# Mass transfer evaluation of the hydration of an olivine-bearing monzo-diorite from the Mid-Continent Rift in Kansas (USA)

Valentine Combaudon, Stephen Centrella, Kanchana Kularatne, Olivier Sissmann, Eric Deville, et al..

Sissmann, Eric Deville, Hannelore Derluyn

► To cite this version:

Valentine Combaudon, Stephen Centrella, Kanchana Kularatne, Olivier Sissmann, Eric Dev-ille, et al.. Mass transfer evaluation of the hydration of an olivine-bearing monzo-diorite from the Mid-Continent Rift in Kansas (USA). Global and Planetary Change, 2026, 256, pp.105158. 10.1016/j.gloplacha.2025.105158. hal-05503641

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Submitted on 10 Feb 2026

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# Mass transfer evaluation of the hydration of an olivine-bearing monzo-diorite from the Mid-Continent Rift in Kansas (USA)

Valentine Combaudon<sup>a,b,*</sup>, Stephen Centrella<sup>c</sup>, Kanchana Kularatne<sup>b</sup>, Olivier Sissmann<sup>a</sup>, Eric Deville<sup>a</sup>, Hannelore Derluyn<sup>b,d</sup>

## ARTICLE INFO

Editor: Prof. Liviu Matenco

**Keywords:**

Mass transfer

Hydration reaction

Mineral replacement

Serpentinization

Amphibole formation

## ABSTRACT

We studied a rock sample belonging to the Mid-Continent Rift in Kansas (USA) that is partially hydrated with the formation of serpentine and amphibole. At the mineral scale, the original mineral assemblage of olivine and orthopyroxene (+ plagioclase + alkali feldspar) is partially replaced by serpentine, magnetite, biotite, amphibole, with locally some apatite and quartz. The replacement tends to preserve the shape of the original olivine and orthopyroxene, supporting a pseudomorphic replacement. Scanning Electron Microscopy (SEM), Electron Probe Micro-Analyzer (EPMA) mineral compositional map, and mass balance equations were used to quantify the transfers of mass during fluid infiltration. Here, the hydration reactions involve the decrease of rock density and, under the assumption of volume preservation, require a significant loss of mass through the fluid phase. Indeed, this study demonstrates that the formation of the newly formed hydrated minerals results from similar mass transfers between parent mineral, fluid, and product minerals. This study also demonstrates that the redox conditions are similar during the formation of all of these hydrated minerals. These observations confirm the close relationship between local stress generated by hydration reactions, mass transfers, and the scale over which the system is open.

## 1. Introduction

During hydration reactions, the disequilibrium conditions reigning between the rock and the infiltrating aqueous phase have a strong influence on the thermal and chemical evolution of the system (Manning, 1994; Putnis and Austrheim, 2012; Centrella et al., 2015; Bebout and Penniston-Dorland, 2016; Centrella et al., 2016; Manning, 2018). Hydration reactions induce changes in the mineralogy and thus the density of the rock. To accommodate the important change in density between primary anhydrous phases and secondary hydrous phases, pore and fracture formation are both associated with the process (e.g. Plümper et al., 2012). They enhance the transport of fluid (Jamtveit et al., 2008) and accelerate the reaction rate by increasing the reactive surface area (Røyne et al., 2008; Rudge et al., 2010; Malvoisin et al., 2012). However, Carmichael (1987) proposed that during the alteration of periclase to brucite, associated with density decrease, the overall volume of the solid was maintained (called pseudomorphic replacement leading to

pseudomorphosis) and the stress due to the density decrease was dissipated by the mass transfer, thus reducing the effects of reaction-induced stress variations. More recently, the same reaction was performed in a closed system (inducing mass preservation) and supported a positive solid volume variation and generation of induced differential stress around 1.5 GPa (Plümper et al., 2012). The question of how a system can accommodate mineralogical and density changes as well as stress generated by hydration reaction remains an open question (e.g. Macdonald and Fyfe, 1985; Klein and Le Roux, 2020).

The fluid phase becomes an important parameter to consider during fluid-rock interaction. One way to understand how a system can accommodate mineralogical and density changes is by quantifying the mass transfer associated with a fluid-mediated rock/mineral transformation. Unlike mass redistribution, which implies mass preservation during a reaction, mass transfer allows the system to lose or gain mass. By quantifying the mass transfer, it becomes possible to constrain the behavior of the elements within a fluid phase, its redox conditions

<footnote><sup>a</sup> IFPEN, Rueil-Malmaison, France</footnote>

<footnote><sup>b</sup> Université de Pau et des Pays de l'Adour, E2S UPPA, CNRS, LFCR, Pau, France</footnote>

<footnote><sup>c</sup> Centre for Exploration Targeting, School of Earth and Oceans, University of Western Australia, Crawley, Australia</footnote>

<footnote><sup>d</sup> Université de Pau et des Pays de l'Adour, E2S UPPA, CNRS, DMEX, Pau, France</footnote>

<footnote>* Corresponding author at: Department of Geological Sciences, University of Colorado, Boulder, CO, United States.
E-mail address: combaudon.valentine@gmail.com (V. Combaudon).</footnote>
(Wogelius et al., 1992; Huang et al., 2011; Merkulova et al., 2017; Centrella et al., 2020; Centrella et al., 2023a; Centrella et al., 2023b; Centrella, 2024), the direction of the fluid flow (Baumgartner and Olsen, 1995; Centrella, 2019; Centrella et al., 2023a) and the associated rheology evolution (Jamtveit et al., 2000; Plümper et al., 2012).

An intensively studied example of a hydration reaction is the reaction of serpentinization. It is a fluid-mediated reaction coupling dissolution of olivine or pyroxene and precipitation of serpentine through mineral interfaces (Moody, 1976; Putnis, 2009, Evans et al., 2013, Escario et al., 2018 amongst many others). Some authors have shown that serpentinization is an isochemical process with the increase of volume being accommodated only by fracturation and vein formation (Coleman and Keith, 1971; O'Hanley, 1992; Evans, 2004; Plümper et al., 2012; Rouméjon and Cannat, 2014; Klein and Le Roux, 2020). Authors commonly assume serpentinizing systems as closed systems when modeling or experimentally reproducing such reactions. This is justified by the proposed idea of a lack of strong diffusion gradients able to remove elements during serpentinization (Malvoisin et al., 2012, Okland et al., 2014, McCollom et al., 2016, McCollom et al., 2020a, McCollom et al., 2020b, Klein and Le Roux, 2020 amongst others). Some authors have also proposed that the overall water-rock (WR) ratios are not high enough to destroy the closed-system behavior and change the abundance of major elements (Evans, 2008). However, considering natural serpentinizing systems as open systems, involving slow but pervasive diffusive fluid flow through the rocks, cannot be completely ruled out, especially since the effects of fluid circulation on mafic and ultramafic rocks remain poorly understood. For instance, some experimental studies have been conducted on peridotites under open-system conditions (Normand et al., 2002; Godard et al., 2013; Escario et al., 2018). They concluded that element concentrations were decreasing through the reaction to precipitate secondary phases and then reached steady-state values while being mostly dependent on flow rate. Nevertheless, most of the serpentinization models and experiments are based on thermodynamic equilibrium conditions, and reactive transport is not considered (Godard et al., 2013). A study from Tutolo et al. (2018) conducted closed and open-system experiments and showed that serpentinization is a reactive transport process and thus must be considered as an open system. This is further supported by the study of Evans et al. (2017) conducted on peridotite samples from the Duluth Complex, part of the North American Mid-Continent Rift (MCR, Minnesota, USA). Authors highlight a significant Fe-enrichment of secondary hydrous and anhydrous products relative to the protolith, and a correlated Mg-depletion of the system due to leaching during a low-temperature serpentinization event. This petrographic study is of prime importance and shows that serpentinization of continental mafic and ultramafic rocks should not be regarded as an isochemical process. Open-system consequences on the natural serpentinizing continental systems need to be better characterized. In addition, since several authors are proposing that the volume increase during overall hydration reactions is mainly accommodated by significant mass transfer (Carmichael, 1987; Manning, 1994; Putnis and Austrheim, 2012; Centrella et al., 2015; Bebout and Penniston-Dorland, 2016; Centrella et al., 2016; Manning, 2018), one could wonder if it is also the case for the serpentinization reaction.

