Assessing the Effect of Mineralogy and Reaction Pathways on Geological Hydrogen (H₂) Generation in Ultramafic and Mafic (Basaltic) Rocks

Abstract

This study evaluates the impact of mineralogy, elemental composition, and reaction pathways on hydrogen (H₂) generation in seven ultramafic and mafic (basaltic) rocks. Experiments were conducted under typical low-temperature hydrothermal conditions (150 °C) and captured early and evolving stages of fluid–rock interaction. Pre- and post-interactions, the solid phase was analyzed using X-ray Diffraction (XRD) and X-ray Photoelectron Spectroscopy (XPS), while Inductively Coupled Plasma Mass Spectrometry (ICP-MS) was used to determine the composition of the aqueous fluids. Results show that not all geologic H₂-generating reactions involving ultramafic and mafic rocks result in the formation of serpentine, brucite, or magnetite. Our observations suggest that while mineral transformation is significant and may be the predominant mechanism, there is also the contribution of surface-mediated electron transfer and redox cycling processes. The outcome suggests continuous H₂ production beyond mineral phase changes, indicating active reaction pathways. Particularly, in addition to transition metal sites, some ultramafic rock minerals may promote redox reactions, thereby facilitating ongoing H₂ production beyond their direct hydration. Fluid–rock interactions also regenerate reactive surfaces, such as clinochlore, zeolite, and augite, enabling sustained H₂ production, even without serpentine formation. Variation in reaction rates depends on mineralogy and reaction kinetics rather than being solely controlled by Fe oxidation states. These findings suggest that ultramafic and mafic rocks may serve as dynamic, self-sustaining systems for generating H₂. The potential involvement of transition metal sites (e.g., Ni, Mo, Mn, Cr, Cu) within the rock matrix may accelerate H₂ production, requiring further investigation. This perspective shifts the focus from serpentine formation as the primary driver of H₂ production to a more complex mechanism where mineral surfaces play a significant role. Understanding these processes will be valuable for refining experimental approaches, improving kinetic models of H₂ generation, and informing the site selection and design of engineered H₂ generation systems in ultramafic and mafic formations.

1. Introduction

The geological process known as serpentinization has garnered significant attention in recent years due to its potential as a method for generating low-carbon hydrogen (H₂). This process involves the reaction of water with ultramafic rocks, primarily composed of olivine and pyroxene, leading to the oxidation of ferrous iron (Fe²⁺) and the production of hydrogen gas as a byproduct. Hydrogen generation through fluid–rock interactions has been extensively studied, with a focus on ultramafic and mafic rocks due to their rich Fe-bearing mineral compositions. An extreme case is the hydrous alteration of ultramafic rocks, known as serpentinization, where most of the original rock minerals are replaced by serpentine. The alteration of mafic rocks, such as basalts, also produces H₂, but in smaller amounts compared to those generated in ultramafic environments [1,2,3,4]. Experimental studies suggest that the alteration of felsic rocks under hydrothermal conditions can also lead to significant H₂ production [5].

Early studies, such as those by [6], primarily investigated serpentinization kinetics, focusing on the hydration rates of forsterite and enstatite, which led to the formation of serpentine, brucite, and talc. However, these studies did not assess H₂ production, leaving a gap in understanding the direct role of mineral transformation in hydrogen generation. Kita et al. [7] explored H₂ generation in granite and quartz systems at temperatures ranging from 25 to 270 °C, attributing H₂ production to reactions involving Si- and Si-O-based transformations. These studies laid the foundation for understanding geochemical transformations, but did not fully address the iron (Fe) oxidation pathways as a major driver of H₂ evolution.

Recent discoveries underscore the growing importance of natural hydrogen systems in diverse geological contexts. For example, studies in the Songliao Basin [8] and Bohai Bay Basin [9] in China demonstrate that large-scale H₂ generation is actively occurring in mafic and ultramafic lithologies, highlighting their potential as sustainable subsurface energy resources. Similarly, well-characterized sites in Oman, the Philippines, and Turkey [10] confirm that serpentinization-driven H₂ production is a global phenomenon with significant implications for future energy portfolios. Despite these advances, fundamental mechanisms governing H₂ fluxes remain poorly constrained, particularly the interplay of mineralogy, geochemical conditions, and fluid–rock interactions. Such knowledge is critical for assessing the viability of natural hydrogen as a sustainable energy resource and for developing strategies to optimize its production, utilization, and storage in subsurface systems.

More recent research has demonstrated that H₂ generation is strongly linked to the oxidation of Fe²⁺ to Fe³⁺ in ultramafic systems. Marcaillou et al. [11] investigated peridotite at 300 °C and 300 bar, demonstrating that Fe²⁺ oxidation in olivine and pyroxene, resulting in Fe³⁺ in magnetite and serpentine, contributed to H₂ production. Mayhew et al. [12] expanded upon these findings by examining a broader range of ultramafic rocks and minerals, including fayalite, pyroxene, olivine, and magnetite, under ambient pressure and at temperatures of 55 °C and 100 °C over a period of 100 days. Their results indicated a direct correlation between H₂ production and metal oxide minerals with the general formula [M²⁺M₂³⁺]O₄, though their experiments did not reach equilibrium conditions. Similarly, [13] conducted computational simulations on various compositions of forsterite-fayalite, establishing a direct correlation between the Fe content of olivine and the extent of H₂ generation across a temperature range of 25 °C to 400 °C at 50 MPa.

Experimental studies by [14,15] have further refined our understanding of H₂ production mechanisms in rocks. Using synthetic samples at 230 °C and 35 MPa over 146 days, the authors of [14] found that H₂ generation increased significantly with pH and that the reaction of olivine and orthopyroxene proceeded at a slower rate initially. Their later study [15] focused on high-Mg, low-Fe minerals, demonstrating that Fe-rich olivine produced more H₂ per mole than Fe-poor olivine. These results reinforce the importance of Fe oxidation states in controlling H₂ generation and suggest that mineral composition plays an important role in determining reaction kinetics and efficiency. However, these studies examined fewer rock types, which limits their broader applicability.

Some previous studies used only major elements data, including MgO, SiO₂, FeO, CaO, Na₂O, and K₂O, to estimate H₂ generation potential. Although these estimates of past ultramafic rock distribution rely on assumptions based on a single bulk rock composition [16,17], natural ultramafic rocks exhibit compositional variability, leading to differences in H₂ production potential. Leong et al. [18] addressed this variability by incorporating rocks with a range of compositions, from Mg-rich to Mg-depleted. However, variations in Fe and other trace minerals were not considered, and the simulations were conducted only under ambient temperature conditions.

Leong et al. [18] utilized thermodynamic simulations of water–rock interactions combined with mass-transport calculations to estimate H₂ production from the alteration of Fe-bearing igneous rocks under low-temperature conditions (25 °C). Their simulations predicted H₂-generation potential for the hydrous alteration of Fe-rich igneous rocks, ranging from ultramafic rocks with high MgO content, such as peridotites, to those with lower MgO content, like basalts. The authors observed a decrease in H₂-generation potential between rocks with MgO content above and below 35 wt.%, even when FeO concentrations were similar. Additionally, their model results indicated a gradual reduction in H₂ generation as rock compositions became less Mg-rich, approaching 20 wt.% MgO. They do not promote serpentine formation during rock alteration. Instead, the alteration of these Mg-deficient but Si- and Al-rich rocks favors the stabilization of minerals like chlorite, talc, and clay minerals.

Recent research has also emphasized the impact of environmental conditions on serpentinization and H₂ evolution. Leong et al. [19] investigated H₂ outgassing from the Samail ophiolite in Oman under ambient temperature and pressure conditions (25 °C, 1 bar). Their study reported serpentinization rates and H₂ yields of approximately 8 × 10⁻¹⁴ s⁻¹ and 0.3 mol H₂ kg⁻¹ of ultramafic rock, respectively, emphasizing the potential for natural H₂ reservoirs in ultramafic formations. The most prominent reactions being reported in the literature involve the reaction of olivine and water to form serpentine [14,15] or serpentine, magnetite, brucite, and hydrogen [11,14,20,21,22], while others reported the formation of magnetite, silica, and hydrogen generation, as shown in reactions (1) through (4).

Olivine (Fosterite) + Water → Serpentine

3Mg₂SiO₄ + SiO₂ + 4H₂O → 2Mg₃Si₂O₅(OH)₄

(1)

Olivine (Fosterite) + Water → Serpentine + Brucite + Hydrogen

2Mg₂SiO₄ + 3H₂O → Mg₃Si₂O₅(OH)₄ + Mg(OH)₂

(2)

2Mg₂SiO₄ + 4H₂O + H⁺ → Mg₃Si₂O₅(OH)₄ + Mg(OH)₂ + H₂ + OH⁻

(3)

Olivine (Fayalite) + Water → Magnetite + Silica + Hydrogen

3Fe₂SiO₄ + 2H₂O → 2Fe₃O₄ + 3SiO₂ + 2H₂

(4)

The amount of H₂ generated through Fe oxidation, as seen in reaction (4), is influenced far more by the bulk composition of the rock [18,23]. Furthermore, the thermodynamic modeling by [23,24] suggests that other reactions may be contributing to H₂ production. These authors indicate that iron partitioning and oxidation state are highly dependent on reaction conditions, while the availability of an external silica source influences magnetite formation.

The body of literature demonstrates that Fe oxidation and partitioning, rock mineral composition, and the impact of pH and temperature are important factors in determining H₂ yields from ultramafic and mafic rocks. Rock mineralogy is a fundamental factor influencing H₂ generation during serpentinization and hydro-geochemical reactions in mafic and ultramafic rocks [18,23,25]. Accordingly, rock minerals may not only participate as reactants but could also act as goecatalysts, influencing reaction rates and pathways without necessarily being consumed or transformed [26,27,28,29,30,31,32]. The potential for catalysis in geochemical reactions due to the rock minerals is likely to have significant implications for their progression.

The hypothesis of this study is that H₂ generation from ultramafic and mafic rocks is not solely controlled by mineral hydration or serpentinization, but is also influenced by Fe oxidation–reduction cycles, and the presence of trace reactive minerals in ultramafic and mafic rocks. The objective of this study is to test this hypothesis by investigating what mineralogy, elemental composition, and reaction mechanisms influence H₂ generation in ultramafic and basaltic mafic rocks. It aims to identify key mineral phases, and geochemical interactions that control H₂ production during water–rock interactions in these natural rock systems.

The current literature primarily focuses on either a limited selection of rock samples or relies heavily on simulations to study hydrogen generation processes. The novelty of this work lies in its comprehensive experimental investigation of a diverse range of rock types, up to seven, including ultramafic and mafic rocks sourced from different geological regions. This broader scope allows for a more in-depth understanding of H₂ generation beyond the widely studied serpentinization process, revealing additional reaction pathways and complexities that may not have been extensively explored. Employing X-ray Photoelectron Spectroscopy (XPS) for surface elemental analysis provides detailed information on the elemental composition and oxidation states of these rocks. This involves assessing natural ultramafic and mafic (basaltic) rocks obtained from various locations, comparing H₂ generation potential, mineral phase modifications, reaction rates, and changes in the oxidation states of some major elements via spectroscopic methods. Understanding the role of minerals in accelerating reactions or acting as buffers could lead to new insights into natural hydrogen generation, mineral transformations, and broader geochemical processes to improve hydrogen generation yield from mafic and ultramafic rocks.

