# Insights into the acquisition of kinetic iron isotope signatures of pyrite precipitated through the polysulfide pathway

Franck Poitrasson, Pierre Le Pape, Guillaume Morin, Romain Guilbaud, Camille Baya, Qianyu Deng, Jeremy Rabone, Julie Aufort, Marc Blanchard

<figure><img src="image_1.png" /></figure>

## To cite this version:

Franck Poitrasson, Pierre Le Pape, Guillaume Morin, Romain Guilbaud, Camille Baya, et al.. Insights into the acquisition of kinetic iron isotope signatures of pyrite precipitated through the polysulfide pathway. Earth and Planetary Science Letters, 2025, 674, pp.119729. 10.1016/j.epsl.2025.119729. hal-05386336

HAL Id: hal-05386336

https://hal.science/hal-05386336v1

Submitted on 28 Nov 2025

HAL is a multi-disciplinary open access archive for the deposit and dissemination of scientific research documents, whether they are published or not. The documents may come from teaching and research institutions in France or abroad, or from public or private research centers.

L'archive ouverte pluridisciplinaire HAL, est destinée au dépôt et à la diffusion de documents scientifiques de niveau recherche, publiés ou non, émanant des établissements d'enseignement et de recherche français ou étrangers, des laboratoires publics ou privés.

<figure><img src="image_2.png" /></figure>

Distributed under a Creative Commons Attribution 4.0 International License
<figure><img src="image_3.png" /></figure>

# Insights into the acquisition of kinetic iron isotope signatures of pyrite precipitated through the polysulfide pathway

Franck Poitrasson<sup>a,*</sup>, Pierre Le Pape<sup>b</sup>, Guillaume Morin<sup>b,†</sup>, Romain Guilbaud<sup>a,‡</sup>, Camille Baya<sup>b</sup>, Qianyu Deng<sup>b</sup>, Jeremy Rabone<sup>a</sup>, Julie Aufort<sup>b</sup>, Marc Blanchard<sup>a</sup>

## ARTICLE INFO

Editor: Dr H Bao

**Keywords:**

Iron isotopes

Pyrite

Polysulfide pathway

Kinetic isotope partitioning

Ancient Earth

## ABSTRACT

Sedimentary pyrite (FeS₂) chemistry and isotope compositions have often been used to reconstruct the paleo-redox conditions of the ancient Earth's surface. However, the reaction pathways leading to the formation of this mineral in sedimentary environments remain debated and therefore, its elemental and isotope record is still ambiguous. We thus produced iron isotope time series of pyrite precipitation experiments at ambient temperature. Pyrite minerals were produced through the polysulfide pathway, one of the main mechanisms of pyrite production on Earth. The analysis of precipitated solids and coexisting solutions show that the iron isotope partitioning departs increasingly from equilibrium as the reaction proceeds until the end of our experiments (1510 h). We observe a strong inversion of the isotope partitioning between solid and solution coinciding with the transformation of disordered mackinawite to pyrite. Experiments with added nickel or arsenic, which impact the nucleation kinetics of pyrite, shows that the timing of pyrite nucleation is faithfully recorded by the iron isotope partitioning. However, whatever the timing of the isotope inversion, the final partitioning between pyrite and solution leads to the same Δ⁵⁷Fe ~ -3.5 ‰. This apparent fractionation is similar to that previously found in pyrite precipitation experiments through the H₂S pathway. Combined with new ab initio equilibrium fractionation factors computation and previous X-ray absorption spectroscopic measurements, we propose that the observed fractionation may be explained by isotope exchange with the same key, isotopically light iron molecular intermediate, leading to pyrite nucleation and growth in both the so-called H₂S and polysulfide pathways. At ambient temperature, this fractionation process corresponds to the lightest δ⁵⁷Fe isotope signatures found in the Archean pyrite sedimentary record and it may sign a specific kinetic and irreversible reaction common to both main abiotic pyrite production pathways.

## 1. Introduction

Pyrite (FeS₂) is the most stable and most common iron sulfide mineral in sedimentary environments (Rickard and Luther, 2007). As such, it is often considered as a valuable archive of ancient ocean chemistry, through its trace element contents (e.g., Gregory et al., 2017; Large et al., 2017), its sulfur (S) isotope compositions (Strauss, 1997) and its iron (Fe) isotope compositions (Rouxel et al., 2005). In low-temperature, sedimentary environments, the largest Fe isotope fractionations occur in association with pyrite formation (e.g., Rolison et al., 2018), dissimilatory iron reduction (DIR) (e.g., Beard et al., 1999; Icopini et al., 2004;

Crosby et al., 2005; Johnson et al., 2005) and Fe(II) oxidation (e.g., Johnson et al., 2002; Welch et al., 2003; Croal et al., 2004; Beard et al., 2010; Kappler et al., 2010; Amor et al., 2016).

However, the use of the geochemical record of pyrite is plagued by its complex and still debated crystallization mechanism. Pyrite formation is considered to be often biologically mediated in nature, as marine sulfide production is tightly linked to bacterial sulfate reduction. However, once aqueous sulfide is produced, pyrite can form abiotically, and two main possible pathways have been proposed for its nucleation and growth: the polysulfide pathway and the H₂S pathway, respectively (Rickard and Luther, 2007). Understanding how elements and S and Fe

<footnote><sup>a</sup> Laboratoire Géosciences Environnement Toulouse, UMR 5563, Centre National de la Recherche Scientifique - Université de Toulouse - Institut de Recherche pour le Développement - Centre National d'Etudes Spatiales, 14-16b, avenue Edouard Belin, 31400 Toulouse, France</footnote>

<footnote><sup>b</sup> Institut de Minéralogie, de Physique des Matériaux et de Cosmochimie, UMR 7590, Centre National de la Recherche Scientifique - Sorbonne Université - Museum National d'Histoire Naturelle - Institut de Recherche pour le Développement, 4, place Jussieu, 75252, Paris cedex 05, France</footnote>

<footnote>* Corresponding author.</footnote>

<footnote>E-mail address: Franck.Poitrasson@cnrs.fr (F. Poitrasson).</footnote>
isotopes partition between forming pyrite, its precursors and the ambient fluid is essential to interpret properly the geological record. Sulfur isotope compositions of sedimentary and synthetic pyrite produced via the polysulfide and H₂S pathways reflect isotope contributions from S reservoirs (Butler et al. 2004) with implications on the tracing of microbial activities (Marin-Carbonne et al. 2014). Similarly, microbial processes such as DIR are recurrently invoked in the literature to explain the very light iron isotope compositions of sedimentary pyrite (e.g., Agangi et al., 2015; Archer and Vance, 2006; Czaja et al., 2010; Galic et al., 2017; Johnson et al., 2008; Marin-Carbonne et al., 2014; Nishizawa et al., 2010; Yamaguchi et al., 2005; Yoshiya et al., 2015a; 2015b). However, experimental evidence demonstrates that abiotic sulfidation processes also result in isotopically light products due to kinetic effects (Guilbaud et al., 2011a; Heard et al., 2020; Mansor and Fantle, 2019). Hence, iron isotope partitioning during pyrite formation appears as a promising redox paleo-proxy (Guilbaud et al. 2011a; Mansor and Fantle 2019; Heard et al., 2020), which deserves further mechanistical investigations.

Previous studies have shown that abiotic pyrite production through the H₂S pathway without Fe redox change may generate a large kinetic Fe isotope fractionation (Guilbaud et al., 2011a; Heard et al., 2020), which may be recorded in Archean sedimentary pyrite (e.g., Rouxel et al., 2005). Yet, at marine pH, the speciation of aqueous sulfide may include increasing proportions of polysulfide, which also play a role in pyrite formation (Rickard and Luther, 2007). However, the isotope effect of this other major abiotic route to pyrite production, the polysulfide pathway, remains unknown at ambient temperature.

Previous pyrite synthesis experiments conducted at ambient temperature through the polysulfide pathway have shown that the occurrence of aqueous trace elements may accelerate (nickel, Ni) or delay (arsenic, As) the kinetics of the reaction (Baya et al., 2021), due to the complex role of these elements in pyrite nucleation (Morin et al. 2017; Le Pape et al. 2017). Whether such a trace element-induced kinetic effect on pyrite formation could also affect the Fe isotope signature recorded in pyrite is currently unknown. To tackle this, and given the natural enrichment of sedimentary pyrite in arsenic and nickel (Gregory et al. 2015) and the implication this may have on our understanding on the Fe isotope record of pyrite from paleo-ocean sediments, we studied the Fe isotope partitioning of a series of pyrite precipitation experiments conducted at 25 °C previously studied by Extended X-ray Absorption Spectroscopy (EXAFS) (Baya et al., 2021). We explored the isotope evolution of solids and solutions during pyrite formation, with and without addition of aqueous Ni or As at trace concentrations. Combined with new ab initio calculation to back up the interpretation of our experimental isotope results (see, e.g., Pokrovski et al., 2021), our isotope approach provides better insights into the reaction mechanisms involved in pyrite production at low temperature through the polysulfide pathway.

