# Mineralogical Changes and H2 Generation Yield During Hydrothermal Alteration of a Magnetite-Siderite Assemblage

U. Geymond, L. Truche, O. Sissmann, D. Kubániová, N. Recham, I. Martinez

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## To cite this version:

U. Geymond, L. Truche, O. Sissmann, D. Kubániová, N. Recham, et al., Mineralogical Changes and H2 Generation Yield During Hydrothermal Alteration of a Magnetite-Siderite Assemblage. Journal of Geophysical Research: Solid Earth, 2025, 130 (8), pp.e2024JB030724. 10.1029/2024JB030724. hal-05212982

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Submitted on 19 Aug 2025

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Distributed under a Creative Commons Attribution 4.0 International License
# JGR Solid Earth

## RESEARCH ARTICLE

10.1029/2024JB030724

**Special Collection:**

Solid Earth Geophysics as a means to address issues of global change

**Key Points:**

* We study mineralogical changes and H₂ generation resulting from hydrothermal alteration of a siderite-magnetite assemblage at 200°C

* Non-stoichiometric magnetite traps dissolved Fe²⁺ to recover its expected stoichiometry, according to experimental results

* H₂ generation yields are mainly constrained by the fluid-rock ratio, according to thermodynamic modeling

**Supporting Information:**

Supporting Information may be found in the online version of this article.

**Correspondence to:**

U. Geymond, geymond@ipgp.fr

**Citation:**

Geymond, U., Truche, L., Sissmann, O., Kubániová, D., Recham, N., & Martinez, I. (2025). Mineralogical changes and H₂ generation yield during hydrothermal alteration of a magnetite-siderite assemblage. *Journal of Geophysical Research: Solid Earth, 130*, e2024JB030724. https://doi.org/10.1029/2024JB030724

Received 17 NOV 2024
Accepted 28 JUL 2025

**Author Contributions:**

**Conceptualization:** U. Geymond, L. Truche, O. Sissmann, I. Martinez

**Data curation:** U. Geymond

**Formal analysis:** U. Geymond, L. Truche, O. Sissmann, D. Kubániová, N. Recham, I. Martinez

**Funding acquisition:** L. Truche, D. Kubániová, I. Martinez

**Investigation:** U. Geymond, L. Truche, O. Sissmann

© 2025. The Author(s). This is an open access article under the terms of the Creative Commons Attribution License, which permits use, distribution and reproduction in any medium, provided the original work is properly cited.

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## Mineralogical Changes and H₂ Generation Yield During Hydrothermal Alteration of a Magnetite-Siderite Assemblage

U. Geymond¹,², L. Truche³, O. Sissmann², D. Kubániová⁴, N. Recham⁵, and I. Martinez¹

¹Université Paris Cité, Institut de Physique du Globe de Paris, CNRS, Paris, France, ²IFP Energies Nouvelles, Rueil-Malmaison, France, ³Université Grenoble Alpes, ISTERRE, CNRS, Grenoble, France, ⁴Department of Low Temperature Physics, Faculty of Mathematics and Physics, Charles University, Prague, Czech Republic, ⁵Université de Picardie Jules Verne—LRCS, CNRS, Amiens, France

**Abstract** To date, the generation of natural hydrogen (H₂) from the alteration of FeII-bearing minerals has mainly been studied through serpentinization in (ultra)mafic rocks. This study explores Banded Iron Formations (BIF), which are rich in FeII minerals, as a potential source for H₂. We conducted a hydrothermal experiment at 200°C with a combined magnetite-siderite assemblage, two major components of BIF. The experiment, designed with a high water-to-rock ratio and a gas phase (W-R = 300) enabled to assess mineral transformations during the alteration. Thermodynamic simulations were finally conducted, to explore H₂-generating yields in more realistic geological scenarios (no gas phase, lower W-R). Our experimental findings show that H₂ is produced through complete siderite dissolution and magnetite precipitation. Concomitantly, non-stoichiometric primary magnetite (FeII/FeIII < 0.5) did not enhance H₂ yield; instead, it acted as a sink for dissolved Fe²⁺, sequestrating about 10% of the iron from siderite without oxidation to recover a more ideal stoichiometry. This suggests that abundant, non-stoichiometric magnetite in natural settings may reduce H₂ generation yields. Mass balance calculations indicate that 83% of the expected H₂ generated was unaccounted for, consistent with suspected CO₂ reduction and formation of dissolved organic compounds in the fluids. Thermodynamic simulations at varying W-R ratios (from 300 to 1) reveal that H₂ yield ranges widely (39 mmol–73 µmol H₂ per kg of siderite), since lower W-R prevent from siderite dissolution and enhance H₂ consumption via carbon reduction. These findings imply that low W-R ratios in BIF -and other siderite-bearing lithologies, may limit H₂ resources.

**Plain Language Summary** This study explores how ancient iron-rich rocks, called Banded Iron Formations (BIF), might naturally generate hydrogen gas (H₂) through iron oxidation and water reduction, when altered by water in oxygen-free conditions, thus offering insights into potential clean energy sources. Previous research mainly focused on different iron-rich rocks with minerals like olivine and pyroxene. Here, we targeted BIF for their high content of siderite and magnetite, two iron-based minerals potentially involved in H₂ generation. In the lab, we heated these minerals in water at 200°C to track mineral changes and H₂ generation. Results showed that siderite generates H₂ as it dissolves, while magnetite does not; instead, magnetite actually lowers H₂ levels by capturing some of the iron released from siderite, preventing iron to get oxidized and trigger H₂ generation. Additionally, much of the expected H₂ and CO₂ are missing, likely transformed into other carbon-bearing compounds. The extent of this conversion seems to depend on the water-to-rock ratio, with less water favoring reactions that consume H₂. These findings suggest that BIF and similar rocks might have limited hydrogen potential in natural settings where water supply is limited. This research enhances our understanding of natural H₂ generation factors, informing future clean energy exploration.

## 1. Introduction

Understanding how natural H₂ forms within the Earth's crust has become a major field of research, as H₂ might have played a role in the appearance of primitive life on Earth (McCollom & Seewald, 2013; Sherwood Lollar et al., 2007), and may guide the search for life beyond Earth, particularly on extraterrestrial bodies (Shock & Schulte, 1998). In addition, natural H₂ is currently attracting attention to face the growing global demand for H₂—rising from 60 Mt/year in 2000 to 100 Mt/year in 2024, with projections reaching 500 Mt/year by 2050 (IEA, 2024)—as it may be used as a putative energy source and sustain the steel and fertilizing sectors (Greenwald et al., 2024; Zgonnik, 2020). Natural H₂ is now known to occur in specific geological settings, such as (ultra)mafic
**Methodology:** U. Geymond, L. Truche
**Software:** U. Geymond, L. Truche
**Supervision:** L. Truche, O. Sissmann, I. Martinez
**Validation:** L. Truche, O. Sissmann, I. Martinez
**Visualization:** U. Geymond, L. Truche
**Writing – original draft:** U. Geymond, L. Truche
**Writing – review & editing:** U. Geymond

environments, Precambrian crystalline rocks, and occasionally coal bed formations (Lévy et al., 2023; Truche et al., 2020; Zgonnik, 2020). Since the discovery of H₂-rich vents at mid-oceanic ridges, the serpentinization of ultramafic rocks has been recognized as the main H₂ production mechanism in the oceanic crust, as well as in uplifted orogenic and ophiolitic bodies. The hydration and oxidation of ferrous iron (Fe¹¹) contained in mafic minerals such as Fe-olivine (Fe¹¹SiO₄) induce the precipitation of secondary Fe³⁺-bearing minerals along with H₂ generation. It has been demonstrated that Fe³⁺ is mostly incorporated into magnetite (α-Fe³⁺₂Fe¹¹O₄) at temperature above 200°C (Klein et al., 2014). At lower temperatures, magnetite is no longer produced during olivine alteration, but it is suspected to catalyze the serpentinization reaction and H₂ generation by enhancing Fe oxidation at its surface (Mayhew et al., 2013).