2. Methodology

2.1. Rock Samples

This study utilized a total of seven rock samples, consisting of five ultramafic and two mafic (basaltic) rock types (see Figure 1), to investigate their potential for H₂ generation under controlled experimental conditions. The ultramafic rocks consisted of two samples sourced from the Trinity Ophiolite, at Eunice Bluff, California, designated as Samples A and E. Additionally, the ultramafic category included three peridotite samples from different sources: Peridotite Sample B, obtained via Ward’s Science from Balsam, North Carolina; Olivine Sample C obtained via Northern Geological Supplies Ltd., Bolton, UK, from the Gusdal Olivine Pit in Åheim, Norway; and Sample D, acquired from the Twin Sisters Peridotite in Washington. The mafic (basaltic) group included two samples: Sample F obtained via Northern Geological Supplies Ltd., Bolton, UK, from Finland, and Sample G, sourced through a supplier, EISCO, on Amazon.com, processed in Victor, New York, whose exact source location could not be determined. These samples represent a diverse selection of specimens collected from both within and outside the United States.

The rock samples were initially crushed using a porcelain mortar and pestle to avoid any potential metal contamination that could influence subsequent geochemical reactions. The crushed material was then dry-sieved to obtain a uniform particle size distribution. To remove adhering fines and dust that might artificially enhance reaction kinetics, the sieved fractions were carefully rinsed with de-ionized (DI) water and subjected to ultrasonic agitation (sonication). This cleaning step ensured that only well-defined grains were retained for experimentation. A target particle size range of 150–300 µm was selected, as this size is commonly used in laboratory-scale fluid–rock interaction studies to balance sufficient surface reactivity with representative bulk rock properties. The prepared samples were characterized to confirm a consistent specific surface area of approximately 12 m²/kg of rock, which was maintained across all experiments to ensure comparability of results.

2.2. Experimental Description

The experiments were conducted in a 150 mL Hastelloy mini-autoclave using a water–rock ratio of approximately 2. The water–rock mixture was placed in the Hastelloy autoclave reactor. The system was pressurized to 320 psi using high-purity nitrogen (99.99%). Before pressurization, nitrogen gas was used to purge the free space, ensuring the removal of air from the system. The autoclave was then heated in an oven set to 150 °C at a heating rate of 10 °C per minute, with a precision of ±1 °C. As the temperature reached 150 °C, the reaction reactor pressure increased to 500 psi. The rock samples were subjected to temperature and pressure hydrothermal conditions reported in some serpentinized systems [12,22,33], simulating natural subsurface environments where natural hydrogen is produced through rock–water interactions. The hydrothermal rock–fluid interactions were carried out for 3 days in the first batch and 7 days in the second batch of the experiment. The experiments and measurements were conducted at 3 and 7 days because the samples were crushed to provide high-surface-area powders, which accelerate serpentinization. Under these conditions, diagnostic changes such as reactive minerals dissolution and H₂ release can be detected, effectively capturing the early stages of reaction rates. Extended experiments were also performed for Sample A, with reaction durations of 14 and 28 days. Low-temperature serpentinization, mineral alterations, and H₂ production over these time durations have been reported elsewhere [14,34,35,36].

2.2.1. Fluids Analysis

At the end of the experiment, the system was allowed to cool at an ambient temperature for one hour, and the H₂ generated was measured. Hydrogen gas detection was performed using an electrochemical gas analyzer with a measurement accuracy of ±2 ppm. The H₂ gas was carefully collected in a transparent, sealed plastic bag to ensure reliable and contamination-free measurements. Inside the sealed plastic bag, the H₂ detector was placed to measure the gas concentration in a controlled environment. A known volume of gas was introduced, ensuring that the concentration remained within the detection range of the analyzer. Once the gas was stabilized, the H₂ concentration was recorded.

The supernatant aqueous fluid was collected in separate vials and analyzed using Inductively Coupled Plasma Mass Spectrometry (ICP-MS). This technique analyzes the chemical composition by measuring the levels of various elements. This was used to detect elements in the liquids, including those present in trace amounts. ICP-MS to assess the elemental compositions necessary for understanding the precipitation/dissolution processes during the rock–fluid interactions. In addition, the total dissolved solids (TDS), conductivity, and pH of the reaction fluids were measured using a pH/TDS/conductivity meter to monitor changes in the geochemical properties of the system over time. These parameters serve as indicators of fluid chemistry evolution and potential mineral dissolution or precipitation processes. Measuring TDS provides the solution’s overall concentration of dissolved ions, while conductivity reflects the ionic strength and mobility of charged species. Meanwhile, pH variations can indicate acid–base reactions and shifts in equilibrium conditions. These measurements were conducted to evaluate the precipitation and dissolution kinetics of rock–fluid interactions during hydrogen production.

2.2.2. Mineral Characterization Using X-Ray Diffraction (XRD)

X-ray Diffraction (XRD) analysis was performed on the unreacted rock samples to establish a baseline mineralogical composition, which was then compared to the mineralogical changes observed in the hydrothermally reacted samples. This enables the identification of any phase transformations, mineral dissolution, or secondary mineral formations that result from hydrothermal reactions. At the end of the experiment, the solids were retrieved and dehydrated in an oven at 100 °C for 2 h. Alteration of mineral phases was evaluated through XRD analysis using a Malvern Empyrean PANalytical diffractometer equipped with Ni-filtered Cu Kα radiation. The dried bulk solids were homogenized into fine powder using a porcelain mortar and pestle and analyzed under operating conditions of 45 kV and 40 mA, over a 2θ range of 4.5–70°, with a scan time of 8.7 s. The resulting data were analyzed using the HighScore Plus (v. 4.9) search-match module, integrated with the ICDD PDF-4-2025 mineralogical library.

2.2.3. X-Ray Photoelectron Spectroscopy (XPS)

The X-ray Photoelectron Spectroscopy (XPS) spectra of the Fe 2p, Si, Mg, Al 2p, and O 1s photoelectron lines were recorded, and data were acquired using an Omicron ESCA+ system equipped with a magnesium (Mg) X-ray source. The emission current was set to 20 mA, while the applied voltage was maintained at 15 kV to ensure optimal excitation of core electrons. The pass energy was set at 20 eV, providing high-resolution spectra with enhanced peak definition and minimal background noise. The system was calibrated before analysis, and data acquisition was conducted under ultra-high-vacuum (UHV) conditions to minimize contamination and ensure reliable results. The collected spectra were used to determine the elemental composition, oxidation states, and chemical bonding environments of the rock samples before and after the reaction. Deconvolution and interpretation of the components’ peaks were performed using CasaXPS software, version 2.3.26PR1.0. The determination of the oxidic Al 2p core level binding energies and the measured spectra were resolved into oxidic and metallic components in their respective binding energy regions. This procedure is discussed in detail in the works by [37,38,39].

3. Results

3.1. Hydrogen Generation

Hydro-geochemical reactions, often broadly referred to as serpentinization, are commonly associated with H₂ generation across various rock types. In the ultramafic and mafic rocks, both high and low levels of H₂ production have been observed, as shown in Figure 2. Among ultramafic rocks, Sample C generated as little as 1.2 mmol/kg of rock, while Sample B produced up to 9.6 mmol/kg of rock. This variation shows that H₂ generation can vary significantly within the same rock class. Despite both being ultramafic, there is an apparent disparity in their H₂ output.

Similarly, for mafic rocks, Sample F produced up to 10.8 mmol/kg of rock, whereas Sample G yielded only 0.76 mmol/kg of rock. These results indicate that H₂ generation is not uniform across ultramafic and mafic rocks, suggesting the influence of mineralogical and geochemical differences on H₂ production. The mineralogical composition of rocks plays a significant role in governing H₂ generation during serpentinization and hydro-geochemical reactions in mafic and ultramafic formations. The primary difference between mafic and ultramafic rocks lies in their silica and magnesium content [18]; however, this is likely not the only factor influencing H₂ production in these rocks. Replicate experiments, particularly the 3-day batch tests, were carried out on selected rock types under identical conditions. The results showed strong consistency in hydrogen production patterns and associated mineralogical changes. The subsequent sections examine how the mineralogy, elemental composition, and reaction pathways influence H₂ production in these rock types. The key mineral phases and geochemical processes that regulate H₂ release during water–rock interactions in ultramafic and mafic environments are further analyzed.

3.2. Mineralogical Characterization

3.2.1. Elemental Composition

The elemental composition of the rock samples using X-Ray Fluorescence (XRF) analysis is presented in Figure 3. Although both ultramafic and mafic rocks exhibit comparable concentrations of silicon (Si), iron (Fe), manganese (Mn), and zinc (Zn), they display notable differences in the composition of other major cations. Specifically, the mafic rocks (Sample F and G) contain substantially higher levels of aluminum (Al), calcium (Ca), sodium (Na), and potassium (K), while exhibiting significantly lower concentrations of magnesium (Mg) in comparison to the ultramafic rocks (Sample A to E). These variations in elemental composition likely reflect differences in the mineralogical makeup and geochemical evolution of the rocks, which could, in turn, result in varying hydrogen H₂ generation potentials across the different rock types.

3.2.2. Ultramafic Rocks

There are significant variations in the mineral composition of the assessed rocks, with each rock exhibiting unique mineralogical alterations over different treatment durations. For ultramafic rocks, notable mineralogical changes occur during the generation of hydrogen. For example, significant mineral alteration was observed in Sample A (Figure 4). The olivine content, initially 59.5 wt.%, decreased to 57.3 wt.% after 3 days and decreased to 55.3 wt.% after 7 days of interaction, and no brucite formation was detected in this sample over this reaction period. In Sample B (Figure 5), a substantial increase in the lizardite phase was observed after 3 days (64.3 wt.%) and 7 days (68.5 wt.%) of hydrothermal exposure, compared to its initial composition (38.4 wt.%). This aligns with the commonly reported serpentinization reaction responsible for H₂ production in hydrothermal systems. Additionally, Fe-rich minerals such as olivine, hematite, and clinopyroxene showed a decrease, indicating their transformation into lizardite, which correlates with H₂ generation. Furthermore, birnessite, a manganese hydroxide (Mn(OH)₄), was also observed to substantially decrease after 7 days of exposure, suggesting additional geochemical interactions during the reaction process.

