# Carbon trapping during contact metamorphism in volcanic basins: example of the Guaymas basin

Alban Cheviet, Philippe Goncalves, Flavien Choulet, Christophe Galerne, Wolfgang Bach, Armelle Riboulleau, Torsten Vennemann, Martine Buatier

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Alban Cheviet, Philippe Goncalves, Flavien Choulet, Christophe Galerne, Wolfgang Bach, et al. Carbon trapping during contact metamorphism in volcanic basins: example of the Guaymas basin. Contributions to Mineralogy and Petrology, 2025, 180 (10), pp.74. 10.1007/s00410-025-02262-0. hal-05318373

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# Carbon trapping during contact metamorphism in volcanic basins: example of the Guaymas basin

Alban Cheviet¹,² ID · Philippe Goncalves² · Flavien Choulet² · Christophe Galerne¹,³ · Wolfgang Bach¹,³ ·

Armelle Riboulleau⁴ · Torsten Vennemann⁵ · Martine Buatier²

Received: 9 January 2025 / Accepted: 25 August 2025 / Published online: 17 September 2025
© The Author(s) 2025

## Abstract

Magmatic activity in young oceanic sedimentary basins impacts the transfer of life-essential volatiles, such as C, S, H and Cl between the Earth's interior and the surface. The intrusion of magmas into sediments leads to significant remobilization and/or sequestration of carbon and sulphur in the metamorphic aureoles. Core samples from the Guaymas Basin, collected during the IODP 385 expedition include Site U1546, where a tholeiitic sill was encountered. The metamorphic aureole below the sill is characterized by a pyrrhotite-pyroxene-calcite-siderite assemblage in a quartz-bearing matrix. Combining petrological and geochemical analyses, coupled with thermodynamic modelling, we show that cooling under a quasi-closed system conditions controls the precipitation of carbonates in equilibrium with sulphides. In contrast, the upper aureole remains unaffected by retrograde phase precipitation, suggesting that fluids have been totally evacuated from the system. While degassing toward the surface is not in question, we estimate that up to 38% of the carbon (equivalent to 1.26 Mt) may have been trapped in the metamorphic aureoles of the sill U1546 in the Guaymas Basin. At the basin scale, the volume of these metamorphic reservoirs, that are directly correlated to the magmatic activity, and their ability to trap volatiles, are important considerations for global cycles of the volatile elements. These findings contrast with the common view that metamorphic aureoles are always a source of volatiles, which has implications for estimating the volatile release fluxes during past episodes of increased magma-sediment interaction in analogous basin settings and their consequences on the atmospheric carbon cycle.

**Keywords** Contact metamorphism · Carbon sequestration · Fluid-rock interactions · Magma-sediment · IODP expedition 385 · Thermodynamic models

Communicated by Othmar Müntener.

Alban Cheviet
alban.cheviet@univ-fcomte.fr

¹ Department of Geosciences, University of Bremen,
Klagenfurter Straße 2-4, 28359 Bremen, Germany

² Université Marie et Louis Pasteur, CNRS, Chrono-
environnement (UMR 6249), Besançon F-25000, France

³ MARUM Center for Marine Environmental Sciences,
University of Bremen, 28359 Bremen, Germany

⁴ Laboratoire d'Océanologie et de Géosciences, UMR 8187,
Université de Lille, CNRS, Université du Littoral Côte 10
d'Opale, Villeneuve d'Ascq 59655, France

⁵ Laboratory of Biogeosciences, Institute of Earth Surface
Dynamics, University of Lausanne, Lausanne, Switzerland

## Introduction

Sedimentary basins act as major reservoirs of carbon and sulphur in the Earth's upper crust, and they also play an important role in the chemical exchange between the surface and the Earth's interior. Understanding the impact of magmatic sill emplacement in these basins on the carbon and sulphur cycles is essential for estimating the global carbon and sulphur cycles and their influence on the past climate changes and mass extinctions (Ernst and Youbi 2017; Ganino and Arndt 2009; Shen et al. 2011; Yallup et al. 2013).

Many studies have quantified carbon emissions related to the emplacement of magmatic intrusions and contact metamorphism in sediments (Aarnes et al. 2010, 2011a; Armstrong McKay et al. 2014; Bédard et al. 2023; Brune et al. 2017; Capriolo et al., 2021; Deegan et al. 2022; Ernst and Youbi 2017; Galerne and Hasenclever 2019; Ganino
and Berndt 2009; Heimdal et al. 2018; Jones et al. 2016; Svensen et al. 2004). For instance, during the Permian–Triassic crisis, the emplacement of sills in the Tunguska Basin related to the Siberian traps potentially contributed to the release of 12,000 to 66,000 Gt of CH$_4$ into the oceans and the atmosphere (Aarnes et al. 2010; Kontorovich et al. 1997; Svensen et al. 2009), leading to a major change in climate and significant mass extinction (Burgess et al. 2017). The Paleocene–Eocene thermal maximum coincides with a significant increase of the concentration of carbon in the atmosphere, which could be related to contact metamorphism and subsequent devolatilization around sills in the North Atlantic Igneous Province (Svensen et al. 2004). In all of these studies, carbon outgassing budgets were estimated by using numerical heat transfer models to quantify the carbon released via organic matter [OM] cracking and host rock dehydration and decarbonatization (Aarnes et al. 2010, 2011a; Galerne and Hasenclever 2019; Iyer et al. 2018). Simplified quantitative models use the assumption that the totality of the produced fluids is emitted to the seafloor through hydrothermal vents (Aarnes et al. 2010, 2011a; Iyer et al. 2018). In contrast, more complex 2D fluid circulation models have shown that the fluids under the sills can remain partly trapped, limiting upward advection (Galerne and Hasenclever 2019; Iyer et al. 2013). Interaction of the metamorphic fluids confined under the intrusion and the host rock leads to the precipitation of metamorphic carbonates possibly sequestrating a significant part of the carbon released during the peak of temperature (Cheviet et al. 2023).

In this contribution, we propose that a metamorphic carbon sink is induced by magmatic sill emplacement in an oceanic sedimentary basin. We determine the most favourable chemical (e.g. redox conditions) and environmental (e.g. temperature) conditions to stabilize carbonates together with sulphides and quantify the amount of carbon sequestered in the metamorphic aureoles. Our study focuses on Site U1546 of IODP Expedition 385 (Teske et al., 2020; Teske et al. 2019), where a 76-m thick magmatic sill was emplaced in soft sediments at 354 m below the seafloor [mbsf] (Cheviet et al. 2023; Teske et al. 2021a).

## Geological setting

The Guaymas Basin is located in the Gulf of California, at the Pacific-North America tectonic plate boundary (Fig. 1a). At present, the plate boundary in the central to southern Gulf of California comprises several narrow spreading segments in the form of axial troughs separated by transform faults (Fig. 1a). Two spreading centres in the Guaymas Basin, physically manifested by graben structures developed in the

sedimentary basin fill (Fig. 1b). New igneous oceanic crust formation has accommodated the Pacific/North American plate motion (Lizarralde et al. 2007; Miller and Lizarralde 2013). The sedimentary succession on the oceanic crust consists of mainly organic-rich biogenic sediments derived from highly productive overlying waters and minor terrigenous sediments from the Baja California peninsula and the Sonora margin (Einsele et al. 1980; Saunders et al. 1982).

The north-western Guaymas Basin was the primary focus of IODP Expedition 385 (Teske et al., 2020; Teske et al. 2021a). It is characterized by a large off-axis sill network emplaced at shallow depth (Berndt et al., 2016). During the expedition, eight sites were drilled, covering an 81 km-long northwest-southeast transect (Fig. 1b). Site U1546 recovered the oldest and thickest sediment succession (to ~540 mbsf) of IODP Expedition 385. The sediment succession is mainly composed of clay-rich diatom ooze disrupted by a ~74–76 m-thick sill intrusion located between 354.6 and 430.0 mbsf (Teske et al. 2021b).

A previous study (Cheviet et al. 2023) has determined the thickness of the metamorphic aureole using mineralogical transformations and OM degradation as proxies. The main mineralogical transformation is the conversion of the biogenic/diagenetic silica into metamorphic quartz (Cheviet et al. 2023). According to the appearance of the quartz and the decrease of TOC around the sill at Site U1546, the upper metamorphic aureole is about 8 m thick, while the thickness of the lower aureole may reach 16 m (Fig. 1c).

