# Intense biogeochemical iron cycling revealed in Neoarchean micropyrites from stromatolites

Marie-Noëlle Decraene, Johanna Marin-Carbonne, Christophe Thomazo, Nicolas Olivier, Pascal Philippot, Harald Strauss, Etienne Deloule

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Marie-Noëlle Decraene, Johanna Marin-Carbonne, Christophe Thomazo, Nicolas Olivier, Pascal Philippot, et al.. Intense biogeochemical iron cycling revealed in Neoarchean micropyrites from stromatolites. Geochimica et Cosmochimica Acta, 2021, 312, pp.299-320. 10.1016/j.gca.2021.07.020. hal-03329228

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Distributed under a Creative Commons Attribution 4.0 International License
Intense biogeochemical iron cycling revealed in
Neoarchean micropyrites from stromatolites

Marie-Noëlle Decraene<sup>a,b,*</sup>, Johanna Marin-Carbonne<sup>a</sup>, Christophe Thomazo<sup>c,d</sup>,
Nicolas Olivier<sup>e</sup>, Pascal Philippot<sup>f,g</sup>, Harald Strauss<sup>h</sup>, Etienne Deloule<sup>b</sup>

Received 22 March 2021; accepted in revised form 19 July 2021; available online xxxx

**Abstract**

Iron isotope compositions of sedimentary pyrites (FeS₂) are used to constrain the redox evolution of the Precambrian ocean and early Fe-based metabolisms such as Dissimilatory Iron Reduction (DIR). Sedimentary pyrites can record biotic and abiotic iron reduction, which have similar ranges of Fe isotopic fractionation, as well as post-depositional histories and metamorphic overprints that can modify Fe isotope compositions. However, some exceptionally well-preserved sedimentary records, such as the stromatolite-bearing Tumbiana Formation (ca. 2.7 Ga, Western Australia) have been proven to retain primary information on Early Neoarchean microbial ecosystems and associated metabolic pathways. Here, we present *in situ* Fe isotope measurements of micropyrites included in four stromatolites from the Tumbiana Formation in order to assess iron respiration metabolism using Fe isotope signatures. A set of 142 micropyrites has been analyzed in three lamina types, *i.e.* micritic, organic-rich and fenestral laminae, by Secondary Ion Mass Spectrometry (SIMS), using a Hyperion radio-frequency plasma source. The diversity of laminae is attributed to specific depositional environments, leading to the formation of Type 1 (micritic laminae) and Type 2 (organic-rich laminae) and early diagenetic effects (Type 3, fenestral laminae). Type 1 and 2 laminae preserved comparable δ⁵⁶Fe ranges, respectively from -1.76‰ to +4.15‰ and from -1.54‰ to +4.44‰. Type 3 laminae recorded a similar range, although slightly more negative δ⁵⁶Fe values between -2.20‰ and +2.65‰. Globally, our data show a large range of δ⁵⁶Fe values, from -2.20‰ to +4.44‰, with a unimodal distribution that differs from the bimodal distribution previously reported in the Tumbiana stromatolites. Such a large range and unimodal distribution cannot be explained by a unique process (*e.g.*, biotic/abiotic Fe reduction or pyrite formation only controlled by the precipitation rate). It rather could reflect a two-step iron cycling process in the sediment pore water including *i*) partial Fe oxidation forming Fe (OH)₃ with positive δ⁵⁶Fe values followed by *ii*) partial, possibly microbially induced, Fe reduction leading to Fe²⁺ availability for pyrite formation by sulfate reducers carrying both negative δ⁵⁶Fe and δ³⁴S signatures. In this model, the buildup and subsequent reduction through time of a residual Fe(OH)₃ reservoir arising from the activity of methanotrophs, can explain the strongly positive δ⁵⁶Fe<sub>Fe(OH)₃</sub> values up to 4‰. These results indicate that Archean microbial mats have been the site of the

Abbreviations: DIR, Dissimilatory Iron Reduction; MSR, Microbial Sulfate Reduction; AOM, Anaerobic Oxidation of Methane; SEM, Scanning Electron Microscopy; SIMS, Secondary Ion Mass Spectrometry

* Corresponding author at: Institut des Sciences de la Terre, Université de Lausanne, Lausanne, Suisse.

E-mail address: marie-noelle.depraene@unil.ch (M.-N. Decraene).

