# Magnetite production in mesh texture during serpentinization, a marker of H_2 diffusion

Benjamin Malvoisin, Paula Dörfler, Anne-Line Auzende, Fabrice Brunet, Mathilde Cannat, Håkon Austrheim, Mary-Alix Kaczmarek

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Benjamin Malvoisin, Paula Dörfler, Anne-Line Auzende, Fabrice Brunet, Mathilde Cannat, et al. Magnetite production in mesh texture during serpentinization, a marker of H_2 diffusion. Journal of Petrology, In press, 10.1093/petrology/egag021. hal-05523873

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# Magnetite production in mesh texture during serpentinization, a marker of H₂ diffusion

Benjamin Malvoisin<sup>1,*</sup> Paula Dörfler,<sup>1</sup> Anne-Line Auzende,<sup>1</sup> Fabrice Brunet,<sup>1</sup> Mathilde Cannat,<sup>2</sup> Håkon Austrheim<sup>3</sup> and Mary-Alix Kaczmarek<sup>4</sup>

<sup>1</sup>ISTerre, Univ. Grenoble Alpes, Univ. Savoie Mont Blanc, CNRS, IRD, Univ. Gustave Eiffel, rue de la Piscine, 38000, Grenoble, France, <sup>2</sup>Institut de Physique du Globe, Université Paris Cité, CNRS, Paris, France, <sup>3</sup>Physics of Geological Processes (PGP), The Njord Centre, Department of Physics, University of Oslo, P.O. Box 1048, NO-0316, Oslo, Norway and <sup>4</sup>Géosciences Environnement Toulouse (GET), CNRS-CNES-IRD-Université Toulouse III, Observatoire Midi-Pyrénées, Toulouse, France

*Corresponding author. benjamin.malvoisin@univ-grenoble-alpes.fr

## Abstract

The detailed mineralogical study of serpentinized peridotites collected at 13 localities at mid-ocean ridges, in ophiolites and ultramafic bodies shows the formation of two alteration zones around olivine grains. At the olivine contact, a fine-grained mixture of serpentine, Fe-brucite ($\frac{Mg}{Mg+Fe}$ molar ratio comprised between 0.66 and 0.82) and awaruite occurs (Reaction Zone 1). X-ray mapping indicates limited mass transfer during Reaction Zone 1 formation, suggesting isochemical serpentinization except for the addition of water. The measured distribution of Fe and Mg between serpentine and brucite in Reaction Zone 1 is well reproduced with thermodynamic modelling incorporating the latest data for the $Fe(OH)_2$ endmember. Thermodynamic modelling also reveals that, at low water to rock ratio, awaruite formation limits H₂ production to values more than one order of magnitude lower than previous estimates. The predicted $H_{2,aq}$ concentrations are comprised between $10^{-3}$ and $10^{-2}$ mol/kg, that is in the same range than the maximum values measured in fluids expelled at ultramafic-hosted hydrothermal sites. At a water to rock ratio of 1, the updated thermodynamic model only predicts magnetite formation after olivine at temperatures above 300 °C, that is above the temperature estimates for serpentinization based on published oxygen isotope data. Nevertheless, a second assemblage composed of Ni-bearing magnetite, serpentine and Mg-brucite ($\frac{Mg}{Mg+Fe}$ molar ratio comprised between 0.83 and 0.98) can be found at the mesh rim (Reaction Zone 2). Reaction Zone 2 can display a symplectite microtexture. Transmission electron microscopy reveals the presence of porosity and Fe-brucite relics in the symplectite, suggesting Reaction Zone 2 formation after Reaction Zone 1 by dissolution-precipitation. Significant mass transfer is observed in Reaction Zone 2 at the mesh rim, probably associated with the solid volume decrease of 7 % during Reaction Zone 1 reaction to form Reaction Zone 2. The mineralogy and the composition of Reaction Zone 2 are reproduced with thermodynamic modelling of olivine serpentinization not by increasing the water to rock ratio but rather by removing H₂ from the system. This indicates that H₂ diffusion is the main driver for magnetite formation during serpentinization. The $H_{2,aq}$ concentrations at the equilibrium with Reaction Zone 2 fall in the $10^{-7} - 10^{-3}$ mol/kg range. The relative proportion of Reaction Zone 1 and Reaction Zone 2 in serpentinized peridotites has first-order impacts on H₂ production.

## Introduction

The interaction of ultramafic rocks with water (i.e. serpentinization) occurs at mid-ocean ridges [Cannat, 1993], in ophiolites [Neal and Stanger, 1983],
in subduction zones [Vitale Brovarone et al., 2020] and on extraterrestrial bodies [Vance and Daswani, 2020]. During this reaction, Fe²⁺-bearing minerals (olivine, pyroxene) react to form both Fe²⁺- and Fe³⁺-bearing minerals such as serpentine ((Mg, Fe²⁺, Fe³⁺)₃(Si, Fe³⁺)₂O₅(OH)₄), magnetite (Fe₃O₄), brucite ((Mg, Fe)(OH)₂), clinopyroxene (Ca(Mg, Fe)Si₂O₆) and garnet (Ca₃Fe₂³⁺(SiO₄)₃) [Moody, 1976, McCollom and Bach, 2009, Plümper et al., 2014]. Among the products of serpentinization, magnetite has attracted significant attention. This mineral is a ferromagnetic mineral producing a strong magnetic signal (92 A.m².kg⁻¹). Rock magnetism studies provide insights on reaction extent, magnetite structure and Fe mobility during serpentinization from the micrometer and centimeter scale [Toft et al., 1990, Oufi et al., 2002, Maffione et al., 2014, Bonnemains et al., 2016, Hong et al., 2022] up to the hundred of meters to kilometer scale above ophiolites and hydrothermal sites [Tominaga et al., 2017, Szitkar et al., 2017, Szitkar and Murton, 2018, Michels et al., 2018]. Magnetite (U-Th-Sm)/He thermochronometry can also be used as a geochronology tool in ultramafic rocks [Cooperdock and Stockli, 2016, Schwartz et al., 2020, Corre et al., 2023]. Finally, the electrons released during Fe oxidation associated with magnetite formation can be transferred to water, leading to the formation of natural hydrogen (H₂), a typical component of natural fluids in serpentinizing environments [Neal and Stanger, 1983, Abrajano et al., 1990, Charlou et al., 2002, Leong et al., 2023, Etiope, 2023]. H₂ is an energy source for chemosynthetic communities which could be involved in the emergence of life [Shock, 1990, Martin et al., 2008]. H₂ is also a strong reducing agent allowing for the formation of a reduced environment, marked by the presence of Fe-Ni alloys such as awaruite in serpentinized peridotites [Schwarzenbach et al., 2014, Evans et al., 2023].

The fate of magnetite (and H₂) during serpentinization depends on Fe distribution. Thermodynamic modelling of olivine (Mg/Mg+Fe) ratio, X<sub>Mg</sub> = 0.9) serpentinization at temperatures below 350 °C allows prediction of the Fe content of serpentinization
products [McCollom and Bach, 2009]. Brucite displays the largest range of $X_{Mg}$ among the reaction products. $X_{Mg}$ of brucite varies from ~ 0.8 to ~ 0.95 as the temperature increases from 25 to 350 °C or the water to rock ratio increases from 0.2 to 100 [McCollom and Bach, 2009]. The abundances of brucite and magnetite also display the largest variations. Whereas the mode of brucite decreases as temperature and water to rock ratio increase, the mode of magnetite increases [Klein et al., 2009]. The thermodynamic studies thus indicate that Fe distribution during serpentinization is mainly controlled by the following reaction relationship:

$$3 [Fe(OH)_2]^{\text{brucite}} = Fe_3O_4 + H_2 + 2H_2O \quad (1)$$

Klein et al. [2014] combined thermodynamic modelling of reaction 1 with petrographic, isotopic and magnetic data to propose that magnetite production during serpentinisation is mainly controlled by temperature, with magnetite production mainly occurring above 200 °C. This statement is seemingly in agreement with magnetization measurements one order of magnitude lower at the low-temperature (90 °C) Lost City hydrothermal field than at the high-temperature (350 °C) Rainbow hydrothermal field [Szitkar et al., 2017, Szitkar and Murton, 2018]. However, recent experimental work on reaction 1 by Carlin et al. [2024] has allowed an update on thermodynamic parameters for $Fe(OH)_2$. The calculated standard state enthalpy and third law entropy differ by more than 1 % from the values given in McCollom and Bach [2009], used in most thermodynamic studies of ultramafic rocks [Klein et al., 2009, 2013, 2014, Leong and Shock, 2020, Ely et al., 2023]. Such large changes in thermodynamic properties modify the predicted stability and composition of brucite, and may thus involve strong modifications of Fe distribution predictions for serpentinization. In particular, this may have important consequences for magnetite stability. The impact of the new thermodynamic data on Fe distribution still needs to be clarified.
At the micrometer to centimeter scale, olivine serpentinization at low pressure at mid-ocean ridges and in ophiolites has been observed to produce two reaction zones. Fe-bearing brucite (Fe-brucite) and serpentine are observed at the contact with olivine, whereas magnetite and serpentine are reported at the mesh rim [Bach et al., 2006, Katayama et al., 2010, Beard et al., 2009, Frost et al., 2013]. The parameter controlling magnetite formation in such reaction zones remains under debate. Modifications of temperature [Li et al., 2020] and the achievement of increasingly open system conditions during reaction [Bach et al., 2006, Beard et al., 2009, Frost et al., 2013] have been proposed to explain the observed microtexture. The strong increase in magnetic susceptibility at reaction progress above ~ 60 % measured in serpentinized peridotites collected at the Mid-Atlantic Ridge [Oufi et al., 2002], in the Mirdita ophiolite [Maffione et al., 2014] and in Oman [Hong et al., 2022] supports the model of increasingly open system conditions during reaction [Bach et al., 2006, Frost et al., 2013]. However, the exact parameter or combination of parameters controlling magnetite formation under fluid-dominated conditions remains unclear. Increase in silica activity due to orthopyroxene breakdown [Bach et al., 2006, Beard et al., 2009], increase in fluid flow (i.e. increase in water to rock ratio) and H₂ diffusion out of the system [Frost et al., 2013] have been proposed as potential drivers for magnetite formation. Weathering on the seafloor and in ophiolites could also play a role in magnetite formation. Brucite is found to dissolve during weathering [Ulven et al., 2017, Klein et al., 2020], and the released Fe can precipitate in Fe (hydr-)oxides in the formed pores [Jöns et al., 2017]. Finally, a model of magnetite formation should also explain the wide diversity of magnetite content-serpentinization degree relationship measured in natural samples. Susceptibility increase with serpentinization degree is three times more pronounced in samples from Mid-Atlantic Ridge [Oufi et al., 2002] than from Oman [Hong et al., 2022]. Magnetic susceptibility can also linearly increase with serpentinization degree
[Li et al., 2020] or remains low at all serpentinization degrees in Pindos ophiolite
[Bonnemains et al., 2016] and Mariana forearc [Klein et al., 2014] samples.

Here, we use observations in natural samples collected at thirteen locations at mid-ocean ridges, in ophiolites and in ultramafic bodies to determine the parameter(s) controlling Fe distribution between magnetite and Fe-brucite in serpentinized peridotites. We show that the alteration products of olivine serpentinization are generally an assemblage of Fe-rich brucite + serpentine + awaruite in which the phases are mixed at the submicrometer scale. The measured compositions of the Fe-brucite + serpentine assemblages are compared to thermodynamic predictions with different thermodynamic properties for Fe(OH)₂. The serpentine + Fe-brucite assemblage is sometimes bounded by veins of magnetite + Mg-rich brucite at the mesh rim. The observed alteration sequence of reaction products is only reproduced with thermodynamic modelling when considering equilibration under a RedOx gradient maintained at the microscale through hydrogen diffusion.

# Methods

## Natural samples

We study thin sections of ultramafic samples from five ophiolites (Marum ophiolite and Papuan ultramafic belt in Papua New Guinea, Leka ophiolite complex in Norway, Chenaillet ophiolite in the French Alps, Oman ophiolite, and Mirdita ophiolite in Albania), one mantle fragment (Feragen ultramafic body in Norway) and three mid-ocean ridges (Southwest Indian Ridge, Mid-Atlantic Ridge and Hess Deep in the Pacific at the East of the Galapagos triple junction; Table 1). In all the studied samples, fluid-rock interaction is dominated by serpentinization. Samples collected at mid-ocean ridges were exhumed along large offset normal faults, associated with the Cocos-Nazca rift in the case of Hess Deep [Lonsdale, 1988], and detachment faults in the case of the Mid-Atlantic Ridge and Southwest Indian Ridge samples [Cann
et al., 1997, Sauter et al., 2013]. Detailed microtextural observations and oxygen isotope data in such samples suggest that serpentinization probably started during exhumation with hydrothermal fluids introduced in the system soon after thermal cracking [Plümper et al., 2012, Rouméjon and Cannat, 2014]. Among the studied samples, some samples from ODP Leg 209 (Mid-Atlantic Ridge) are the only ones in which fluid-rock interaction associated with black smoker activity has been recorded [Paulick et al., 2006]. Samples from ophiolites experienced serpentinization both on the seafloor and later after ophiolite emplacement during meteoritic water infiltration. An overprint by high-pressure metamorphism is recorded by antigorite laths and secondary olivine in the samples from the Leka ophiolitic complex [Iyer et al., 2008]. The areas containing antigorite and secondary olivine are not selected for the analysis performed here.

The studied samples have dunite and harzburgite protoliths displaying a varied degree of serpentinization ranging from approximately 5 to 100 % based on modal analysis with a microscope. We focus here on serpentine + brucite ± magnetite assemblages in the vicinity of olivine grains (mesh texture) to determine the composition and microtexture of the minerals formed during serpentinization. Fluid-rock interaction during serpentinization may lead to serpentine recrystallization in later vein generations [Andreani et al., 2007, Rouméjon et al., 2015]. These vein generations are not investigated here. Studied samples were collected during fieldwork on-land, during dives with submersibles, by dredging and by drilling in the framework of international drilling programs (Oman drilling project and ODP Leg 209). More details on the geology of the studied samples can be found in the references provided in Table 1.