An olivine-bearing monzo-diorite has been studied recently in Combaudon et al. (2024). This rock belongs to the south extension of the Mid-Continent Rift (MCR) and presents an interesting iron-rich character linked to anhydrous and hydrous minerals. The identified hydration reaction of the monzo-diorite seems to proceed along grain boundaries (Fig. 5. in Combaudon et al., 2024), preserving the grain boundaries (shape) and the volume of the parent mineral associated with an important mass transfer between the initial mineral assemblage, the fluid phase, and the newly formed minerals. The present work aims to examine the genesis relationships between the initial minerals (olivine, orthopyroxene, and feldspars) and the product minerals (serpentine, biotite, amphibole, magnetite, quartz, and apatite), and to quantify the mass transfers using the map approach of Centrella (2024).

Using compositional maps for major elements, this method allows for the quantification of the partial fluid chemical composition and the redox conditions associated with a specific reaction.

## 2. Geologic context

Our study focused on the Precambrian basement of Kansas (USA) and more specifically on the Mid-Continent Rift (MCR), an aborted Precambrian continental rift whose formation started more than 1.1 Ga ago and extends along more than 1000 km from Kansas to the Lake Superior province (Woelk and Hinze, 1991). Based on the work of Ojakangas et al. (2001), the rocks belonging to the MCR have been identified as a lower volcanic syn-rift sequence in which 90% of the rocks are known to be flood basalt and 10% to be rhyolite. A few magmatic complexes are present. On top of this syn-rift sequence is the post-rift subsidence phase characterized by a clastic sedimentary sequence (Ojakangas et al., 2001).

A recent study shows that, in Kansas, rocks belonging to the MCR can be found east of the well-mapped portion of the rift (Combaudon et al., 2024). The authors studied the basement rocks from the DR1-A well, which is located above the Nemaha topographic high (Fig. 1). This well is one of the very few ones in Kansas that reach the basement and presents a unique opportunity to study the rock of the MCR in Kansas. The authors described, at a depth of 450 m, an olivine-bearing monzo-diorite made of the olivine endmember fayalite, ferrosilite, augite, Fe-rich pargasite in a matrix made of albite, and orthoclase. A complex association of Fe³⁺ and Fe²⁺ rich phyllosilicates was also described as associated with the olivines. A first generation made of a nanometric scale mixture of serpentine, chlorite, and mica, rich in reduced iron, was described as propagating through the fayalite but also the other mafic minerals and, to a lesser extent, the feldspathic matrix. The redox state of these secondary phases was assessed by Scanning Transmission X-ray Mapping coupled with X-ray Absorption Near-Edge Spectroscopy (STXM-XANES). Quartz, magnetite, and apatite are associated with these reduced hydrous species. A second, more oxidized generation of phyllosilicates made of Fe³⁺-rich interstratified chlorite-smectite and Fe³⁺-rich serpentine was described at the interface of the fayalite only. Contacts between the two generations of phyllosilicates and the fayalite are sharp. Textures and micro textures of these complex phyllosilicates suggest they have formed during a late magmatic event involving deuteric alteration (oxidation reaction) of olivine and could have produced H₂. These phyllosilicates are present as micrometric patches in the fayalite. As described by the same authors, the preservation of the parent olivine and orthopyroxene shape by the product minerals supports the hypothesis of a pseudomorphic replacement.

## 3. Materials and methods

### 3.1. Mineralogical characterization

We used a fayalite-bearing monzo-diorite from a core sample (named K2) obtained from the basement rocks from the DR1-A well, described in Section 2. Mineralogical analysis was performed on a cylindrical subsample (d = 8.75 mm and h = 4 mm), which was previously reported by the authors (Kularatne et al., 2024). One of the circular surfaces of the sample was polished using a diamond polishing paste of decreasing grit sizes: 1 µm, 0.25 µm, and 0.03 µm. The polished surface of the cylinder was first studied by Scanning Electron Microscopy (SEM) using a FEG JEOL JSM 7100F SEM instrument operated at an accelerating voltage of 15 keV and a beam current of 48 nA, at the Centre Castaing (Toulouse, France). SEM-Backscattered electron (BSE) analysis and SEM-Energy dispersive X-ray spectroscopy (EDS) element distribution maps allowed choosing an area of interest (≈ 4 mm²) with a representative mineral assemblage containing olivine, orthopyroxene, feldspar, serpentine, amphibole, biotite, quartz, and magnetite present in the monzo-diorite sample for further study using Electron Probe Micro-
<figure><img src="image_3.png" /><figcaption>Fig. 1. Simplified west-East cross-section of Kansas and geological map of the basement top with the location of the DR1-A well (modified from Combaudon et al., 2024, please see references therein).</figcaption></figure>

Analyzer (EPMA). Data acquisition was performed using a CAMECA SXFive instrument at the Centre Castaing (Toulouse, France), operated at 15 keV and 48 nA. Ten elements were analyzed for standardization: Na, Mg, Al, Si, K, Ca, Ti, Fe, Mn, and Ba using a beam size of 2 nm and 0.125 dwell time. Compositional maps were performed using the XMapTools 4.1 package (Lanari et al., 2014, 2019).

Phase segmentation between olivine, orthopyroxene, feldspar, serpentine, amphibole, biotite, quartz, and magnetite was performed with selections of region-of-interest and using the random forest machine learning algorithm in XMapTools 4.1 package (Lanari et al., 2014, 2019).

As an electron probe microanalyzer does not allow for differentiating the ferric and ferrous iron contents, the estimation of the FeO and Fe₂O₃ contents of amphibole was performed based on Leake et al. (1997). The Fe²⁺/Fe³⁺ ratio of serpentine has been acquired by STXM-XANES (Combaudon et al., 2024). Knowing the Fe²⁺/Fe³⁺ ratio and the actual FeO₃ in serpentinite, specific FeO and Fe₂O₃ content was recalculated.

