the achievement of isotopic equilibrium during serpentinization processes
153 (e.g., Lafay et al., 2019).

154 In this study we apply SIMS techniques in combination with new reference
155 materials (Scicchitano et al., 2020) to analyze oxygen and carbon isotopes in
156 serpentine, carbonate (calcite, dolomite) and magnetite in order to unravel the
157 complex fluid-rock history recorded at the microscale in ophicarbonate samples from
158 the Chenaillet ophiolite. The fine-grained textures observed in the investigated
159 samples can only be targeted by SIMS, which requires the use of matrix-matched
160 reference materials (Valley and Kita, 2009). The new analytical capabilities provide
161 unique constraints on oceanic hydrothermal alteration recorded in the Chenaillet
162 ophiolite that were not resolved by earlier studies based on conventional data.
163 Furthermore, the ability to target mineral phases that are in textural equilibrium (e.g.,
164 serpentine-carbonate, serpentine-magnetite) allowed us to constrain independently
165 both alteration temperatures and fluid sources.

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## 2. GEOLOGICAL SETTING AND PREVIOUS STUDIES

168

169 The Chenaillet ophiolite in the Franco-Italian Alps (Fig. 1) is interpreted as a
170 remnant of embryonic oceanic crust formed at the distal part of a Magma-Poor-Ocean
171 Continent Transition (MP-OCT) margin (Manatschal and Müntener, 2009;
172 Manatschal et al., 2011; Li et al., 2013; Lafay et al., 2017) resembling actual slow
173 spreading centers (Lagabrielle and Cannat, 1990; Lagabrielle and Lemoine, 1997). It
174 consists mainly of pillow lavas with Mid Ocean Ridge Basalt (MORB) geochemical
175 signature superimposed on tectonically exhumed gabbros and serpentinized mantle
176 rocks comparable to oceanic core complexes (OCC) exposed in modern slow- to
ultraslow-spreading ridges (Manatschal and Müntener, 2009; Manatschal et al., 2011; Li et al., 2013; Lafay et al., 2017). Recent U-Pb zircon ages in the gabbros suggest that the Piemont-Liguria oceanic domain formed between 158-166 Ma (Li et al., 2013). No sheeted dike complex and post-rift sedimentary cover are observed in the Chenaillet ophiolite; sediments occur locally and consist of reworked mafic and ultramafic rocks covered by pelagic deposits, radiolarite and limestones (Lemoine et al., 1983; Manatschal et al., 2011). In contrast to other ophiolitic terrains in the Alps, the Chenaillet ophiolite escaped high-pressure metamorphism during the Late Cretaceous to Eocene Alpine subduction and was only weakly overprinted by subgreenschist to lower greenschist facies metamorphism during thrusting onto the European margin (Mével et al., 1978; Goffé et al., 2004; Manatschal and Müntener, 2009; Manatschal et al., 2011). High- to low-temperature hydrothermal processes in the Chenaillet ophiolite occurred in an oceanic setting prior to obduction and are represented by mylonitized amphibole-bearing gabbros (Mével et al., 1978), serpentinites and ophicarbonates (Manatschal and Müntener, 2009; Lafay et al., 2017). Serpentinization of mantle rocks is pervasive and original silicate minerals from the igneous stage are usually replaced by mesh and bastite textures typical of seafloor alteration under static conditions (Andreani et al., 2007; Manatschal and Müntener, 2009; Manatschal et al., 2011; Lafay et al., 2017). The pioneering work of Lemoine et al. (1983) and Lagabrielle and Lemoine (1997) has been recently reconsidered by Lafay et al. (2017) who provided a detailed petrological, geochemical and stable isotope (O, C, H) study of ophicarbonates from the Chenaillet ophiolite. These rocks consist of few tens of meter thick layers (Fig. 1b) of breccia found atop massive serpentinites near the serpentinite-basalt transition. Ophicarbonate breccias are made of slightly rounded serpentinite clasts cemented by a matrix of small (~12
µm in average) calcite grains (Lemoine et al., 1983; Lafay et al., 2017) (Fig. A.1). The diversity of microstructures (veins, replacement, matrix) allowed the authors to identify multiple events of carbonation following initial serpentinization of peridotites, which were attributed to a progressive cooling during exhumation of the mantle at the seafloor (Lafay et al., 2017). However, values of δ¹⁸O in serpentine (3.3 to 5.1‰) and carbonates (11 to 16‰) could suggest that both serpentinization and carbonation occurred at higher temperatures (T = 155-190 °C and 110-180 °C, respectively) than seafloor alteration (Lafay et al., 2017). These temperature estimates should be reexamined. They were calculated assuming a constant unaltered seawater composition (Lafay et al., 2017) and homogenous mineral grains. However, the microstructures and chemical zoning suggest that the micro-drilling technique used to select serpentine and carbonate for stable isotope analyses has mixed different generations resulting in average compositions. In light of recent analytical developments allowing accurate *in situ* analysis of oxygen and carbon isotopes in serpentine, magnetite and carbonate by SIMS (Valley and Kita, 2009; Huberty et al., 2010; Kita et al., 2011; Śliwiński et al., 2015a,b; Scicchitano et al., 2018a, 2020), a number of ophicarbonates studied by Lafay et al. (2017) were reanalyzed for stable isotope ratios (O, C) by SIMS to: (i) target the various serpentine and carbonate generations identified by Lafay et al. (2017) that could not be separated with micro-drilling techniques due to their fine-grained size, (ii) identify systematic isotopic zoning recorded at the microscale in both serpentine and carbonate, (iii) couple *in situ* analysis of oxygen isotopes in potentially co-crystallized minerals (serpentine-calcite, serpentine-dolomite, calcite-dolomite, serpentine-magnetite) to estimate the equilibrium temperature independently of the fluid composition. To the best of our knowledge, this study represents the first successful attempt to applying the
227 serpentine-magnetite geothermometer at the microscale and provides the first stable
228 isotope data in dolomite in these samples that could not be targeted by previous
229 studies due to its small grain size.

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231 **3. PREPARATION OF GRAIN MOUNTS AND ANALYTICAL METHODS**

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233 Five ophicarbonate samples (RL3, RL6, RL7, RL9 and RL10) from the Chenaillet
234 ophiolite previously characterized by Lafay et al. (2017) were prepared for oxygen
235 and carbon isotope analysis by SIMS, as described in the Electronic Appendix A.

236 A Hitachi S3400 Variable Pressure (VP) Scanning Electron Microscope (SEM) at
237 the Department of Geoscience (University of Wisconsin-Madison, UW-Madison) was
238 operated at low (P = 30 Pa) and high vacuum in backscattered electron (BSE) mode to
239 identify the different generations of serpentine, carbonate and magnetite in the
240 Chenaillet ophicarbonates described by Lafay et al. (2017). Carbonate minerals were
241 also imaged using a cathodoluminescence (CL) Gatan PanaCL/F system. Careful
242 SEM imaging in BSE, secondary electron (SE) and CL modes was performed
243 following each SIMS session to check the location and features of all SIMS pits. We
244 rejected analyses where pits were located on fractures, mixed domains or other
245 features that could alter isotope ratios. Prior to SEM analysis, mounts were coated
246 with carbon or a thin layer of gold; to avoid affecting carbon isotope ratios during
247 SIMS analysis in carbonates, no carbon coating was used on the corresponding
248 mounts. The accelerating voltage and working distance were usually ~15 or 10 kV
249 and ~10 or 15 mm for BSE and CL imaging, respectively.

250 Serpentine, calcite, dolomite and magnetite from the Chenaillet ophicarbonates
251 were analyzed for major and minor element mineral chemistry using a CAMECA
252 SXFive Field Emission Electron Microprobe (EPMA) at the Department of
253 Geoscience (UW-Madison). Details of EPMA methods are given in the Electronic
254 Appendix A. Direct analysis of oxygen with EPMA was necessary for accurate
255 quantification of variable Fe oxidation in serpentine and magnetite and variable water
256 contents in serpentine. All serpentine analyses were normalized to 18 (O + OH)
257 according to the general formula Mg₆Si₄O₁₀(OH)₈. Approximate H₂O content (wt.%)
258 was calculated based on excess oxygen values measured by EPMA assuming total
259 iron as Fe²⁺ (Table A.1). Carbonate analyses were normalized to 6 O according to the
260 general formula (Ca,Mg)₂(CO₃)₂ (Table A.2). Magnetite analyses were normalized to
261 3 cations according to the general formula Fe₃O₄, and approximate Fe³⁺ and Fe²⁺
262 contents were estimated by charge balance (Table A.3).

263 A CAMECA IMS-1280 large radius multicollector secondary ion mass
264 spectrometer at the WiscSIMS laboratory (UW-Madison) was used to perform *in situ*
265 oxygen and carbon isotope analyses of serpentine, carbonate and magnetite in the
266 Chenaillet ophicarbonates during seven analytical sessions (Tables A.1 to A.4).
267 Details on analytical settings are given in the Electronic Appendix A. We monitored
268 bias, drift and external precision during each analytical session by bracketing 10-20
269 unknown analyses with at least eight analyses of the reference material located in the
270 same grain mount. Matrix effects on bias due to variations in Mg/Fe ratios in
271 serpentine and dolomite-ankerite solid solutions, which are observed during oxygen
272 and carbon isotope analysis by SIMS (Śliwiński et al., 2015a,b; Scicchitano et al.,
273 2020), were monitored and quantified for each analytical session by constructing a
274 calibration curve using a suite of serpentine and carbonate reference materials (Table
275 A.4). These calibration curves were then used to correct δ¹⁸O and δ¹³C values
276 measured in serpentine and dolomite from the Chenaillet ophicarbonates. Orientation
effects in chrysotile were estimated by analyzing two different orientations of the chrysotile C22908 reference material as described in Scicchitano et al. (2018a, 2020), while orientation effects in magnetite were minimized using a special 3 kV protocol (see Electronic Appendix A for details) and evaluated by analyzing ~30 grains with different orientation of the magnetite M7-5830 reference material (Huberty et al., 2010; Kita et al., 2011).

