# Constraints on Fe-Ca metasomatism in mineralized mantle rocks: Insights from in-situ geochemistry and thermodynamic modeling

Rémi Coltat, C.G.C. Patten, J.A. Padrón-Navarta, F. Hochscheid, Marc Ulrich, Yannick Branquet, Philippe Boulvais, C.J. Garrido

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## To cite this version:

Rémi Coltat, C.G.C. Patten, J.A. Padrón-Navarta, F. Hochscheid, Marc Ulrich, et al., Constraints on Fe-Ca metasomatism in mineralized mantle rocks: Insights from in-situ geochemistry and thermodynamic modeling. Chemical Geology, 2025, 687, pp.122826. 10.1016/j.chemgeo.2025.122826. insu-05064010v2

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Research Paper

# Constraints on Fe-Ca metasomatism in mineralized mantle rocks: Insights from in-situ geochemistry and thermodynamic modeling☆

R. Coltat<sup>a,b,*</sup>, C.G.C. Patten<sup>c</sup>, J.A. Padrón-Navarta<sup>a</sup>, F. Hochscheid<sup>d</sup>, M. Ulrich<sup>e</sup>, Y. Branquet<sup>f</sup>, P. Boulvais<sup>f</sup>, C.J. Garrido<sup>a</sup>

## ARTICLE INFO

Editor: S Aulbach

Keywords:

Fluid-rock interactions

Fe-Ca metasomatism

Ultramafic-hosted volcanogenic massive sulfide deposits

Ilvaite

Platta nappe

Thermodynamic modeling

## ABSTRACT

Ultramafic-hosted seafloor massive sulfide deposits have been reported in present-day oceanic settings for nearly thirty years. However, the development of comprehensive genetic models that account for deep-seated hydrothermal processes is largely hindered by the limited availability of seafloor observations and their reliance on large-scale geophysical studies. The Platta nappe (Swiss Alps) preserves a Jurassic hydrothermal system (the Marmorera-Cotschen Hydrothermal System; MCHS), where Cu-Fe-Co-Zn-Ni mineralization is associated with Fe-Ca silicates (ilvaite, hydrogarnet, and diopside). Petrographic analyses and thermodynamic modeling indicate that Fe-Ca metasomatism occurred between 300 and 360 °C and at low fO₂ (from FMQ –6 to +1), likely coeval with early-stage serpentinization. The composition of Fe-Ca silicates (Co, Ni, and REE contents, measured by in-situ LA-ICP-MS) indicates fluid-rock interaction from an ultramafic-dominated system to an open-system, involving fluids derived from both mafic and ultramafic rocks. Mineralogical and geochemical signatures of Fe-Ca silicates in the MCHS do not support genetic relationships with common rodingitization. Our results highlight that Fe-Ca metasomatism may be a widespread deep-seated alteration along mafic-ultramafic rock contacts or in mantle rocks modified through melt-rock interaction accompanying mantle exhumation.

## 1. Introduction

Mantle exhumation occurs at extensional oceanic settings (e.g., (ultra)slow-spreading Mid-Oceanic Ridges, MORs; Ocean-Continent Transitions, OCTs; fore-arc and back-arc spreading centers), where the magmatic supply is not high enough to accommodate the total amount of divergence. It generally involves the formation of high-offset normal fault (e.g., detachment), exhuming deep lithospheric rocks (e.g., peridotites and gabbros) in dome-like massif at the seafloor (Escartín and Olive, 2022). The coupling between extensional tectonics, magmatic inception and seawater circulation triggers hydrothermal circulation leading to numerous low- (< 300 °C) to high-temperature (HT, > 300 °C) alterations and related hydrothermal fluids (e.g., Schwarzenbach and Steele-MacInnis, 2020). Among alterations, the most

prominent consist of the serpentinization and carbonation of mantle rocks (Schwarzenbach and Harris, 2025), while formation of mineral deposits—the so-called ultramafic-hosted seafloor massive sulfide deposits (UM-SMS; Fouquet et al., 2010)—is common.

UM-SMS have been studied at MORs through geophysical, tectonic, petrographic, and geochemical tools (e.g., Andreani et al., 2014). Genetic models of UM-SMS generally involve downward seawater circulation along detachment or high-angle normal faults (McCaig et al., 2007; Andreani et al., 2014). At depth, seawater, along with possible subordinate amounts of magmatic fluids (Tivey, 2007), extracts heat from cooling magmatic bodies and reacts with the surrounding rocks. This process transforms it into a hot, acidic and reduced hydrothermal fluid that easily leaches metals from the surrounding rocks (Fouquet et al., 2010) and transports them as complexes (e.g., Alt and Shanks,

<footnote><sup>a</sup> Instituto Andaluz de Ciencias de la Tierra (IACT), CSIC, Armilla, Granada, Spain</footnote>

<footnote><sup>b</sup> ISTO, UMR 7327, Univ Orléans, CNRS, BRGM, OSUC, F-45071 Orléans, France</footnote>

<footnote><sup>c</sup> Institut de Mineralogy and Petrography, University of Innsbruck, A-6020 Innsbruck, Austria</footnote>

<footnote><sup>d</sup> University of California, Berkeley, CA, USA</footnote>

<footnote><sup>e</sup> Institut Terre et Environnement de Strasbourg, Université de Strasbourg, EOST, CNRS UMR 7063, F-67084 Strasbourg, France</footnote>

<footnote><sup>f</sup> Geosciences Rennes, Université de Rennes, CNRS UMR 6118, F-35000 Rennes, France</footnote>

<footnote>* This article is part of a Special issue entitled: 'DasOphiolite' published in Chemical Geology.</footnote>

<footnote>* Corresponding author at: Institut des Sciences de la Terre d'Orléans (ISTO), Orléans, France.
E-mail address: remi.coltat@univ-orleans.fr (R. Coltat).</footnote>
2003; Douville et al., 2002; Debret et al., 2018). On the way back to the seafloor, the mineralizing fluid mixes with seawater, which destabilizes metal complexes and triggers the precipitation of metal-bearing phases (sulfides and oxides), forming UM-SMS (Fouquet et al., 2010; Andreani et al., 2014). However, limited observations and drill cores in present-day seafloor environment prevent a comprehensive understanding of these hydrothermal systems.

A complementary approach is to investigate fossil ultramafic-hosted volcanogenic massive sulfide (UM-VMS) deposits that share structural, petrographic, and geochemical affinities with UM-SMS. These UM-VMS deposits are preserved in ophiolites worldwide (Patten et al., 2022 and references therein). Here, we investigate a fossil ultramafic-hosted hydrothermal mineralized system in the Platta nappe (Swiss Alps), associated with a Fe-Ca metasomatic mineral assemblage consisting of

<figure><img src="image_3.png" /><figcaption>**Fig. 1.** (A) Simplified map of the major paleogeographic units of the Western and Central Alps, and the Apennines. (B) Geological map of the Platta nappe (modified after Schaltegger et al., 2002). The white inset represents the field view shown in C. (C) Satellite view (Google Earth) of the mineralization in the Cotschen area. Mineralized and metasomatized serpentinized peridotites are in contact with mafic intrusions. A detailed view of the mineralized system and its relationships with mafic intrusions is given in A. 1.</figcaption></figure>
ilvaite, hydrogarnet, and diopside (Coltat et al., 2019b). Although this mineralogical assemblage is rather typical of skarn deposits (e.g., Sale-mink, 1985), several ultramafic-hosted occurrences have been documented in both fossil (Naslund et al., 1983; Sakai and Kuroda, 1983; Agata and Adachi, 1995; Bocchio et al., 2010) and present-day (Gaggero et al., 1997; Coltat et al., 2023) oceanic settings, and are locally associated with mineralization (Peltonen et al., 2008; Toffolo et al., 2017; Vezzoni et al., 2023; Dini et al., 2024). Thus, Fe-Ca silicates might serve as proxies for constraining the formation of UM-VMS deposits and their modern analogs. Unlike mafic-hosted mineralized systems, little is known about the deep-seated metasomatic processes that enable metal mobilization and precipitation in ultramafic-dominated environments. In this study, we seize the opportunity to better constrain the ore-forming processes through petrographic observations, in-situ geochemical analyses, and thermodynamic modeling of syn-mineralization Fe-Ca silicates in a fossil UM-VMS system exposed in the Swiss Alps (Platta nappe).

## 2. Geological context

The Platta nappe is located in the southeastern part of the Swiss Alps and belongs to the South Pennine Alpine units (Fig. 1A, B). It is interpreted as the fossil Ocean-Continent Transition (OCT) of the Adriatic margin, formed during the Jurassic opening of the Alpine Tethys Basin (Froitzheim and Manatschal, 1996; Müntener et al., 2004, 2010). The nappe is primarily composed of serpentinites, commonly overlain by mafic volcanics and/or post-rift sediments (Fig. 1B; Epin et al., 2019).

During the Jurassic, E-W-oriented mantle exhumation was accommodated along detachment faults (Fig. 1C; Froitzheim and Manatschal, 1996; Epin et al., 2019), and was accompanied by mafic magmatism (Desmurs et al., 2002). Zircons from syn-kinematic gabbroic rocks and plagiogranite yielded U-Pb ages of 161 ± 1 Ma (Schaltegger et al., 2002). Syn-tectonic fluid circulation is recorded by: i) serpentinization of mantle rocks (Desmurs et al., 2002; Müntener et al., 2010); ii) alteration of mafic rocks (e.g., rodingitization of mafic intrusions and epidotization-chloritization of mafic lava flows; Desmurs et al., 2001); iii) formation of mineralized systems (Coltat et al., 2019b; Coltat et al., 2021a); and iv) late carbonation of serpentinites and basalts (Coltat et al., 2019a, 2020).

During Alpine orogenic overprint, several deformation phases affected the Platta nappe. The Late Cretaceous D1 phase (Froitzheim et al., 1994) produced nappe stacking along thrusts with top-to-the-W to NW vergence, locally reactivating Jurassic extensional structures (Epin et al., 2019). A subsequent Upper Cretaceous extensional D2 phase formed top-to-the-E to SE normal faults (Froitzheim et al., 1994), which locally truncated D1 structures. During the Cenozoic, N-S-oriented shortening (D3 phase) resulted in the formation of large-scale E-W folds and N-vergent thrusts (Froitzheim et al., 1994). The northern part of the Platta nappe, which hosts the Marmorera-Cotschen Hydrothermal System (MCHS), was only affected by low-temperature prehnite-pummpellyite facies metamorphism (< 300–325 °C; Ferreiro Mählmann, 1995). Most primary oceanic-related mineralogical and geochemical signatures were preserved through this metamorphism (see discussion in Coltat et al., 2021b; Bernoulli and Weissert, 2021).

### 2.1. The Marmorera-Cotschen Hydrothermal System (MCHS)

The MCHS has been identified in the serpentinized footwall of the Jurassic detachment (Fig. 1B). Coltat et al. (2019b) showed that the overall system underwent some tilting during the Alpine orogeny, without significant internal deformation or Alpine metamorphism. This allows for restoring and studying the Jurassic system in its original geometry. The mineralization is spatially associated with mafic dykes and sills within the serpentinites, suggesting that lithological contacts acted as preferential fluid pathways (Coltat et al., 2019b). The Cu-Fe-Co-Zn-Ni mineralization mainly consists of chalcopyrite, pyrrhotite, magnetite,

isocubanite, and pentlandite, with minor sphalerite, pyrite, and late supergene phases. The mineralization was accompanied by Fe-Ca metasomatism (ilvaite, hydrogarnet, and diopside), either replacing ultramafic rocks or forming veins (Coltat et al., 2019b). The Sr isotope compositions of Fe-Ca silicates suggest interactions between a seawater-derived fluid and both mafic and ultramafic reservoirs (Hochscheid et al., 2022). Fe-Ca silicates are abundant in the deepest parts of the system and disappear towards the detachment (Coltat et al., 2019b).

The evolution of geochemical tracers (e.g., an increase in Se content and a decrease in the Co/Ni ratio) is consistent with the progressive upward mixing of hydrothermal fluid with seawater (Coltat et al., 2021a). In addition, at a given structural level, both the bulk rock and mineral Co/Ni ratios increase in the most metasomatized and/or mineralized zones, as observed in present-day UM-SMS (Marques et al., 2006) and other fossil UM-VMS (Toffolo et al., 2017; Dini et al., 2024). This suggests that the Co/Ni ratio could serve as a proxy for the intensity of fluid-rock interactions. Based on field and mineralogical evidence, the MCHS was interpreted as an analog to the root zone of present-day ultramafic-hosted black smokers, such as those formed at slow-spreading MORs (Coltat et al., 2019b; Coltat et al., 2021a).

## 3. Material and methods

### 3.1. Sampling strategy

We investigated six samples of metasomatized mantle rocks from the MCHS (Table 1, Fig. 2): one clinopyroxene-rich sample corresponding to a former pyroxenite or wehrlite layer (or a clinopyroxene-rich lherzolite; Fig. 2B), and five nearly fully serpentinized peridotites (likely lherzolites).

Fe-Ca silicates are either associated with mineralization (Fig. 2A-F, labeled “M” in Table 1) or occur in barren assemblages (Fig. 2G, H, named “B” in Table 1). They display a large variety of textures, regardless of their association with mineralization or not, as a pervasive replacement (P), vein infillings (V) and disseminated grains (D; Fig. 2). We performed in-situ analyses by EPMA and LA-ICP-MS on Fe-Ca silicates (ilvaite, hydro-andradite and clinopyroxene) from both mineralized and barren assemblages.