## Materials and methods

### Organic-inorganic carbon and sulphur measurements

Measurements of organic and inorganic carbon and total sulphur in sediments and metasediments are a compilation of data acquired during IODP Expedition 385 and detailed in the Expedition 385 Methods section (Teske et al. 2021b). Additional analyses were carried out on samples close to the contact. Inorganic carbon, TOC and total sulphur were measured by Rock-eval analysis (Baudin 2024; Baudin et al. 2015; Espitalie et al. 1985), carried out by the RE7S instrument in “Recent Sediment” mode with measurement of total S, pyrolysis starting at 180 °C and increasing to 650 °C at a rate of 30 °C/min, followed by oxidation from 300 °C to 1200 °C at a rate of 20 °C/min. These data were completed with data from the literature (Cheviet et al. 2023). The TOC was measured by HCl (25%) leaching for 1 h to expel carbonates as CO$_2$, and the residue was analysed for C and S by induction furnace on the same samples, these analyses
<figure><img src="image_2.png" /><figcaption>**Fig. 1** a Satellite view and tectonic setting of western North America and the Gulf of California. Modified from Google Earth with information from Teske et al. (2020). Boxed insert shows the location of the lower map. b Bathymetric map of Guaymas Basin, Mexico, with Baja California to the southwest, the Sonora margin to the northeast and the</figcaption></figure>

location of the Deep-Sea Drilling Project (DSDP) Leg 64 and IODP Expedition 385 sites. modified Modified from Teske et al. (2020). c Lithostratigraphic columns of Site U1546, showing the depth of the sill and the thickness of the metamorphic aureoles determined in Cheviet et al. 2023. Modified from Cheviet et al. 2023

were carried out at the laboratory ALS Dublin accredited to ISO=IEC 17,025: 2017 standards (Table. SI1).

Stable isotope analyses

Carbonate $\delta^{18}$O and $\delta^{13}$C isotope measurements were carried out on bulk rock powders and micro-drilled mineral fractions. Metamorphic calcite-siderite corona in the lower aureole were separated from the quartz matrix by micro-drilling using a DREMEL 3000JD. For sediments and metasediments in the upper aureole, where carbonates are

disseminated, analyses were carried out on whole rock powders. The carbonates composition of sediments (calcite and/or dolomite) was determined by XRD (Cheviet et al. 2023). Oxygen ($^{18}$O/$^{16}$O) and carbon ($^{13}$C/$^{12}$C) isotopic analyses were performed using a Gas Bench II coupled to a Thermo Finnigan Delta V mass spectrometer.

Isotopic analyses of $\delta^{18}$O of silica were performed on the 20–50 µm fraction of the whole rock to remove clays minerals, detrital grains, pyrrhotite, concentrating the metamorphic quartz for the aureoles and the opal-A/CT for the sediments. The 20–50 µm fraction was separated by decantation in a water column, and the purity of the quartz fraction
was verified by microscopy. Oxygen analyses were carried out using a laser CO₂ extraction line and fluorination using a Nu-Instruments Perspective coupled to a MAT253 mass spectrometer.

All analyses were performed at the Isolab, University of Lausanne, Switzerland. Subsequently, δ¹⁸O values were normalised to the V-SMOW standard and δ¹³C to the V-PDB standard (Craig 1961; Friedman and O'Neil 1977).

## Microscopy

Thin sections of epoxy-impregnated sediments were assessed using a numerical Keyence optical microscope (OM) and a ThermoFisher APREO S scanning electron microscope (field emission gun (FEG)-SEM), hosted at Université Marie et Louis Pasteur, France. For SEM observations, polished thin-sections were carbon-coated, and the images were acquired in high vacuum mode with a backscatter detector, an acceleration voltage of 15 kV and a beam current of 1.6nA.

## Microprobe

Electron probe microanalysis (EPMA) of pyroxene, calcite and siderite minerals was done on carbon-coated polished thin-sections using a JEOL JXA-8530 F HyperProbe microprobe, equipped with five wave-dispersive X-ray spectrometers (WDS), hosted at University of Lausanne. We operated at 15 kV with a beam current of 10 nA and a spot size of 1 µm for the pyroxenes and 5 µm for calcite and siderite. Standards used for the pyroxenes analyses were: Na₂O: albite, Al₂O₃: andalusite, MgO: forsterite, TiO₂: ilmenite, SiO₂: wollastonite, CaO: wollastonite, FeO: fayalite, MnO: tephroite, NiO₂: Ni oxide, Cr₂O₃: chromite The standards for calcite and siderite analyses were: SiO₂: wollastonite, MgO: dolomite, SrO: Sr glass C5_stron, FeO: siderite, CaO: calcite, MnO: rhodochrosite, BaO: Ba oxide ben2.

## Heat transfer modelling

To simulate the thermal evolution associated with the sill emplacement in the sediments, we used the SILLi code developed by Iyer et al. (2018), which solves conductive heat transfer during contact metamorphism. Model parameters were adjusted to reflect the geological and thermal conditions of the Guaymas Basin, with a 75 m-thick sill intruding sediments at a paleo-depth of 250 mbsf (Galerne et al. 2025). Following the sill emplacement, an additional ~100 m of sediments were deposited. In Fig. 6, the depth of the upper margin of the sill accounts for this extra

sedimentation to fit the present-day depth of the sill at Site U1546. The emplacement temperature of the basaltic melt was set at 1050 °C, a relatively low value intended to account for heat lost during basaltic magma accent through the sediments. The thermal and physical properties of the system were defined based on a compilation of the IODP Expedition 385 data and from literature compiled in Table 1.

## Thermodynamic modelling

Phase diagram projections and sections were used to study the effect of a COHS fluid on stable mineralogical assemblages. Phase relations were computed in the Ca-Fe-Mg-Si-C-H₂-O₂ system using Perple_X_7.1.5 program (Connolly 2005) with the thermodynamic database from Holland and Powell (2011). Solution models and end-member phases considered in the models are listed in Table 2.