In addition to the natural samples studied here, compositional data for serpentine/brucite mixture from the literature are also used (Table 1). They are provided by the studies of Klein et al. [2014] (samples from international drilling
<table><thead><tr><th>Sample names</th><th>Geological context</th><th>Sampling</th><th>Location</th><th>Protolith</th><th>Presence of brucite</th><th>Olivine alteration zone(s)</th><th>References</th></tr></thead><tbody><tr><td>A5, A4b, B1, 5</td><td>Ophiolite</td><td>Fieldwork</td><td>Marum ophiolite and Papuan ultramafic belt (Papua New Guinea)</td><td>Dunite</td><td>x</td><td>RZ1 + RZ2</td><td>Kaczmarek et al. (2015), Malvoisini et al. (2017)</td></tr><tr><td>LE26, LE27, LE29</td><td>Ophiolite</td><td>Fieldwork</td><td>Leka ophiolite complex (Norway)</td><td>Dunite</td><td>x</td><td>RZ1</td><td>yer et al. (2008)</td></tr><tr><td>PR1, PR2, CD01</td><td>Ophiolite</td><td>Fieldwork</td><td>Chenailet ophiolite (France)</td><td>Harzburgite</td><td></td><td></td><td>Schwartz et al. (2013)</td></tr><tr><td>BA4A-81-1-1-17</td><td>Ophiolite</td><td>Drilling</td><td>Oman Drilling Project Hole BA4A</td><td>Dunite</td><td>x</td><td>RZ1</td><td>Malvoisini et al. (2020), Kelemen et al. (2021)</td></tr><tr><td>807-C57088-55-02-45-50, 807-C57088-56-03-21-26, 807-C57088-57-03-43-48, 807-C57088-57-04-35-40, 807-C57088-58-01-2-6, 807-C57088-61-01-33-38</td><td>Ophiolite</td><td>Drilling</td><td>Oman Drilling Project Hole CM2B</td><td>Dunite</td><td>x</td><td>RZ1 + RZ2</td><td>Kelemen et al. (2021)</td></tr><tr><td>96AL-8, 96AL-10B, 96AL-10dl, 96AL-39C</td><td>Ophiolite</td><td>Fieldwork</td><td>Mirdita ophiolite (Albania)</td><td>Dunite</td><td>x</td><td>RZ1</td><td>Nicolas et al. (1999)</td></tr><tr><td>NC09-4A</td><td>Ophiolite</td><td>Fieldwork</td><td>New Caledonia ophiolite</td><td>Dunite</td><td>x</td><td>RZ1</td><td>Frost et al. (2013)</td></tr><tr><td>SE10-12</td><td>Ophiolite</td><td>Fieldwork</td><td>Santa Elena ophiolite (Costa Rica)</td><td>Harzburgite</td><td>x</td><td>RZ1 + RZ2</td><td>Schwarzenbach et al. (2016)*</td></tr><tr><td>10ZD-6-18, 15ZD-06, 10ZD-6-7, 10ZD-6-19, 10ZD-6-6, 15ZD-03</td><td>Ophiolite</td><td>Fieldwork</td><td>Zedang ophiolite (Southern Tibet)</td><td>Harzburgite + Hlerzolite</td><td>x</td><td>RZ1</td><td>Li et al. (2020)*</td></tr><tr><td>un.</td><td>Ophiolite</td><td>Fieldwork</td><td>Mineoka ophiolite</td><td>Harzburgite + Dunite</td><td>x</td><td>RZ1</td><td>Katayama et al. (2010)*</td></tr><tr><td>R@DR03-07, R@DR03-11A</td><td>Ophiolite</td><td>Fieldwork</td><td>Feragen ultramafic body (Norway)</td><td>Dunite</td><td>x</td><td>RZ1 + RZ2</td><td>Pryadunenko et al. (2022), Beilich and Austrheim (2012)</td></tr><tr><td>SMS-DR05-03-26A, SMS-DR11-03-10, SMS-DR34-2-25, SMS-DR35-5-9, SMS-DR10-4-5, SMS-DR28-4-8, SMS-DR29-6-1, SMS-DR17-4-29, SMS-DR17-4-48, SMS-DR17-4-32A, SMS-DR17-4-66-C</td><td>Mid-ocean ridge</td><td>Dredging</td><td>Southwest Indian Ridge</td><td>Harzburgite</td><td>x</td><td>RZ1</td><td>Rouméjon et al. (2014), Rouméjon et al. (2015), Bickert et al. (2021)</td></tr><tr><td>ODM-ROC-V561-OR420, ODR-V561-421, ODM-DR15-2-05</td><td>Mid-ocean ridge</td><td>Submersible + Dredging</td><td>Mid-Atlantic Ridge (Odemar Cruise)</td><td>Dunite</td><td></td><td></td><td>Escartin et al. (2017)</td></tr><tr><td>SE-DR01-7-7, SE-DR05-1-11, SE-DR05-2-16a, SE-DR02-19, SE-DR06-1-3, SE-DR07-1-10, SE-DR07-1-13</td><td>Mid-ocean ridge</td><td>Dredging</td><td>Mid-Atlantic Ridge (Serpentine Cruise)</td><td>Harzburgite</td><td></td><td></td><td>Picazo et al. (2012)</td></tr><tr><td>1274A-SR1-10, 1274A-SR1-16</td><td>Mid-ocean ridge</td><td>Drilling</td><td>Mid-Atlantic Ridge (ODP Leg 209)</td><td>Dunite + Harzburgite</td><td>x</td><td>RZ1 + RZ2</td><td>Bach et al. (2004)</td></tr><tr><td>HS-13-5, HS-13-3C, HS-13-5b, HS-13-12, HS-13-3C</td><td>Mid-ocean ridge</td><td>Submersible</td><td>Mid-Atlantic Ridge (MARK area, Hydrosnake cruise)</td><td>Harzburgite</td><td></td><td></td><td>Mével et al. (1991)</td></tr><tr><td>N2-9-7, N2-17-2b</td><td>Mid-ocean ridge</td><td>Submersible</td><td>Pacific - Hess Deep (East of Galapagos Triple Junction)</td><td>Dunite</td><td>x</td><td>RZ1</td><td>Hekininan et al. (1993)</td></tr><tr><td>82-55BZ-42R-1W-107-109</td><td>Mid-ocean ridge</td><td>Drilling</td><td>Mid-Atlantic Ridge (Azores Platform; DSDP Leg 82)</td><td>Harzburgite + Dunite</td><td>x</td><td>RZ1</td><td>Klein et al. (2014)*</td></tr><tr><td>153-920B-1W-3W-64-66, 153-920B-10R-1W-82-86, 153-920B-10R-2W-68-70, 153-920B-2R-1W-80-82, 153-920B-12R-2W-140-143</td><td>Mid-ocean ridge</td><td>Drilling</td><td>Mid-Atlantic Ridge (MARK area, ODP Leg 153)</td><td>Harzburgite + Dunite</td><td>x</td><td>RZ1 + RZ2</td><td>Klein et al. (2014)*</td></tr><tr><td>147-895D-2R-1W-29-31, 147-895D-7R-3W-13-15</td><td>Mid-ocean ridge</td><td>Drilling</td><td>Pacific - Hess Deep (East of Galapagos Triple Junction, ODP Leg 147)</td><td>Harzburgite + Dunite</td><td>x</td><td>RZ1</td><td>Klein et al. (2014)*</td></tr><tr><td>107-651A-58R-2W-40-42</td><td>Passive margin</td><td>Drilling</td><td>Tyrhrenian sea (ODP Leg 107)</td><td>Harzburgite + Dunite</td><td>x</td><td>RZ1</td><td>Klein et al. (2014)*</td></tr><tr><td>149-897D-17R-6W-1-3, 149-897D-23R-1W-17-20, 149-897D-24R-3W-103-106, 173-1068A-24R-1W-73-74, 173-1068A-24R-2W-56-59, 173-1068A-25R-1W-104-106</td><td>Passive margin</td><td>Drilling</td><td>Iberia margin (ODP Legs 149 and 173)</td><td>Harzburgite + ODP Leg 173</td><td>x (only in Harzburgite)</td><td>RZ1</td><td>Klein et al. (2014)*</td></tr><tr><td>125-779A-05R-2W-35-38, 125-779A-10R-2W-51-53, 125-779A-26R-3W-50-52</td><td>Forearc</td><td>Drilling</td><td>Mariana forearc (ODP Leg 125)</td><td>Harzburgite + Dunite</td><td>x</td><td>RZ1</td><td>Klein et al. (2014)*</td></tr></tbody></table>

Notes: un.: unknown sample name; *: reference from which data of serpentine/brucite mixture compositions were retrieved. All the other data were acquired in this study.

<caption>**Table 1.** List and general characteristics of the samples used in this study.</caption>

programs of DSDP Leg 82, ODP Leg 107, ODP Leg 125, ODP Leg 147, ODP Leg 153, ODP Leg 173), Schwarzenbach et al. [2016] (Santa Elena ophiolite, Costa Rica), Frost et al. [2013] (New Caledonia ophiolite), Li et al. [2020] (Zedang ophiolite, Southern Tibet) and Katayama et al. [2010] (Mineoka ophiolite complex, Japan). The analyses of Klein et al. [2014] for ODP Leg 153 are split in two datasets due to two clearly distinguishable trends on a Fe vs. Si diagram. The analyses of Klein et al. [2014] for ODP Leg 149 are excluded from the compilation due to the absence of brucite-rich
compositions in the dataset (Si (apfu, on the basis of 7 oxygens) > 1.7 for all the analyses in the dataset). The analyses of Schwarzenbach et al. [2016] corresponding to mixed analyses between magnetite and serpentine/brucite mixture are excluded from the data analysis performed here.

## Isotope temperature estimates

A compilation of oxygen isotope data for serpentine (lizardite and chrysotile) is built to provide estimates of serpentinization temperatures at low pressure at mid-ocean ridges, passive margins and in ophiolites (Supplementary Tab. S1, references therein).

$\delta^{18}\text{O}_{\text{VSMOW}}$ data are converted to temperatures with the calibration of Vho et al. [2019] by assuming serpentinization by seawater with $\delta^{18}\text{O}_{\text{VSMOW}} = 0‰$. The isotopic composition of serpentinizing fluids can evolve during serpentinization and hydrothermal processes at mid-ocean ridges and passive margins [Alt et al., 2007].

Vesin et al. [2023, 2024] suggested that temperature is the dominant factor controlling $\delta^{18}\text{O}$ values of serpentine in samples collected in these environments. However, one should keep in mind that $\delta^{18}\text{O}$ serpentine geothermometers are not applicable at the low fluid to rock ratios sometimes inferred in serpentinized peridotites [Rouméjon et al., 2015]. In ophiolites, fluid $\delta^{18}\text{O}$ can be influenced by meteoritic water influx [Scicchitano et al., 2021]. The temperature estimates calculated here are thus first order estimates associated with significant uncertainty [up to 70 °C; Rouméjon et al., 2015].

## Raman spectroscopy

Raman spectroscopy point analysis is performed on petrographic thin sections with a LabRAM Soleil confocal Raman microscope at ISTerre (Grenoble). We use an excitation laser beam of 532 nm, a laser power comprised between 1.8 and 11 mW at the sample surface and a diffraction grating of 1800 gr/mm. Spectra are acquired both in the 100-1200 and in the 3550-3800 cm$^{-1}$ regions.
Electron microprobe analysis and X-ray maps

Point analyses and X-ray maps are acquired at 15 keV with a JEOL JXA-iHP200F field emission electron probe microanalyzer (EPMA) at ISTerre (Grenoble). Magnetite is analysed at 20 nA with a spot-shaped beam. Point analyses of silicates (serpentine/brucite mixture and olivine) are acquired at 2 nA with a beam size equal or above 2 µm to limit surface damage during analysis. Structural formulae of olivine and serpentine/brucite mixtures are calculated on a basis of 4 and 7 oxygens, respectively. Magnetite structural formulae are calculated on a basis of 3 cations and 4 oxygens. Based on the work of Beard and Frost [2017], Al in serpentine is considered to be equally distributed between tetrahedral and octahedral sites (Tschermak's substitution). As a result, the sum of cations in the octahedral site is expressed here as: Mg + Fe + Mn + Ca + Ni + Al/2 (apfu), and the sum of cations in the tetrahedral site as: Si + Al/2 (apfu). Serpentine is known to incorporate significant amounts of Fe³⁺ through ferri-Tschermak's or cronstedtite substitutions [Beard and Frost, 2017]. Fe³⁺ in serpentine is not considered here in the calculation of serpentine structural formulae but it is taken into account in the thermodynamic model (see section 2.6). Assuming that the composition of brucite does not vary spatially in each thin section, the mean brucite and serpentine compositions can be determined at the intercepts of the linear fit of the Fe vs. Si + Al/2 data at Si + Al/2 = 0 and Si + Al/2 = 2.0, respectively [Frost et al., 2013, Beard and Frost, 2017]. Datasets acquired on different thin sections at the same locality and providing similar brucite and serpentine compositions are grouped together to provide greater statistical power on composition. This strategy is applied at all studied localities except the Mirdita ophiolite for which two distinct mixture compositions are measured. Additionally, datasets for which calculated Si (apfu) is always above 1.7 are not considered for composition calculation due to high fitting error with a lack of data points near the pure brucite endmember.
Errors for all the parameters determined above by fitting are given as the 95 % confidence interval bounds, or as propagated errors with these bounds.

The X-ray maps are acquired with a probe current and a dwell time of either 30 nA and 60 ms or 100 nA and 20 ms. The number of pixels and the size of each map are provided in Supplementary Tab. S2. In the map regions, spot analyses are also acquired to provide internal standards. They are then used to convert the count number of the X-ray maps into oxide mass and molar compositional maps with the XMapTools software v. 4.3. [Lanari et al., 2019]. The software allows extraction of bulk rock composition from a selected area and structural formula at each pixel. The structural formulae of serpentine/brucite mixtures are used to determine the molar fraction of brucite in the mixture (X<sub>Br</sub>) by using the formula provided in Malvoisin et al. [2020]:

$$X_{\text{Br}} = \frac{n_{\text{Mg}} + n_{\text{Fe}} + n_{\text{Ni}} - 3/2n_{\text{Si}}}{n_{\text{Mg}} + n_{\text{Fe}} + n_{\text{Ni}} - n_{\text{Si}}} \quad (2)$$

with n<sub>Mg</sub>, n<sub>Fe</sub>, n<sub>Ni</sub> and n<sub>Si</sub> the number of moles of each element in the mixture calculated with XMapTools software. For the maps, Mg over (Mg + Fe) molar ratio of brucite in the mixture (X<sup>b</sup><sub>Mg</sub>) is determined by assuming fixed serpentine composition in the mixture. Mg fraction in serpentine (X<sup>s</sup><sub>Mg</sub>) is determined based on the regression of point analyses as described above for each locality. X<sup>b</sup><sub>Mg</sub> is then calculated with the following formula:

$$X_{\text{Mg}}^{\text{b}} = \frac{n_{\text{Mg}} - \frac{3n_{\text{Si}}}{2} X_{\text{Mg}}^{\text{s}}}{n_{\text{Mg}} + n_{\text{Fe}} - \frac{3n_{\text{Si}}}{2}} \quad (3)$$

X<sup>b</sup><sub>Mg</sub> is only computed for n<sub>Si</sub> < 1.7, that is excluding mixtures largely dominated by serpentine. For these latter mixtures, small variations in X<sup>s</sup><sub>Mg</sub> can induce large errors on X<sup>b</sup><sub>Mg</sub> estimated from equation 3.
XMapTools software is also used to determine local bulk rock compositions of olivine and its reaction products based on X-ray maps. The local bulk rock composition of the reaction products is used to determine changes in composition with respect to olivine. This change is quantified by defining an anhydrous composition change in element i ($\Delta X_i$). Water addition during hydration leads to apparent dilution of the other oxides in the reaction products. To circumvent this effect, each oxide concentration in the reaction products is multiplied by the ratio of the total oxide concentration in olivine over the total oxide concentration in the reaction products ($R_{tot}$). $\Delta X_i$ is then calculated as the mass fraction of change in composition with respect to olivine, leading to:

$$ \Delta X_i = \frac{X_i^{rp} R_{tot} - X_i^{ol}}{X_i^{ol}} \quad (4) $$

with $X_i^{rp}$ and $X_i^{ol}$ the mass concentrations in oxide i in the reaction products and in olivine, respectively.

## Transmission electron microscopy (TEM)

Focused Ion Beam (FIB) thin sections are extracted and thinned down to electron transparency (< 100 nm) with a Zeiss Crossbeam 5550 L (IEM, Lille, France) using a gallium source. TEM observations (bright field imaging and Selected Area Electron Diffraction - SAED) are carried out on a JEOL 2010 microscope (CINAM, Marseille, France) operating at 200 kV and equipped with a LaB6 crystal, providing a point-to-point resolution of 0.23 nm. The high-sensitivity low light level camera (GATAN™ Ultrascan 1000XP -2k*2k) allow very low illumination of the sample, keeping beam damage as low as possible on sensitive minerals. X-ray energy dispersive (EDX) microanalyses are conducted with a X-Flash Silicon Drift Detector 5030 (Bruker™) with an active area of 30 mm² (slew AP3.3 window) and a 133 eV resolution allowing detection of light elements with a good sensitivity. Data collection parameters are
set as follows: magnification of 50 000 x, 4L spot size, angular tilt of 20° toward the detector, time constant of 60 kcp.s⁻¹, energy range of 40 keV, and corrected counting time of 30 s. The beam diameter is set to ~ 20 nm in order to reach the smallest particles. The constant beam density is ~ 63.5 pA.cm⁻². O, Na, Mg, Al, Si, K, Ca, and Fe are quantified by applying the Bruker AXS MET line mark data quantification procedure, which is close to the original Cliff and Lorimer method [Cliff and Lorimer, 1975]. In this procedure, the acquired EDX spectra are corrected by background subtraction (Bremsstrahlung calculation), Gaussian deconvolution, and k factor corrections using values previously calculated on layer silicate standards with known homogeneous compositions [Berthonneau et al., 2014].

## Thermodynamic modelling

EQ3/6 [Wolery, 1992] is commonly used to model ultramafic rock alteration [McCollom and Bach, 2009, Klein et al., 2009, 2013, Ely et al., 2023]. The study of fluid/rock interactions requires modelling stable mineralogical assemblages for large ranges of system compositions. Change of system composition with EQ3/6 requires the execution of independent calculations, which can be done automatically by coupling softwares [Malvoisin, 2015, Ely et al., 2023]. An alternative approach, adopted here, consists of using thermodynamic calculations based on Gibbs energy minimization with Perple_X [Connolly, 2005], allowing for efficient investigation of the compositional space. Another advantage of Perple_X is the availability of a large range of solid solutions, that have to be considered for an accurate modelling of ultramafic rock alteration [Leong and Shock, 2020].