<caption>**Table 1**</caption>

Metasomatized mantle rocks from the Marmorera-Cotschen hydrothermal system investigated in this study.

<table><thead><tr><th rowspan="2">Sample</th><th rowspan="2">Lat.</th><th rowspan="2">Long.</th><th colspan="2">Mineral modal proportions (%)<sup>a/b</sup></th><th rowspan="2">Type of Fe-Ca silicates</th></tr><tr><th>Mantle silicates</th><th>Hydrothermal assemblages</th></tr></thead><tbody><tr><td>Cot18_40a</td><td>46°29′56″N</td><td>9°37′16″E</td><td>–</td><td><sup>b</sup>Ilv (38), Adr (15), Srp (40), Mag (6)</td><td>PM</td></tr><tr><td>Cot18_46b</td><td>46°29′56″N</td><td>9°37′16″E</td><td>–</td><td><sup>b</sup>Ilv (45), Adr (11), Srp (19), Mag (25)</td><td>PM</td></tr><tr><td>Mar18_39a</td><td>46°30′18″N</td><td>9°37′40″E</td><td>–</td><td><sup>a</sup>Ilv (20), Di (35), Srp (40), Mag (5)</td><td>VM</td></tr><tr><td>Cot16_61</td><td>46°29′54″N</td><td>9°36′58″E</td><td>–</td><td><sup>b</sup>Srp (59), Mag (41)</td><td>DM</td></tr><tr><td>Cot16_66</td><td>46°29′54″N</td><td>9°36′58″E</td><td><sup>a</sup>Cpx (15)</td><td><sup>a</sup>Ilv (15), Adr (15), Di (5), Srp (50)</td><td>PB</td></tr><tr><td>Cot18_57</td><td>46°29′54″N</td><td>9°36′54″E</td><td>–</td><td><sup>a</sup>Ilv (9), Di (15), Srp (76)</td><td>VB</td></tr></tbody></table>

P = pervasive, V = vein, D = disseminated, M = mineralized, B = barren, Ilv = ilvaite, Adr = hydro-andradite, Di = diopside, Cpx = clinopyroxene, Srp = serpentine, Mag = magnetite.

<sup>a</sup>Estimations made from petrographic observations (semiquantitative),

<sup>b</sup>Estimations acquired by Rietveld refinement from XRD analysis. Minor phases (< 5 %) not detected.
<figure><img src="image_4.png" /><figcaption>**Fig. 2.** Microphotographs of Fe-Ca silicates in mantle rocks of the Marmorera-Cotschen hydrothermal system. (A) Ilvaite and minor hydro-andradite associated with sulfides (chalcopyrite and sphalerite) and magnetite pervasively replacing mantle rock in sample Cot18_46b. (B) Ilvaite after olivine in a mineralized mantle rock. (C) Metasomatic banding composed of ilvaite bands alternating with Fe-serpentine and diopside in sample Cot18_46b. (D) Late hydro-andradite-Fe-serpentine vein cutting through ilvaite and magnetite in sample Cot18_40a. (E) Ilvaite-diopside veins with minor magnetite and chalcopyrite within Fe-serpentine in sample Mar18_39a. (F) Disseminated ilvaite associated with Cu-sulfides in sample Cot16_61. (G) Ilvaite-hydro-andradite-clinopyroxene at apparent textural equilibrium in the barren sample Cot16_66. Hydro-andradite replaces bastitized pyroxene. Late serpentine veins crosscut Fe-Ca silicates. (H) Ilvaite-diopside-Fe-serpentine vein within Fe-serpentine in sample Cot18_57. Srp = serpentine, Fe-Srp = Fe-serpentine, Spl = spinel, Ilv = ilvaite, Cpx = clinopyroxene, Di = diopside, Adr = hydro-andradite, Mag = magnetite, Ccp = chalcopyrite, Cv = covellite, Bn = bornite, Sp = sphalerite, Sulf = undetermined sulfides.</figcaption></figure>
## 3.2. EPMA analyses

EPMA analyses were carried out at the Microsonde Ouest Centre (Brest, France) using a Cameca SX100 Electron Probe Micro Analyzer to determine the major and minor element compositions of Fe-Ca silicates used as internal standard during LA-ICP-MS analyses. An accelerating voltage of 15 kV, a beam current of 15 nA, and a beam diameter of 5 µm were used. The count time for all analyzed elements was 10s. Standards (element) were albite (Na, Si), InP (P), forsterite (Mg), corundum (Al), orthoclase (K), wollastonite (Ca), synthetic manganese titanate (Mn, Ti), and pure iron (Fe).

## 3.3. LA-ICP-MS analyses

In situ trace element analyses by Laser Ablation-Inductively Coupled Plasma-Mass Spectrometry (LA-ICP-MS) were carried out at the Laboratory of Environment and Raw material Analysis (LERA), Karlsruhe Institute of Technology (Karlsruhe, Germany), using a Teledyne 193 nm Excimer Laser coupled to an Element XR ThermoFisher ICP-MS. Analyses of Fe-Ca silicates were performed with a spot size of 35 µm, a laser frequency of 10 Hz, a fluence of 5 J cm⁻², and helium and nitrogen flow rates of 0.3 L min⁻¹ and 10 mL min⁻¹, respectively. The Si contents from EPMA analyses were used for internal standard calibration. The following isotopes were measured: 7Li, 29Si, 31P, 32S, 45Sc, 49Ti, 51V, 53Cr, 55Mn, 57Fe, 59Co, 60Ni, 63Cu, 66Zn, 75As, 85Rb, 88Sr, 89Y, 90Zr, 93Nb, 95Mo, 107Ag, 118Sn, 121Sb, 137Ba, 139La, 140Ce, 141Pr, 146Nd, 147Sm, 151Eu, 157Gd, 159Tb, 161Dy, 165Ho, 167Er, 169Tm, 172Yb, 175Lu, 178Hf, 181Ta, 197Au, 205Tl, 208Pb, 209Bi, 232Th, 238U. Calibration and data quality control were performed using NIST-612 and NIST-614 glasses, as well as the USGS reference glasses BHVO-2G, BCR-2G, and BIR-1G. Data reduction was done using the Iolite software with the 3DRS plugin, considering NIST-612 as the reference material. This procedure yielded an accuracy better than 5 % for all elements. Accuracy, precision, and detection limits for the reference materials are detailed in Appendix A.2.

## 3.4. Thermodynamic modeling

Forward thermodynamic modeling was conducted in the CFMASHO system using elements as systems components (Ca–Fe–Mg–Al–Si–H₂–O₂). Phase relations were computed at three constant pressures (800, 2000 and 3000 bars) and using temperature (T) and chemical potential of O₂ (and therefore fugacity of O₂, fO₂) as thermodynamic variables. The other system components were constrained based on the measured composition of a representative sample Cot18_46b (Table 5, Coltat et al., 2021a). Gibbs energy minimization was computed using PerpleX version 7.1.5 (Connolly, 2005; Connolly, 2009) and the thermodynamic database of Holland and Powell (2011) (version TC-DS633). The considered solid solutions models were: olivine (Ol), orthopyroxene (Opx), clinopyroxene (Cpx), and garnet (Grt) from (Tomlinson and Holland, 2021), antigorite (Atg) that includes a ferric iron component from Eberhard et al. (2023), chlorite (Chl) from White et al. (2014) with the ferri-clinochlore endmember (fclin3) modified as indicated by Eberhard et al. (2023), and an ideal ferro-tremolite–tremolite solution. Hematite, magnetite, and iron are considered pure phases. The fluid was modeled as an electrolytic O-H fluid using a generic hybrid molecular fluid EoS solution model (Connolly, 1995). Thermodynamic data for aqueous species were taken from the latest version of the Deep Earth Water (DEW) model (Huang and Sverjensky, 2019).

The thermodynamic data for stoichiometric orthorhombic ilvaite, CaFe³⁺Fe²⁺Si₂O₈(OH), were incorporated in the database of Holland and Powell (2011, version TC-DS633) by fitting the original experimental data as follows. At room temperature the orthorhombic structure of ilvaite is metastable relative to the monoclinic one, therefore, the standard state molar volume of the orthorhombic structure (10.089

$J\bar{a}^{-1}$) was found by extrapolating the cell volume data (expressed as $p$) from Robie et al. (1988) from the 373–858 K range to 298.15 using their expression $p = p_0(1 + [T-273],$ where subscript refers to 0 K) and the preferred $p_0$ and the linear expansion coefficient ( ) for that temperature range (669.54 Å³ and $3.39 \times 10^{-5}$ T⁻¹ respectively). The temperature dependence of the thermal expansion and the bulk modulus at high temperatures were fitted to the expressions of Holland and Powell (1998) using the molar volume data as a function of temperature from Robie et al. (1988) and pressure from Finger and Hazen (1987), respectively. The lattice parameters at high pressure at room temperature obtained by the diamond-anvil cell are necessarily monoclinic (Finger and Hazen, 1987). Therefore, the derived bulk modulus is assumed to apply also to the modeled orthorhombic ilvaite. The constant pressure heat capacity was fitted to $C_p \rightarrow a + bT + cT^{-2} + dT^{-1/2}$ using the data from Robie et al. (1988). The reference Gibbs energy (HSC convention) was estimated based on the reaction ilvaite + O₂ = hedenbergite + magnetite + H₂O at 2000 bars and the FMQ and NNO buffers from Gustafson (1974) and HM and NNO buffers from Ghazi-Bayat et al. (1987). The resulting thermodynamic data for ilvaite are in agreement with those obtained by Delgado Martin and Soleri Gil (2010) that were compatible with the GEOe-CALC software (Berman et al., 1987). Note that the position of ilvaite relative in the T-logfO₂ space is incorrectly labeled in the Fig. 7 of Delgado Martin and Soleri Gil (2010) because ilvaite is the reduced phase relative to the magnetite + hedenbergite + H₂O assemblage in the CMSHO system, thus plotting at lower fO₂ relative to it.

Phase relations for a fixed bulk composition in the Fe–Mg–Cu–Zn–O₂ system were also computed using PerpleX and the SUPCRT92 thermodynamic database (Johnson et al., 1992). Isocubanite was added to this database (F. Klein, pers. com.) The bulk composition in this system is based on a modal proportion of magnetite (55 vol%), sphalerite (40 vol %), isocubanite (4 vol%) and pyrrhotite (1 vol%).

# 4. Results

## 4.1. Petrographic description

In mineralized samples (Fig. 2A–F), Fe-Ca silicates occur as pervasive replacement, veining, and disseminated grains (Table 1). Pervasive Fe-Ca silicate assemblages mainly consist of ilvaite, minor hydro-andradite, and diopside (Table 1, Fig. 2A–D). Hydro-andradite forms coronas around sphalerite–magnetite–chalcopyrite blebs, which are themselves surrounded by ilvaite (Fig. 2A), although ilvaite may also occur in direct contact with sulfides and magnetite. Hydro-andradite is also associated with a green, pleochroic, fan-shaped silicate in late veins crosscutting ilvaite (Fig. 2D). Ilvaite appears to replace olivine (Fig. 2B) and is associated with the formation of Fe-rich antigorite or greenalite (hereafter referred to as Fe-serpentine; Fig. 2B, C, A.1), locally forming banded textures with ilvaite (Fig. 2C). Similar banding occurs between ilvaite and magnetite, with both phases associated with minor chalcopyrite. Fe-serpentine in bands is richer in Fe (average Mg# = 0.32; i.e., molar Mg/(Mg + Fe); $n = 3$) than serpentine in other textural positions (average Mg# = 0.82, $n = 4$, SEM-EDX measurements). In addition to pervasive replacement, ilvaite, diopside, and minor magnetite and chalcopyrite form a network of veins (Fig. 2E, Table 1), locally associated with diopside-rich pockets. Ilvaite also surrounds magnetite after spinel (Fig. 2E). Finally, disseminated grains of ilvaite, associated with Cu-sulfides, also occur at contacts with magnetite (Fig. 2F).

In non-mineralized samples (Fig. 2G–H), Fe-Ca silicates occur as pervasive replacement and veins (Table 1). In one sample (Cot16_66), pervasive Fe-Ca silicate assemblages consist of hydro-andradite, ilvaite, and clinopyroxene. SEM-EDX analyses reveal two types of clinopyroxene: a predominant one enriched in Al, Cr, Na, and Ti, often rimmed by a secondary clinopyroxene of diopside–hedenbergite composition. This secondary clinopyroxene may also crystallize as newly-formed euhedral grains associated with other Fe-Ca silicates and serpentine.
andradite replaces bastitized pyroxene (Fig. 2G), whereas the replacement of olivine by ilvaite remains uncertain. The rock is crosscut by late serpentine veins (Fig. 2G). SEM-EDX analyses show that serpentine in bastite is more Fe-rich than serpentine in late veins. Sample Cot18_57 consists of large veins (up to 1 cm wide) of ilvaite, diopside, and Fe-serpentine cutting through a serpentinite made of Fe-serpentine (Fig. 2H).

In the following sections, the terminology of the different types of grains is simplified. For a given Fe-Ca silicate (e.g., ilvaite, hydro-andradite and clinopyroxene), we first refer to its textural habit using the abbreviation P for grains in pervasive replacement, V for vein-infilling grains, D for disseminated grains. Additionally, we will use M and B to refer to the association of Fe-Ca silicates with the mineralization or not, respectively. For instance, “PM ilvaite” refers to ilvaite that pervasively replaces mineralized mantle rocks, while “VB diopside” refers to diopside filling veins that cut through the barren mantle rock (Table 1). “PB augitic clinopyroxene” refers to clinopyroxene rimmed by secondary clinopyroxene in the barren metasomatized ultramafic rocks.