Similar mineral alterations were observed for Sample C (Figure 6), with a decrease in olivine content from the initial 82.4 wt.% to 77.6 wt.% and 74.6 wt.% after 3 and 7 days of reaction, respectively. Correspondingly, lizardite increased from 4.8 wt.% to 7.3 wt.% and 8.2 wt.%. Additionally, this sample showed an increase in brucite content from 1.8 wt.% to 2.6 wt.% after 7 days. Contrary to Sample A, which showed a decreasing trend in hematite, Sample C exhibited an increase in hematite content from 1.6 wt.% to 3.9 wt.% and 5.7 wt.% after 3 and 7 days of hydrothermal interaction, respectively. No magnetite was observed in this sample. XRD analysis revealed changes in the composition of iron-rich minerals such as olivine, hematite, and magnetite, indicating their role as major reactants in the transformation process.

For Sample D (Figure 7) and Sample E (Figure 8), although a decent amount of H₂ was generated by these samples (3.2 and 5.4 mmol/kg of rock, respectively), there was no notable magnetite presence as shown in the XRD. Considering Sample D (Figure 4), olivine content decreased from 70.7 wt.% to 66.8 wt.% and 62.8 wt.% after 3 and 7 days, respectively. Meanwhile, lizardite increased from 15.2 wt.% in the untreated rock to 23.9 wt.% after 7 days of interaction, accompanied by the formation of brucite, which reached 1.4 wt.% after 7 days of exposure. This is the well-established pathway for H₂ generation during serpentinization. It involves serpentine (lizardite) formation through the hydration of olivine, along with brucite precipitation, where Mg-rich forsterite reacts with water to produce Mg(OH)₂. Additionally, the oxidation of Fe-rich wuestite (FeO) facilitates H₂ release as Fe²⁺ undergoes oxidation.

For Sample E (Figure 8), where larger H₂ was produced (compared to Sample D), there was considerable mineral phase transformation. Specifically, the lizardite content increased from 19.6 wt.% in the untreated rock to 36.7 wt.% after 7 days of hydrothermal exposure. No brucite or magnetite was formed. This suggests that not all geological H₂-generating reactions in ultramafic rocks result in the formation of brucite or magnetite.

3.2.3. Mafic (Basaltic) Rocks

The mafic rocks investigated primarily comprise plagioclase minerals, with calcic labradorite and anorthite being the dominant phases. In addition to these, olivine and trace amounts of silicate minerals, including biotite, orthopyroxene (enstatite), clinopyroxene (diopside), and zeolite, were also identified. Despite the general similarity in mineralogical composition between the two mafic rock samples examined, a significant difference in H₂ generation was observed. Sample F showed a remarkably high H₂ production, reaching up to 10.8 mmol/kg of rock, whereas Sample G yielded only 0.76 mmol/kg of rock. This stark contrast suggests that factors beyond mere bulk mineral composition influence H₂ production, with the reaction pathway and mineral transformation mechanisms likely playing an important role.

Before hydrothermal treatment, both rock types contained similar primary mineral phases, including labradorite, andesine, biotite, olivine, and enstatite. However, after 7 days of reaction, Sample F exhibited a transformation of plagioclase labradorite into andesine, suggesting a gain of Na or a loss of Ca from the solid matrix (Figure 9). In contrast, sample G transformed into anorthite, indicating Na depletion or Ca enrichment within the crystalline phase (Figure 10). These variations in element redistribution suggest distinct geochemical processes that may have influenced the extent of H₂ production in each rock type.

Interestingly, Fe-rich mineral phases, including clinopyroxene (enstatite) and maghemite, were observed in Sample F, while olivine decreased after 7 days of reaction. This transformation was also accompanied by the consumption of cristobalite (SiO₂), a silica polymorph that may indicate enhanced reaction kinetics. The presence of these Fe-rich phases is significant, as iron oxidation and redox cycling are known to play a critical role in H₂ generation [11,14,15,40,41]. This suggests that the higher H₂ production observed in Sample F could be attributed to active Fe-driven redox reactions and the formation of reactive mineral surfaces, including maghemite and biotite. This could be the primary reason for the high H₂ generation observed in this mafic rock type, surpassing even the H₂ produced by all the ultramafic rocks studied. In contrast, Sample G exhibited the lowest H₂ production among all the rocks investigated, which could be linked to its lower total content of Fe-rich minerals in the unreacted state compared to the other analyzed rocks, as shown in Figure 11.

3.3. Elemental Analysis and Iron (Fe) Oxidation State

3.3.1. XPS Elemental Analysis

X-ray Photoelectron Spectroscopy (XPS) was performed to determine the elemental composition and chemical state of the rock samples before and after reaction by measuring the binding energies. The XPS atomic ratios of the major elements in the ultramafic and mafic rocks are analyzed. The mafic rock samples, Sample F and Sample G, exhibited a higher silicon (Si) content compared to the ultramafic rocks. At the same time, their magnesium (Mg) composition is relatively lower in the bulk surface. Basalts originate from smaller extents of mantle melting, which usually results in rocks with lower Mg and higher Si, Fe, and Al contents [18,42]. This compositional variation is significant because Mg-rich minerals, such as olivine and pyroxenes, are commonly associated with serpentinization reactions in ultramafic rocks. However, the higher Fe and Al content in these mafic rocks, compared to their ultramafic counterparts, suggests that factors beyond just Mg content play an important role in determining the reactivity of these rocks in H₂ generation.

The increasing Fe content with treatment duration observed in the bulk surface of these mafic rocks is believed to be a key factor influencing their H₂ generation potential, specifically in Sample F. Iron is a key component in redox reactions, particularly in converting Fe²⁺ to Fe³⁺, a process closely linked to H₂ production. The availability of Fe-bearing minerals, such as clinopyroxene, maghemite, and biotite, may facilitate surface-mediated electron transfer, potentially enhancing H₂ evolution in this basaltic rock. Additionally, aluminum (Al) in the crystalline structure can influence mineral stability and reaction pathways by affecting cation exchange and fluid–mineral interactions. The compositional differences between mafic and ultramafic rocks suggest that while Mg content is traditionally associated with H₂ production through serpentinization, the presence of Fe-rich phases in mafic rocks could provide an alternative mechanism for sustained hydrogen generation.

3.3.2. Iron (Fe) Oxidation States and H₂ Generation

Iron (Fe) can exist in multiple oxidation states, with Fe²⁺ and Fe³⁺ being the most common in geochemical systems. Different iron oxides, such as magnetite (Fe₃O₄), hematite (Fe₂O₃), and wuestite (FeO), exhibit distinct binding energy (B.E) peaks in X-ray Photoelectron Spectroscopy (XPS). However, due to the overlap of peak positions among these oxides, accurately determining the Fe oxidation state requires careful analysis. In this study, the Fe oxidation state was identified using XPS by referencing literature-established peak, ensuring precise differentiation between Fe²⁺ and Fe³⁺ species. The input parameters for the deconvolution of the XPS spectra [37,38,39] are detailed in

Table 1. Peak model input parameters for deconvolution of XPS spectra using CasaXPS software [37,38,39].

S/#	Name	Pos. Constraint (eV)	FWHM Constraint (eV)	Area Constraint	Position (eV)	FWHM (eV)
A	Fe	729.108, 704.398	0.9, 0.95	0.0, 10,000,000.0	706.70	0.95
B	Fe2+ 1	708.75, 708.25	1.35, 1.45	0.0, 10,000,000.0	708.48	1.45
C	Fe2+ 2	710.05, 708.258	1.55, 1.75	0.0, 12,400,000.0	710.05	1.75
D	Fe2+ 3	711.25, 708.25	1.55, 1.75	0.0, 5,990,000.0	708.25	1.75
E	Fe2+ 4	712.45, 708.25	2.85, 3.05	0.0, 10,580,000.0	709.42	3.05
F	Fe2+ 5	715.75, 708.25	2.45, 2.55	0.0, 2,310,000.0	714.36	2.5
G	Fe3+ 1	710.35, 709.75	1.15, 1.35	0.0, 10,000,000.0	710.35	1.15
H	Fe3+ 2	711.35, 709.75	1.25, 1.35	0.0, 9,520,000.0	711.13	1.25
I	Fe3+ 3	712.25, 709.75	1.35, 1.55	0.0, 7,320,000.0	712.25	1.55
J	Fe3+ 4	713.35, 709.75	1.35, 1.55	0.0, 3,980,000.0	709.75	1.55
K	Fe3+ 5	714.45, 709.75	1.85, 1.95	0.0, 3,010,000.0	713.65	1.95
L	Fe3+ 6	719.85, 709.75	2.65, 2.75	0.0, 3,270,000.0	719.85	2.75

.

This approach helps understand the redox transformations and their role in H₂ generation and mineral reactivity. The peak models provide constraints for the peak position, area, and the full width at half maximum (FWHM) for each of the Fe components input into the model. With known peak position and FWHM, the components are quantified as area percentages. The XPS spectra reveal variations in Fe oxidation states, with each rock type displaying distinct oxidation behavior.

The resolved components in the Fe core level peak, obtained through curve-fitting of the various iron oxides, are identified. The XPS spectra of the ultramafic rocks with different Fe-oxide components are provided in Figure 12, Figure 13 and Figure 14. Distinct variations in Fe oxidation were observed across different rock types and at various interaction times. This suggests that the hydro-geochemical reactions responsible for H₂ production differ among rock types, indicating that multiple reaction pathways are involved. These differences arise because each rock type exhibits unique mineralogical and geochemical properties, as well as crystallinity and surface reactivity, which influence how iron undergoes Fe oxidation–reduction cycles under identical reaction conditions.

In addition, the presence of trace minerals, such as biotite and zeolite, as identified by XRD, can further modulate reaction kinetics by acting as mediators, buffers, or inhibitors, which influence reactivity and the potential for H₂ generation. Even minor variations in mineral composition can introduce changes, alter reaction kinetics, and influence Fe²⁺/Fe³⁺ cycling. Consequently, the observed differences in Fe oxidation states across rock types reflect the complex interplay of bulk mineralogy, trace element chemistry, and reaction dynamics, ultimately leading to distinct pathways for hydrogen production.

The Fe²⁺/Fe³⁺ ratios obtained through deconvolution of the XPS spectra were plotted against the amount of H₂ generated by the various rocks, as presented in Figure 15. This is to investigate potential correlations between iron oxidation states and hydrogen production. Sample A showed a direct correlation between its Fe²⁺/Fe³⁺ ratio and hydrogen production. Although Sample B produced the highest amount of H₂ among the ultramafic rocks analyzed, no clear correlation was observed between its Fe²⁺/Fe³⁺ ratio and the hydrogen generated. This lack of correlation suggests that Fe oxidation alone is unlikely to be the primary controlling factor for H₂ generation in this rock type within the investigated time frame and reaction conditions. Among the other ultramafic samples, only Samples A and C exhibited an observable relationship between Fe²⁺/Fe³⁺ ratios and H₂ production, which will be discussed in detail in subsequent sections. However, no direct correlation was observed for most ultramafic rocks analyzed, indicating that additional geochemical factors, such as trace mineral composition, or secondary reaction pathways, may play a role in H₂ production.