## 3.2. Method of mass transfer calculation

The mass balance equation used by Gresens (1967) allows for to estimation of element mobility, especially to quantify gains and losses of elements of interest, during a fluid-mediated rock transformation with a homogeneous fluid phase. Eq. 1 is used for the case of two minerals and n components, but it can also be used for a more complex assemblage of minerals:

$$f_v \left( \frac{d_{\text{mineral 2}}}{d_{\text{mineral 1}}} \right) c_{\text{mineral 2}} - c_{\text{mineral 1}} = x_n \quad (1)$$

fv is the volume factor, which can be related to the solid volume variation (fv = 1 means no volume variation, fv < 1 means the volume is decreasing, fv > 1 means the volume is increasing), dmineral is the density

of the mineral considered, cmineral is the weight fraction of component n in the mineral of interest, and xn is the gain or the loss for each component in gram per 100 g of the initial mineral. Positive values correspond to a gain of mass, while negative values correspond to a loss of mass through the fluid phase.

To convert the mass change in percent of mass change, xn must be divided by the element weight percent of the same element. For each pixel that composes a mineral, mineral densities were calculated based on a linear combination of end members assuming an ideal solid solution between each end member of a mineral as described in Eq. 2 (Centrella et al., 2018).

$$d = \sum [(d_1 \times w_1) + \dots + (d_n \times w_n)] \quad (2)$$

d_n is the density of n mineral, and w_n is the weight fraction for n mineral.

Olivine density has been estimated for each pixel based on a linear combination between fayalite [Fe₂SiO₄], forsterite [Mg₂SiO₄], and tephroite [Mn₂SiO₄] with respective densities of 4.39, 3.27, and 4.25 g. cm⁻³ (Birle et al., 1968). Then, based on the pixels of the compositional map, the average olivine density is estimated at around 4.32 g.cm⁻³ (Fig. 2). Orthopyroxene density has been estimated based on a linear combination between enstatite ([Mg₂Si₂O₆] - Carlson et al., 1988) and ferrosilite ([Fe₂Si₂O₆] - Burnham et al., 1971), with respective densities of 3.20 and 3.95 g.cm⁻³. Based on the acquired compositional maps, orthopyroxene density is estimated to be around 3.81 g.cm⁻³ (Fig. 2). Amphibole solid solutions are more complex, and it can be difficult to estimate their correct density. Based on compositional maps and the nomenclature of Leake (1978), amphibole density is estimated at around 3.58 g.cm⁻³ (Fig. 2). Taking into account the Fe-rich nature of the serpentine, the density of the serpentine has been estimated based on a linear combination between lizardite ([Mg₃Si₂O₅(OH)₄], 2.57 g.cm⁻³ - Mellini, 1982), greenalite ([Fe³⁺Fe²⁺)₂₃Si₂O₅(OH)₄], 3 g.cm⁻³ - Shirouz and Bailey, 1965), and cronstedtite ([Fe₂⁺Fe³⁺)₂Si₂O₅(OH)₄], 3.59 g.
<figure><img src="image_4.png" /><figcaption>**Fig. 2.** a- BSE-SEM image of the area studied; b- phases and; c- density maps of the studied area obtained from mineral compositional maps. Black pixels are mixed pixels, i.e., their chemical composition can't be attributed to only one mineral constitutive of the system. This could be due to the resolution of the analysis being lower than the size of the feature studied. The Density map is obtained by chemical composition and linear relationship between each endmember of olivine, orthopyroxene, amphibole, biotite, serpentine, plagioclase, alkali feldspar, quartz, apatite and magnetite (see Table 1 for more details on the endmembers used). Plots are obtained from XMapTools 4.1 (Lanari et al., 2014, 2019).</figcaption></figure>

cm⁻³ – Hendricks, 1939). The Fe³⁺/Fe²⁺ ratio of the serpentine was estimated by Combaudon et al., 2024. Based on compositional maps, serpentine density is estimated at around 2.70 g.cm⁻³ (Fig. 2). K-feldspar and plagioclase densities have been estimated based on a linear combination between albite [NaAlSi₃O₈], anorthite [CaAl₂Si₃O₈], and orthoclase [KAlSi₃O₈] with respective densities of 2.62 (Prewitt et al., 1976), 2.73 (Foit and Peacor, 1973), and 2.52 g.cm⁻³ (Colville and Ribbe, 1968). Based on compositional maps, their values are estimated at around 2.53 g.cm⁻³ and 2.64 g.cm⁻³, respectively (Fig. 2). Biotite density has been estimated based on a linear combination between annite [KFe₃²⁺AlSi₃O₁₀(OH, F)₂], phlogopite [KMg₃(AlSi₃O₁₀(OH, F)₂], and siderophyllite [KFe₂²⁺Al(Si₂Al₂)O₁₀(OH, F)₂] with respective

densities of 3.34 (Hazen and Burnham, 1973), 2.83 (Hendricks and Jefferson, 1939), and 3.00 g.cm⁻³ (Brigatti et al., 2000). Based on compositional maps, biotite density is estimated at around 3.2 g.cm⁻³ (Fig. 2). For magnetite, a pure endmember is selected (Wechsler et al., 1984), and the density is estimated to be 5.15 g.cm⁻³ (Fig. 2). Density-corrected local bulk composition and associated density were extracted from the compositional map in XMapTools 4.1 (Lanari et al., 2014, 2019).

To quantify the mass transfer associated with the hydration reaction, the approach of Centrella (2024) was used. The original local bulk chemical composition and density used as referential material correspond to the average chemical composition of the assemblage olivine (Ol) + orthopyroxene (Opx) and olivine + orthopyroxene + plagioclase (Pl) + alkali-feldspar (Afs). Average values are presented in Table 1.

As described by Combaudon et al., 2024, the specific replacement of the parent minerals, especially olivine and orthopyroxene, suggests a pseudomorphic replacement. The implication of such observation implies a volume factor equal to 1, meaning no solid volume variation. Mass balance calculations were therefore performed using a volume factor of 1.

The mass balance calculation is performed on each pixel of the product mineral assemblage: serpentine + amphibole + biotite + magnetite + quartz + apatite. Results represent the percent of mass change of a specific element for each pixel of the product reaction. In the case of a pseudomorphic replacement reaction, the solid volume variation is not represented since it is equal to 0 %. The map of the sum of mass transfer allows for evaluating whether the mineral assemblage behaves as an open or as a closed system for the specific reaction at hand, by examining the sum of all elements gained and lost during the reaction. When the sum is negative, the reaction tends to lose mass; when the sum is positive, mass is gained; in both cases, we are dealing with an open system. Conversely, when the sum of mass transfer is equal to zero, the system is closed, with mass being preserved.

# 4. Results

## 4.1. Sample description and textural relationships

The recent petrographic study of the olivine-bearing sample focused on a micro to nanometric scale of observation and performed a very high-resolution characterization of the mineralogical assemblage (Combaudon et al., 2024). Fe-rich clay minerals are present mainly through the olivine but are also observed to a lesser extent through the orthopyroxene and to an even lesser extent through the feldspars (Fig. 3).