## 4.2. Major element compositions of Fe-Ca silicates

Major element compositions of Fe-Ca silicates are reported in Tables 2 to 4 and in Fig. 3. Ilvaite ($CaFe_2^{2+}Fe^{3+}Si_2O_8(OH)$) displays compositions close to the ideal end-member composition (Table 2). It has an average $FeO_t$ content ranging from 46.4 to 50.5 wt%, the lowest values being measured in PM and PB ilvaite. The latter displays high MgO and MnO contents with average values up to 3 and 1.8 wt%, respectively (Table 1, Fig. 3A, B). Conversely, VM and VB ilvaite display low MgO and MnO contents (Fig. 3A, B). Ilvaite of the MCHS has MgO and MnO contents comparable to the ones measured in ultramafic-hosted ilvaite (Fig. 3B). Their MgO content is higher (0.83 to 3.01 wt%) than in skarn- and mafic-hosted ilvaite (Fig. 3A). The $Cr_2O_3$ content is low (0.01 to 0.09 wt%) in all samples (the highest values being measured in VB ilvaite). The CaO content is homogenous, ranging between 14.0 and 14.4 wt%. The $TiO_2$ content is higher in non-mineralized samples (up to 0.37 wt% in average) than in mineralized samples (up to 0.02 wt% in average).

Hydrogarnet is almost pure hydro-andradite ($Ca_3Fe_2^{3+}(SiO_4)_3$.

<table><caption>**Table 2**
Major element compositions of ilvaite (in wt%).</caption><thead><tr><th>Sample</th><th colspan="2">Cot16_66</th><th colspan="2">Cot18_57</th><th colspan="2">Cot16_61</th><th colspan="2">Mar18_39</th><th colspan="2">Cot18_40a</th><th colspan="2">Cot18_46b</th><th>Theoretical</th></tr><tr><th>n</th><th>27</th><th>28</th><th>27</th><th>28</th><th>27</th><th>28</th><th>27</th><th>28</th><th>27</th><th>28</th><th>27</th><th>28</th><th></th></tr><tr><th></th><th>Avg</th><th>2σ</th><th>Avg</th><th>2σ</th><th>Avg</th><th>2σ</th><th>Avg</th><th>2σ</th><th>Avg</th><th>2σ</th><th>Avg</th><th>2σ</th><th></th></tr></thead><tbody><tr><td>SiO<sub>2</sub></td><td>30.00</td><td>1.05</td><td>29.61</td><td>0.57</td><td>29.52</td><td>0.55</td><td>29.51</td><td>0.38</td><td>29.94</td><td>0.58</td><td>29.88</td><td>0.67</td><td>29.40</td></tr><tr><td>Cr<sub>2</sub>O<sub>3</sub></td><td>0.03</td><td>0.12</td><td>0.15</td><td>0.69</td><td>0.09</td><td>0.25</td><td>0.04</td><td>0.06</td><td>0.01</td><td>0.04</td><td>bdl</td><td>bdl</td><td>-</td></tr><tr><td>TiO<sub>2</sub></td><td>0.25</td><td>0.98</td><td>0.37</td><td>1.39</td><td>bdl</td><td>bdl</td><td>0.02</td><td>0.04</td><td>bdl</td><td>bdl</td><td>0.01</td><td>0.06</td><td>-</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>0.08</td><td>0.07</td><td>0.18</td><td>0.36</td><td>0.11</td><td>0.29</td><td>0.07</td><td>0.13</td><td>0.09</td><td>0.06</td><td>0.12</td><td>0.41</td><td>-</td></tr><tr><td>*FeO<sub>(tot)</sub></td><td>46.43</td><td>1.28</td><td>50.47</td><td>2.56</td><td>50.13</td><td>0.62</td><td>50.30</td><td>0.66</td><td>47.75</td><td>1.63</td><td>48.59</td><td>1.58</td><td>-</td></tr><tr><td>**FeO</td><td>30.96</td><td></td><td>33.65</td><td></td><td>33.42</td><td></td><td>33.53</td><td></td><td>31.83</td><td></td><td>32.39</td><td></td><td>35.15</td></tr><tr><td>**Fe<sub>2</sub>O<sub>3</sub></td><td>17.21</td><td></td><td>18.70</td><td></td><td>18.57</td><td></td><td>18.64</td><td></td><td>17.69</td><td></td><td>18.00</td><td></td><td>19.53</td></tr><tr><td>MnO</td><td>1.77</td><td>0.24</td><td>0.43</td><td>0.09</td><td>1.29</td><td>0.21</td><td>0.60</td><td>0.12</td><td>1.39</td><td>0.38</td><td>1.45</td><td>0.66</td><td>-</td></tr><tr><td>MgO</td><td>2.75</td><td>0.65</td><td>0.83</td><td>0.70</td><td>1.44</td><td>0.29</td><td>0.98</td><td>0.38</td><td>3.01</td><td>1.31</td><td>1.98</td><td>1.07</td><td>-</td></tr><tr><td>CaO</td><td>14.40</td><td>0.73</td><td>14.11</td><td>0.50</td><td>13.93</td><td>1.23</td><td>13.99</td><td>0.31</td><td>14.05</td><td>0.52</td><td>14.05</td><td>0.82</td><td>13.72</td></tr><tr><td>Total</td><td>95.74</td><td></td><td>96.18</td><td></td><td>96.50</td><td></td><td>95.54</td><td></td><td>96.28</td><td></td><td>96.12</td><td></td><td>97.8</td></tr><tr><td colspan="14">Cations per 8.5 oxygens</td></tr><tr><td>Si</td><td>2.01</td><td></td><td>2.00</td><td></td><td>1.98</td><td></td><td>2.00</td><td></td><td>2.00</td><td></td><td>2.00</td><td></td><td>2.00</td></tr><tr><td>Cr</td><td>0.00</td><td></td><td>0.01</td><td></td><td>0.00</td><td></td><td>0.00</td><td></td><td>0.00</td><td></td><td>0.00</td><td></td><td>-</td></tr><tr><td>Ti</td><td>0.01</td><td></td><td>0.02</td><td></td><td>0.00</td><td></td><td>0.00</td><td></td><td>0.00</td><td></td><td>0.00</td><td></td><td>-</td></tr><tr><td>Al</td><td>0.01</td><td></td><td>0.01</td><td></td><td>0.01</td><td></td><td>0.01</td><td></td><td>0.01</td><td></td><td>0.01</td><td></td><td>-</td></tr><tr><td>**Fe<sup>3+</sup></td><td>0.87</td><td></td><td>0.94</td><td></td><td>0.94</td><td></td><td>0.95</td><td></td><td>0.89</td><td></td><td>0.91</td><td></td><td>1.00</td></tr><tr><td>**Fe<sup>2+</sup></td><td>1.73</td><td></td><td>1.89</td><td></td><td>1.89</td><td></td><td>1.90</td><td></td><td>1.78</td><td></td><td>1.82</td><td></td><td>2.00</td></tr><tr><td>Mn</td><td>0.10</td><td></td><td>0.02</td><td></td><td>0.07</td><td></td><td>0.03</td><td></td><td>0.08</td><td></td><td>0.08</td><td></td><td>-</td></tr><tr><td>Mg</td><td>0.27</td><td></td><td>0.08</td><td></td><td>0.14</td><td></td><td>0.10</td><td></td><td>0.30</td><td></td><td>0.20</td><td></td><td>-</td></tr><tr><td>Ca</td><td>1.03</td><td></td><td>1.02</td><td></td><td>1.00</td><td></td><td>1.02</td><td></td><td>1.00</td><td></td><td>1.01</td><td></td><td>1.00</td></tr></tbody></table>

<table><caption>**Table 3**
Major element compositions of hydrogarnet (in wt%).</caption><thead><tr><th>Sample</th><th colspan="2">Cot16_66</th><th colspan="2">Cot18_46b</th></tr><tr><th>n</th><th>28</th><th>22</th><th>28</th><th>22</th></tr><tr><th></th><th>Avg</th><th>2σ</th><th>Avg</th><th>2σ</th></tr></thead><tbody><tr><td>SiO<sub>2</sub></td><td>35.68</td><td>0.64</td><td>35.35</td><td>0.59</td></tr><tr><td>Cr<sub>2</sub>O<sub>3</sub></td><td>0.25</td><td>0.47</td><td>bdl</td><td>bdl</td></tr><tr><td>TiO<sub>2</sub></td><td>0.07</td><td>0.26</td><td>bdl</td><td>bdl</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>1.56</td><td>0.24</td><td>0.98</td><td>1.59</td></tr><tr><td>*FeO</td><td>25.71</td><td>0.85</td><td>27.16</td><td>1.95</td></tr><tr><td>MnO</td><td>0.11</td><td>0.05</td><td>0.07</td><td>0.07</td></tr><tr><td>MgO</td><td>0.43</td><td>1.07</td><td>0.35</td><td>1.59</td></tr><tr><td>CaO</td><td>33.41</td><td>1.28</td><td>33.37</td><td>2.10</td></tr><tr><td>Total</td><td>97.24</td><td></td><td>97.31</td><td></td></tr><tr><td colspan="5">Cations per 12 oxygens</td></tr><tr><td>Si</td><td>2.99</td><td></td><td>2.97</td><td></td></tr><tr><td>Cr</td><td>0.02</td><td></td><td>0.00</td><td></td></tr><tr><td>Ti</td><td>0.00</td><td></td><td>0.00</td><td></td></tr><tr><td>Al</td><td>0.15</td><td></td><td>0.10</td><td></td></tr><tr><td>**Fe<sup>3+</sup></td><td>1.80</td><td></td><td>1.91</td><td></td></tr><tr><td>Mn</td><td>0.01</td><td></td><td>0.01</td><td></td></tr><tr><td>Mg</td><td>0.05</td><td></td><td>0.04</td><td></td></tr><tr><td>Ca</td><td>3.00</td><td></td><td>3.00</td><td></td></tr><tr><td colspan="5">Modal abundances</td></tr><tr><td>Adr</td><td>0.91</td><td></td><td>0.95</td><td></td></tr><tr><td>Grs</td><td>0.08</td><td></td><td>0.05</td><td></td></tr><tr><td>Uv</td><td>0.01</td><td></td><td>0.00</td><td></td></tr></tbody></table>

$x(OH)_{4x}$; Adr > 0.91) with a small hydro-grossular component for PB grains (average $Al_2O_3$ = 1.6 wt%, Table 3). The $Cr_2O_3$ and $TiO_2$ contents are low, as low as 0.3 wt% and 0.1 wt% in average, respectively. The highest values are found in PB grains. Finally, the MnO and CaO contents are quite homogenous at about 0.1 wt% and 33.4 wt% in average, respectively.

<footnote>*Total Fe as FeO. “-” = not measured, bdl = below detection limits. The low totals are due to the presence of non-measured hydroxide ions. **Fe<sup>2+</sup>/Fe<sup>3+</sup> assumed to be 2/1.</footnote>

<footnote>*Total Fe as FeO. The low totals are due to the presence of hydroxide ions. Bdl = below detection limits.</footnote>

<footnote>**All Fe is assumed to be Fe<sup>3+</sup>.</footnote>
<table><caption>**Table 4**
Major element compositions of clinopyroxene (in wt%).</caption><thead><tr><th rowspan="2">Sample</th><th colspan="2">Cot16_66 (I)</th><th colspan="2">Cot16_66 (II)</th><th colspan="2">Cot18_57</th></tr><tr><th>15</th><th>4</th><th>15</th><th>4</th><th>15</th><th>4</th></tr><tr><th>n</th><th>Avg</th><th>2σ</th><th>Avg</th><th>2σ</th><th>Avg</th><th>2σ</th></tr></thead><tbody><tr><td>SiO<sub>2</sub></td><td>51.15</td><td>0.68</td><td>54.19</td><td>0.90</td><td>51.86</td><td>1.42</td></tr><tr><td>Cr<sub>2</sub>O<sub>3</sub></td><td>0.86</td><td>0.10</td><td>0.02</td><td>0.04</td><td>bdl</td><td>bdl</td></tr><tr><td>TiO<sub>2</sub></td><td>0.83</td><td>0.16</td><td>0.01</td><td>0.03</td><td>0.02</td><td>0.05</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>4.64</td><td>0.66</td><td>0.12</td><td>0.13</td><td>0.28</td><td>0.39</td></tr><tr><td>FeO*</td><td>3.07</td><td>0.48</td><td>3.50</td><td>1.79</td><td>14.12</td><td>5.00</td></tr><tr><td>MnO</td><td>0.10</td><td>0.06</td><td>0.18</td><td>0.06</td><td>0.33</td><td>0.68</td></tr><tr><td>MgO</td><td>16.04</td><td>1.27</td><td>16.78</td><td>0.75</td><td>9.49</td><td>3.96</td></tr><tr><td>CaO</td><td>22.65</td><td>1.79</td><td>25.31</td><td>0.76</td><td>24.41</td><td>1.15</td></tr><tr><td>Na<sub>2</sub>O</td><td>0.60</td><td>0.08</td><td>0.02</td><td>0.04</td><td>0.07</td><td>0.14</td></tr><tr><td>Total</td><td>99.97</td><td></td><td>100.13</td><td></td><td>100.59</td><td></td></tr></tbody><tbody><tr><td colspan="7">Cations per 6 oxygens</td></tr><tr><td>Si</td><td>1.87</td><td></td><td>1.98</td><td></td><td>1.99</td><td></td></tr><tr><td>Cr</td><td>0.02</td><td></td><td>0.00</td><td></td><td>0.00</td><td></td></tr><tr><td>Ti</td><td>0.02</td><td></td><td>0.00</td><td></td><td>0.00</td><td></td></tr><tr><td>Al</td><td>0.20</td><td></td><td>0.01</td><td></td><td>0.01</td><td></td></tr><tr><td>**Fe<sup>2+</sup></td><td>0.09</td><td></td><td>0.11</td><td></td><td>0.45</td><td></td></tr><tr><td>Mn</td><td>0.00</td><td></td><td>0.01</td><td></td><td>0.01</td><td></td></tr><tr><td>Mg</td><td>0.87</td><td></td><td>0.92</td><td></td><td>0.54</td><td></td></tr><tr><td>Ca</td><td>0.89</td><td></td><td>0.99</td><td></td><td>1.00</td><td></td></tr><tr><td>Na</td><td>0.04</td><td></td><td>0.00</td><td></td><td>0.01</td><td></td></tr><tr><td>Mg#</td><td>0.90</td><td></td><td>0.90</td><td></td><td>0.54</td><td></td></tr><tr><td>Full range</td><td>0.89-0.91</td><td></td><td>0.87-0.92</td><td></td><td>0.38-0.69</td><td></td></tr></tbody></table>