For the mafic (basaltic) rocks (Figure 15 (Samples E and F)), Sample F showed an increasing trend between the Fe²⁺/Fe³⁺ ratio and the amount of H₂ generated. A positive relationship suggests that a lower proportion of Fe³⁺ was detected at the surface after reaction, as observed in the XPS analysis. This finding indicates that the oxidation of Fe²⁺ to Fe³⁺ was not the dominant mechanism driving mineral phase transformations and H₂ production in this sample over the investigated period (0 to 7 days). In contrast, Sample G did not show such a correlation, emphasizing the notion that variations in bulk and surface Fe oxidation states do not uniformly control H₂ generation across all rock types. Instead, other geochemical and mineralogical factors likely influence the reaction pathways involved in serpentinization and H₂ production.

3.4. Aqueous Fluids Analysis

3.4.1. Dissolution/Precipitation and Components Partitioning

The composition of the major elements in the supernatant, including Al, Mg, K, Ca, Zn, and Na, was analyzed to assess the geochemical evolution of the aqueous fluids. ICP-MS analysis of the aqueous fluids revealed that both dissolution and precipitation processes occurred in most of the samples, including both ultramafic and mafic rocks. However, each rock type exhibited distinct dissolution/precipitation kinetics, indicating that different reaction pathways contribute to H₂ generation.

For the ultramafic rocks, Sample A (Figure 16) showed an increase in Mg and Ca concentrations in the aqueous solution, accompanied by a slight decrease in Zn content. In contrast, Sample B showed minimal modification of the aqueous fluid composition after reaction. A reduction in Mg and Zn concentrations suggests that these elements were removed from the solution through precipitation, potentially forming secondary mineral phases. Despite their low concentrations, this implies that Mg- and Zn-bearing minerals in the bulk solids are more likely to precipitate than dissolve. Importantly, no significant dissolution was observed in this rock, indicating limited release of elements into the fluid phase. Sample C displayed the opposite trend, with a substantial increase in Mg concentration in the supernatant, suggesting that dissolution of Mg-bearing minerals occurred, releasing Mg into the fluid. The most notable change for Sample D (Figure 16) was an increase in K concentration in the reacted water compared to the initial unreacted water, indicating the dissolution of K-bearing minerals into solution, possibly due to phyllosilicate dissolution. For Sample E, a significant increase in Ca concentration after 7 days of hydrothermal reaction and a slight increase in Mg concentration were observed. This suggests that Ca-rich phases, such as clinopyroxenes or apatite, underwent dissolution, enriching the fluid with Ca.

Regarding the mafic rocks, notable geochemical changes were observed in the aqueous fluid composition, particularly in Sample G, where significant increases in K, Ca, and Zn concentrations were detected. This suggests that the dissolution of K-, Ca-, and Zn-bearing minerals occurred, releasing these elements into the fluid phase. The increase in Ca concentration is likely due to the breakdown of calcic silicates, resulting in the transformation of labradorite to andesine or the dissolution of clinopyroxene (diopside). The presence of high K levels points to the leaching of K-rich feldspars. In contrast, Sample F exhibited a more selective geochemical response, characterized by an increase in K concentration. This suggests that K-bearing minerals, such as phyllosilicate or mica, specifically biotite (K(Mg,Fe)₃AlSi₃O₁₀(OH,F)₂), were more susceptible to dissolution in this rock type. It has been argued that an elevated Si concentration in the reacting fluid can suppress H₂ production [25]. Aside from other solutes in the reacting fluid, the initial dissolved Si concentrations [25] in the reacting fluids can also influence the overall process of rock alteration [18,25]. Notably, no Si was detected in the aqueous phase for any of the rock–water systems investigated herein.

The variation in elemental release between these mafic and ultramafic rocks shows differences in mineral composition, solubility, and fluid–rock interactions, which influence the pathways of element mobilization and potential contributions to hydrogen generation. The differences in reaction kinetics emphasize that H₂ production is influenced not only by Fe oxidation but also by mineralogy, which impacts the solubility and mobility of key elements within the rock–fluid system.

3.4.2. Other Trace Elements

Traces of additional elements, including Ni, Cu, and Mo, were detected in the supernatant, indicating their mobilization during fluid–rock interactions, as presented in Figure 17. These elements are transition metals, whose oxides are commonly known for their geocatalytic properties in redox and hydrothermal reactions [29,30,31]. Their presence in the reaction fluid suggests potential geochemical transformations involving metal dissolution and reprecipitation.

Interestingly, apart from Samples D and F, all other rocks exhibited increased concentrations of one or more of these trace elements, implying that their presence was more prominent in certain rock types. The varied release of these metals could be linked to differences in mineral composition and changes in oxidation state. Given that Ni, Cu, and Mo are often associated with oxides or silicate minerals, their mobilization may result from the breakdown of specific mineral phases under hydrothermal conditions. We speculate that the presence of these transition metals could contribute to geocatalytic mechanisms in hydrogeochemical H₂ generation. Transition metals are known to facilitate electron transfer reactions, potentially accelerating the oxidation of Fe²⁺ to Fe³⁺ and enhancing H₂ production [29,30]. Their role in surface-mediated redox reactions and mineral–fluid interactions may influence reaction kinetics, affecting the overall efficiency of hydrogen generation in ultramafic and mafic rock systems.

While the leaching of trace elements such as Ni, Cu, and Mo coincided with enhanced H₂ yields in the same samples, these observations should be regarded as correlative rather than conclusive. Leaching data alone does not establish direct catalytic involvement, and other mineralogical or geochemical processes may contribute to the observed variability. Therefore, we interpret the role of trace elements as a potential but unconfirmed factor in H₂ production. Further investigation into the speciation, coordination environment, and reactivity of these metals could provide a deeper understanding of their influence on natural hydrogen production.

3.4.3. Physicochemical Properties

Figure 18 presents the total dissolved solids (TDS), conductivity, and pH of the reaction fluids. The TDS represents the total concentration of dissolved ions in the aqueous system, which originates from the dissolution of minerals during the rock–fluid interaction process. Electrical conductivity is directly related to the ionic strength of the fluid, as the presence of more dissolved ions enhances the conductivity of water. Conventionally, as TDS increases, conductivity also increases, and vice versa. This is because the breakdown of minerals releases ions such as Mg²⁺, Ca²⁺, K⁺, and Na⁺, as observed in their elemental form, which contribute to higher ionic concentrations and thereby increase conductivity. Conversely, a decrease in TDS and conductivity suggests that precipitation processes are occurring, where dissolved ions are removed from the solution and incorporated into newly formed mineral phases.

The pH of the supernatant water is another important parameter that reflects the nature of dissolution and precipitation reactions. For example, an increase in pH may indicate the precipitation of hydroxides or the consumption of protons in the reactions. Among the ultramafic rocks, the most pronounced pH increase was observed in Sample A and Sample C, while among the mafic rocks, Sample F exhibited a similar trend. However, using pH, TDS, and conductivity alone to predict the H₂ generation potential of rocks is challenging. Although both Sample A and Sample C showed an increase in pH, Sample A generated significantly more H₂, whereas Sample C produced the lowest amount of H₂ among the ultramafic rocks. Similarly, Sample G, a mafic rock, exhibited increased TDS, conductivity, and water pH, yet it produced less H₂ than Sample F, another mafic rock. These observations suggest that multiple geochemical factors, beyond just fluid pH and ion concentration, influence the efficiency of H₂ production in hydrothermal systems.

The observed changes in TDS, conductivity, and pH across different rock types suggest that hydrothermal dissolution and precipitation processes vary, influencing reaction kinetics and hydrogen generation efficiency. For instance, ultramafic rocks rich in olivine and pyroxene may show high Mg²⁺ and Fe²⁺ release, driving H₂ production via Fe oxidation, whereas mafic rocks with plagioclase and amphiboles may exhibit different dissolution patterns. Correlating TDS, conductivity, and pH trends with mineralogical changes, as observed in the aqueous fluids’ composition, is important in fluid-driven geochemical reactions that control hydrogen production in natural rock systems.

4. Discussion

4.1. H₂ Generation in Ultramafic and Mafic (Basaltic) Rocks

The hydrothermal reactions of different rock types exhibit distinct reactivity patterns. Even trace mineral components in the bulk rock composition are likely to significantly influence overall reaction kinetics and H₂ production. The study revealed that both bulk mineralogy and elemental-scale transformations contribute to the variability in H₂ generation pathways across different rock types. Although H₂ generation is linked to the oxidation of ferrous iron in rocks, the hydrous alteration of igneous rocks with similar ferrous iron content does not always result in equivalent H₂ amounts. For instance, basaltic and ultramafic rocks contain similar amounts of ferrous iron (with XPS atomic ratio 1 to 4% Fe), and one might expect their hydrous alteration to generate comparable amounts of H₂ through the oxidation of ferrous to ferric iron. However, both high and low H₂ yields were observed in ultramafic and mafic rocks, with ultramafic rocks producing as little as 1 mmol and as much as 9.6 mmol of H₂ per kilogram of rock. In contrast, the mafic rocks exhibited a broader range, from 0.36 mmol to 10.8 mmol, the latter being larger than most hydrogen generation rates of the ultramafic rocks in this study. This difference is primarily attributed to distinct reaction pathways, influenced by the presence of trace minerals that affect hydrogen generation mechanisms.

4.1.1. Reaction Pathways in Ultramafic Rocks

Among the ultramafic rocks, variations in H₂ generation rates and reaction mechanisms were apparent. Our results suggest that rocks with a higher abundance of Fe-rich minerals tend to exhibit greater H₂ generation, confirming that iron availability plays a vital role in hydrothermal hydrogen production. This trend was consistent across both ultramafic and basaltic rock types, reinforcing the significance of Fe-bearing minerals in driving redox reactions that lead to the release of H₂. However, despite this broad trend, individual rock types display variations in reaction pathways. For example, Sample B of ultramafic rock did not generate any H₂ during the short-term hydrothermal exposure (3 days). The noticeable mineral transformation was the conversion of clinopyroxene ($\mathrm{C}\mathrm{a}\left(\mathrm{M}\mathrm{g},\mathrm{F}\mathrm{e}\right){\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_6)$ and hematite $\left({\mathrm{F}\mathrm{e}}_2{\mathrm{O}}_3\right)$ with some olivine into increased amounts of lizardite (${\mathrm{M}\mathrm{g}}_3{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_5{\left(\mathrm{O}\mathrm{H}\right)}_4$), as indicated in the XRD, with this reaction defined by reaction 5. Significant mineral alterations and H₂ generation (the largest among the ultramafic rocks) were observed after 7 days of interaction (reaction 6). Most Fe-rich minerals, including olivine, hematite, and clinopyroxene (which decreased from 24.9 to 21.9 wt.%, 5.8 to 2.0 wt.%, and 19 to 3.1 wt.%, respectively), were depleted as shown in the XRD analysis. The primary newly formed minerals were serpentine (lizardite), which increased from 38.4 to 68.5 wt.%, and magnetite, which rose from 1.8 to 2.57 wt.%. Notably, no brucite was formed. Detailed information about the XRD mineral transformations is provided in Figure 4, Figure 5, Figure 6, Figure 7, Figure 8, Figure 9 and Figure 10 above. The mineral names and nominal stoichiometry are assigned based on a restricted compositional range within a “solid solution series,” characterized by the free isomorphic substitution of two elements (e.g., Fe and Mg) within the crystal lattice (see [43]). The reaction equations are approximated based on the nominal stoichiometry of each mineral identified through XRD; the elements in parentheses may freely substitute for one another.