# 4. RESULTS

## 4.1. Petrography and mineral chemistry

For the purpose of this study, we summarize the main textures targeted for *in situ* analysis by SIMS, which are essential for putting into petrological context the new oxygen and carbon isotope data.

### 4.1.1. Serpentine

Serpentine is the main constituent of the serpentinite clasts where it forms either mesh textures characterized by well-defined cores and rims (Fig. 2a-c) or hourglass textures in which a distinction between mesh rim and mesh center is not possible (Fig. 2d). Raman spectroscopy showed that mesh cores consist of a mixture between lizardite and chrysotile (hereafter referred to as “serpentine”), whereas mesh rims consist of pure lizardite (Lafay et al., 2017). Magnetite (up to ~20 µm large) is often observed around mesh cores in sample RL6 (Fig. 2b), while it is usually lacking or very limited in mesh textures of other samples. It is occasionally possible to identify an inner and outer mesh
core both surrounded by magnetite (Fig. 2b). A complex network of mesh veinlets (V-1) consisting of lizardite+magnetite follows original (pristine silicate) grain boundaries (Fig. 2e-f), while larger lizardite+magnetite veins 250 µm in width (V-2) crosscut mesh and hourglass textures (Fig. 2g). In sample RL6, lizardite+dolomite layers are observed within the carbonate matrix (Fig. 2j-k); this serpentine generation is hereafter referred to as “matrix serpentine”. Finally, small chrysotile aggregates up to a few tens of µm large (Lafay et al., 2017) are widespread in the carbonate matrix surrounding serpentinite clasts (Fig. 2k-l).

Mesh cores have higher Mg numbers [Mg# = 100 × Mg/(Mg+Fe), molar, with Fe = Fe(total)] and silica contents relative to mesh rims in all samples (Fig. A.2). In particular, mesh cores in samples RL6, RL7 and RL9 have Mg# ranging between 96 and 99 mole % and SiO₂ contents ranging between 40.6 and 43.4 wt.%, whereas mesh rims have Mg# of 90 to 97 mole % and SiO₂ contents of 38.0 to 42.0 wt.% (Fig. A.2, Table A.1). A similar trend between Mg# and silica is observed in mesh textures from samples RL10 and RL3, which however show lower Mg# (91 to 95 mole % in mesh cores, 85 to 95 mole % in mesh rims) and comparable SiO₂ contents (39.9 to 42.5 wt.% in mesh cores, 36.0 to 42.8 wt.% in mesh rims) relative to mesh serpentine in the other three samples (Fig. A.2, Table A.1). The variability in alumina content is lower in mesh serpentine from samples RL6, RL7 and RL9 (Al₂O₃ = 0.2-1.3 wt.%) relative to mesh serpentine from samples RL10 and RL3 (Al₂O₃ = 0.1-4.3 wt.%) (Fig. A.2, Table A.1). Serpentine in hourglass textures has a similar chemical composition as serpentine in mesh rims (Fig. A.2, Table A.1). Lizardite in V-1 veins has similar silica (SiO₂ = 41.2-43.1 wt.%) and alumina contents (Al₂O₃ = 0.1-0.6 wt.%) across different samples, whereas Mg# decreases from sample RL9 (Mg# = 96-97) to samples RL10 and RL3 (Mg# = 94-95) (Fig. A.2, Table A.1). Lizardite in V-2 veins and “matrix serpentine”
observed in sample RL6 have similar chemical compositions as those of mesh cores in the same sample (Fig. A.2, Table A.1). Chrysotile in the carbonate matrix surrounding serpentinite clasts is also characterized by higher Mg# in samples RL6, RL7 and RL9 (Mg# = 94-98) relative to samples RL10 and RL3 (Mg# = 88-95; Fig. A.2, Table A.1); in contrast, both silica and alumina contents are similar across different samples ($SiO_2$ = 39.5-44.2 wt.%, $Al_2O_3$ = 0.2-2.2 wt.%).

### 4.1.2. Carbonates within serpentinite clasts

Serpentine in mesh cores is often partially or completely replaced by calcite (hereafter referred to as C2, Lafay et al., 2017) and relict serpentine inclusions are frequently observed in calcite mesh (Fig. 2a-c). C2 calcite is sometimes connected to calcite veinlets ~10 µm in width (hereafter C1; Lafay et al., 2017); the largest C1 calcite veins (up to 500 µm in width) crosscut serpentine in mesh and hourglass textures as well as V-1 and V-2 lizardite+magnetite veins (Fig. 2h-i). Although C1 calcite veins are usually restricted to individual clasts and are often truncated at the edges of the clasts, they sometimes propagate into the surrounding carbonate matrix (Fig. A.1). SEM-CL imaging shows complex patchy zoning and resorption features in both C1 veins and C2 mesh suggesting multiple events of recrystallization (Fig. 3). As previously shown by Lafay et al. (2017), the edges of C1 calcite veins are sometimes delimited by distinct CL calcite rims (Fig. 3b-c).

In addition to calcite, subordinate dolomite is observed in some samples as (i) rims around C2 calcite mesh (Fig. 3e, 4a), (ii) borders along C1 veins at the boundary with serpentine (Fig. 3a, 4b), (iii) subhedral grains associated with "matrix serpentine" in sample RL6 (Fig. 2k, 4c) and (iv) veinlets in the C3 calcite matrix.
No significant variation is observed in the chemical composition of calcite in different textures across different samples (Table A.2). FeO and MnO contents are usually <0.5 wt.% with the exception of calcite in C2 mesh (FeO = 0.1-1.3 wt.%) from sample RL6. In contrast, the MgO content in calcite from different textures has a larger variability from 0.02 to 1.6 wt.% (Table A.2).

Dolomite has a similar chemical composition in all samples and textural locations with FeO content varying from 0.1 to 1.2 wt.% and MnO content <0.3 wt.%; Fe# is <3 mole %.

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### 4.1.3. Carbonates surrounding serpentinite clasts

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Serpentinite clasts are embedded in a calcite matrix (hereafter C3; Lafay et al., 2017) that appears either “turbid”, when intergrown with fine-grained chrysotile aggregates as documented through Raman spectroscopy (Lafay et al., 2017), or “clear”, when no chrysotile is present (Fig. A.3). Clear calcite domains (hereafter C1') in the C3 calcite matrix were interpreted by Lafay et al. (2017) as remnants of C1 veins channeling fluids. Calcite veins ~100 µm in width (hereafter C3-veins) sometimes crosscut the C3 calcite matrix but not the lizardite+dolomite layers (Fig. 2j). Finally, a large calcite vein ~5 mm in width (hereafter C4; Lafay et al., 2017) crosscuts both C3 calcite matrix and serpentinite clasts in sample RL9 (Fig. A.1). SEM-CL imaging shows that C1' clear calcite domains consist of large calcite grains that sometimes exhibit a concentric (Fig. 5a-b) or oscillatory CL zoning (Fig. 5c) truncated by the surrounding aggregates of fine-grained C3 calcite matrix; the latter also shows a similar concentric CL zoning (Fig. 5a, d). Calcite grains in C1' are often crosscut or bounded by later calcite veinlets (hereafter C1'-veins) (Fig. 5a, c). In contrast, C3-veins do not exhibit complex CL zoning.
zoning (Fig. 5d). In addition to calcite, dolomite was observed at the boundary between
C4 vein and serpentinite clast in sample RL9 (Fig. 4d).

FeO and MnO contents are usually <0.5 wt.% with the exception of calcite in C3-
veins (MnO = 0.1-0.7 wt.%) from sample RL6. In contrast, the MgO content in calcite
from different textures has a larger variability from 0.03 to 1.9 wt.% (Table A.2). No
significant variation is observed between different textures across different samples.

383

384 **4.2. Oxygen and carbon isotopes by SIMS**

385

386 **4.2.1. Serpentine**

387

In situ oxygen isotope analyses of serpentine were performed in two analytical
sessions (S1 and S2; Table A.1). A suite of six (S1) to seven (S2) antigorite reference
materials was analyzed at the beginning of each analytical session in order to build a
calibration curve for the correction of polymorph-specific matrix effects due to Mg/Fe
variations in the compositional range Mg# = 77.5-99.5 (Table S2 in Scicchitano et al.,
2020). Lizardite (L3431) and chrysotile (C22908) reference materials (Scicchitano et
al., 2018a, 2020) were also analyzed to correct for the bias relative to the normalizing
antigorite reference material Al06-44A (Table S2 in Scicchitano et al., 2020). A bias
of ~1‰ was measured between lizardite L3431 and antigorite Al06-44A in both
analytical sessions. Orientation effects in chrysotile C22908 were ~1‰ in all
analytical sessions, hence this is considered as a limitation on the accuracy of
analyses done in this polymorph, but no orientation effects are detected for lizardite or
antigorite. Measured δ¹⁸O values in serpentine from the Chenaillet ophicarbonates

389

390

391

392

393

394

395

396

397

398

399

400
401 were corrected according to the method described in Scicchitano et al. (2020) and are
402 reported relative to the V-SMOW scale (Table A.1).