<footnote><sup>a</sup> Institut des Sciences de la Terre, Université de Lausanne, Lausanne, Switzerland</footnote>

<footnote><sup>b</sup> Université de Lorraine, CNRS, CRPG, 54500 Vandoeuvre les Nancy, France</footnote>

<footnote><sup>c</sup> UMR CNRS 5561 Biogéosciences, Université de Bourgogne Franche Comté, Dijon, France</footnote>

<footnote><sup>d</sup> Institut Universitaire de France</footnote>

<footnote><sup>e</sup> Université Clermont Auvergne, CNRS, IRD, Laboratoire Magmas et Volcans, F-63000 Clermont-Ferrand, France</footnote>

<footnote><sup>f</sup> Géosciences Montpellier, CNRS UMR 5243, Université de Montpellier, 34000 Montpellier, France</footnote>

<footnote><sup>g</sup> Institute of Astronomy, Geophysics and Atmospheric Sciences, University of São Paulo, Rua do Matão 1226, Cidade Universitária, 05508-090 São Paulo, Brazil</footnote>

<footnote><sup>h</sup> Westfälische Wilhelms-Universität, Institut für Geologie und Paläontologie, Corrensstr, 24, 48149 Münster, Germany</footnote>
interaction of several closely linked biogeochemical cycles involving Fe, S and C.

© 2021 The Author(s). Published by Elsevier Ltd. This is an open access article under the CC BY license (http://creativecommons.org/licenses/by/4.0/).

*Keywords:* Early life; Stromatolites; Tumbiana Formation; Iron isotopes; Ion microprobe; Iron respiration

# 1. INTRODUCTION

Stromatolites are laminated organo-sedimentary struc-
tures, formed by microbial activities in response to environ-
mental and sediment dynamics, and are considered to be
the most undisputable evidence of early life on Earth
(Allwood et al., 2006; Awramik, 2006; Schopf et al.,
2007). The 2.724 Ga old Tumbiana Formation, Western
Australia, contains some of the best preserved and most
diverse Archean stromatolitic carbonate structures (Buick,
1992; Sakurai et al., 2005; Awramik and Buchheim, 2009;
Coffey et al., 2013), including domal, conical, ridged and
tufted mm- to m-scale morphologies (Buick, 1992;
Flannery and Walter, 2012). A variety of isotopic proxies
such as δ¹³C org, δ⁹⁸/⁹⁵Mo and δ⁸²/⁷⁸Se have been interpreted
to reflect the buildup of oxidized dissolved species that
could result from the activity of oxygenic photosynthetizers
(Coffey et al., 2013; Stüeken et al., 2015a; Williford et al.,
2016; Stüeken et al., 2017). The record of one of the largest
organic carbon isotope excursions on Earth, the Fortescue
Excursion, with δ¹³C org values as low as −60‰ (Hayes,
1994; Hinrichs, 2002) has been interpreted to represent
either the incorporation of ¹²C in organic matter during
oxidation of methane through methanotrophy
(Schidlowski, 1982; Hayes, 1994; Eigenbrode and
Freeman, 2006; Thomazo et al., 2009; Lepot et al., 2019)
or linked to acetogenesis using Acetyl-coA metabolisms
(Slotznick and Fischer, 2016), although this latter pathway
has been considered unlikely in the Tumbiana stromatolites
by other workers (Lepot et al., 2019). Other biological path-
ways reported in these stromatolites include Microbial Sul-
fate Reduction (Thomazo et al., 2009; Marin-Carbonne
et al., 2018), Dissimilatory Iron Reduction (Yoshiya
et al., 2012), arsenic cycling involving both As(III) oxida-
tion and As(V) reduction (Sforna et al., 2014),
microbially-induced ammonia oxidation leading to one of
the most extreme isotopic ¹⁵N enrichment (δ¹⁵N up to
+50‰) ever recorded on Earth (Thomazo et al., 2011;
Stüeken et al., 2015b), and microbially-derived nano-
aragonite precipitates associated with organic globules con-
taining aromatic, aliphatic and carboxyl groups (Lepot
et al., 2008; Lepot et al., 2009).