## 2. Samples studied

The solid and aqueous samples investigated in this work were selected from the set of pyrite synthesis experiments conducted by Baya et al. (2021) in JACOMEX glovebox (<5 ppm O₂), for which Fe and S

<table><caption>**Table 1**
Iron isotope compositions of solids and solutions and pyrite fraction in the solid of precipitation experiments at 25 °C.</caption><thead><tr><th>Sample<sup>a</sup>:</th><th>Time (h)</th><th>δ<sup>57</sup>Fe (‰)<sup>b</sup></th><th>2SE</th><th>δ<sup>56</sup>Fe (‰)<sup>b</sup></th><th>2SE</th><th>n</th><th>Δ<sup>57</sup>Fe<sub>solid-liq</sub> (‰)</th><th>2SE</th><th>FeS<sub>2</sub> (%)<sup>c</sup></th><th>3SE<sup>d</sup></th></tr></thead><tbody><tr><td colspan="10"><strong>Starting iron:</strong></td></tr><tr><td>Fe(III)Cl<sub>3</sub> init</td><td></td><td>-0.381</td><td>0.058</td><td>-0.260</td><td>0.040</td><td>6</td><td></td><td></td><td></td><td></td></tr><tr><td colspan="10"><strong>Experimental products:</strong></td></tr><tr><td colspan="10"><strong>Fe-S experiment:</strong></td></tr><tr><td>Fe-S_0h</td><td>0</td><td>0.182</td><td>0.076</td><td>0.138</td><td>0.050</td><td>3</td><td>0.593</td><td>0.079</td><td>0</td><td>3</td></tr><tr><td>Fe-S LIQ_0h</td><td>0</td><td>-0.411</td><td>0.022</td><td>-0.278</td><td>0.041</td><td>3</td><td></td><td></td><td></td><td></td></tr><tr><td>Fe-S_252h</td><td>252</td><td>0.110</td><td>0.141</td><td>0.063</td><td>0.116</td><td>3</td><td>0.304</td><td>0.321</td><td>0</td><td>3</td></tr><tr><td>Fe-S LIQ_252h</td><td>252</td><td>-0.194</td><td>0.288</td><td>-0.381</td><td>0.190</td><td>3</td><td></td><td></td><td></td><td></td></tr><tr><td>Fe-S_324h</td><td>324</td><td>0.123</td><td>0.198</td><td>0.065</td><td>0.107</td><td>3</td><td>0.587</td><td>0.408</td><td>11</td><td>2</td></tr><tr><td>Fe-S LIQ_324h</td><td>324</td><td>-0.463</td><td>0.356</td><td>-0.334</td><td>0.230</td><td>3</td><td></td><td></td><td></td><td></td></tr><tr><td>Fe-S_492h</td><td>492</td><td>-2.898</td><td>0.051</td><td>-1.942</td><td>0.027</td><td>6</td><td>-3.163</td><td>0.068</td><td>76</td><td>5</td></tr><tr><td>Fe-S LIQ_492h</td><td>492</td><td>0.265</td><td>0.045</td><td>0.186</td><td>0.020</td><td>6</td><td></td><td></td><td></td><td></td></tr><tr><td>Fe-S_1510h</td><td>1510</td><td>-3.299</td><td>0.126</td><td>-2.233</td><td>0.049</td><td>3</td><td>-3.665</td><td>0.271</td><td>100</td><td>0</td></tr><tr><td>Fe-S LIQ_1510h</td><td>1510</td><td>0.366</td><td>0.240</td><td>0.233</td><td>0.233</td><td>3</td><td></td><td></td><td></td><td></td></tr><tr><td colspan="10"><strong>Fe-S-Ni experiment:</strong></td></tr><tr><td>Fe-S:Ni_0h</td><td>0</td><td>0.184</td><td>0.023</td><td>0.139</td><td>0.044</td><td>3</td><td>0.580</td><td>0.033</td><td>1</td><td>5</td></tr><tr><td>Fe-S:Ni LIQ_0h</td><td>0</td><td>-0.396</td><td>0.024</td><td>-0.279</td><td>0.024</td><td>3</td><td></td><td></td><td></td><td></td></tr><tr><td>Fe-S:Ni_114h</td><td>114</td><td>-0.177</td><td>0.068</td><td>-0.133</td><td>0.028</td><td>3</td><td>0.132</td><td>0.148</td><td>16</td><td>1</td></tr><tr><td>Fe-S:Ni LIQ_114h</td><td>114</td><td>-0.309</td><td>0.131</td><td>-0.218</td><td>0.045</td><td>3</td><td></td><td></td><td></td><td></td></tr><tr><td>Fe-S:Ni_228h</td><td>228</td><td>-3.259</td><td>0.073</td><td>-2.212</td><td>0.043</td><td>6</td><td>-3.531</td><td>0.376</td><td>88</td><td>4</td></tr><tr><td>Fe-S:Ni LIQ_228h</td><td>228</td><td>0.273</td><td>0.369</td><td>0.178</td><td>0.215</td><td>3</td><td></td><td></td><td></td><td></td></tr><tr><td>Fe-S:Ni_1510h</td><td>1510</td><td>-3.356</td><td>0.443</td><td>-2.258</td><td>0.319</td><td>3</td><td>-3.641</td><td>0.462</td><td>86</td><td>2</td></tr><tr><td>Fe-S:Ni LIQ_1510h</td><td>1510</td><td>0.285</td><td>0.131</td><td>0.163</td><td>0.063</td><td>3</td><td></td><td></td><td></td><td></td></tr><tr><td colspan="10"><strong>Fe-S-As experiment:</strong></td></tr><tr><td>Fe-S:As_0h</td><td>0</td><td>0.210</td><td>0.129</td><td>0.142</td><td>0.076</td><td>3</td><td>0.770</td><td>0.135</td><td>0</td><td>0</td></tr><tr><td>Fe-S:As LIQ_0h</td><td>0</td><td>-0.560</td><td>0.040</td><td>-0.376</td><td>0.061</td><td>3</td><td></td><td></td><td></td><td></td></tr><tr><td>Fe-S:As_492h</td><td>492</td><td>0.128</td><td>0.200</td><td>0.086</td><td>0.130</td><td>3</td><td>0.512</td><td>0.205</td><td>1</td><td>2</td></tr><tr><td>Fe-S:As LIQ_492h</td><td>492</td><td>-0.384</td><td>0.047</td><td>-0.262</td><td>0.022</td><td>6</td><td></td><td></td><td></td><td></td></tr><tr><td>Fe-S:As_758h</td><td>758</td><td>-3.037</td><td>0.039</td><td>-2.046</td><td>0.028</td><td>5</td><td>-3.336</td><td>0.065</td><td>86</td><td>2</td></tr><tr><td>Fe-S:As LIQ_758h</td><td>758</td><td>0.298</td><td>0.052</td><td>0.194</td><td>0.058</td><td>3</td><td></td><td></td><td></td><td></td></tr><tr><td>Fe-S:As_860h</td><td>860</td><td>-3.359</td><td>0.219</td><td>-2.279</td><td>0.142</td><td>6</td><td>-3.623</td><td>0.221</td><td>90</td><td>1</td></tr><tr><td>Fe-S:As LIQ_860h</td><td>860</td><td>0.264</td><td>0.034</td><td>0.163</td><td>0.040</td><td>3</td><td></td><td></td><td></td><td></td></tr><tr><td>Fe-S:As_1510h</td><td>1510</td><td>-3.111</td><td>0.140</td><td>-2.112</td><td>0.077</td><td>3</td><td>-3.492</td><td>0.155</td><td>92</td><td>8</td></tr><tr><td>Fe-S:As LIQ_1510h</td><td>1510</td><td>0.381</td><td>0.066</td><td>0.242</td><td>0.068</td><td>3</td><td></td><td></td><td></td><td></td></tr></tbody></table>

<footnote><sup>a</sup> liq: aqueous solutions, the rest are recovered solids. <sup>b</sup> The iron isotope composition and two standard error (2SE) uncertainties quoted are calculated from the number of analyses indicated (n) and using the Student's t-correcting factors (Platzner, 1997). <sup>c</sup> Molar percentage of iron in the form of pyrite in the solid, estimated from linear combination fitting analysis of Fe K-edge EXAFS spectra, uncertainty on the last digit value is given under bracket (Baya et al., 2021). <sup>d</sup> Uncertainty on the last digit rounded to the closest integer.</footnote>
K-edge X-ray absorption analysis was previously performed. Briefly, pyrite syntheses were performed at room temperature (25±3 °C) in continuously shaken glass vials under anoxic conditions (< 1 ppm O₂ in a N₂ atmosphere). Three types of experiments were conducted: one with 0.1 % (mol/mol) of Ni added relative to Fe (experiment named Fe-S:Ni in Table 1), one with As added (experiment Fe-S:As) in the same proportions, and a last one without trace metal addition (experiment Fe-S). For all experiments, the initial Fe³⁺Cl₃ was added to a concentration of 0.016 mol/L. The experiments lasted up to 1510 h and solids and supernatant solutions were regularly sampled in the glove box where the experiments were conducted. At each sampling time, suspensions were centrifugated 5–7 min at 6000 rpm. The recovered solids were washed twice with O₂-free milliQ water (N₂ bubbling for 45 min at 80 °C) and vacuum-dried down to 0.1 mbar for ≥ 24 h in a desiccator within the glove box. The remaining supernatants were filtered through 0.2 µm cellulose filters and stored for further analysis. The initial and final pH values were in the 5.6–5.8 and 3.7–4.5 ranges, respectively, this decrease being interpreted as resulting from polysulfide cycling during pyrite formation (Baya et al. 2021).

The mineralogical composition and crystal chemistry of the solid samples were analyzed by Baya et al. (2021) using powder X-ray diffraction, scanning electron microscopy coupled to energy dispersive analysis (SEM-EDX) and X-ray absorption spectroscopy at the Fe- and S K-edges. According to these analyses, the set of samples selected in the present study consists in two main types (Table 1): i) the samples of the first type contained mainly (>80 %) the FeS(nano,clusters) precursor phase (Baya et al., 2021) and were collected at 0, 252, and 324 h for the Fe-S experiment, at 0 and 114 h for the Fe-S:Ni experiment and at 0 and 492 h for the Fe-S:As experiment; (ii) the samples of the second type contained mainly (>76 %) pyrite (FeS₂) and were collected at 492 and 1510 h, 228 and 1510 h, and 758, 860 and 1510 h, for the Fe-S, Fe-S:Ni and Fe-S:As experiments, respectively. Both the solid samples and their filtered supernatant counterparts were analyzed for iron isotope composition in this study, as detailed hereafter.

Typically, in this protocol, initial Fe³⁺ was fully and immediately reduced to Fe²⁺ by H₂S that was partially oxidized to polysulfides and S⁰. Then, Fe²⁺ partially precipitated with H₂S to form FeS. Pyrite finally formed via the classical polysulfide pathway, through the reaction of FeS with S⁰/polysulfides (Morin et al. 2017; Le Pape et al. 2017; Baya et al. 2021, 2022; Deng et al. 2025). According to the obtained pyrite to total Fe ratio, our synthesis protocol (Baya et al. 2021) is consistent with anoxic and Fe(II)-rich (ferruginous) conditions (0 ≤ FePy/FeHR ≤ 0.7/0.8; where subscripts "py" and "HR" stand for "pyrite" and "highly reactive", respectively; Poulton and Canfield, 2011). The time-scale for pyrite formation in our experiment is typically within the range of the estimated time for frambooidal pyrite formation during early-diagenesis of sediments (e.g., Rickard 2019).