Based on Combaudon and co-authors (2024), the minerals within the olivines are described as a complex nanometric scale mixture of serpentine, smectite, mica, and chlorite. They can be divided into two generations, and they are believed to be formed during a late magmatic event leading to the deuteric alteration of the olivine. When looking at the studied sample from a much broader perspective, it appears that some additional textural relationships can be identified between minerals. Indeed, the Fe-rich amphibole is often present as a rim around the olivine and the orthopyroxene (Fig. 4a and b), indicating that a metamorphic reaction has happened and led to additional mineral replacements. In addition, the biotite is also closely associated with the Fe-rich amphibole and the orthopyroxene. The serpentine is directly identified as propagating into the olivine, but also the orthopyroxene, with sharp contacts (Fig. 4c and d). Magnetite, even if present in minor amounts, is also identified at the interface between the serpentine, the olivine, and the orthopyroxene (at this scale of observation). The mineral interfaces are generally sharp and irregular, particularly between olivine and serpentine (Fig. 3a), but also amongst serpentine-biotite-amphibole and orthopyroxene (Figs. 2 and 3). These textural relationships suggest that olivine and orthopyroxene are replaced by amphibole, biotite, serpentine, and magnetite. The partial
<caption>**Table 1**</caption>

Average chemical composition of olivine, orthopyroxene, amphibole, serpentine, biotite, alkali feldspar, plagioclase, and magnetite with their respective densities. FeO and Fe₂O₃ for amphibole have been determined following Leake et al. (1997). FeO and Fe₂O₃ for serpentine are from Combaudon et al., 2024. XFo: forsterite, XFa: fayalite, XTep: tephroide, XEn: enstatite, XFs: ferroenstatite, XLiz: lizardite, XGre: greenalite, XCron: cronstedtite, XAn: annite, XPhl: phlogopite, XSid: siderophilite, XAb: albite, Xan: anorthite, XOrth: orthoclase from Whitney and Evans, 2009. For hydrous phases, the H₂O content was adjusted to bring the totals to 100 %.

<table><thead><tr><th>(ox.wt%)</th><th>Olivine</th><th>Orthopyroxene</th><th>Amphibole</th><th>Serpentine</th><th>Biotite</th><th>K-feldspar</th><th>Plagioclase</th><th>Magnetite</th><th>Ol + Opx + Pl + Afs</th><th>Ol + Opx</th></tr></thead><tbody><tr><td>SiO<sub>2</sub></td><td>29.72</td><td>50.01</td><td>41.52</td><td>38.88</td><td>35.62</td><td>64.80</td><td>62.25</td><td>0.00</td><td>51.98</td><td>36.56</td></tr><tr><td>TiO<sub>2</sub></td><td>0.03</td><td>0.02</td><td>1.00</td><td>0.02</td><td>4.27</td><td>0.03</td><td>0.02</td><td>0.55</td><td>0.03</td><td>0.05</td></tr><tr><td>Al<sub>2</sub>O<sub>3</sub></td><td>0.01</td><td>0.17</td><td>9.56</td><td>1.49</td><td>12.70</td><td>18.62</td><td>22.89</td><td>1.06</td><td>12.10</td><td>0.12</td></tr><tr><td>FeO</td><td>65.04</td><td>42.88</td><td>28.20</td><td>46.91</td><td>28.15</td><td>0.12</td><td>0.29</td><td>100.00</td><td>25.53</td><td>57.21</td></tr><tr><td>MnO</td><td>2.66</td><td>2.15</td><td>0.48</td><td>1.35</td><td>0.26</td><td>0.03</td><td>0.03</td><td>0.16</td><td>1.10</td><td>2.4</td></tr><tr><td>MgO</td><td>2.12</td><td>5.70</td><td>4.65</td><td>1.27</td><td>6.42</td><td>0.02</td><td>0.01</td><td>0.08</td><td>1.48</td><td>3.34</td></tr><tr><td>CaO</td><td>0.04</td><td>0.74</td><td>10.40</td><td>0.27</td><td>0.04</td><td>0.21</td><td>4.82</td><td>0.08</td><td>2.02</td><td>0.28</td></tr><tr><td>Na<sub>2</sub>O</td><td>0.03</td><td>0.02</td><td>1.56</td><td>0.63</td><td>0.05</td><td>2.25</td><td>9.24</td><td>0.10</td><td>3.97</td><td>0.04</td></tr><tr><td>K<sub>2</sub>O</td><td>0.03</td><td>0.03</td><td>1.22</td><td>0.05</td><td>8.79</td><td>13.22</td><td>0.25</td><td>0.07</td><td>1.92</td><td>0.05</td></tr><tr><td>BaO</td><td>0.04</td><td>0.05</td><td>0.03</td><td>0.05</td><td>0.37</td><td>1.00</td><td>0.03</td><td>0.16</td><td>0.18</td><td>0.04</td></tr><tr><td>H<sub>2</sub>O</td><td></td><td></td><td>1.38</td><td>9.08</td><td>3.33</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Total</td><td>99.72</td><td>101.75</td><td>100.00</td><td>100.00</td><td>100.00</td><td>100.30</td><td>99.83</td><td>102.26</td><td>100.28</td><td>100.09</td></tr><tr><td>FeO</td><td>65.04</td><td>42.88</td><td>20.67</td><td>18.23</td><td>28.15</td><td>0.12</td><td>0.29</td><td>31.00</td><td></td><td></td></tr><tr><td>Fe<sub>2</sub>O<sub>3</sub></td><td>0.00</td><td>0.00</td><td>7.53</td><td>28.68</td><td>0.00</td><td>0.00</td><td>0.00</td><td>0.00</td><td>69.00</td><td></td></tr><tr><td>FeO/(Fe<sub>2</sub>O<sub>3</sub> + FeO)</td><td>1.00</td><td>1.00</td><td>0.73</td><td>0.39</td><td>1.00</td><td>1.00</td><td>1.00</td><td>0.31</td><td></td><td></td></tr><tr><td>Normalized to:</td><td>4Ox</td><td>3Ox</td><td>23Ox</td><td>14Ox</td><td>11Ox</td><td>8Ox</td><td>8Ox</td><td></td><td></td><td></td></tr><tr><td>Si</td><td>1.00</td><td>1.00</td><td>6.59</td><td>4.51</td><td>2.81</td><td>2.98</td><td>2.77</td><td></td><td></td><td></td></tr><tr><td>Al</td><td>0.00</td><td>0.01</td><td>1.79</td><td>0.20</td><td>1.18</td><td>1.01</td><td>1.20</td><td></td><td></td><td></td></tr><tr><td>Ti</td><td>0.00</td><td>0.00</td><td>0.12</td><td>0.00</td><td>0.25</td><td>0.00</td><td>0.00</td><td></td><td></td><td></td></tr><tr><td>Fe</td><td>1.81</td><td>0.77</td><td>3.60</td><td>4.20</td><td>1.86</td><td>0.00</td><td>0.01</td><td></td><td></td><td></td></tr><tr><td>Mn</td><td>0.07</td><td>0.03</td><td>0.06</td><td>0.13</td><td>0.02</td><td>0.00</td><td>0.00</td><td></td><td></td><td></td></tr><tr><td>Mg</td><td>0.10</td><td>0.17</td><td>1.10</td><td>0.22</td><td>0.76</td><td>0.00</td><td>0.00</td><td></td><td></td><td></td></tr><tr><td>Ca</td><td>0.00</td><td>0.01</td><td>1.77</td><td>0.03</td><td>0.00</td><td>0.01</td><td>0.23</td><td></td><td></td><td></td></tr><tr><td>Na</td><td>0.00</td><td>0.00</td><td>0.48</td><td>0.14</td><td>0.01</td><td>0.20</td><td>0.80</td><td></td><td></td><td></td></tr><tr><td>K</td><td>0.00</td><td>0.00</td><td>0.251</td><td>0.01</td><td>0.89</td><td>0.78</td><td>0.01</td><td></td><td></td><td></td></tr><tr><td>Ba</td><td>0.00</td><td>0.00</td><td>0.00</td><td>0.00</td><td>0.01</td><td>0.02</td><td>0.00</td><td></td><td></td><td></td></tr><tr><td>Sum</td><td>2.98</td><td>1.99</td><td>15.51</td><td>9.44</td><td>7.79</td><td>5.00</td><td>5.02</td><td></td><td></td><td></td></tr><tr><td>End-members:</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>XFo</td><td></td><td>0.05</td><td>XEn</td><td>0.15</td><td>XLiz</td><td>0.03</td><td>XAn</td><td>0.38</td><td>XAb</td><td>0.74</td></tr><tr><td>XFa</td><td></td><td>0.91</td><td>XFs</td><td>0.85</td><td>XGre</td><td>0.44</td><td>XPhl</td><td>0.15</td><td>XAn</td><td>0.23</td></tr><tr><td>XTep</td><td></td><td>0.04</td><td></td><td></td><td>XCron</td><td>0.53</td><td>XSid</td><td>0.61</td><td>XOrth</td><td>0.03</td></tr><tr><td>Calculated density (g. cm-3)</td><td>4.33</td><td>3.84</td><td>3.58</td><td>3.00</td><td>3.50</td><td>2.74</td><td>2.64</td><td>5.15</td><td>3.14</td><td>4.32</td></tr></tbody></table>