Onshore, apart from the serpentinization process that can occur within ophiolitic massifs and exhumed mantellic bodies (Lefeuvre et al., 2021; Truche et al., 2024; Vacquand et al., 2018), few studies exist on H₂ generation originating from the hydration of other Fe-bearing minerals, such as Fe-biotite or Fe-amphibole (Murray et al., 2020; Truche et al., 2021). The H₂ potential of siderite (Fe¹¹CO₃) is worth questioning since H₂ concentration anomalies in soil are reported near siderite-rich lithologies, such as Banded Iron Formations (Geymond et al., 2022; Moretti et al., 2022; Roche et al., 2024) and gold mineralization zones (Malvoisin & Brunet, 2023). Experimental studies conducted at 300°C have demonstrated that water-siderite interactions generate H₂ and CO₂, along with the precipitation of magnetite, according to equation Equation 1 (McCollom, 2003; Milesi et al., 2015).

$$3\text{FeCO}_3 + \text{H}_2\text{O} \leftrightarrow \alpha - \text{Fe}_3\text{O}_4 + 3\text{CO}_2 + \text{H}_2 \quad (1)$$

The presence of reduced carbon compounds by-products was also evidenced in these experimental studies involving siderite hydrothermal alteration at 300°C, occurring as traces of CH₄ in the gas phase, as graphite-like precipitates in the solid phase, and as alkylated aromatic compounds in the fluid phase. It was proposed that the presence of these various forms of reduced carbon compounds could be promoted by magnetite, a presumed catalyst of carbon reduction and hydrogen oxidation (Fu et al., 2007; Horita & Berndt, 1999; Voglesonger et al., 2001). Recently, magnetite oxidation was also suggested to be a primary source of H₂ (Geymond et al., 2023). Indeed, an experimental study aiming to alter synthetic magnetite under anoxic conditions at 80°C and 200°C proposed that the remnant budget of Fe¹¹ carried by magnetite is prone to oxidize and form maghemite (γ-Fe³⁺₂O₃) according to equation Equation 2.

$$2(\alpha - \text{Fe}_3\text{O}_4) + \text{H}_2\text{O} \leftrightarrow 3(\gamma - \text{Fe}_2\text{O}_3) + \text{H}_2 \quad (2)$$

Magnetite and siderite are commonly associated in Banded Iron Formations (BIF) (Ewers & Morris, 1981; Klein, 2005). The H₂ generating potential of these two minerals, as well as the catalytic effect of magnetite, have been investigated separately in a few experimental studies. However, their interdependence in terms of H₂ generation and CO₂ reduction when they are associated have not been tested until now. The objective of this study is to evaluate the reactivity of a magnetite-siderite assemblage and characterize the mineral transitions occurring during the hydrothermal alteration of this simplified system that mimics BIF. The reactivity of a magnetite-siderite assemblage with respect to H₂ generation and consumption was studied (a) experimentally by reacting mineral powders under anoxic condition in a batch reactor at 200°C, and (b) numerically by investigating thermodynamic equilibrium of the involved reactions at various temperatures, and alternatively at 200°C with/without headspace and with varying water-to-rock ratios (W-R).

# 2. Methods

## 2.1. Starting Material

In this study, synthetic magnetite and siderite were chosen to simplify interpretations of the results. For magnetite, a highly pure (99.99% trace metals basis) manufactured material from Sigma-Aldrich® was selected (CAS number: 1317-61-9), which was already used in previous magnetite alteration experiments (Geymond et al., 2023). Siderite was synthesized in-house, by reacting Fe-sulfate (FeSO₄) and urea (CO[NH₂]₂), following a protocol described by Milesi et al. (2015). Subsequently, a starting mineral assemblage was prepared, consisting in a mix of 22 wt% of siderite and 78 wt% of magnetite, determined through quantitative Mössbauer Spectroscopy (MS) analysis, as detailed in Section 3.3, Texts S2 and S3 in Supporting Information S1. The mineral assemblage
<figure><img src="image_4.png" /><figcaption>**Figure 1.** (a) XRD pattern and (b–e) SEM images of starting material obtained in secondary electron mode. (b, c) Large scale view of the round-shaped siderite grains and irregular-shaped magnetite grains. Note the cohesiveness and smoothness of siderite surface (d) compared to the rough and highly porous magnetite texture (e).</figcaption></figure>

was controlled prior to experiments using X-ray diffraction (XRD) and scanning electron microscopy (SEM) to ensure the quality of the starting material (Figure 1). According to SEM images, the starting powder consisted in grains of siderite and magnetite of few tens of micrometers, identified by their respective shapes: sub-spherical for siderite and more chaotic and angular for magnetite. Their surface textures appeared different as well: Dense and smooth for siderite, rough and highly porous for magnetite (Figures 1b–1e). Further details concerning the characterization of the starting material are presented in Section 3.3, along with the characterization of products of alteration. The specific surface area of the starting material was measured following the BET method and using Kr as the adsorbate gas, providing the following estimates: 2.0 ± 0.1 m²/g for magnetite alone and 2.25 ± 0.75 m²/g for the magnetite—siderite assemblage. The uncertainty of the BET measurement of the magnetite-siderite assemblage is relatively high due to the very low amount of material available to perform the analysis.

## 2.2. Batch Reactor Setup

The experiment was conducted at 200°C, in a fully-passivated titanium reactor (Parr®) of 500 mL total volume, constantly stirred with a magnetically driven impeller and equipped with several connections enabling water and gas sampling. The autoclave was filled with 990 mg of the starting powder and 300 mL of milliQ water (conductivity of 18.2 MΩ.cm). The presence of a headspace and the very high W-R likely provided unrealistic geological conditions by imposing the system to remain far from equilibrium, but offered comfortable conditions to characterize products of alteration due to increased yields of reaction.
The autoclave was sealed and flushed with argon during 20 min to remove O₂ within the system. The hydrothermal experiment was then launched for 40 days, and was stopped when no more significant change in H₂ level was observed, inferring no more on-going mineral reaction. Initial absolute pressure of 3.5 bar of Ar imposed in the headspace at the beginning of the experiment rapidly reached up 16.5 bar due to the water saturation pressure at 200°C.

The gas phase was sampled over the course of the experiment, to monitor changes in gas composition through time. At the end of the experiment, aqueous fluids were sampled to measure the pH and later investigate the presence of dissolved iron as well as any organic compounds. Temperature and pressure were then rapidly decreased in the reactor by quenching it in a water bath at ambient temperature, to prevent from the precipitation of new solid phases that could occur during a slower cooling. Finally, solid products of reaction were retrieved and dried at room temperature in a dry anoxic glove box ([O₂] < 1 ppmv) to explore mineral transformations that could have occurred during the experiment.