Sample B 3 days

$$\mathrm{O}\mathrm{l}\mathrm{i}\mathrm{v}\mathrm{i}\mathrm{n}\mathrm{e}+\mathrm{L}\mathrm{i}\mathrm{z}\mathrm{a}\mathrm{r}\mathrm{d}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{B}\mathrm{i}\mathrm{r}\mathrm{n}\mathrm{e}\mathrm{s}\mathrm{s}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{H}\mathrm{e}\mathrm{m}\mathrm{a}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{W}\mathrm{a}\mathrm{t}\mathrm{e}\mathrm{r}\to \mathrm{L}\mathrm{i}\mathrm{z}\mathrm{a}\mathrm{r}\mathrm{d}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{O}\mathrm{l}\mathrm{i}\mathrm{v}\mathrm{i}\mathrm{n}\mathrm{e}+\mathrm{B}\mathrm{i}\mathrm{r}\mathrm{n}\mathrm{e}\mathrm{s}\mathrm{s}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{H}\mathrm{e}\mathrm{m}\mathrm{a}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}$$

$$\begin{gathered} {\left(\mathrm{M}\mathrm{g},\mathrm{F}\mathrm{e}\right)}_2\mathrm{S}\mathrm{i}{\mathrm{O}}_4+{\mathrm{M}\mathrm{g}}_3{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_5{\left(\mathrm{O}\mathrm{H}\right)}_4+\mathrm{M}\mathrm{n}{\left(\mathrm{O}\mathrm{H}\right)}_4+{\mathrm{F}\mathrm{e}}_2{\mathrm{O}}_3+{\mathrm{H}}_2\mathrm{O} \\ \to {\mathrm{M}\mathrm{g}}_3{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_5{\left(\mathrm{O}\mathrm{H}\right)}_4+{\left(\mathrm{M}\mathrm{n},\mathrm{F}\mathrm{e}\right)}_2\mathrm{S}\mathrm{i}{\mathrm{O}}_4+\mathrm{M}\mathrm{n}{\left(\mathrm{O}\mathrm{H}\right)}_4+{\mathrm{F}\mathrm{e}}_2{\mathrm{O}}_3 \end{gathered}$$

(5)

Sample B 7 days

$$\begin{gathered} \mathrm{L}\mathrm{i}\mathrm{z}\mathrm{a}\mathrm{r}\mathrm{d}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{B}\mathrm{i}\mathrm{r}\mathrm{n}\mathrm{e}\mathrm{s}\mathrm{s}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{O}\mathrm{l}\mathrm{i}\mathrm{n}\mathrm{i}\mathrm{v}\mathrm{e}+\mathrm{C}\mathrm{l}\mathrm{i}\mathrm{n}\mathrm{o}\mathrm{p}\mathrm{y}\mathrm{r}\mathrm{o}\mathrm{x}\mathrm{e}\mathrm{n}\mathrm{e}+\mathrm{M}\mathrm{a}\mathrm{g}\mathrm{n}\mathrm{e}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{H}\mathrm{e}\mathrm{m}\mathrm{a}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{W}\mathrm{a}\mathrm{t}\mathrm{e}\mathrm{r} \\ \to \mathrm{L}\mathrm{i}\mathrm{z}\mathrm{a}\mathrm{r}\mathrm{d}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{B}\mathrm{i}\mathrm{r}\mathrm{n}\mathrm{e}\mathrm{s}\mathrm{s}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{O}\mathrm{l}\mathrm{i}\mathrm{v}\mathrm{i}\mathrm{n}\mathrm{e}+\mathrm{C}\mathrm{l}\mathrm{i}\mathrm{n}\mathrm{o}\mathrm{p}\mathrm{y}\mathrm{r}\mathrm{o}\mathrm{x}\mathrm{e}\mathrm{n}\mathrm{e}+\mathrm{m}\mathrm{a}\mathrm{g}\mathrm{n}\mathrm{e}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{H}\mathrm{e}\mathrm{m}\mathrm{a}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{H}\mathrm{y}\mathrm{d}\mathrm{r}\mathrm{o}\mathrm{g}\mathrm{e}\mathrm{n} \end{gathered}$$

$$\begin{gathered} {\mathrm{M}\mathrm{g}}_3{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_5{\left(\mathrm{O}\mathrm{H}\right)}_4+\mathrm{M}\mathrm{n}{\left(\mathrm{O}\mathrm{H}\right)}_4+{\left(\mathrm{M}\mathrm{g},\mathrm{F}\mathrm{e}\right)}_2\mathrm{S}\mathrm{i}{\mathrm{O}}_4+\mathrm{C}\mathrm{a}\left(\mathrm{M}\mathrm{g},\mathrm{F}\mathrm{e}\right){\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_6+{\mathrm{F}\mathrm{e}}_3{\mathrm{O}}_4+{\mathrm{F}\mathrm{e}}_2{\mathrm{O}}_3+{\mathrm{H}}_2\mathrm{O}+{\mathrm{H}}^{+} \\ \to {\mathrm{M}\mathrm{g}}_3{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_5{\left(\mathrm{O}\mathrm{H}\right)}_4+\mathrm{M}\mathrm{n}{\left(\mathrm{O}\mathrm{H}\right)}_4+{\left(\mathrm{M}\mathrm{g},\mathrm{F}\mathrm{e}\right)}_2\mathrm{S}\mathrm{i}{\mathrm{O}}_4+\mathrm{C}\mathrm{a}\left(\mathrm{M}\mathrm{g},\mathrm{F}\mathrm{e}\right){\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_6+{\mathrm{F}\mathrm{e}}_3{\mathrm{O}}_4+{\mathrm{F}\mathrm{e}}_2{\mathrm{O}}_3+{\mathrm{H}}_2+{\mathrm{O}\mathrm{H}}^{-} \end{gathered}$$

(6)

For Sample C, this rock type produced the lowest H₂ yield among the ultramafic rocks, generating only 0.76 mmol of H₂ per kilogram of rock. After 7 days of reaction, the most notable mineral transformations included the decrease in olivine and vermiculite ((Mg,Fe,Al)₆(Si,Al)₄O₁₀(OH)₈), accompanied by the formation of hematite and brucite through the alteration of Fe-Mg silicates like olivine. These mineral modifications contributed to the changes in Mg concentration in the aqueous solution, as shown in Figure 16 (Sample B). The approximate chemical reaction pathways for these transformations can be represented by Reaction 7. Although serpentine formation was not the sole mineral transformation, H₂ generation in these rocks appears to be primarily driven by the presence of Fe-containing minerals.

Sample C 3 to 7 days interaction

$$\mathrm{O}\mathrm{l}\mathrm{i}\mathrm{v}\mathrm{i}\mathrm{n}\mathrm{e}+\mathrm{L}\mathrm{i}\mathrm{z}\mathrm{a}\mathrm{r}\mathrm{d}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{h}\mathrm{e}\mathrm{m}\mathrm{a}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{b}\mathrm{r}\mathrm{u}\mathrm{c}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{w}\mathrm{a}\mathrm{t}\mathrm{e}\mathrm{r}\to \mathrm{O}\mathrm{l}\mathrm{i}\mathrm{v}\mathrm{i}\mathrm{n}\mathrm{e}+\mathrm{L}\mathrm{i}\mathrm{z}\mathrm{a}\mathrm{r}\mathrm{d}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{h}\mathrm{e}\mathrm{m}\mathrm{a}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{b}\mathrm{r}\mathrm{u}\mathrm{c}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{H}\mathrm{y}\mathrm{d}\mathrm{r}\mathrm{o}\mathrm{g}\mathrm{e}\mathrm{n}$$

$$\begin{gathered} (\mathrm{F}\mathrm{e},{\mathrm{M}\mathrm{g})}_2\mathrm{S}\mathrm{i}{\mathrm{O}}_4+{\mathrm{M}\mathrm{g}}_3{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_5{\left(\mathrm{O}\mathrm{H}\right)}_4+{\mathrm{F}\mathrm{e}}_2{\mathrm{O}}_3+\mathrm{M}\mathrm{g}{\left(\mathrm{O}\mathrm{H}\right)}_2+{\mathrm{H}}_2\mathrm{O}+{\mathrm{H}}^{+} \\ \to (\mathrm{F}\mathrm{e},{\mathrm{M}\mathrm{g})}_2\mathrm{S}\mathrm{i}{\mathrm{O}}_4+{\mathrm{M}\mathrm{g}}_3{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_5{\left(\mathrm{O}\mathrm{H}\right)}_4+{\mathrm{F}\mathrm{e}}_2{\mathrm{O}}_3+\mathrm{M}\mathrm{g}{\left(\mathrm{O}\mathrm{H}\right)}_2+{\mathrm{H}}_2+{\mathrm{O}\mathrm{H}}^{-} \end{gathered}$$

(7)

Although serpentine formation was not the only mineral transformation observed, hematite was also formed, as evidenced by its increased XRD composition in Sample C. It is important to note that this is the only rock sample that exhibited an increase in hematite content, whereas all other analyzed samples showed a decreasing trend compared to the untreated rocks. In these rocks, H₂ generation is primarily attributed to the presence of Fe-containing minerals based on their mineralogical composition. The ultramafic rocks did not exhibit a consistent trend between H₂ production and the formation of brucite or magnetite, as shown in Figure 4, Figure 5, Figure 6, Figure 7, Figure 8, Figure 9 and Figure 10. Specifically, magnetite decreased (or was consumed) in Sample A, increased in Sample B, was not detected in Samples C and D, and showed no significant change in Sample E. For brucite, precipitation occurred in Samples A, C, and D after hydrothermal reaction, while Samples B and E showed no change over the 3- and 7-day interaction periods. Particularly, the reaction pathway for Sample A is similar to that of Sample E, with the exception of greater serpentine formation in Sample E over the 3- and 7-day reaction durations. Similarly, the reaction pathway or mechanism responsible for H₂ generation in Sample C resembles that of Sample D. The key difference lies in the transformation of hematite, which increased in Sample C but decreased in Sample D from the untreated state through the 3- and 7-day hydrothermal reactions. This unique feature in the mineralogical transformation of Sample C, among all the examined ultramafic rocks, could be the primary reason for the lowest H₂ generation observed compared to the other ultramafic samples.