<figure><img src="image_5.png" /><figcaption>**Fig. 3.** Additional BSE-SEM images of the studied sample showing a- the textural relationship between the primary minerals olivine and orthopyroxene, and the secondary minerals serpentine, amphibole, and magnetite; b- the absence of serpentine propagation through the feldspar matrix. (ol: olivine, opx: orthopyroxene, srp: serpentine, mag: magnetite, amp: amphibole, kfs:kfeldspar, pl: plagioclase from Whitney and Evans, 2009).</figcaption></figure>
<figure><img src="image_6.png" /><figcaption>Fig. 4. a and b- Simplified sketch illustrating the pseudomorphic replacement of olivine and orthopyroxene by amphibole, serpentine, and, to a lesser extent, by biotite, magnetite, c- zoom on the olivine and serpentine grain boundary, d- zoom on the serpentine, orthopyroxene, and amphibole grain boundaries, showing that the hydration reaction proceeds by interface-coupled dissolution-precipitation mechanism (c) with a sharp contact between minerals (d).</figcaption></figure>

replacement of the parent olivine and orthopyroxene by these hydrous phases (together with magnetite) tends to preserve the overall morphology of the original minerals without any visible porosity or edge pit. The following sketch sum up the textural relationships observed in the samples of the hydrated monzo-diorite (Fig. 4).

Based on the previous authors' observations and the updated observations made in this study, two potential metamorphic reactions are considered:

1) olivine + orthopyroxene have already crystallized and are in disequilibrium with an aqueous fluid. This aqueous fluid could be a late residual magmatic fluid (Combaudon et al., 2024). The hydration reaction forms the Fe-rich serpentine, the Fe-rich amphibole and

biotite, as well as minor amounts of magnetite, quartz, and apatite. The feldspathic matrix is not considered in this scenario. It is assumed it is not reactive and is at equilibrium with the percolating aqueous fluid:

$$[\text{Ol} + \text{Opx} + \text{fluid} \rightarrow \text{Srp} + \text{Amp} + \text{Bt} + \text{Mag} + \text{Qtz} + \text{Ap}] \quad (3)$$

2) the second scenario considers the presence of olivine, orthopyroxene, plagioclase, and alkali feldspar as parent minerals. The feldspathic matrix is assumed to be reactive and not at equilibrium with the percolating aqueous fluid. The hydration reaction can then be written:

$$[\text{Ol} + \text{Opx} + \text{Pl} + \text{Afs} + \text{fluid} \rightarrow \text{Serp} + \text{Amp} + \text{Bt} + \text{Mag} + \text{Qtz} + \text{Ap}] \quad (4)$$

## 4.2. Minerals chemical composition

The average chemical composition of each phase is presented in Table 1. It is obtained from an average compositional map. The olivine has been identified as the iron-rich endmember fayalite [Fa$_{0.91}$]. The orthopyroxene, with 43.8 ox.wt% of FeO and only 5.3 ox.wt% of MgO, has been identified as ferrosilite [Fs$_{0.81}$]. The biotite, with 28.1 ox.wt% of FeO, 6.45 ox.wt% of MgO, and 4.52 ox.wt% of TiO$_2$ corresponds to a siderophyllite [Sid$_{0.55}$]. The plagioclase has been identified as albite [Ab$_{0.74}$] and K-feldspar has been identified as orthoclase [Or$_{0.87}$]. The iron oxide has been identified as magnetite Fe$_3$O$_4$. The amphibole is iron, calcium, and aluminum-rich with 28.05 ox.wt% of FeO$_{Total}$ and 10.45 ox.wt% of CaO and 1.67 ox.wt% of Na$_2$O. The estimation of ferric iron in amphibole was made following the calculation described by Leake et al. (1997). Thus, amphibole is identified as hastingsite and is composed of 23.51 ox.wt% and 4.54 ox.wt% of FeO and Fe$_2$O$_3$. The serpentine is iron-rich with 42.09 ox.wt% of FeO$_{Total}$ associated with 2.85 ox.wt% of Al$_2$O$_3$, 1.44 ox.wt% of MgO, 1.28 ox.wt% of MnO, and 1.09 ox.wt% of Na$_2$O. H$_2$O (or LOI content) reaches 10.1 ox.wt%. The ferric iron content from STXM-XANES data ($[\text{Fe}^{3+}] / \sum \text{Fe} = 0.27 \pm 0.02$) in Combaudon et al. (2024) was used to estimate ferric iron content in serpentine. FeO$_{Total}$ is about 42.39 ox.wt% and was converted into element weight percent by multiplying by the molar weight ($Fe_{Total} = 32.95 \text{ wt%}$). Based on STXM-XANES, this mass is composed of 27 % of Fe$^{3+}$ and 73 % of Fe$^{2+}$. The specific Fe$^{2+}$ and Fe$^{3+}$ were then converted into FeO and Fe$_2$O$_3$, respectively.