### 3. Analytical methods

The iron isotope composition of the experimental solids and solutions were determined by MC-ICP-MS after sample mineralization and Fe purification. Aliquots of experimental solutions were pipetted and taken to dryness, whereas around 1 mg of solid samples was weighed using a balance accurate to <100 µg. Samples were decomposed in aqua regia produced from bi-distilled acids at 120 °C on a hotplate during 24 h. After evaporation, samples were redissolved in 5 ml HCl 6 M for the same time and temperature followed by evaporation to dryness. The final residue was redissolved in 6 M HCl, and an aliquot containing 10 to 100 µg of Fe was loaded onto homemade thermoretractable Teflon columns filled with 0.5 ml of Bio Rad AG1 × 4 (200–400 mesh) resins, following the protocol described in Poitrasson et al. (2004). As noted in a previous pyrite precipitation study, but conducted under high temperature hydrothermal conditions (Pokrovski et al., 2021), it was found that this column extraction step should be repeated once to ensure that the purified Fe fraction was exempt from most of the initial sulfur. Iron was

then redissolved in 0.05 M HCl and stored in cupboards for subsequent isotope analysis. Overall procedural blanks were found to be <3–4 ng of Fe in the course of this study, which is <1000 times the amount of experimental iron processed. This therefore does not affect the uncertainties of our isotope measurements.

Iron isotope measurements were subsequently conducted on a Thermo Neptune+ MC-ICP-MS at GET-OMP following the methods described in Poitrasson and Freyder (2005). Mass bias correction involved a combination of sample-standard bracketing and daily regression method using Ni doping. Each sample was analyzed typically between 3 and 6 times, preferably over different analytical sessions. The Fe isotope data were reported in the standard delta notation relative to the European reference material IRMM-14, expressed in permil (‰). For the ⁵⁷Fe/⁵⁴Fe ratio, this yields:

$$ \delta^{57}\text{Fe} = \left( \frac{^{57}\text{Fe}/^{54}\text{Fe}_{\text{sample}}}{^{57}\text{Fe}/^{54}\text{Fe}_{\text{IRMM-14}}} - 1 \right) \times 1000 $$

Data quality was checked by repeated analyses of our in-house hematite standard every 6 samples during the analytical sequence. Long-term external reproducibility was estimated from replicate analyses of this hematite in every analytical session. The mean of δ⁵⁷Fe individual hematite values pooled by groups of 6 yielded δ⁵⁷Fe = 0.758 ± 0.047 ‰ (2SD, N = 14), in good agreement with previous works (see Akerman et al., 2024, and references therein). Uncertainties on δ⁵⁷Fe and δ⁵⁶Fe values reported in this study are expressed as two standard errors (2SE) calculated from the number of replicates and using the Student's t-correcting factor (Platzner, 1997). The international reference material IF-G analyzed during this work gave δ⁵⁷Fe = 0.929 ± 0.113 ‰ (2SE), in good agreement with previous determinations (e.g., Busigny and Dauphas, 2007; Fabre et al., 2011).

### 4. Atomistic modelling

Measured isotope fractionations between aqueous species and solid phases can be compared with thermodynamic equilibrium fractionations, αA-B, between two phases (or species) A and B, determined from the reduced partition function ratios, also called β-factors, as:

$$ 1000\ln\alpha_{A-B} = 1000\ln\beta_A - 1000\ln\beta_B \approx \Delta^{57}\text{Fe}_{A-B} \text{ (in \textperthousand}} $$

Here we supplement the data from Fujii et al. (2014) and Pokrovski et al. (2021) with additional aqueous species, using strictly the same calculation parameters, in order to build up a consistent database of β-factors (Table 2). DFT calculations as implemented in the Gaussian 16 code (Frisch et al., 2016) were performed with the hybrid DFT B3LYP functional (Becke, 1993; Lee et al., 1988). Electronic wave functions were described by the 6–311+G(d,p) basis set for all atoms. An ‘ultra-fine’ numerical integration grid was used, and the self-consistent field (SCF) convergence criterion was set to 10⁻⁹. Only the first coordination sphere around the Fe atom was considered, without any forced symmetry.

As thoroughly described by Blanchard et al. (2017), the β-factor can be determined by computing vibrational frequencies using a harmonic approach (Bigeleisen and Mayer, 1947):

$$ \beta = \left[ \prod_{i=7}^{3N} \frac{\nu_i}{\nu_i} \times \frac{\exp(-h\nu_i/2kT)}{1 - \exp(-h\nu_i/kT)} \times \frac{1 - \exp(-h\nu'_i/kT)}{\exp(-h\nu'_i/2kT)} \right]^{1/n} $$

where N is the number of atoms in the molecule and n the number of exchangeable atoms. νi and ν'_i correspond to the harmonic frequency of the vibrational mode i before and after the isotope substitution. T is the temperature in Kelvin, and h and k are the Planck and Boltzmann constants, respectively. The main structural and electronic factors affecting theoretical equilibrium isotope fractionation properties, including bond-length, coordination and spin-state are discussed in e.g., Schauble (2004) and Blanchard et al. (2017).
<caption>**Table 2**</caption>

Calculated ⁵⁷Fe/⁵⁴Fe β-factors at 298 K (25 °C) for FeS₂ pyrite, FeS as ordered or disordered mackinawite, and Fe(III) and Fe(II) ions. Polynomial fits of Fe β-factors as a function of temperature for aqueous species calculated in this study, 1000lnβ = ax + bx² with x = 10⁶/T² (T in K). Molecular structures of the aqueous species investigated in the present study are described in Fig. S1 and Table S1.

<table><thead><tr><th>Species / Phases</th><th>ox. state</th><th>CN</th><th>Mag. spin <sup>c</sup></th><th>⁵⁷/⁵⁴Fe β (%)</th><th>a</th><th>b × 10<sup>-3</sup></th></tr></thead><tbody><tr><td>FeS<sub>2</sub> pyrite</td><td>II</td><td>6</td><td>ls</td><td>13.05 <sup>d</sup></td><td></td><td></td></tr><tr><td>FeS as ordered / disordered mack.</td><td>II</td><td>4</td><td>hs</td><td>11.47 / 6.78 <sup>e</sup></td><td></td><td></td></tr><tr><td>FeS as arbitrary mixture of 20 % ordered &amp; 80 % disordered mack.</td><td>II</td><td>4</td><td>hs</td><td>7.72</td><td></td><td></td></tr><tr><td>[Fe(H<sub>2</sub>O)<sub>6</sub>]<sup>3+</sup></td><td>III</td><td>6</td><td>hs</td><td>11.85</td><td>1.1316</td><td>-6.99</td></tr><tr><td>[FeCl<sub>3</sub>(H<sub>2</sub>O)<sub>2</sub>]</td><td>III</td><td>6</td><td>hs</td><td>9.30</td><td>0.8717</td><td>-4.03</td></tr><tr><td>[Fe(H<sub>2</sub>O)<sub>6</sub>]<sup>2+</sup></td><td>II</td><td>6</td><td>hs</td><td>7.47 <sup>e</sup></td><td></td><td></td></tr><tr><td>[FeCl(H<sub>2</sub>O)<sub>5</sub>]<sup>3+</sup></td><td>II</td><td>6</td><td>hs</td><td>7.18 <sup>e</sup></td><td></td><td></td></tr><tr><td>[FeCl<sub>2</sub>(H<sub>2</sub>O)<sub>4</sub>]</td><td>II</td><td>6</td><td>hs</td><td>6.82 <sup>e</sup></td><td></td><td></td></tr><tr><td>Average Fe(II) spec.</td><td>II</td><td>6</td><td>hs</td><td>7.16</td><td></td><td></td></tr><tr><td>[Fe(SH)<sub>2</sub>(H<sub>2</sub>O)<sub>2</sub>]</td><td>II</td><td>4</td><td>hs</td><td>6.64 <sup>e</sup></td><td></td><td></td></tr><tr><td>[Fe(SH)<sub>4</sub>]<sup>2-</sup></td><td>II</td><td>4</td><td>hs</td><td>5.08 <sup>e</sup></td><td></td><td></td></tr><tr><td>[Fe(SH)<sub>2</sub>(S<sub>3</sub>)<sub>2</sub>]<sup>4-</sup></td><td>II</td><td>4</td><td>hs</td><td>4.07</td><td>0.3690</td><td>-0.66</td></tr><tr><td>[Fe(SH)<sub>2</sub>(S<sub>3</sub>)<sub>2</sub>]<sup>4+</sup></td><td>II</td><td>6</td><td>ls</td><td>9.29</td><td>0.8555</td><td>-2.72</td></tr><tr><td>[Fe(S<sub>3</sub>)<sub>2</sub>]<sup>2-</sup></td><td>II</td><td>6</td><td>ls</td><td>7.65</td><td>0.6963</td><td>-1.54</td></tr><tr><td>[Fe(S<sub>3</sub>)<sub>2</sub>]</td><td>II</td><td>6</td><td>ls</td><td>11.55</td><td>1.0645</td><td>-3.41</td></tr><tr><td>[Fe(S<sub>3</sub>)<sub>3</sub>]<sup>4+</sup></td><td>II</td><td>6</td><td>ls</td><td>8.50</td><td>0.7765</td><td>-1.94</td></tr><tr><td>[Fe(S<sub>3</sub>)<sub>3</sub>]<sup>-</sup></td><td>II</td><td>6</td><td>hs</td><td>9.41</td><td>0.8629</td><td>-2.42</td></tr></tbody></table>

## 5. Isotope results

The iron isotope compositions of experimental solids and coexisting solutions for pyrite precipitation experiments without element added, and with Ni or As added are reported in Table 1, and represented graphically as a function of experimental time on Fig. 1a. The Fe isotope partitioning between solids and solutions (liq), expressed as:

$$ \Delta^{57}\text{Fe}_{\text{solid-liq}} = \delta^{57}\text{Fe}_{\text{solid}} - \delta^{57}\text{Fe}_{\text{liq}} $$

is also reported on Table 1 and Fig. 1b The fraction of solid consisting of crystalline pyrite determined from Linear Combination fitting analysis of Fe K-edge EXAFS spectra (Baya et al., 2021) is given in Table 1 for reference.