Two types of clinopyroxene can be distinguished. First, most of the grains from the PB sample have high Cr<sub>2</sub>O<sub>3</sub>, Al<sub>2</sub>O<sub>3</sub>, TiO<sub>2</sub> and Na<sub>2</sub>O contents (type I in Table 4, Fig. 3D), overlapping the composition of mantle augitic clinopyroxene found in ultramafic rocks of the Platta nappe and other abyssal peridotites (Fig. 3D). Some other clinopyroxene present chemical affinities with the diopside-hedenbergite series and have low Cr<sub>2</sub>O<sub>3</sub>, Al<sub>2</sub>O<sub>3</sub>, TiO<sub>2</sub> and NaO contents (Table 4, Fig. 3D). In the PB sample, type II clinopyroxene makes overgrowths around type I clinopyroxene. The FeO content of diopside-like pyroxene is highly heterogeneous, even in a single vein (sample Cot18_57; full range between Dp<sub>38</sub> and Dp<sub>69</sub>), with high variability from grain to grain (2σ on the FeO content around 5 wt%).

### 4.3. Trace element concentrations of Fe-Ca silicates

#### 4.3.1. Ilvaite

The Ni and Co contents of ilvaite are variable (Fig. 4A). PB ilvaite displays the highest Ni and Co contents (up to 3416 and 186 ppm, respectively), excluding five grains that present very low contents without specific textural feature. This ilvaite type overlaps the concentrations of serpentine grains of the MCHS and lies on the serpentinization array, defined by the progressive loss of Co and Ni in mantle rocks during serpentinization; the Co/Ni ratio remains constant at about 0.05 (green arrow in Fig. 4A). The Co and Ni concentrations decrease in ilvaite from mineralized samples, with minimal values recorded in PM ilvaite (up to 1 and 0.3 ppm of Ni and Co, respectively), defining a mineralization trend (red arrow in Fig. 4A) along which the Co/Ni ratio increases in the most mineralized and altered samples.

In the mineralized samples, the Cu content decreases with the

<figure><img src="image_5.png" /><figcaption>**Fig. 3.** EPMA major element compositions of Fe-Ca silicates in mantle rocks of the Marmorera-Cotschen hydrothermal system. Data from the literature are plotted for comparison for ultramafic-hosted settings (Naslund et al., 1983; Sakai and Kuroda, 1983; Agata and Adachi, 1995; Gaggero et al., 1997; Bocchio et al., 2010), mafic-hosted settings (Wise and Moller, 1990; Ixer et al., 2005; Endo, 2017) and skarn-related occurrences (Pesquera and Velasco, 1986; Bocchio et al., 2010; Delgado Martin and Soleri Gil, 2010; Yu et al., 2019). Clinopyroxene of the Lower Platta is from Desmurs et al. (2002); Müntener et al. (2004), (Müntener et al., 2010). Clinopyroxene of the Mid-Atlantic Ridge Kane (MARK) area is from Stephens (1997); Werner et al. (1997).</figcaption></figure>

<footnote>*Total Fe as FeO. **All Fe is assumed to be Fe<sup>2+</sup>. Bdl = below detection limits.</footnote>
<figure><img src="image_6.png" /><figcaption>**Fig. 4.** LA-ICP-MS trace element compositions of ilvaite from the Marmorera-Cotschen hydrothermal system. Serpentine mesh and veins from the Marmorera-Cotschen area are displayed (light green dots, Srp MCHS; data from Hochscheid et al., 2024). The average geochemical composition of abyssal and fossil serpentinites (full green dot, data from Paulick et al., 2006; Delacour et al., 2008; Müntener et al., 2010; Kodolanyi et al., 2012; Barnes et al., 2014; Mazzeo et al., 2017; Cooperdock et al., 2018), and Rainbow hydrothermal fluid (full red dot, from Douville et al., 2002) are plotted. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>

increase of the Co/Ni ratio from 2087 ppm in DM ilvaite to 0.08 ppm in PM ilvaite (Fig. 4B). In the barren samples, VB ilvaite has heterogenous Cu content comprised between 0.07 and 789 ppm, regardless its Co/Ni ratio, whereas PB ilvaite has a quite homogenous Cu content (between 3.4 and 54 ppm), similar to the one of serpentinites (green point, Fig. 4B). The Zn content of ilvaite is homogenous in all grains between 60 and 100 ppm (Fig. 4C), and the La content decreases significantly with the Co/Ni ratio in mineralized samples (Fig. 4D).

Finally, Ti and V display a positive correlation (Fig. 4E). PM ilvaite

has the lowest Ti and V contents, whereas other ilvaite types have V and Ti contents that overlap with generally high V content in ilvaite associated with the mineralization.

### 4.3.2. *Hydro-andradite*

As for ilvaite, Ni and Co show a positive correlation in hydro-andradite (Fig. 5A). PB hydro-andradite displays high and heterogeneous Ni and Co contents, overlapping those of serpentine from the MCHS, and following the serpentinization array (Fig. 5A). In
<figure><img src="image_7.png" /><figcaption>Fig. 5. LA-ICP-MS trace element compositions of hydro-andradite from the Marmorera-Cotschen hydrothermal system. Full green and red dots, and light green dots as in Fig. 4. Garnet in skarn-related mineralized systems (beige envelop; data from Gaspar et al., 2008; Bocchio et al., 2010) and garnet in serpentinites (green envelop; data from Bocchio et al., 2010) are plotted for comparison. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>

mineralized samples, hydro-andradite displays low Co and Ni contents, lying on a second trend (red arrow on Fig. 5A) along which the Co/Ni ratio increases (up to 0.5). The Cu content varies by two orders of magnitude in both mineralized and barren samples (Fig. 5B). The Zn content increases by one order of magnitude with the Co/Ni ratio (Fig. 5C) while the La content decreases in PM hydro-andradite (Fig. 5D).

The Ti and V contents of PB hydro-andradite are comparable to those measured in serpentinites (Fig. 5E) and they decrease by two to three

orders of magnitude in mineralized samples.

### 4.3.3. Clinopyroxene

Co and Ni display a roughly positive correlation in clinopyroxene (Fig. 6A), with the highest values measured in PB augitic clinopyroxene (Fig. 6A, Co/Ni ratio of 0.06). The Ni and Co contents decrease in VM and VB diopside, increasing their Co/Ni (up to 2.77). There is no clear relationship between the Co/Ni ratio and the Cu, Zn, La and Sr contents in clinopyroxene (Fig. 6B–E). PB augitic clinopyroxene displays the
<figure><img src="image_8.png" /><figcaption>Fig. 6. LA-ICP-MS trace element compositions of clinopyroxene from the Marmorera-Cotschen hydrothermal system. Full green and red dots, and light green dots as in Fig. 4. In (E) the green box represents the average Sr content of the upper mantle (from Delacour et al., 2008). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>

lowest Cu, Zn and Sr contents and the highest La content. The Cu, Zn and Sr contents increase in other samples by three orders of magnitude for Cu and by one order of magnitude for Zn and Sr, while the La content decreases by one to two orders of magnitude.

Finally, Pb augitic clinopyroxene has the highest Ti and V values, while those values decrease up to two orders of magnitude in other samples (Fig. 6F).

4.4. *Rare Earth element concentrations of Fe-Ca silicates*

The chondrite-normalized REE patterns of ilvaite, garnet and diop-
side range between those of mafic and ultramafic rocks (Fig. 7). PM
ilvaite and PM hydro-andradite display the lowest values similar to the
serpentine of the MCHS (Fig. 7A, B). A strong positive Eu anomaly is
present in most silicates.

Ilvaite (Fig. 7A) presents a moderate U-shaped REE pattern with
limited LREE to HREE fractionation: LaN/YbN values are between 0.93
<figure><img src="image_9.png" /><figcaption>**Fig. 7.** LA-ICP-MS chondrite-normalized REE compositions of Fe-Ca silicates of the Marmorera-Cotschen hydrothermal system compared with literature data (chondrite compositions from McDonough and Sun, 1995). REE compositions of serpentine of the Marmorera-Cotschen hydrothermal system are from Hochscheid et al. (2024).</figcaption></figure>

and 1.28 for all samples but grains forming veins where LaN/YbN reaches 16.7. Generally, the REE concentration of ilvaite decreases with the intensity of mineralization.

PM hydro-andradite (Fig. 7B) has a roughly U-shaped REE pattern, whereas PB grains display a convex-up shape for LREE and a roughly flat pattern for MREE and HREE. Both have a weak LREE to HREE fractionation (LaN/YbN = 0.75 and 0.62 for PM and PB hydro-andradite, respectively).

PB augitic clinopyroxene (open squares in Fig. 7C) displays the highest REE concentrations, comparable to those of mantle clinopyroxene, with a LREE depletion (LaN/YbN = 0.19), a convex-up REE pattern shape, and a slight negative Eu anomaly. VM and VB diopside present a U-shaped REE pattern, with a moderate LREE to HREE fractionation (LaN/YbN = 1.14 and 1.76 in VM and VB grains, respectively).

### 4.5. Thermodynamic modeling in the CFMASHO system

Fig. 8 shows a T-log fO₂ pseudosection for the investigated sample (Cot18_46b). At 2000 bars, ilvaite has a rather restricted stability field and coexists with the other Fe-Ca silicates (andradite and diopside), antigorite and magnetite at temperatures of ~240–410 °C for a range in oxygen fugacity. The observed MCHS assemblage consisting of ilvaite-andradite-clinopyroxene-magnetite-antigorite-chlorite is stable at fluid-saturated conditions for a fO₂ ranging from FMQ –6 to FMQ +0 at the lowest temperature for this assemblage (~300 °C, Fig. 8C). Interestingly, this assemblage is predicted to be stable only along a univariant line between ~300 and ~340 °C, along which fO₂ decreases with increasing temperature (blue line dividing the fields 1 and 2 in Fig. 8). At

lower temperatures, andradite replaces clinopyroxene (field 1, Fig. 8A). This 300–340 °C range is lower than the 425 ± 75 °C estimated by Gustafson (1974) in the simpler Ca-Fe-Si-H-O₂ system at the same pressure. Ilvaite is stable at higher temperature when oxygen fugacity increases but its stability field is narrower than at low fO₂. Overall, ilvaite is mostly stable together with magnetite, below the FMQ buffer (Fig. 8C). It has a very restricted stability field with native iron buffer at low fO₂ where the stable assemblage consists of antigorite-ilvaite-clinopyroxene-native Fe-chlorite-fluid (field 4 in Fig. 8A). Interestingly, the model predicts that chlorite should be stable throughout, alongside ilvaite, although its presence has not been confirmed in the mineral assemblages. In our model, chlorite is predicted to be stable in small quantities (< 7 vol%) when the antigorite model from Eberhard et al. (2023) is used instead of the model from Padrón-Navarta et al. (2013). In the latter case, most of the Al is incorporated into antigorite, rendering chlorite unstable. However, the model by Eberhard et al. (2023) is preferred as it aligns better with the apparent XMg (where all iron is considered ferrous, Fig. 8). The occurrence of chlorite in Fig. 8 is, therefore, a consequence of the current limitations of the antigorite model in accounting for both the ferric iron and the alumina content. This may also explain the somewhat higher andradite component of garnet (> Adr₉₆) predicted by the model than the measured compositions (average of Adr₉₅). Noteworthy is that in the stability field of ilvaite, regardless of the stable mineral assemblage, the model predicts ilvaite can be stable at significantly low fO₂ where the fluid can contain a significant amount of molecular H₂ (0.02 mol fraction of H₂) and MgOaq (around 700 mmol.kg⁻¹) at the lowest fO₂.
<figure><img src="image_10.png" /><figcaption>**Fig. 8.** (A) Stability fields for a fixed bulk composition (pseudosection) for the pervasive mineralized sample Cot18_46b of the Marmorera-Cotschen hydrothermal system computed in the Ca-Fe-Mg-Al-Si-H₂ system, P = 2000 bars, as a function of fO₂ and T°C computed with Perple_X. Colored solid lines mark the appearance or disappearance of some key phases: clinopyroxene (orange solid line), andradite (blue solid line), olivine (green solid line) and chalcopyrite (red solid line). For reference (black solid and dotted lines) some relevant oxygen buffers are shown Hematite-Magnetite (HM), Fayalite-Magnetite-Quartz (FMQ), and Magnetite-Iron (MI). (B) Stability fields of ilvaite as a function of varying pressure (800, 2000 and 3000 bars). Note that the stability field of ilvaite shrinks with decreasing pressure towards lower oxygen fugacity without a strong incidence on temperatures. (C) Stability field of ilvaite-bearing assemblages (labeled “ilvaite”), selected buffers, and mineral reactions as a function of T°C and fO₂ relative to the FMQ buffer, P = 2000 bars. (D-F) Modal abundances of selected minerals predicted by the thermodynamic model: antigorite (green lines), ilvaite (grey lines), magnetite (black lines), garnet (blue lines) and clinopyroxene (orange lines). (G-I) Compositions of solid-solution predicted by the thermodynamic model for antigorite (green lines), garnet (blue lines) and clinopyroxene (orange lines). Atg = antigorite, Ilv = ilvaite, Cpx = clinopyroxene, Adr = andradite, Chl = chlorite, Mag = magnetite, Grt = garnet, Ol = olivine, Fe = native iron, Ccp = chalcopyrite, Bn = bornite, Trov = pyrrhotite. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>
# 5. Discussion