## 4.3. Mass transfer associated with the hydration reaction

The mass balance equation (Eq. 1) was applied to each pixel of the studied area, which allows us to determine the gain or loss for major elements during the mineralogical replacement occurring in the olivine-bearing monzo-diorite. The sum of mass transfer (or mass change), which represents the sum of all elements gained and lost by the system during the reaction, is presented in Fig. 5. When mass is preserved, its values are null. For more clarity, only Si, Al, Mg, Ca, Fe$^{2+}$, and Fe$^{3+}$ are presented in Figs. 6 and 7. Variations of other elements (Mn, Ti, Na, K, and Ba) are presented in Supplementary 1 and 2.

The simultaneous replacement of olivine and orthopyroxene by serpentine, magnetite, biotite, amphibole, quartz, and apatite (Reaction 1) induces a global loss of mass for the system of about 30 % (Fig. 5). The formation of quartz is responsible for a mass loss of around 60 %. It means some elements are not incorporated by these secondary products and leave the system through the circulating aqueous fluid. In contrast, the magnetite formation is associated with a net gain of mass of about 20 %. If the plagioclase and the alkali-feldspar are considered, the system tends to gain about 5 % of mass.

Regarding the mass change of Si (Figs. 6a and b), assuming (or not) the plagioclase and the alkali-feldspar as part of the original mineral assemblage, all product minerals are associated with a net loss of mass ($\approx -20 \%$). It is different for Al and Mg, where all product minerals gain
<figure><img src="image_7.png" /><figcaption>Fig. 5. a- Results of the sum of mass change (%) for reaction 1 (Ol + Opx = Amp + Bt + Srp + Mag + Qtz + Ap) and b- reaction 2 (Ol + Opx + Pl + Afs = Amp + Bt + Srp + Mag + Qtz + Ap). Positive values indicate that the system gains mass, while negative values indicate a loss.</figcaption></figure>

Al and lose Mg if olivine and orthopyroxene are the only parent minerals (Fig. 6c-f). If the plagioclase and the alkali-feldspar are considered in the original mineral assemblage, then almost all minerals lose Al, and only serpentine loses Mg. Apatite and Quartz appear to lose Mg in Fig. 5e and f, as there's no Mg in either mineral. Overall, similar observations can be made between Si and Ca (Fig. 6g and h). Fig. 7, being more focused on the redox-sensitive element Fe, represents the mass change of Fe²⁺ and Fe³⁺ for each mineral. Note that no Fe³⁺ was considered for biotite; all iron is considered as Fe²⁺. The assumption of considering or not the plagioclase and the alkali-feldspar as being part of the original mineral assemblage doesn't change the final elemental composition. All product minerals lose almost the same amount of Fe²⁺ (≈ 70 %) and gain the same amount of Fe³⁺ (≈ 2 × 106 %).

Based on average values from the mass transfer approach, the hydration reaction can be expressed as follows, considering only olivine and orthopyroxene as the original reactants:

$$ 
\begin{aligned}
100 \text{ g (olivine} - \text{orthopyroxene)} + 0.30 \text{ g Ti} + 1.54 \text{ g Al} + 3.87 \text{ g Fe}^{3+} \\
+ 3.27 \text{ g Ca} + 0.33 \text{ g Na} + 0.46 \text{ g K} = 71.54 \text{ g (serpentine} \\
- \text{amphibole} - \text{biotite} - \text{magnetite} - \text{quartz}) \\
+ 2.98 \text{ g Si} + 33.7 \text{ g Fe}^{2+} + 1.43 \text{ g Mn} + 0.28 \text{ g Mg}
\end{aligned}
\qquad (5) $$

If considering olivine, orthopyroxene, and the two feldspars, the reaction can be expressed as follows:

$$ 
\begin{aligned}
100 \text{ g (olivine} - \text{orthopyroxene} - \text{plagioclase} \\
- \text{AFeldspar)} + 0.44 \text{ g Ti} + 5.33 \text{ g Fe}^{3+} + 1.49 \text{ g Mg} + 3.33 \text{ g Ca} \\
= 98.6 \text{ g (serpentine} - \text{amphibole} - \text{biotite} - \text{magnetite} - \text{quartz}) \\
+ 4.85 \text{ g Si} + 1.03 \text{ g Al} + 4.98 \text{ g Fe}^{2+} + 0.25 \text{ g Mn} + 0.99 \text{ g Na} + 0.14 \text{ g K}
\end{aligned}
\qquad (6) $$

Eq. 5 shows that about 9.77 g of element mass reacts with olivine-orthopyroxene to produce 71.54 g of serpentine, amphibole, biotite, magnetite, and quartz. At the same time, the fluid leaving the system carries around 38.4 g of element mass. Eq. 6 shows that about 10.6 g of element mass reacts with olivine-orthopyroxene to produce 98.6 g of serpentine, amphibole, biotite, magnetite, and quartz. At the same time, the fluid leaving the system carries around 12.24 g of element mass.

# 5. Discussion

## 5.1. Evolution of the olivine-bearing monzo-diorite during the hydration reaction at serpentinization/amphibolite facies conditions

The formation of serpentine, amphibole, biotite, quartz, apatite, and magnetite through reaction one induces a density decrease (4.32 g.cm⁻³

to 3.24 g.cm⁻³). The formation of the same product minerals through the second reaction induces a density increase (3.14 g.cm⁻³ to 3.24 g.cm⁻³). For most elements, there is a clear correlation between gains and losses, where, in general, all elements behave the same way during the hydration reactions. Furthermore, there is also a similar element behavior for Fe²⁺ and Fe³⁺ for all Fe-bearing minerals, suggesting a similar oxidizing environment. In this study, while we do not have access to the post-reaction fluid, we looked at the chemical composition of the reaction products with respect to the chemical composition of the protolith and calculated the gain and loss of elements during a hydration reaction in an open system.

Reaction 1 induces a global loss of mass in the system post-reaction, except for the magnetite formation, which induces a local gain of mass. Significant transfer of elements happens during reaction 1. The serpentine does not incorporate Mg, but the biotite and amphibole do, which balances the amount of Mg present in the system. In the same way, the serpentine and biotite do not incorporate Ca, but the amphibole does. Finally, the system loses Al, Si, and Fe²⁺ but gains a significant amount of Fe³⁺. It seems that a local mass balance occurs for specific elements but not for all.

Reaction 2 induces overall a conservation of the mass post-reaction except for the magnetite formation, which, once again, induces a local gain of mass, and the quartz formation, which induces a local loss of mass. Significant redistribution of elements still happens during reaction 2. The serpentine does not incorporate Mg, but the biotite and amphibole do. The biotite and serpentine do not incorporate Ca, but amphibole does. A portion of Al is incorporated into the biotite, but all secondary minerals are responsible for Al loss. Finally, the system loses Si and Fe²⁺, but gains Fe³⁺. As for reaction 1, it seems that a local mass balance occurs only for some specific elements.