At the start of the experiment, the dissolved Fe pool is isotopically similar to the starting Fe<sup>III</sup>Cl<sub>3</sub> solution used for all three experiments (δ<sup>57</sup>Fe = -0.381±0.058 ‰; Table 1), and then becomes heavier towards the end of the experiments. Conversely, experimental solids start with δ<sup>57</sup>Fe values that are higher than the starting Fe<sup>III</sup>Cl<sub>3</sub> solution and then sharply drop down to values clustering around δ<sup>57</sup>Fe = -3.26±0.13 ‰ (Fig. 1). Note that for all experiments, this drop in the solid δ<sup>57</sup>Fe composition is in concert with the onset of pyrite formation (Table 1; Fig. 1a), i.e., for the samples taken ≥228 h after the start of the Fe-S:Ni-doped experiment, ≥492 h after the start of the experiment Fe-S without trace element added and ≥758 h after the start of the Fe-S:As-doped experiment. These strong decreases in δ<sup>57</sup>Fe<sub>solid</sub> coincide with a more moderate increase in δ<sup>57</sup>Fe<sub>liq</sub> values, resulting in a sharp decrease of the Fe isotope partitioning between solid and solution expressed as Δ<sup>57</sup>Fe<sub>solid-liq</sub> (Table 1 and Fig. 1b). Indeed, at the start of each experiment, Δ<sup>57</sup>Fe<sub>solid-liq</sub> values range from 0.5 to 0.7 ‰ for all three sets of experiments and end up with remarkably similar Δ<sup>57</sup>Fe<sub>solid-liq</sub>, between -3.5 and -3.7 ‰ at the end of the experiments.

## 6. Discussion

### 6.1. Iron isotope inversion

A first striking feature of these experimental results is the inversion of the Fe isotope partitioning, with an overall Δ<sup>57</sup>Fe<sub>solid-liq</sub> range of ~4 ‰ for all three experiments (Table 1; Fig. 1b), which is at the upper end of the range according to recent compilations of Fe isotope partitioning experiments between solids and liquids at ambient temperature (Dauphas et al., 2017; Johnson et al., 2020; Wu et al., 2019). Given the closed system nature of the experimental set-up, this inversion of the isotope partitioning implies isotope exchange between solid and solution over the course of the experiment. The much stronger shift of the solid Fe isotope composition relative to that remaining in solution (Fig. 1a) implies that iron dominantly remains in aqueous form in these closed system throughout the experimental duration. Assuming the samplings were representative of all experimental solids and liquids, we compute by mass balance using Fe isotope signatures that 20±2 ‰ (1SD) of the starting Fe was sequestered in pyrite at the end of the experiments (1510 h), the rest remaining in solution. This percentage is close to the value of 17±1 ‰ obtained from iron mass balance in Baya et al. (2022) for similar experimental conditions. According to Baya et al. (2021), the solid precursor that forms during the lag phase of the synthesis experiment, i.e., before pyrite nucleation, consists in a FeS<sub>(nano,cluster)</sub> phase that is interpreted as a mixture of nanocrystalline mackinawite and aggregated FeS clusters. For comparison with computed β-factors and for the sake of simplicity, such a complex FeS precursor phase may be initially assimilated to poorly-crystalline mackinawite (FeS), which dissolves and further reacts to form pyrite (FeS<sub>2</sub>) towards the end of the synthesis experiments. Comparing the computed β-factors at 25 °C of these two phases, respectively 6.78 ‰ (estimation for disordered mackinawite; Pokrovski et al., 2021) and 13.05 ‰ (Blanchard et al., 2009), with those of Fe<sup>3+</sup> aqueous complexes in the starting Fe<sup>III</sup>Cl<sub>3</sub> solution (9.30 ‰ for [FeCl<sub>3</sub>(H<sub>2</sub>O)<sub>3</sub>]<sup>0</sup>, Table 2) would lead to equilibrium Δ<sup>57</sup>Fe partitioning between solid and solution of -2.5 ‰ at the start of the experiment, increasing to +3.8 ‰ towards the end of the experiment, when pyrite has formed (Table 3). Hence, the observed Δ<sup>57</sup>Fe<sub>solid-liq</sub> isotope partitioning measured in our experiments, starting with values of 0.5 to 0.7 ‰ at the FeS precursor stage, and ending up with values of -3.5 to -3.6 ‰ at the pyrite stage, are unlikely to reflect isotope equilibration with pyrite according to the corresponding β-factor.

In our pyrite synthesis protocol, the entire pool of aqueous Fe<sup>3+</sup> is immediately reduced by aqueous sulfide to Fe<sup>2+</sup>, a proportion of which (here ~20±2 ‰) precipitates with H<sub>2</sub>S to form the FeS precursor and finally pyrite. The other part of the iron pool remains dissolved, mostly as aqueous Fe(II), the major fraction of the FeS clusters being aggregated due to the high ionic strength (Morin et al., 2017). In line with the experimental conditions, we considered the following aqueous Fe(II) species for the DFT calculations: [Fe(H<sub>2</sub>O)<sub>6</sub>]<sup>2+</sup>, [FeCl(H<sub>2</sub>O)<sub>5</sub>]<sup>+</sup> and [FeCl<sub>2</sub>(H<sub>2</sub>O)<sub>4</sub>]<sup>0</sup>. β-factors tend to decrease with Cl content, but over a small range, with an average value of ~7.2 ‰ at 25 °C (Table 2). Hence, assuming an early and full conversion of Fe<sup>3+</sup> into Fe<sup>2+</sup> in the synthesis, computed equilibrium Δ<sup>57</sup>Fe between solid and solution would still be low at the start of the experiment (-0.4 ‰ vs. 0.5 to 0.7 ‰, respectively) and would shift further away from equilibrium (+5.9 ‰ vs. -3.5 to -3.6 ‰), at the end of the experiment, when pyrite has formed (Table 3). It should be noted that the β-factor calculated for the FeS precursor is uncertain. Whereas the β-factor for crystalline mackinawite (11.47 ‰ at 25 °C, Table 2) was calculated from DFT frequencies, the β-factor for disordered mackinawite (6.78 ‰ at 25 °C) was estimated by applying a scaling factor to the frequencies (Pokrovski et al., 2021). The theoretical and experimental fractionations between FeS precursor and Fe species in solution could be easily reconciled either by using a slightly smaller scaling factor or by considering a mixture of disordered and crystalline mackinawite (to mimic the mixture of aggregated FeS clusters and

<footnote><sup>c</sup> ls = low spin, hs= high spin.</footnote>

<footnote><sup>d</sup> Blanchard et al., 2009.</footnote>

<footnote><sup>e</sup> Pokrovski et al., 2021.</footnote>
<figure><img src="image_4.png" /><figcaption>Fig. 1. a. Iron isotope compositions relative to IRMM-14 (δ<sup>57</sup>Fe) of solid and solution along pyrite precipitation experiments. The black horizontal line represents the Fe isotope composition measured from the starting FeCl<sub>3</sub> solution. b. The fractionation factor (Δ<sup>57</sup>Fe) between solid and solution (noted Δ<sup>57</sup>Fe<sub>solid-liq</sub> in the text for simplicity). Dashed curves are only guides for the eyes: Their subvertical sections correspond to pyrite nucleation timings in the different experiments. Data are from Table 1.</figcaption></figure>

nanocrystalline mackinawite described by Baya et al., 2021). An arbitrary mixture of 80 % of disordered mackinawite and 20 % of crystalline mackinawite would give a β-factor of ~7.7 ‰ (Table 2) and a Δ<sup>57</sup>Fe solid-solution fractionation of +0.6 ‰ in agreement with the measurements (Table 3).

These slightly positive δ<sup>57</sup>Fe values in solids at the FeS precursor stage are consistent with experimentally-determined equilibrium Fe isotope fractionation between FeS and Fe(II)<sub>aq</sub> species at ambient temperature (Guilbaud et al., 2011b; Wu et al., 2012). Specifically, Guilbaud et al. (2011b) reported an equilibrium Δ<sup>57</sup>Fe value of +0.78 ± 0.24 ‰ between FeS and Fe<sup>2+</sup>, while Wu et al. (2012) reported an equilibrium Δ<sup>57</sup>Fe value of +0.96 ± 0.54 ‰ between FeS and FeHS<sup>+</sup>. Similarly, strongly negative δ<sup>57</sup>Fe values in solids at the pyrite stage and apparent steady-state partitioning are in good agreement with previous studies reporting large kinetic isotope effects during pyrite formation, which argue that the extremely low solubility of pyrite at ambient temperature greatly hampers atomic exchange at the mineral surface, preventing the system from reaching equilibrium (Guilbaud et al., 2011a; Heard et al., 2020). This specific isotope composition seems to be robustly locked in the pyrite crystal once formed, as suggested by the apparent steady-state

isotope partitioning observed after pyrite formation (Fig. 1).

6.2. *A key isotopically light intermediate*

We now explore the specific mechanism that imprints such a strong and light Fe isotope signal to pyrite. In a kinetically controlled crystalization regime, lighter isotopes are favored (e.g., Criss, 1999). Given the closed system nature of the experiments, this would therefore imply an extended exchange between the Fe atoms of the aqueous pool and those of the solid FeS precursor pool, with a specific selection of aqueous light iron isotopes to produce pyrite.

Interestingly, low-temperature pyrite formation through the other main abiotic pyrite formation process, the H₂S pathway, also leads to this strong kinetic imprint on the Fe isotope composition of pyrite (Guilbaud et al., 2011a). The authors surmise it could be related to a shift from FeII in a tetrahedrally coordinated high-spin d⁶ configuration in FeSₐq to a FeII in an octahedrally coordinated low-spin d⁶ configuration in pyrite (Guilbaud et al., 2011a). Such a preferential incorporation of isotopically light Fe in newly formed pyrite would be favored by the irreversible pyrite formation reaction, that is accordingly sparingly
<caption>**Table 3**</caption>

Theoretical equilibrium ⁵⁷Fe/⁵⁴Fe α fractionation factors discussed in the text. All FeS phases defined as “20:80 arbitrary mixtures of mack.” correspond to a mixture of 20 % of ordered mackinawite and 80 % of disordered mackinawite. Values are given in ‰ at 298 K (25 °C). For comparison, our experimental Δ⁵⁷Fe/⁵⁴Fe solid-liq isotope partitioning starts with values of 0.5 to 0.7 ‰ at the FeS precursor stage, and ends up with values of −3.5 to −3.6 ‰ at the pyrite stage.