Calculations are performed with Perple_X version 7.1.12 [Connolly, 2005]. The fluid is modelled with a generic hybrid fluid equation of state solution model [WADDAH; Connolly, 1995]. The concentrations of aqueous species are calculated by using lagged speciation [Connolly and Galvez, 2018]. Two thermodynamic databases are created from the database of Holland and Powell [2011] for solid
phases and the Deep Earth Water model of Huang and Sverjensky [2019] for
aqueous species (DEW19HP6222ver_elements.dat in Perple_X). A new Fe(OH)₂
endmember is introduced in the databases. The Cₚ function of Fe(OH)₂ is
calculated, following McCollom and Bach [2009], as: Cₚ,Fe(OH)₂ = Cₚ,Mg(OH)₂ +
1/3 (Cₚ,greenalite - Cₚ,lizardite). The α₀, κ₀, κ'₀ and κ''₀ parameters in the equation
of state of Fe(OH)₂ [Holland and Powell, 2011] are set as equal to those of brucite
(Mg(OH)₂). For one database, we use the standard state enthalpy and third law
entropy provided in McCollom and Bach [2009] based on the experimental data of
Sverjensky and Molling [1992] and Wagman [1982], respectively. These data are used
to calculate equilibrium constants at 50 MPa in the 25-400°C temperature range for
reaction 1 with SUPCRT92 [Johnson et al., 1992]. We use the llnl.dat thermodynamic
database distributed along with SUPCRT92 for water, magnetite and dihydrogen.
Enthalpy and third law entropy for Fe(OH)₂ are then calculated by fitting the
equilibrium constants determined with SUPCRT92 with the thermodynamic data of
magnetite, H₂O and H₂ from the DEW19HP622ver_elements database. The obtained
thermodynamic database is referred to as MC database in the following (Table 2).
Another database is built with standard state enthalpy and third law entropy for
Fe(OH)₂ fitted with the experimental dataset of Carlin et al. [2024] for reaction
1 (CA database in the following). The procedure for fitting the parameters is the
same of the one described in Carlin et al. [2024] but with the thermodynamic data
for magnetite, H₂O and H₂ from the DEW19HP622ver_elements database. The MC
and CA databases are thus internally consistent. Ni-bearing alloy (awaruite, Ni₃Fe),
Ni ferrite (trevorite, NiFe₂O₄), Ni(OH)₂, nepouite (Ni₃Si₂O₅(OH)₄) and kerolite
(Ni₃Si₄O₁₀(OH)₂) properties are also added to the databases. The thermodynamic
properties of awaruite are from Evans et al. [2017]. For trevorite, the enthalpy of
formation and the entropy are taken from Lemire et al. [2020]. The cₚ data of
trevorite measured by Nelson et al. [2014] are fitted with a Landau model for the
<table><thead><tr><th>Endmember</th><th>Δ<sub>i</sub>H° (kJ/mol)</th><th>Δ<sub>i</sub>G° (kJ/mol)</th><th>S° (J/mol/K)</th><th>a</th><th>b (×10<sup>-5</sup>)</th><th>c</th><th>d</th></tr></thead><tbody><tr><td>Fe(OH)<sub>2,MC</sub></td><td>-573.18</td><td>-491.35</td><td>88.7</td><td>0.1456</td><td>0.3842</td><td>-1730.7</td><td>-0.6869</td></tr><tr><td>Fe(OH)<sub>2,CA</sub></td><td>-580.38</td><td>-497.82</td><td>86.2</td><td>0.1456</td><td>0.3842</td><td>-1730.7</td><td>-0.6869</td></tr></tbody></table>

Notes: MC for fitting with McCollom and Bach (2009)'s data; CA for Carlin et al. (2024)'s data. ΔᵢH° is the standard enthalpy of formation from the elements at 298.15 K and 1 bar; ΔᵢG° is the standard Gibbs energy of formation from the elements at 298.15 K and 1 bar; S° is the third law molar entropy; a, b, c and d are molar heat capacity (Cₚ in kJ/mol/K) polynomial coefficients where Cₚ = a + bT + cT² + dT¹/².

<caption>**Table 2.** Thermodynamic properties of Fe(OH)₂ endmembers used in this study.</caption>

lambda transition [Holland and Powell, 1998]. The thermal expansion coefficient is determined by fitting the data of Nelson et al. [2014]. The isothermal bulk modulus and its temperature derivative are taken from Liebermann [1972]. For Ni(OH)₂, we use the thermodynamic parameters provided in Thermoddem database [Blanc et al., 2012] based on experimental data of Palmer and Gamsjäger [2010]. Nepouite and kerolite thermodynamic properties are calculated as linear combinations of brucite, forsterite and Ni₂SiO₄ thermodynamic properties.

We use solid solution models for olivine [Holland and Powell, 1998], orthopyroxene [Holland and Powell, 1996], brucite (ideal with brucite, amakinite and Ni(OH)₂ as endmembers), talc (ideal), magnetite/trevorite (ideal) and lizardite modified from Evans et al. [2013] by adding nepouite as ideal endmember. Experimental data indicate an ideal behaviour of the magnetite/trevorite solid solution for magnetite fractions above ~ 85 mol.% [Lemire et al., 2020]. The lizardite solid solution model of Evans et al. [2013] considers ferri-Tschermak's substitution. The combined analysis of microprobe and Mössbauer data reveal that Tschermak's substitution is the main substitution involved in Fe³⁺ incorporation in serpentine [Beard and Frost, 2017].

We test the influence of five parameters on the predicted mineralogy and composition: the thermodynamic data for Fe(OH)₂, the presence of Ni in the system, temperature, mass water to rock ratio, protolith composition and initial fluid $\frac{O_2}{H_2+O_2}$ ratio (XOᵢ). We consider a minimum water to rock ratio of 0.2, approximately corresponding to the minimum amount of water required for complete olivine
hydration [Klein et al., 2009]. As in previous thermodynamic studies [McCollom and Bach, 2009, Klein et al., 2009], simulations are performed here in a single step during which equilibrium composition is calculated for fixed amounts of fluid and rock. We therefore do not take into account the fact that fluid/rock interactions can occur though continuous fluid transport into the rock. To test the influence of assuming a single step of fluid/rock interaction, we compare our simulation outputs with a model considering incremental renewing of the fluid (Supplementary Fig. S1). This latter model uses a loop of 1) addition of an increment of water amount (Δw) to a fixed amount of rock, 2) equilibrium calculation, and 3) fluid removal. The loop is stopped when the sum of increments of water amount is equal to the targeted water to rock ratio ($\Sigma_{r} \Delta w = w/r$ with w and r the mass of water and rock, respectively). The results are similar for both models (Supplementary Fig. S1), indicating that considering a single interaction step instead of several steps of fluid renewal has minor effects on the computed equilibrium composition. The effect of protolith composition is investigated by using bulk rock compositions corresponding to olivine/orthopyroxene mixtures. According to the results from this study, olivine composition is considered as $(Mg_{0.91}Fe_{0.09})_2SiO_4$ when Ni is not included in the simulation, and $(Mg_{0.9064}Fe_{0.0896}Ni_{0.004})_2SiO_4$ when Ni is included in the simulation (equivalent to a NiO concentration of 0.37 wt.%). Orthopyroxene composition is fixed to $(Mg_{0.91}Fe_{0.09})SiO_3$ following Klein et al. [2009]. The effects of changes in $X_{Oi}$ are investigated to determine the role played by $H_2$ removal from the system without change in water content (i.e. diffusion). $X_{Oi}$ is only increased compared to its value for pure water (1/3); this is equivalent to simulate the effect of $H_2$ loss only.

To determine which thermodynamic database (MC or CA) better reproduces observations in natural samples, we compare calculated and measured distribution coefficients of $Fe^{2+}$ and Mg between brucite and serpentine ($K_D$). We thus consider
divalent cation exchange and have to correct for the presence of Fe³⁺ in serpentine. K_D is defined according to Frost et al. [2013] as:

$$K_D = \frac{X_{Mg}^b (1 - X_{Mg}^{s2+})}{(1 - X_{Mg}^b) X_{Mg}^{s2+}} \quad (5)$$

where X<sub>Mg</sub><sup>s2+</sup> is the M<sub>Mg</sub>/M<sub>Mg+Fe2+</sub> ratio of serpentine which can be calculated from X<sub>Mg</sub><sup>s</sup> as:

$$X_{Mg}^{s2+} = \frac{X_{Mg}^{s}}{1 - X_{Fe3}^{s} (1 - X_{Mg}^{s})} \quad (6)$$

where X<sub>Fe3</sub><sup>s</sup> is the molar fraction of Fe³⁺ over Fe²⁺ + Fe³⁺ in serpentine. X<sub>Fe3</sub><sup>s</sup> can be calculated with the solid solution composition output with Perple_X. In this study, X<sub>Fe3</sub><sup>s</sup> is not measured in natural samples. We thus assume values for natural serpentinites based on previous measurements. Measured X<sub>Mg</sub><sup>s</sup> in natural serpentinized peridotites are converted into X<sub>Mg</sub><sup>s2+</sup> with equation 6 by considering X<sub>Fe3</sub><sup>s</sup> = 1/3 or 2/3. The 1/3 value approximately corresponds to the mean composition reported for chrysotile in Mayhew and Ellison [2020] and to the value measured by Mössbauer spectroscopy in experimental products of serpentinization in McCollom et al. [2016]. The 2/3 value approximately corresponds to the mean composition reported for lizardite in Mayhew and Ellison [2020].

# Results

## Reaction zones around olivine grains

## Olivine texture and composition

The primary minerals consist of olivine and spinel in serpentinized dunites and olivine, orthopyroxene, clinopyroxene and spinel in serpentinized harzburgites. All the studied samples contain olivine relics. Primary grains of olivine are divided into < 5 to 200 µm wide smaller fragments (referred to as olivine grains in the following) separated
by a network of veins. This produces the typical mesh texture of serpentinized
peridotites. Olivine grains can display saw-toothed boundaries on scanning electron
images, indicating the presence of etch pits at the grain surface (Fig. 1A). Mg/Mg+Fe
molar ratio (X<sub>Mg</sub><sup>ol</sup>) varies from 0.900 ± 0.004 in the olivine of the samples from the
Southwest Indian Ridge to 0.919 ± 0.003 in the samples from the Leka ophiolite (Table
3). The standard deviation of X<sub>Mg</sub><sup>ol</sup> is below 0.009 at all studied localities, indicating
limited spatial variation in olivine Mg content. Olivine also contains minor amounts
of Mn at all localities with MnO between 0.04 ± 0.07 wt.% in samples from Hess Deep
and 0.14 ± 0.12 wt.% in samples from ODP Leg 209. The Ni content is also always
above the detection limit with values ranging between 0.36 ± 0.07 wt.% in samples
from the Mirdita ophiolite and 0.41 ± 0.08 wt.% in samples from the Leka ophiolite.
Ca displays the highest absolute variation between the analyzed olivines. Measured
CaO contents are systematically below the detection limit in samples from the Mirdita
and Leka ophiolites. They are above 0.05 wt.% in the samples from ODP Leg 209,
Hess Deep and Papua New Guinea. No systematic variation in olivine composition
is observed between dunite and harzburgite. Olivines are devoid of inclusions except
for samples from CM2B hole of Oman drilling project and from Hess Deep in which
numerous needle-shaped and parallel inclusions are found. The inclusions are enriched
in Fe, Al, Ca and Cr with respect to olivine.

Reaction products at the olivine contact (Reaction Zone 1 in the mesh
center)

In the center of the meshes at the contact with olivine, a reaction zone composed of
serpentine, Fe-brucite and Ni-alloy is systematically found (Reaction Zone 1; Fig. 1).
Raman spectra of serpentine mineral in the high frequency range (3550-3850 cm⁻¹;
OH region) reveal a main peak at ~ 3685 cm⁻¹ and a minor peak at ~ 3703 cm⁻¹,
indicating that the serpentine mineral is lizardite [Auzende et al., 2004]. At most
localities (9 over the 13 localities studied here, Table 1), strong brightness contrasts
<table><thead><tr><th rowspan="2">Wt.%<br>No.</th><th colspan="2">Mirdiza ophiolites</th><th colspan="2">OOP Leg 209</th><th colspan="2">Leka ophiolites</th><th colspan="2">Hess Deep</th><th colspan="2">Papua New Guinea</th><th colspan="2">Feragen ultramafic body</th><th colspan="2">Southwest Indian Ridge</th><th colspan="2">Oman ophiolites</th></tr><tr><th>Average</th><th>Stand. Dev.</th><th>Average</th><th>Stand. Dev.</th><th>Average</th><th>Stand. Dev.</th><th>Average</th><th>Stand. Dev.</th><th>Average</th><th>Stand. Dev.</th><th>Average</th><th>Stand. Dev.</th><th>Average</th><th>Stand. Dev.</th><th>Average</th><th>Stand. Dev.</th></tr></thead><tbody><tr><td>SiO<sub>2</sub></td><td>41.08</td><td>0.34</td><td>41.33</td><td>0.40</td><td>41.08</td><td>0.29</td><td>41.10</td><td>0.33</td><td>41.40</td><td>0.32</td><td>41.35</td><td>0.31</td><td>41.19</td><td>0.36</td><td>40.74</td><td>0.52</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>bdl</td><td>bdl</td><td>0.03</td><td>0.03</td><td>bdl</td><td>bdl</td><td>0.01</td><td>0.02</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td></tr><tr><td>Fe<sup>2+</sup>(T)</td><td>8.92</td><td>0.73</td><td>8.30</td><td>0.73</td><td>7.93</td><td>0.27</td><td>8.70</td><td>0.24</td><td>7.88</td><td>0.86</td><td>7.88</td><td>0.38</td><td>9.68</td><td>0.35</td><td>8.89</td><td>0.25</td></tr><tr><td>MgO</td><td>0.07</td><td>0.09</td><td>0.14</td><td>0.12</td><td>0.06</td><td>0.09</td><td>0.04</td><td>0.07</td><td>0.10</td><td>0.09</td><td>0.08</td><td>0.09</td><td>0.12</td><td>0.10</td><td>0.08</td><td>0.09</td></tr><tr><td>Co<sub>2</sub>O<sub>3</sub></td><td>49.42</td><td>0.63</td><td>49.76</td><td>0.48</td><td>50.33</td><td>0.47</td><td>50.25</td><td>0.70</td><td>50.01</td><td>0.70</td><td>49.93</td><td>0.38</td><td>48.76</td><td>0.26</td><td>49.48</td><td>0.74</td></tr><tr><td>Na<sub>2</sub>O</td><td>bdl</td><td>bdl</td><td>0.16</td><td>0.08</td><td>bdl</td><td>bdl</td><td>0.11</td><td>0.01</td><td>0.06</td><td>0.16</td><td>0.02</td><td>0.03</td><td>0.05</td><td>0.04</td><td>0.03</td><td>0.03</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>0.01</td><td>0.03</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td></tr><tr><td>Cr<sub>2</sub>O<sub>3</sub></td><td>bdl</td><td>bdl</td><td>0.01</td><td>0.03</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td></tr><tr><td>Ni</td><td>0.36</td><td>0.07</td><td>0.38</td><td>0.08</td><td>0.41</td><td>0.08</td><td>0.39</td><td>0.04</td><td>0.33</td><td>0.16</td><td>0.37</td><td>0.08</td><td>0.39</td><td>0.09</td><td>0.37</td><td>0.06</td></tr><tr><td>TiO<sub>2</sub></td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td></tr><tr><td>Total</td><td>100.09</td><td>0.14</td><td>100.39</td><td>0.74</td><td>100.12</td><td>0.25</td><td>100.68</td><td>0.55</td><td>100.10</td><td>0.26</td><td>100.13</td><td>0.23</td><td>100.32</td><td>0.45</td><td>100.12</td><td>0.23</td></tr><tr><td colspan="19">Basis</td></tr><tr><td>Si</td><td>1.004</td><td>0.005</td><td>1.006</td><td>0.004</td><td>1.001</td><td>0.006</td><td>0.997</td><td>0.011</td><td>1.008</td><td>0.006</td><td>1.007</td><td>0.005</td><td>1.007</td><td>0.005</td><td>0.999</td><td>0.013</td></tr><tr><td>Al</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.001</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td></tr><tr><td>Fe</td><td>0.182</td><td>0.016</td><td>0.169</td><td>0.003</td><td>0.362</td><td>0.005</td><td>0.177</td><td>0.004</td><td>0.160</td><td>0.008</td><td>0.182</td><td>0.008</td><td>0.138</td><td>0.008</td><td>0.182</td><td>0.005</td></tr><tr><td>Mn</td><td>0.001</td><td>0.002</td><td>0.003</td><td>0.002</td><td>0.002</td><td>0.001</td><td>0.001</td><td>0.002</td><td>0.002</td><td>0.002</td><td>0.002</td><td>0.002</td><td>0.002</td><td>0.002</td><td>0.002</td><td>0.002</td></tr><tr><td>Mg</td><td>1.801</td><td>0.017</td><td>1.805</td><td>0.007</td><td>1.828</td><td>0.014</td><td>1.817</td><td>0.019</td><td>1.814</td><td>0.019</td><td>1.813</td><td>0.011</td><td>1.777</td><td>0.013</td><td>1.809</td><td>0.024</td></tr><tr><td>Co</td><td>0.000</td><td>0.000</td><td>0.004</td><td>0.002</td><td>0.000</td><td>0.000</td><td>0.003</td><td>0.006</td><td>0.004</td><td>0.006</td><td>0.000</td><td>0.001</td><td>0.001</td><td>0.001</td><td>0.000</td><td>0.001</td></tr><tr><td>Na</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.001</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td></tr><tr><td>Cr</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.001</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td></tr><tr><td>Ni</td><td>0.000</td><td>0.000</td><td>0.007</td><td>0.002</td><td>0.000</td><td>0.002</td><td>0.008</td><td>0.001</td><td>0.006</td><td>0.003</td><td>0.002</td><td>0.002</td><td>0.002</td><td>0.002</td><td>0.007</td><td>0.001</td></tr><tr><td>Ti</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.000</td></tr><tr><td>X<sub>Mg</sub></td><td>0.908</td><td>0.008</td><td>0.914</td><td>0.001</td><td>0.919</td><td>0.003</td><td>0.911</td><td>0.001</td><td>0.919</td><td>0.009</td><td>0.917</td><td>0.004</td><td>0.900</td><td>0.004</td><td>0.908</td><td>0.002</td></tr><tr><td>X<sub>Fe</sub></td><td>0.004</td><td>0.001</td><td>0.004</td><td>0.001</td><td>0.004</td><td>0.001</td><td>0.004</td><td>0.000</td><td>0.002</td><td>0.002</td><td>0.004</td><td>0.001</td><td>0.004</td><td>0.001</td><td>0.004</td><td>0.001</td></tr></tbody></table>

Notes: No.: number of analyses; FeO(T): total iron; bdl: below detection limit; X<sub>Mg</sub> = Mg/(Mg + Fe); X<sub>Fe</sub> = Ni/(Ni + Fe + Mg).