403 The highest average δ¹⁸O values are recorded in the core of mesh textures where
404 oxygen isotope ratios vary between 4.4 and 8.5‰ with no distinction between calcite-
405 free and calcite-bearing mesh cores (Fig. 6). Although mesh rims show a large
406 variability in oxygen isotope ratios with δ¹⁸O values varying from 4.5 to 8.8‰, their
407 average δ¹⁸O values are generally lower compared to mesh cores (Fig. 6). No
408 correlation between δ¹⁸O value and Mg# was observed in mesh textures. Lizardite in
409 V-1 veins has an oxygen isotope composition varying from 3.5 to 6.4‰ with average
410 δ¹⁸O values lower relative to mesh rims (Fig. 6). Lizardite in V-2 veins crosscutting
411 both V-1 veins and mesh textures has low δ¹⁸O values of 3.2-4.2‰ (Fig. 6). Similar
412 low oxygen isotope ratios relative to mesh cores are recorded in hourglass serpentine
413 (δ¹⁸O = 3.8-5.7‰), “matrix serpentine” (δ¹⁸O = 3.1-5.4‰) and serpentine in the C4
414 calcite vein (δ¹⁸O = 4.0-4.8‰) (Fig. 6). Chrysotile in the C3 calcite matrix has a large
415 variability in δ¹⁸O values from 3.0 to 5.7‰ (Fig. 6).

416 Overall, SIMS analyses allow identifying a larger variation in serpentine δ¹⁸O
417 values (from 3.0 to 8.8‰) relative to previous studies on powdered micro-drilled
418 samples (δ¹⁸O = 3.3-5.1‰; Lafay et al., 2017) as well as a systematic variation in
419 δ¹⁸O values between different serpentine generations.

420

421 **4.2.2. Carbonate**

422

423 *In situ* oxygen and carbon isotope analyses in carbonates were performed in four
424 analytical sessions (S3 to S5 for O, S6 for C; Table A.2). Calcite UWC-3 (Kozdon et
425 al., 2009) was used as a working reference material. A suite of seven to ten carbonate
reference materials was analyzed at the start of each session to calibrate matrix effects in samples with chemical composition along the dolomite-ankerite solid solution series (Table A.4). Measured δ¹⁸O and δ¹³C values in dolomite from the Chenaillet ophicarbonates were corrected according to the method described in Śliwiński et al. (2015a,b) and for the bias between dolomite UW6220 and calcite UWC-3 reference materials (Table A.2). Values of δ¹⁸O and δ¹³C are reported relative to the V-SMOW and V-PDB scales, respectively (Table A.2).

Values of δ¹⁸O for calcite in C1 veins mostly range between 9.4 and 13.9‰ and occasionally up to 22.3‰ (Fig. 7). The highest δ¹⁸O values are commonly observed at the edges of large C1 veins (Fig. 3b). Analyses performed across one of the largest C1 veins (Fig. 3c, A.4a) showed a nearly symmetric edge-to-edge oxygen isotope zoning up to ~3‰ with the lowest δ¹⁸O values in the core of the vein (Fig. A.4). Values of δ¹³C for C1 calcite veins vary between -0.6 and 2.4‰ but no systematic correlation is observed with textures (Fig. A.5). Calcite in C2 mesh textures shows the largest variability in oxygen isotope composition with δ¹⁸O values ranging from 9.8 to 30.6‰ (Fig. 7), whereas δ¹³C values vary between -2.6 and 1.9‰ (Fig. A.5). The highest δ¹⁸O values in C2 calcite mesh are typically observed at the core of mesh textures, whereas the edges usually have lower δ¹⁸O values (Fig. 3d-i). It is sometimes possible to observe resorption textures where calcite with low δ¹⁸O values, similar to those measured in nearby C1 veinlets, replaces C2 calcite mesh cores with higher δ¹⁸O values (Fig. 3f-i). Calcite in C1' textures also shows a large variability in oxygen and carbon isotopes similar to C2 calcite mesh with δ¹⁸O and δ¹³C values of 10.6 to 24.9‰ and -1.1 to 1.6‰, respectively (Fig. 7, A.5). Similar to calcite in C2 mesh textures, the highest δ¹⁸O values in C1' textures are commonly observed in the cores of large calcite grains (Fig. 5 and Electronic Appendix B). Late C1'-veins
crosscutting or bounding C1' calcite grains have low δ¹⁸O values of 10.9 to 11.8‰ (Fig. 5c, 7 and Electronic Appendix B) and δ¹³C values of -0.6 to 0.7‰ (Fig. A.5).

Similarly, calcite in the fine-grained C3 matrix and crosscutting C3-veins also have low δ¹⁸O values of 9.8-12.9‰ and 10.5-11.5‰, respectively (Fig. 5d, 7 and B.1 in Electronic Appendix B); δ¹³C values vary from -0.1 to 2.2‰ and -0.6 to 0.6‰, respectively (Fig. A.5). Finally, calcite in the latest C4 vein has the highest average oxygen isotope composition (average δ¹⁸O = 21.1‰, Fig. 7) and carbon isotope values similar to calcite in other textural locations (δ¹³C = -0.7 to 1.7‰, Fig. A.5).

Dolomite associated with mesh textures in serpentinite clasts has δ¹⁸O values of 10.7 to 14.1‰ (n = 20) that are comparable to the oxygen isotope composition of dolomite bordering C1 calcite veins (δ¹⁸O = 10.1-11.9‰, n = 9) and the C4 vein (δ¹⁸O = 11.2-12.5‰, n = 2) (Table A.2). Similar δ¹⁸O values are also recorded in subhedral dolomite associated with matrix serpentine in the layers at the boundary between serpentinite clasts and C3 calcite matrix (δ¹⁸O = 11.4-13.2‰, n = 5) (Table A.2).

SIMS analyses show a much larger variability (~21‰) in calcite δ¹⁸O values relative to ~6‰ variation described in Lafay et al. (2017). Additionally, the ability to place SIMS data into a petrological context highlights that higher δ¹⁸O values in zoned C1 veins, C2 mesh and C1' calcite domains likely represent an older signature that is progressively replaced by younger calcite generations with lower δ¹⁸O values (core of C1 veins, C1'-veins, C3 matrix and C3-veins). SEM-CL imaging linked to oxygen isotope ratios also allowed distinguishing C1' domains from C1 open veins.

Values of δ¹⁸O increase again in the latest C4 calcite vein (average δ¹⁸O = 21.1‰) crosscutting serpentinite clasts and C3 calcite matrix (average δ¹⁸O = 11.1‰). Values of δ¹³C measured by SIMS vary by ~4.5‰ at most in C2 calcite mesh, similar to the variation observed by Lafay et al. (2017). Nevertheless, SIMS analyses are generally
476 higher relative to those reported in Lafay et al. (2017) (–2.6 to +2.4‰ vs. –5.2 to
477 +0.4‰, respectively).

478

479 **4.2.3. Magnetite**

480

481 *In situ* oxygen isotope analyses in magnetite were performed in one analytical
482 session (S7; Table A.3). The extent of crystal orientation effects in magnetite was
483 determined by bracketing analyses of the M7-5830 magnetite reference material with
484 eight analyses of the UWQ-1 quartz reference material (Kelly et al., 2007) (Table
485 A.4). A grain-to-grain reproducibility of ~1‰ (2 standard deviations, 2SD) was
486 obtained (using the low-kV protocol of Kita et al., 2011) in both M7-5830 and UWQ-
487 1 (Table A.4); hence this represents the maximum precision achievable during oxygen
488 isotope analyses of magnetite in this study. Magnetite analyses in the ophicarbonate
489 samples from the Chenaillet ophiolite were bracketed with eight analyses in the
490 UWC-3 calcite reference material (Kozdon et al., 2009) located in the same grain
491 mounts in order to monitor bias, drift and external precision (Table A.3). Magnetite
492 analyses were corrected for the bias between UWC-3 calcite and M7-5830 magnetite
493 reference materials measured in different mounts, and they are reported relative to the
494 V-SMOW scale (Tables A.3 and A.4).

495 Magnetite grains larger than ~10 µm located in different textures (mesh, V-1 and
496 V-2 veins, hourglass serpentine) were targeted for SIMS oxygen isotopes analysis in
497 samples RL6, RL3 and RL9 (Table A.3). Magnetite in mesh textures (Fig. 2b and
498 A.6a) shows the largest variability in δ¹⁸O values from -10.6 to -5.0‰ (n = 32).
499 Magnetite in V-1 veins (Fig. 2f and A.6b) and V-2 veins (Fig. 2g and A.6c) has
500 oxygen isotope ratios varying from -8.5 to -7.3‰ (n = 3) and -7.7 to -10.6‰ (n = 2),
respectively. Magnetite associated with hourglass serpentine (Fig. 2d and A.6d) has an oxygen isotope composition of –9.6 to –8.0‰ (n = 4).