The presence of sulfate and Fe-oxides at the time of
deposition and/or early diagenesis may have fueled the
reaction of anaerobic oxidation of methane (AOM) in the
Tumbiana Formation (Marin-Carbonne et al., 2018;
Lepot et al., 2019). Sulfato-reducing metabolizers reduce
sulfate into dissolved S species (e.g., H₂S, HS⁻, S₀), a pro-
cess accompanied by a large fractionation of sulfur isotopes
between sulfide and sulfate of ~70‰ (e.g., Johnston, 2011;
Sim et al., 2011). This metabolic effect has been identified
in the Tumbiana Formation by bulk rock (Thomazo
et al., 2009) and *in situ* (Marin-Carbonne et al., 2018) S-

isotope analyses, although the extent of the measured iso-
topic fractionation differed from one technique to another.
The reduced δ³⁴S range of ~8‰ measured by bulk analyses
has been interpreted as MSR completed under low sulfate
concentrations, as expected during the Archean (Crowe
et al., 2014). In contrast, the larger δ³⁴S range of ~84‰
recorded at the lamina scale confirms the activity of MSR
associated with locally high sulfate concentrations
(Marin-Carbonne et al., 2018). DIR reduces Fe-oxides into
dissolved Fe²⁺ and is associated with Fe isotopic fractiona-
tion of -2.9‰ between ferrous iron and Fe-oxides (Crosby
et al., 2005; Crosby et al., 2007). Although a bulk rock
study reported a limited δ⁵⁶Fe range in the Tumbiana For-
mation interpreted to reflect a limited Fe cycle (Czaja et al.,
2010), *in situ* laser ablation analyses evidenced a larger Fe
isotopic range of both negative and positive δ⁵⁶Fe values
between -4.2‰ and +2.1‰, interpreted as reflecting an
active Fe cycling (Nishizawa et al., 2010; Yoshiya et al.,
2012). Ideally, coupled analyses of Fe and S isotopes of pyr-
ites can provide informations on the intimate link between
Fe and S cycling in microbial mats (Archer and Vance,
2006; Marin-Carbonne et al., 2014). While δ³⁴S analyses
at the scale of individual microscopic pyrites have been per-
formed in stromatolitic laminae of the Tumbiana Forma-
tion (Marin-Carbonne et al., 2018), this approach was not
available until recently for Fe isotopes.

Here we use the new procedure developed for micromet-
ric scale analyses of Fe isotopes by SIMS (Decraene et al.,
2021) to the same micropyrite-bearing laminae present in
Tumbiana stromatolites previously investigated by Marin-
Carbonne et al. (2018). Results show that a large range of
positive and negative δ⁵⁶Fe recorded in these micropyrites
is best attributed to an intense oxidative and reductive Fe
cycling, parsimoniously interpreted as resulting from DIR
activity, coupled with methanotrophy and MSR.

# 2. MATERIAL AND METHODS

## 2.1. Geological setting and sample collection for SIMS analyses and Fe speciation

The Tumbiana Formation is a sedimentary succession of
about 200 m thick, 680 km wide, with an WNW-ESE exten-
sion through the Pilbara Craton (Thorne and Trendall,
2001). The Tumbiana Formation is part of the Neoarchean
Fortescue Group (Hamersley Basin, Western Australia)
and consists of alternating low-metamorphic grade volcanic
and sedimentary rocks deposited during a period of conti-
nental extension and rifting (Blake and Barley, 1992;
Blake, 1993). It has been subdivided in two main members,
the Mingah Member, composed of volcanic tuffs, accre-
tionary lapillis, conglomerates and sandstones, and the
Meentheena Member, composed of several horizons of
stromatolitic limestones, sandstones and mudstones (Thorne and Trendall, 2001; Sakurai et al., 2005; Flament et al., 2011; Flannery et al., 2014). U-Pb dating on zircons and volcanic rocks yielded depositional ages between 2724 ± 5 Ma and 2715 ± 6 Ma (Arndt et al., 1991; Blake et al., 2004). The Tumbiana Formation overprinted low-grade metamorphism < 300 °C, allowing the preservation of cell-like organic globules in micrites (Lepot et al., 2008). The depositional environment has been interpreted either as a giant lake system (Buick, 1992; Bolhar and Van Kranendonk, 2007; Awramik and Buchheim, 2009; Coffey et al., 2013; Stüeken et al., 2015b) or a shallow marine setting (Thorne and Trendall, 2001; Sakurai et al., 2005).