## 5.1. *T*, *P* and *fO*₂ constraints on the stability of Fe-Ca silicates

This study marks the first attempt to model mineral assemblages formed during Fe-Ca metasomatism in ultramafic rocks using thermodynamic modeling, incorporating fO₂ as an independent variable. We acknowledge that Fe-Ca metasomatism occurs in an inherently open system, making it a complex process that cannot be fully constrained by thermodynamic modeling alone due to the involvement of mass transfer in a broad sense. Nonetheless, we simplified the system to six main representative components (Ca-Fe-Mg-Al-Si-H) and employed both existing and newly compiled thermodynamic properties of relevant minerals (including those of ilvaite) to investigate the stability conditions of Fe-Ca silicates. In particular, these models extend existing experimental data in the CFSHO subsystem into ultramafic compositions by incorporating solid solution models that include the Mg-endmembers of all ferromagnesian silicates (notably antigorite and clinopyroxene), which had previously not been considered (cf. Gustafson, 1974). Our model thus provides a robust framework for discussing the (P)-T-fO₂ conditions associated with Fe-Ca metasomatism in ultramafic rocks, which typically lack barometers and thermometers.

At the Marmorera-Cotschen Hydrothermal System, no barometer is available to constrain the pressure conditions under which Fe-Ca silicates formed, and the deep geometry of the paleo-detachment fault remains poorly constrained. To assess the potential role of pressure, we investigated a discrete range of values (800, 2000, and 3000 bars; Fig. 8) which encompasses reasonable depths for fluid circulation associated with mantle exhumation at (ultra-)slow spreading centers, based on seismic data (Demartin et al., 2007; Tao et al., 2020). With decreasing pressure, the stability field of ilvaite slightly narrows in terms of temperature range and shifts towards lower oxygen fugacity. Additionally, the stability field of the ilvaite + clinopyroxene assemblage expands relative to that of ilvaite + andradite. Notably, the field where ilvaite coexists with olivine (field 3 in Fig. 8A) disappears at pressures below approximately 1100 bars. Ilvaite and magnetite replacing olivine have been reported in tectonically exhumed serpentinites from the Kane area along the Mid-Atlantic Ridge (Coltat et al., 2023), suggesting that this reaction occurs at depth during mantle exhumation, at least ~2.5 km below the seafloor, considering a 3000 m water column and a rock density of 2.5 g.cm⁻³. This estimate aligns well with the depth proposed by Coltat et al. (2023) for Fe-Ca metasomatism at Kane, based on P-wave velocity profiles (Canales et al., 2000). Taken together, these observations indicate that the Fe-Ca silicate assemblages observed at the MCHS can be stable across a broad pressure range.

Thermodynamic modeling suggests that Fe-Ca metasomatism and the formation of the observed mineral assemblages occurred at temperatures of 320–340 °C and fO₂ conditions between FMQ –6 and FMQ +0. At these conditions, the model predicts ~40 vol% of ilvaite, closely matching the value obtained by Rietveld refinement (45 %). Similarly, the modeled proportions of clinopyroxene (5–10 vol%) and its Mg# number (0.88–0.93) are consistent with petrographic observations and measured compositions (Mg# of 0.89). However, the model underestimates the abundance of andradite and magnetite relative to Rietveld data and overestimates the modal abundance of serpentine. Although the predicted Mg# of serpentine (0.80–0.85) is comparable to that measured in serpentine away from banding-like textures, the model fails to capture the composition of Fe-rich serpentine. These discrepancies likely arise from several factors: i) intrinsic model limitations—primarily related to Al content, as discussed in the results section, ii) local mineralogical and chemical heterogeneities between in situ mineral analyses and samples used for bulk rock geochemistry, and iii) the lack of apparent textural equilibrium between Fe-Ca silicates. Indeed, although ilvaite, andradite, clinopyroxene, magnetite, and Fe- serpentine coexist in the sample and are part of the same metasomatic event, textural evidence indicate that they did not necessarily form

under identical conditions. In sample Cot18_46b (Fig. 2A), ilvaite surrounds andradite and sulfides-oxide blebs, suggesting that ilvaite might have formed later. On the other hand, late hydro-andradite veins cut across ilvaite in sample Cot18_40a (Fig. 2D), indicating dynamic shifts in fluid chemistry during metasomatism, allowing different Fe-Ca silicates to crystallize in succession. The banding-like texture involving ilvaite, magnetite, Fe-serpentine, and diopside likely reflects chemical gradients during episodic fluid flow typical of metasomatic processes (e.g., Brady, 1977). Ilvaite and diopside are stable in barren veins (sample Cot18_57) and coexist with magnetite and chalcopyrite in VM (sample Mar18_39a), indicating formation temperature exceeding ~300 °C. The occurrence of ilvaite with bornite in sample Cot16_61 (Fig. 2F) further suggests that ilvaite can form at even higher temperatures, within the bornite stability field, up to 410 °C (Fig. 8). In PB sample (Cot16_66), ilvaite, clinopyroxene, and hydro-andradite appear texturally equilibrated, narrowing down the formation temperatures at ~300–340 °C.

To better constrain the genetic links between Fe-Ca metasomatism and mineralization, we conducted a thermodynamic model of oxide and sulfide phases in the Fe-Cu-Zn-S system, using the same T-logfO₂ conditions as for the CFMASH system (Fig. 9). In PM samples, the mineralization consists of magnetite, sphalerite, pyrrhotite, isocubanite, and chalcopyrite. The model successfully reproduces this assemblage, except for chalcopyrite, which is predicted to be unstable relative to isocubanite. The overall consistency between the predicted and observed mineral assemblages provides additional constraints on the temperatures and redox conditions prevailing during Fe-Ca metasomatism and associated ore deposition (Fig. 9).

Finally, we propose that Fe-Ca metasomatism at the MCHS occurred between 300 and 360 °C under overall reducing conditions, between FMQ and MI buffers. During this process, mineral growth induced local chemical gradients and transient shifts in stable mineral assemblages. However, the progressive replacement of olivine by ilvaite may have initiated at slightly higher temperatures (~370–410 °C at 2000 bars, Fig. 8). These temperatures and redox conditions are consistent with

<figure><img src="image_11.png" /><figcaption>**Fig. 9.** Stability fields of opaque phases computed in the Fe-Cu-Zn-S system, P = 2000 bars, as a function of log fO₂ and T°C and calculated using solution models in Perple_X. Note that the stability field of opaque mineral assemblages reported in pervasively mineralized samples (Cot18_40a and Cot18_46b) coincide with the stability field of Fe-Ca silicates, referred to as “ilvaite stability field”. Py = pyrite, Sp = sphalerite, Cb = isocubanite, Hem = hematite, Mag = magnetite, Po = pyrrhotite, Fe = native iron.</figcaption></figure>
previous studies that have predicted or calculated the formation of Fe-Ca
silicates in mantle rocks within a similar thermal range (~300–440 °C;
Agata and Adachi, 1995; Evans et al., 2017). Under such conditions, the
formation of Fe-Ca silicates is accompanied by strong Mg depletion, due
to Fe-Ca enrichment, which is characteristic of low fO₂ conditions in
ultramafic rocks, in agreement with the observations of Evans et al.
(2017) on altered peridotites from the Duluth Complex (Minnesota).

## 5.2. Temporal relationships between Fe-Ca metasomatism and serpentinization of mantle rocks

Several lines of evidence suggest that Fe-Ca metasomatism and mineralization deposition took place during an early stage of serpentinization at the MCHS.

Ilvaite, together with magnetite, locally replaces olivine, whereas Fe-
serpentine occurs at grain boundaries in the absence of magnetite
(Fig. 2B). A similar texture was documented at the Kane area along the
slow-spreading Mid-Atlantic Ridge (Coltat et al., 2023). Given that the
serpentinization rate of olivine is significantly slower than that of py-
roxene at high temperatures (> 350 °C; Martin and Fyfe, 1970; Allen and
Seyfried Jr., 2003), this texture was interpreted as the result of olivine
alteration at T > 350 °C. Our thermodynamic model supports this
interpretation, indicating that olivine replacement by ilvaite and
magnetite occurs at high temperature (~370–410 °C; Fig. 8). At 1100
bars, this reaction becomes restricted to a narrow stability field around
~335–370 °C. These conditions align with those proposed for ilaite-
-andradite–diopside–magnetite assemblages in the ultramafic rocks of
the Asama Igneous Complex (Agata and Adachi, 1995), where Fe-Ca
silicates are closely associated with ferritchromite and ilmenite. Using
the Stormer (1983) geothermometer, Agata and Adachi (1995) inferred
formation temperatures of 360–440 °C under fO₂ conditions between
10⁻³¹ and 10⁻³⁸ bars. At the MCHS, Fe-serpentine is restricted to met-
asomatized and mineralized zones (Dietrich, 1972; Coltat et al., 2019b),
contrasting with the lizardite and chrysotile assemblages typical of
serpentinites in the Platta nappe (Dietrich, 1972). The co-occurrence of
ilvaite and Fe-serpentine at the MCHS therefore suggests that olivine
hydration occurred during Fe-Ca metasomatism. Furthermore, ilvaite
and diopside veins crosscutting Fe-serpentine (Fig. 2H), and the pres-
ence of Fe-serpentine clasts embedded in massive sulfides (Coltat et al.,
2019b), indicate that part of the Fe-Ca metasomatic and mineralizing
event postdates Fe-serpentine formation.

The oxygen isotope compositions of serpentinites from the Platta
nappe range between 6 and 10 ‰ (Burkhard and O'Neil, 1988; Früh-
Green et al., 1990). Using the oxygen isotope fractionation factors be-
tween serpentine and H₂O provided by Saccocia et al. (2009), and
assuming an initial seawater composition of δ¹⁸O = 0 ‰ with suffi-
ciently high fluid/rock (F/R) ratios to preserve the fluid's isotopic
signature, this corresponds to serpentinization temperatures of approx-
imately 150–200 °C. These values are significantly lower than the
300–360 °C temperature range estimated for Fe-Ca metasomatism based
on our thermodynamic modeling. Furthermore, in the PB sample, the Fe-
Ca silicate assemblage (ilvaite–hydrogarnet–diopside) is crosscut by
serpentine veins (Fig. 2G), indicating that serpentinization locally
postdates the formation of Fe-Ca silicates. However, it remains unclear
whether these late serpentine veins formed during the same serpenti-
nization event that produced the widespread chrysotile and lizardite
assemblage characteristic of the Platta nappe. This ambiguity notwith-
standing, the data suggest that the main serpentinization event occurred
after the development of Fe-Ca metasomatism and the associated for-
mation of Fe-serpentine.

Altogether, petrographic and geochemical evidence suggests that Fe-
Ca metasomatism took place during an early stage of serpentinization, at
least at a point when olivine was still stable relative to serpentine, while
pyroxene had already begun to alter. This early alteration stage is
notably supported by the replacement of bastitized pyroxene by hydro-
andradite (Fig. 2G). The main pervasive serpentinization event likely

occurred later, during the progressive cooling of the system associated
with mantle exhumation and unroofing. While we cannot exclude that
the shallow portions of the mantle, close to the seafloor, were already
serpentinized at that time, the depth interval where Fe-Ca meta-
somatism took place probably still recorded only incipient serpenti-
nization. This highlights the spatial and temporal complexity of alteration
processes accompanying tectonic mantle exhumation.

## 5.3. Is Fe-Ca metasomatism of ultramafic rocks genetically linked with rodingitization?

Ca-rich and Si-poor metasomatism affecting mafic intrusions within
mantle rocks referred to as rodingitization (Coleman, 1967; O'Hanley
et al., 1992). This process has been extensively documented in various
geological settings (e.g., Bach and Klein, 2009 and references therein).
Rodingites are interpreted to form through the interaction between
mafic (or felsic) intrusions and Ca-rich, Si-poor, alkaline fluids gener-
ated during the serpentinization of ultramafic rocks (Janecky and Sey-
fried Jr, 1986; O'Hanley et al., 1992). As a result, rodingitization is
typically constrained to thermal window in which serpentinization
proceeds, generally below 400 °C (Bach and Klein, 2009). Rodingitiza-
tion reactions lead to the development of secondary mineralogical as-
semblages in the mafic protolith, including clinopyroxene, grossular,
vesuvianite, prehnite, epidote, and chlorite (e.g., Frost, 1975; O'Hanley
et al., 1992; Laborda-López et al., 2018), reflecting a substantial Ca
input and Si depletion (Li et al., 2004; Laborda-López et al., 2020).