503

504

# 5. DISCUSSION

505

## 5.1. Estimating alteration temperatures and water $\delta^{18}$O values in co-crystallized mineral pairs

508

Based on textural observations, we identified seven pairs of minerals that appear to be in textural equilibrium along the metasomatic sequence and, therefore, assumed to have co-crystallized: (i) lizardite+magnetite in mesh textures with a single mesh core, V-1 veins (Fig. 2e, f), V-2 veins (Fig. 2g), and in the hourglass matrix (Fig. 2d); (ii) lizardite+dolomite in the layers at the boundary between serpentinite clast and carbonate matrix (Fig. 2k and 4c) as well as with hourglass serpentine (Fig. 4b, d); (iii) chrysotile+calcite in the C3 calcite matrix (Fig. 2l). Coupled oxygen isotope analysis by SIMS in these mineral pairs allows us to test whether isotopic equilibrium was achieved during hydration and carbonation reactions, in which case equilibrium temperatures and water $\delta^{18}$O values are independently estimated.

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517

518

519

### 5.1.1. Serpentine-magnetite

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The majority of targeted serpentine-magnetite pairs (n = 17, Table 1) are in mesh textures (sample RL6) because of the larger (>10 µm) size of magnetite grains. Only a limited number of analyses was possible in V-1 veins (n = 1), V-2 veins (n = 2) and hourglass matrix (n = 4) because of the small grain size of magnetite (Table 1).
A large uncertainty exists on the serpentine-water oxygen isotope fractionation because different calibrations (Wenner and Taylor, 1971; Savin and Lee, 1988; Früh-Green et al., 1996; Zheng, 1993; Saccocia et al., 2009) disagree by >4‰ and up to ~20‰ when extrapolated to temperatures <200 °C (see Fig. 10 and discussion in Scicchitano et al., 2020). In this study, the preferred calibrations for calculating the equilibrium temperature for each serpentine-magnetite pair are the serpentine-water oxygen isotope fractionation factors of Früh-Green et al. (1996) and the magnetite-water oxygen isotope fractionation factors of Bottinga and Javoy (1973). The same calibrations were then used to calculate the oxygen isotope composition of water in equilibrium with both serpentine and magnetite at each equilibrium temperature (Table 1 and Fig. 8). For mesh cores, equilibrium temperatures vary from 204 °C to 132 °C and δ¹⁸O(water) values decrease with temperature from 5.2 to 1.8‰ (Fig. 8). For serpentine-magnetite pairs in V-1 veins, V-2 veins and hourglass textures, equilibrium temperatures vary from 173 °C, 169-215 °C and 162-198 °C, respectively, while water δ¹⁸O values vary between ~1 and 3‰ (Fig. 8).

For comparison, equilibrium temperatures and δ¹⁸O(water) values were also calculated using other calibrations. Estimates calculated by combining the serpentine-water oxygen isotope fractionation factors of Saccocia et al. (2009) with the magnetite-water oxygen isotope fractionation factors of Bottinga and Javoy (1973) differ by ~50 °C (from 261 to 191 °C) and ~1‰ (from ~4 to 0‰), respectively, relative to estimates derived from the calibration by Früh-Green et al. (1996) (Fig. 8 and Table 1). Temperatures calculated with the calibration of Saccocia et al. (2009) could be overestimated of up to 40 °C (at exchange fraction < 40%) due to a bias in their serpentine-water calibration resulting from extrapolation of equilibrium fractionation factors in three-exchange site systems using the methods described by
Northrop and Clayton (1966) or Zheng et al. (1994), which are valid for two-exchange site systems (Agrinier and Javoy, 2016). Equilibrium temperatures calculated by combining the serpentine-water and magnetite-water oxygen isotope fractionation factors of Zheng (1991, 1993) as well as the serpentine-magnetite calibration by Früh-Green et al. (1996; their Eq. 3) are ~50-70 °C lower than those estimated using the serpentine-water and magnetite-water calibrations by Früh-Green et al. (1996) and Bottinga and Javoy (1973) (Table A.5). Considering the maximum uncertainties derived from the various calibration equations and δ¹⁸O measurements, we estimate a total uncertainty of ±80 °C on the calculated equilibrium temperatures. The uncertainty on δ¹⁸O(water) values is estimated to be in the order of ±1‰.

561

### 5.1.2. Serpentine-dolomite

563

Lizardite and subhedral dolomite grains in the layers at the boundary between serpentinite clast and carbonate matrix are relatively homogenous in oxygen isotopes with average δ¹⁸O values of 4.7±1.4‰ (2SD, n = 10) and 12.2±1.3‰ (2SD, n = 5), respectively, suggesting that they are in isotopic equilibrium. Using the average δ¹⁸O values measured by SIMS, we calculated an oxygen isotope fractionation (Δ¹⁸O<sub>Dol-Srp</sub>) of 7.5‰, which corresponds to an equilibrium temperature of 229 °C (Fig 9a). In order to calculate the temperature at which serpentine and dolomite formed, we used the empirical calibration of the serpentine-water oxygen isotope fractionation by Früh-Green et al. (1996) and the calibration of the dolomite-water oxygen isotope fractionation (experiments at T = 252-292 °C, Matthews and Katz, 1977).

Extrapolation of the experimental calibration by Saccocia et al. (2009) would result in an unrealistically high equilibrium temperature (even if accounting for the bias

575
described by Agrinier and Javoy, 2016) given a maximum calculated $\Delta^{18}O_{Dol-Srp}$ of 5.6‰ is obtained assuming $T = 500$ °C, which is the maximum thermal stability of serpentine in oceanic settings (e.g., Mével, 2003). Consequently, we consider extrapolation of the empirical calibration by Früh-Green et al. (1996) more reliable. At a temperature of 229 °C, serpentine and dolomite would be in equilibrium with a fluid of 3.3‰ (Fig. 9b). Similar equilibrium temperature of 240 °C and water $\delta^{18}O$ value of 3.3‰ are calculated using the average oxygen isotope ratios measured in hourglass serpentine and dolomite pairs (Table 2, Fig. 9). These estimates are within uncertainty of those calculated for lizardite+magnetite pairs in hourglass textures.

### 5.1.3. Serpentine-calcite

Calcite in the fine-grained C3 matrix and the co-crystallized fine-grained chrysotile have average $\delta^{18}O$ values of 11.1±1.2‰ (2SD, $n = 57$) and 3.9±1.4‰ (2SD, $n = 22$), respectively, corresponding to a $\Delta^{18}O_{Cal-Srp}$ of 7.2‰ (Fig. 9a). Combining the empirical calibration of the serpentine-water system by Früh-Green et al. (1996) with the experimental calibration of the calcite-water oxygen isotope fractionation (experiments at $T = 0-500°C$, O'Neil et al., 1969), yields an equilibrium temperature of 208 °C and a $\delta^{18}O_{(water)}$ of 2.0‰ (Fig. 9). The calibration by O'Neil et al. (1969) is in good agreement with experimental calibrations by Golyshev et al. (1981) and Kim and O'Neil (1997). As for the dolomite-serpentine pair, the experimental calibration by Saccocia et al. (2009) for the calcite-serpentine pair would result in unrealistically high equilibrium temperature not consistent with seafloor serpentinization ($T > 500$ °C).
601 **5.2. Evolution model for the hydrothermal alteration in the Chenaillet ophiolite**

602
603 In the following discussion, stable isotope data are combined with petrographic
604 observations and well-documented geological context to reconstruct the main phases
605 of fluid-rock interaction recorded in the Chenaillet ophicarbonates (Fig. 10, 11).
606 Alteration temperatures and water oxygen isotope compositions estimated
607 independently from mineral-mineral pairs (see above) are used as anchor points for
608 the interpretations.

609

610 **5.2.1. Exhumation to the seafloor: early serpentinization followed by low-T**
611 **carbonation**

612

613 Oxygen isotope values of serpentine increase from V-1 mesh veins (δ¹⁸O = 3.5 to
614 6.4‰) to mesh cores (δ¹⁸O = 4.4 to 8.5‰), with no distinction between carbonate-free
615 and carbonate-bearing mesh textures (Fig. 6, 10). Additionally, serpentine in mesh
616 textures associated with magnetite has similar oxygen isotope and chemical
617 compositions as serpentine in magnetite-free mesh textures from the same sample.
618 Hence, we assume that the formation of serpentine in different mesh cores (i.e.,
619 magnetite- and calcite-bearing versus magnetite- and calcite-free) is attributable to the
620 same hydrothermal event. The physical-chemical conditions at which serpentine in V-
621 1 mesh veins and mesh rims formed cannot be independently constrained, but it is
622 likely that they record the earliest stages of alteration with V-1 veins acting as
623 conducts that allowed progressive fluid percolation and pervasive replacement in
624 mesh cores. The variation in Mg# and Al₂O₃ observed in mesh textures from samples
625 RL6, RL7 and RL9 versus samples RL10 and RL3 (Fig. A.2) suggests formation of
serpentine after olivine versus pyroxene, respectively (e.g., Deschamps et al., 2013).
Because no relict pyroxene and olivine are observed in any of the investigated samples, it is possible that serpentinization started at T>400 °C (favoring rapid dissolution of orthopyroxene) and then proceeded at T<300 °C (favoring rapid dissolution of olivine), as documented by experimental studies (e.g., Martin and Fyfe, 1970; Allen and Seyfried, 2003). The limited occurrence of antigorite, chlorite and tremolite in the Chenaillet ophicarbonates (Lafay et al., 2017) is consistent, in fact, with greenschist facies conditions (T = 300-500 °C). The abundant precipitation of massive magnetites in mesh cores would explain the formation of serpentine with higher Mg# in response to changes in the conditions of hydrothermal alteration. It should be noted that the marked crystallization of magnetite in mesh cores could be a consequence of the combined variation in fluid temperature and δ¹⁸O composition, water/rock ratio (W/R) and redox conditions (Klein et al., 2009, 2013).