A blank experiment was carried out at 200°C over 3 weeks using the same autoclave filled with 300 mL milliQ water only (anoxic condition, Ar flushing and degassing). The gas phase was sampled over time to measure the background level of H₂. Hydrogen was detected after 1 week, but its concentration remained consistently below the quantification limit of GC-TCD (5 ppm). Finally, the background level of H₂ measured in the blank experiment is by at least two orders of magnitude lower that the amount of H₂ measured during the siderite-magnetite experiment.

## 2.3. Analytical Tools

### 2.3.1. Gas and Liquid Phases Analyses

The gas phase was sampled from the reactor headspace using a syringe equipped with a three-way valve, one exit of the valve being equipped with a rubber septum. This design allowed water to condense while sampling and gas to expend in a known volume. The sampled gas was then immediately transferred to a gas chromatograph (GC) by the mean of a Hamilton® gas tight syringe of calibrated volume. Analyses were performed on a Perkin Elmer® Clarus 500 GC, equipped with a thermal conductivity detector (TCD) and a 2-m long column Shin Carbon ST 80/100 from Restek® traversed by Ar as carrier gas. The GC-TCD system was calibrated using certified standard gas mixtures supplied by Air Liquide of known H₂, CO₂, CH₄ and Ar concentrations. All analyzed samples fell within the concentration ranges covered by these standards. The detection limits were 5 ppmv for H₂, 100 ppmv for CH₄, and 500 ppmv for CO₄. The amounts of H₂, CH₄ and CO₂ present in the autoclave were calculated using the ideal gas law. As the water vapor condenses during sampling, the pressure used in the calculations corresponds to the total pressure (measured with the gauge mounted on the autoclave) minus the vapor pressure at 200°C. The volume of the gas phase at 200°C is equal to 155 mL (considering the total volume of the reactor, the initial volume of water and the thermal expansion of water). The overall amount of gas produced accounts for the quantity extracted out of the autoclave at each sampling steps and was corrected from minor air contamination occurring during gas transfer from the autoclave to the GC. More details regarding gas sampling and GC quantifications are provided in Text S1 and Figure S1 of Supporting Information S1.

The fluid phase was sampled only at the end of the experiment to ensure that all experimental parameters (e.g., the W-R ratio) remained unchanged throughout, allowing for more accurate estimates of H₂, CO₂, and CH₄ concentrations in the reactor. Just before quenching the reactor, the fluid sampling valve of the reactor was opened and the hot raw fluid was collected under atmospheric air. The first third was pH-measured right after to limit pH changes due to cooling and redox re-equilibration between the atmosphere and the fluid. The remaining two-third of the fluid was filtered at 0.2 µm (PTFE filter), and diluted either in tenfold in milliQ water or tenfold in HNO₃ (0.2 vol%) to prevent any iron precipitation. The two diluted solution were then stored in falcon tubes in the fridge before analyses. The total organic carbon (TOC) concentration was determined in the milliQ-diluted fluids by the mean of a TOC-analyzer VCP-H from Shimadzu®. Finally, total dissolved iron concentrations (ΣFe) were measured in HNO₃-diluted fluids by mass spectrometry using an Agilent Technologies® 7900 ICP-MS. Due to the sampling under atmospheric air, accurately measuring the respective concentrations of Fe²⁺ and Fe³⁺ would have appeared uncertain. Moreover, since the ΣFe concentration appeared extremely low in the fluid (see results in Section 3.2), no efficient technique could provide an accurate estimate of dissolved Fe²⁺/ΣFe, thus no analysis was conducted in that sense.
### 2.3.2. Solid Phase Analyses

The starting powder as well as the reacted one were observed by scanning electron microscopy (SEM) to assess the surface texture of the grains, and image the precipitation of newly-formed solid phases. A Zeiss® Auriga40 SEM was used for that purpose. It is equipped with a 1 nm resolution FEG as well as a Bruker® Quantax 800 EDS and a Bruker® XFlash 410-M detector. In parallel, room-temperature ⁵⁷Fe Mössbauer spectroscopy (MS) analyses were performed before and after the experiments to study changes in iron speciation. In the following sections, for clarity, Fe¹¹ (and Fe¹³) refer to iron immobilized in the solid phases, whereas Fe²⁺ refers to dissolved iron. The MS acquisition of the spectra was done in a transmission mode with the ⁵⁷Co source diffused into a Rh matrix, moving with a constant acceleration. The velocity calibration of WissEL spectrometer was performed using a standard α-Fe foil and the isomer shifts are given relative to this standard at 296 K. The fitting of the hyperfine parameters, using a standard least-squares fitting routine with pseudo-Voigt spectral lines, was performed with Confit2000 (Žák, 2001) based on spectral models described in detail elsewhere (Chistyakova et al., 2023; Novák et al., 2000). The uncertainty is estimated at ~1 at%. Finally, changes in mineral compositions were determined by means of X-ray diffraction (XRD). Measurements were conducted using a Malvern-Panalytical® Empyrean diffractometer equipped with a copper tube (Kα = 1.541874 Å) and a Malvern-Panalytical® multi-channel PIXcel detector. The chosen angular range was between 5° and 80°, the step size was 0.007°, and the time per step was 80 s (i.e., 1 hr measurement per sample). Diffractograms were then processed using Highscore Plus software from Malvern-Panalytical®.

### 2.3.3. Thermodynamic Modeling

Numerical simulations were conducted using the PhreeqC software (Parkhurst & Appelo, 1999) together with the LLNL thermodynamic database (Johnson et al., 2000). A first series of simulations was conducted to investigate the thermodynamic equilibria of the reactions involved in the experiment and evaluate the effect of varying temperatures. Similar experimental conditions were replicated in these simulations: 300 mL of pure water reacted with 990 mg of mineral powder composed of 78 wt% magnetite and 22 wt% siderite, with a 200 mL headspace to allow gases to exsolve from the water. A second series of simulations was then conducted at 200°C, maintaining all the other parameters constant while varying the W-R and headspace volumes to explore more realistic geological scenarios. All input and output files for both series of simulations are available in a permanent online repository (Geymond, 2025). Given that the results of thermodynamic simulations are highly sensitive to the selected input parameters, it is important to briefly detail and justify these choices:

*   **Fluid:** The pH and pe (i.e., the redox) were treated as free parameters in all the simulations, with pe being adjusted according to redox equilibrium, to account for chemical changes occurring in the system over time. Indicative initial values (pH = 7 and pe = 4) were provided for the starting solution, as expected for an ideal pure water.

*   **Starting material:** All the simulations were designed to provide a first-order representation of the reacted system without introducing unnecessary complexity. For this reason, only stoichiometric magnetite (Fe₃O₄) was used as the starting solid phase, even though the magnetite used in the experiment was slightly non-stoichiometric (see solid material characterization in Section 3.3). This simplification is justified by two main considerations. First, both the starting and final magnetite compositions closely approach the pure magnetite endmember (89 mol% and 94 mol%, respectively). Second, accurately representing the exact stoichiometry of magnetite in thermodynamic models is highly challenging. The thermodynamic properties of the maghemite endmember (i.e., fully oxidized magnetite) have been only sparsely investigated in the literature, and available estimates vary significantly depending on the degree of vacancy ordering in the crystal lattice (see for instance Diakonov, 1998). Furthermore, building such a model requires, for instance, to account for non-ideal mixing behavior between the magnetite and maghemite endmembers, influenced by several factors including particle size (Jungcharoen et al., 2022). For these reasons, non-stoichiometric magnetite was not considered in this work, despite the availability of maghemite data in some databases (e.g., Thermoddem: https://thermoddem.brgm.fr/databases/phreeqc). Effectively modeling the evolution of non-stoichiometric magnetite in natural H₂-rich systems is a valuable objective but would require a dedicated thermodynamic study of its own.