Thermodynamic simulations yielded similar results, indicating that magnetite is unlikely to form during the serpentinization of orthopyroxene, as all Fe initially present in orthopyroxene is instead incorporated into serpentine [13]. Meanwhile, ferroan brucite oxidation has been proposed as a major contributor to H₂ production in ophiolites at temperatures below 423 K [44,45]. This aligns with petrographic observations of natural samples, which indicate that brucite undergoes oxidation to form magnetite under open-system conditions [46,47].

The thermodynamic simulation work of [18] demonstrated that H₂ production is strongly influenced by rock composition. In particular, the serpentinization process in lithologies with larger magnesium content results in significantly greater H₂ generation, attributing higher H₂ generation to increased MgO composition. Specifically, rocks with MgO content exceeding 35 wt.% have the highest potential for H₂ generation. However, the amounts vary significantly due to differences in their ability to stabilize various Fe(III)-bearing phases during the process of hydrous alteration. Even among rocks with similar MgO content, variations in SiO₂ levels influence the precipitation of Fe-bearing secondary minerals. Higher SiO₂ concentrations promote the formation of greenalite over cronstedtite, resulting in lower H₂ production, as Fe(II) from primary minerals is incorporated into greenalite without oxidation [24,25]. In contrast, rocks with MgO content below 20 wt.% do not produce significant H₂ at any water-to-rock ratio in their simulations [18].

It is important to mention that the thermodynamic computations by [13] reported a somewhat contrasting finding: the serpentinization of Fe-rich olivine within the forsterite–fayalite (Fo-Fa) solid solution led to the formation of a free H₂ gas phase, with H₂ generation increasing from Fo₁₀-Fa to Fo₇₀-Fa. The complete diadochy between Mg²⁺ and Fe²⁺ results in a broad range of olivine compositions, ranging from forsterite to fayalite [48]. The chemical makeup of olivine significantly affects its stability when exposed to water [13]. Consistent with our experimental findings, mineral alteration and H₂ generation are highly dependent on the specific starting minerals.

It was documented that in the early stages of serpentinization, Fe-bearing serpentine minerals commonly form, incorporating ferric iron within their structure [18]. Interestingly, not all H₂-generating hydrothermal reactions resulted in serpentine ((Mg_2.70Fe_0.18Al_0.11)(Si_1.81Al_0.19O₅)(OH)₄) formation. This challenges the conventional assumption that H₂ generation in Fe-rich rocks is always linked to serpentinization. Instead, in some cases, birnessite (a Mn-rich layered hydroxide mineral, (H_4.424MnO_4.212)), and hematite (Sample C), could form as a secondary mineral in hydrothermal environments, particularly from the alteration of manganoan olivine and pyroxene, was identified as a major reaction products. This distinction shows that alternative mineralogical pathways contribute to H₂ generation beyond traditional serpentinization reactions.

4.1.2. Reaction Pathways in Mafic (Basalts) Rocks

For the mafic rocks (Samples F and G), there is no evidence of serpentine formation. H₂ production in these systems follows a different mineralogical transformation path. Instead, the secondary minerals formed were andesine and anorthite, both of which belong to the plagioclase feldspar group. This aligns with recent findings that indicate H₂ can be generated from basalts [49]. Basaltic rocks generate H₂ through mechanisms that differ from those governing Fe-rich ultramafic systems, involving feldspar alteration and Fe-rich mineral transformations rather than typical ultramafic serpentinization pathways.

The reactive mineral assemblage in basalts is dominated by Mg- and Fe-bearing phases, with olivine present in lower modal abundance relative to ultramafic counterparts. These compositional differences significantly influence mineral–fluid reactivity and the efficiency of subsequent H₂ generation during water–rock interaction. These minerals include orthopyroxene ($\left({\mathrm{M}\mathrm{g},\mathrm{F}\mathrm{e})}_2\right){\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_6$), enstatite (${\mathrm{M}\mathrm{g}}_2{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_6$), diopside ($\mathrm{C}\mathrm{a}\mathrm{M}\mathrm{g}{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_6),$ and biotite ($\mathrm{K}{\left(\mathrm{M}\mathrm{g},\mathrm{F}\mathrm{e}\right)}_3{\mathrm{A}\mathrm{l}\mathrm{S}\mathrm{i}}_3{\mathrm{O}}_{10}{\left(\mathrm{O}\mathrm{H}\right)}_2$). The increased concentrations of K and Ca observed in the aqueous solution, as shown in Figure 16 (Samples F and G), provide clear evidence of the dissolution of biotite and diopside, respectively. Biotite, a K-rich phyllosilicate mineral, releases potassium ions (K⁺) into the solution as it undergoes chemical breakdown. Similarly, diopside, a Ca-Mg pyroxene, contributes to the increased Ca²⁺ concentrations through its dissolution. This facilitates geochemical reactions that drive the generation of H₂. The breakdown of Fe- and Mg-bearing silicates promotes redox reactions where ferrous iron (Fe²⁺) within the mineral structure can be oxidized, concurrently reducing water to produce molecular H₂ as presented in Reactions 8 and 9. This process is particularly significant in mafic rock systems, where Fe²⁺-rich phases readily participate in redox-driven H₂ generation.

Sample F 3 days

$$\mathrm{L}\mathrm{a}\mathrm{b}\mathrm{r}\mathrm{a}\mathrm{d}\mathrm{o}\mathrm{r}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{A}\mathrm{n}\mathrm{d}\mathrm{e}\mathrm{s}\mathrm{i}\mathrm{n}\mathrm{e}+\mathrm{E}\mathrm{n}\mathrm{s}\mathrm{t}\mathrm{a}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{W}\mathrm{a}\mathrm{t}\mathrm{e}\mathrm{r}\to \mathrm{L}\mathrm{a}\mathrm{b}\mathrm{r}\mathrm{a}\mathrm{d}\mathrm{o}\mathrm{r}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{A}\mathrm{n}\mathrm{d}\mathrm{e}\mathrm{s}\mathrm{i}\mathrm{n}\mathrm{e}+\mathrm{E}\mathrm{n}\mathrm{s}\mathrm{t}\mathrm{a}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}$$

$$\begin{gathered} 2{\mathrm{A}\mathrm{l}}_{0.81}{\mathrm{C}\mathrm{a}}_{0.33}{\mathrm{N}\mathrm{a}}_{0.16}{\mathrm{S}\mathrm{i}}_{1.19}{\mathrm{O}}_4+{\mathrm{A}\mathrm{l}}_{0.74}{\mathrm{C}\mathrm{a}}_{0.24}{\mathrm{N}\mathrm{a}}_{0.26}{\mathrm{S}\mathrm{i}}_{1.27}{\mathrm{O}}_4+{\mathrm{M}\mathrm{g}}_2{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_6+{\mathrm{H}}_2\mathrm{O} \\ \to {\mathrm{A}\mathrm{l}}_{0.81}{\mathrm{C}\mathrm{a}}_{0.33}{\mathrm{N}\mathrm{a}}_{0.16}{\mathrm{S}\mathrm{i}}_{1.19}{\mathrm{O}}_4+{\mathrm{A}\mathrm{l}}_{0.74}{\mathrm{C}\mathrm{a}}_{0.24}{\mathrm{N}\mathrm{a}}_{0.26}{\mathrm{S}\mathrm{i}}_{1.27}{\mathrm{O}}_4+{\mathrm{M}\mathrm{g}}_2{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_6+{\mathrm{H}}_2\mathrm{O} \end{gathered}$$

(8)

Sample F 7 days

$$\mathrm{L}\mathrm{a}\mathrm{b}\mathrm{r}\mathrm{a}\mathrm{d}\mathrm{o}\mathrm{r}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{B}\mathrm{i}\mathrm{o}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{E}\mathrm{n}\mathrm{s}\mathrm{t}\mathrm{a}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{O}\mathrm{l}\mathrm{i}\mathrm{v}\mathrm{i}\mathrm{n}\mathrm{e}+\mathrm{W}\mathrm{a}\mathrm{t}\mathrm{e}\mathrm{r}\to \mathrm{L}\mathrm{a}\mathrm{b}\mathrm{r}\mathrm{a}\mathrm{d}\mathrm{o}\mathrm{r}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{A}\mathrm{n}\mathrm{d}\mathrm{e}\mathrm{s}\mathrm{i}\mathrm{n}\mathrm{e}+\mathrm{B}\mathrm{i}\mathrm{o}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}\left(\mathrm{M}\mathrm{i}\mathrm{c}\mathrm{a}\right)+\mathrm{E}\mathrm{n}\mathrm{s}\mathrm{t}\mathrm{a}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{O}\mathrm{l}\mathrm{i}\mathrm{v}\mathrm{i}\mathrm{n}\mathrm{e}+\mathrm{H}\mathrm{y}\mathrm{d}\mathrm{r}\mathrm{o}\mathrm{g}\mathrm{e}\mathrm{n}$$

$$\begin{gathered} 2{\mathrm{A}\mathrm{l}}_{0.81}{\mathrm{C}\mathrm{a}}_{0.33}{\mathrm{N}\mathrm{a}}_{0.16}{\mathrm{S}\mathrm{i}}_{1.19}{\mathrm{O}}_4+\mathrm{K}{\left(\mathrm{M}\mathrm{g},\mathrm{F}\mathrm{e}\right)}_3\left(\mathrm{A}\mathrm{l}{\mathrm{S}\mathrm{i}}_3{\mathrm{O}}_{10}\right){\left(\mathrm{O}\mathrm{H}\right)}_2+{\mathrm{M}\mathrm{g}}_2{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_6+{\left(\mathrm{M}\mathrm{g},\mathrm{F}\mathrm{e}\right)}_2\mathrm{S}\mathrm{i}{\mathrm{O}}_4+{\mathrm{H}}_2\mathrm{O}+{\mathrm{H}}^{+} \\ {\to \mathrm{A}\mathrm{l}}_{0.81}{\mathrm{C}\mathrm{a}}_{0.33}{\mathrm{N}\mathrm{a}}_{0.16}{\mathrm{S}\mathrm{i}}_{1.19}{\mathrm{O}}_4+{\mathrm{A}\mathrm{l}}_{0.74}{\mathrm{C}\mathrm{a}}_{0.24}{\mathrm{N}\mathrm{a}}_{0.26}{\mathrm{S}\mathrm{i}}_{1.27}{\mathrm{O}}_4+\mathrm{K}{\left(\mathrm{M}\mathrm{g},\mathrm{F}\mathrm{e}\right)}_3\left(\mathrm{A}\mathrm{l}{\mathrm{S}\mathrm{i}}_3{\mathrm{O}}_{10}\right){\left(\mathrm{O}\mathrm{H}\right)}_2+{(\mathrm{M}\mathrm{g},\mathrm{F}\mathrm{e})}_2\mathrm{S}\mathrm{i}{\mathrm{O}}_4+{\mathrm{M}\mathrm{g}}_2{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_6+{\mathrm{O}\mathrm{H}}^{-}+{\mathrm{H}}_2 \end{gathered}$$

(9)

Thus, the increasing K and Ca concentrations not only confirm the reactivity of biotite and diopside in the rock–fluid system but also serve as indirect indicators of mineral–fluid interactions that contribute to enhanced H₂ production. As primary sources of Mg and Fe, these minerals are essential for H₂ generation in basalts. Notably, Sample E contains a higher quantitative abundance of these reactive Mg- and Fe-rich phases compared to Sample G (refer to Figure 9 and Figure 10). This greater mineral availability enhances H₂ generation in Sample E. The difference in H₂ yield between the two mafic samples can thus be attributed to variations in mineral composition and reactivity, emphasizing the significance of Fe-Mg silicates in influencing H₂ production rates during rock–fluid interactions. The major minerals’ transformation pathway for Sample G is illustrated in Reaction 10.