In light of the above assumptions, the physical-chemical conditions at which the serpentine-magnetite pairs in mesh textures formed (i.e., T = 204 to 132 °C and δ¹⁸O(water) = 5 to 2‰) can be considered valid for serpentine in all mesh cores. It is worth noting that the concomitance between cooling and decrease in water δ¹⁸O values results in a net variation of magnetite and serpentine δ¹⁸O values by ~5.5‰ and ~0.4‰, respectively (Fig. A.7); this is consistent with the variability in δ¹⁸O values measured by SIMS in serpentine-magnetite pairs (Table 1). The decrease in fluid δ¹⁸O values during cooling is consistent with a progressive evolution from a rock-dominated system during early stages of serpentinization (or low W/R), with water δ¹⁸O values similar to the oxygen isotopic composition of the upper mantle (δ¹⁸O_mantle = 5.5‰, δ¹⁸O_olivine = 5.18±0.28‰; e.g., Mattey et al., 1994; Chazot et al., 1997), to a fluid-dominated system during later stages of serpentinization (or high
W/R) and exhumation of the mantle near the seafloor. Calculated water $\delta^{18}$O values of ~2‰ agree well with $\delta^{18}$O values of 1.2-1.5‰ (Logatchev hydrothermal field) and 2.2-2.4‰ (Nibelungen hydrothermal field) measured in hot ($T = 192-365^{\circ}$C)) ultramafic-hosted vent fluids sampled along the Mid-Atlantic slow-spreading ridge (e.g., Charlou et al., 2010 and references therein; Schmidt et al., 2011), which are higher in $\delta^{18}$O than present-day pure seawater (0‰) due to interaction with the oceanic crust (e.g., Mével, 2003). Similar temperatures ($T = 265-364^{\circ}$C) and $\delta^{18}$O values (1.1-2.4‰) are also recorded in mafic-hosted vent fluids (e.g., Jean-Baptiste et al., 1997).

Our new *in situ* data are overall in agreement with the original suggestion of Lafay et al. (2017) who, in continuity with the tectonic models proposed by Manatschal et al. (2011), proposed that exhumation of oceanic units occurred along detachment faults favoring hydration and carbonation of mantle rocks (Fig. 12a). The authors proposed that serpentinization initiated in a shallow seafloor environment at conditions of at least 150 °C, with reactive fluids assumed to have $\delta^{18}$O values of 0-2‰. The new *in situ* data suggest comparable temperature estimates within uncertainty, but independently constrain the composition and origin of the interacting fluids as well as identify that early serpentinization was followed by early low-temperature carbonation due to exposure of the mantle rocks on the paleo seafloor and interaction with low-temperature seawater (Fig. 12a). This is documented by the zonation in $\delta^{18}$O values observed in some carbonate generations (Fig. 7, 10), with texturally older calcite domains having high $\delta^{18}$O values in zoned C2 calcite mesh (18.0-30.6‰), at the edges of zoned C1 veins (18.7-22.3‰) and in the cores of zoned C1' domains (18.2-24.9‰), which are comparable to those reported for marine carbonates and metacarbonates from the Lago Nero unit (Lemoine et al., 1983). Such
carbonates likely formed at temperatures as low as 10 °C assuming interaction with
pure seawater in an ice-free world (δ¹⁸O = -1‰). Such low carbonation temperatures
relative to serpentinization are also consistent with the observation that calcite in
mesh textures is clearly replacing (hence postdating) serpentine.

680

## 5.2.2. An active hydrothermal cell: coupled high-T carbonation and serpentinization

683

An overall decrease in oxygen isotope values with textural progression is recorded
by both serpentine and carbonate in all investigated samples (Fig. 6, 7, 10). In
particular, δ¹⁸O values of serpentine in V-2 veins (δ¹⁸O = 3.2-4.2‰) and hourglass
textures (δ¹⁸O = 3.8-5.7‰) are ~3‰ lower on average relative to serpentine in mesh
cores (Fig. 6). Similarly, texturally younger calcite in zoned C2 mesh, C1 veins and
C1' domains has δ¹⁸O values down to ~10‰ (Fig. 7, 10). Low oxygen isotope ratios
are also recorded in lizardite at the boundary between serpentinite clast and carbonate
matrix (δ¹⁸O = 3.1-5.4‰), chrysotile intergrown with fine-grained C3 calcite (δ¹⁸O =
3.0-5.7‰), C3 calcite matrix (δ¹⁸O = 9.8-12.8‰), C3-veins (δ¹⁸O = 10.7-11.4‰) and
C1'-veins (δ¹⁸O = 10.9‰). Oxygen isotope thermometry from mineral pairs (see
above) suggests that such a drop in δ¹⁸O values in both serpentine and carbonates is
consistent with a crystallization temperature of up to 240 °C and water δ¹⁸O value of
2±1‰. The smooth oxygen isotope zoning in both C1 veins and C1' domains as well
as the similarity in fluid δ¹⁸O compositions between early and late stages of
hydrothermal alteration point to a continuous process that likely occurred entirely in a
seafloor environmental setting. The pervasive carbonation textures often overprint
serpentinization ones, but we do not exclude that carbonation initiated while residual

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700
anhydrous silicates were still present. The calculated oxygen isotope compositions of water are consistent with isotopically shifted seawater rather than indicative of interaction with meteoric water or metamorphic fluids in continental settings. Oxygen isotope values measured in serpentine fall in the range of oceanic serpentinites (Wenner and Taylor, 1973) and all values of δ¹³C in carbonates (from -2.6 to +2.4‰) suggest a seawater origin for carbon. It is worth noting that in a δ¹⁸O versus δ¹³C plot, the carbonates from the Chenaillet ophiolite show a distinct horizontal trend parallel to the δ¹⁸O axis (Fig. A.5) suggesting that the fraction of CO₂ in the fluid was not significant. Calcite can precipitate at a mole fraction of CO₂/(CO₂ + H₂O) as low as 0.01, but in such case a high water/rock ratio is required for a significant carbonation.

Carbonate cements commonly form at low temperature from seawater (biogenic and abiogenic) and freshwater (calcretes and speleothems). High-δ¹⁸O carbonates are assumed to form in equilibrium with very low-T seawater and our new thermometry estimates suggest that low-δ¹⁸O carbonates in the Chenaillet ophicarbonates formed at high temperature. Consequently, both temperature and water δ¹⁸O must have changed.

The flat trend can result from the dissolution and reprecipitation of pre-existing carbonates that interacted with nearly pure water such that the δ¹³C remained constant. This agrees with observations of dissolution and reprecipitation textures in the carbonates of the investigated samples.

Heating associated with burial during oceanic spreading is an unlikely scenario to explain this high-T hydrothermal event due to the intense fracturing of the ophicarbonates that points to active tectonics, accordingly to Lafay et al. (2017) who also suggested that authigenic carbonation, partially replacing serpentinites and achieving mantle minerals alteration, occurred under brittle conditions favored by the likely fluid-dominated system and tectonic context. In fact, the authors hypothesized
that the composition of the fluid and the temperature conditions evolved in a
hydrothermal environment comparable to that observed in the vicinity of mid-ocean
hydrothermal vents, i.e. active tectonic and magmatic conditions. The occurrence of a
warmer (200-240 °C) stage of hydrothermal alteration due to interaction with evolved
seawater-derived fluids (δ¹⁸O = 2-3‰) documented by the new *in situ* data could be
the result of burial by lavas causing reheating and channeling of upwelling
hydrothermal fluids along high-angle faults (Fig. 12b). Such a scenario would be
consistent with the tectonic models of Manatschal et al. (2011) for the Chenaillet
ophiolite suggesting that a first stage of mantle exhumation along a detachment fault
was followed by a stage of ‘high-angle extrusion’, when magma and fluids were
remobilized along high-angle faults acting also as feeder channels for the overlying
pillow basalts. A similar model of high-temperature hydrothermal discharge, driven
by mafic intrusions, along normal faults lying at high angle to an oceanic detachment
surface was proposed by Alt et al. (2018) in the northern Appenine ophiolites. The
authors documented an evolution of the fault system at the paleo-seafloor from talc
alteration and sulfide deposits, typical of high-temperature hydrothermal discharge (T
≈ 350 °C), to carbonated serpentinites formed at T ≈ 100 °C (Alt et al., 2018).
Carbonates formed at temperatures as high as ~ 225 °C are observed also in the
serpentinite basement of the Lost City Hydrothermal Field in the Mid-Atlantic Ridge,
where they precipitate from warm hydrothermal fluids with high concentrations of
methane and hydrogen derived from serpentinization (Früh-Green et al., 2003).