*   **Solid products:** In addition to siderite and magnetite, goethite and hematite were authorized to precipitate in all simulations, in equilibrium with the solution (i.e., with a saturation index equal to zero). This choice was motivated by two main considerations. First, it simplifies the system by considering only the most
<figure><img src="image_5.png" /><figcaption>**Figure 2.** Amount of CO₂, CH₄ and H₂ produced in the gas phase of the autoclave through time at 200°C. Error bars for H₂ are within the size of the symbol. Dotted lines stand for the initial and long-term trends for both H₂ and CO₂ during the experiments. Their intersections correspond to the abrupt change of reaction regime.</figcaption></figure>

thermodynamically stable Fe-oxides (Cornell & Schwertmann, 2003) while excluding metastable phases such as lepidocrocite or ferrihydrite, which were not observed in our experiments (see result in Section 3.3). Second, it is consistent with the limited experimental literature available on hydrothermal alteration of siderite, which reports the formation of hematite and/or goethite between 50 and 250°C, yet under uncontrolled anoxic conditions (Maddin et al., 2024), and magnetite at 200–300°C, under strictly anoxic conditions (Milesi et al., 2015) Finaly, solid carbon (i.e., graphite-like) was considered as a possible alteration product only in the second series of simulations (performed at 200°C with varying headspace volumes and W-R ratios). This choice is justified by pre-existing experimental observations of such kind of precipitate, only reported at T ≥ 200°C (Milesi et al., 2015).

# 3. Results

## 3.1. Gas

A rapid increase of both H₂ and CO₂ amounts in the gaseous phase was observed during the first days of the experiment (Figure 2). After 2–3 days, H₂ level reached a maximum of 0.10 × 10⁻³ mol in the autoclave's headspace and became stable around this value until the end of the experiment at 41 days. In parallel, the generation of CO₂ slowed down until reaching a maximum level of 0.34 × 10⁻³ mol after 10 days. After this time point, the amount of CO₂ in the headspace of the autoclave slowly decreased until the end of the experiment. These changes of trends for both CO₂ and H₂ levels within the gas phase of the autoclave appear temporally correlated at 2–3 days, suggesting at least one common mechanism controlling the fate of these gases. During this experiment, only one analysis revealed the presence of CH₄ at 21 days, in very low amounts estimated at about 7.18 × 10⁻⁷ mol.

## 3.2. Aqueous Solution

At the end of the run, the solution appeared more acidic than the starting one (MilliQ, pH ≈ 6.5), with a measured pH of 5.05 ± 0.1. ICP-MS analyses revealed a very low concentration of dissolved ΣFe of 0.87 ± 0.01 µmol/L. Interestingly, TOC-analyses showed the presence of dissolved organic compounds at a total concentration of 0.73 ± 0.03 mmol/L. However, no dissolved organic compounds analysis was performed on the blank experiment to set the background chemistry and quantify eventual contamination along the entire experimental chain.
<figure><img src="image_6.png" /><figcaption>**Figure 3.** Interpreted (a) XRD patterns and (b) Room temperature MS spectra for both the starting and reacted powders. Siderite is no longer detected after the experiments in the reacted solids.</figcaption></figure>

### 3.3. Solids

#### 3.3.1. Fe Speciation and Mineralogy

XRD patterns showed that the starting powder only contained siderite and magnetite (Figure 3a). This result was confirmed by room-temperature Mössbauer Spectroscopy (MS) analyses. Iron within the starting powder was exclusively distributed between siderite and magnetite, 16 at% and 84 at%, respectively (Figures 3b, Table 1, Table S2 in Supporting Information S1). This corresponds to a 22–78 wt% distribution in the siderite-magnetite assemblage (see Texts S2 and S3 in Supporting Information S1 for the detailed calculations). In magnetite, Fe was distributed between both octahedral sites, as FeII and FeIII, and tetrahedral sites, as FeIII only, following a $[Fe^{II+III}]/(Fe^{III})$ ratio of 60.2/39.8, where [] and () refers to the octahedral and tetrahedral sites, respectively. Here, the ratio $x$, defined in Equation 3 and already introduced in the literature (Gorski & Scherer, 2010; Latta et al., 2012), can be used to evaluate the stoichiometry of magnetite, that is, the abundance of FeII in the mineral lattice. For magnetite in the starting powder, the ratio yielded $x_{start} = 0.431$. This value is far from the expected ratio of $x_{mgt} = 0.5$ for a pure magnetite of ideal formula $(Fe^{III}_{1.00})[Fe^{II}_{1.00}, Fe^{III}_{1.00}]O_4$. This implies that our starting magnetite was partially oxidized and already consisted in a solid solution between pure magnetite and pure maghemite with $x_{mgh} = 0$ of ideal formula $(Fe^{III}_{1.00})[Fe^{II}_{0.00}, Fe^{III}_{1.67}, O_4]$. The symbol [] stands for vacancies and occurs specifically in octahedral cation positions of the mineral lattice to ensure the overall neutrality of the mineral. Such solid solution has been widely described in the literature (Castro et al., 1996; Doriguetto et al., 2003; Jungcharoen et al., 2022). Our

<table><caption>**Table 1**

Room-Temperature MS Results for Analyses Conducted on the Starting Powder and Solid Products After Hydrothermal Alteration at 200°C</caption><thead><tr><th rowspan="3">Sample</th><th colspan="3">Raw data room-temperature MS analyses (at%)</th><th colspan="2">Calculated</th></tr><tr><th rowspan="2">Fe<sup>II</sup><sub>siderite</sub></th><th colspan="2">Fe magnetite</th><th rowspan="2">x</th><th rowspan="2">Fe<sup>2+</sup>/ΣFe<sub>bulk</sub></th></tr><tr><th>[Fe<sup>II+III</sup>]</th><th>(Fe<sup>III</sup>)</th></tr></thead><tbody><tr><td>Starting</td><td>16</td><td>60.2</td><td>39.8</td><td>0.431</td><td>0.413</td></tr><tr><td>Sidmag200</td><td>0</td><td>63.3</td><td>36.7</td><td>0.463</td><td>0.317</td></tr></tbody></table>

Note. For the details on $x$ and $Fe^{2+}/\Sigma Fe_{bulk}$ ratios see Equation 3 and Equation 4.
$x_{start}$ ratio allowed to calculate a structural formula for the starting magnetite following the method detailed in Texts S2 and S3 of Supporting Information S1: ($Fe^{III}_{1.00}$)[$Fe^{II}_{0.89}$, $Fe^{III}_{1.07}$, $\ll 0.04$]O$_4$. It corresponds to a solid solution 11 mol% maghemite–89 mol% magnetite. A second ratio, $Fe^{2+}/\Sigma Fe_{bulk}$, defined in Equation 4 and already introduced in the literature (for instance see Marcaillou et al., 2011), describes the abundance of $Fe^{II}$ in the whole powder. It yields $Fe^{2+}/\Sigma Fe_{bulk,start} = 0.413$ for our starting material.