Our samples are from the PDP1 drill core of the Pilbara Drilling Project, which intercepts the base of the Maddina basalt, the Meentheena Member and the top of the Mingah Member (Philippot et al., 2009). This drill core has been extensively studied over the past decade for its mineralogy and S, C, N isotopic compositions and represents a unique time capsule in the search for Earth's earliest metabolic activities (Thomazo et al., 2007; Lepot et al., 2008; Lepot et al., 2009; Philippot et al., 2009; Thomazo et al., 2011; Marin-Carbonne et al., 2018; Lepot et al., 2019). This SIMS study focuses on four stromatolites from the Meentheena Member, sampled at 67.8 m (GIS67.8), 68.1 m (GIS68.1), 68.9 m (GIS68.9) and 70.8 m (GIS70.8) depth in stromatolite-rich intervals of the PDP1 drill core (Fig. 1). Samples GIS67.8 and GIS68.9 have been previously investigated for their C<sub>organic</sub> and S<sub>pyrite</sub> isotopes compositions of carbonaceous matter and nano-scale sulfides at high spatial resolution (Marin-Carbonne et al., 2018; Lepot et al., 2019). A detailed sample description is further provided in Section 3.1.

## 2.2. Optical and SEM observations

The samples were first described by optical microscopy. Scanning Electron Microscopes (SEM) were used at CRPG Nancy (JEOL JSM 6510) and at ISTE Lausanne (Tescan Mira LMU) to produce backscattered electron images at a working distance of 21 mm. Analysis were performed with a 1.5nA primary beam at 20 kV. Energy dispersive X ray spectrometry was used for semi quantitative punctual analysis and mapping to determine elemental composition of the sulfides and to select pyrite rich areas for subsequent SIMS analysis.

## 2.3. Iron speciation

Identification of Fe-mineral species follows the sequential iron extraction procedure proposed by Poulton and Canfield (2005). Four different mineral phases were extracted: (1) carbonate-associated Fe (Fe<sub>carb</sub>) present in calcite, siderite and ankerite; (2) Fe contained in oxides, hydroxides and oxyhydroxides (Fe<sub>ox</sub>) including hematite, goethite, akaganeite, lepidocrocite and ferrihydrite; (3) magnetite (Fe<sub>mag</sub>) and (4) Fe bound to poorly reactive sheet silicates (Fe<sub>PRS</sub>) including biotite, chlorite and glauconite. Furthermore, in order to evaluate total iron contents

<figure><img src="image_2.png" /><figcaption>Fig. 1. Simplified log of the PDP1 drill core modified after Thomazo et al. (2009). Black arrows represent the location of the studied samples.</figcaption></figure>