**5.2.3. Late stage veining**
The youngest C4 calcite vein records a large variability in δ¹⁸O values (12.3 to 23.9‰) and the highest average oxygen isotope composition among different carbonate generations (Fig. 7). No independent calculations of temperature and fluid oxygen isotope composition are possible due to the lack of mineral pairs. It is, therefore, necessary to assume that the variation in δ¹⁸O values is due to a variation in temperature, fluid oxygen isotope composition or more likely a combination of both.
Given the lowest δ¹⁸O values are comparable to those measured in the “high-temperature” carbonates, it is plausible that formation of the C4 vein started in similar physical-chemical conditions and continued due to interaction with progressively cooler and low- δ¹⁸O hydrothermal fluids (down to a minimum of ~40 °C, assuming δ¹⁸O(water) of -1‰) due to mixing with fresh seawater channeled along fractures at shallower depths. Such estimates are comparable, for instance, to the active vent fluids sampled at the Lost City Hydrothermal Field characterized by low temperature (T = 55-75 °C) and δ¹⁸O values (0-0.1‰), and are interpreted to result from mixing between seawater and high-pH fluids derived from subsurface serpentinization reactions (Früh-Green et al., 2003). The evolutionary model suggested here for the hydrothermal alteration of the Chenaillet ophicarbonates is consistent with the overall observation that this ophiolitic terrain was only partly affected by low-grade metamorphism during obduction.

# 6. CONCLUSIONS

We performed *in situ* analysis by SIMS of oxygen and carbon isotopes in serpentine, calcite, dolomite and magnetite from ophicarbonate samples from the Chenaillet ophiolitic complex (western Alps). SIMS analysis revealed unprecedented
zoning in oxygen isotope ratios of serpentine (3.0 to 8.8‰), calcite (9.4 to 30.6‰) and magnetite (–5.0 to –10.6‰) relative to previous studies, which relied on conventional bulk analysis. To the best of our knowledge, we document the first evidence of ‘cold carbonates’ (i.e., carbonates with δ¹⁸O > 25‰; e.g., Fig. 6 in Picazo et al., 2020 and references therein) in Alpine ophiolites. The lack of such high oxygen isotope signatures in other Alpine ophiolitic terrains could be due to the much larger sampling scale of the bulk techniques that average δ¹⁸O values of multiple mineral generations. The high spatial resolution of SIMS allowed direct quantification of oxygen isotope ratios in mineral pairs that are in textural equilibrium and are interpreted to have co-crystallized. The relative oxygen isotope homogeneity of the targeted mineral pairs is consistent with isotopic equilibrium allowing for the independent calculation of equilibrium temperatures and water oxygen isotope values. Considering that the temperature-pressure stabilities of serpentine minerals are not well constrained, *in situ* stable isotope thermometry presents a powerful tool to constrain the conditions at which alteration of ultramafic rocks occurs in various tectonic settings. Additionally, numerical models by Klein and Garrido (2011) show that temperature-dependent activity gradients of both aqueous carbon dioxide (aCO₂,aq) and aqueous silica (aSiO₂,aq) determine the carbonation potential of serpentinite. Independently constraining the formation temperature is therefore essential to estimating aSiO₂(aq), aCO₂(aq), oxygen fugacity and pH, which in turn would allow estimating the phase equilibria with implications for the monitoring of reaction progress during CO₂ sequestration of ultramafic rocks (Klein and Garrido, 2011). Nevertheless, a major source of uncertainty on the calculation of alteration temperatures, particularly for the serpentine-water system, remains the large discrepancy between different calibrations of the oxygen isotope fractionation at T <
200 °C, hence future studies should aim to improve such calibrations. The new dataset and thermometry estimates in the Chenaillet ophiolite are interpreted as indicative of a continuous hydrothermal alteration starting during uplift and exposure of the lithospheric mantle to the seafloor, subsequently followed by further carbonation and serpentinization induced by hot (up to 240 °C) and isotopically shifted (δ¹⁸O ~2‰) seawater, and finally by channeling of and mixing with low-T and low-δ¹⁸O seawater along fractures. This scenario is consistent with alteration occurring in long-lived hydrothermal systems controlled by tectonic activity where multiple metasomatic events can explain effective mantle hydration and carbonation over a few tens of thousands of years. Carbonation of ultramafic rocks in oceanic settings plays an important role for the global carbon cycle as abyssal peridotites and serpentinites might act as a sink for CO₂. Unraveling the complexity and sequence of reactions, temperatures and fluid compositions responsible for peridotite or serpentinite carbonation in oceanic and continental settings is important for correctly estimating rates of mineral carbonation and potential for CO₂ intake.

815

816

ACKNOWLEDGEMENTS

817

We thank Mike Spicuzza and Brian Hess for assistance with sample preparation, Bil Schneider and John Fournelle for assistance with SEM and EPMA analyses, Kouki Kitajima for assistance in the SIMS lab and data processing. We are grateful to Eugenio Piluso for fruitful discussions and for reading an earlier version of the manuscript. Jeffrey Alt, Frieder Klein and an anonymous reviewer are acknowledged for constructive reviews that helped improve the manuscript. We also thank Frieder Klein for editorial handling. This work is supported by the U.S.

824
825 Department of Energy, Office of Science, Office of Basic Energy Sciences

826 (Geosciences) under Awards Number DE-FG02-93ER14389 and DE-SC0020666.

827 WiscSIMS is supported by the U.S. National Science Foundation (EAR-1658823, -

828 2004618) and the University of Wisconsin- Madison.
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1103 silicates. *Earth Planet. Sci. Letters* **120**, 247–263.
1104 Zheng Y.-F., Metz P., Satir M. and Sharp Z. D. (1994) An experimental calibration of

1105 oxygen isotope fractionation between calcite and forsterite in the presence of a

1106 CO₂-H₂O fluid. *Chem. Geol.* **116**, 17–27.

1107
FIGURE AND TABLE CAPTIONS

1108

1109

1110 **Fig. 1.** (a) Tectonic map of the Piemont domain of the western Alps. (b). Geological map of the Chenaillet massif (star in Fig 1a, modified after Lafay et al., 2017) showing the location of the ophicarbonate outcrops and investigated samples.

1113 **Fig. 2.** SEM-BSE images of representative textures observed in the Chenaillet ophicarbonates. (a) Serpentine mesh textures with and without C2 calcite partially replacing serpentine in mesh cores (sample RL6). (b) Complex mesh textures showing inner and outer cores surrounded by magnetite; inner cores are locally fully replaced by C2 calcite (sample RL6). (c) Mesh inner and outer cores completely replaced by C2 calcite (sample RL6). (d) Hourglass serpentine associated with magnetite (sample RL9). (e, f) Lizardite+magnetite veinlets (V-1) following original grain boundaries (samples RL3 and RL10). (g) Large lizardite+magnetite vein (V-2) crosscutting mesh textures (sample RL6). (h) C1 calcite vein crosscutting V-2 veins (sample RL6). (i) Textural link between C2 calcite in mesh textures and C1 calcite veins (sample RL10). (j, k) Lizardite+dolomite layers within the C3 calcite matrix containing a relict spinel (Spl); C3 calcite matrix is associated with small chrysotile grains (Ctl) and sometimes crosscut by calcite veins free of serpentine (C3-vein) (sample RL6). (l) Chrysotile grains in C3 calcite matrix (sample RL6). White and red circles are oxygen (δ¹⁸O~VSMOW~, ‰) and carbon (δ¹³C~VPDB~, ‰, in parentheses) isotope SIMS pits, respectively. Srp = serpentine, Ctl = chrysotile, Lz = lizardite, Cal = calcite, Dol = dolomite, Mag = magnetite, Spl = Cr-spinel.

1130 **Fig. 3.** SEM-CL images showing representative CL features in (a-c) C1 calcite veins (samples RL9, RL3 and RL10) and (d-i) C2 calcite mesh (samples RL6, RL10 and RL9). C1 veins are sometimes rimmed by a distinct brighter (b) or darker (c) CL layer.
1133 C2 calcite mesh have marked euhedral (g, i) to anhedral (f, h) CL zonation. Resorption
1134 features are sometimes observed in C2 calcite mesh (f-i). White and red circles are
1135 locations of oxygen (δ¹⁸O<sub>VSMOW</sub>, ‰) and carbon (δ¹³C<sub>VPDB</sub>, ‰, in parentheses) isotope
1136 SIMS pits, respectively.

1137 **Fig. 4.** SEM-BSE images of representative dolomite textures in the Chenaillet
1138 ophicarbonates. Dolomite is observed as (a) rims around C2 calcite mesh (sample RL9),
1139 (b) borders along C1 calcite veins at the contact with serpentine (sample RL7), (c)
1140 subhedral grains in the lizardite layers at the boundary between serpentinite clasts and
1141 C3 calcite matrix (sample RL6), (d) borders between C4 calcite vein and serpentinite
1142 clast (sample RL9). White and red circles are oxygen (δ¹⁸O<sub>VSMOW</sub>, ‰) and carbon
1143 (δ¹³C<sub>VPDB</sub>, ‰, in parentheses) isotope SIMS pits, respectively.

1144 **Fig. 5.** SEM-CL images showing representative CL features in the carbonate matrix
1145 surrounding serpentinite clasts (samples RL6 and RL10). Domains with large calcite
1146 grains (C1') showing concentric (a-b) or oscillatory CL zoning (c) are surrounded by
1147 an aggregate of fine-grained calcite+chrysotile (a-b). (a, c) Large calcite grains in C1'
1148 domains are crosscut or bounded by calcite veinlets (C1'-veins). (d) Sharp contact
1149 between a calcite C3-vein and the surrounding fine-grained calcite matrix (C3). White
1150 and red circles in (c) and (d) are locations of oxygen (δ¹⁸O<sub>VSMOW</sub>, ‰) and carbon
1151 (δ¹³C<sub>VPDB</sub>, ‰, in parentheses) isotope SIMS pits, respectively. Dashed lines in (a) to
1152 (c) are locations of SIMS oxygen and carbon isotope traverses (see Electronic Appendix
1153 B).