At the MCHS, Fe-Ca metasomatic assemblages developed within
mantle rocks and show mineralogical similarities with those produced
during rodingitization, except the presence of ilvaite, which has never
been reported in rodingites to our knowledge. This arises the question of
a possible genetic relationship between these two processes. However,
several lines of evidence indicate that Fe-Ca metasomatism is distinct
from typical rodingitization.

First, mafic intrusions found in contact with the mineralization and Fe-Ca silicates do not exhibit petrographic or geochemical characteristics typical of rodingites. While rodingites may show compositional variability due to multistage hydrothermal alterations (e.g., O'Hanley et al., 1992; Li et al., 2004; Laborda-López et al., 2018), the mafic intrusions at the MCHS are consistently depleted in CaO (0.4–7.3 wt%) and enriched in FeO<sub>tot</sub> (11.2–26.9 wt%, supplementary material A. 3) relative to typical rodingites (Fig. 10A–B). Likewise, Fe-Ca silicates-bearing mantle rocks display distinct chemical signatures, with low Ca and high Fe contents, that sharply contrast with those of rodingites (Fig. 10A–B). Additionally, the mafic intrusions consist of chlorite, pumpellyite, augite, titanite, and apatite (A. 1D), whereas common rodingite minerals such as garnet, diopside, and vesuvianite are absent (Bach and Klein, 2009).

Second, rodingites typically form where hydrothermal fluids have
not interacted with crustal mafic rocks and have instead been buffered
by ultramafic rocks, producing alkaline fluids such as those venting at
the Lost City hydrothermal field (e.g., Bach and Klein, 2009; Seyfried Jr.
et al., 2015). Such fluids are inefficient at mobilizing metals (Palandri
and Reed, 2004), in stark contrast with the extensive mineralization
developed contemporaneously with Fe-Ca metasomatism at the MCHS
(Coltat et al., 2019b). Massive sulfide at the MCHS contain up to 27 wt%
Cu, 36 wt% Fe, 0.28 wt% Zn, and 0.16 wt% Co (Coltat et al., 2021a).
Moreover, both the Fe-Ca silicates-bearing mantle rocks and adjacent
mafic intrusions display elevated Cu concentrations (up to 0.9 wt% Cu,
Fig. 10C), indicating interaction with a metal-rich, ore-forming fluid. In
contrast, rodingites are systematically devoid of Cu enrichment. Effi-
cient metal mobilization is favored by hot (> 350 °C), reduced, and
acidic fluid (e.g., Reed and Palandri, 2006), akin to those venting at the
Rainbow hydrothermal field (Seyfried Jr. et al., 2011; Evans et al.,
2024). Such fluids, not genetically linked to rodingitization, likely drove
both mineralization and Fe-Ca metasomatism at the MCHS. Notably, in
one diabase sample, the Cu content increases by an order of magnitude
<figure><img src="image_12.png" /><figcaption>**Fig. 10.** Major and trace element compositions of Fe-Ca silicates-bearing mantle rocks (data from Coltat et al., 2021a) and mafic intrusive rocks (this study, A. 3) of the Marmorera-Cotschen hydrothermal system. The compositions are plotted against worldwide orogenic and oceanic rodingites (data from Coleman, 1967; Honorez and Kirst, 1975; Schandl et al., 1989; O'Hanley et al., 1992; Li et al., 2004; Li et al., 2007; Austrheim and Prestvik, 2008; Duan et al., 2021; Rogkala et al., 2022).</figcaption></figure>

from the core to the rim, suggesting that fluid flow was focused along the diabase-mantle rocks interface, enhancing reaction with the ore-forming fluid.

Third, the scale of alteration observed at the MCHS further argues against a genetic link with rodingitization of mafic intrusions. Mafic dikes are typically less than 1 m wide, whereas the Fe-Ca silicate alteration zone exceeds 100 m in width (Fig. 1C, supplementary material A. 1 A–B). It is therefore unlikely that fluids released from the rodingitization of such small mafic intrusions could account for the chemical budget required to form the extensive Fe-Ca silicate assemblages.

Finally, thermodynamic modeling of Fe-Ca silicates at the MCHS points to formation temperatures above 300 °C, significantly higher than typical rodingitization temperatures observed in oceanic settings, which generally range between 200 and 300 °C (Bach and Klein, 2009; Fig. 11). These temperatures are consistent with those estimated for ilvaite formation in serpentinized wehrlites of the Asama igneous complex (Agata and Adachi, 1995). While we cannot fully exclude that high-temperature (> 300 °C) alteration of mafic intrusions may have contributed locally to the formation of Fe-Ca silicates, the overall evidence indicates that Fe-Ca metasomatism at the MCHS is not a direct analogue of rodingitization processes.

### 5.4. Geochemical record of Fe-Ca metasomatism

At first order, two types of Fe-Ca silicate assemblages can be distinguished at the MCHS: pervasive mineralized (PM) and pervasive barren (PB). In the PB sample Cot16_66, magmatic augitic clinopyroxene is preserved, whereas ilvaite, hydro-andradite, and diopside-hedenbergite are of clearly hydrothermal origin. Both ilvaite and hydro-andradite exhibit heterogeneous compositions and follow the serpentinization trend defined by Co and Ni variations in serpentine (Fig. 4A, 5A). This

trend reflects the chemical evolution of serpentine during ongoing serpentinization of mantle rocks, suggesting that ilvaite and hydro-andradite formed under conditions similar to those prevailing during high-temperature serpentinization (see also the first section of the discussion). A marked positive Eu anomaly in ilvaite and hydrogarnet further attests to fluid-rock interactions involving plagioclase (Douville et al., 1999; Andreani et al., 2014). This feature likely reflects interactions with plagioclase-bearing lherzolites, which are abundant in the Platta nappe (Müntener et al., 2010).

In mineralized samples (Table 1), ilvaite—and to a lesser extent hydro-andradite and diopside—shows a progressive increase in the Co/Ni ratio that correlates with mineralization (Figs. 4A, 5A, 6A). This trend cannot be solely explained by partitioning effect between Fe-Ca silicates and coeval metal-bearing phases, as a similar increase in Co/Ni is also observed in sulfides and oxides at the MCHS (Coltat et al., 2019b; Coltat et al., 2021b). Instead, it reflects distinctive geochemical characteristics of the ore-forming fluid, likely driven by variations in the F/R ratio and sources during Fe-Ca metasomatism. In DM sample, ilvaite exhibits normalized REE compositions higher than those of the Platta nappe ultramafic rocks, yet intermediate between ultramafic and mafic lithologies (Fig. 7). Concomitantly, the Co/Ni ratio rises (from 0.1 to 0.5, Fig. 4A), approaching values typical of mafic rocks in the Platta nappe (0.07–1.23, average 0.48; Desmurs et al., 2002). These data support a transition from a system dominated by ultramafic rock-fluid interaction to one involving both mafic and ultramafic rocks. Further increase in Co/Ni and decrease in REE concentrations in mineralized samples may be attributed to: i) progressive fluid evolution, ii) elevated F/R ratios during fluid-mafic rock interaction (leaching) and/or Fe-Ca silicate precipitation in ultramafic rocks (trapping), iii) a combination of these processes. The mineralogical evolution of Fe-Ca silicate assemblages, from ilvaite alone to ilvaite–diopside, and finally to ilvaite–hydro-
<figure><img src="image_13.png" /><figcaption>Fig. 11. Schematic diagram illustrating the mineralogical assemblages formed in mafic intrusions and their host mantle rock during fluid-rock interactions as a function of temperature. Serpentinization is enhanced at T < 400 °C (Klein et al., 2013). Fe-Ca metasomatism occurs in mantle rocks at temperatures >300 °C. Rodingitization of mafic intrusions is expected between 200 and 300 °C during the serpentinization of mantle rocks (Bach and Klein, 2009). Srp = serpentine, Mag = magnetite, Brc = brucite, Spl = spinel, Ol = olivine, Px = pyroxene, Ilv = ilvaite, Adr = (hydro-)andradite, Di = diopside, Chl = chlorite, Cpx = clinopyroxene, Ep = epidote, Tr = tremolite, Pl = plagioclase, Grt = garnet, Prh = prehnite.</figcaption></figure>

andradite–diopside in disseminated, veined, and pervasively altered samples, may reflect cooling of the fluid (Fig. 8) and/or increasing Ca/Si activity. There is no compelling reason to invoke distinct fluids resulting from variable interactions with mafic and ultramafic rocks at different F/R ratios. Rather, the observations suggest a single fluid with progressively increasing F/R ratios from disseminated to vein to pervasive alteration settings. Overall, the Ni, Co, and REE contents of Fe-Ca silicates in mineralized samples converge towards those measured in fluids from the Rainbow hydrothermal field (Figs. 4A, 5A, 6A, and 7, Douville et al., 2002), suggesting formation from a fluid similar to typical black smoker systems. The positive Eu anomaly in mineralized Fe-Ca silicates is interpreted as resulting either from plagioclase breakdown during fluid circulation in mafic rocks (Douville et al., 1999), and/or the redox state of the fluid (Sverjensky, 1984). Although precise partition coefficients for REEs and trace elements in this mineral-fluid system remain elusive, the consistent decrease in REE concentrations and increase in Co/Ni ratios associated with mineralization strongly support a scenario involving elevated F/R ratios and significant mafic lithology contribution during Fe-Ca metasomatism. We thus propose that mineralized Fe-Ca silicates formed in an open, fluid-dominated system from a reduced and acidic hydrothermal fluid derived from interactions with both mafic and ultramafic rocks. This interpretation is further supported by the Sr isotope composition of Fe-Ca silicates at the MCHS (Hochscheid et al., 2022).

At the MCHS, part of the geochemical variability of the Fe-Ca silicates may come from primary mantle heterogeneity (e.g., presence of pyroxenite layers, refractory versus fertile mantle domains, or depleted dunites), known to impart significant variations in element compositions in mantle rocks (e.g., Godard et al., 2008). However, given the

limited spatial scale over which this variability is observed, we suggest that it primarily reflects differences in fluid-rock interaction conditions rather than initial mantle composition. Trace element data indicate that PB Fe-Ca silicates likely formed in a system dominated by interaction with ultramafic rocks, whereas PM Fe-Ca silicates developed in a more open, fluid-dominated system, where the fluid had previously interacted with both mafic and ultramafic rocks. In this framework, Fe-Ca silicates record distinct stages of ultramafic lithology hydration: PB Fe-Ca silicates formed from a fluid buffered by ultramafic rocks alone, while Fe-Ca silicates in mineralized samples reflect fluid compositions modified by prior reactions with both mafic and ultramafic rocks.

### 5.5. Possible mineral reactions during Fe-Ca metasomatism

Petrographic observations indicate that olivine turns to ilvaite and Fe-serpentine (Fig. 2B, C). Similar occurrences of olivine ilvaitization have been documented in both fossil (Naslund et al., 1983) and present-day oceanic systems (Coltat et al., 2023). In an open system, this reaction may be written in the general form:

$$n_1[A] + n_2X + n_3H_2O = n_4[B] + n_5Y$$

where A and B represent primary mantle and metasomatic mineral phases, respectively. X corresponds to a cation consumed by the reaction (e.g., Fe, Ca, Mn), Y represents released species such as cations or gases (e.g., H₂). The stoichiometric coefficients n₁ to n₅ are critical as they determine both the elemental fluxes (i.e., supply vs. removal) and the associated volume changes (expansion or shrinkage). In open systems such as at the MCHS, it is particularly difficult to distinguish between elements sourced from the dissolution of primary minerals and those
introduced by external hydrothermal fluids, which themselves may have variable chemical compositions over time and space. An alternative approach involves estimating the reaction stoichiometry by considering the volume changes associated with the transformation. For example, experimental and natural observations show that serpentinization of mantle rocks typically results in a net volume expansion, up to 44 % of the original rock volume (e.g., Coleman and Keith, 1971; Andréani et al., 2007; Klein and Le Roux, 2020). In contrast, volumetric constraints for ilvaitization of mantle rocks are less well-documented. However, textural evidence from other settings (e.g., the presence of cracks in igneous plagioclase adjacent to altered olivine in the Duluth Complex) suggests that Fe-Ca metasomatism may also be accompanied by volume expansion (Evans et al., 2017). At the MCHS, the advanced degree of alteration and the absence of preserved primary minerals prevent direct identification of deformation features classically associated with volume increase (e.g., crack formation or mineral offsets; Coleman and Keith, 1971; Evans et al., 2017). Given the lack of direct volumetric constraints, we considered two end-member scenarios to balance the reactions: one assuming isovolumetric replacement and the other volume expansion during mineral transformations.