(Fe<sub>T</sub>) the technique developed by Aller et al. (1986); see Poulton and Canfield (2005) was applied. Around 200 mg of each sample were ashed (at 450 °C for 8 h) in ceramic crucibles and reacted with 10 mL 6 N HCl. This extraction was conducted at 80 °C for 24 h. Each extracted species of iron described above was reacted with phenantroline and the concentration of the resulting Fe<sup>2+</sup>-phenantroline complex was measured using a Genesys spectral photometer Series 10 at 515 nm, at the Institut für Geologie und Paläontologie, Westfälische Wilhelms-Universität Münster, Germany, following the method proposed in Reuschel et al. (2012) and described in detail in Sauvage et al. (2013). The concentration of S associated with pyrite (S<sub>py</sub>) was evaluated using gravimetric quantification after chromium reducible sulfur extraction (Canfield et al., 1986). Iron associated with pyrite (Fe<sub>py</sub>) was calculated from the S<sub>py</sub> assuming a stoichiometric ratio. The total iron concentrations (Fe<sub>T</sub>) measured during the course of this study using the sequential extraction method were also compared to previous ICPMS measurements reported in Thomazo et al. (2010). The intercalibration between both methods displays a good correlation (R<sup>2</sup> = 0.97, Table 1, Electronic Annex (EA); Fig. EA-1).
<table><caption>Table 1
Iron speciation data for stromatolite samples, from 42.7 m to 85.7 m depth of the drill core PDPI.</caption><thead><tr><th>Sample</th><th>Lithology</th><th>Al (ppm)</th><th>Fe<sub>T</sub> (ppm)</th><th>Fe<sub>ICPMS</sub> (ppm)</th><th>Fe<sub>py</sub> (ppm)</th><th>Fe<sub>carb</sub> (ppm)</th><th>Fe<sub>ox</sub> (ppm)</th><th>Fe<sub>mag</sub> (ppm)</th><th>Fe<sub>PRS</sub> (ppm)</th><th>Fe<sub>U</sub> + Fe<sub>PRS</sub> (ppm)</th><th>Fe<sub>HR</sub> (ppm)</th><th>Fe<sub>HR</sub>/ Fe<sub>T</sub></th></tr></thead><tbody><tr><td>GIS42.7</td><td>Stromatolite</td><td>3963.11</td><td>10,260.00</td><td>11,354</td><td>348</td><td>3640</td><td>178</td><td>522</td><td>1100</td><td>5572.00</td><td>4688</td><td>0.46</td><td>0.07</td></tr><tr><td>GIS68.0</td><td>Stromatolite</td><td>14,257.09</td><td>21,443.36</td><td>21,707</td><td>958</td><td>3887</td><td>90</td><td>316</td><td>1194</td><td>16,192.36</td><td>5251</td><td>0.24</td><td>0.18</td></tr><tr><td>GIS68.2</td><td>Stromatolite</td><td>22,435.73</td><td>n.a</td><td>25,459</td><td>1655</td><td>2530</td><td>193</td><td>464</td><td>4109</td><td>n.a</td><td>4842</td><td>0.19</td><td>0.34</td></tr><tr><td>GIS68.9</td><td>Stromatolite</td><td>55,316.89</td><td>n.a</td><td>42,581</td><td>3745</td><td>n.a</td><td>n.a</td><td>n.a</td><td>n.a</td><td>n.a</td><td>n.a</td><td>n.a</td><td>n.a</td></tr><tr><td>GIS69.2</td><td>Stromatolite</td><td>17,074.17</td><td>26,845.63</td><td>27,440</td><td>3745</td><td>4207</td><td>120</td><td>826</td><td>1104</td><td>17,947.63</td><td>8898</td><td>0.33</td><td>0.42</td></tr><tr><td>GIS70.4</td><td>Stromatolite</td><td>15,332.43</td><td>18,687.26</td><td>18,949</td><td>1742</td><td>3967</td><td>41</td><td>312</td><td>755</td><td>12,625.26</td><td>6062</td><td>0.32</td><td>0.29</td></tr><tr><td>GIS70.8</td><td>Stromatolite</td><td>18,578.64</td><td>22,503.36</td><td>22,960</td><td>87</td><td>3642</td><td>204</td><td>813</td><td>1306</td><td>17,757.36</td><td>4746</td><td>0.21</td><td>0.02</td></tr><tr><td>GIS72.1</td><td>Stromatolite</td><td>8976.31</td><td>12,697.44</td><td>12,446</td><td>523</td><td>2794</td><td>12</td><td>149</td><td>284</td><td>9219.44</td><td>3478</td><td>0.27</td><td>0.15</td></tr><tr><td>GIS77.1</td><td>Stromatolite</td><td>4483.11</td><td>n.a</td><td>12,628</td><td>1045</td><td>3860</td><td>324</td><td>n.a</td><td>509</td><td>n.a</td><td>n.a</td><td>n.a</td><td>n.a</td></tr><tr><td>GIS77.2</td><td>Stromatolite</td><td>51,989.90</td><td>n.a</td><td>57,246</td><td>1391</td><td>325</td><td>571</td><td>2741</td><td>571</td><td>n.a</td><td>3245</td><td>0.06</td><td>n.a</td></tr><tr><td>GIS78.1</td><td>Stromatolite</td><td>6275.34</td><td>10,370.15</td><td>15,330</td><td>1132</td><td>4018</td><td>114</td><td>171</td><td>510</td><td>4935.15</td><td>5435</td><td>0.52</td><td>0.21</td></tr><tr><td>GIS78.4</td><td>Stromatolite</td><td>7123.50</td><td>17,622.55</td><td>16,177</td><td>697</td><td>3649</td><td>82</td><td>371</td><td>659</td><td>12,823.55</td><td>4799</td><td>0.27</td><td>0.15</td></tr><tr><td>GIS83.2</td><td>Stromatolite</td><td>21,143.30</td><td>46,724.23</td><td>43,029</td><td>958</td><td>4765</td><td>188</td><td>1605</td><td>2277</td><td>39,208.23</td><td>7516</td><td>0.16</td><td>0.13</td></tr><tr><td>GIS85.7</td><td>Stromatolite</td><td>n.a</td><td>n.a</td><td>n.a</td><td>2003</td><td>4947</td><td>277</td><td>1620</td><td>2959</td><td>25,190.77</td><td>8847</td><td>0.26</td><td>0.23</td></tr></tbody></table>