<table><thead><tr><th>Species / Phases A</th><th>Species / Phases B</th><th>⁵⁷/⁵⁴Fe α<sub>A-B</sub></th></tr></thead><tbody><tr><td>FeS as disordered mack.</td><td>[FeCl<sub>3</sub>(H<sub>2</sub>O)<sub>3</sub>]</td><td>−2.5</td></tr><tr><td>FeS<sub>2</sub> pyrite</td><td>[FeCl<sub>3</sub>(H<sub>2</sub>O)<sub>3</sub>]</td><td>3.8</td></tr><tr><td>FeS as disordered mack.</td><td>Average Fe(II) spec.</td><td>−0.4</td></tr><tr><td>FeS as 20:80 arbitrary mixture of mack</td><td>Average Fe(II) spec.</td><td>0.6</td></tr><tr><td>FeS<sub>2</sub> pyrite</td><td>Average Fe(II) spec.</td><td>5.9</td></tr><tr><td>[Fe(SH)<sub>2</sub>(H<sub>2</sub>O)<sub>2</sub>]</td><td>FeS as 20:80 arbitrary mixture of mack.</td><td>−1.1</td></tr><tr><td>[Fe(SH)<sub>4</sub>]<sup>2-</sup></td><td>FeS as 20:80 arbitrary mixture of mack.</td><td>−2.6</td></tr><tr><td>[Fe(SH)<sub>2</sub>(H<sub>2</sub>O)<sub>2</sub>]</td><td>Average Fe(II) spec.</td><td>−0.5</td></tr><tr><td>[Fe(SH)<sub>4</sub>]<sup>2-</sup></td><td>Average Fe(II) spec.</td><td>−1.6</td></tr><tr><td>[Fe(SH)<sub>2</sub>(S<sub>3</sub>)<sup>4-</sup>]</td><td>FeS as 20:80 arbitrary mixture of mack.</td><td>−3.6</td></tr><tr><td>[Fe(SH)<sub>2</sub>(S<sub>3</sub>)<sup>4-</sup>]</td><td>FeS as 20:80 arbitrary mixture of mack.</td><td>1.6</td></tr><tr><td>[Fe(S<sub>5</sub>)<sub>2</sub>]<sup>2-</sup></td><td>FeS as 20:80 arbitrary mixture of mack.</td><td>−0.1</td></tr><tr><td>[Fe(S<sub>5</sub>)<sub>2</sub>]</td><td>FeS as 20:80 arbitrary mixture of mack.</td><td>3.8</td></tr><tr><td>[Fe(S<sub>3</sub>)<sub>3</sub>]<sup>4+</sup></td><td>FeS as 20:80 arbitrary mixture of mack.</td><td>0.8</td></tr><tr><td>[Fe(S<sub>3</sub>)<sub>3</sub>]</td><td>FeS as 20:80 arbitrary mixture of mack.</td><td>1.7</td></tr></tbody></table>

soluble (Rickard and Luther, 2007). Such type of mechanism involving specific isotope composition of putative aqueous FeSaq precursors is also possible in the present set of experiments, and would strongly support the hypothesis of a kinetic imprint of iron isotope composition during low temperature pyrite formation (Guilbaud et al., 2011a).

Indeed, the expected β-factor of tetrahedrally coordinated Fe(II) in high-spin aqueous FeSaq complexes is in the 5.1–6.6 ‰ range, as estimated from DFT calculations (Pokrovski et al., 2021), to be compared to the β-factor of FeS precursor (7.7 ‰ if we take the value of the arbitrary mixture of disordered and crystalline mackinawite; Table 2) and aqueous Fe(II) (~ 7.2 ‰). Thus, depending on Fe utilization from these two sources (i.e., FeS precursor or aqueous Fe(II)), the fractionation toward tetrahedrally coordinated FeSaq complexes would then range between −0.5 and −2.6 ‰, the latter maximum value corresponding to a sole FeS precursor source (Table 3). This source is in agreement with our synthesis reaction. After FeS precipitation, the sulfur necessary to form FeSaq from aqueous Fe(II) would only be supplied by poorly-crystalline FeS dissolution. Thus, assuming a strong preference of Fe for such complexes and a fast and irreversible formation of pyrite from them, leading to a complete locking-in of this fractionation in the final pyrite, these values are still below the observed fractionation, by about 1 ‰. Consequently, this systematic Δ⁵⁷Fe solid-liq ~ −3.5 to −3.6 ‰ observed at low temperature for both the polysulfide and H₂S pathway experiments conducted in a variety of conditions (Guilbaud et al., 2011a; Heard et al., 2020; This study) strongly suggest that a yet unidentified mechanism drives the kinetic iron isotope fractionation during pyrite formation.

We propose that the key molecular iron intermediate to pyrite is an octahedrally-coordinated Fe polysulfide (S<sub>n</sub><sup>2-</sup>) precursor, with an even lighter isotope composition than the FeSaq complexes, which may imprint the kinetic isotope fractionation to pyrite. Indeed, an octahedrally-coordinated Cobalt polysulfide precursor has recently been observed in the polysulfide pathway to pyrite and has been proposed as a possible alternative intermediate to tetrahedrally coordinated precursors in pyrite formation (Deng et al., 2025). In such an octahedral FeS<sub>n</sub> precursor, Fe would be partly bound to polysulfides chains that are longer than S<sub>2</sub><sup>2-</sup> as in pyrite, with a resulting lower electron density on the apical S atoms of the octahedra connected to the polysulfide chains (Deng et al., 2025). In contrast to the configuration of pyrite, this one with a lower crystal-field splitting would favor a high-spin configuration, with a lower expected β-factor.

To test the stability of octahedrally-coordinated iron polysulfide clusters and calculate the associated β-factors, several hypothetical configurations of iron with SH⁺, S<sub>3</sub><sup>•</sup>, S<sub>3</sub><sup>•</sup>, S<sub>5</sub><sup>•</sup> and S<sub>5</sub><sup>•</sup> ions were investigated (Table 2, Table S1, Fig. S1 from the ESM, Fig. 2). The stability of such complexes is strongly influenced by the crowding of the sulfur atoms around iron. Indeed, of the clusters investigated, only those in which polysulfide ions are multidentate and those which are less charged than −4 are found to be stable towards dissociation or reconfiguration. The spin state of iron also plays a role; high-spin octahedral precursors tend to reconfigure to tetrahedral coordination (or exhibit electronic instabilities) while the equivalent low-spin clusters remain octahedral. All of the Fe β-factors of these FeS<sub>n</sub> polysulfide precursors are too large compared to the β-factor of the FeS precursor, with the exception of the tetrahedral [Fe(SH)<sub>2</sub>(S<sub>3</sub>)<sub>2</sub>]<sup>4-</sup> complex, which has a β-factor of 4.07 ‰ at 25 °C (Table 2, Fig. S1). Such a complex would explain the fractionations measured between pyrite and the solution (~ −3.6 ‰, Table 3) in our experiments and those of previous studies (Guilbaud et al., 2011a; Heard et al., 2020), but does not agree with the octahedral coordination suggested by the spectroscopic data (Deng et al., 2025). However, these calculations do not rule out the involvement of octahedral FeS<sub>n</sub> polysulfide precursors as they are not exhaustive. Thus far, we have considered only isolated clusters containing iron, sulfur species and water. Calculations indicate that the instability of these clusters increases with their charge and S atoms crowding, resulting in the cluster dissociation. In calculations of an octahedral, low-spin [Fe(S<sub>3</sub>)<sub>2</sub>(SH)<sub>4</sub>]<sup>6-</sup> complex the two S<sub>3</sub><sup>2-</sup> ions detach to leave the square planar [Fe(SH)<sub>4</sub>]<sup>2-</sup> ion before the electronic convergence becomes unstable (Fig. 2). However, the possibility that larger, anionic clusters are stabilised in solution, either by the surrounding water molecules or through association with counter-ions is anticipated. The larger size of such clusters, and the corresponding expansion of the configuration space, necessitate a separate theoretical study.

## 6.3. A unified mechanistic view of the polysulfide and H₂S pyrite formation pathways

Interestingly, the proposed involvement of such Fe-polysulfide precursors in both the so-called polysulfide and H₂S pathways could on one hand explain similar kinetic isotope fractionation effects in both pathways (Guilbaud et al., 2011a; Pokrovski et al., 2021; This study). On the other hand, it calls for an explanation of the occurrence and importance of polysulfides in the H₂S pathway. To this regard, as suspected from the occurrence of S<sup>0</sup> in the ambient temperature H₂S-pathway pioneering experiment of Rickard (1997), we propose that this pathway likely proceeds first via the corrosion of poorly-crystalline FeS by H⁺ in acidic medium, which leads to the production of Fe(III) from FeS in strictly anoxic conditions (Bourdoiseau et al., 2011; Le Pape et al., 2025), and the subsequent production of polysulfides by the reaction of Fe(III) with sulfide (Baya et al., 2021; Deng et al., 2025; Morin et al., 2017).