1154 **Fig. 6.** Summary of oxygen isotope data in serpentine from the Chenaillet
1155 ophicarbonates measured by SIMS.

1156 **Fig. 7.** Summary of oxygen isotope data in carbonate from the Chenaillet
1157 ophicarbonates measured by SIMS.
1158 **Fig. 8.** Serpentine-magnetite thermometry. (a) Calculation of equilibrium temperatures for serpentine-magnetite pairs in different textures based on SIMS-measured mineral $\delta^{18}$O values. (b) Calculation of water oxygen isotope values in equilibrium with both serpentine and magnetite in different textures based on calculated equilibrium temperatures. Solid and dashed lines are calculated equilibrium temperatures based on serpentine-water oxygen isotope fractionation factors by Früh-Green et al. (1996) and Saccocia et al. (2009), respectively. Magnetite-water oxygen isotope fractionation factors are after Bottinga and Javoy (1973).

1166 **Fig. 9.** Serpentine-carbonate thermometry. (a) Calculation of equilibrium temperatures for lizardite+dolomite and chrysotile+calcite pairs based on SIMS-measured mineral $\delta^{18}$O values. (b) Calculation of water oxygen isotope values in equilibrium with serpentine, dolomite and calcite in each pair based on calculated equilibrium temperatures. Solid and dashed lines are calculated equilibrium temperatures for dolomite-serpentine (Matthews and Katz, 1977; Früh-Green et al., 1996) and calcite-serpentine (O'Neil et al., 1969; Früh-Green et al., 1996), respectively.

1174 **Fig. 10.** Summary of the different mineral generations and corresponding $\delta^{18}$O values ($\text{‰}$). Formation is attributed to four stages of hydrothermal alteration in the Chenaillet ophicarbonates.

1177 **Fig. 11.** Schematic illustration showing the progression of textural and $T-\delta^{18}$O(water) variations in the Chenaillet ophicarbonates. $\text{Ol}$ = olivine, $\text{Opx}$ = orthopyroxene, $\text{Cr-Spl}$ = Cr-spinel, $\text{Spc}$ = serpentine mesh core, $\text{SrP}$ = serpentine mesh rim, $\text{Mag}$ = magnetite, $\text{Lz}$ = lizardite, $\text{Dol}$ = dolomite, $\text{Ctrl}$ = chrysotile. $\delta^{18}$Ow = oxygen isotope values of water.
1182 **Fig. 12.** Simplified cross section showing the evolution of the hydrothermal system presented in this study in a seafloor tectonic setting (modified after Manatschal et al., 2011 and Lafay et al., 2017). Numbers correspond to the stages of alteration shown in Fig. 11.

1186

1187 **Table 1.** Serpentine-magnetite thermometry for ophicarbonates from the Chenaillet ophiolite.

1189 **Table 2.** Serpentine-dolomite thermometry in hourglass textures for ophicarbonates from the Chenaillet ophiolite.

1190
<table><caption>1191 Table 1. Serpentine-magnetite thermometry for ophicarbonates from the Chenaillet ophiolite.</caption><thead><tr><th>Analysis ID</th><th>Texture</th><th>δ<sup>18</sup>O (‰)<br>serpentine</th><th>δ<sup>18</sup>O (‰)<br>magnetite</th><th>Δ<sup>18</sup>O(Srp-Mag)</th><th>T (°C)*<br>calculated</th><th>δ<sup>18</sup>O<sub>w</sub> (‰)*<br>calculated</th><th>T (°C)**<br>calculated</th><th>δ<sup>18</sup>O<sub>w</sub> (‰)**<br>calculated</th></tr></thead><tbody><tr><td>RL6-1-(Srp-18b/Mag-1)</td><td>mesh</td><td>6.6</td><td>-6.0</td><td>12.7</td><td>196</td><td>4.3</td><td>245</td><td>3.1</td></tr><tr><td>RL6-1-(Srp-18b/Mag-2)</td><td>mesh</td><td>6.6</td><td>-5.6</td><td>12.2</td><td>204</td><td>4.5</td><td>252</td><td>3.4</td></tr><tr><td>RL6-1-(Srp-20/Mag-3)</td><td>mesh</td><td>6.7</td><td>-5.7</td><td>12.5</td><td>199</td><td>4.5</td><td>248</td><td>3.4</td></tr><tr><td>RL6-1-(Srp-34/Mag-4)</td><td>mesh</td><td>6.6</td><td>-10.2</td><td>16.9</td><td>136</td><td>2.1</td><td>194</td><td>0.0</td></tr><tr><td>RL6-1-(Srp-24/Mag-5)</td><td>mesh</td><td>7.8</td><td>-6.3</td><td>14.2</td><td>172</td><td>4.7</td><td>225</td><td>3.2</td></tr><tr><td>RL6-1-(Srp-25/Mag-7)</td><td>mesh</td><td>7.4</td><td>-5.5</td><td>12.9</td><td>192</td><td>5.0</td><td>242</td><td>3.7</td></tr><tr><td>RL6-1-(Srp-27/Mag-8)</td><td>mesh</td><td>6.3</td><td>-9.8</td><td>16.3</td><td>144</td><td>2.2</td><td>201</td><td>0.3</td></tr><tr><td>RL6-1-(Srp-29/Mag-9)</td><td>mesh</td><td>6.8</td><td>-6.3</td><td>13.2</td><td>188</td><td>4.2</td><td>239</td><td>2.9</td></tr><tr><td>RL6-1-(Srp-41/Mag-17)</td><td>mesh</td><td>7.2</td><td>-5.1</td><td>12.3</td><td>202</td><td>5.1</td><td>250</td><td>3.9</td></tr><tr><td>RL6-1-(Srp-42/Mag-18)</td><td>mesh</td><td>7.4</td><td>-5.6</td><td>13.1</td><td>188</td><td>4.9</td><td>239</td><td>3.6</td></tr><tr><td>RL6-1-(Srp-50/Mag-19)</td><td>mesh</td><td>5.2</td><td>-9.6</td><td>15.0</td><td>161</td><td>1.8</td><td>216</td><td>0.1</td></tr><tr><td>RL6-1-(Srp-49/Mag-21)</td><td>mesh</td><td>6.9</td><td>-10.2</td><td>17.3</td><td>132</td><td>2.2</td><td>191</td><td>0.1</td></tr><tr><td>RL6-1-(Srp-49/Mag-22)</td><td>mesh</td><td>6.9</td><td>-9.5</td><td>16.5</td><td>141</td><td>2.6</td><td>198</td><td>0.6</td></tr><tr><td>RL6-2-(Srp-28/Mag-13)</td><td>mesh</td><td>7.4</td><td>-5.0</td><td>12.6</td><td>198</td><td>5.2</td><td>247</td><td>4.0</td></tr><tr><td>RL6-2-(Srp-27/Mag-14)</td><td>mesh</td><td>7.2</td><td>-7.6</td><td>14.9</td><td>162</td><td>3.7</td><td>217</td><td>2.1</td></tr><tr><td>RL6-2-(Srp-26/Mag-15)</td><td>mesh</td><td>7.5</td><td>-7.9</td><td>15.5</td><td>154</td><td>3.7</td><td>210</td><td>2.0</td></tr><tr><td>RL6-2-(Srp25/Mag-16)</td><td>mesh</td><td>7.4</td><td>-8.0</td><td>15.6</td><td>153</td><td>3.7</td><td>209</td><td>1.9</td></tr><tr><td>RL6-2-(Srp-13/Mag-5)</td><td>V-2 vein</td><td>4.1</td><td>-10.1</td><td>14.4</td><td>169</td><td>1.0</td><td>223</td><td>-0.6</td></tr><tr><td>RL6-2-(Srp-18/Mag-10)</td><td>V-2 vein</td><td>3.8</td><td>-7.7</td><td>11.6</td><td>215</td><td>2.0</td><td>261</td><td>1.1</td></tr><tr><td>RL3-(Srp-43/Mag-5)</td><td>V-1 vein</td><td>5.8</td><td>-8.2</td><td>14.1</td><td>173</td><td>2.8</td><td>226</td><td>1.3</td></tr><tr><td>RL9-(Srp-54/Mag-16)</td><td>hourglass Srp</td><td>5.2</td><td>-9.6</td><td>14.9</td><td>162</td><td>1.8</td><td>216</td><td>0.1</td></tr><tr><td>RL9-(Srp-58/Mag-18)</td><td>hourglass Srp</td><td>4.5</td><td>-8.4</td><td>13.1</td><td>190</td><td>2.0</td><td>240</td><td>0.7</td></tr><tr><td>RL9-(Srp-57/Mag-19)</td><td>hourglass Srp</td><td>4.6</td><td>-9.1</td><td>13.8</td><td>178</td><td>1.7</td><td>230</td><td>0.3</td></tr><tr><td>RL9-(Srp-56/Mag-20)</td><td>hourglass Srp</td><td>4.5</td><td>-8.0</td><td>12.5</td><td>198</td><td>2.3</td><td>247</td><td>1.1</td></tr></tbody></table>