<caption>**Table 3.** Averages and standard deviations of electron microprobe analyses of olivine composition (wt.%).</caption>

are observed on back-scattered electron images (Fig. 1), suggesting compositional changes at the micrometer scale, confirmed by X-ray mapping (Figs. 3, 4, and Supplementary Fig. S2). Electron probe microanalyses performed in the whole mesh texture reveal a composition of the reaction products lying between pure serpentine and pure brucite (Fig. 6A). This indicates that serpentine is intermixed at the submicrometer scale with nanometer-sized brucite. This is confirmed by Raman microanalysis showing the presence of peaks associated with brucite at ~ 277, 442 and 3645 cm⁻¹.

Brucite abundance estimated with X<sub>Br</sub> strongly varies at the micrometer scale (Figs. 3, 4, and Supplementary Fig. S2). Single olivine grains are generally rimmed by both serpentine-dominated and Fe-brucite-dominated mixtures (Fig. 4). The complete replacement of olivine grains leads to the concentric alternation at the tens of micrometer scale of serpentine-dominated and Fe-brucite-dominated mixtures forming an onion-like texture (Fig. 4C). The mean value of X<sub>Br</sub> in Reaction Zone 1 is of 0.19, 0.39, 0.39 and 0.48 in samples from the Feragen ultramafic body, ODP Leg 209, Oman Drilling Project (CM2B hole) and Papua New Guinea, respectively.

For the samples analyzed here, Al₂O₃ contents in serpentine/brucite mixtures are generally below the detection limit except for samples from hole CM2B of the Oman
drilling project in which Al₂O₃ can reach 1.75 wt.%. Al₂O₃ contents are also low in the compilation of analyses from the literature except for DSDP Leg 82, ODP Leg 153 and ODP Leg 173 with maximum concentrations of 2.16, 1.17 and 0.81 wt.%, respectively. Electron microprobe analyses of the mixtures are aligned in Fe vs. Si + Al/2 diagrams (Fig. 5 and Supplementary Figs. S3 and S4). A single alignment is observed at most localities. In the Mirdita ophiolite, two trends are observed for two distinct groups of samples (Fig. 5A). Fe content decreases from Si + Al/2 = 0 to Si + Al/2 = 2, indicating that brucite is enriched in Fe compared to lizardite. Consistently, the measured X<sub>Mg</sub><sup>b</sup> are lower than X<sub>Mg</sub><sup>s</sup>. X<sub>Mg</sub><sup>b</sup> varies between 66 and 82 mol.% with a mean value of 76 mol.% whereas X<sub>Mg</sub><sup>s</sup> varies between 92 and 98 mol.% with a mean value of 95 mol.% (Table 4). Serpentine/brucite mixtures show noticeable dispersion around the linear fits in Fe vs. Si + Al/2 diagrams (Fig. 5 and Supplementary Figs. S3 and S4), indicating spatial variations of X<sub>Mg</sub><sup>b</sup> and X<sub>Mg</sub><sup>s</sup> at least at the thin section scale. This is confirmed with X-ray mapping in which variations in X<sub>Mg</sub><sup>b</sup> can reach up to 15 mol.% (Fig. 4E and F and Supplementary Fig. S2C). The variations in X<sub>Mg</sub><sup>b</sup> occur between ~ 20 µm wide clusters of homogeneous composition. These clusters appear as randomly oriented with respect to the olivine contacts. A notable exception consists in sample 1274A-8R1-16 (ODP Leg 209) for which X<sub>Mg</sub><sup>b</sup> increases from 60 to 75 mol.% with the distance from olivine contact (Fig. 4 E).

Mn, Ca and Ni, the three other elements detected in olivine composition analysis,
display different distributions in the reaction products of serpentinization. Mn and
Ca distributions mimic Fe distribution on X-ray compositional maps (Figs. 3I, J, 4I
to L and Supplementary Fig. S2E). Point analyses confirm this relationship for Mn
with a positive correlation between Mn and Fe in the serpentine/brucite mixture,
with a Mn:Fe ratio of 0.0246 ± 0.0006 (Fig. 6B).

Even though Ni molar concentrations in olivine are at least two times higher
than Mn concentrations (Table 3), the measured Ni contents in serpentine + brucite
<table><thead><tr><th rowspan="2">Locality</th><th colspan="3">Reaction Zone 1</th><th colspan="3">Reaction Zone 2</th><th rowspan="2">Reference</th></tr><tr><th>N°</th><th>X<sub>Mg</sub><sup>b</sup></th><th>X<sub>Mg</sub><sup>c</sup></th><th>N°</th><th>X<sub>Mg</sub><sup>b</sup></th><th>X<sub>Mg</sub><sup>c</sup></th></tr></thead><tbody><tr><td>Papua New Guinea</td><td>237</td><td>0.794 ± 0.005</td><td>0.942 ± 0.007</td><td>82</td><td>0.976 ± 0.003</td><td>0.937 ± 0.010</td><td>This study</td></tr><tr><td>Feragen ultramafic body</td><td>22</td><td>n.f.</td><td>n.f.</td><td>57</td><td>0.915 ± 0.004</td><td>0.962 ± 0.004</td><td>This study</td></tr><tr><td>Leka ophiolite complex</td><td>79</td><td>0.674 ± 0.015</td><td>0.957 ± 0.008</td><td></td><td></td><td></td><td>This study</td></tr><tr><td>CM2B (Oman)</td><td>116</td><td>0.664 ± 0.019</td><td>0.969 ± 0.019</td><td>50</td><td>0.829 ± 0.011</td><td>0.961 ± 0.016</td><td>This study</td></tr><tr><td>Mirdita ophiolite (96AL-8)</td><td>30</td><td>0.780 ± 0.020</td><td>0.953 ± 0.047</td><td></td><td></td><td></td><td>This study</td></tr><tr><td>Mirdita ophiolite (other)</td><td>99</td><td>0.659 ± 0.012</td><td>0.920 ± 0.017</td><td></td><td></td><td></td><td>This study</td></tr><tr><td>Southwest Indian Ridge</td><td>83</td><td>0.698 ± 0.010</td><td>0.924 ± 0.012</td><td></td><td></td><td></td><td>This study</td></tr><tr><td>ODP Leg 209</td><td>96</td><td>0.765 ± 0.013</td><td>0.951 ± 0.007</td><td>56</td><td>0.947 ± 0.004</td><td>0.964 ± 0.023</td><td>This study</td></tr><tr><td>Hess Deep</td><td>33</td><td>0.740 ± 0.018</td><td>0.938 ± 0.023</td><td></td><td></td><td></td><td>This study</td></tr><tr><td>DSDP Leg 82</td><td>56</td><td>0.786 ± 0.004</td><td>0.980 ± 0.004</td><td></td><td></td><td></td><td>Klein et al. (2014)</td></tr><tr><td>ODP Leg 107</td><td>32</td><td>0.760 ± 0.009</td><td>0.944 ± 0.016</td><td></td><td></td><td></td><td>Klein et al. (2014)</td></tr><tr><td>ODP Leg 125</td><td>345</td><td>0.821 ± 0.004</td><td>0.933 ± 0.002</td><td></td><td></td><td></td><td>Klein et al. (2014)</td></tr><tr><td>ODP Leg 147</td><td>177</td><td>0.773 ± 0.006</td><td>0.940 ± 0.008</td><td></td><td></td><td></td><td>Klein et al. (2014)</td></tr><tr><td>ODP Leg 153</td><td>127</td><td>n.f.</td><td>n.f.</td><td>88</td><td>0.941 ± 0.005</td><td>0.978 ± 0.004</td><td>Klein et al. (2014)</td></tr><tr><td>ODP Leg 173</td><td>79</td><td>0.811 ± 0.007</td><td>0.934 ± 0.011</td><td></td><td></td><td></td><td>Klein et al. (2014)</td></tr><tr><td>Santa Elena Ophiolite</td><td>651</td><td>0.789 ± 0.006</td><td>0.950 ± 0.002</td><td>65</td><td>0.936 ± 0.005</td><td>0.964 ± 0.006</td><td>Schwarzenbach et al. (2016)</td></tr><tr><td>New Caledonia</td><td>238</td><td>0.815 ± 0.007</td><td>0.956 ± 0.004</td><td></td><td></td><td></td><td>Frost et al. (2013)</td></tr><tr><td>Zedang ophiolite</td><td>30</td><td>0.822 ± 0.015</td><td>0.959 ± 0.020</td><td></td><td></td><td></td><td>Li et al. (2020)</td></tr><tr><td>Mineoka ophiolite</td><td>96</td><td>0.699 ± 0.021</td><td>0.960 ± 0.005</td><td></td><td></td><td></td><td>Katayama et al. (2010)</td></tr></tbody></table>

N°: number of analyses; n.f.: dataset not fitted due to the absence of analyses of serpentine/brucite mixture with Si (apfu) < 1.7.

<caption>**Table 4.** Molar $\frac{Mg}{Mg+Fe}$ ratio in brucite ($X_{Mg}^b$) and serpentine ($X_{Mg}^s$) calculated with the data obtained here and datasets from the literature.</caption>

mixtures are lower than the measured Mn contents. Moreover, Ni is not correlated with Fe in the serpentine/brucite mixtures (Fig. 6C). Ni rather concentrates in submicrometer wide grains homogeneously distributed in Reaction Zones 1 (Fig. 3G, H, 4G and H and Supplementary Fig. S2D). Ni-bearing grains are observed as numerous bright spots on back-scattered electron images, leading to a starry sky appearance (Fig. 1A, E and F). Sulfur is not detected with X-ray mapping in Reaction Zone 1. In the Ni-rich grains, X-ray maps show a clear positive correlation between Ni and Fe. Ni maps are poorly calibrated at high Ni concentration due to the difficulty in analyzing submicrometer grains, providing only semi-quantitative Ni concentration estimates. The estimated Ni:Fe ratio is comprised between 2 and 3. This suggests that the Ni-bearing phase is awaruite. No other phases (and in particular magnetite) are observed in Reaction Zone 1.

Anhydrous mass changes in Reaction Zone 1 are both positive and negative and below 11 wt.% for SiO₂ and MgO in all acquired X-ray maps (Fig. 7). ΔX<sub>Fe</sub> is only small for Reaction Zone 1 in samples from CM2B hole of the Oman drilling project. At all the other localities, both positive (ODP Leg 209 and Papuan ultramafic belt)
and negative (Feragen ultramafic body) changes in FeO content of several tens of percent with respect to olivine are observed.

## Reaction products at the mesh rim (Reaction Zone 2)

At four localities studied here (Papua New Guinea, Feragen ultramafic body, CM2B hole in Oman and ODP Leg 209), Reaction Zone 1 can display a sharp contact with another reaction zone (Reaction Zone 2). This latter assemblage occurs in veins forming the mesh rim in a distance ranging from less than 20 µm to more than 200 µm from the olivine contact (Fig. 1A, B, E and F). The veins are ~ 10 to ~ 100 µm wide. The vein walls are generally enriched in serpentine with respect to brucite. The vein centers are made of brucite, magnetite and minor amount of serpentine (Figs. 3C, D, 4C, D and Supplementary Fig. S2B). Ni-rich spots are observed in Reaction Zone 2 at only one locality (CM2B hole). Raman spectra indicate that serpentine in Reaction Zone 2 is a mixture between chrysotile and lizardite. $X_{Mg}^s$ in Reaction Zone 2 are comprised between 94 and 98 mol.% and systematically equal or higher than $X_{Mg}^s$ measured in Reaction Zone 1 (Table 4). Brucite/serpentine mixing is less pronounced than in Reaction Zone 1, and single brucite grains can be found (Figs. 1F, G and H and 4E and F).

Two microstructures are associated with the presence of brucite in Reaction Zone 2. In Mg-rich areas, a first type of mosaic microtexture is observed (Fig. 1F, G and H, details in Fig. 2A, B, C). It consists of large brucite grains up to 20 µm in size, which appear dark on back-scattered electron images. They are intersected by thin, 100 nm wide vein-like domains, which appear brighter on back-scattered electron images. TEM bright-field imaging show that these vein-like domains consist of a mix of chrysotile and lizardite (planar spacing of 7.4 Å). The brucite chemistry characterized using TEM-coupled EDX-spectroscopy show a Mg-rich composition with $X_{Mg}^b$ >0.95 (Fig. 5E). At the Mg-brucite/serpentine contact, Fe-brucite is in places identified due to its basal spacing of 4.6 Å and its chemical composition of $X_{Mg}^b$ ~ 0.8 (Fig.
5E). Fe-brucite is more stable under the TEM-beam than the large Mg-brucite crystals which quickly amorphize during analysis. Mg-brucite is thus characterized using TEM mainly through its chemical composition measured using coupled EDX spectroscopy. In areas with higher Si-content, a second type of microtexture is observed. Elongated brucite crystals and serpentine form a complex assemblage which resembles a symplectitic microtexture. Brucite crystals have a typical aspect ratio (length over width) of 10 and a length ranging from the submicrometer scale to ~ 10 µm (Fig. 1H, Fig. 2D, E, F). They are often found without a preferred orientation, but may appear parallel in places. TEM-analysis shows that these brucite crystals are Mg-rich ($X_{Mg}^b$ ~ 0.95). The only serpentine mineral found in this assemblage is lizardite. Furthermore, Fe-brucite is identified. It often occurs as irregularly shaped grains together with Mg-brucite grains and lizardite, and has a small grain size of 10 - 100 nm. In places, Fe-brucite forms patches along the corners and edges of Mg-brucite; very small grains occur within lizardite (Supplementary Fig. S5). Between the Fe-brucite planes, undulating pore spaces are identified due to their high transmittance of the TEM electron beam.