Fluid composition has been modelled using a generic fluid solution model COHS-Fluid (Connolly and Cesare 1993). The bulk composition used for phase diagram sections corresponds to the average of twelve compositions of metasediments given in Table SI4 from Cheviet et al. (2023) and shown on the phase diagrams. To simplify the models, we have used a fluid with H₂O-CO₂-H₂S species, excluding CH₄ to maintain the models in oxidizing conditions. This choice is justified by the presence of pyrrhotite indicating that the COHS metamorphic fluid has an H/O close to 2:1 (Connolly and Cesare 1993). Furthermore, we observed carbonates, indicating that the system has to have a H/O<2 to favour carbon oxidation. Although the sediment contains OM that can produced CH₄ by cracking, the presence of sulphur can buffer the H₂. The amount of carbon was fixed based on the TOC measured in the sediments but also the inorganic carbon in the biogenic and diagenetic carbonates measured in the sediments (Table SI1). Due to the high pore-water content in the sediments, the sediment can be considered water saturated. The amount of water in the models corresponds to the sum of loss of ignition [LOI] measured in the sediments (15 wt%) from Cheviet et al. (2023) given in Table SI4, which is enough to saturate the system. The phase diagrams were computed at three distinct fixed P-T conditions corresponding to the peak of temperature during the contact metamorphism (600 °C), and the temperature of precipitation estimated for carbonates (400–200 °C) based on heat transfer models (see main text for more details). The water-rich sediments are unconsolidated with a high porosity, so to constrain the pressure of thermodynamic computation because the porosity is connected in sediments, we consider an hydrostatic pressure estimated at 200 bar, which corresponds to a water column of 1,586 m of seawater above the sediment and 430 m of porewater in sediment,
<table><caption>**Table 1** Input parameters for the SILLi model</caption><thead><tr><th>Symbol</th><th>Value</th><th>Description</th><th>Units</th><th>References</th></tr></thead><tbody><tr><td>T0</td><td>2</td><td>Temperature at the seafloor</td><td>°C</td><td>This study</td></tr><tr><td>T500</td><td>89</td><td>Temperature at 500 mbsf</td><td>°C</td><td>Teske et al. (2021)</td></tr><tr><td>th</td><td>75</td><td>Sill thickness</td><td>m</td><td>Teske et al. (2021)</td></tr><tr><td>dTop</td><td>250</td><td>Sill upper contact at time of emplacement</td><td>m</td><td>Galerne et al. 2025</td></tr><tr><td>cpf</td><td>3000</td><td>Fluid (water) heat capacity</td><td>J.kg<sup>-1</sup>.K<sup>-1</sup></td><td>Grunberg (1970)</td></tr><tr><td>cpr</td><td>1000</td><td>Rock heat capacity</td><td>J.kg<sup>-1</sup>.K<sup>-1</sup></td><td>Hadgu et al. (2007)</td></tr><tr><td>Lc</td><td>400</td><td>Latent heat of crystallization</td><td>KJkg<sup>-1</sup></td><td>Aarnes et al. (2010)</td></tr><tr><td>TL</td><td>1100</td><td>Liquidus temperature</td><td>°C</td><td>Aarnes et al. (2010)</td></tr><tr><td>TS</td><td>1000</td><td>Solidus temperature</td><td>°C</td><td>Aarnes et al. (2010)</td></tr><tr><td>pm</td><td>2600</td><td>Melt (basalt) density</td><td>kgm<sup>-3</sup></td><td>Aarnes et al. (2010)</td></tr><tr><td>pc</td><td>2700</td><td>Solidified basalt density</td><td>kgm<sup>-3</sup></td><td>Aarnes et al. (2010)</td></tr><tr><td>ThC</td><td>1.8</td><td>Melt thermal conductivity</td><td>W.m<sup>-1</sup>.K<sup>-1</sup></td><td>Aarnes et al. (2010)</td></tr><tr><td>Hm</td><td>820</td><td>Heat Capacity melt (basalt)</td><td>J.kg<sup>-1</sup>.C<sup>-1</sup></td><td>Aarnes et al. (2010)</td></tr><tr><td>Hs</td><td>850</td><td>Heat Capacity solid (basalt)</td><td>J.kg<sup>-1</sup>.C<sup>-1</sup></td><td>Iyer et al. (2018)</td></tr><tr><td>T</td><td>1050</td><td>Sill emplacement temperature</td><td>°C</td><td>This study</td></tr><tr><td>Z</td><td>0–500</td><td>Depth</td><td>km</td><td></td></tr><tr><td>φ</td><td>0.8–0.45</td><td>Rock porosity</td><td>Fraction</td><td>Teske et al. (2019)</td></tr><tr><td>κr</td><td>0.75</td><td>Rock thermal conductivity</td><td>W.m<sup>-1</sup>.K<sup>-1</sup></td><td>Teske et al. (2019)</td></tr><tr><td>κf</td><td>0.6</td><td>Fluid (water) thermal conductivity</td><td>W.m<sup>-1</sup>.K<sup>-1</sup></td><td>Aarnes et al. (2010)</td></tr><tr><td>ρf</td><td>1051</td><td>Fluid (water) density</td><td>kgm<sup>-3</sup></td><td>Aarnes et al. (2010)</td></tr><tr><td>pr</td><td>1000–4000</td><td>Rock density</td><td>kgm<sup>-3</sup></td><td>Teske et al. (2019)</td></tr><tr><td>Lo</td><td>375</td><td>Latent Heat of Organic Maturation</td><td>kJ.kg<sup>-1</sup></td><td>Aarnes et al. (2010)</td></tr><tr><td>Ld</td><td>2800</td><td>Latent Heat of Dehydration</td><td>kJ.kg<sup>-1</sup></td><td>Aarnes et al. (2010)</td></tr><tr><td>TOC</td><td>2</td><td>Total Organic Content in sediments</td><td>wt%</td><td>This study</td></tr><tr><td>t</td><td></td><td>Time</td><td>s</td><td></td></tr></tbody></table>

with an average water density of 1.025 g.cm⁻³ (Teske et al. 2021a). Overpressure resulting from compaction during the emplacement of the sill and fluid pressure generated during dehydration reactions has been neglected.

## Results

### Host rock composition and metamorphic reactions

The sediments are unconsolidated and mostly consist of biogenic silica (opal-A), which is progressively transformed into diagenetic silica (opal-CT) with increasing depth and temperature towards the sill. The sediments also contain framboidal pyrite (FeS₂), organic matter (OM), and a small amount of detrital silicates (clays, feldspars, quartz). Biogenic and diagenetic carbonates are locally present. The average total organic carbon (TOC) content in the sediments is 2.1±0.8 wt%, while the total sulphur (TS) content ranges from 1 to 2 wt% (Fig. 2a). The amount of inorganic carbon (IC) is between 0.13 and 0.48 wt%. Carbonate nodules, mostly composed of diagenetic dolomite, were observed in distinct sections of the core throughout the Guaymas Basin and are interpreted as diagenetic dolomite (Kelts and McKenzie 1982; Teske et al. 2021c). One of these enriched horizons is present at 348.09 mbsf, in the transition zone between the sediments and the metamorphic aureole (Figs. 2b).

In the upper metamorphic aureole, sedimentary laminae are still preserved (Fig. 3a) and are characterised by the presence of small (~20 µm) grains of metamorphic quartz and pyrrhotite (FeS) (Fig. 3). The polymorphic transformation of opal-A/CT into quartz occurs about 8 m above the sill and the reduction of pyrite into pyrrhotite was reported 2 m below, closer to the sill contact (Cheviet et al. 2023). Detrital grains are still present, and the porosity is filled by clay minerals (Fig. 3b). The TOC and TS contents clearly drop with decreasing the distance to the sill, and the TOC content tends to zero at the contact. The minimum TS content at the contact is 0.54 wt% in the upper aureole (Fig. 2a). The amount of IC in the upper metamorphic aureole does not exceed 0.46 wt% near the contact excluding the diagenetic horizon just above the contact aureole limit (Fig. 2b).

The lower metamorphic aureole is thicker, with quartz appearing 16 m below the sill, while pyrrhotite occurs only 6 m away from the contact (Cheviet et al. 2023). Downcore changes in TOC and TS mirror those above the sill with an almost total depletion at the contact (Fig. 2a). In contrast, the amount of IC in the lower metamorphic aureole is much higher, reaching up to 1.68 wt% (Fig. 2b). The increase of the IC content is observed in the first 6 m-thick metasediment horizon below the sill. In the lower metamorphic aureole, the metasediments are recrystallised with numerous millimetre-large patches of carbonates and sulphides (Fig. 4b, c, d). The patches are disseminated in a matrix of metamorphic quartz, minor plagioclase feldspars and clay minerals (Fig. 4c, d). The presence of these patches is consistent with the IC-enriched layers described previously. These patches consist of a partially resorbed pyrrhotite
<table><caption>**Table 2** Solid solutions models and phases used in the phase diagrams</caption><thead><tr><th>Phases</th><th>Solid solution label used in PerleX</th><th>Pure phase and solid solution label used in texte</th><th>Independent end-members</th><th>Independent end-members labels used in figures</th><th>References</th></tr></thead><tbody><tr><td>Clinopyroxene</td><td>Cpx(stx7)</td><td>cpx</td><td>Hedenbergite-diopside</td><td>hed-di</td><td>Holland & Powell (1998)</td></tr><tr><td>Amphibole</td><td>Tr</td><td>amph</td><td>Tremolite-ferroactinolite</td><td>tr-ftr</td><td>Holland & Powell (1998)</td></tr><tr><td>Carbonate</td><td>oCcM(EF)</td><td>carb</td><td>Calcite-magnesite-siderite</td><td>cc-mgs-sid</td><td>Massonne & Willner (1998)</td></tr><tr><td>COHS-Fluid</td><td>COH-Fluid</td><td>FH2O,CO2,H2S</td><td>H2O-CO2-H2S-SO2</td><td>H2O-CO2-H2S-SO2</td><td>Connoly & Cesare (1993)</td></tr><tr><td>Quartz</td><td>q</td><td>qz</td><td></td><td></td><td></td></tr><tr><td>Pyrite</td><td>py</td><td>py</td><td></td><td></td><td></td></tr><tr><td>Pyrrhotite</td><td>po</td><td>po</td><td></td><td></td><td></td></tr><tr><td>Magnetite</td><td>mt</td><td>mag</td><td></td><td></td><td></td></tr><tr><td>Iron</td><td>iron</td><td>iron</td><td></td><td></td><td></td></tr><tr><td>Anhydrite</td><td>any</td><td>any</td><td></td><td></td><td></td></tr></tbody></table>

core separated from the quartz-bearing matrix by a two-layer corona (Fig. 4d, e). The inner corona is formed by a clinopyroxene-calciteª-sideriteª assemblage without quartz (Fig. 4d, e). Clinopyroxene is a hedenbergite-diopside solid solution (60% Di and 40% Hed; Table 3). The outer corona is composed of metamorphic calciteª, sideriteᵇ and quartz. The average carbonate proportion of the patches (including both coronas) is 80% calcite and 20% siderite (Fig. SI1). Metamorphic calciteª and calciteᵇ in respectively inner and outer corona have the same composition (Table 3). Metamorphic sideriteª and sideriteᵇ display an intermediate composition between siderite and ankerite (Ca₀.₁₂⁻₀.₁₉Fe₀.₄₆⁻₀.₄₉Mg₀.₃₀(CO₃) (Table 3).