Altogether and according to these considerations, the H₂S and polysulfide pathways could be tentatively expressed as follows (Baya et al., 2021):

1) Initiation of the so-called H₂S pathway (Bourdoiseau et al., 2011; Le Pape et al., 2025):

$$ \mathrm{Fe}^{\mathrm{II}}\mathrm{S} + 2\mathrm{xH}^{+} \rightarrow \mathrm{Fe}_{1-3x}\mathrm{Fe}_{\frac{3}{2}x}\mathrm{S} + x\mathrm{Fe}^{2+}_{(\mathrm{aq})} + \mathrm{xH}_{2} $$

2) Initiation of the so-called polysulfide pathway, derived from Morin et al. (2017), Baya et al. (2021):

$$ \mathrm{Fe}_{1-3x}\mathrm{Fe}_{\frac{3}{2}x}\mathrm{S} + \mathrm{x/(n-1)} \mathrm{H}_{2}\mathrm{S}_{(\mathrm{aq})} \rightarrow (1-x) \mathrm{Fe}^{\mathrm{II}}\mathrm{S} + \mathrm{x/(n-1)} \mathrm{H}_{2}\mathrm{S}_{(\mathrm{aq})} $$
<figure><img src="image_5.png" /><figcaption>Fig. 2. Proposed mechanism of iron polysulfide to pyrite transformation that accounts for the evolution of the Fe isotope partitioning observed in our experiments and supported by our ab initio calculation. (Is: low spin; hs: high spin).</figcaption></figure>

3) Production of an octahedral metal-polysulfide intermediate, i.e., “FeS_n” (Deng et al., 2025), either from the FeS or aqueous Fe(II) source:

$$Fe^{II}S + H_2S_{n(aq)} \rightarrow [FeS(H_2S_n)] \rightleftharpoons FeS_n$$

$$Fe^{2+}_{(aq)} + H_2S_{n(aq)} \rightarrow FeS_n + 2H^+$$

4) Putative polymerization to form pyrite, derived from Luther's rationale (Luther, 1991):

$$FeS(H_2S_n) \rightarrow FeS_2 + H_2S_{n-1}$$

This mechanism, which would need to be further confirmed by direct theoretical, spectroscopic and isotope investigations would likely unify our understanding of pyrite formation routes in reconciling the so-called polysulfide and H₂S pathways. Moreover, assuming this mechanism, isotope fractionation would be acquired upon rapid equilibration of aqueous Fe(II) with the metastable FeS_n molecular intermediate during step 3. The acquired signature would then not further depend on the kinetic of polymerization leading to the nucleation and growth described step 4. These latter processes are sufficiently fast to lock the isotope signature of the isotopically light intermediate in the resulting pyrite.

Our proposed mechanism has important geochemical implications and thus calls for further studies. For instance, direct identification of the hypothetical FeS_n octahedral precursor would require to separate such species from the dominant FeS tetrahedral matrix compound. This would represent a complex analytical challenge in order to identify such minor colloidal or surface bound species, requiring XAS approaches combined with physical or chemical separative techniques.

## 6.4. Timing of the isotope partitioning inversion as a function of pyrite trace element content

Another noticeable feature from these experiments is the contrasted timing of the inversion of the iron isotope partitioning between solid and solution depending on the experiments. This isotope inversion was

observed from the t = 228 h sample in the Fe-S:Ni-added experiment, from t = 492 h for the Fe-S experiment without trace element added and from t = 758 h for the Fe-S:As added experiment (Table 1 and Fig. 1). The isotope inversion observed between the solid and solution closely matches the onset of pyrite formation (Table 1), as determined by Baya et al. (2021) from fitting analysis of the K-edge EXAFS spectra. This can also be seen from a graph showing the correlation of Δ⁵⁷Fe between solid and solution and the proportion of pyrite in the precipitated solid (Fig. 3). According to Morin et al. (2017) and Baya et al. (2021; 2022), Ni rapidly forms Ni-rich pyrite nuclei and thus accelerates pyrite formation. In contrast, As has an opposite effect and delays pyrite nucleation (Baya et al., 2021; 2022; Le Pape et al., 2017), possibly through the formation of thiosulfate-As that may co-precipitate with or bind to FeS precursors and thereby slow down pyrite nucleation. The kinetic effects of trace elements on the rate of pyrite formation thus explain that the sharp inversion of the isotope partitioning between solid of solution Fe isotope fractionation (Fig. 1) was observed to faithfully mark the timing of pyrite nucleation at different times from the start, depending on the adding of Ni, As or no trace element (Fig. 1; Table 1). Further exploration of the effect of Ni and As on pyrite nucleation and isotope inversion would benefit from being performed on modern analogues such as mangrove sediments, where rapid pyritization occurs (Noël et al. 2014) in the presence of As (Nguyen et al. 2018; Wang et al., 2018) and Ni (Noël et al. 2015, 2017). In addition, such an isotope inversion is also expected for other chalcophile elements, irrelative of the kinetic effects they exert on pyrite nucleation (Baya et al. 2022).

Nevertheless, the final partitioning remained the same (Δ⁵⁷Feₚy·liq ~ -3.5 to -3.6 ‰) for all three sets of experiments, whatever the duration of the lag time before pyrite formation. This indicates that the Fe isotope partitioning between final pyrite and aqueous Fe did not depend on the nucleation or on the growth rate. Accordingly, within the unified pyrite formation framework proposed in section 6.3 above, iron isotope fractionation is acquired at step 3 upon equilibration with the molecular octahedral FeS_n intermediate, whereas trace elements accelerate or delay pyrite nucleation and growth in step 4. To this regard, Ni²⁺ and Co²⁺ likely exhibit a strong affinity for polysulfides to form NiS_n and CoS_n intermediates, the former rapidly nucleating to pyrite (Baya et al., 2021; Morin et al., 2017), the later rapidly forming an amorphous phase (Deng et al., 2025). Arsenic may interfere with the polymerization step
<figure><img src="image_6.png" /><figcaption>**Fig. 3.** Iron isotope fractionation factors (δ<sup>57</sup>Fe) between solid and solution (noted Δ<sup>57</sup>Fe<sub>solid-liq</sub> in the text) against the proportion of iron in the form of pyrite in the solid, as determined by linear decomposition of Fe K-edge EXAFS spectra from Baya et al. (2021). Data are from Table 1.</figcaption></figure>

since it has been proposed to replace S in the polysulfide chains (Le Pape
et al., 2017).

6.5. Implications for the pyrite δ<sup>57</sup>Fe geological record

Very light pyrite δ<sup>57</sup>Fe values akin to those produced in the course of
the present experiments are commonly observed in Archean sedimen-
tary rocks (Rouxel et al., 2005). This contrasts with the more recent
geological record, which tends to show isotopically heavier pyrite
(Rouxel et al., 2005; Duan et al., 2010), closer to the continental crust
values (i.e., δ<sup>57</sup>Fe<sub>IRMM-14</sub> ~0.1 ‰; Poitrasson, 2006). Although recent
reviews on the topic show that the reasons for such difference remain
debated (Dauphas et al., 2017; Johnson et al., 2020), very light pyrite Fe
isotope signatures certainly reflect the anoxic and Fe-rich conditions
that prevailed in Archean oceans, similar to the anoxic conditions of the
pyrite precipitation experiments presented in this work. Yet, these sig-
natures are still largely attributed to microbial processes such as
dissimilatory iron reduction in the literature (e.g., Agangi et al., 2015;
Archer and Vance, 2006; Czaja et al., 2010; Galic et al., 2017; Johnson
et al., 2008; Marin-Carbonne et al., 2014; Nishizawa et al., 2010;
Yamaguchi et al., 2005; Yoshiya et al., 2015a; 2015b), although a
growing body of evidence demonstrates that sulfidation processes alone
produce isotopically light Fe(II) (McAnena et al., 2024) and pyrite
(Guilbaud et al., 2011a; Heard et al., 2020; Mansor and Fantle, 2019).
The present experiments further illustrate that it is possible to achieve
such light pyrite δ<sup>57</sup>Fe signatures through abiotic reaction involving the
polysulfide pyrite formation pathway. This work therefore supports
earlier studies (e.g., Guilbaud et al., 2011a) that also showed it is
possible to produce isotopically light pyrite through the other main
abiotic, H<sub>2</sub>S pathway. Moreover, although the isotope ranges obtained
from biotic and abiotic pathways largely overlap, our results suggest
that the kinetics of the isotope exchange between the aqueous and solid
precursors of pyrite likely controls the final iron isotope composition of
pyrite at low temperature. The molecular-level mechanisms of this effect
remain, however, incompletely understood and this calls for further
research.

Another interesting outcome of the present study is that although
some trace elements (Ni and As) may variably affect the speed of
acquisition of the pyrite final isotope composition (at the time-scale of a
few weeks; see Fig. 1), the final Δ<sup>57</sup>Fe partitioning between pyrite and
aqueous Fe remain the same. Chemical compositions or ratio of element
concentrations are dependent on the conditions of past pyrite

precipitation and of the composition of the source fluids and sediments. In contrast, it appears that the Fe isotope partitioning between pyrite and aqueous Fe seems independent from the pyrite production pathway, i.e., polysulfide (this work) and H₂S (Guilbaud et al., 2011a), and from the presence of minor amounts of aqueous Ni or As in the formation medium. If we surmise that the Fe isotope composition of the Archean oceans had an important end-member dominated by Fe from hydrothermal fluids having a δ<sup>57</sup>Fe close to the continental crust (Johnson et al., 2008), our experimental results could explain the most negative δ<sup>57</sup>Fe from Archean pyrite, contrasting from those coming from more recent sediments (Rouxel et al., 2005).

CRediT authorship contribution statement

**Franck Poitrasson:** Writing – original draft, Visualization, Validation, Methodology, Investigation, Formal analysis, Data curation, Conceptualization. **Pierre Le Pape:** Writing – review & editing, Visualization, Methodology, Funding acquisition, Formal analysis, Conceptualization. **Guillaume Morin:** Writing – original draft, Formal analysis, Conceptualization. **Romain Guilbaud:** Writing – review & editing, Methodology, Investigation, Formal analysis, Conceptualization. **Camille Baya:** Writing – review & editing, **Qianyu Deng:** Writing – review & editing, Conceptualization. **Jeremy Rabone:** Writing – review & editing, Investigation, Formal analysis, Conceptualization. **Julie Aufort:** Writing – review & editing, Visualization, Investigation, Formal analysis, Conceptualization. **Marc Blanchard:** Writing – original draft, Visualization, Validation, Project administration, Methodology, Investigation, Funding acquisition, Formal analysis, Conceptualization.

Declaration of competing interest

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

Acknowledgements

Eric Gardrat and Ghylaine Quitté are thanked for allowing access to, respectively, well maintained LAERO balance and IRAP clean laboratory during the two years period when the GET clean laboratory was down. Manuel Henry and then Laetitia Leroy are thanked for their support during difficult times with the GET clean lab. The GET-OMP ICP team
Jérôme Chmeleff, Camille Duquenoy and Aurélie Marquet are also
thanked for their support when issues occurred with the MC-ICP-MS and
for maintaining it in good working order. Three EPSL referees are
gratefully acknowledged for their insightful and detailed constructive
reviews. This research was funded by a CNRS/INSU-CNES PNP grant
(ISOPYR project) and an ANR PRC grant (PYRISOFE project, ANR-22-
CE49-0012), both coordinated by MB. DFT calculations were performed
using the HPC resources from CALMIP (grant 2025 - P1037).