$X_{Mg}^b$ values in Reaction Zone 2 vary between 92 and 98 mol.% except in samples from the CM2B hole of Oman drilling project in which a significantly lower value of 83 mol.% is measured (Fig. 5, Supplementary Figs. S3 and S4 and Table 4). $X_{Mg}^b$ values are systematically between 15 and 19 mol.% higher than in Reaction Zone 1. As in Reaction Zone 1, point analyses indicate a positive correlation between Mn and Fe in brucite/serpentine mixtures in Reaction Zone 2. However, the Mn:Fe ratio is 1.8 times higher than in Reaction Zone 1 at 0.045 ± 0.002 (Fig. 6B). Magnetite occurs in the center of veins in association with brucite, generally as aggregates of several micrometer wide subgrains. The aggregates are elongated parallel to the vein with a maximum length reaching several hundreds of micrometers and a maximum width of 20 µm (Figs. 1B, E and G, 3A and B, 4A and Supplementary Fig. S2A).
In samples from Papua New Guinea and the Leka ophiolite, isolated idiomorphic magnetite grains with a hexagonal section also occur. They can reach a size of 200 µm and can be fractured. Point analyses of magnetite reveal, that, in addition to FeO, magnetite can also contain more than 0.05 wt.% of MgO, NiO, SiO₂ and MnO (Table 5). MgO, NiO and SiO₂ contents can be above 1 wt.%. Two types of magnetite compositions can be distinguished based on the maximum measured MnO and NiO contents. Maximum measured MnO and NiO contents are both above 0.08 wt.% in samples from the Feragen ultramafic body, ODP Leg 209 and the Papuan ultramafic belt, whereas they are both lower than this value in samples from CM2B hole in Oman. The maximum measured MnO contents are of 0.42 and 0.29 wt.% in samples from ODP Leg 209 and Papua New Guinea. The maximum measured NiO contents are of 1.12 and 1.75 wt.% in samples from the Feragen ultramafic body and Papua New Guinea, respectively. X-ray maps also reveal a difference in the spatial distribution of Ni between the samples. Ni is only concentrated in magnetite in samples from the Feragen ultramafic body, ODP Leg 209 and the Papuan ultramafic belt (Figs. 3H and 4G and H). In samples from CM2B hole of the Oman drilling project, Ni is not concentrated in magnetite but in micrometer-wide grains as in Reaction Zone 1. The semi-quantitative relationship between Ni and Fe in these grains suggests a Ni:Fe ratio comprised between 2 and 3, compatible with awaruite.

The local bulk rock composition of Reaction Zones 2 indicate significant differences in composition with respect to olivine. Anhydrous mass changes in SiO₂ and MgO contents are limited to negative values ranging from -13 to -5 wt.% for samples from ODP Leg 209 and CM2B hole (Fig. 7). FeO content in these samples is approximately doubled in Reaction Zone 2 with respect to olivine. In samples from the Feragen and Papuan ultramafic bodies, all ΔXᵢ are above 30 % in absolute values. FeO content quadruples compared to olivine in the analyzed map of the Feragen ultramafic body. In the analyzed maps of the Papuan ultramafic belt, MgO content increases by more
<table><thead><tr><th rowspan="2">Wt.%</th><th colspan="2">Papua New Guinea</th><th>Oman</th><th>ODP Leg 209</th><th>Feragen</th></tr><tr><th>A5</th><th>A5</th><th>807-58-01-2-6</th><th>1274A-8R1-16bis</th><th>ROR0307</th></tr></thead><tbody><tr><td>SiO<sub>2</sub></td><td>0.03</td><td>0.03</td><td>0.12</td><td>0.13</td><td>0.14</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>bdl</td><td>bdl</td><td>bdl</td><td>0.02</td><td>bdl</td></tr><tr><td>FeO(T)</td><td>89.93</td><td>89.22</td><td>90.90</td><td>89.42</td><td>89.56</td></tr><tr><td>MnO</td><td>bdl</td><td>0.19</td><td>bdl</td><td>0.42</td><td>0.07</td></tr><tr><td>MgO</td><td>0.86</td><td>1.12</td><td>0.15</td><td>0.57</td><td>0.22</td></tr><tr><td>Cr<sub>2</sub>O<sub>3</sub></td><td>bdl</td><td>bdl</td><td>bdl</td><td>1.24</td><td>bdl</td></tr><tr><td>NiO</td><td>1.48</td><td>0.85</td><td>bdl</td><td>0.24</td><td>1.13</td></tr><tr><td>TiO<sub>2</sub></td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td></tr><tr><td>V<sub>2</sub>O<sub>3</sub></td><td>n.a.</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td></tr><tr><td>ZnO</td><td>n.a.</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td></tr><tr><td>CoO</td><td>n.a.</td><td>bdl</td><td>bdl</td><td>bdl</td><td>bdl</td></tr><tr><td>FeO<sub>calc</sub></td><td>28.01</td><td>27.75</td><td>30.17</td><td>29.14</td><td>28.87</td></tr><tr><td>Fe<sub>2</sub>O<sub>3</sub>calc</td><td>68.81</td><td>68.32</td><td>67.50</td><td>66.99</td><td>67.45</td></tr><tr><td>Total<sub>calc</sub></td><td>99.20</td><td>98.25</td><td>97.94</td><td>98.76</td><td>97.88</td></tr><tr><td>Basis</td><td colspan="5">3 cations and 4 oxygens</td></tr><tr><td>Si</td><td>0.001</td><td>0.001</td><td>0.005</td><td>0.005</td><td>0.005</td></tr><tr><td>Al</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.001</td><td>0.000</td></tr><tr><td>Fe<sup>2+</sup></td><td>0.904</td><td>0.902</td><td>0.991</td><td>0.946</td><td>0.949</td></tr><tr><td>Fe<sup>3+</sup></td><td>1.999</td><td>1.999</td><td>1.995</td><td>1.956</td><td>1.995</td></tr><tr><td>Mn</td><td>0.000</td><td>0.006</td><td>0.000</td><td>0.014</td><td>0.002</td></tr><tr><td>Mg</td><td>0.050</td><td>0.065</td><td>0.009</td><td>0.033</td><td>0.013</td></tr><tr><td>Cr</td><td>0.000</td><td>0.000</td><td>0.000</td><td>0.038</td><td>0.000</td></tr><tr><td>Ni</td><td>0.046</td><td>0.026</td><td>0.000</td><td>0.007</td><td>0.036</td></tr><tr><td>XNi</td><td>0.048</td><td>0.029</td><td>0.000</td><td>0.008</td><td>0.036</td></tr></tbody></table>

Notes: FeO(T): total iron; FeO<sub>calc</sub>, Fe<sub>2</sub>O<sub>3</sub>calc and Total<sub>calc</sub>: calculated based on stoichiometry, bdl: below detection limit; n.a.: not analyzed; XNi = Ni/(Ni + Fe<sup>2+</sup>).

<caption>**Table 5.** Representative electron microprobe analyses of magnetite in Reaction Zone 2. Low totals reveal high Fe<sup>3+</sup>/Fe<sup>2+</sup> ratio probably due to partial magnetite oxidation during sample preparation.</caption>

than 50 % whereas FeO content decreases by 34 %. Large magnetite grains are also found in the center of veins outside the acquired X-ray maps in these samples, suggesting that the important variations in composition measured here in Reaction Zone 2 reflect mass transfer and segregation at a scale larger than the acquired maps. The largest acquired map has a width of 750 µm (Supplementary Tab. S2), indicating that mass transfer could occur above the millimeter scale. CaO contents
measured with X-ray maps are significantly lower in Reaction Zone 2 than in olivine and Reaction Zone 1 (Figure 4K and L).

In the samples from the Southwest Indian Ridge, reaction products observed at the mesh rim slightly differ from Reaction Zone 2 at other localities. They consist of magnetite and serpentine with a microtexture similar to that observed in Reaction Zone 2. However, the mesh rim lacks Mg-brucite and contains numerous pores reaching a size of several micrometers (Supplementary Fig. S6).

## Thermodynamic modelling

### Fe partitioning between serpentine and brucite

The distribution coefficients of Fe²⁺ and Mg between serpentine and brucite are computed for the two databases studied here (MC based on the thermodynamic data of McCollom and Bach [2009] and CA based on the thermodynamic data of Carlin et al. [2024]; Fig. 8). Between 25 and 350 °C, K_D increases from 0.33 to 0.89 and from 0.02 to 0.30 with the MC and CA databases, respectively. There is thus no overlap between the K_D values computed with the two databases, allowing for easily determining which thermodynamic database is consistent with composition measurements in experiments and natural samples.

K_D values calculated with composition of serpentine and brucite from laboratory experiments [McCollom et al., 2016, Malvoisin et al., 2012, Okamoto et al., 2011, Ogasawara et al., 2013, Lafay et al., 2012] vary from 0.09 to 0.5 and from 0.04 to 0.25 by assuming X<sub>Fe3</sub><sup>s</sup> = 1/3 and 2/3, respectively (Fig. 8B). They are thus mainly consistent with the K_D values computed with the CA database.

K_D values calculated with the mineral composition measured here in natural samples vary from 0.04 to 0.36 and from 0.02 to 0.18 by assuming X<sub>Fe3</sub><sup>s</sup> = 1/3 and 2/3, respectively (Fig. 8A). Over the 22 values of K_D calculated in natural samples, only one calculated assuming X<sub>Fe3</sub><sup>s</sup> = 1/3 is consistent with calculations with the MC database and would correspond to a serpentinization temperature lower than 50
°C (Fig. 8B). All the other K_D values are consistent with calculations with the CA database. Using the CA database, the temperatures corresponding to the average K_D values ± their standard deviation calculated in natural samples range from 100 to 270 °C and from 50 to 160 °C for X<sub>Fe3</sub><sup>s</sup> = 1/3 and 2/3, respectively. The temperature estimates based on oxygen isotope measurements in serpentine fall in the same range (Fig. 8C). In ophiolites and at passive margins, the estimated temperatures ± their standard deviation vary between 100 and 170 °C. At mid-ocean ridges, they vary between 140 and 250 °C.

Based on K<sub>D</sub> comparison between thermodynamic predictions and natural samples above, thermodynamic calculations are performed with the CA database in the following.

Effect of the presence of Ni on olivine serpentinization

Olivine serpentinization is modelled as a function of temperature with the CA database for Ni-bearing and Ni-free systems (Fig. 9).

Qualitatively, the effects of temperature are the same for both models: with
increasing temperature, a decrease of brucite mode is accompanied by increases in
magnetite mode (Fig. 9A and B), and in X<sub>Mg</sub><sup>b</sup> (Fig. 9C and D). As Fe is divalent in
brucite and both trivalent and divalent in magnetite, these trends lead to an increase
in dissolved H<sub>2,aq</sub> concentration in the fluid with temperature (Fig. 9E and F). At
temperatures above 350 °C, olivine starts to form and serpentinization progressively
stops.

From 25 to 350 °C for a water to rock ratio of 1 without considering Ni, X<sub>Mg</sub><sup>b</sup> increases from 0.67 to 0.87. Magnetite formation occurs at temperatures above 310 °C (Fig. 9A). The addition of Ni in the system has three main consequences. First, awaruite is predicted to form at temperatures below 370 °C. The stability of awaruite at high temperature is limited by the formation of olivine incorporating most of the Ni. Then, magnetite stability field grows to temperatures as low as 295 °C (Fig.
9B). Finally, the predicted amount of produced hydrogen is lower than in a Ni-free system, especially at low temperature (Fig. 9E and F). At temperatures below 150 °C, predicted H₂ concentrations are approximately 4 times lower in a Ni-bearing system than in a Ni-free system.

Based on the observation of Ni-bearing phases in natural samples, simulations are performed in the following by considering Ni in the system.

## Effect of water to rock ratio on olivine serpentinization

Simulations as a function of water to rock ratio and temperature indicate that, in addition to serpentine, brucite and awaruite are the stable phases at temperatures below 360 °C and low water to rock ratio (Fig. 10). Brucite and awaruite are replaced by magnetite at higher water to rock ratio. The water to rock ratio of the transition from brucite and awaruite to magnetite decreases with temperature from 20 at 200 °C to 2 at 350 °C. Magnetite contains less than 1 mol.% of the Ni-endmember (trevorite; Fig. 10E) in all the investigated parameter range. $X_{Mg}^b$ and $H_{2,aq}$ concentration in the fluid mainly depend on temperature (Fig. 10D and F). At temperatures below 130 °C, $X_{Mg}^b$ does not depend on water to rock ratio (Fig. 10D). At temperatures between 130 and 290 °C, $X_{Mg}^b$ decreases with increasing water to rock ratio with a difference between low and high water to rock ratio of up to 0.09 at 220 °C. $X_{Mg}^b$ increases with water to rock ratio at higher temperature. Mg solubility in the performed simulations mainly depends on temperature. The dominant aqueous species is $MgO_{aq}$ with concentrations increasing with temperature from $5.10^{-4}$ mol/kg at 50 °C to more than 0.4 mol/kg at temperatures above 250 °C. At temperatures ranging from 50 to 300 °C, the dominant aqueous Si and Fe species are $SiO_{2,aq}$ and $Fe(HSiO_3)^+$, respectively, with concentrations 4 to 5 orders of magnitude lower than $MgO_{aq}$. At temperatures below 350 °C in the brucite- and awaruite-bearing domain, $H_{2,aq}$ concentration is mainly controlled by temperature. Variations of $H_{2,aq}$ concentration as a function of water to rock ratio at a fixed
temperature do not exceed 0.5 log units. Fe³⁺ over Feₜot molar ratio in serpentine varies in a similar way to H₂,aq concentration. At a fixed temperature below 300 °C, it varies with water to rock ratio by less than 10 % at values below 0.35 (Supplementary Fig. S7A).

## Effect of orthopyroxene fraction in the protolith

Simulation at 250 °C and a water to rock ratio of 1 leads to the progressive decrease in brucite amount as the orthopyroxene fraction increases (Fopx; Fig. 11A). At Fopx = 0.5, Fe-brucite has completely disappeared and serpentine is the only mineral phase. Serpentine coexists with talc at Fopx > 0.5. X<sub>Mg</sub><sup>s</sup> and X<sub>Mg</sub><sup>b</sup> slightly decrease from 0.94 to 0.9 and from 0.8 to 0.7, respectively, as Fopx increases from 0 to 0.5 (Fig. 11B). Magnetite is not predicted to form in the simulation and H<sub>2</sub> production is only related to Fe³⁺ incorporation into serpentine. H<sub>2</sub> concentration almost doubles from ~ 0.03 to ~ 0.06 mol/kg as Fopx increases from 0 to 0.5. It strongly decreases by one order of magnitude at Fopx > 0.5 where it reaches values of approximately 0.005 mol/kg (Fig. 11C).

## Effect of X<sub>Oi</sub>

The influence of bulk H₂ loss is investigated by determining mineralogy and composition as a function of X<sub>Oi</sub> (Fig. 12). Awaruite is formed at temperatures below 360 °C at X<sub>Oi</sub> below approximately 0.334. Magnetite is present at all temperatures for 0.3336 < X<sub>Oi</sub> < 0.3357. Before disappearing at high X<sub>Oi</sub>, it coexists with hematite and goethite at temperatures below 330 °C (Fig. 12B). X<sub>Mg</sub><sup>b</sup> strongly increases with X<sub>Oi</sub>. The dependency of X<sub>Mg</sub><sup>b</sup> on X<sub>Oi</sub> almost vanishes at X<sub>Oi</sub> > 0.335 where X<sub>Mg</sub><sup>b</sup> values are of approximately 0.95. Fe³⁺ over Feₜot molar ratio in serpentine varies in a similar way to X<sub>Mg</sub><sup>b</sup>. It strongly increases from values below 0.4 at X<sub>Oi</sub> = 1/3 to 1 at X<sub>Oi</sub> > 0.335 (Supplementary Fig. S7B). Ni-endmember molar fraction in the magnetite solid solution is below approximately 0.03 at X<sub>Oi</sub> < 0.335.
It increases above 0.08 at temperature above 250 °C and $X_{Oi} > 0.3355$. $H_{2,aq}$ concentration always shows an increase with temperature in the investigated $X_{Oi}$ range at temperatures below 350 °C. It decreases by approximately 2 orders of magnitude as $X_{Oi}$ increases from 1/3 to 0.335.

## Discussion

We discuss serpentinization processes in the observed samples, and provide quantitative constraints on the evolution of phase mode and composition. Chemical reactions illustrate each process and the associated mass transfer. One should keep in mind that they only represent some of the equilibria occurring in the rock without considering all possible mass transfers and mineral phases.

### Fe-rich brucite formation at the olivine contact: insights from a new thermodynamic model

Microprobe analyses of olivine show limited variation in major element compositions at the investigated localities, and in particular in the $\frac{Mg}{Mg+Fe}$ (mean value of 0.91 ± 0.04) and $\frac{Ni}{Ni+Fe+Mg}$ (mean value of 0.04 ± 0.001) molar ratios. At all localities, olivine grains are homogeneously fractured, leading to the typical mesh texture during alteration. Fracturing is probably associated with thermal cracking during mantle rock exhumation [Jamtveit et al., 2009, Rouméjon and Cannat, 2014]. The main difference between olivine grains consists in the presence of oriented parallel inclusions in the samples from the Oman ophiolite and ODP Leg 209. The shape and composition of the inclusions suggest that they are exsolutions of clinopyroxene and spinel which are commonly described in olivine [Moseley, 1984]. Such inclusions have already been observed in samples from BA4A hole (Oman drilling project) and appear to be inert during serpentinization [Malvoisin et al., 2020].