$$x = \frac{Fe_{mgt}^{II}}{Fe_{mgt}^{III}} = \frac{\frac{1}{2}[Fe]}{\frac{1}{2}[Fe] + (Fe)} \quad (3)$$

$$Fe^{2+}/\Sigma Fe_{bulk} = \frac{Fe_{mgt}^{II} + Fe_{siderite}^{II}}{Fe_{total}} \quad (4)$$

After alteration, the X-ray diffractogram presented no remaining peak corresponding to siderite and all peaks matched magnetite (Figure 3a). Mössbauer spectroscopy analysis confirmed this result with no $Fe_{siderite}$ detected in the altered powder, all Fe being carried by magnetite (Figures 3b, Table 1). Magnetite exhibits a [$Fe^{II+III}$]/($Fe^{III}$) ratio of 63.3/36.7, which corresponds to a stoichiometry of $x_{200} = 0.463$ and a structural formula of ($Fe^{III}_{1.00}$)[$Fe^{II}_{0.94}$, $Fe^{III}_{1.04}$, $\ll 0.02$]O$_4$. Interestingly, the stoichiometry of magnetite after alteration appears closer to the ideal stoichiometry of a pure magnetite than the starting magnetite. The abundance of $Fe^{II}$ in the reacted solids is consistently lower at the end of the experiment with a calculated $Fe^{2+}/\Sigma Fe_{bulk,200} = 0.317$ (compared to 0.413 prior to alteration). Additional details concerning the hyperfine parameters and the MS spectra of the starting assemblage acquired at 4.2 K are presented in Text S4, Figure S2 and Table S1 in Supporting Information S1.

### 3.3.2. Microscopic Observations

After alteration, the textural aspect of the powder remained similar at large scale, with both sub-spherical grains and rough-angular grains (Figure 4a). However, a zoom on the sub-spherical grains reveal that they no longer correspond to siderite. The bi-pyramidal shape of the newly-formed minerals of few micrometers (Figures 4a–4c), labeled Mgt(II), are typical of magnetite (Cornell & Schwertmann, 2003). Since no siderite was observed along with the new magnetite octahedrons, it can be inferred that the whole siderite grains were in situ converted into Mgt(II) during the reaction, thus allowing the aggregate of new magnetite to preserve the specific shape of starting siderite grains. Concerning the reactivity of primary magnetite, labeled Mgt(I), porosity of initial grains appears lower after alteration, associated with an increase of the crystallinity at grain surface (Figure 4d). The nature of the mineral phase filling up the porosity cannot be determined unambiguously. This new texture at the surface of reacted Mgt(I) is homogeneous, suggesting the mineral reactivity has been ubiquitous in the whole powder during the experiment.

## 4. Discussion

### 4.1. Experimental Results Compared to Thermodynamic Numerical Simulations

To help interpret the experimental results, some of the physicochemical parameters studied during the experimental run (pH, dissolved Fe, gas phase composition, and mineralogy) were computed using Phreeqc software as a function of temperature (within the 20–220°C range) assuming global redox equilibrium among all redox couples (Figure 5). To assess any effect of the presence of magnetite in the mineral assemblage, the results of two distinct models are presented: a first one with a starting mineral assemblage made up exclusively of siderite, and another made up of pure magnetite (78 wt%)-siderite (22 wt%) as a simplified representation of our experimental setup. It is interesting to note that whatever the temperature investigated, expected equilibrium of the two models are almost identical, suggesting a very limited impact of magnetite on the final state of the system (Figure 5). The numerical simulations performed show that 91% of the total $H_2$ produced in the system is exsolved in the gas phase at 200°C (64.3 µmol in the gas phase vs. 5.82 µmol in the aqueous phase). Thus, the discussion and mass balance calculations in the following sections only account for $H_2$ in the gas phase and neglect dissolved $H_2$.

The modeling results closely match most of the experimental parameters (pH, dissolved Fe, $H_2$ partial pressure and mineralogy) measured at the end of the 200°C experiment. Notably, XRD and MS results (100 wt% magnetite) align with model predictions (Figure 5c). Similarly, the amount of $H_2$ produced at the end of the experiment, labeled $H_2_{exp}$, is consistent with the predicted $H_2$ generation amount (64.3 µmol) in the gas phase
<figure><img src="image_7.png" /><figcaption>**Figure 4.** SEM images of the solid products of alteration at 200°C performed in secondary electron mode. Sid refers to siderite, Mgt(I) refers to primary magnetite, Mgt(II) refers to secondary magnetite. (a) Large-scale view and (b) zoom on Mgt(II) grains. Note that the whole spherical-shaped siderite grain has been converted to newly-formed magnetite octahedrons. (c) Large-scale view of the contact between two generations of magnetite: altered Mgt(I) and newly-formed Mgt(II). (d) Primary magnetite Mgt(I) after alteration. Note the increase of crystallinity and decrease of porosity at grain surfaces.</figcaption></figure>

from numerical simulations, labeled H₂ simul (Figure 5d). In contrast, the experimentally measured amounts of CO₂ exp and CH₄ exp significantly differ from thermodynamic calculations. CO₂ exp is fourfold lower than CO₂ simul, with 0.25 mmol measured versus 1.11 mmol modeled. Methane was not even detected at the end of the experiment, despite simulations predicting a CH₄ production of 0.13 mmol if thermodynamic equilibrium was achieved. This discrepancy is likely due to kinetic factors, an incomplete consideration of the involved species, or a combination of both, as detailed below:

1. The reaction kinetic is slow and the experiment did not reach the modeled thermodynamic equilibrium before ending. This could explain the very low level of CH₄ in the experiment, in case its generation is a slow process under strong kinetic control.

2. One (or several) chemical reaction has not been taken into account during numerical simulations, which should then involve a shift in thermodynamic equilibrium within the reacted system. For instance, dissolved organic compounds are not accounted for in the numerical modeling (since most of the related thermodynamic data at 200°C are lacking), although their presence was detected after quenching the experiment. The formation of such compounds could explain the consumption of CO₂, and limit/inhibit the presence of CH₄.

This study was neither aimed at nor designed to investigate the kinetic aspects of CO₂ and CH₄ formation and consumption. As a consequence, carbon speciation and distribution were not deeply studied in this work. It prevents us to make any mass balance calculation and to draw definitive conclusions on the lack of CO₂ and CH₄ at the end of the experiment with respect to the thermodynamic equilibrium. However important concentrations of TOC were detected in fluids, already reported in H₂- and CO₂-rich hydrothermal environments (McCollom & Seewald, 2001, 2003). It is thought that dissolved carbon compounds prevail in kinetically controlled systems, rather than other thermodynamically more stable phases like CH₄ (Andreani & Ménez, 2019).
<figure><img src="image_8.png" /><figcaption>**Figure 5.** Experimental results of the 200°C autoclave run plotted along with numerically predicted parameters from thermodynamic simulations closely mimicking the experimental conditions. Dash lines correspond to numerical modeling with siderite only in the starting mineral assemblage. Plain lines correspond to numerical modeling with siderite + magnetite in the starting mineral assemblage. (a) Final pH as a function of temperature. (b) Final Fe<sub>tot</sub> (aq) as a function of temperature. (c) Final mineral assemblage as a function of temperature. (d) Gas generation (H<sub>2</sub>, CO<sub>2</sub> and CH<sub>4</sub>) as a function of temperature. The uncertainty of experimental measurements is within the size of the symbols.</figcaption></figure>