Sample G 3 to 7 days

$$\mathrm{L}\mathrm{a}\mathrm{b}\mathrm{r}\mathrm{a}\mathrm{d}\mathrm{o}\mathrm{r}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{A}\mathrm{n}\mathrm{o}\mathrm{r}\mathrm{t}\mathrm{h}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{A}\mathrm{u}\mathrm{g}\mathrm{i}\mathrm{t}\mathrm{e}\left(\mathrm{E}\mathrm{n}\mathrm{s}\mathrm{t}\mathrm{a}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{D}\mathrm{i}\mathrm{o}\mathrm{p}\mathrm{s}\mathrm{i}\mathrm{t}\mathrm{e}\right)+\mathrm{w}\mathrm{a}\mathrm{t}\mathrm{e}\mathrm{r}\to \mathrm{L}\mathrm{a}\mathrm{b}\mathrm{r}\mathrm{a}\mathrm{d}\mathrm{o}\mathrm{r}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{A}\mathrm{n}\mathrm{d}\mathrm{e}\mathrm{s}\mathrm{i}\mathrm{n}\mathrm{e}+\mathrm{A}\mathrm{n}\mathrm{o}\mathrm{r}\mathrm{t}\mathrm{h}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{A}\mathrm{u}\mathrm{g}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{H}\mathrm{y}\mathrm{d}\mathrm{r}\mathrm{o}\mathrm{g}\mathrm{e}\mathrm{n}$$

$$\begin{gathered} 2{\mathrm{A}\mathrm{l}}_{0.81}{\mathrm{C}\mathrm{a}}_{0.33}{\mathrm{N}\mathrm{a}}_{0.16}{\mathrm{S}\mathrm{i}}_{1.19}{\mathrm{O}}_4+\mathrm{C}\mathrm{a}{\mathrm{A}\mathrm{l}}_2{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_8+{\mathrm{M}\mathrm{g}}_2{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_6+\mathrm{C}\mathrm{a}\mathrm{M}\mathrm{g}{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_6+{\mathrm{H}}_2\mathrm{O} \\ \to {\mathrm{A}\mathrm{l}}_{0.81}{\mathrm{C}\mathrm{a}}_{0.33}{\mathrm{N}\mathrm{a}}_{0.16}{\mathrm{S}\mathrm{i}}_{1.19}{\mathrm{O}}_4+{\mathrm{A}\mathrm{l}}_{0.74}{\mathrm{C}\mathrm{a}}_{0.24}{\mathrm{N}\mathrm{a}}_{0.26}{\mathrm{S}\mathrm{i}}_{1.27}{\mathrm{O}}_4+\mathrm{C}\mathrm{a}{\mathrm{A}\mathrm{l}}_2{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_8+(\mathrm{C}\mathrm{a},\mathrm{N}\mathrm{a})(\mathrm{M}\mathrm{g},\mathrm{F}\mathrm{e},\mathrm{A}\mathrm{l}){(\mathrm{S}\mathrm{i},\mathrm{A}\mathrm{l})}_2{\mathrm{O}}_6+{\mathrm{H}}_2 \end{gathered}$$

(10)

To determine the reactive components in basalts responsible for H₂ production, ref. [49] conducted experiments on typical basaltic mineral phases, including olivine, pyroxene, and feldspar. Specifically, the authors investigated basalt minerals, including labradorite, augite, and olivine (fayalite and forsterite). Their findings indicated that minor ferrous minerals present in basalts, such as magnetite (Fe₃O₄), ilmenite (FeTiO₃), hematite (Fe₂O₃, a ferric oxide), pyrite (FeS₂), and pyrrhotite (FeS), did not generate detectable levels of H₂. However, olivine end-members—fayalite and forsterite—and augite generated H₂. In addition, ferrous iron-bearing silicate minerals, including biotite, produced only trace amounts of H₂. The study concluded that ferrous silicate minerals, particularly pyroxene and olivine, were the primary contributors to H₂ generation in basalts, suggesting that Fe²⁺ was a key reactant in the H₂-generating reaction.

Considering the elemental and atomic-scale processes, the results revealed that Fe²⁺ oxidation to Fe³⁺ correlated with H₂ generation in only certain rock types or at a later stage of the reaction progression. This suggests that while Fe oxidation is an important mechanism, it is likely not the sole driver of H₂ production; additional geochemical factors, such as the availability of trace elements and mineral dissolution/precipitation rates, also contribute to H₂ generation in certain rocks.

4.2. Extended Reactions of Sample A and Its Impact on H₂ Generation

Sample A (dunite, ultramafic rock) was subjected to extended reaction durations to better understand mineral phase transformations and H₂ generation dynamics over time. This approach aimed to capture both early-stage reactivity and intermediate to long-term trends in mineral-fluid interactions. A comparison of the Fe oxidation states using the Fe²⁺/Fe³⁺ ratio, mineral transformations, and H₂ generation is presented in Figure 19a,b.

4.2.1. Early-Stage H₂ Generation (0–7 Days)

During the early hydrothermal exposure (0–7 days), H₂ generation was rapid, reaching up to 8 mmol/kg of rock (see Figure 19a). The high rate of hydrogen generation in this early stage suggests that multiple geochemical mechanisms are actively contributing to H₂ release. First, the oxidation of Fe²⁺ to Fe³⁺: The conversion of Fe-bearing minerals, including olivine and lizardite, to Fe³⁺-rich phases (including magnetite, Fe₃O₄) releases electrons, which can facilitate water reduction to H₂. The approximate reaction at this stage is represented by Reaction 11.

Sample A 0–7 days

$$\mathrm{O}\mathrm{l}\mathrm{i}\mathrm{v}\mathrm{i}\mathrm{n}\mathrm{e}+\mathrm{L}\mathrm{i}\mathrm{z}\mathrm{a}\mathrm{r}\mathrm{d}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{M}\mathrm{a}\mathrm{g}\mathrm{n}\mathrm{e}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{H}\mathrm{e}\mathrm{m}\mathrm{a}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{w}\mathrm{a}\mathrm{t}\mathrm{e}\mathrm{r}\to \mathrm{O}\mathrm{l}\mathrm{i}\mathrm{v}\mathrm{i}\mathrm{n}\mathrm{e}+\mathrm{L}\mathrm{i}\mathrm{z}\mathrm{a}\mathrm{r}\mathrm{d}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{M}\mathrm{a}\mathrm{g}\mathrm{n}\mathrm{e}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{H}\mathrm{e}\mathrm{m}\mathrm{a}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{W}\mathrm{u}\mathrm{e}\mathrm{s}\mathrm{t}\mathrm{i}\mathrm{t}\mathrm{e}+\mathrm{H}\mathrm{y}\mathrm{d}\mathrm{r}\mathrm{o}\mathrm{g}\mathrm{e}\mathrm{n}$$

$$\begin{gathered} {\left(\mathrm{M}\mathrm{g},\mathrm{F}\mathrm{e}\right)}_2\mathrm{S}\mathrm{i}{\mathrm{O}}_4+2{\mathrm{M}\mathrm{g}}_3{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_5{\left(\mathrm{O}\mathrm{H}\right)}_4+{2\mathrm{F}\mathrm{e}}_3{\mathrm{O}}_4+4{\mathrm{F}\mathrm{e}}_2{\mathrm{O}}_3+{\mathrm{H}}_2\mathrm{O}+{3\mathrm{H}}^{+} \\ \to {\left(\mathrm{M}\mathrm{g},\mathrm{F}\mathrm{e}\right)}_2\mathrm{S}\mathrm{i}{\mathrm{O}}_4+2{\mathrm{M}\mathrm{g}}_3{\mathrm{S}\mathrm{i}}_2{\mathrm{O}}_5{\left(\mathrm{O}\mathrm{H}\right)}_4+{3\mathrm{F}\mathrm{e}}_3{\mathrm{O}}_4+{\mathrm{F}\mathrm{e}}_2{\mathrm{O}}_3+3\mathrm{F}\mathrm{e}\mathrm{O}+{\mathrm{H}}_2+{3\mathrm{O}\mathrm{H}}^{-} \end{gathered}$$

(11)

Additionally, the presence of preformed reactive intermediates, such as metastable Fe-bearing phases or partially oxidized minerals, enhances H₂ generation by promoting electron transfer reactions. These intermediates include wuestite (FeO), clinopyroxenes, and brucite, which were precipitated and then dissolved as the interaction proceeded (see Figure 4). Their presence suggests that these Fe-rich minerals exist in a reactive, transient state, allowing them to rapidly participate in redox reactions with water. As these metastable phases undergo further oxidation and Fe partitioning, along with olivine conversion (decreasing from 59.5 wt.% in the unreacted rock to 49.8 wt.% after 7 days of interaction) into lizardite, they contribute to an initial burst of H₂ production, increasing H₂ output in the early stages of the hydrothermal interactions. However, as these reactive intermediates are depleted or converted into more stable mineral phases, such as lizardite and birnessite, the rate of H₂ generation slows down. This highlights the significance of mineral composition and oxidation state variability in controlling the temporal dynamics of hydrothermal hydrogen production.

This stage is characterized by an increase in the Fe²⁺/Fe³⁺ ratio (see Figure 19a), suggesting that the oxidation of ferrous iron (Fe²⁺) to ferric iron (Fe³⁺) was not the dominant reaction pathway during this period of interaction. Instead, other geochemical processes, such as mineral dissolution, solid-state Fe redistribution, Fe²⁺ incorporation into newly forming phases, may have played a more significant role in controlling the Fe redox balance and resulting H₂ generation. The persistence of Fe²⁺ suggests that either oxidation was kinetically limited or competing reduction processes were occurring, maintaining a higher Fe²⁺/Fe³⁺ ratio.

Additionally, the presence of metastable Fe-bearing minerals, such as wuestite and apatite, alongside increasing serpentinization (evidenced by the rise in lizardite formation, as shown in Figure 19b), may have contributed to buffering Fe oxidation. This, in turn, could have slowed or altered the typical redox progression associated with Fe-driven H₂ generation.