## Isotopic data and peak metamorphism temperature

The temperature of the metamorphic peak (Tmax) was estimated using heat transfer modelling (Silli code from Iyer et al. 2018), with parameters specific to the Guaymas setting (Table 1). The maximum temperature reached at the sill is 715 °C and 720 °C for the upper and lower metamorphic aureole respectively (Fig. 6a). The evolution of the temperature over time was extracted at 3 different depths (Fig. 6b): 1 m below the sill (431 mbsf), which corresponds to the metasediments the closest to the sill, 6 m below the sill (436 mbsf), corresponding to the upper limit of observed carbonates and 16 m meters below the sill (446 mbsf), which is the limit of the metamorphic aureole based on the presence of quartz. Modelled peak temperature extracted at these three depths are: 720 °C, 605 °C and 460 °C (Fig. 6b). At the maximum temperature (i.e. 720 °C), the quartz-pyrrhotite-fluid (H₂O, CO₂, H₂S) metamorphic assemblage is produced

(Cheviet et al. 2023). Modeling shows that retrogression of quartz-pyrrhotite high-temperature assemblage into pyroxene and carbonate (Fig. 4), takes place 150 years after sill emplacement. During the same time frame, peak temperature at 6 and 16 m are reached (Fig. 6b) and all metasediments recorded the high-temperature event characterized by the transformation of opal-A/CT into quartz and pyrite into pyrrhotite (Cheviet et al. 2023). Thermal relaxation, corresponding to a temperature of ca. 200 °C in the whole aureole is reached 2000 years after sill emplacement. We conclude that the retrograde path in the metamorphic aureole lasts ~ 1850 years over a temperature range between 600 and 200 °C.

The isotopic composition of the biogenic/diagenetic silica (opal-A/opal-CT) shows a decrease in δ¹⁸O from 23.72 to 19.29‰ (V-SMOW), with increasing depth and temperature (Fig. 5). In the metamorphic aureoles, the δ¹⁸O values of metamorphic quartz are significantly lower (Fig. 5). Above the sill, the δ¹⁸O decreases from 15.62‰ to 13.07‰ close to the sill (Fig. 5). The measured δ¹⁸O values of quartz below the sill range from 10.6 at the sill contact to 14.6‰ with an average of 12.4‰ (Fig. 5). The δ¹⁸O of the carbonates in the sediments, range from 30.78‰ to 25.25 and the δ¹³C is between 1.23 and -0.67‰ (Fig. 5). Above and below the sill, the δ¹³C display small variations with values between -1.88 and -0.20‰ for the upper aureole and -1.71 and 1.60‰ for the lower aureole. The δ¹⁸O values of metamorphic calcite and siderite, which are widespread in the lower metamorphic aureole, range from 14.02 to 23.5‰, consistent with a metamorphic origin. Their isotopic composition is clearly different from biogenic/diagenetic carbonates with δ¹⁸O values > 25‰ (Fig. 5).
<figure><img src="image_3.png" /><figcaption>**Fig. 2** a Concentration of total organic carbon (TOC), and total sulphur (TS) of the sediments and metasediments. b Inorganic carbon (IC) content in wt% of the sediments and metasediments. The dark</figcaption></figure>

The isotopic composition of the metamorphic fluid in equilibrium with quartz was estimated using the maximum temperature (Tmax) obtained by heat transfer modelling at each sample depth (Table SI3). The oxygen isotopic fractionation between quartz and water using Clayton et al. (1972) calibration shows a δ¹⁸O fluid composition decreasing from 13.4 to 10.2‰ towards the sill from 436.6 to 432.7 mbsf, (Fig. 6c). Two samples closer to the sill, at depth lower than 432.06 mbsf have higher δ¹⁸O fluid composition of 11.27 and 13.33‰ near the sill (Fig. 6c).

The oxygen isotopic fluid composition in equilibrium with carbonate was estimated by fixing the temperature at 600 °C, 400 °C and at 200 °C, corresponding to the temperature range defined for retrogression (Fig. 6d). Based on the isotopic fractionation calibration of Friedman and O'Neil (1977) between a metamorphic fluid and the carbonate corona consisting of 80% calcite and 20% siderite, we obtained an oxygen isotopic composition of the fluid

<figure><img src="image_4.png" /><figcaption>blue outlier corresponds to the diagenetic dolomite nodule. Grey areas correspond to the metamorphic aureoles deduced from the appearance of quartz. All values are in Table SI1</figcaption></figure>

ranging from 13.24 to 22.73‰ at 600 °C and 4.07 to 8.56‰ at 200 °C (Fig. 6d). The δ¹⁸O of the metamorphic fluid in equilibrium with quartz and carbonates in the corona are in the same range (Fig. 6d). Hence, we consider a closed system, in which the metamorphic fluid produced at high temperature remains trapped in the lower metamorphic aureole during retrograde metamorphism. In the following section, phase equilibrium models were conducted at 600 °C, 400 °C and 200 °C to follow the evolution of the metamorphic fluid and mineralogic assemblage during the cooling of metasediments.

## Thermodynamic modelling of pyrrhotite-pyroxene-carbonates equilibrium

The petrological observations indicate that the dissolution of pyrrhotite in the presence of quartz enables the formation of pyroxene, calcite and siderite (inner corona) and calcite,
<figure><img src="image_5.png" /><figcaption>Fig. 3 a Optical microscope scan of a metasediment above the sill (353.72 mbsf) showing sedimentary laminae. b SEM backscattered image of the fine-grained matrix of the metasediment above the sill. Mineral abbreviations: Qz quartz, Fds feldspar, Po pyrrhotite, Cal calcite, Dol dolomite</figcaption></figure>

siderite, quartz (outer corona). This is the first time that these reactions involving silicates, sulphides and carbonates with a fluid have been observed in this type of system, which is why these reactions are investigated through thermodynamic modelling. This mineralogical reaction is likely driven by changes in environmental (i.e. temperature) and chemical variables, i.e., the chemical potential (or fugacity) of O₂ and S₂ (Poulson 1996). Phase relations are described in isothermal and isobaric µO₂-µS₂ phase diagrams projections computed for simplified chemical systems (Fe-S₂-O₂ and Si-Fe-Ca-H₂-C-O₂-S₂) using end-member minerals. Temperature used are 600, 400 and 200 °C, corresponding to the estimated temperature of quartz and carbonate precipitation during cooling. Pressure is set at 200 bars, which corresponds to the hydrostatic pressure at sill emplacement depth (Fig. 1).

In the Fe-S₂-O₂ system (Fig. 7), the stability of sulphides strongly depends on the chemical potential of sulphur. At high µS₂ values, pyrite is stable. For a constant µO₂, but at lower µS₂, pyrrhotite becomes stable, followed by native iron at very low µS₂. The stability of sulphides also depends on µO₂, controlling the sulphide/oxide transition. Under more oxidising conditions, wustite (at high temperature), magnetite or even hematite are stable (Fig. 7).