## 4.2. H₂ Generation Mechanism

The trends of H₂ generation as a function of temperature obtained after numerical simulations performed with siderite as the only reacting solid, and with a siderite-magnetite mixture are overlapping (Figure 5d). This indicates a unique process to generate H₂ that is the alteration of siderite and associated Fe oxidation following Equation 1. During this process, oxidized Fe is expected to precipitate as magnetite exclusively within the whole range of temperature investigated (Figure 5c). Experimentally, the full conversion of siderite into magnetite according to SEM, XRD, and MS results validate the conclusion that siderite alteration is the main and probably the only reaction generating H₂. In addition, the similar trends observed for H₂ and CO₂ generations over time during the experiment (Figure 2) also support the conclusion that H₂ generation originates from the alteration of siderite. Interestingly, we cannot point any positive or negative effect of magnetite in the starting assemblage on H₂ levels, as the comparison between experimental results and thermodynamic modeling do not exhibit significant differences (Figure 5d). Such a comparison suggests that magnetite does not enhance the overall amount of H₂ produced. Kinetically, primary magnetite might still have played an indirect role, for instance by accelerating siderite dissolution or by promoting CO₂ reduction, but such combined effects cannot be unambiguously resolved from our experimental results.

## 4.3. Iron Partitioning During Siderite Alteration

Raw MS results allow to calculate the molar budget of siderite and magnetite in the starting powder: 1.88 and 3.37 mmol respectively, for a total Fe budget of 11.98 mmol (see Texts S2 and S3 in Supporting Information S1 for the detailed calculations). From this budget, 9.6% Fe was oxidized during the experiment (Fe²⁺/ΣFe<sub>bulk,200</sub>
<figure><img src="image_9.png" /><figcaption>**Figure 6.** (a) Bimodal partitioning of dissolved Fe²⁺ from siderite within two different magnetite reservoirs. The mean stoichiometry *x*<sub>mgt,expect</sub> of the bulk magnetite at the end of the reaction depends on the partitioning, and is modeled in panel (b). Plotting the experimental *x*<sub>200</sub> along the partitioning curve allows to determine the amount Φ of Fe²⁺ incorporated into new magnetite, Mgt(II), during the experiment.</figcaption></figure>

minus Fe²⁺/ΣFe<sub>bulk,start</sub>), corresponding to 1.15 mmol of Fe<sup>III</sup>. Assuming that all Fe oxidized (1.15 mmol) originated from the starting siderite and was incorporated into newly-formed magnetite Mgt(II) during H₂ generation following Equation 1 (with a stoichiometric ratio of 1 Fe<sup>II</sup> and 2 Fe<sup>III</sup>), it follows that the total amount of Fe incorporated within magnetite reached 1.73 mmol. Thus, 0.150 mmol of Fe present in the starting siderite remain missing (1.88 mmol of siderite = 1.88 mmol Fe) and cannot be accounted for by the very low amount of dissolved ΣFe present in fluid (0.87 µmol/L corresponding to 0.28 µmol in the experimental solution).

A second mechanism must thus be invoked to sequester Fe coming from the dissolution of siderite. In the literature, uptake of dissolved Fe²⁺ by non-stoichiometric magnetite has been widely studied (Gorski et al., 2012; Gorski & Scherer, 2009; Pasakarnis et al., 2013). This uptake leads to a reorganization within the mineral lattice and finally to a more stoichiometric magnetite. Our SEM observations (Figure 4d) revealed an important reactivity at primary magnetite surfaces Mgt(I), suggesting that this mechanism is at play for Fe²⁺ immobilization within the final mineral assemblage. A model was built to investigate this assumption, consisting in calculating the expected stoichiometry of bulk magnetite (*x*<sub>mgt,expect</sub>) in the final mineral assemblage as a function of the partitioning of Fe<sub>siderite</sub> in the starting assemblage. The theoretical partitioning, labeled Φ, can range from 0.00 to 1.88 mmol. This Fe<sub>siderite</sub> could be incorporated into two sinks, either at Mgt(I) surface or in newly-formed Mgt(II). Visually, the model corresponds to a curve that spans between two endmembers. If all Fe<sub>siderite</sub> is incorporated into newly-formed Mgt(II), the final stoichiometry of the bulk magnetite reaches *x*<sub>mgt,expect</sub> = 0.441 (Figure 6, Φ = 1.88). On the contrary, if all Fe from siderite is incorporated at primary Mgt(I) surfaces, the final stoichiometry of the bulk magnetite reaches *x*<sub>mgt,expect</sub> = 0.701 (Equation five and Figure 6, Φ = 0), which is thermodynamically impossible since *x*<sub>mgt</sub> cannot exceed 0.5 (Pasakarnis et al., 2013). By plotting on the curve our experimental stoichiometry *x*<sub>200</sub> = 0.463, we find a total Fe amount incorporated into newly-formed Mgt(II) of Φ = 1.72 mmol. This result falls very close to the 1.73 mmol of magnetite formed according to the percentage of Fe oxidized during the experiment. As a consequence, we propose that the slightly non-stoichiometric primary magnetite acted as a minor Fe²⁺ sink during the experiment. The sink of Fe²⁺ at Mgt(I) surface apparently did not impact significantly the generation of H₂ (Figure 5d), presumably because this effect was limited (only 8% of the Fe budget from siderite).

## 4.4. Influence of the Headspace and Water-Rock Ratio on the H₂ Generating Potential of Siderite

Our hydrothermal experiment, designed with both a headspace and a very high W-R represents an extreme scenario to investigate siderite behavior during hydrothermal alteration. To evaluate siderite reactivity at lower W-R and without headspace, typical of most semi-closed geological settings, two additional thermodynamic simulations were conducted at 200°C using almost the same starting mineral assemblage as in our experiment: 990 mg made up of 22 wt% siderite (~1.88 mmol) and 78 wt% magnetite (~3.37 mmol, stoichiometric here). The first simulation consisted in modulating only the headspace volume, from 0 to 200 mL, with a constant W-R of
<figure><img src="image_10.png" /><figcaption>**Figure 7.** Numerical simulation results illustrating (a, b) solid phase assemblage and (c, d) total H₂ and CH₄ (mol/kg<sub>siderite</sub>) formed at thermodynamic equilibrium after alteration of a siderite-magnetite assemblage at 200°C. Results are shown for (a, c) constant W-R = 1 and headspace volumes ranging from 200 to 1 mL and (b, d) various W-R ranging from 300 to 1, without headspace. Our experimental results serve as one endmember for comparison with simulations. Total H₂ and CH₄ generations include both dissolved and gaseous species in simulations with headspaces (solid lines), and only dissolved species for simulations without headspace (dashed lines).</figcaption></figure>

300 (Figures 7a and 7c). The second simulation consisted in modulating only the W-R, from 1 to 300, without headspace at the beginning of the run (Figures 7b and 7d). For this simulation, an external pressure of 100 bar, equivalent to 1 km depth of hydrostatic pressure, was imposed. A gas phase was allowed to form if the partial pressures of dissolved gases reached the bubbling point, but this condition was never met. Thus, all gaseous species remained dissolved in this second simulation. As previously discussed, the CH₄ formation predicted by these two thermodynamic models might appear somewhat unrealistic if the formation of dissolved organic compounds predominates due to kinetic constraints (Section 4.1). Nevertheless, these models provide a useful first-order estimate of the impact of the W-R and the presence of a gas phase on H₂ generation yields within the studied system.