Observations in natural samples reveal the presence of Reaction Zone 1 composed of Fe-rich brucite, lizardite, awaruite at the contact with olivine at 9 of the 13 studied
localities (Table 1). The Fe-brucite + lizardite assemblage has already been reported in numerous localities at mid-ocean ridges, passive margins [Klein et al., 2014] and in ophiolites such as New Caledonia [Frost et al., 2013, Mothersole et al., 2017] or Oman [Malvoisin et al., 2021a]. The samples from three of the four localities where the assemblage Fe-brucite + lizardite is not found were collected by dredging or with submersible on the seafloor (Table 1). This seems to confirm the observation of Klein et al. [2020] that dredged serpentinites are devoid of brucite whereas most of drilled serpentinites contain brucite. Nevertheless, one should note that Fe-brucite is also found here in samples collected by dredging at the Southwest Indian Ridge and with submersible at the Hess Deep triple junction in the Pacific. The presence of brucite in these samples indicates that seafloor weathering does not necessary induce complete brucite dissolution. The analysis of Fe-brucite/lizardite mixture composition in Reaction Zone 1 reveals $\frac{Mg}{Mg+Fe^{2+}}$ molar ratios comprised between 0.66 and 0.82 for brucite and above ~ 0.95 for serpentine (Fig. 8D and E). The minimum measured $X_{Mg}^b$ of ~ 0.66 in samples from Oman, the Mirdita ophiolite and the Leka ophiolite complex is close to the minimum expected value of 0.64 if all the Fe initially contained in olivine is incorporated into brucite according to the reaction:

$$2(\text{Mg}_{0.91}, \text{Fe}_{0.09})_2\text{SiO}_4 + 3\text{H}_2\text{O} = \text{Mg}_3\text{Si}_2\text{O}_5(\text{OH})_4 + (\text{Mg}_{0.64}\text{Fe}_{0.36})(\text{OH})_2 \quad (7)$$

Equation 7 shows that a variation in $X_{Mg}^{ol}$ leads to a variation in $X_{Mg}^{b}$ four times higher. As a result, variations in olivine composition in the protolith may lead to significant changes in brucite composition. Interestingly, the samples from the Mirdita ophiolite display one of the two highest standard deviation for $X_{Mg}^{ol}$ (0.08; Table 3), which is sufficient to explain the two brucite compositions measured in Reaction Zone 1 at this locality (Table 4).
More generally, measured $X_{Mg}^b$ and $X_{Mg}^s$ in Reaction Zone 1 display an average value around 0.8 and 0.95, respectively (Table 4). Moreover, mass transfer is limited during Reaction Zone 1 formation (Fig. 7), indicating reaction in closed system conditions. Assuming a $\frac{Fe^{3+}}{Fe^{2+}+Fe^{3+}}$ of 1/3 in serpentine and neglecting Ni in olivine, this suggests that the reaction leading to Reaction Zone 1 formation is:

$$2(\text{Mg}_{0.91}, \text{Fe}_{0.09})_2\text{SiO}_4 + 3.024\text{H}_2\text{O} = (\text{Mg}_{0.95}\text{Fe}_{0.034}^{2+}\text{Fe}_{0.016}^{3+})_{2.988}\text{Si}_2\text{O}_5(\text{OH})_4 + 1.012(\text{Mg}_{0.79}\text{Fe}_{0.21})(\text{OH})_2 + 0.012\text{H}_2 \quad (8)$$

Temperature estimates of serpentinization with $K_D$ measured in natural samples falls below 50 °C and in the 50-275 °C range with the databases of McCollom and Bach [2009] (MC database) and Carlin et al. [2024] (CA database), respectively. $K_D$ calculations thus confirm the better fit with the CA database of temperature estimates of serpentinization based on oxygen isotopes (Fig. 8A, B and C). Finally, the $K_D$ values estimated with the experimental data of McCollom et al. [2016], Malvoisin et al. [2012], Okamoto et al. [2011], Ogasawara et al. [2013] and Lafay et al. [2012] are also better fitted with the CA database than with the MC database. Altogether, this indicates that the CA database obtained from experimental data on reaction 1 should be preferred to the MC database derived from solubility measurements to investigate Fe distribution and $H_2$ production during serpentinization. Even though measured $K_D$ fall in the same range as predicted $K_D$ with the CA database, important variations at a single temperature prevent any precise temperature estimate based on serpentine and brucite composition. This is mainly due to the lack of precise measurements of the $\frac{Fe^{2+}}{Fe^{2+}+Fe^{3+}}$ ratio of serpentine.
Awaruite production after olivine, a strong control on H₂ production

The presence of alloys is commonly reported in serpentinites [Frost, 1985, Alt and Shanks, 1998, Schwarzenbach et al., 2014, Eslami et al., 2021, Evans et al., 2023]. Alloys are found both as a reaction product of olivine alteration, and of primary Ni-sulfides such as pentlandite [Klein et al., 2009, Schwarzenbach et al., 2014, 2021, Evans et al., 2023, Steinthorsdottir et al., 2022]. The latter observation indicates formation by desulfidation in the presence of a reduced and S-poor fluid.

Reaction Zone 1 includes submicrometer awaruite grains. They are not associated with sulfides, suggesting formation by precipitation of the Ni released by olivine during serpentinization. Evans et al. [2023] proposed that the formation of such alloys involves disproportionation of the Fe²⁺ initially contained in olivine to form Fe⁰ and Fe³⁺. As we do not observe magnetite in Reaction Zone 1, Fe³⁺ is probably hosted in serpentine in the samples studied here. Regarding Ni, Ni reduction requires ferric serpentine formation, leading to the following reaction for awaruite formation:

$$8 [(Fe_{0.8125}Ni_{0.1875})_2SiO_4]^{\text{olivine}} + 12H_2O = Ni_3Fe + 4 [Fe_2^{3+}Si_2O_5(OH)_4]^{\text{serpentine}} + 4 [Fe(OH)_2]^{\text{brucite}} \quad (9)$$

In equation 9, olivine serpentinization does not produce H₂ while Fe³⁺ is incorporated into serpentine. Each mole of Ni reacting to form awaruite produces 8/3 moles of Fe³⁺ into serpentine. Considering olivine composition of (Mg₀.9064Fe₀.0896Ni₀.004)₂SiO₄ and a Mg/Mg+Fe in serpentine of 0.95 consistent with measurements in natural samples (Tables 3 and 4), this corresponds to a Fe³⁺/Fe²⁺+Fe³⁺ molar ratio in serpentine of ~ 0.28. Interestingly, the lower bound for Fe³⁺/Fe²⁺+Fe³⁺ ratio is of 0.3 for lizardite in natural samples in the compilation of Mayhew and Ellison [2020], and of 0.2 in the study of Andreani et al. [2013] where Fe oxidation variations
were investigated at the thin section scale. This indicates that a significant part of the Fe³⁺ incorporated into serpentine might not be associated with H₂ production but rather to Fe²⁺ disproportionation in association with awaruite production. Fe³⁺ in serpentine is thus both related to H₂ production and Fe²⁺ disproportionation; in the presence of awaruite, it is therefore not a direct proxy for H₂ production during serpentinization. Thermodynamic calculations show that H₂ production (Figs. 10F and 12F) is a complex function of serpentine composition and awaruite, magnetite and serpentine modes.

Awaruite formation implies relatively high activity of H₂ and low activity of sulfur in the fluid [Evans et al., 2017]. Such high activities are achieved through oxidation of the Fe²⁺ contained in olivine to form Fe³⁺-serpentine and/or magnetite. When the field of awaruite is entered, equilibria between awaruite and Ni-bearing phases such as brucite and serpentine limit the increase in H₂aq concentration through reactions such as:

$$3 [Ni(OH)_2]^{\text{brucite}} + [Fe(OH)_2]^{\text{brucite}} + 4H_{2,aq} = Ni_3Fe + 8H_2O$$

(10)

As a result, predicted H₂aq concentrations are always smaller than for models in which awaruite is not considered (Fig. 9). This effect is more pronounced at low temperature (Fig. 9F) and at low water to rock ratio (Fig. 10C and F), where awaruite modes are the highest. At 150 °C and the lowest water to rock ratio for complete olivine hydration, predicted H₂aq concentration is for example of 4 mmol/kg when considering awaruite (Fig. 10) and above 100 mmol/kg for a model without awaruite [Klein et al., 2009]. Ni content and X<sub>Mg</sub><sup>ol</sup> of olivine in mantle peridotites vary between
2500 and 4000 ppm and 0.89 and 0.92, respectively [Herzberg et al., 2013]. These
limited variations suggest that the calculations performed here for a Ni content of
3700 ppm and $X_{Mg}^{ol} = 0.91$ prevail for most mantle rocks, and thus that awaruite is
generally produced in contact with olivine during serpentinization.

Reaction Zone 2 formation: temperature or fluid flow is not the main driver

Reaction Zone 2 contains magnetite + Mg-brucite in the interior of veins bordering
the mesh texture. Such magnetite micro textures have been described in many natural
samples, but not necessarily in association with Mg-brucite [Bach et al., 2006, Beard
et al., 2009, Miyoshi et al., 2014]. Its spatial relationship with Reaction Zone 1 has
been interpreted by previous workers as the result of more open-system conditions at
the mesh rims. In particular, higher silica activity have been proposed to favor the
reaction of the Fe-brucite found in Reaction Zone 1 to form serpentine and magnetite
[Bach et al., 2006]. Frost et al. [2013], Schwarzenbach et al. [2016] and Li et al. [2020]
described an Mg-brucite + magnetite assemblage at the mesh rim similar to Reaction
Zone 2 in samples from New Caledonia, the Santa Elena ophiolite (Costa Rica) and
the Zedang ophiolite, respectively. In details, Reaction Zones 1 and 2 identified here
correspond to type 1 and 2/3 veins in Frost et al. [2013] and to zones 2/3 and zones
4-6 in Schwarzenbach et al. [2016], respectively. Bach et al. [2006] also reported the
presence of both Fe- and Mg- rich brucites in samples from ODP Leg 209. Two models
have been proposed to explain Reaction Zone 2 formation.

First, Li et al. [2020] proposed that the two types of reaction zones are formed
at the contact with olivine during two successive stages at different temperatures.
Reaction Zone 2 would form first at the mesh rim during a high-temperature alteration
event at temperature above 250 °C. Then, lower temperature conditions below 200
°C would favor Reaction Zone 1 formation. This latter interpretation of Reaction
Zone 2 formation relies on the thermodynamic predictions performed with the MC
database [Klein et al., 2014]. However, as shown above, the thermodynamic properties
of Carlin et al. [2024] should be preferentially used for thermodynamic modelling. For a mass water to rock ratio of 1, the minimum temperature of magnetite formation is of ~ 300 °C with this latter database (Fig. 9), which is not compatible with magnetite formation at the typical temperatures of serpentinization calculated with oxygen isotopes (Fig. 8 C). The stability field of magnetite increases with the water to rock ratio. Magnetite can thus be stable at temperatures as low as 200 °C for water to rock ratios above 10 (Fig. 10B). However, the predicted $X_{Mg}^b$ is constant or slightly decreases with water to rock ratio at these temperatures (Fig. 10D). Moreover, the concentration in $H_{2,aq}$ is approximately constant as the water to rock ratio increases (Fig. 10F), involving limited change in oxygen fugacity and $Fe^{3+}$ content in serpentine (Supplementary Fig. S7). Magnetite formation without significant changes in brucite composition as the fluid to rock ratio increases is due to the high solubility of Mg in the fluid, allowing for the following reaction to take place:

$$10 (Mg_{0.7}Fe_{0.3}) (OH)_2 = Fe_3O_4 + 7MgOaq + 9H_2O + H_{2,aq} \quad (11)$$

Increasing the fluid to rock ratio provides insight into the effect of high fluid flow in the rock. High fluid to rock ratios promote magnetite formation through Fe-brucite dissolution.

Such a reaction may explain the observation of brucite dissolution and Fe (hydroxide formation in pores in weathered serpentinized peridotites [Ulven et al., 2017, Jöns et al., 2017, Klein et al., 2020]. Reaction 11 suggests significant $H_{2,aq}$ production during weathering. However, we do not consider the presence of $O_{2,aq}$ in the serpentinizing fluid in the thermodynamic model. This latter consideration is relevant for near-surface conditions and would lead to a reaction similar to reaction 11 but with $O_{2,aq}$ consumption instead of $H_{2,aq}$ production. Despite $Fe^{2+}$ oxidation, weathering is thus not expected to produce significant $H_2$ in serpentinized peridotites.
Reaction 11 and thermodynamic model do not predict the highly magnesian composition of brucite measured in Reaction Zone 2 (Table 4). More generally, in the simulation performed here as a function of water to rock ratio, the maximum predicted $X_{Mg}^b$ is of 0.88 at 350 °C and of 0.78 at 200 °C. The predicted $X_{Mg}^b$ are below 5 out of the 6 values of $X_{Mg}^b$ determined here in Reaction Zone 2 in natural samples (Table 4). In addition, magnetite is predicted to be stable together with awaruite in parts of the investigated temperature - water to rock ratio space. Moreover, predicted molar $\frac{Ni}{Ni+Fe^{2+}}$ ratios in magnetite are below 0.01 (Fig. 10E), that is lower than most of the measured values in Reaction Zone 2 (Table 5). As a result, simulations as a function of water to rock ratio alone do not reproduce the observed mineralogical assemblages and mineral compositions. Altogether this suggests that Reaction Zone 2 formation cannot be explained only by changes in fluid to rock ratio, that is by significant fluid flow on olivine grain margins.

## Reaction Zone 2 formation: hydrogen diffusion as the main driver

Frost et al. [2013] proposed that magnetite formation in Reaction Zone 2 is associated with a fluid-dominated system in which Fe extraction from brucite in Reaction Zone 1 is responsible for magnetite formation in Reaction Zone 2. Such a reaction is consistent with the observation in Reaction Zone 2 of ~100 nm-wide domains containing Fe-rich brucite interpreted here as remnants of Reaction Zone 1 (Fig. 2). In samples from Papua New Guinea and ODP Leg 209, we also observe symplectites composed of fine-grained intergrowths of serpentine and Mg-brucite with a vermicular shape in Reaction Zone 2 (Fig. 2). Symplectites are common in retrogressed metamorphosed rocks [Vernon, 2018]. Their origin is debated and generally thought to involve volume-diffusion allowing for solid-state replacement reactions [Field, 2008, Gaidies et al., 2017]. Extrapolation of measured diffusion coefficients for oxygen and divalent cations in minerals at temperatures below 250 °C [Brady and Cherniak, 2010] indicates typical timescale above the age of the Earth for observed intergrowths formation by
volume diffusion. Putnis and Austrheim [2010] suggested that symplectite formation can involve a fluid phase. This was later confirmed with experiments showing that symplectite formation can be controlled by a dissolution-precipitation mechanism [Spruzeniece et al., 2017]. In this mechanism, a parent mineral first reacts with a fluid through a dissolution-precipitation reaction to form a porous reaction product. Then, the generated porosity is filled by a secondary product mineral. The microstructural features indicating such a replacement process consist in a sharp nanometer-scale reaction interface between the parent and product minerals, and the presence of porosity in the replacement rim. These two features are observed here with sharp boundaries at the nanoscale between Fe-brucite and Reaction Zone 2, and numerous pores in the symplectite (Fig. 2). Moreover, Mg-brucite grains appear to be enclosed in a matrix of serpentine. This could indicate that they form through precipitation in a porous serpentine matrix, in agreement with the mechanism proposed by Spruzeniece et al. [2017]. Interestingly, the calculated Mg solubility is orders of magnitude higher at more than 0.1 mol/kg than for the other cations involved in the reaction. This is consistent with precipitation of Mg-brucite as a secondary product mineral from the fluid phase. Using average brucite compositions measured in natural samples (Table 4) and assuming no mass transfer, Reaction Zone 2 formation through symplectite precipitation translates in the following reaction:

$$1.012 (Mg_{0.79}Fe_{0.21}) (OH)_2 = 0.85 (Mg_{0.94}Fe_{0.06}) (OH)_2 + 0.0538Fe_3O_4 + 0.109H_2O + 0.0538H_2 \quad (12)$$

The solid volume decreases by 7 % during reaction 12, probably leading to porosity generation during symplectite formation, in agreement with the model of Spruzeniece et al. [2017]. To our knowledge, symplectite formation has never been observed at temperatures as low as 200 °C. This further challenges the classical view on symplectite and their use to estimate P-T conditions or reaction rates.
In Frost et al. [2013]'s model, fO₂ increase at the mesh rim drives magnetite formation. In agreement with such a model, the symplectic microstructures and nanoscale observations discussed above show evidence for Reaction Zone 1 reaction to form Reaction Zone 2. Several other evidence indicates that Reaction Zone 2 is formed as a result of an increase in fO₂. First, at most of the studied localities, awaruite is present in Reaction Zone 1 and disappears in Reaction Zone 2 where magnetite contains up to 5 mol.% of the trevorite end-member (Table 5). This indicates the occurrence of the following reaction:

$$Ni_3Fe + 5 [Fe(OH)_2]^{\text{brucite}} + 2H_2O = 3 [NiFe_2O_4]^{\text{magnetite}} + 7H_2 \quad (13)$$

In this latter reaction, trevorite formation at the expense of awaruite is favored by a decrease in hydrogen activity (equivalent to an increase in fO₂). Reaction 13 is observed in all the samples with Reaction Zone 2 except in samples from the CM2B hole collected during the Oman drilling project. In Oman samples, awaruite is observed in Reaction Zone 2 and magnetite is Ni-free (Table 5). Interestingly, Mg-brucite in Oman samples also displays the lowest $X_{\text{Mg}}^b$ measured here in Reaction Zone 2 (0.83 whereas $X_{\text{Mg}}^b$ in other samples are comprised between 0.92 and 0.98; Table 4). According to the brucite / magnetite equilibrium of reaction 1, lower $X_{\text{Mg}}^b$ in brucite involves higher hydrogen fugacity, which is consistent with the preservation of awaruite in Reaction Zone 2 from Oman. Another difference between Reaction Zone 2 in Oman and at other locations is the Mn:Fe positive correlation in serpentine/brucite mixtures (Fig. 6B), suggesting that Fe-brucite incorporates most of the Mn during incipient serpentinization. Mn content as a function of Fe content displays a lower slope in Oman than in other Reaction Zones 2 studied here; this slope is similar to the slope measured in Reaction Zone 1. Mn content in magnetite is low compared to Fe content (Table 5). As a result, magnetite formation probably concentrates Mn into brucite. This suggests that less magnetite has precipitated during Reaction Zone
2 formation in Oman, in agreement with the measured $X_{Mg}^b$ (Fig. 8D) and Mn:Fe ratio in brucite (Fig. 6B).