At the MCHS, Fe-serpentine has variable Fe content with an average Mg# of about 0.8. No remnant of pristine olivine was found in serpentinites (Müntener et al., 2010). For first order mass-balance calculations related to ilvaitization, we used olivine compositions (Mg# between 0.871 and 0.911) documented in serpentinized peridotites from the Central Alps ophiolites, such as the Malenco and Totalp units (Müntener et al., 2010). Using a representative Mg# of 0.9 for olivine, the quasi-isovolumetric and non-isovolumetric transformation of olivine into ilvaite and Fe-serpentine in reactions 1, 2, 3, 4, 1' and 2' (reactions 1 and 2, respectively) can be written:

$$ 
\begin{aligned}
5(\text{Mg}_{1.8}\text{Fe}_{0.2})\text{SiO}_4 [\text{Olivine}] &+ \text{Ca}^{2+} (\text{aq}) + 2.3\text{Fe}^{2+} (\text{aq}) + 0.1\text{Mn}^{2+} (\text{aq}) \\
&+ 6\text{H}^{+} (\text{aq}) \\
&= \text{Ca}(\text{Fe}_{2.7}\text{Mg}_{0.2}\text{Mn}_{0.1})\text{Si}_2\text{O}_8(\text{OH}) [\text{Ilvaite}] \\
&+ (\text{Mg}_{2.4}\text{Fe}_{0.6})\text{Si}_2\text{O}_5(\text{OH})_4 [\text{Fe} - \text{serpentine}] \\
&+ 6.4 \text{Mg}^{2+} (\text{aq}) + \text{SiO}_2 (\text{aq}) + 0.5\text{H}_2
\end{aligned} 
\quad (1) $$

$$ 
\begin{aligned}
4(\text{Mg}_{1.8}\text{Fe}_{0.2})\text{SiO}_4 [\text{Olivine}] &+ \text{Ca}^{2+} (\text{aq}) + 2.5\text{Fe}^{2+} (\text{aq}) + 0.1\text{Mn}^{2+} (\text{aq}) \\
&+ 2\text{H}_2\text{O} + 2\text{H}^{+} (\text{aq}) = \text{Ca}(\text{Fe}_{2.7}\text{Mg}_{0.2}\text{Mn}_{0.1})\text{Si}_2\text{O}_8(\text{OH}) [\text{Ilvaite}] \\
&+ (\text{Mg}_{2.4}\text{Fe}_{0.6})\text{Si}_2\text{O}_5(\text{OH})_4 [\text{Fe} - \text{serpentine}] + 4.6 \text{Mg}^{2+} (\text{aq}) + 0.5\text{H}_2
\end{aligned} 
\quad (2) $$

Reaction (1) implies a nearly constant volume ($\Delta V = -2$ %; Vm of reactants ~225 cm³/mol, while Vm of products ~220 cm³/mol), while reaction (2) reflects a net increase in volume (~22 %), a value lower than the 44 % observed during serpentinization (Klein and Le Roux, 2020).

Where magnetite is also formed during olivine breakdown (Fig. 2B), two additional isovolumetric and non-isovolumetric reactions, respectively reaction (3) and (4), can be written:

$$ 
\begin{aligned}
6(\text{Mg}_{1.8}\text{Fe}_{0.2})\text{SiO}_4 [\text{Olivine}] &+ \text{Ca}^{2+} (\text{aq}) + 5.3\text{Fe}^{2+} (\text{aq}) + 0.1\text{Mn}^{2+} (\text{aq}) \\
&+ 2\text{H}_2\text{O} (\text{aq}) + 4\text{H}^{+} (\text{aq}) = \text{Ca}(\text{Fe}_{2.7}\text{Mg}_{0.2}\text{Mn}_{0.1})\text{Si}_2\text{O}_8(\text{OH}) [\text{Ilvaite}] \\
&+ (\text{Mg}_{2.4}\text{Fe}_{0.6})\text{Si}_2\text{O}_5(\text{OH})_4 [\text{Fe} - \text{serpentine}] + \text{Fe}_3\text{O}_4 [\text{Magnetite}] \\
&+ 8.4 \text{Mg}^{2+} (\text{aq}) + 2\text{SiO}_2 (\text{aq}) + 1.5\text{H}_2
\end{aligned} 
\quad (3) $$

$$ 
\begin{aligned}
5(\text{Mg}_{1.8}\text{Fe}_{0.2})\text{SiO}_4 [\text{Olivine}] &+ \text{Ca}^{2+} (\text{aq}) + 5.3\text{Fe}^{2+} (\text{aq}) + 0.1\text{Mn}^{2+} (\text{aq}) \\
&+ 4\text{H}_2\text{O} (\text{aq}) = \text{Ca}(\text{Fe}_{2.7}\text{Mg}_{0.2}\text{Mn}_{0.1})\text{Si}_2\text{O}_8(\text{OH}) [\text{Ilvaite}] \\
&+ (\text{Mg}_{2.4}\text{Fe}_{0.6})\text{Si}_2\text{O}_5(\text{OH})_4 [\text{Fe} - \text{serpentine}] + \text{Fe}_3\text{O}_4 [\text{Magnetite}] \\
&+ 6.4 \text{Mg}^{2+} (\text{aq}) + \text{SiO}_2 (\text{aq}) + 1.5\text{H}_2
\end{aligned} 
\quad (4) $$

The reaction (4) is accompanied with a volume expansion of ~22 %.

In all reactions, Fe and Ca are externally sourced, whereas Mg and aqueous silica (reaction 1, 3 and 4) are released. The release of SiO₂ and Mg could have contributed to the stabilization of diopside (and locally tremolite and talc), which are occasionally observed in minor amount in the PM sample (Coltat et al., 2019b). Apart from reaction (4), these transformations also require an acidic hydrothermal fluid capable of transporting metals. Such fluids, especially when Cl-rich, are efficient at mobilizing base and transition metals (Reed and Palandri, 2006; Seyfried Jr. et al., 2011; Debret et al., 2018; Hodel et al., 2018). These fluids may originate from the interaction with mafic rocks (Alt and Shanks, 2003; Debret et al., 2018), or through high-temperature alteration (> 350 °C) of ultramafic rocks involving pyroxene breakdown and talc/ amphibole formation (e.g., reaction 8 in Allen and Seyfried Jr., 2003). Proton-releasing reactions such as plagioclase chloritization (Tivey, 2007) may also enhance fluid acidity. Such alteration has been documented in the mafic intrusions of the MCHS, where plagioclase is replaced by a fine-grained matrix of chlorite and pumpellyite (Fig. A.1D).

The original protolith composition plays a major role in mass balance calculations. Formation of ilvaite after Fe-rich olivine (e.g., MgO₂Fe₁.₈) has been documented in the ultramafic Skaergaard intrusion (Naslund et al., 1983), where external Ca was inferred to be supplied by the fluid while Fe was released during the reaction. In mantle rocks, the Mg# of olivine is influenced by magmatic processes and tends to decrease during melt-rock interaction (i.e., melt impregnation). This process is well documented at the Mid-Atlantic Ridge Kane area, where olivine in contact with mafic intrusions can display Mg# values as low as 0.63 (Cannat et al., 1997; Niida, 1997; Coltat et al., 2023). Although no primary silicates were preserved at the MCHS, the close spatial association between Fe-Ca silicates and mafic intrusions suggests that ilvaitization may have involved the replacement of Fe-rich olivine. Assuming an olivine Mg# of 0.7, the mass balance equations for olivine replacement can be rewritten as follows:

$$ 
\begin{aligned}
5(\text{Mg}_{1.4}\text{Fe}_{0.6})\text{SiO}_4 [\text{Olivine}] &+ \text{Ca}^{2+} (\text{aq}) + 0.3\text{Fe}^{2+} (\text{aq}) + 0.1\text{Mn}^{2+} (\text{aq}) \\
&+ 6\text{H}^{+} (\text{aq}) = \text{Ca}(\text{Fe}_{2.7}\text{Mg}_{0.2}\text{Mn}_{0.1})\text{Si}_2\text{O}_8(\text{OH}) [\text{Ilvaite}] \\
&+ (\text{Mg}_{2.4}\text{Fe}_{0.6})\text{Si}_2\text{O}_5(\text{OH})_4 [\text{Fe} - \text{serpentine}] + 4.4 \text{Mg}^{2+} (\text{aq}) \\
&+ \text{SiO}_2 (\text{aq}) + 0.5\text{H}_2
\end{aligned} 
\quad (1) $$

$$ 
\begin{aligned}
4(\text{Mg}_{1.4}\text{Fe}_{0.6})\text{SiO}_4 [\text{Olivine}] &+ \text{Ca}^{2+} (\text{aq}) + 0.9\text{Fe}^{2+} (\text{aq}) + 0.1\text{Mn}^{2+} (\text{aq}) \\
&+ 2\text{H}_2\text{O} + 2\text{H}^{+} (\text{aq}) = \text{Ca}(\text{Fe}_{2.7}\text{Mg}_{0.2}\text{Mn}_{0.1})\text{Si}_2\text{O}_8(\text{OH}) [\text{Ilvaite}] \\
&+ (\text{Mg}_{2.4}\text{Fe}_{0.6})\text{Si}_2\text{O}_5(\text{OH})_4 [\text{Fe} - \text{serpentine}] + 3 \text{Mg}^{2+} (\text{aq}) + 0.5\text{H}_2
\end{aligned} 
\quad (2) $$

As in the case of peridotite serpentinization at spreading centers (e. g., Klein et al., 2013) and the Iberia passive margin (Albers et al., 2021), ilvaitization of olivine involves the oxidation of ferrous iron, leading to the production of H₂-rich reduced fluids. This process is also predicted by our thermodynamic modeling. In addition, these reactions consume protons, which causes an increase in fluid pH. Such a pH increase is considered a key mechanism for destabilizing metal complexes and promoting metal precipitation as sulfides (Reed and Palandri, 2006). Altogether, these mineralogical reactions show that mafic-ultramafic rock interfaces act as pH and redox boundaries during fluid-rock interactions, dramatically affecting the chemical evolution of hydrothermal fluids.

## 5.6. Where do Fe and Ca come from?

The availability of Ca and Fe is a key parameter controlling the formation of Fe-Ca silicates during metasomatic alteration. In peridotites, the main Ca-bearing phases are clinopyroxene and plagioclase represent, both commonly found in ilherzolites of the lower Platta nappe (Müntener et al., 2010). In contrast, orthopyroxene has too little CaO (<
1.7 wt%; Müntener et al., 2010) to significantly influence the Ca budget. Therefore, part of the Ca mobilized during Fe-Ca metasomatism likely originates from the dissolution of mantle clinopyroxene and plagioclase (see also reaction 2 in the previous section). At 400 °C and 500 bars, clinopyroxene dissolution efficiently releases Ca into hydrothermal fluids (Allen and Seyfried Jr., 2003), a process invoked to explain the formation of andradite veins in the Feragen ultramafic body in Norway (Plümper et al., 2014).

In mafic rocks, calcium can also be mobilized through albitization of plagioclase and amphibolitization of clinopyroxene (Coltat et al., 2019b, 2023). These reactions are supported by the observed Ca depletion in altered mafic rocks near the MCHS, compared to those in the Platta nappe (A. 3; Desmurs et al., 2002). Furthermore, Sr isotopic compositions of Fe-Ca silicates at the MCHS indicate that the hydrothermal fluid interacted with both mafic and ultramafic lithologies, implying that Ca (as well as Sr) was sourced from all three reservoirs: mafic rocks, mantle rocks, and seawater-derived fluids (Hochscheid et al., 2022).

While iron is a major constituent of peridotite, typical serpentinization processes do not result in significant Fe mobility because Fe released during olivine breakdown is typically fixed in magnetite or sulfides forming along with serpentine. Under high-temperature conditions, as expected during Fe-Ca metasomatism, serpentinization of pyroxenite and lherzolite generates Si-rich and Fe-poor fluids relative to those derived from dunite (Allen and Seyfried Jr., 2003). Thus, the alteration of ultramafic rocks alone likely cannot account for the total Fe required during Fe-Ca metasomatism, unless large volumes of peridotite were altered, an option that cannot be entirely excluded.

Altered mafic intrusions adjacent to Fe-Ca silicates at the MCHS are not plausible Fe sources, as they are enriched in Fe relative to typical diabase (A. 3; Desmurs et al., 2002), indicating Fe addition during hydrothermal alteration. One mechanism that may explain localized Fe enrichment in mantle rocks is melt-rock interaction. This process is well documented at the Mid-Atlantic Ridge Kane area (e.g., Cannat et al., 1997; Niida, 1997; Coltat et al., 2023, 2025), where the extent of Fe enrichment correlates with the size of magmatic intrusions. At the MCHS, the metric to pluri-metric size of mafic bodies would imply Fe enrichment over several meters, consistent with the thickness of the observed metasomatic halo (Coltat et al., 2019b). However, textural and geochemical evidence for such melt-rock interaction may have been overprinted during the late crystallization of Fe-Ca silicates and subsequent serpentinization.

An alternative and likely more significant Fe source is provided by differentiated mafic rocks. Magmatic differentiation leads to Fe enrichment in residual melts, and Fe-Ti-(P) gabbros may act as major Fe reservoirs. Such gabbros are found in the Val da Natons and Fuorcla da Faller areas of the Platta nappe, containing up to 3.30 wt% FeO and 11.52 wt% $Fe_2O_3$ (Desmurs et al., 2002). These rocks underwent syntectonic retrograde alteration from amphibolite to greenschist facies conditions consistent with those inferred for Fe-Ca silicate formation. Furthermore, gabbros from the Val da Natons area are part of the same detachment system as the MCHS (Fig. 1B). Their alteration during exhumation likely produced Fe-rich fluids, which were later channeled along the permeable detachment surface at diabase-mantle rock contacts, precipitating the observed Fe-Ca silicates.

In summary, we propose that Ca at the Marmorera-Cotschen hydrothermal system was sourced from the alteration of mafic intrusions (e.g., plagioclase albitization and/or chloritization, clinopyroxene amphibolitization), the high-temperature dissolution of mantle clinopyroxene, and seawater-derived fluids. Iron was likely contributed by both Fe-enriched mantle rocks, possibly modified during melt-rock interaction, and Fe-rich differentiated gabbroic rocks altered during exhumation and fluid flow through the detachment system.