Supplementary materials

Supplementary material associated with this article can be found, in
the online version, at doi:10.1016/j.epsl.2025.119729.

Data availability

Data will be made available on request.

References

Agangi, A., Hofmann, A., Rollion-Bard, C., Marin-Carbonne, J., Cavalazzi, B., Large, R., Meffre, S., 2015. Gold accumulation in the Archaean Witwatersrand Basin, South Africa - evidence from concentrically laminated pyrite. Earth-Sci. Rev. 140, 27-53.

Akerman, A., Poitrasson, F., Oliva, P., Seyler, P., Souza, V.D., Ponzevera, E., Santos, R.V., 2024. Impact of deforestation on the hydrogen, oxygen and iron isotope compositions of Amazonian streams. Chem. Geol. 668, 122296.

Amor, M., Busigny, V., Louvat, P., Gélabert, A., Cartigny, P., Durand-Dubief, M., Ona-Nguema, G., Alphandéry, E., Chebbi, I., Guyot, F., 2016. Mass-dependent and-independent signature of Fe isotopes in magnetotactic bacteria. Science 352, 705-708.

Archer, C., Vance, D., 2006. Coupled Fe and S isotope evidence for archean microbial Fe (III) and sulfate reduction. Geology 34 (3), 153-156.

Baya, C., Le Pape, P., Baptiste, B., Brest, J., Landrot, G., Elkaim, E., Noel, V., Blanchard, M., Ona-Nguema, G., Juillot, F., Morin, G., 2021. Influence of trace level As or Ni on pyrite formation kinetics at low temperature. Geochim. Cosmochim. Acta 300, 333-353.

Baya, C., Le Pape, P., Baptiste, B., Menguy, N., Delbes, L., Morand, M., Rouelle, M., Aubry, E., Ona-Nguema, G., Noël, Juillot, F., Morin, G., 2022. A methodological framework to study the behavior and kinetic influence of V, Mn, Co, Ni, Cu, Zn, As, Se and Mo during pyrite formation via the polysulfide pathway at ambient temperature. Chem. Geol. 613, 13-18.

Beard, B.L., Handler, R.M., Scherer, M.M., Wu, L.L., Czaja, A.D., Heimann, A., Johnson, C.M., 2010. Iron isotope fractionation between aqueous ferrous iron and goethite. Earth Planet Sci. Lett. 295, 241-250.

Beard, B.L., Johnson, C.M., Cox, L., Nealson, K.H., Aguilar, C., 1999. Iron isotope
biosignatures. Science 285, 1889-1892.

Becke, A.D., 1993. Density-functional thermochemistry. 3. The role of exact exchange. J. Chem. Phys. 98 (7), 5648-5652.

Bigeleisen, J., Mayer, B.G., 1947. Calculation of equilibrium constants for isotopic exchange reactions. J. Chem. Phys. 15, 261-267.

Blanchard, M., Balan, E., Schauble, E.A., 2017. Equilibrium fractionation of non-traditional isotopes: a molecular modeling perspective. In: Teng, F.Z., Watkins, J., Dauphas, N. (Eds.), Non-Traditional Stable Isotopes. Reviews in Mineralogy & Geochemistry. Mineralogical Soc Amer & Geochemical Soc, Chantilly, pp. 27-63.

Blanchard, M., Poitrasson, F., Méheut, M., Lazzeri, M., Mauri, F., Balan, E., 2009. Iron isotope fractionation between pyrite (FeS₂), hematite (Fe₂O₃) and siderite (FeCO₃): a first-principles density functional theory study. Geochim. Cosmochim. Acta 73, 6565-6578.

Bourdoiseau, J.A., Jeannin, M., Rémazeilles, C., Sabota, R., Refait, P., 2011. The transformation of mackinawite into greigite studied by Raman spectroscopy. J. Raman Spectrosc. 42 (3), 496-504.

Butler, I.B., Böttcher, M.E., Rickard, D., Oldroyd, A., 2004. Sulfur isotope partitioning during experimental formation of pyrite via the polysulfide and hydrogen sulfide pathways: implications for the interpretation of sedimentary and hydrothermal pyrite isotope records. Earth Planet, Sci. Lett. 228 (3-4), 495-509. https://doi.org/10.1016/j.epsl.2004.10.005.

Busigny, V., Dauphas, N., 2007. Tracing paleofluid circulations using iron isotopes: a study of hematite and goethite concretions from the Navajo Sandstone (Utah, USA). Earth Planet Sci. Lett. 254 (3-4), 272-287.

Croal, L.R., Johnson, C.M., Beard, B.L., Newman, D.K., 2004. Iron isotope fractionation by Fe(II)-oxidizing photoautotrophic bacteria. Geochim. Cosmochim. Acta 68, 1127-1242.

Crosby, H.A., Johnson, C.M., Roden, E.E., Beard, B.L., 2005. Coupled Fe(II)-Fe(III) electron and atom exchange as a mechanism for Fe isotope fractionation during dissimilatory iron oxide reduction. Env. Sci. Technol. 39, 6698-6704.

Criss, R.E., 1999. Principles of Stable Isotope Distribution. Oxford University Press, Oxford, p. 254.

Czaja, A.D., Johnson, C.M., Beard, B.L., Eigenbrode, J.L., Freeman, K.H., Yamaguchi, K. E., 2010. Iron and carbon isotope evidence for ecosystem and environmental

diversity in the ~2.7 to 2.5 Ga Hamersley Province. West Aust., Earth Planet. Sci. Lett. 292 (1-2), 170-180.

Dauphas, N., John, S.G., Rouxel, O., 2017. Iron isotope systematics. In: Teng, F.Z., Watkins, J., Dauphas, N. (Eds.), Non-Traditional Stable Isotopes. Reviews in Mineralogy & Geochemistry. Mineralogical Soc Amer & Geochemical Soc, Chantilly, pp. 415-510.

Deng, Q.Y., Le Pape, P., Aufort, J., Blanchard, M., Baptiste, B., Delbes, L., Baya, C., Juillot, F., Ona-Nguema, G., Menguy, N., Guigner, J.M., Proux, O., Duquenoy, C., Guilbaud, R., Poitrasson, F., Morin, G., 2025. HERFD-XAS evidence for an octahedrally coordinated CoSn polysulfide precursor as a probe for the mechanism of pyrite formation. Geochim. Cosmochim. Acta 401, 104-121.

Duan, Y., Severmann, S., Anbar, A.D., Lyons, T.W., Gordon, G.W., Sageman, B.B., 2010. Isotopic evidence for Fe cycling and repartitioning in ancient oxygen-deficient settings: examples from black shales of the mid-to-late devonian Appalachian basin. Earth Planet Sci. Lett 290 (3-4), 244-253.

Fabre, S., Nédélec, A., Poitrasson, F., Strauss, H., Thomazo, C., Nogueira, A., 2011. Iron and sulphur isotopes from the Carajas mining province (Para, Brazil): implications for the oxidation of the ocean and the atmosphere across the Archaean-Proterozoic transition. Chem., Geol. 289 (1-2), 124-139.

Frisch, M.J., et al., 2016. Gaussian 16. Gaussian, Inc., Wallingford CT. Revision B.01.

Fujii, T., Moynier, F., Blichtert-Toft, J., Albarede, F., 2014. Density functional theory estimation of isotope fractionation of Fe, Ni, Cu, and Zn among species relevant to geochemical and biological environments. Geochim. Cosmochim. Acta 140, 553-576.

Galic, A., Mason, P.R.D., Mogollón, J.M., Wolthers, M., Vroon, P.Z., Whitehouse, M.J., 2017. Pyrite in a sulfate-poor Paleozoarchean basin was derived predominantly from elemental sulfur: evidence from 3.2 Ga sediments in the Barberton Greenstone Belt. Kaapvaal Craton, Chem. Geol. 449, 135-146.

Gregory, D.D., Large, R.R., Halpin, J.A., Baturina, E.L., Lyons, T.W., Wu, S., Danyushesky, L., Sack, P.J., Chappaz, A., Maslennikov, V.V., Bull, S.W., 2015. Trace element content of sedimentary pyrite in black shales. Econ. Geol. 110 (6), 1389-1404.

Gregory, D.D., Lyons, T.W., Large, R.R., Jiang, G.Q., Stepanov, A.S., Diamond, C.W., Figueroa, M.C., Olin, P., 2017. Whole rock and discrete pyrite geochemistry as complementary tracers of ancient ocean chemistry: an example from the Neoproterozoic Doushantuo Formation. Geochim. Cosmochim. Acta 216, 201-220.

Guilbaud, R., Butler, I.B., Ellam, R.M., 2011a. Abiotic pyrite formation produces a large Fe isotope fractionation. Science 332 (6037), 1548-1551.

Guilbaud, R., Butler, I.B., Ellam, R.M., Rickard, D., Oldroyd, A., 2011b. Experimental determination of the equilibrium Fe isotope fractionation between 11b-aq(2+) and FeSm (mackinawite) at 25 and 2 degrees C. Geochim. Cosmochim. Acta 75 (10), 2721-2734.

Heard, A.W., Dauphas, N., Guilbaud, R., Rouxel, O.J., Butler, I.B., Nie, N.X., Bekker, A., 2020. Triple iron isotope constraints on the role of ocean iron sinks in early atmospheric oxygenation. Science 370 (6515), 446-449.

Icopini, G.A., Anbar, A.D., Ruebush, S.S., Tien, M., Brantley, S.L., 2004. Iron isotope fractionation during microbial reduction of iron: the importance of adsorption. Geology 32, 205-208.

Johnson, C.M., Beard, B., Weyer, S., 2020. Iron geochemistry: an isotopic perspective. Advances in Isotope Geochemistry. Springer Nature, Cham, p. 360.

Johnson, C.M., Beard, B.L., Klein, C., Beukes, N.J., Roden, E.E., 2008. Iron isotopes constrain biologic and abiologic processes in banded iron formation genesis. Geochim. Cosmochim. Acta 72 (1), 151-169.