According to our first numerical simulation, the mineralogy after alteration remains unchanged regardless of the headspace volume. All siderite is consumed, resulting in the precipitation of a constant amount (0.63 mmol) of new magnetite (Figure 7a). However, the total amount of H₂ produced (dissolved + gaseous) decreases from 318 to 47 mmol<sub>H₂</sub>/kg<sub>siderite</sub> as the headspace volume decreases from 200 to 1 mL (Figure 7c). In parallel, more CH₄ is generated through CO₂-H₂ interactions, as evidenced by the steadily increasing CH₄ generation, from 634 to 699 mmol<sub>CH₄</sub>/kg<sub>siderite</sub> (Figure 7c).

In scenarios where the absence of headspace prevents gas exsolution, complete siderite dissolution is still observed, and the same amount of magnetite (0.63 mmol) precipitates for W-R > 50 (Figure 7b). However, in the W-R range of 300 to 50, H₂ generation decreases progressively from 38.6 to 6.45 mmol<sub>H₂</sub>/kg<sub>siderite</sub>, buffered by CH₄ production (Figures 7d and 7c). For W-R between 50 and 15, the full siderite dissolution still operates, and the amount of magnetite remains unchanged (Figure 7b). However, graphite becomes stable and its amount increases as W-R diminishes. The formation of such a reduced carbonaceous compound introduces an additional mechanism for H₂ consumption. As a result, H₂ concentrations are no longer constrained by the H₂-CH₄-CO₂ equilibrium but by the H₂-CH₄-CO₂-graphite equilibrium. Consequently, both CH₄ and H₂ concentrations
<figure><img src="image_11.png" /><figcaption>**Figure 8.** Qualitative schematic synthesis of experimental and numerical results, highlighting (a) the dissolution of siderite and associated release of CO₃²⁻ and Fe²⁺ in water, (b) the oxidation of Fe²⁺ associated with the formation of secondary magnetite accompanied by (c) the generation of H₂, (d) the uptake of Fe²⁺ at primary magnetite surfaces, and (e) the formation of dissolved organic compounds, graphite or CH₄. Note that all these reactions occur concomitantly. Their yields are highly impacted by the water-to-rock ratio. Dots refer to dissolved species, while any other features refer to solid materials. Figure inspired by Gorski et al., 2012.</figcaption></figure>

decrease in parallel, as they are now controlled by graphite precipitation (Figure 7d). Finally, a significant change in the behavior of the system occurs at W-R = 15, since siderite is only partially dissolved below this threshold and the system is governed by the H₂-CH₄-CO₂-graphite-siderite equilibrium. The molar fraction of siderite remaining stable at equilibrium ranges from 0% at W-R = 15 to 90% at W-R = 1 (1.70 mmol after alteration vs. 1.88 mmol before alteration). As a consequence, a lower W-R increases siderite stability, reduces graphite and magnetite precipitation, and decreases total H₂ and CH₄ generations, yielding only 73.4 µmol·H₂/kg_siderite and 6.45 mmol·CH₄/kg_siderite, respectively, at W-R = 1.

**4.5. Implications for Natural Environments**

During this numerical and experimental study, several geochemical processes were identified (Figure 8), likely operating concomitantly. These processes can be summarized and extrapolated to natural environments to broaden the perspectives:

1. *Destabilization of siderite.* Siderite is prone to destabilize at 200°C, completely at W-R > 15 and partially below, with only 10 mol% dissolved at W-R = 1. In the context of H₂ generation, this implies that significant reactivity of siderite-dominated rocks requires higher W-R compared to olivine-dominated rocks, which dissolve completely at W-R = 1 at the same temperature (Klein et al., 2013). Reciprocally, this could explain the relatively good preservation of siderite-rich lithologies, such as Precambrian BIF, through the geological history, as their alteration necessitates higher water volumes.

2. *Formation of new magnetite.* Fe²⁺ released in solution is oxidized to Fe³⁺ and subsequently precipitates as iron oxides. The amount of Fe-oxide formed depends on the total Fe²⁺ budget available for oxidation, itself controlled by both the extent of siderite dissolution, presented in (a), and the uptake of Fe²⁺ by non-stoichiometric magnetite, presented in (d). The nature of the Fe-oxide formed depends on temperature and oxygen fugacity. A previous study conducted under uncontrolled redox conditions reported goethite (α-FeO(OH)) and hematite (α-Fe₂O₃) formation during experimental siderite alteration at 50–250°C (Maddin et al., 2024). Our experimental results show that under anoxic conditions at 200°C, magnetite formation is favored, consistent with thermodynamic simulations. In contrast, our simulation results indicate
that hematite could co-exist at the thermodynamic equilibrium with magnetite at near-ambient temperatures under the same experimental conditions (<80°C, see Figure 5c).

3. *Generation of H₂*. The oxidation of Fe²⁺ and precipitation of magnetite is accompanied by the reduction of water and the formation of H₂. This process, already well documented at various temperatures in serpentinization systems (see Barbier et al., 2020 and references therein), remained poorly studied in other contexts. Until recently, hydrothermal experiments involving siderite had only demonstrated H₂ generation at T > 300°C (Milesi et al., 2015). The present study shows that siderite is also capable of generating H₂ at lower temperatures, expanding the range of potential rocks to look for natural H₂ exploration. Our findings can be directly extrapolated to BIF, which are currently under investigation for their H₂ potential (Geymond et al., 2022; Roche et al., 2024). This lithology frequently exhibits high concentrations of Fe-carbonate and are occasionally associated with late-stage magnetite precipitates, which may reflect H₂ generation through Fe²⁺ oxidation (Rasmussen & Muhling, 2018). More generally, this study infers that all Fe-carbonate bearing lithologies could also represent H₂ generating rocks. Notably, most of the significant siderite-ankerite occurrences apart from BIF are reported in carbonatites (Buckley & Woolley, 1990; Yaxley et al., 2022), recently proposed as possible H₂-generating rocks (Malvoisin & Brunet, 2023). To a lesser extent, skarns can also present important amounts of Fe-carbonates (Castorina & Masi, 2008; Choi et al., 2003).

4. *Uptake of Fe²⁺ by non-stoichiometric magnetite*. A previous work proposed that the anoxic alteration of stoichiometric magnetite could lead to non-stoichiometric magnetite formation and H₂ generation (Geymond et al., 2023). Our experimental results indicate that non-stoichiometric magnetite can capture Fe²⁺ from the fluid to recover its stoichiometry. These two findings align with earlier studies suggesting that magnetite-to-hematite (or maghemite, γ-Fe₂O₃) transformation is reversible and may proceed via redox (O₂/H₂ exchange) or non-redox (Fe²⁺ mobilization/immobilization) pathways (Mücke & Cabral, 2005; Ohmoto, 2003; Otake et al., 2007). Despite this Fe²⁺ uptake observed during the experiment, the amount of H₂ generated closely matches the H₂ generation predicted by a simplified model assuming stoichiometric magnetite (see Section 4.1). However, it is important to note that both experimental and modeled H₂ production values carry inherent uncertainties (e.g., related to gas chromatography quantification or to thermodynamic data accuracy) which could explain the apparent agreement between experimental and computed results. Most importantly, Mössbauer data indicate that the Fe²⁺ uptake was very limited, as the starting magnetite was only slightly non-stoichiometric. On the contrary, the uptake of Fe²⁺ by highly-oxidized magnetite could substantially affect the Fe²⁺ budget available for oxidation, thus representing a key control on H₂ generation yields in natural systems. Noteworthy, maghemite has been documented (to various extent) in all the siderite-bearing lithologies aforementioned, including BIF (Morris & Kneeshaw, 2011; Morris & Ramanaidou, 2007), carbonatites (Giovannini et al., 2021) and Fe-skarns (Ünlü et al., 2019). Most interestingly, maghemite and titanomaghemite (i.e., vacancy-bearing Ti-magnetite) have been reported in ultramafic rocks such as peridotites (Nazarova, 1994; Nazarova et al., 2000), gabbros (Jesus et al., 2003), and basalts (Oliva-Urcia et al., 2011), inferring their potential influence on H₂ generation yields during serpentinization.