For the purpose of this study, we have included Si, C, Ca, and H₂ in the system to account for the presence of Ca-bearing silicates, carbonates and COHS-fluids. At 600 °C, In the stability field of pyrrhotite, an increase of µO₂ coupled with a decrease in µS₂ can destabilise pyrrhotite and quartz to form hedenbergite and CO₂ according to the equilibrium (1) (Fig. 7b):

$$2qz + 2po + cal + O_{2(g)} = hed + CO_{2} + S_{2(g)} \quad (1)$$

At 400 °C, carbonates become stable, as the destabilisation of pyrrhotite and quartz form hedenbergite and siderite according the equilibrium (2) (Fig. 7d):

$$2qz + 2po + cal + O_{2(g)} = hed + sid + S_{2(g)} \quad (2)$$

This equilibrium is consistent with the phase assemblage observed in the inner corona, which shows abundant pyroxene and siderite, while quartz is absent (Fig. 4).

Under similar conditions, pyrrhotite can also be dissolved in the presence of CO₂ and O₂ to produce siderite and release S₂ through reaction (3) (Fig. 7d):

$$2po + 2CO_{2} + O_{2} = 2sid + S_{2} \quad (3)$$

Equilibrium (3) can explain the formation of the outer corona that is composed of siderite without pyroxene and where quartz is not dissolved (Fig. 4).

The position of equilibrium (2) appears to be almost insensitive to changes in temperature (Fig. 7). In contrast, equilibrium (3) is strongly temperature-dependent and shifts towards more reducing conditions (i.e. lower chemical potential of O₂), and a high sulfidation (Fig. 7), when temperature drop from 600 °C to 400 °C and finally 200 °C. This evolution on phase diagram topology suggests that during a temperature decrease from 600 to 200 °C, the system evolves towards more reducing conditions driven by the reaction (2) consuming O₂ (and/or higher chemical potential of S₂). At this stage, O₂ and S₂ chemical potential conditions are buffered by equilibrium (3). Similar reactions
<figure><img src="image_6.png" /><figcaption>**Fig. 4** a Totally recrystallized metasediments below the sill (431.00 mbsf), showing patches of carbonates around pyrrhotite below the sill. (b) Reflected light image of the metasediments showing the distribution of the patches of carbonates into the quartz-rich matrix. c-e) SEM backscattered image of the carbonate patches below the sill showing a first corona rim (in red) of calciteª, sideriteᵇ, pyroxene with a dissolved</figcaption></figure>

pyrrhotite at its edges; the quartz matrix is completely absent from this rim. A second corona rim (in blue) with calciteᵇ, sideriteᵇ and undisolved quartz is also present. Mineral abbreviations: : *Qzt* quartz (dark grey), *Po*: pyrrhotite (white), *Cal* calcite (medium grey), *Sd* siderite (light grey), *Px* pyroxene (medium grey)
<table><caption>**Table 3** Average compositions of pyroxenes and carbonates in the carbonate patches in the lower aureole measured by microprobe EPMA. Calcitea and sideritea corresponds to the siderite carbonates in the inner corona close to the pyrrhotite cores, and calciteb-siderite are in the outer corona around the pyrrhotites (see Fig. 4). The structural formulae were calculated on the basis of 6 oxygens for the pyroxenes and 3 oxygens for the carbonates. The detailed oxides and sum of the analyses are available Table SI4</caption><thead><tr><th></th><th>Si<br>(Cation)</th><th>Al<br>(Cation)</th><th>Ca<br>(Cation)</th><th>Mg<br>(Cation)</th><th>Fe<br>(Cation)</th><th>Mn<br>(Cation)</th><th>Na<br>(Cation)</th><th>Ti<br>(Cation)</th><th>Cr<br>(Cation)</th><th>Sr<br>(Cation)</th><th>Ba<br>(Cation)</th><th>number of<br>analyses</th><th>Structural formula</th></tr></thead><tbody><tr><td colspan="13"><strong>Inner corona</strong></td></tr><tr><td>Pyroxene</td><td>1.95</td><td>0.03</td><td>0.97</td><td>0.65</td><td>0.41</td><td>0.01</td><td>0.026</td><td>0.004</td><td>0</td><td></td><td></td><td>7</td><td>[Na<sub>0.026</sub>Ca<sub>0.97</sub>Mg<sub>0.65</sub>Fe<sub>0.41</sub>][Si<sub>1.95</sub>Al<sub>0.03</sub>O<sub>6</sub>]</td></tr><tr><td>Calcite<sup>a</sup></td><td>0.00</td><td></td><td>0.98</td><td>0.01</td><td>0.01</td><td>0.00</td><td></td><td></td><td></td><td>0.00</td><td>0.00</td><td>48</td><td>Ca<sub>0.98</sub>Fe<sub>0.01</sub>Mg<sub>0.01</sub>(CO<sub>3</sub>)</td></tr><tr><td>Siderite<sup>a</sup></td><td>0.06</td><td></td><td>0.12</td><td>0.30</td><td>0.46</td><td>0.00</td><td></td><td></td><td></td><td>0.00</td><td>0.01</td><td>27</td><td>Ca<sub>0.12</sub>Fe<sub>0.46</sub>Mg<sub>0.30</sub>(CO<sub>3</sub>)</td></tr><tr><td colspan="13"><strong>Outer corona</strong></td></tr><tr><td>Calcite<sup>b</sup></td><td>0.01</td><td></td><td>0.96</td><td>0.01</td><td>0.01</td><td>0.00</td><td></td><td></td><td></td><td>0.00</td><td>0.00</td><td>48</td><td>Ca<sub>0.96</sub>Fe<sub>0.01</sub>Mg<sub>0.01</sub>(CO<sub>3</sub>)</td></tr><tr><td>Siderite<sup>b</sup></td><td>0.01</td><td></td><td>0.19</td><td>0.30</td><td>0.49</td><td>0.00</td><td></td><td></td><td></td><td>0.00</td><td>0.01</td><td>31</td><td>Ca<sub>0.19</sub>Fe<sub>0.49</sub>Mg<sub>0.30</sub>(CO<sub>3</sub>)</td></tr></tbody></table>

implying CO₂ and siderite in the field of pyrite, iron, magnetite and hematite are observed (reaction (6), (9), (17) and (19); Fig. SI2).

Equilibrium (4) shows the relationships between the fugacities of O₂ and S₂ and activities of H₂O and H₂S in the fluid:

$$2H_2S + O_2 = 2H_2O + S_2 \quad (4)$$

Assuming that the activity of H₂O(l) is 1, the equilibrium constant K of the reaction (4) can be expressed as:

$$\log K = \log f O_2 + 2 \log a H_2S - \log f S_2$$

The equation can be rearranged to:

$$\log \left( \frac{f S_2}{f O_2} \right) = 2 \log a H_2S - \log K$$

Thus, μO₂ and μS₂ (μ and

f

are related, by μᵢ = μᵢ° + RT ln

fᵢ

) can develop in different directions simply by varying the amount of dissolved H₂S. Combining equilibrium (4) with (2) and (3) gives equations (5) and (6) :

$$2qz + 2po + cal + 2H_2O_{(aq)} = hed + sid + 2H_2S_{(aq)} \quad (5)$$

$$2po + 2CO_2 + 2H_2O_{(aq)} = 2sid + 2H_2S_{(aq)} \quad (6)$$

These relations show that a loss of H₂S will drive both reactions from left to right.