4.2.2. Intermediate-Stage H₂ Generation (7–28 Days)

Between 7 and 28 days of the hydrothermal interaction, a decline in the Fe²⁺/Fe³⁺ ratio was observed, coinciding with a decrease in the rate of hydrogen generation. This shift suggests that the early-stage reactive intermediates were depleted, leaving Fe²⁺ oxidation to Fe³⁺ as the dominant mechanism driving H₂ production in the intermediate phase. The reduction in Fe²⁺ availability may have limited further electron transfer reactions, leading to a gradual decrease in H₂ output. The formation of a stable Fe³⁺-rich phase (such as serpentine) may have reduced the reactivity of iron minerals, further slowing the rate of hydrogen generation. Precipitation of secondary mineral phases, including lizardite and birnessite, may have altered the surface reactivity of the rock, modifying the kinetics of Fe oxidation and H₂ release, as well as the approximate hydrogeothermal reaction for this stage. In rock systems, this precipitation could lead to surface passivation.

Hydrogen generation in ultramafic rock systems is a complex process, indicating that different reaction pathways dominate at various stages of hydrothermal alteration. Serpentinization and Fe oxidation contribute to H₂ generation through a time-dependent, multi-step reaction process. In the early phase, both Fe oxidation and reactive intermediates contribute to the release of H₂. In the intermediate phase, Fe²⁺ oxidation becomes the primary mechanism, but reaction rates slow as Fe²⁺ is gradually depleted. While the early stage appears to involve a combination of Fe²⁺ oxidation and additional reactive intermediates, the long-term reaction dynamics are primarily controlled by iron oxidation alone. As a result, less H₂ is produced since iron oxidation occurs to a lesser extent compared to scenarios where serpentine forms extensively. Recent studies have highlighted that non-redox-sensitive elements in rocks, such as Mg, Si, Al, and Ca, play a significant role in controlling the distribution and transformation of Fe during secondary mineralization, thereby influencing the redox processes that generate reduced volatiles during fluid–rock interactions. Most of the iron released from primary minerals is integrated into secondary minerals with minimal oxidation, as certain minerals preferentially incorporate ferrous iron into their crystal structures [18].

The availability of Fe²⁺-bearing minerals and the reaction pathway are important in sustaining H₂ production over extended periods. The mineralogical transformations, effects of trace metals, and changes in fluid composition could provide a deeper understanding and optimize conditions for prolonged and efficient H₂ generation from ultramafic and basaltic rock systems.

4.3. Potential Geocatalysis During H₂ Generation

The presence of transition metals such as Ni, Mo, and Cu within ultramafic and mafic rocks suggests a potential geocatalytic effect [29,30,31] that enhances H₂ generation during fluid–rock interactions (see Section 3.4.2). These metals can facilitate electron transfer reactions, accelerating Fe²⁺ oxidation and promoting sustained H₂ production beyond typical serpentinization pathways. Additionally, the structural properties of secondary minerals, such as birnessite, vermiculite, and zeolite (as observed by XRD in the basalts), may further support geocatalytic pathways by preserving reactive surfaces, thereby enabling continuous hydrothermal hydrogen evolution over extended periods [27,50]. However, additional research is needed to fully assess the extent of such reactive mediators in geologic H₂ generation.

Recent studies indicate that transition-metal sites within ultramafic and mafic rocks can act as true geocatalysts that enhance abiotic H₂ production beyond what is expected from simple Fe(II) oxidation and serpentinization alone [51,52]. Nickel and Ni–Fe alloys (e.g., awaruite and native Fe–Ni phases) promote H₂ evolution by providing surface sites for electron transfer and by lowering activation barriers for H-forming redox reactions. Experimental and theoretical studies show that Ni²⁺ and Ni–Fe nanoparticles markedly increase H₂ generation rates under low-temperature aqueous conditions. The introduction of only 1% Ni²⁺ boosted the H₂ generation rate by nearly two orders of magnitude at 90 °C [53,54]. Similarly, Fe-rich phases and Fe–Ni alloys formed during alteration can catalyze hydrogen-producing reactions and facilitate redox cycling between reduced metal sites and aqueous species, sustaining H₂ production even after initial mineral transformations. These catalytic pathways are complementary to, and sometimes kinetically dominant over, mineral transformation-controlled H₂ release, meaning that rock compositions rich in catalytic transition metals (Ni, Fe, and trace PGE (platinum-group elements)) may produce sustained H₂ fluxes through surface-mediated reactions and alloy catalysis [51,52,53,54].

Meanwhile, aluminum, when present in the reactants, generally acts to stabilize secondary mineral formation and buffers silica activity in olivine-rich serpentinization process. It serves as a viable means of accelerating serpentinization reactions toward an economically feasible timescale for industrial hydrogen production, including geologic means [35,55]. In comparison with other technologies (like steam reforming) for H₂ production, serpentinization also introduces energy minimization through its exothermic nature, and the presence of aluminum can further ensure its occurrence even at lower temperatures. X-ray photoelectron spectroscopy, an effective method for differentiating between Al-hydroxide and Al-oxide phases, utilizes the ratio of oxidic-bound Al to O to analyze Al oxidation and hydration behavior [56]. The investigation of Al-oxide and Al-hydroxide transformations revealed more pronounced alterations in basalts compared to ultramafic rocks. The O/Al atomic ratios of aluminum oxide layers were determined based on O 1s and Al 2p photoelectron intensities, showing an increase from 5.87 to 6.07 in the unreacted and reacted ultramafic Sample E, respectively, and from 3.74 to 4.62 in the unreacted and reacted basalt Sample E, respectively.

In an experimental study, [55] demonstrated a correlation between the onset of serpentinization and the dissolution of aluminum oxide (Al₂O₃) microspheres introduced into the reacting fluid. Importantly, the very first signs of serpentine minerals appeared simultaneously on both the olivine grains and the surface of these dissolving Al₂O₃ microspheres. This suggested that the presence of dissolved aluminum in the fluid significantly increased the solubility of olivine. Additionally, dissolved Al played a significant role in the formation (nucleation) and subsequent growth of serpentine minerals, particularly those rich in aluminum.

Pens et al. [35] also illustrated the contrasting effect of aluminum in both olivine-rich and orthopyroxene-rich serpentinization processes. Their results indicated that in the presence of aluminum (Al), olivine rapidly converts to lizardite with a half-life of only 7 h. In contrast, orthopyroxene shows minimal reaction (only 11% conversion) even after 6 days under the same conditions. Interestingly, orthopyroxene reacts faster without the presence of aluminum (48% conversion in 6 days). Olivine experiments resulted in the formation of lizardite and magnetite, while Orthopyroxene experiments exclusively produced proto-serpentine. The explanation for these contrasting effects lies in how aluminum interacts differently with the surfaces of olivine and orthopyroxene. Aluminum in solutions primarily exists as negatively charged complexes, such as Al(OH)₄-. The positively charged surface of olivine attracts and adsorbs these negatively charged Al complexes. This adsorption process likely enhances the rate at which olivine dissolves and also helps in the nucleation and growth of lizardite, which is enriched in aluminum. On the other hand, the neutral surface of orthopyroxene does not significantly interact with the negatively charged Al complexes. This lack of interaction explains the slower reaction rate of orthopyroxene in the presence of aluminum.

5. Conclusions

This study examined the potential for hydrothermal H₂ generation in seven ultramafic and mafic rocks with diverse mineralogical compositions. The study aimed to identify key reaction pathways and geochemical factors influencing H₂ generation efficiency. The results showed how differences in rock composition, Fe oxidation states, and secondary mineral formation influence hydrogen yield and reaction kinetics.

Our findings demonstrate that not all geologic H₂-generating reactions involving ultramafic and mafic rocks result in the formation of serpentine, brucite, or magnetite. Instead, the results suggest that H₂ generation in these systems is a strong function of mineral transformation. Surface-mediated electron transfer and redox cycling play an important role in sustaining hydrogen production beyond the initial mineral phase changes.

The diversity in reaction pathways, especially the formation of birnessite and/or clinochlore instead of serpentine in some instances, suggests that not all Fe-driven hydrogen production processes should be classified strictly as serpentinization. These findings also suggest that categorizing all H₂-producing hydrothermal reactions as “serpentinization” can be misleading, as alternative mineral assemblages such as birnessite, feldspars (andesine and anorthite), and other phyllosilicates (clinochlore) may be more representative of the dominant reaction mechanisms in certain rock types.

Magnetite experienced minimal transformation and, in some cases, was actually formed as a secondary mineral from precursor olivine during hydrothermal exposure. This suggests that the rock mineral may act both as a product and a reactant, indicating that its transformation was not the primary mechanism for H₂ generation. Further investigations into the kinetics of Fe oxidation and the role of trace mineral interactions are needed to fully understand how different geological settings influence natural H₂ production potential in both ultramafic and mafic rock systems.

Major Fe-rich ultramafic minerals, such as olivine and pyroxenes, seem to support sustained H₂ production through continuous solid–solution exchange and phase transformation, while aluminosilicates and transition metal sites (e.g., Ni, Mo, Mn, Cu) likely enhance H₂ yields. Additionally, fluid–rock interactions regenerate reactive surfaces, supporting prolonged H₂ generation even in the absence of serpentine formation. These results suggest that reaction kinetics and mineralogical composition, rather than Fe oxidation states alone, dictate the rates and efficiency of H₂ production.

Ultimately, these insights shift the focus from serpentine-driven mechanisms to a broader mineralogical framework, where transition metal sites and mineral surface properties may play a central role. Understanding these processes can enhance subsurface hydrogen exploration and facilitate the development of engineered H₂ production systems that utilize minerals known to occur within ultramafic and mafic rock formations.

6. Future Works

Future work should investigate longer reaction durations to assess the sustained potential of H₂ generation and mineral transformation over extended time scales. Reactions carried out beyond four-week periods could reveal slower geochemical processes and the stability of hydrogen-producing pathways. Additionally, experiments conducted at a broader range of temperatures and pressures will help determine the thermal thresholds that optimize specific reaction mechanisms such as serpentinization, brucite precipitation, and iron oxidation.

Further investigation is also needed to evaluate the influence of varying fluid chemistries, particularly salinity and pH, on H₂ production efficiency and mineral alteration. These parameters are important in replicating realistic subsurface environments and understanding their role in modulating reaction rates. In addition to serpentinization, this work suggests that surface redox reactions and geocatalytic processes may play a significant role in H₂ production. However, the catalytic contributions of trace minerals and naturally occurring transition metals presence in ultramafic and mafic rocks remain uncertain and warrant further investigation in future studies. Future investigations should focus on explicitly resolving the mechanistic role of trace elements in H₂ production. This could include experiments with isotopically labeled water to track redox processes, selective mineral doping to isolate catalytic effects, and advanced surface-sensitive spectroscopic analyses (e.g., XPS, XAS) to monitor the oxidation states of Ni, Cu, and Mo during reaction. Such approaches will help decouple trace element effects from broader mineralogical pathways and establish whether these elements act as true catalysts or simply correlate with other controlling factors.

Complementing experimental efforts with geochemical modeling will enable the prediction of reaction mechanisms, fluid–mineral equilibria, and long-term system evolution under various conditions, ultimately aiding in the design and optimization of field-scale hydrogen production strategies.