5. *Consumption of H₂*. The amount of H₂ in the system is balanced by several consumption mechanisms. At W-R > 50, the formation of CH₄ only is expected to control H₂ concentration. During the experiment, the formation of dissolved organic carbon compounds was observed rather than CH₄, which could be due to kinetic effects. This consumption implied a lack of about 83% of H₂ within the experimental system (0.1 mmol measured vs. 0.58 mmol generated theoretically, considering 1.73 mmol FeII from siderite was incorporated into new magnetite, see Section 4.3). The presence of graphite (or carbonaceous-type precipitates) at the end of the experiment was not investigated, but is not expected at W-R = 300. In fact, numerical simulations indicate that graphite becomes stable at W-R < 50. It constitutes an additional process for H₂ consumption during hydrothermal alteration at lower W-R, as proposed by a previous experimental study performed on siderite at W-R = 2 (Milesi et al., 2015). Initially thought to constitute a marker of primitive life (Mojzsis et al., 1996), origin of graphite in BIF is now controversial as it might have formed abiotically during siderite thermal decomposition at 500-600°C according to isotopic data (Van Zuilen et al., 2002, 2003). Interestingly, our study suggests that hydrothermal alteration of siderite at lower temperature (200°C) can also form graphite, triggered by an elevated H₂ activity in the system originating from Fe²⁺ oxidation and magnetite precipitation. This finding, akin to (ultra)mafic-hosted hydrothermal systems were graphite is commonly associated with H₂-rich fluid (Andreani & Ménez, 2019; Heijlen et al., 2006), raises the question of using graphite as a possible marker to track pristine H₂ generation in siderite-rich lithologies.
## 5. Conclusion

This experimental study aimed to evaluate the mineralogical changes occurring during the hydrothermal alteration of a magnetite-siderite assemblage at 200°C. These two minerals are known to generate H₂ during water-rock interactions, either as source mineral or catalyst, but have not been tested in combination until now. The methodology involved investigating Fe speciation and partitioning using ⁵⁷Fe Mössbauer spectroscopy, a crucial method enabling the detection of slight changes in Fe speciation within the altered solids and providing precise constraints to refine H₂ generation pathway and mass balance calculations.

During the experiment, H₂ generation was observed to originate from the alteration of siderite, which was replaced in situ by magnetite incorporating the oxidized iron. This reaction, previously identified at higher temperatures, is now confirmed at 200°C. The presence of magnetite in the starting assemblage affected negatively the H₂ production by incorporating in its structure about 8% of FeII from siderite. In fact, a key observation of this study is that the slightly non-stoichiometric initial magnetite acts as a sink for dissolved Fe²⁺ released by siderite dissolution, allowing the reacted magnetite to recover a more ideal stoichiometry. In previous studies, magnetite was found to generate H₂ when reacting with water alone, but its role may differ in the presence of an external source of both Fe²⁺ and H₂. Magnetite may also have catalyzed CO₂ reduction in the presence of H₂, by kinetically promoting the formation of reduced dissolved carbon species rather than CH₄, as revealed by our measurements, which differ from the expected thermodynamic equilibrium. This process further contributed to decreasing the overall potential of H₂ generation from siderite alteration, but it remains to be better documented.

Finally, thermodynamic simulations provide interesting insights into the influence of the water-to-rock ratio (W-R), tested without headspace from 300 to 1, on H₂ generation yields, which range accordingly from 38.6 mmol/kg<sub>siderite</sub> to 73.4 µmol<sub>H₂</sub>/kg<sub>siderite</sub>. Additionally, solid-phase equilibria are also strongly affected, with siderite remaining stable for W-R < 15 and graphite precipitating for W-R < 50.

The presence of graphite in siderite-rich Banded Iron Formations (BIF) is common, and some authors have proposed siderite thermal decomposition during peak metamorphism (500–600°C) to explain the presence of graphite (Van Zuilen et al., 2002, 2003). Altogether, our results provide an additional explanation for graphite presence in BIF, that is, low temperature (200°C) hydrothermal alteration, and demonstrate that such lithology can generate H₂.

To summarize, the hydrothermal alteration at 200°C of a siderite-magnetite assemblage, and by extension siderite/magnetite-bearing lithologies such as BIF, occurs through the following steps: (a) The siderite dissolution releasing Fe²⁺ in solution, (b) the oxidation of Fe²⁺ leading to the formation of new magnetite and H₂, (c) the sequestration of part of the Fe²⁺ by the initial non-stoichiometric magnetite, and (d) the consumption of H₂ associated with the reduction of CO₂ to form reduced carbon phases.

The conclusions of this study do not preclude the possibility of higher H₂ generation yields at lower temperatures, especially if H₂ consumption processes are restricted. This possibility will need to be further investigated in future studies. This study has significant implications for our understanding of Earth's geochemical cycles and the future of natural H₂ exploration as an energy resource.

## Data Availability Statement

Details on the gas sampling methodology and gas chromatography (GC) analyses are provided in Text S1 and Figure S1 of Supporting Information S1. Calculations for mineral formula units are described in Texts S2 and S3 of Supporting Information S1. Additional information on the characterization of the starting material using Mössbauer Spectroscopy (MS) can be found in Text S4, Table S1, and Figure S2 of Supporting Information S1. The hyperfine parameters used to fit the MS data at 295 K are compiled in Table S2 of Supporting Information S1. All raw data supporting this study, including MS and GC datasets as well as input and output files from numerical simulations, are freely and permanently available in a public repository at (Geymond, 2025).

## References

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* Barbier, S., Huang, F., Andreani, M., Tao, R., Hao, J., Eleish, A., et al. (2020). A review of H₂, CH₄, and hydrocarbon formation in experimental serpentinization using network analysis. *Frontiers in Earth Science*, 8, 209. https://doi.org/10.3389/feart.2020.00209

## Acknowledgments

All authors want to express their gratitude to Sophie Nowak (UPC) for performing XRD analyses, Stephan Borenstjan (IPGP) for sharing his knowledge on SEM, Dan Lévy (IPGP) and Isabelle Moretti (UPPA-LFCR) for their help on final interpretations and corrections on the manuscript. This work is part of the PhD project of UG, funded by MERSI and INSU-CNRS (Tellus). Mössbauer analyses realized in this work were supported by the Charles University, project GA UK no. 410622. Laurent Truche acknowledges support from the Institut Universitaire de France-CNRS. We also sincerely thank Stéphanie Vialle, the second anonymous reviewer, and the associate editor for their valuable comments, which helped improve the quality of the manuscript.
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