## Quantification of carbon sequestration

Isochemical potential phase diagrams have been computed at the same temperature conditions, including Mg into the system and using the metasediment chemical composition (Fig. 8) (table SI4). The aim is to quantify the amount of carbonates, which is expected to precipitate by equilibrium (2) and (3) with decreasing temperature. To determine the most appropriate conditions consistent with the observed assemblages in the coronitic texture, phase diagram has been contoured for S₂ content (Fig. 8). The sediments have a composition in sulphur ranging from 0 to 2 wt%. The area depicted in red in Fig. 8d and f, corresponds to the formation conditions of the inner and outer coronas, respectively.
<figure><img src="image_7.png" /><figcaption>**Fig. 5** Evolution of the δ¹⁸O of the SiO₂ polymorphs (quartz, opal-A, opal-CT; square), δ¹³C (triangle) and δ¹⁸O (circle) of carbonate thought with depth (mbsf). Orange for unmetamorphosed Sediments sediments are in orange; green and light blue for metasediments (upper and lower aureole, respectively) and, dark blue for (dolomite nodule) for upper metamorphic aureole and light blue for lower aureole. The δ¹⁸O of the quartz metasediments showing a decrease near the sill due to increase of temperature and the impact of metamorphic</figcaption></figure>

fluid. The carbonates in the sediments have a typical marine carbonate signature and there is a decrease of δ¹⁸O for the carbonates in the metamorphic aureole due to the increase of temperature. Carbonates below the sill have a totally different signature which can be due to the impact of the metamorphic fluid trapped below the sill. Grey areas represent the metamorphic aureoles; their boundary is based on the appearance of quartz

At 600 °C, less than 1 wt% of carbonates are stable due to the high temperature and the sulphur is mostly in the fluid. As temperature decreases to 400 °C, the predicted conditions become consistent with the observed mineralogical assemblage in the inner corona, consisting of quartz-pyroxene-pyrrhotite-calcite in equilibrium with a fluid rich in CO₂, H₂S and H₂O (Fig. 8c). For these conditions, the maximum amount of calcite does not exceed 5%, suggesting that carbon sequestration is inefficient. At lower temperature (200 °C), the predicted stable assemblage consists of quartz-calcite-ankerite-pyrrhotite in equilibrium with two immiscible fluids: a H₂O-dominated fluid and a COHS fluid representing a mixture of CO₂-H₂S-H₂O, the separation of such two fluids probably produced by fluid boiling. At 200 °C, the carbonate composition predicted by the model corresponds to ankerite (Ca(Fe, Mg)(CO₃)₂, which is slightly different from the siderite (Ca₀.₁₂Mg₀.₃Fe₀.₄₆CO₃) found in the corona. This is probably due to the excess of Ca in the calculated models. The amount of carbonates (calcite

and ankerite) that precipitate is significantly higher than at 600–400 °C and reaches 15 to 17 wt% making it a viable mechanism to store carbon during cooling (Fig. 8b).

## Discussion

### Isotopic evolution of the fluid

In the studied sediments, the initial δ¹⁸O values are dominated by Opal-A/CT at 23.7‰ in the shallowest sediments. The δ¹⁸O value progressively decreases during burial and diagenesis to values around 20‰. This signature remains in strong disequilibrium with the pore water, whose isotopic composition is close to that of seawater (around 0‰). During early diagenesis, the progressive conversion of opal-A to opal-CT and then to quartz is associated with progressive equilibration of δ¹⁸O values of metasediments towards interstitial fluids (Arnórsson et al. 2006; Sheppard 1981).
<figure><img src="image_8.png" /></figure>

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

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

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<figure><img src="image_12.png" /></figure>
**Fig. 6** a Curve of the maximum temperature (Tmax) reached at each depth following sill emplacement, as calculated from thermal modeling using SILLi code (Iyer et al. 2018). b Temporal evolution of the temperature at different distances below the sill (1 m, 6 m, and 16 m), showing a sharp temperature peak followed by a prolonged slow cooling period. The peak of contact metamorphism (Qtz+Po+fluid) is constrained by previous petrographic observations (red line, Cheviet et al. 2023), while retrograde metamorphism associated with Po destabilization and carbonate precipitation is recorded during the later cooling phase (blue line, this study). c δ¹⁸O values of fluids in equilibrium with quartz at Tmax. (d) δ¹⁸O values of fluids in equilibrium with carbonates, estimated at 600 °C (open circles), 400 °C (open diamonds) and 200 °C (open squares). Precipitation temperature of metamorphic quartz in equilibrium with metamorphic fluid in metamorphic aureoles and seawater at δ¹⁸O = 0‰ for sediments (see text). Grey areas represent metamorphic aureoles, their boundary is based on the appearance of quartz. Dashed line represent the actual geothermal gradient. (b) Precipitation temperature of carbonates in equilibrium with metamorphic fluid in metamorphic aureoles and seawater at δ¹⁸O = 0‰ for sediments

During the intrusion of the sill the increase of temperature in the surrounding sediments induces metamorphic dehydration reactions, organic matter cracking, and destabilization of sedimentary carbonates and pyrites. Consequently, water-rock ratio increases, and the fluid becomes a mixture of initial pore water and additional volatiles released by metamorphism. At high temperature, the oxygen isotope exchange between minerals and interstitial fluids becomes more efficient (Clayton et al. 1972; Sheppard 1981), driving the system towards equilibrium. The δ¹⁸O values of quartz decreases systematically, from 13.4 to 10.2‰, while temperature increases approaching the sill. This temperature-dependence interpretation of quartz δ¹⁸O isotopic composition is supported by most of the data, except for the two samples the closest to the sill. They show anomalously high δ¹⁸O values around 11–13‰ that deserve a discussion. Two distinct hypothesis are proposed. (1) basaltic sill emplacement and contact metamorphism may have reduced permeability of the sediments located just below the sill due to compaction and metamorphism (Galerne and Hasenclever 2019). This newly formed low permeability horizon may have restricted fluid/rock interactions and therefore limited isotopic re-equilibration between the pore water (i.e. see water at 0‰) and the sediments. (2) Alternatively, the system may have been locally open, allowing some of the metamorphic fluid to escape upward to the sill (Galerne and Hasenclever 2019). Similarly to the previous hypothesis, interactions between the underlying sediments and the metamorphic fluid would be limited. In such a scenario, quartz may have retained a δ¹⁸O signature closer to its pre-intrusion sedimentary value, with limited influence from low δ¹⁸O porewater.

In the 600–200 °C temperature range, the calculated δ¹⁸O values for fluids in equilibrium with carbonates span a wide compositional range (Fig. 6). Below 400 °C, the calculated

δ¹⁸O values of these fluids are comparable to those inferred for the quartz-hosting metamorphic fluids. This supports the hypothesis that the carbonates crystallized during the retrograde evolution of the system and are isotopically continuous with the quartz. The observed variability in carbonate δ¹⁸O values can be attributed to several processes, including fluid fractionation during progressive mineral crystallization, heterogeneous fluid migration within the aureole, or boiling that fractionate oxygen isotopes (Arnórsson et al., 2007). However, due to the bulk nature of our analyses, these processes cannot be discriminated at the scale of individual mineral phases.

## Sources of carbon in the metamorphic aureoles

In the Guaymas Basin, four possible sources of carbon in the metamorphic aureole can be inferred: (1) marine carbonates with an average δ¹³C near 0‰ (Keith and Weber 1964a, b), (2) organic carbon from sediments with a δ¹³C value around –20‰ (Gerlach and Taylor 1990a, b), (3) dissolved carbon in the porewater at 11‰ (Torres and Kim 2022a, b) and (4) carbon from magma degassing between –5 to –3‰ (Simoneit 1984; Simoneit et al. 1979a, b). Carbonates resulting from OM degradation display negative values ranging from –13.9 to –9.6‰ in a hydrothermal chimney in the Southern Trough of the Guaymas Basin (Peter and Shanks 1992a, b). At Site U1546, the sediments contain IC, OM and dissolved carbon in porewater, and the resulting metamorphic carbonates have δ¹³C values around 1‰ (Fig. 5). The initial IC in carbonates in sediment (about 0.13–0.48 wt%) with a δ¹³C value around 0‰ (Fig. 5) and the dissolved carbon (~3 mM) with a positive δ¹³C value averaging 11‰ (Torres and Kim 2022a, b) are not sufficient to buffer the light carbon from the cracking of OM, which produced the highest amount of carbon in the sediment (2.1 wt%). It can be hypothesised that during the cracking of OM the main carbon species are CO₂ and CH₄ and during methanogenesis carbon fractions making CO₂ molecules with a negative δ¹³C value close to –5‰ and CH₄ with a δ¹³C around –30 to –40‰ (Bottinga 1969; Chacko et al. 2001; Craig 1953; Kueter et al. 2019). Carbon isotope fractionation studies between CO₂/CH₄ are done on single molecules and not kerogen, which can modify the fractionation and produce CO₂ with a δ¹³C value close to 0‰ (Bertassoli et al. 2016; Kelley and Früh-Green 1999). In addition, carbon can be fractionated during calcite and siderite precipitation increasing the carbonate δ¹³C content. The combination of IC, dissolved carbon in porewater and CO₂ from cracking of organic matter may explain the positive δ¹³C values of carbonates. There is no evidence of fluid transfer from the magma to the metasediments. However, during the IODP 385 expedition, carbonates veins at the bottom of the sill were described (Teske et al. 2021a,
<figure><img src="image_13.png" /></figure>