Johnson, C.M., Beard, B.L., Welch, S., Croal, L., Newman, D., Nealson, K., 2005. Experimental constraints on Fe isotope fractionations during biogeochemical cycling of Fe. Geochim. Cosmochim. Acta 66, A371.

Johnson, C.M., Skulan, J.L., Beard, B.L., Sun, H., Nealson, K.H., Braterman, P.S., 2002. Isotopic fractionation between Fe(III) and Fe(II) in aqueous solutions. Earth Planet. Sci. Lett. 195, 141-153.

Kappler, A., Johnson, C., Crosby, H., Beard, B., Newman, D., 2010. Evidence for equilibrium iron isotope fractionation by nitrate-reducing iron (II)-oxidizing bacteria. Geochim. Cosmochim. Acta 74, 2826-2842.

Large, R.R., Mukherjee, I., Gregory, D.D., Steadman, J.A., Maslennikov, V.V., Meffre, S., 2017. Ocean and atmosphere geochemical proxies derived from trace elements in marine pyrite: implications for ore genesis in sedimentary basins. Econ. Geol. 112 (2), 423-450.

Le Pape, P., Baptiste, B., Radtke, G., Cabaret, D., Aufort, J., Brest, J., Baya, C., Elkaïm, E., Ona-Nguema, G., Juillot, F., Morin, G., 2025. Mackinawite transformation to greigite at room temperature under anoxic and acidic conditions: a corrosion pathway?. https://hal.science/hal-05286552.

Le Pape, P., Blanchard, M., Brest, J., Boulliard, J.C., Ikogou, M., Stetten, L., Wang, S., Landrot, G., Morin, G., 2017. Arsenic incorporation in pyrite at ambient temperature at both tetrahedral S¹ and octahedral Fe¹¹ sites: evidence from EXAFS-DFT analysis. Env., Sci, Technol 51 (1), 150-158.

Lee, C.T., Yang, W.T., Parr, R.G., 1988. Development of the coll-salvetti correlation-energy formula into a functional of the electron-density. Phys. Rev. B 37 (2), 785-789.

Luther, G.W., 1991. Pyrite synthesis via polysulfide compounds. Geochim. Cosmochim. Acta 55 (10), 2839-2849.

Mansor, M., Fantle, M.S., 2019. A novel framework for interpreting pyrite-based Fe isotope records of the past. Geochim. Cosmochim. Acta 253, 39-62.

Marin-Carbonne, J., Rollion-Bard, C., Bekker, A., Rouxel, O., Agangi, A., Cavalazzi, B., Wohlgemuth-Ueberwasser, C.C., Hofmann, A., McKeegan, K.D., 2014. Coupled Fe and S isotope variations in pyrite nodules from Archean shale. Earth Planet Sci. Lett. 392, 67-79.
McAnena, A., Severmann, S., Guilbaud, R., Poulton, S.W., 2024. Iron isotope
fractionation during sulfide-promoted reductive dissolution of iron (oxyhydr)oxide
minerals. Geochim. Cosmochim. Acta 369, 17–34.

Morin, G., Noel, V., Menguy, N., Brest, J., Baptiste, B., Tharaud, M., Ona-Nguema, G., Ikogou, M., Viollier, E., Juillot, F., 2017. Nickel accelerates pyrite nucleation at ambient temperature. Geochem. Perspect. Lett. 5, 6–11.

Nishizawa, M., Yamamoto, H., Ueno, Y., Tsuruoka, S., Shibuya, T., Sawaki, Y., Yamamoto, S., Kon, Y., Kitajima, K., Komiya, T., Maruyama, S., Hirata, T., 2010. Grain-scale iron isotopic distribution of pyrite from precambrian shallow marine carbonate revealed by a femtosecond laser ablation multicollector ICP-MS technique: possible proxy for the redox state of ancient seawater. Geochim. Cosmochim. Acta 74 (9), 2760–2778.

Nguyen, V.T., Osanai, Y., Adachi, T., Thai, P.K., Nakano, N., Ozaki, A., Kuwahara, Y., Kato, R., Makio, M., Kurosawa, K., 2018. Chemical speciation and bioavailability concentration of arsenic and heavy metals in sediment and soil cores in estuarine ecosystem, Vietnam. Microchem. J. 139, 268–277.

Noël, V., Marchand, C., Juillot, F., Ona-Nguema, G., Viollier, E., Marakovic, G., Olivi, L., Delbes, L., Gelebart, F., Morin, G., 2014. EXAFS analysis of iron cycling in mangrove sediments downstream of a lateritized ultramafic watershed (Vavouto Bay, New Caledonia). Geochim. Cosmochim. Acta 136, 211–228.

Noël, V., Morin, G., Juillot, F., Marchand, C., Brest, J., Bargar, J.R., Muñoz, M., Marakovic, G., Ardo, S., Brown Jr., G.E., 2015. Ni cycling in mangrove sediments from New Caledonia. Geochim. Cosmochim. Acta 169, 82–98.

Noël, V., Juillot, F., Morin, G., Marchand, C., Ona-Nguema, G., Viollier, E., Prevot, F., Dublet, G., Maillot, F., Delbes, L., Marakovic, G., Bargar, J.R., Brown Jr., G.E., 2017. Oxidation of Ni-rich mangrove sediments after isolation from the sea (Dumbea Bay, New Caledonia): fe and Ni behavior and environmental implications. ACS Earth Space Chem. https://doi.org/10.1021/acsearthspacechem.7b00005.

Platzner, I.T., 1997. Modern Isotope Ratio Mass Spectrometry. John Wiley & sons,
Chichester, p. 514.

Poitrasson, F., 2006. On the iron isotope homogeneity level of the continental crust. Chem. Geol. 235, 195–200.

Poitrasson, F., Freydier, R., 2005. Heavy iron isotope composition of granites determined by high resolution MC-ICP-MS. Chem. Geol. 222, 132–147.

Poitrasson, F., Halliday, A.N., Lee, D.C., Levasseur, S., Teutsch, N., 2004. Iron isotope
differences between Earth, Moon, Mars and Vesta as possible records of contrasted
accretion mechanisms. Earth Planet Sci. Lett. 223, 253–266.

Pokrovski, G.S., Blanchard, M., Saunier, G., Poitrasson, F., 2021. Mechanisms and rates of pyrite formation from hydrothermal fluid revealed by iron isotopes. Geochim. Cosmochim. Acta 304, 281–304.

Poulton, S.W., Canfield, D.E., 2011. Ferruginous conditions: a dominant feature of the
ocean through Earth's history. Elements 7 (2), 107–112.

Rickard, D., 1997. Kinetics of pyrite formation by the H2S oxidation of iron (II) monosulfide in aqueous solutions between 25 and 125 degrees C: the rate equation. Geochim. Cosmochim. Acta 61 (1), 115–134.

Rickard, D., 2019. How long does it take a pyrite framboid to form? Earth Planet. Sci. Lett. 513, 64–68. https://doi.org/10.1016/j.epsl.2019.02.019.

Rickard, D., Luther, G.W., 2007. Chemistry of iron sulfides. Chem. Rev. 107 (2), 514–562.

Rolison, J.M., Stirling, C.H., Middag, R., Gault-Ringold, M., George, E., Rijkenberg, M.J., 2018. Iron isotope fractionation during pyrite formation in a sulfidic precambrian ocean analogue. Earth Planet Sci. Lett. 488, 1–13.

Rouxel, O.J., Bekker, A., Edwards, K.J., 2005. Iron isotope constraints on the archean and paleoproterozoic ocean redox state. Science 307, 1088–1091.

Schauble, E.A., 2004. Applying stable isotope fractionation theory to new systems. In: Geochemistry of Non-Traditional Stable Isotopes: Reviews in Mineralogy & Geochemistry, 55, pp. 65–111.

Strauss, H., 1997. The isotopic composition of sedimentary sulfur through time. Palaeogeogr. Palaeoclimatol. Palaeoecol. 132 (1-4), 97–118.

Wang, Y., Le Pape, P., Morin, G., Asta, M., King, G., Bartova, B., Suvorova, E., Frutschi, M., Ikogou, M., Pham, V.H.C., Vo, P.L., Herman, F., Charlet, L., Bernier-Latmani, R., 2018. Arsenic speciation in Mekong Delta sediments depends on their depositional environment. Environ. Sci. Technol. 52, 3431–3439. https://doi.org/10.1021/acs.est.7b05177.

Welch, S., Beard, B.L., Johnson, C.M., Braterman, P.S., 2003. Kinetic and equilibrium Fe isotope fractionation between aqueous Fe(II) and Fe(III). Geochim. Cosmochim. Acta 67, 4231–4250.

Wu, B., Amelung, W., Xing, Y., Bol, R., Berns, A.E., 2019. Iron cycling and isotope fractionation in terrestrial ecosystems. Earth-Sci. Rev. 190, 323–352.

Wu, L.L., Druschel, G., Findlay, A., Beard, B.L., Johnson, C.M., 2012. Experimental determination of iron isotope fractionations among Fe-aq(2+)-FeSaq-Mackinawite at low temperatures: implications for the rock record. Geochim. Cosmochim. Acta 89, 46–61.

Yamaguchi, K.E., Johnson, C.M., Beard, B.L., Ohmoto, H., 2005. Biogeochemical cycling of iron in the Archean-paleoproterozoic Earth: constraints from iron isotope variations in sedimentary rocks from the Kaapvaal and Pilbara Cratons. Chem. Geol. 218, 135–169.

Yoshiya, K., Sawaki, Y., Hirata, T., Maruyama, S., Komiya, T., 2015a. In-situ iron isotope analysis of pyrites in ~3.7 Ga sedimentary protoliths from the Isua supracrustal belt, southern West Greenland. Chem. Geol. 401, 126–139.

Yoshiya, K., Sawaki, Y., Shibuya, T., Yamamoto, S., Komiya, T., Hirata, T., Maruyama, S., 2015b. In-situ Iron Isotope Analyses of Pyrites from 3.5 to 3.2 Ga sedimentary Rocks of the Barberton Greenstone Belt, 403. Kaapvaal Craton. Chemical Geology, pp. 58–73.