## 5.7. Fe-Ca metasomatism: an overlooked alteration of ultramafic rocks?

Fe-Ca metasomatism, including the formation of ilvaite formation,

has been primarily described in skarn deposits (Gustafson, 1974; Pesquera and Velasco, 1986; Yu et al., 2019). However, a limited number of studies have also reported its occurrence in fossil (Agata and Adachi, 1995; Naslund et al., 1983; Sakai and Kuroda, 1983; Graser and Markl, 2008) and present-day ultramafic-hosted systems (e.g., Gaggero et al., 1997; Coltat et al., 2023). In these cases, Fe-Ca metasomatism is commonly associated with mafic-ultramafic and alkaline intrusions (e. g., wehrlite, diorite, syenite), or occurs at lithological interfaces between mafic and ultramafic rocks. Such settings are typically enriched in Fe and, to a lesser extent, in Ca compared to typical peridotites. Although part of the Fe may derive from the magmatic protolith itself (e.g., Fe-rich wehrlite of the Skaergaard intrusion; Naslund et al., 1983), a significant proportion is likely externally introduced by hydrothermal fluids. The mineral assemblages formed during Fe-Ca metasomatism reflect interaction with Mg- and Si-poor, but Fe- and Ca-rich fluids, chemically distinct from those buffered solely by reactions with peridotite. Thus, Fe-Ca metasomatism does not spontaneously develop during typical serpentinization of mantle rocks, but rather under specific conditions that we explore below.

Ca-metasomatism in ultramafic rocks is more commonly associated with rodingitization (Bach and Klein, 2009), formation of talc-chlorite-tremolite schists (Boschi et al., 2006), or carbonate-brucite chimneys at the seafloor, such as the Lost City hydrothermal field (Mid-Atlantic Ridge, 30°N). The latter is linked to low-T (< 116 °C), alkaline (pH ~10–11) fluids (Seyfried Jr. et al., 2011), conditions that strongly contrast with those of Fe-Ca metasomatism at the MCHS. Rodingitization, commonly reported in ophiolitic sequences worldwide (Oman diopsidite, Python et al., 2007; Alps, Li et al., 2004; Greece, Rogkala et al., 2022; Betic Cordillera, Laborda-López et al., 2018; Tianshan and Tibet areas, Li et al., 2007; Li et al., 2017), as well as in present-day fracture zones (Honnorez and Kirst, 1975) and oceanic core complexes (Marques et al., 2006; Früh-Green et al., 2018), typically results from the interaction of mafic rocks with Ca-rich, metal-poor and alkaline fluids. These fluids may derive from clinopyroxene breakdown during serpentinization (Frost and Beard, 2007) and/or from gradients in aqueous silica and proton concentrations (Bach and Klein, 2009). However, the mineralogical and chemical signatures of rodingites remain fundamentally distinct from those observed at the MCHS.

In contrast, iron mobility poses a greater challenge. Part of the Fe may originate from magmatic intrusions and/or melt-rock interactions (see the previous section). According to Evans et al. (2017), Fe mobilization can be enhanced through the circulation of mineralizing fluids and/or fluid interaction with sulfide mineralization. At the MCHS, the synchronous formation of Fe-Ca silicates and sulfides suggests that they both precipitated from the same Fe-rich ore-forming fluid. A putative natural analog for such a fluid is provided by the Rainbow hydrothermal field, where black smoker fluids circulate through an ultramafic basement (Evans et al., 2024). There, Si-Cl thermometry indicates fluid buffering by plagioclase-talc-chlorite-tremolite assemblages at shallow depths and temperatures of 430–470 °C. These conditions differ markedly from those inferred for the MCHS. However, one could propose that the reaction zone of the ore-forming fluid at the MCHS was located deeper, under higher temperatures, and buffered by somehow similar reactions within mafic rocks. Fe-Ca metasomatism and sulfide precipitation would then occur during the ascent and cooling of the fluid. Notably, Evans et al. (2024) estimated a 24–30 % heat loss during fluid ascent at Rainbow. Applying a 30 % cooling from the original temperature 430–470 °C yields a temperature range of ~300–330 °C, consistent with the conditions inferred from thermodynamic modeling of Fe-Ca silicates stability at the MCHS (Fig. 8). In deeper structural levels of the MCHS, Fe-rich gabbroic rocks altered to actinolite-albite-chlorite ± biotite are documented (Desmurs et al., 2002), indicating lower amphibolite to greenschist facies conditions comparable to those in the proposed reaction zone beneath Rainbow. Although the protolith of these gabbros is uncertain, the replacement of Fe-Ti oxides suggests Fe was released during hydrothermal alteration. In such a scenario, fluid
flow would be favored at permeable mafic-ultramafic interfaces (Coltat et al., 2019b), similar to recently proposed hydrothermal pathways at the Rainbow Massif (Jian et al., 2024). If this model holds, Fe-Ca metasomatism together and associated sulfide mineralization should be detectable at Rainbow, especially along mafic-ultramafic contacts (Jian et al., 2024).

The plagioclase-chlorite-tremolite-talc assemblages mentioned above are commonly observed in altered mafic rocks from oceanic core complexes and mantle-dominated environments at slow to ultraslow spreading centers. Such assemblages have been reported from Holes U1309D (Ildefonse et al., 2007) and U1601C (Lissenberg et al., 2024), as well as from dive samples at the Atlantis Massif (Boschi et al., 2006), the 15°N area (Kelemen et al., 2007), the Mid-Atlantic Ridge Kane area (Cannat et al., 1995), and the Atlantis Bank on the South-West Indian Ridge (Dick et al., 2019). These settings host different magmatic lithologies—from juvenile olivine gabbros to evolved Fe-Ti oxide gabbros—intruded into a peridotitic basement. Recent drilling at the Atlantis Massif confirmed that magmatic rocks comprise ~30 % of the total rock volume (Lissenberg et al., 2024), though local proportions can approach 100 % (Ildefonse et al., 2007; Dick et al., 2019). The geochemical diversity of these rocks, and especially the occurrence of Fe-rich lithologies, may promote the production of Fe- and Ca-rich fluids during fluid-rock interaction, while buffering Si activity to low levels via plagioclase-chlorite-tremolite-talc assemblages (Evans et al., 2024). The interaction of such fluids with partially serpentinized peridotites near intrusive contacts could generate the observed Fe-Ca silicates and potentially associated sulfide mineralization, as documented at the MCHS. In this model, it remains unclear whether the peridotites must be pre-enriched in Fe, via prior melt-rock interactions (Coltat et al., 2023, 2025), before reacting with the hydrothermal fluid. We do not argue that Fe-Ca metasomatism is a systematic process in mantle rocks, but rather suggest that it is far from being exceptional, and has likely been overlooked in many systems. This under-recognition may stem from ilvaite's opaque appearance in optical microscopy, which hampers identification under transmitted light, and from the common dichotomy between studies of ore deposits and metamorphic petrology of mantle rocks.

## 5.8. Fe-Ca metasomatism as a vector for ultramafic-hosted VMS exploration

UM-VMS have been recognized worldwide (Patten et al., 2022) and, among them, many are associated with Fe-Ca silicates in ultramafic rocks located near mafic lithologies, such as in the Limassol Forest Complex (Thalhammer et al., 1986), Outokumpu (Peltonen et al., 2008), the Duluth igneous complex (Evans et al., 2017), the Western Alps (Toffolo et al., 2017) and the Apennines (Dini et al., 2024). The close spatial association between Fe-Ca silicates and sulfide mineralization makes the former a powerful petrological indicator for studying the formation of UM-VMS.

There is no universal metal endowment related to specific Fe-Ca silicate assemblages in UM-VMS. For instance, calc-silicates associated with the Cu-Zn-Co-Ni mineralization in Outokumpu mostly contain diopside and tremolite, while andradite and ilvaite are absent (Peltonen et al., 2008). On the other hand, ilvaite is the only Fe-Ca silicate observed in association with the Ni-Co-Cu mineralization of the Limassol Forest Complex (Thalhammer et al., 1986). This variability reflects the influence of local geological contexts, suggesting that specific Fe-Ca silicate assemblages depend on the nature and chemistry of the surrounding lithologies as well as the physico-chemical conditions (T, P, pH, $fO_2$) of the ore-forming environment, which can vary significantly from one system to another.

Similar to the MCHS, a genetic relationship between Fe-Ca metasomatism and metal deposition is reported in the fossil Cogne deposit (Western Alps). There, a magnetite-rich mineralized system associated with minor Cu-sulfides occurs in serpentinized harzburgites, diopside-

rich layers, and troctolites (Toffolo et al., 2017). In diopside-rich layers, magnetite occurs with andradite and antigorite in replacement of former diopside (Toffolo et al., 2017), whereas in serpentinized troctolites, nodular ores of magnetite replace former olivine grains. Toffolo et al. (2017) proposed that hydrothermal Fe-metasomatism occurred in an oceanic setting during mantle exhumation at the seafloor, following serpentinization of mantle rocks, at temperatures around 300–400 °C. This temperature range is in good agreement with our thermodynamic calculations (300–360 °C, Fig. 8). Excluding the absence of ilvaite, the petrographic and geochemical signatures of the mineralization at Cogne show strong similarities with those observed at the MCHS, notably an increase of the Co/Ni ratio with the extent of mineralization. A similar trend is reported in Ligurian Cu-Zn VMS genetically associated with ilvaite and andradite (Dini et al., 2024), consistent with trends observed in present-day seafloor massive sulfides (Marques et al., 2006).

Fe-Ca silicates can form together with hydrothermal sulfide and oxide in ultramafic-hosted deposits. The recognition of ilvaite, andradite, and/or diopside in mantle rocks near mafic intrusions should be considered reliable petrographic vectors towards Cu-Co-Zn-Ni mineralization, thereby enhancing target identification in mineral exploration. When available, in situ trace element geochemical analyses of Fe-Ca silicates may provide valuable additional insights into the “metal fertility” of the system. In particular, the Co/Ni ratio of single minerals and whole rocks tends to increase with the extent of mineralization, and values greater than 0.05 (i.e., average Co/Ni in serpentinite) should be considered as potential indicators of ore-forming processes.

## 6. Conclusions

In the fossil Platta nappe, Swiss Alps, the Marmorera-Cotschen Hydrothermal System recorded oceanic hydrothermalism responsible for Fe-Ca metasomatism and metal deposition. Ilvaite, hydro-andradite, and diopside crystallized in mantle rocks at the contact with mafic intrusions, coeval with Cu-Fe-Co-Zn-Ni mineralization. Petrographic and geochemical analyses, combined with thermodynamic modeling of Fe-Ca silicates, indicate that Fe-Ca metasomatism occurred during early stages of serpentinization of mantle rocks, at temperatures ranging from 300 to 360 °C, during mantle exhumation.

Geochemical signatures of Fe-Ca silicates (Co, Ni, and REE concentrations) reveal distinct fluid-rock interactions between barren and mineralized samples. In mineralized mantle rocks, Fe-Ca silicates crystallized in an open system, from a fluid that interacted with both mafic and ultramafic rocks. In contrast, pervasive Fe-Ca silicates in barren metasomatized mantle rock formed in a rock-dominated system, originating from a fluid that had previously reacted with ultramafic rocks.

The study of the Marmorera-Cotschen hydrothermal system suggests that Fe-Ca metasomatism is likely a widespread process in the shallow lithospheric mantle during tectonic exhumation, particularly at the contact zones between magmatic and mantle rocks or in mantle rocks altered by melt-rock interaction. Fe-Ca metasomatism is often accompanied by sulfide mineralization and could represent an efficient mechanism for pre-concentrating metals in the ultramafic basement. These metals might later be remobilized during subsequent magmatic events and black smoker-like hydrothermal activity, contributing to the formation of seafloor massive sulfides. This suggests a potential genetic continuum between deep-seated hydrothermal alteration and seafloor mineralization processes. Ultimately, Fe-Ca metasomatism, particularly when marked by the presence of ilvaite, andradite, or diopside, emerges as a valuable proxy in the exploration of ultramafic-hosted volcanogenic massive sulfide deposits.

### CRediT authorship contribution statement

R. Coltat: Writing – review & editing, Writing – original draft, Software, Project administration, Methodology, Investigation, Funding
acquisition, Formal analysis, Conceptualization. C.G.C. Patten: Writing – review & editing, Formal analysis, Funding acquisition. J.A. Padrón-Navarta: Writing – review & editing, Software, Funding acquisition. F. Hochscheid: Writing – review & editing. M. Ulrich: Writing – review & editing. Y. Branquet: Writing – review & editing. P. Boulvais: Writing – review & editing, Funding acquisition. C.J. Garrido: Writing – review & editing, Funding acquisition.

## Declaration of competing interest

The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

## Acknowledgments

We thank X. Lecoz and J. Langlade for thin-section preparation and EPMA analyses, respectively. R. Coltat acknowledges "Juan de la Cierva" Fellowship (FJC2021-047190-I), funded by MCIN/AEI/10.130 39/501100011033 with EU Next Generation Funds. Additionally, sam- ple collecting and preparation, and EPMA & LA-ICP-MS analyses were initially supported by French CNRS-INSU-CESSUR and University of Rennes 1 "Défis Scientifiques" grants (P. Boulvais and R. Coltat), the French Ministry of Research (pH.D. of R. Coltat), and the German Connecting Young Scientists Program (C. Patten). C.J. Garrido, J.A. Padrón-Navarta, and R. Coltat acknowledge funding from project PID2022-136471NB-C21 "The role of ultramafic rocks and related rocks in the sulfur and water cycles and their implications for the redox state of subduction zones (RUSTED)", funded by MCIN/AEI/10.130 39/501100011033 and by the ESF+.

## Appendix A. Supplementary data

Supplementary data to this article can be found online at https://doi. org/10.1016/j.chemgeo.2025.122826.

## Data availability

Data will be made available on request.

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