The mass transfer maps in Figs. 6 and 7 illustrate how the fluid phase redistributed elements during the reactions, both introducing and removing them for each mineral formation. Although the two reactions (Eq. 3 and Eq. 4) display minor differences in the behavior of major elements, they both show a consistent trend of Fe²⁺ loss and Fe³⁺ gain. Moreover, biotite, amphibole, and serpentine exhibit similar patterns by losing Si, Mg, Ca, and Fe²⁺ and gaining Fe³⁺, except for Al. Because the mass transfer behavior of these minerals is so similar, it is reasonable to infer that they formed contemporaneously under comparable conditions. This further suggests that the fluid's chemical composition and redox state remained stable throughout the reactions. The sharp reaction interfaces (replacement occurring along the grain boundaries) between parent and product phases and the lack of chemical zoning in the parent phases support a replacement mechanism by an interface-coupled dissolution-precipitation mechanism (Putnis, 2009; Putnis and Austrheim, 2012). The preservation of the shape (and thus the volume) requires mass transfer of elements between parents, product phases, and
<figure><img src="image_8.png" /><figcaption>**Fig. 6.** a, c, e- Results of the percent of the mass change of Si, Al, Mg, and Ca for reaction 1 (Ol + Opx = Amp + Bt + Srp + Mag + Qtz + Ap) and b, d, f, h- reaction 2 (Ol + Opx + Pl + Afs = Amp + Bt + Srp + Mag + Qtz + Ap). Positive values indicate that the system gains a specific mass, while negative values indicate a loss. Of note, the changes are relative to each other and not absolute values.</figcaption></figure>

the fluid to accommodate the decrease in density (Table 1).

Pseudomorphic replacement associated with fluid-rock interaction is not unusual. This process has been highlighted in various studies at different P-T conditions with for example, the preservation of a gabbro

texture at eclogite-facies conditions (John and Schenk, 2003, Fig. 3a; Putnis and John, 2010), or the granulite texture preserved after fluid infiltration at amphibolite-facies conditions (Centrella et al., 2015; Centrella et al., 2016). The primary factor supporting volume
<figure><img src="image_9.png" /><figcaption>**Fig. 7.** a and c- Results of the percent of mass change of Fe²⁺ and Fe³⁺ for reaction 1 (Ol + Opx = Amp + Bt + Srp + Mag + Qtz + Ap) and b and d- reaction 2 (Ol + Opx + Pl + Afs = Amp + Bt + Srp + Mag + Qtz + Ap). Positive values indicate that the system gains a specific mass, while negative values indicate a loss. Of note, the changes are relative to each other and not absolute values.</figcaption></figure>

preservation is the retention of the parent minerals' shape.

The specific replacement of olivine by serpentine, magnetite, and orthopyroxene by amphibole and biotite implies a net density loss (Fig. 2c). If the volume is preserved and the molar volume of each phase increases, then some material must be lost to the fluid phase. By considering the density-corrected bulk chemical composition of olivine + orthopyroxene, all product phases lose about 30 % of the mass (Fig. 5a). If the plagioclase and alkali-feldspar are also considered part of the original mineral assemblage, then the system behaves as a “closed system” where mass tends to be preserved (of about 0 % of mass gain/loss, Fig. 5b). Despite the overall mass loss, some elements are gained during the hydration (Figs. 6 and 7).

This naturally raises the question of reaction-induced fracturing during hydration reaction, especially in the case of serpentinization of olivine (Plümper et al., 2012). In this study and at the resolution of the SEM, no etch pit was visible in the original olivine, suggesting that no fractures were induced by the reaction. This implies that the stress generated by the hydration reaction can be balanced by removing material into the fluid, allowing the original solid volume to be preserved.

This fundamental problem of mass transfer (open system with no mass preservation) versus mass redistribution (closed system with mass preservation) is also discussed in other hydration reactions, such as the hydration of periclase to brucite (Carmichael, 1987; Plümper et al., 2022). The significant mass transfer implies that the system must be “open” on a larger scale than the reaction site. Fig. 5 presents an interesting result related to whether the system is open or closed. By considering only the olivine + orthopyroxene as the original material, all product minerals lose around 30 % of the mass (1296 kg/m³). If two feldspars are considered, then the total mass transfer of the product serpentine + amphibole + biotite is close to null, implying a relatively closed system at that scale of observation. These results highlight the importance of considering the representative scale of a reactive system

and its fundamental implications for our understanding of metasomatic processes. The system considered here is relatively limited, by its scale but also by the amount of data available. It is difficult to trace such fluid movements at a large scale and assess the regional impact. Nevertheless, hydration reactions of this magnitude can profoundly influence strain localization and broader tectonic movements.

## 5.2. Fe redox state of the system and implications for potential associated H₂ generation

Based on the mass transfer results (Figs. 5, 6, and 7), while being one of the least abundant secondary phases, the magnetite is the only product gaining mass (≈ 20–70 % depending on whether or not alkali feldspar and plagioclase are considered in the parent minerals). It is well-accepted that magnetite is a common by-product of the serpentinization of ultramafic rocks and is associated with H₂ production (Geymond et al., 2023, Geymond et al., 2025, Kularatne and Moretti, 2025, and Deville, 2025). Klein et al. (2009) have investigated iron partitioning and oxidation state evolution during serpentinization of the Mid-Atlantic Ridge (MAR) peridotites and showed that both parameters are sensitive to temperature and water-rock ratio evolution during the reaction (Klein et al., 2009). Indeed, the serpentinization of such peridotites can cause the formation of Fe³⁺-serpentine, with additional brucite when the temperature is below 320 °C and magnetite when the temperature is above 200 °C (Klein et al., 2013). Thus, ferric iron partitioning is dependent on the temperature of the reaction, as well as the mineral composition of the protolith. However, under low to moderate temperatures, serpentinization can lead to most of the Fe³⁺ being incorporated into serpentine while the amount of magnetite is low and still produces non-negligible amounts of hydrogen (Klein et al., 2013).

Such iron enrichment in continental ultramafic-mafic rocks has already been discussed by Evans et al. (2017). They studied peridotites
from the Duluth Complex, part of the Mid-Continent Rift (Minnesota, USA). The rocks are characterized by a significant Fe-enrichment of the secondary phases relative to the protolith, especially the serpentine that is enriched up to 30 % in iron. Authors infer that this enrichment is the consequence of an enhanced Mg depletion due to an increase in the Fe—Mg exchange potential between the primary and secondary phases. In other words, the water/ rock ratio, as well as the percolating fluid chemistry, were effective enough for the system to go from internally to externally controlled by the circulating fluid, which enhanced the dissolution of the olivine, the leaching of Mg from the system, and ultimately produced a relatively iron-enriched mixture of phyllosilicates with a limited amount of magnetite. This process has been known to happen during the deuteric alteration of olivine and is often associated with iron oxidation, which explains why the produced phyllosilicates are particularly ferric iron-rich (Ross and Shannon, 1925; Edwards, 1938; Deer et al., 1962). Authors state that even if the mineral assemblage is relatively magnetite-poor, the hydrogen yield produced during this complex suite of reactions could be 10 times larger than for an ophiolitic dunite with a forsteritic olivine (Evans et al., 2017).