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

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

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

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

<figure><img src="image_18.png" /></figure>
Fig. 7 Quantitative Qualitative µS2-µO2 projections calculated at 600 °C (a-b), 400 °C (c-d) and 200 °C (e-f) at 200 bar. The solid lines correspond to the univariant reactions in the subsystem Fe-S₂-O₂ system, whereas the dashed lines are the univariant reactions in the system Si-Ca-Fe-H₂-C-S₂-O₂ system. Reactions (1) and (2) correspond to the reactions/equilibria used in the text. Phase relation in these simplified systems show the reactions involving silicates, carbonates and sulphides in the presence of a fluid as a function of variations of µS2-µO2. The complete projection with all the reactions is given in the appendix (Fig. SI2). Mineral abbreviations: *qz*:quartz, *hed*:hedenbergite, *ftr*:Fe-tremolite, *tr*:tremolite, *po*:pyrrhotite, *cal*:calcite, *sid*:siderite, *mag*:magnetite

b, c, d), suggesting a fluid transfer from the metamorphic lower aureole toward the sill. The isotopic data support the hypothesis of a closed system in which the organic carbon, the inorganic carbon in the marine carbonates and the dissolved carbon in the porewater within the sediments prior to sill emplacement are locally redistributed, without any external carbon input.

Carbon sequestration in contact metamorphic aureoles

Petrographic observations, geochemical analyses and modelling results all suggest that carbon can be sequestered in the metamorphic aureoles by precipitation of a calcite and Mg-rich siderite assemblage during cooling and the transformation of the host sediment into metasediment after sill emplacement. Due to the sill geometry and the low degree of induration of the sediments, it can be assumed that the fluids, which are rich in water, methane, CO₂ and sulphur produced by the metamorphic reactions in the upper aureole, can be extracted and transferred to the seafloor through hydrothermal vents (Teske et al. 2021a). In this open system, retrograde fluid-rock interactions that trap COHS compounds will be limited (Fig. 9). Conversely, the sill can be considered as an impermeable barrier preventing the ascent of metamorphic fluids produced in the lower metamorphic aureole (Galerne and Hasenclever 2019). In this case, retrograde fluid-rock interactions during cooling of the sill occur in a quasi-closed system (Fig. 9).

Sulphides play an important role in these reactions. Pyrite formed during burial diagenesis is first transformed into pyrrhotite during contact metamorphism. In the presence of water and carbon from the organic matter (OM, commonly represented by CH₂O), pyrite is destabilised via reaction (5) (Cheviet et al. 2023; Yallup et al. 2013):

$$H_2O + CH_2O + 2FeS_2 = CO_2 + 2H_2S + 2FeS \quad (7)$$

leading to the release of H₂S and CO₂. Subsequent pyrrhotite breakdown is controlled by reaction (2'), which produces pyroxenes, H₂S, and minor carbonates. Upon cooling, the CO₂ released from OM-breakdown reacts with leftover

pyrrhotite to form siderite (reaction 2'). The carbon from OM is oxidised by the pyrite breakdown Eq. (5) favouring the carbonates stability. All these reactions are favoured by the removal of H₂S from the aqueous solution.

The models show that at high temperature (600 °C) the reactions cannot proceed, while, when temperature reaches 400 °C, the first reactions implying carbonates are observed. The absence of a high temperature assemblage can be explained by several observations. First, the heat transfer models overestimate the temperature in the aureoles because they do not consider the heat convection, which is a better process to transfer heat compared to only conduction (Aarnes et al., 2011a; Iyer et al. 2018). Second, heat is rapidly removed from the metamorphic aureoles, reaching Tmax in the entire aureole in only 150 years, so that the kinetics of the reactions can limit the high temperature assemblage if the Tmax was really reached.

With decreasing temperature, separation of a single hydrothermal fluid into an H₂O-rich fluid and a CO₂-H₂S-H₂O-rich fluid by boiling is predicted. The boiling of hydrothermal fluids occur at a temperature range around 200–300 °C creating a fluid and vapor phase (Arnórsson et al., 2007; Heinrich, 2007). The H₂O(v) compound from the H₂O-CO₂-H₂S fluid is separated from the initial fluid, leading to a rise of sulphur and carbon in the residual fluid (Fig. 8). Such H₂O removal may trigger the pyrrhotite forming reaction, which started at higher temperature. Efficient H₂S removal into an immiscible CO₂/H₂S-rich fluid at lower temperature will greatly facilitate the formation of hedenbergite and siderite, which would otherwise cease after limited reaction progress once the levels of dissolved H₂S have reached equilibrium. The texture of the metasediments with pyrrhotite-carbonate patches within a clay-rich matrix is consistent with fluid unmixing and suggests that there may be local sub-systems, where fluid exchanges can take place between the corona and the matrix.

The upper aureole, characterized by fluid loss and the absence of carbonate precipitation, evolved under open system conditions. In contrast, the reactions observed in the metasediments below the sill likely occurred in a quasi-closed system and did not require the influx of an externally-derived fluid. The OM was an internal source for CO₂ enrichment, and H₂S was released by the breakdown of pyrite. Fluid unmixing removed dissolved H₂S from the aqueous fluid and allowed the carbonate-forming reactions to progress to the large extent observed. Our closed system model contrasts with suggestions from other studies (Aarnes et al. 2010, 2011a, b; Iyer et al. 2018; Svensen et al. 2007, 2009), where all the fluids produced during contact metamorphism and cooling are assumed to be transferred to the surface of the basin. Our quasi-closed system assumption is consistent with 2D fluid flow models, which show that even
<figure><img src="image_19.png" /></figure>

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<figure><img src="image_24.png" /></figure>
Fig. 8 Quantitative µS2-µO2 projections calculated at 600°C, 400°C and 200°C at 200 bar for the composition of the metasediment’s composition (table SI4). Blue lines correspond to the total carbonates content (sum of calcite, siderite and magnesite) in the system (a, c and e) (sum of calcite, siderite and magnesite). The red area corresponds to the sulphur content range of these sediments including minerals and fluid. Subplots b, d and f contains display the isopleths of S2wt% which help to construct the red area. Mineralogical assemblages in bold are the stable assemblage for the S2 range and observed in the thin sections. Grey lines represent immiscible reaction involving fluids. Mineral abbreviations: qz : quartz; po : pyrrhotite; cal : calcite; sid : siderite; ank : ankerite; mag : magnetite; cpx : solid solution of clinopyroxene (Table 1)

at peak metamorphism, advection does not allow to drain hydrothermal fluids from underneath the flat parts of the sill intrusions (Galerne and Hasenclever 2019; Iyer et al. 2013).

Carbon and sulphur storage in metamorphic aureoles

Contact metamorphic aureoles are commonly considered as settings where volatiles ($H_2O$, $CO_2$ and $H_2S$) are released via dehydration, decarbonation and desulphurization reactions (Aarnes et al. 2010; Jones et al. 2016). However, our models show the possibility of trapping a significant portion of these volatiles in the lower metamorphic aureole. The dissolved $CO_2$ precipitates in the form of carbonates, up to 15–17 wt% of the bulk rock; Fig. 8). The results of the thermodynamic models are consistent with the increase of carbonates below the sill (Fig. 2). It is not excluded that a fraction of these volatiles may also enter into the sill (Ganino and Arndt 2009; Iacono Marziano et al. 2007; Iacono-Marziano et al. 2009; Lenhardt et al. 2023; Svensen et al. 2004, 2007, 2009).