Two drivers are proposed for $fO_2$ increase at the mesh rim in Frost et al. [2013]: (1) the input of a fluid with a higher oxygen fugacity than the surrounding serpentinite, equivalent to increasing the water to rock ratio in the model, or (2) $H_2$ loss by diffusion. As shown above with simulations as a function of the water to rock ratio, (1) induces brucite dissolution without reproducing the low $X_{Mg}^b$ and awaruite reaction observed in natural samples. Simulations as a function of $X_{Oi}$ allow to model stable mineralogical assemblages for systems with $H_2$ loss; they thus provide insights on (2). It appears that, in addition to magnetite formation, $X_{Oi}$ increase induces awaruite disappearance, a strong increase of $X_{Mg}^b$ up to a value of 1 and an increase of the Ni content in magnetite to values comparable to those measured in natural samples (Fig. 12 and Table 5). This suggests that $H_2$ loss by diffusion is the main driver for Reaction Zone 1 reaction to form Reaction Zone 2. In addition to the compositional changes measured here, thermodynamic modelling predicts a strong difference between $Fe^{3+}$ over $Fe_{tot}$ molar ratio in serpentine ($X_{Fe3}^s$) between the two proposed drivers for $fO_2$ decrease. $X_{Fe3}^s$ does not strongly vary with water to rock ratio whereas it strongly increases with $X_{Oi}$ up to values of $X_{Fe3}^s \sim 1$. A way to confirm that $H_2$ loss by diffusion is the main driver for reaction would thus consist in measuring $X_{Fe3}^s$ in the reaction zones.

## Model of serpentinization under a RedOx gradient: implications for $H_2$ production and mass transfer during serpentinization

The observations performed here in serpentinized peridotite samples collected at mid-ocean ridges, in ophiolites and ultramafic bodies as well as thermodynamic modelling allow to propose a general model for olivine serpentinization at low pressure and temperatures lower than $\sim 350$ °C (Fig. 13).
First, a magnetite-free assemblage composed of serpentine, Fe-brucite and awaruite (Reaction Zone 1) is formed after olivine through reactions 8 and 9. These reactions lead to a solid volume increase of ~ 50 %, consistent with measurements for serpentinization in natural samples [Malvoisin et al., 2020] and in experiments [Klein et al., 2020]. Malvoisin et al. [2021a] and Malvoisin et al. [2021b] have shown that Fe-brucite and lizardite are mixed at the nanometer-scale in Reaction Zone 1, and that nanoscale porosity probably plays a key role in allowing for water to continuously flow up to the reaction front even though the solid volume increases. We observe here that Reaction Zone 1 formation does not involve significant mass transfer (Fig. 7), which is consistent with other observations suggesting that serpentinization can be isochemical except for the addition of water [Coleman and Keith, 1971, Paulick et al., 2006, Beard et al., 2009, Malvoisin et al., 2020]. Reaction 8 produces H₂, which will diffuse out of the system, for example up to the seafloor or the continental surface. A concentration gradient thus establishes between the serpentinization front and the mesh rim. If H₂aq diffusion rate is slow compared to H₂ production rate by reaction 8, Reaction Zone 1 will dominate in serpentinized peridotites and H₂aq concentrations will be fixed by equilibrium at values between 10⁻³ and 10⁻² mol/kg in the 100-250 °C temperature range (Fig. 12F), that is for the typical temperatures of serpentinization (Fig. 8C). These concentrations are the highest concentrations expected for H₂aq during serpentinization. Interestingly, they are consistent with the H₂aq concentrations measured in fluids expelled at ultramafic-hosted hydrothermal sites at mid-ocean ridges such as Lost City [0.25 to 14 mmol/kg; Kelley et al., 2001, Proskurowski et al., 2006, Lang et al., 2010], Rainbow [16 mmol/kg; Charlou et al., 2002], Logatchev [12 mmol/kg; Charlou et al., 2002], Nibelungen [11 mmol/kg; Melchert et al., 2008], and Ashadze [10 to 26 mmol/kg; Charlou et al., 2010]. A serpentinizing fluid with the maximum H₂aq concentration predicted here (~ 5·10⁻² mol/kg) reaches saturation at a depth of approximately 1000 m [Brunet and
Malvoisin, 2025]. This is consistent with the observation of H₂ bubbling at the surface in hyperalkaline springs and gas seeps found in ophiolites [Abrajano et al., 1990, Leong et al., 2023, Etiope, 2023, Baciu and Etiope, 2024]. However, this implies that H₂ gas accumulation is unlikely at depth greater than 1000 m with consequences for natural H₂ exploration, and in agreement with recent models of H₂ generation and transport [Brunet and Malvoisin, 2025, Christiansen et al., 2025]. The upper limit fixed here by thermodynamic calculations on H₂,aq concentration is a first-order constraint for assessing H₂ flow in models dealing with H₂ generation and transport in deep aquifers [Brunet and Malvoisin, 2025].

If H₂,aq diffusion rate is faster than H₂ production rate by reaction 8, H₂,aq concentration in pore water will significantly decrease. This will lead to Reaction Zone 1 replacement by Mg-brucite and Ni-bearing magnetite (Reaction Zone 2) according to Reactions 12 and 13. Reaction 12 induces a decrease in solid volume of 7 % favoring symplectite formation through a dissolution-precipitation mechanism and significant mass transfer as observed in Fig. 7. The most obvious transfers consist in Fe and Mg transfers to form magnetite and Mg-brucite grains, reaching several tens of micrometers in width at the mesh rim. We also observe here Ni segregation into magnetite. The studied samples contain a small amount of sulfides. In sulfur-rich rocks, Ni could also be incorporated in secondary sulfides. Finally, we observe a significant loss of Ca in Reaction Zone 2 compared to olivine and Reaction Zone 1. Such a loss may be involved in Ca metasomatism (rodingitization) associated with serpentinization [Bach et al., 2013]. We focus here on Reaction Zone 2 containing Mg-brucite, but mass transfer in pyroxene-bearing rocks will probably involve silica input into the reaction zone leading to serpentine instead of Mg-brucite precipitation. Magnetite + serpentine assemblages at the mesh rim are commonly observed in serpentinized peridotites [Bach et al., 2006, Beard et al., 2009]. Fig. 11 reveals that Si input, modelled here with an increase of the orthopyroxene fraction in the
reacting assemblage, favors serpentine production but does not produce magnetite at temperatures as high as 250 °C. As a result, magnetite + serpentine assemblages are probably formed through a combination of Reaction Zone 1 destabilization due to hydrogen fugacity decrease and Si input due to pyroxene breakdown. Observations of natural samples show more intense fracturing in olivine than in orthopyroxene, leading to faster olivine reaction during incipient serpentinization [Rouméjon and Cannat, 2014]. Moreover, the free energy of olivine hydrolysis is lower than that of orthopyroxene at temperatures below 250 °C, implying faster olivine reaction [Allen and Seyfried, 2003]. This may explain why Reaction Zone 1 can locally develop even in harzburgite before Si input. Samples studied here from the Southwest Indian Ridge and collected by dredging contain Reaction Zone 1 in restricted areas. Reaction Zone 1 is bordered by an association of magnetite, serpentine and pores in a microtexture similar to Reaction Zone 2 (Supplementary Fig. S6). The presence and location of the pores indicate dissolution, probably of Mg-brucite. This suggests that Mg-brucite in Reaction Zone 2 dissolves faster than Fe-brucite intermixed with serpentine in Reaction Zone 1. Mg-brucite may thus not be preserved in weathered samples due to fluid flow [Ulven et al., 2017, Klein et al., 2020]. In addition to Si input, this may explain why Reaction Zone 2 is not so commonly described in serpentinized peridotites.

Mineral compositions measured in Reaction Zone 2 can be compared to thermodynamic calculations to retrieve H₂aq concentrations at equilibrium (Fig. 14). The values retrieved from the brucite and magnetite compositions measured here provide similar estimates of H₂aq. In the 100-250 °C temperature range, they vary between 10⁻⁷ and 10⁻³ mol/kg. These values are generally lower than the values reported for fluids expelled at mid-ocean ridges in ultramafic-hosted hydrothermal vents. This suggests that H₂ production associated with Reaction Zone 2 formation is mainly associated with diffuse transport in the oceanic crust.
Reaction 8 leading to magnetite-free Reaction Zone 1 formation produces ~ 0.04 mol of H₂ per kg of ultramafic rocks. This is approximately 5 times less than H₂ production during magnetite formation according to reaction 12 (0.18 mol of H₂ per kg of ultramafic rocks). As a result, ultramafic rocks can preserve most of their H₂ production potential even when fully serpentinized, as suggested by previous studies [Carlin et al., 2024]. Moreover, this indicates that predicting global H₂ fluxes does not only require to determine serpentinization degree but also the proportions of Reaction Zone 1 and Reaction Zone 2 among the reaction products. In studies of global H₂ flux estimates, H₂ production is generally calculated by assuming either complete Fe oxidation [Sherwood Lollar et al., 2014] or full incorporation in magnetite of the Fe initially contained in olivine [Canfield et al., 2006, Emmanuel and Ague, 2007, Worman et al., 2016, Rüpke and Hasenclever, 2017, Christiansen et al., 2025]. This latter approach is equivalent to assume that serpentinization only leads to Reaction Zone 2. H₂ production could thus be overestimated by almost one order of magnitude if Reaction Zone 1 predominates among the reaction products.

A way to estimate the proportion of Reaction Zone 1 and Reaction Zone 2 is to use magnetic data. Reaction Zone 1 is expected to lead to a limited magnetic signal whereas Reaction Zone 2 contains magnetite with a high magnetization. Cannat et al. [2010] used such an approach with empirical curves fitting the measured magnetite content/serpentinization degree relationships to estimate global H₂ fluxes. Oufi et al. [2002] showed with magnetic data acquired in abyssal serpentinized peridotites that magnetic susceptibility dramatically increases for serpentinization degrees above ~ 70 %. In the model proposed here (Fig. 13), such a behaviour would imply that H₂ diffusion becomes more efficient than H₂ production at high serpentinization degrees. This could result from both a lower amount of olivine producing H₂ in the rock, and better developed H₂ diffusion pathways at high serpentinization degree. Hong et al. [2022] also observed magnetic susceptibility
increase with serpentinization degree in samples from the Oman ophiolite. However, the measured susceptibilities are significantly lower than in abyssal peridotite. Moreover, the magnetite susceptibility/serpentinization degree relationship provided in Oufi et al. [2002] and Hong et al. [2022] does not prevail in all geological settings. For example, low magnetic susceptibility is measured at all serpentinization degrees in the Pindos ophiolite [Bonnemains et al., 2016] and in samples collected at passive margins [Klein et al., 2014]. This variety in magnetite production with serpentinization degree probably arises from the complex interplay between reaction and H₂ diffusion. Depth, tectonic deformation and hydrothermal circulation probably play a key role on H₂ diffusion. As a result, magnetite contents are expected to significantly vary for a single geodynamic setting. For example, at slow-spreading mid-ocean ridges, deformed ultramafic rocks exposed at the surface along detachment faults are expected to contain different amounts of magnetite than rocks occurring in the foot wall and the hanging wall of the fault. A sampling bias exists for serpentinized peridotites from mid-ocean ridges since they are almost exclusively sampled in the damaged zone of detachment faults. More generally, reaction and H₂ diffusion are also expected to depend on the geodynamic setting. Fluid pathways are different at (ultra-)slow-spreading mid-ocean ridges where extension is controlled by tectonic deformation along detachment faults, than at fast-spreading mid-ocean ridges where magmatism mainly accommodates plate divergence. Samples collected at passive margins underwent limited magmatism while tectonic extension is locally important, similar to slow-ultraslow contexts. Similarly, the large range of geodynamic processes leading to ophiolite formation [Dilek and Furnes, 2014] is expected to lead to significantly different spatial and temporal evolutions of fluid pathways, and thus of H₂ diffusion.

The model proposed here applies to low pressure serpentinization in ophiolites and at mid-ocean ridges. It is relevant for olivine- and/or brucite-bearing rocks.
Therefore, it does not directly apply to other reactions potentially producing H₂ in ultramafic rocks (i.e., Fe²⁺-bearing serpentine oxidation, Fe³⁺ garnet formation). However, one should note that, since ~ 70 % of the Fe²⁺ initially contained in olivine is transferred to brucite during serpentinization (equation 8), the remaining H₂ production potential in brucite- and olivine-free rocks is expected to be low. The model does not directly apply to high pressure serpentinization in subduction zones, e.g., in the forearc mantle [Hyndman and Peacock, 2003]. In these conditions, serpentinization is expected to proceed at higher temperatures and will mainly produce antigorite instead of lizardite. Further investigations are therefore required to characterize the products of high pressure serpentinization in conditions relevant to subduction zones.

## Conclusions

The study of serpentinization products at 13 localities at mid-ocean ridges, in ophiolites and in ultramafic bodies has several important consequences for understanding and modelling the serpentinization process:

* Olivine first reacts with water to form an assemblage of serpentine, Fe-brucite and awaruite mixed at the submicrometer scale (Reaction Zone 1). This reaction is isochemical except for the addition of water.

* Awaruite formation involves Fe²⁺ disproportionation with Fe³⁺ incorporation into serpentine. Mass balance considerations indicate that up to 30 % of the Fe contained in serpentine can be Fe³⁺ associated with awaruite production. This reaction does not involve H₂ production. Using Fe³⁺ in serpentine as a proxy for H₂ production during serpentinization thus requires to also consider Ni incorporation in awaruite.

* Mineral compositions in Reaction Zone 1 are better reproduced with thermodynamic modelling by using the thermodynamic data of Carlin et al. [2024]
for Fe(OH)₂. The use of this endmember has strong consequences for predictions of Fe distribution during serpentinization. First, lower $\frac{Mg}{Mg+Fe}$ molar ratios are predicted in brucite than with previous databases. Then, no magnetite is predicted to form at mass water to rock ratios below 1 and at temperatures below 300 °C, that is at the temperatures of serpentinization in natural systems measured with oxygen isotopes. This indicates that temperature is not the main driver for magnetite formation during serpentinization.

* At the mesh rim, another assemblage composed of serpentine, Mg-brucite and magnetite is found (Reaction Zone 2). Nanoscale investigation reveals that Reaction Zone 2 is formed at the expense of Reaction Zone 1. This reaction induces a decrease in volume of 7% and thus generates porosity. Porosity generation leads to symplectite formation through dissolution-precipitation and to mass transfer. To our knowledge, symplectite microtexture formation has never been reported before at temperatures below 300°C.