A more recent study conducted on serpentinized lherzolites from a passive continental margin has highlighted samples with Fe-rich serpentine associated with little magnetite, and no brucite (Albers et al., 2021). Such ultramafic serpentinized rocks are characterized by 2/3 of the total ferric iron being accounted for by the Fe-serpentine. Even though magnetite is not the main phase hosting Fe³⁺, the mass balance calculation performed by the authors established that the hydrogen production due to serpentization could have been consequent (Albers et al., 2021).

Thus, as stated by Evans et al. (2017), the classical model for serpentinized ultramafic rocks is the production of secondary products less enriched in Fe than the initial forsterite olivine, and most of the ferric iron being accounted for by the magnetite. This model appears to be effective for serpentinization of mantle peridotites. It might not always reflect what is happening in all ultramafic-mafic systems, especially the continental ones, where the initial olivine composition ranges between the forsterite and fayalite endmembers.

In this study, the specific chemical composition of the parent and product minerals, especially regarding the Fe content, allows us to better understand the associated redox conditions during pseudomorphic reactions in an open system. Using STXM-XANES on serpentine, the amount of Fe³⁺/Fe<sub>total</sub> was quantified. For amphibole and magnetite, the Fe³⁺/Fe<sub>total</sub> ratios were estimated based on structural formulae. Olivine and orthopyroxene do not contain any Fe³⁺ in opposition to the product phases (Table 1). The overall reaction is oxidizing, but performing a mass balance evaluation allows us to determine the redox of each sub-reaction. Indeed, an advantage of the mass transfer map approach, where mass balance calculation is applied to each pixel of a mineral, is the ability to see the actual redistribution of the chemical

evolution of the system (Centrella, 2024). The balance between the gain and loss of Fe³⁺ and Fe²⁺, respectively, for every product phase, supports similar redox conditions for every sub-reaction. In addition, the actual values of mass change regarding both reactions are also similar. Interpretation of such conditions might have an impact on H₂ generation. Quantifying the actual amount of H₂ generated is complex and requires a multi-variable complex modeling approach. However, the mass transfer map can still be a powerful tool to get a rough estimation of the amount of H (and by extension H₂) generated. Because the changes in element mass are quantified, it is possible to estimate a theoretical H₂ map (Fig. 8) based on the Fe³⁺ mass change following the reaction: 2Fe²⁺ + 2H₂O = 2Fe³⁺ + 2OH⁻ + H₂.

In conclusion, in this monzo-diorite, magnetite occurrence is less evident than in more traditional serpentinized mantle peridotites. However, even if poorly represented, magnetite is still found as rims of the olivine and associated with the orthopyroxene and the Fe-rich amphibole. Its overall amount is very limited compared to the other Fe³⁺-rich secondary phases, which might be explained by the fact that most of the ferric iron was more favorably incorporated into the serpentine and the amphibole. The reasons why ferric iron is somehow more favorably incorporated into those two are not fully deciphered, but, at the system scale, temperature, water/ ratio, kinetic, and initial mineralogy are important factors in controlling Fe partitioning and redox state. Of note, amphibole formation has never been described as an H₂-producing reaction. One could consider the P-T conditions required for such a mineral change to be too high for H₂ to be stable. However, considering that the mass transfer evaluation highlights the net gains of Fe³⁺ and the loss of Fe²⁺, suggesting a clear oxidizing environment, one might question such possibility of producing H₂ through amphibole formation.

## 6. Conclusion

The sample belonging to the Mid-Continent Rift system studied here is partially hydrated with the formation of serpentine and amphibole. The original olivine and orthopyroxene are being replaced by serpentine, amphibole, magnetite, and, to a lesser extent, quartz and apatite. The initial shape of the parent minerals is being preserved, which implies that the hydration reaction occurring here is pseudomorphic.

Mineral compositional maps obtained by electron probe micro-analyzer (EPMA) analysis and mass balance equations allow us to quantify the mass transfer during the hydration of the monzo-diorite. The reaction induces a significant density change, which, coupled with the overall volume preservation, requires the system to lose mass. In addition, the formation of the secondary phases seems to have happened under similar redox and mass transfer conditions, which leads us to believe all secondary minerals were formed at the same time. Regarding the significant iron content of the system and especially the

<figure><img src="image_10.png" /><figcaption>**Fig. 8.** Results of the amount of dihydrogen produced (%) by each mineral for a- reaction 1 (Ol + Opx = Amp + Bt + Srp + Mag + Qtz + (Pl + Afs)) and b- reaction 2 (Ol + Opx + (Pl + Afs) = Amp + Bt + Srp + Mag + Qtz).</figcaption></figure>
ferric iron content of the secondary phases, such as the serpentine and the amphibole, and to a lesser extent the magnetite, one could link this hydration reaction to the potential generation of geologic hydrogen. In other words, the hydration of the olivine-bearing monzo-diorite of the Mid-Continent Rift system in Kansas (USA) presents interesting features of H₂ generation through water-rock interactions. The mineralogical study of such hydrated Precambrian rock and the coupled mass transfer evaluation performed in this work give important clues for future efforts in exploration for geological H₂.

### CRediT authorship contribution statement

**Valentine Combaudon:** Writing – original draft, Methodology, Formal analysis. **Stephen Centrella:** Writing – original draft, Formal analysis, Data curation. **Kanchana Kularatne:** Writing – review & editing, Data curation. **Olivier Sissmann:** Writing – review & editing, Supervision. **Eric Deville:** Writing – review & editing, Supervision. **Hannelore Derluyn:** Writing – review & editing, Supervision, Project administration, Funding acquisition.

### Funding

This research was funded under the framework of E2S UPPA supported by the Nouvelle-Aquitaine Region and by the "France 2030" Investment Program managed by the French National Research Agency (ANR-16-IDEX-0002). H. Derluyn acknowledges the support from the European Research Council (ERC) under the European Union's Horizon 2020 Research and Innovation Program (grant agreement No 850853)".

### Declaration of competing interest

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

### Acknowledgements

The authors thank Arnaud Proietti and Philippe De Parseval for conducting SEM and EMPA analysis at the Centre de Microcaractérisation Raimond Castaing, Université de Toulouse. The authors wish to acknowledge the Kansas Geological Survey and K.D. Newell for their help through sampling. Special thanks to Pascale Sénéchal for assisting in the sample preparation of the K2 sample at DMEX, Université de Pau et des Pays de l'Adour.

## Appendix A. Supplementary data

Supplementary data to this article can be found online at https://doi.org/10.1016/j.gloplacha.2025.105158.

### Data availability

The raw microprobe data used in the current study are available in the Mendeley data repository (https://data.mendeley.com/preview/cp3jf7rpzz?a=0b751948-b231-4a4a-998a-87eb5ae8cdcc).

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