<footnote>* Srp-water after Früh-Green et al. (1996), ** Srp-water after Saccocia et al. (2009). Mag-water are after Bottinga and Javoy (1973). δ¹⁸O<sub>w</sub> = calculated oxygen isotope value of water in equilibrium with serpentine and magnetite.</footnote>
<table><caption>1194 Table 2. Serpentine-dolomite thermometry in hourglass textures for ophicarbonates from the Chenaillet ophiolite.</caption><thead><tr><th>Analysis ID</th><th>Texture</th><th>δ<sup>18</sup>O (‰)<br>serpentine</th><th>δ<sup>18</sup>O (‰)<br>dolomite</th><th>Δ<sup>18</sup>O(Dol-Srp)</th><th>T (°C)<br>calculated</th><th>δ<sup>18</sup>O<sub>w</sub> (‰)<br>calculated</th></tr></thead><tbody><tr><td>RL9-1-(Sr-1/Dol-1)</td><td>hourglass</td><td>4.3</td><td>11.8</td><td>7.5</td><td>228</td><td>2.9</td></tr><tr><td>RL9-1-(Sr-4/Dol-2)</td><td>hourglass</td><td>4.4</td><td>11.7</td><td>7.3</td><td>235</td><td>3.1</td></tr><tr><td>RL9-1-(Sr-6/Dol-4)</td><td>hourglass</td><td>4.1</td><td>12.6</td><td>8.5</td><td>190</td><td>1.6</td></tr><tr><td>RL9-1-(Sr-10/Dol-5)</td><td>hourglass</td><td>4.3</td><td>12.0</td><td>7.6</td><td>220</td><td>2.7</td></tr><tr><td>RL9-1-(Sr-26/Dol-6)</td><td>hourglass</td><td>4.8</td><td>12.6</td><td>7.8</td><td>215</td><td>3.0</td></tr><tr><td>RL9-1-(Sr-29/Dol-10)</td><td>hourglass</td><td>4.3</td><td>12.2</td><td>7.9</td><td>213</td><td>2.5</td></tr><tr><td>RL9-1-(Sr-46/Dol-14)</td><td>hourglass</td><td>4.3</td><td>10.1</td><td>5.8</td><td>315</td><td>4.5</td></tr><tr><td>RL9-1-(Sr-51/Dol-15)</td><td>hourglass</td><td>4.3</td><td>12.5</td><td>8.2</td><td>200</td><td>2.1</td></tr><tr><td>RL9-1-(Sr-59/Dol-19)</td><td>hourglass</td><td>4.2</td><td>11.2</td><td>7.0</td><td>250</td><td>3.2</td></tr><tr><td>RL7-(Sr-4/Dol-4)</td><td>hourglass</td><td>4.6</td><td>10.9</td><td>6.3</td><td>285</td><td>4.3</td></tr><tr><td>RL7-(Sr-10/Dol-2)</td><td>hourglass</td><td>5.6</td><td>11.8</td><td>6.2</td><td>290</td><td>5.4</td></tr><tr><td>RL7-(Sr-9/Dol-1)</td><td>hourglass</td><td>5.0</td><td>11.6</td><td>6.6</td><td>270</td><td>4.4</td></tr></tbody></table>

1195 Srp-water after Früh-Green et al. (1996), Dol-water after Matthews and Katz (1977). δ¹⁸O<sub>w</sub> = calculated oxygen isotope value of water in equilibrium with serpentine and dolomite.

1196 dolomite.

1197
45°20′N

45°00′N

44°40′N

06°50′E

07°00′E

<figure><img src="image_2.png" /></figure>

Chenaillet

Piemont unit

<figure><img src="image_3.png" /></figure>

External crystalline massifs

<figure><img src="image_4.png" /></figure>

External cover

<figure><img src="image_5.png" /></figure>

Briançonnais zone

<figure><img src="image_6.png" /></figure>

Internal crystalline massifs

<figure><img src="image_7.png" /></figure>

Low- to High-temperature Blueschist facies conditions

<figure><img src="image_8.png" /></figure>

Eclogite facies

<figure><img src="image_9.png" /></figure>

Ophiolite bodies

Main faults

River

<figure><img src="image_10.png" /></figure>

Chenaillet Massif

Scree

Ophicalcite

Dolerite / Basalt

<figure><img src="image_11.png" /></figure>

2649

Grand Charvia

FRANCE

2400

2500

500 m
<figure><img src="image_12.png" /></figure>
<figure><img src="image_13.png" /></figure>

<figure><img src="image_14.png" /></figure>

<figure><img src="image_15.png" /></figure>

<figure><img src="image_16.png" /></figure>

<figure><img src="image_17.png" /></figure>

<figure><img src="image_18.png" /></figure>

<figure><img src="image_19.png" /></figure>

<figure><img src="image_20.png" /></figure>

<figure><img src="image_21.png" /></figure>
<figure><img src="image_22.png" /></figure>

<figure><img src="image_23.png" /></figure>

<figure><img src="image_24.png" /></figure>

<figure><img src="image_25.png" /></figure>
<figure><img src="image_26.png" /><figcaption>C1'-vein</figcaption></figure>

<figure><img src="image_27.png" /><figcaption>Srp</figcaption></figure>

<caption>100 µm</caption>

<caption>100 µm</caption>

<figure><img src="image_28.png" /><figcaption>C1'-vein</figcaption></figure>

<figure><img src="image_29.png" /><figcaption>C3-Cal vein</figcaption></figure>

<caption>11.6
(-0.3)</caption>

<caption>11.3
(0.0)</caption>

<caption>50 µm</caption>

<caption>50 µm</caption>

<caption>C3 Cal</caption>
<figure><img src="image_30.png" /></figure>
<figure><img src="image_31.png" /></figure>

RL6

RL10

RL7

RL3

RL9
<figure><img src="image_32.png" /></figure>

<figure><img src="image_33.png" /></figure>
<figure><img src="image_34.png" /></figure>

<figure><img src="image_35.png" /></figure>
<table><thead><tr><th>Stages<br>Minerals</th><th>Early<br>Serpentinization:<br>T<sub>min</sub> = 130-200 °C</th><th>Low-T<br>Carbonation:<br>T<sub>min</sub> = 10 °C</th><th>High-T Serpentinization &amp; Carbonation:<br>T = 160-240 °C</th><th>Late Carbonate<br>Veining:<br>T<sub>min</sub> = 40-70 °C</th></tr></thead><tbody><tr><td>Serpentine V-1 veins</td><td>3.5 to 6.4‰</td><td></td><td></td><td></td></tr><tr><td>Serpentine mesh rims</td><td>4.5 to 8.8‰</td><td></td><td></td><td></td></tr><tr><td>Serpentine mesh cores</td><td>4.4 to 8.5‰</td><td></td><td></td><td></td></tr><tr><td>Magnetite V-1 veins</td><td>-8.5 to -7.3‰</td><td></td><td></td><td></td></tr><tr><td>Magnetite mesh</td><td>-10.6 to -5.0‰</td><td></td><td></td><td></td></tr><tr><td>Calcite C1 veins</td><td></td><td>22.3 to 9.4‰</td><td></td><td></td></tr><tr><td>Calcite C2 mesh</td><td></td><td>30.6 to 9.8‰</td><td></td><td></td></tr><tr><td>Calcite C1'</td><td></td><td>24.9 to 10.6‰</td><td></td><td></td></tr><tr><td>Serpentine V-2 veins</td><td></td><td></td><td>3.2 to 4.2‰</td><td></td></tr><tr><td>Serpentine hourglass</td><td></td><td></td><td>3.8 to 5.7‰</td><td></td></tr><tr><td>Magnetite V-2 veins</td><td></td><td></td><td>-10.6 to -7.7‰</td><td></td></tr><tr><td>Magnetite hourglass</td><td></td><td></td><td>-9.6 to -8.0‰</td><td></td></tr><tr><td>Serpentine matrix</td><td></td><td></td><td>3.1 to 5.4‰</td><td></td></tr><tr><td>Serpentine in C3 matrix</td><td></td><td></td><td>3.0 to 5.7‰</td><td></td></tr><tr><td>Calcite C3 matrix</td><td></td><td></td><td>9.8 to 12.9‰</td><td></td></tr><tr><td>Calcite C3-veins</td><td></td><td></td><td>10.5 to 11.5‰</td><td></td></tr><tr><td>Calcite C1'-veins</td><td></td><td></td><td>10.9 to 11.8‰</td><td></td></tr><tr><td>Dolomite</td><td></td><td>10.1 to 14.1‰</td><td>11.4 to 13.2‰</td><td>11.2 to 12.5‰</td></tr><tr><td>Calcite C4 vein</td><td></td><td></td><td></td><td>12.3 to 26.0‰</td></tr><tr><td>Serpentine in C4 vein</td><td></td><td></td><td></td><td>4.0 to 4.8‰</td></tr></tbody></table>
① peridotite

<figure><img src="image_36.png" /></figure>

② serpentinite
 (mesh texture + V-1 veins)

<figure><img src="image_37.png" /></figure>

⑤ ophicarbonate
(high δ¹⁸O carbonate)

<figure><img src="image_38.png" /></figure>

<figure><img src="image_39.png" /></figure>

④ ophicarbonate
(low δ¹⁸O carb+srp)

<figure><img src="image_40.png" /></figure>

③ ophicarbonate
(high δ¹⁸O carbonate)

<figure><img src="image_41.png" /></figure>
(a)

detachment fault

seawater fluid path

serpentinite

ophicarbonate

3

2

gabbro

peridotite ①

<figure><img src="image_42.png" /></figure>

(b)

hydrothermal discharge

system

seawater fluid path

basalt

5

3

4

high-angle faults

peridotite ①

<figure><img src="image_43.png" /></figure>

gabbro

evolved
fluids

<figure><img src="image_44.png" /></figure>