* Awaruite is generally not observed in Reaction Zone 2 whereas magnetite contains up to 5 mol.% of the Ni endmember. Awaruite thus reacts to form Ni-bearing magnetite. This reaction is predicted to occur through a decrease in H₂ fugacity. Minor element composition of magnetite can thus provide indications on RedOx conditions during serpentinization.

* Mineral compositions in Reaction Zone 2 are only reproduced with thermodynamic modelling when considering H₂ loss. H₂ diffusion in the fluid is thus the main driver for magnetite formation during serpentinization.

* The total amount of produced H₂ during serpentinization as well as H₂,aq concentration at the equilibrium strongly depends on the proportion and type of reaction zone in the reaction products. Reaction Zone 1 formation produces 0.04 mol of H₂ per kg of ultramafic rocks at concentrations of 10⁻³ to 10⁻² mol/kg. Reaction Zone 2 formation produces 0.18 mol of H₂ per kg of ultramafic
rocks at concentrations of 10⁻⁷ to 10⁻³ mol/kg. It is thus essential to consider H₂ diffusion in ultramafic rocks to estimate global natural H₂ production and H₂ concentration in natural fluids. Predicting H₂ fluxes from ultramafic rocks requires consideration of the local environment as well as the geodynamic context.

## Acknowledgments

B. R. Frost and I. Katayama are thanked for providing their electron microprobe data for Reaction Zone 1 in samples from New Caledonia and the Mineoka ophiolite complex, respectively. J.A.D. Connolly is thanked for helping in the selection of the fluid equation of state for H₂-rich fluid in Perple_X. This work is supported by ANR through the NaturaLDH project (ANR-22-CE01-0026). D. Troadec is thanked for the preparation of the FIB sections. FIB experiments are supported by the French RENATECH network, the CPER Hauts de France project IMITECH and the Métropole Européenne de Lille. V. Magnin and V. Batanova are thanked for support during the acquisition of EPMA data. We are grateful to K. Steinthorsdottir and K. Evans for constructive reviews and to T. Mikouchi for editorial handling.

This research uses samples and/or data provided by the Oman Drilling Project. The Oman Drilling Project (OmanDP) has been possible through co-mingled funds from the International Continental Scientific Drilling Project (ICDP; Kelemen, Matter, Teagle Lead PIs), the Sloan Foundation – Deep Carbon Observatory (Grant 2014-3-01, Kelemen PI), the National Science Foundation (NSF-EAR-1516300, Kelemen lead PI), NASA – Astrobiology Institute (NNA15BB02A, Templeton PI), the German Research Foundation (DFG: KO 1723/211, Koepke PI), the Japanese Society for the Promotion of Science (JSPS no:16H06347, Michibayashi PI; and KAKENHI 16H02742, Takazawa PI), the European Research Council (Adv: no.669972; Jamveit PI), the Swiss National Science Foundation (SNF:20FI21_163073, Frü h-Green PI), JAMSTEC, the TAMU-JR Science Operator, and contributions
from the Sultanate of Oman Ministry of Regional Municipalities and Water Resources,
the Oman Public Authority of Mining, Sultan Qaboos University, CRNS-Univ.
Montpellier II, Columbia University of New York, and the University of Southampton.

Data availability

All data are incorporated into the article and its online supplementary material. They
are also available on the EarthChem repository: https://doi.org/10.60520/IEDA/113834.

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<figure><img src="image_2.png" /><figcaption>Fig. 1: Back-scattered electron images of reaction products after olivine in the serpentinized peridotites studied here. A (RØR0307 sample; Feragen ultramafic body, Norway): olivine grains with a saw-toothed boundary altered into a mixture of serpentine and Fe-brucite near the reaction front (Reaction Zone 1). Further away from the reaction front, the reaction products consist in distinct grains of magnetite, Mg-brucite and serpentine (Reaction Zone 2). Note the presence of bright spots corresponding to awaruite grains only occurring in the serpentine/Fe-brucite mixture. B (sample 1274A-8R1-16bis; ODP Leg 209): olivine grain surrounded by, first, a domain with serpentine and Fe-brucite intermixed at the submicrometer scale (Reaction Zone 1) and, then, a vein filled with magnetite and Mg-brucite (Reaction Zone 2). C (sample 96AL-10B; Mirdita ophiolite, Albania): Fe-brucite intermixed with serpentine. D (sample 96AL-8; Mirdita ophiolite, Albania): domains containing high amounts of Fe-brucite. Note the saw-toothed shape of the border of Fe-brucite grains, similar to the boundary of dissolving olivine grains with etch pits. E (sample 807-58-01-2-6 from CM2B borehole, Oman drilling project): reaction products of olivine serpentinization with two distinct reaction zones composed of 1) a mixture of serpentine/Fe-brucite at the contact with olivine (Reaction Zone 1) and 2) distinct grains of serpentine, Mg-brucite and magnetite (Reaction Zone 2). F (sample A5; Papua-New Guinea): olivine surrounded by a mixture of serpentine/Fe-brucite and then by distinct grains of serpentine, Mg-brucite and magnetite. Note that serpentine and Mg-brucite can either occur as 5 to 10 µm-wide grain or in a symplectite texture. G (sample 1274A-8R1-16bis, ODP Leg 209) and H (sample A5, Papua New Guinea): close view of the reaction products in Reaction Zone 2 with Mg-rich brucite either displaying a mosaic (G and H) or a symplectite texture (H). Ol: olivine; Serp: serpentine/Fe-brucite mixture richer in serpentine; Fe-Brc: serpentine/Fe-brucite mixture dominated by Fe-brucite; Mg-Brc: Mg-rich brucite; Mag: magnetite; Symp: symplectite texture. RZ1: Reaction Zone 1. RZ2: Reaction Zone 2.</figcaption></figure>
<figure><img src="image_3.png" /></figure>

Fig. 2: TEM-bright field images of the Papua New Guinea A5 sample (Reaction Zone 2). A, B, C: detail of the grain texture in Fig. 1H. A: Mg-rich brucite (amorphized; aMg-Brc) with intersecting serpentine domains consisting of chrysotile and lizardite. B: Fe-brucite at the contact between Mg-brucite and lizardite + chrysotile domains. C: lizardite and Fe-brucite occurring together. D, E, F: symplectite texture as in Fig. 1F. D: Overview of idiomorphic Mg-brucite and lizardite, Fe-brucite recognized by the wave-shaped pore space. E: Mg-brucite crystals parallel to lizardite. In places, Fe-brucite is located at the corners of the Mg-brucite. White areas on the bottom left are drying cracks produced during sample preparation. F: Detail of the Fe-brucite accompanied by wave-shaped pore space.
<figure><img src="image_4.png" /><figcaption>Fig. 3: Back-scattered electron images and respective compositional maps of serpentinization products at olivine contact measured by wavelength dispersive microscopy in samples 807-C5708B-58-01-2-6 (CM2B hole from Oman Drilling Project; A, C, E, G, I) and RØR0307 (Feragen ultramafic body; B, D, F, H, J). A and B: Back-scattered electron images. C and D: $X_{Br}$, molar brucite fraction in the serpentine/brucite mixture. E and F: $X_{Mg}^b$, Mg/(Mg+Fe) molar ratio in brucite. G and H: NiO content. I and J: MnO content. Ol: olivine; Serp: serpentine/Fe-brucite mixture richer in serpentine; Fe-Brc: serpentine/Fe-brucite mixture dominated by Fe-brucite; Mg-Brc: Mg-rich brucite; Mag: magnetite. RZ1: Reaction Zone 1. RZ2: Reaction Zone 2.</figcaption></figure>
<figure><img src="image_5.png" /><figcaption>Fig. 4: Back-scattered electron images and respective compositional maps of serpentinization products at olivine contact measured by wavelength dispersive microscopy in samples 1274A-8R1-16 (ODP Leg 209, Mid-Atlantic Ridge; A, C, E, G, I, K) and A5 (Papuan ultramafic belt; B, D, F, H, J, L). A and B: Back-scattered electron images. C and D: XBr, molar brucite fraction in the serpentine/brucite mixture. E and F: XMg, Mg/(Mg+Fe) molar ratio in brucite. G and H: NiO content. I and J: MnO content. I and J: CaO content. OI: olivine; Serp: serpentine/Fe-brucite mixture richer in serpentine; Fe-Brc: serpentine/Fe-brucite mixture dominated by Fe-brucite; Mg-Brc: Mg-rich brucite; Mag: magnetite. RZ1: Reaction Zone 1. RZ2: Reaction Zone 2.</figcaption></figure>
<figure><img src="image_6.png" /><figcaption>Fig. 5: Fe vs. Si + Al/2 (apfu) of serpentine/brucite mixtures. The composition is normalized to seven oxygens. A: Mirdita ophiolite (Albania). B: ODP Leg 209. C: Leka ophiolite (Norway). D: Hess Deep. E: Papua New Guinea. F: Feragen ultramafic body ophiolite. G: Southwest Indian Ridge. H: Oman Drilling Project (CM2B hole). Black and red dots correspond to data acquired in the Reaction Zone 1 and Reaction Zone 2, respectively. Blue dots correspond to data acquired in Reaction Zone 1 in sample 96AL-8. Black and red stars correspond to TEM-EDS data. The plain and dashed lines correspond to fits of the electron microprobe and TEM-EDS data, respectively. They are used to determine serpentine and brucite compositions.</figcaption></figure>
<figure><img src="image_7.png" /></figure>

Fig. 6: Composition of all serpentine/brucite mixtures analyzed in this study, and reported in the studies of Klein et al. [2014], Schwarzenbach et al. [2016], Frost et al. [2013], Li et al. [2020], Katayama et al. [2010]. A: Mg + Fe + Mn + Al/2 vs. Si + Al/2 (apfu). B: Mn vs. Fe (apfu). C: Ni vs. Fe (apfu). Black and red dots correspond to analyses performed in Reaction Zone 1 and Reaction Zone 2, respectively. In B and C, data collected in Reaction Zone 2 in samples from CM2B hole of the Oman drilling project are displayed with orange stars.
<figure><img src="image_8.png" /><figcaption>Fig. 7: Anhydrous mass change in SiO₂, MgO and FeO for Reaction Zone 1 (black bars) and Reaction Zone 2 (red bars) with respect to olivine (ΔX, see Methods for definition). ΔX is calculated with local bulk rock composition computed from X-ray maps acquired on samples 807-C5708B-58-01-2-6 (CM2B hole from Oman Drilling Project, first row), 1274A-8R1-16 (ODP Leg 209, Mid-Atlantic Ridge, second row), A5 (Papuan ultramafic belt, third row) and RØR0307 (Feragen ultramafic body, fourth row).</figcaption></figure>
<figure><img src="image_9.png" /><figcaption>Fig. 8: Comparisons of estimated temperature of serpentinization and distribution coefficients of Fe²⁺ and Mg between serpentine and brucite (K_D) for thermodynamic calculations, and experimental and natural data. A: histogram of K_D distribution in the natural samples studied here. Black and red bars correspond to calculated K_D when assuming Fe³⁺/(Fe²⁺+Fe³⁺) in serpentine (X<sub>Fe3</sub><sup>s</sup>) of 1/3 and 2/3, respectively. Horizontal black and red dashed lines are the average ± standard deviation of the K_D values calculated when assuming X<sub>Fe3</sub><sup>s</sup> = 1/3 or 2/3, respectively. B: K_D as a function of temperature. Black and blue curves correspond to computed K_D with the thermodynamic databases of Carlin et al. [2024] and McCollom and Bach [2009], respectively. Red and black symbols correspond to calculated K_D for experimental data assuming X<sub>Fe3</sub><sup>s</sup> = 1/3 and 2/3, respectively. The data of McCollom et al. [2016], Malvoisin et al. [2012], Okamoto et al. [2011], Ogasawara et al. [2013] and Lafay et al. [2012] are displayed with dots, diamonds, stars, triangles, and crosses, respectively. C: distribution of temperature estimates for serpentinization based on oxygen isotope measurements in serpentine from natural samples collected at mid-ocean ridges (gray bars), passive margins (yellow bars) and ophiolites (green bars; see references in the Methods section). Vertical gray, yellow and green dashed lines correspond to the average ± standard deviation values of the temperature estimates for mid-ocean ridges, passive margins and ophiolites, respectively. D and E: molar Mg/(Mg+Fe) ratios measured in natural samples for brucite (D; X<sub>Mg</sub><sup>b</sup>) and serpentine (E; X<sub>Mg</sub><sup>s2+</sup>). Black and red bars in D correspond to values measured in Reaction Zone 1 and Reaction Zone 2, respectively. X<sub>Mg</sub><sup>b</sup> < 0.85 in Reaction Zone 2 is measured here in samples from the Oman Drilling project (CM2B hole). Red and black bars in E correspond to X<sub>Mg</sub><sup>s2+</sup> values calculated by assuming X<sub>Fe3</sub><sup>s</sup> = 1/3 and 2/3, respectively.</figcaption></figure>
<figure><img src="image_10.png" /><figcaption>Fig. 9: Influence of Ni on predicted mineralogy and fluid composition with the CA database as a function of temperature during olivine serpentinization (water to rock ratio of 1 and pressure of 50 MPa). Simulations are performed either without considering Ni in the system (A, C and E) or with Ni in the system (B, D and F). A and B: mineral modes. C and D: Mg/(Mg+Fe²⁺) molar ratio in minerals. The plain, dashed, dotted, dash-dotted and cross lines correspond to brucite, serpentine, olivine, awaruite and magnetite, respectively. E and F: concentration in H₂aq in the fluid.</figcaption></figure>
<figure><img src="image_11.png" /><figcaption>Fig. 10: Predicted mineralogy and fluid composition as a function of temperature and mass water to rock ratio during Ni-bearing olivine serpentinization at 50 MPa. A: brucite mode (mol. fraction). B: magnetite mode (mol. fraction). C: awaruite mode (mol. fraction). D: $X_{Mg}^b$ ($\frac{Mg}{Mg+Fe}$ molar fraction in brucite). E: trevorite (NiFe₂O₄) molar fraction in trevorite/magnetite solid solution. F: decimal logarithm of $H_{2,aq}$ concentration in the fluid.</figcaption></figure>
<figure><img src="image_12.png" /><figcaption>Fig. 11: Predicted mineralogy and fluid composition as a function of orthopyroxene over orthopyroxene + olivine molar ratio at 250 °C and 50 MPa. A: mineral modes (mol. fraction). B: Mg/(Mg+Fe²⁺) molar ratio in minerals. The plain, dashed, dotted and dash-dotted lines correspond to brucite, serpentine, talc and awaruite, respectively. C: concentration in H₂ in the fluid.</figcaption></figure>
<figure><img src="image_13.png" /><figcaption>Fig. 12: Predicted mineralogy and fluid composition as a function of temperature and initial fluid $\frac{\text{oxygen}}{\text{oxygen}+\text{hydrogen}}$ ratio ($X_{Oi}$) during Ni-bearing olivine serpentinization at 50 MPa with a mass water to rock of 1. A: brucite mode (mol. fraction). B: magnetite mode (mol. fraction). The stability fields of hematite (Hem) and goethite (Gth) are also displayed. C: awaruite mode (mol. fraction). D: $X_{Mg}^b$ ($\frac{Mg}{Mg+Fe}$ molar fraction in brucite). E: trevorite ($NiFe_2O_4$) molar fraction in trevorite/magnetite solid solution. F: decimal logarithm of $H_2$ concentration in the fluid.</figcaption></figure>
<figure><img src="image_14.png" /><figcaption>Fig. 13: Scheme of the model of serpentinization proposed here based on observations in natural samples and thermodynamic modelling. Reaction Zone 1 is formed at the contact with olivine. It then reacts at the mesh rim to form Reaction Zone 2. The main driver for reaction is hydrogen diffusion from the serpentinization front towards the mesh rim.</figcaption></figure>
<figure><img src="image_15.png" /><figcaption>Fig. 14: Predicted decimal logarithm of H₂ₐq concentration in the fluid as a function of temperature and initial fluid $\frac{\text{oxygen}}{\text{oxygen}+\text{hydrogen}}$ ratio (XOᵢ) during Ni-bearing olivine serpentinization at 50 MPa with a mass water to rock ratio of 1. The black and gray lines corresponds to isopleths of brucite (X<sub>Mg</sub><sup>b</sup>) and magnetite (trevorite fraction; X<sub>Ni</sub>) compositions, respectively. The isopleths are provided for the minimum and maximum values of X<sub>Mg</sub><sup>b</sup> and X<sub>Ni</sub> measured in natural samples in Reaction Zones 2 containing Ni-bearing magnetite. The highlighted area corresponds to values consistent with measured X<sub>Mg</sub><sup>b</sup> and X<sub>Ni</sub> values in the typical temperature range of serpentinization.</figcaption></figure>