The quantity of carbon trapped in the carbonates was calculated by taking the concentration in wt% of carbonates $m_{carbonate}$ (maximum 17 wt%), the average density $d$ of 2960 kg.m$^{-3}$ (80% of calcite at 2710 kg.m$^{-3}$ and 20% of siderite at 3960 kg.m$^{-3}$), a thickness $h$ of 6 m under the sill, where the carbonate and pyrrhotite are observed and the lateral extension $x$ of the U1546 sill estimated at 4.2 km$^2$ following Eq. (8):

$$mc = \frac{m_{carbonate} \times d \times h \times x}{1000} \qquad (8)$$

The amount of carbon released in the metamorphic aureoles from the studied site within the Guaymas basin, was estimated at 3.29 Mt for both aureoles including 2.19 Mt of carbon for the lower metamorphic aureole (Lizarralde et al. 2023). If we consider the 6 m-thick layer containing calcite-siderite patches in the lower metamorphic aureole that presents a surface of 4.2 km$^2$ (Lizarralde et al. 2023; Teske et al. 2021a) and a maximum carbonate concentration

of 17 wt%, the maximum amount of carbon retained below the sill is 1.26 Mt (Fig. 9). The amount of carbon that can be stored, in this quasi-closed system metamorphic process, is far from being negligible, since it represents up to 38% of the total carbon released during contact metamorphism. The quantity of carbon that can potentially be transmitted to the surface of the ocean floor is 2.03 Mt for this sill (Fig. 9). The stoichiometry of the reactions (2') and (3') controls the amount of dissolved pyrrhotite and precipitated carbonates, which is consistent with the decrease in TS near the sill and the increase in IC (Fig. 2). The models give a maximum estimate of carbonate precipitation, but factors such as reaction kinetics, or availability of sulphur can limit the reactions and carbon sequestration.

Impact of carbon storage on the global carbon cycle

Sedimentary basins with magmatic activity have played an important role in major changes in the atmospheric carbon cycle (Aarnes et al. 2010; Jones et al. 2016; Svensen et al. 2007, 2009). It has been widely demonstrated that, the Siberian Traps at the end of Permian (Heydari et al. 2008; Retallack and Jahren 2008; Svensen et al. 2009), the Central Atlantic Magmatic Province at the end of Triassic (Blackburn et al., 2013; Jones et al. 2016), the Karoo Basin during the Toarcian (Aarnes et al. 2010, 2011b; Burgess et al. 2015; Galerne and Hasenclever 2019; Heimdal et al. 2021; Lenhardt et al. 2023) and, the North Atlantic Igneous Province at the Paleocene-Eocene Thermal Maximum (Svensen et al. 2004), contributed to abrupt climate changes and associated mass extinctions. This study carried out in the Guaymas Basin is one of the rare studies that has considered and quantified metamorphic carbon sequestration. Aarnes et al. (2011) highlighted the possibility of carbonate precipitation during the development of metamorphic aureoles in the Karoo Basin, but this has not been quantified. The mechanisms and reactions proposed here have very little dependency upon pressure (Fig. SI2). In addition, reaction (3) which traps $CO_2$ in the form of siderite has analogues in the stability fields of other sulphides and oxides. Therefore, we suggest that the model developed here may be applicable to deeper sedimentary basins with different $\mu O_2$ and $\mu S_2$. Moreover, in these basins, unlike in the Guaymas Basin, the sills are emplaced at much greater depths in consolidated rocks with lower porosity and permeability (Ishii et al. 2011). Degassing is considered to be restricted to the upper part of the metamorphic aureole, whereas thermogenic gas generated within the lower aureole either remains trapped in situ or migrates into the sill (Galerne and Hasenclever 2019; Lenhardt et al. 2023; Rabbel et al. 2023). A part of the fluids do not
<figure><img src="image_25.png" /><figcaption>**Fig. 9** Model of fluid circulation within the upper and lower metamorphic aureoles. Fluids above the sill can migrate via hydrothermal vents to the seafloor, whereas below the sill they are trapped. The quantities of carbon released during contact metamorphism come from</figcaption></figure>

Lizarralde et al. (2023), the quantity trapped is calculated based on the thermodynamic models of this study thus reducing the quantity of carbon emitted to the surface of the ocean floor
have sufficient time to migrate in the sill (or evacuate) and are therefore trapped in the lower aureole (Rabbel et al., 2023). Although the fluid pressure could favour localised fracturing (Aarnes et al. 2015; Svensen et al., 2006), we speculate that the drainage of the fluids in this context may be limited, favouring carbonate precipitation under quasi-closed system conditions.

The emplacement of sills within the basin is controlled by the opal-A to opal-CT transition occurring around 325 mbsf. Once the ascending magma encounters the unconsolidated sediments (opal-A), its vertical migration is impeded, leading to the sill formation (Aiello et al. 2024; Galerne et al. 2025). This means that the sediments from which the lower metamorphic aureole has developed in indurated opal-CT rich sediment, making it possible to compare the observations made in the Guaymas Basin with other rifting basins or LIPs. Therefore, this model can be extended to other setting, in order to estimate the amount of carbon that could be directly sequestered through contact metamorphism. It was calculated that the Vøring and Møre basins in the Norwegian Sea have released an average of 1540 Gt of carbon (Aarnes et al. 2010), 23,080 Gt in the Karoo (Lenhardt et al. 2023) Basin and 29,250 Gt in the Tunguska basin (Aarnes et al. 2010). Applying the 38% rate of carbon trapped below the sill calculated for Guaymas Basin would induce the sequestration of 585 Gt of carbon in the Vøring and Møre basins, 8770 Gt in the Karoo basin and 11,115 Gt in the Tunguska basin. Of course, this estimated carbon trapping rate may significantly vary across different basins as a function of lithology and redox conditions, with highly reducing environments potentially inhibiting carbonate precipitation. Same reactions can occur in the stability fields of magnetite, pyrite, hematite, iron (Fig. 5, detailed reaction are described in Fig. SI2) indicating that these reactions do not strictly depend on the lithology and can be extended to other settings such as those met in the Tunguska basin where sulphate-carbonate-fluid interactions occur in the metamorphic aureoles (Kontorovich et al. 1997). Despite these limitations, this study provides new insight into the processes of carbon liberation and sequestration in these basins, emphasizing their potential impact on the global carbon cycle and climate regulation throughout geological history.

## Conclusion

This study highlights the significant role of fluid-rich contact metamorphic aureoles in the sequestration of carbon during sill emplacement in sedimentary basins. Thermodynamic modelling and petrographic observations suggest that carbonates such as calcite and siderite can precipitate in quasi-closed system conditions,

trapping up to 38% of the carbon released during contact metamorphism. This process challenges the traditional vision of open-system models, where all released fluids are assumed to escape toward the surface. If similar processes have occurred in other sedimentary basins, we estimate that a substantial amount of carbon could be sequestered globally, underscoring the importance of metamorphic carbon storage in balancing the carbon cycle. Since there is evidence of magma-sediment-fluid interaction in the sills, it is possible that fluid migration into the magma further reduce the carbon budget. Though this aspect was not investigated here, this study provides crucial insights into the overlooked mechanisms of carbon sequestration in sedimentary basins with magmatic activity, emphasizing their potential impact on the global carbon cycle and climate regulation throughout geological history.

**Supplementary Information** The online version contains supplementary material available at https://doi.org/10.1007/s00410-0-25-02262-0.

**Acknowledgements** This research used samples and data provided by the International Ocean Discovery Program (IODP). The authors thank IODP France and the Region Bourgogne-Franche-Comté for their financial support. This study was supported by the German Research Foundation (DFG) funding priority program SPP 527 through the Grant number 447431016 (C.Y.G.). The project received the support of the technological platforms PEAT2 (LCE) and MIMEMTO (FEMTO). Thanks to François Baudin at Sorbonne University for his help on Rock-Eval analysis and Benita Putlitz from the University of Lausanne for her help on isotope measurement on silicates. We would also like to thank the editor and the reviewers, Henrik H. Svensen and the other two anonymous reviewers, for their contributions and comments which helped to improve the article.

**Author contributions** All authors contributed to the study conception and design. Material preparation, data collection and analysis were performed by Alban Cheviet, Martine Buatier and Flavien Choulet contributed on the petrological description, Torsten Vennemann and Wolfgang Bach contributed on the isotopic measurement and interpretation, Christophe Galerne and Armelle Riboulleau contributed on the carbon and sulphur analysis in sediments, Philippe Goncalves contributed on the models design and interpretation. Alban Cheviet and Christophe Galerne contributed on the parametrisation of the SiLLi model. The first draft of the manuscript was written by Alban Cheviet and all authors commented on previous versions of the manuscript. All authors read and approved the final manuscript.

**Funding** Open Access funding enabled and organized by Projekt DEAL.

**Data availability** The authors declare that all data that support the findings of this study are available within the paper and its supplementary data (Tables SI1-SI4 and Figure SI1-SI2).
## Declarations

**Conflict of 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.

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