Because of their high reactivity toward hydrogen sulfide (Canfield et al., 1992; Poulton and Canfield, 2005), the sum of Fe<sub>carb</sub>, Fe<sub>ox</sub>, Fe<sub>mag</sub> and Fe<sub>py</sub> are referred to as ‘Highly Reactive’ Fe fraction (Fe<sub>HR</sub>) of the total iron. Fe<sub>HR</sub>/Fe<sub>T</sub> ratio is commonly used to scale anoxic conditions when values are above a threshold of 0.38 (Raiswell and Canfield, 1998). Fe<sub>HR</sub> enrichment, in excess toward this maximal detrital background, indicates an external source of reactive Fe decoupled from the siliciclastic flux. Euxinic conditions can also be recognized when both Fe<sub>HR</sub>/Fe<sub>T</sub> and Fe<sub>py</sub>/Fe<sub>HR</sub> exceed thresholds of 0.38 and 0.80, respectively. However, in lacustrine environment, bias due to high sedimentation rates makes threshold based on Fe<sub>HR</sub>/Fe<sub>T</sub> questionable (Lyons and Severmann, 2006) and Tumbiana Fe speciation data should be interpreted with caution.

## 2.4. *In situ* Fe isotope analyses of micropyrites by SIMS

142 micropyrites were analyzed over three SIMS sessions (April 2018, July 2018 and October 2018). One reference standard grain was included in each sample mount, and the mounts were carbon-coated before analysis. *In situ* iron isotope analyses were performed with the CAMECA imS 1280HR2 equipped with the new Hyperion radio-frequency source at CRPG-IPNT in Nancy (France). Using the procedure detailed in Decraene et al. (2021), this source allows measurement of micrometric targets (less than 20 microns) with a current density 10 times greater than that of a standard duoplasmatron source (Liu et al., 2018). A 3nA Gaussian <sup>16</sup>O<sup>-</sup> primary beam was focused into a 2.5 to 3 µm spot. Typical <sup>56</sup>Fe intensities ranged between 1.2x10<sup>7</sup> and 8.2x10<sup>7</sup> counts per second. The mass resolution power was set at ~6800 to resolve interferences on <sup>54</sup>Fe<sup>+</sup> (<sup>53</sup>CrH<sup>+</sup>) and on <sup>56</sup>Fe<sup>+</sup> (<sup>55</sup>MnH<sup>+</sup>). The measurement of <sup>52</sup>Cr allow to monitor the <sup>54</sup>Cr<sup>+</sup> isobaric interference on <sup>54</sup>Fe<sup>+</sup> according to the method described in Marin-Carbonne et al. (2011). Simultaneous measurement of <sup>52</sup>Cr, <sup>54</sup>Fe, <sup>56</sup>Fe and <sup>57</sup>Fe was performed in multicollection mode with three off-axis Faraday cups and one electron multiplier for <sup>52</sup>Cr. Finally, iron isotope compositions are reported as permil variations of <sup>56</sup>Fe/<sup>54</sup>Fe ratios measured in the samples normalized to that of the international reference material IRMM-014 as the following:

$$ \delta^{56}\text{Fe} = \left[ \frac{\left( \frac{^{56}\text{Fe}}{^{34}\text{Fe}} \right)_{\text{sample}}}{\left( \frac{^{56}\text{Fe}}{^{34}\text{Fe}} \right)_{\text{IRMM-014}}} - 1 \right] * 1000 $$

Before running analyses, <sup>32</sup>S<sup>+</sup> isotope images were collected via electron multiplier, in order to record the exact location of the micropyrites (Electronic Annex (EA); Fig. EA-2). Afterward, a 90 s presputtering time was applied, followed by data acquisition in multicollection mode for 300 seconds. Backgrounds of the detectors were measured during presputtering and were subtracted from each analysis. Balmat and Spain pyrites (respectively δ<sup>56</sup>-Fe<sub>true</sub> = -0.40 ± 0.01‰ Whitehouse and Fedo, 2007 and δ<sup>56</sup>Fe<sub>true</sub> = +0.52 ± 0.03‰ Decraene et al., 2021) were used as standards to correct the instrumental mass fractionation
<figure><img src="image_3.png" /><figcaption>Fig. 2. Optical microscope large views of the four analyzed samples. (a) GIS68.9; (b) Schematic view of the four different laminae (Type 1: Micritic lamina; Type 2: Organic-rich lamina; Type 3: Fenestral lamina and Type 4: coarse-grained lamina) in GIS68.9; (c) GIS70.8; (d) GIS67.8 and (e) GIS68.1. The different types of laminae are indicated by white arrows. The mineralogy of the different laminae is detailed in Fig. 3. Arrows on the top right give the orientation of the samples. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>

$\alpha$=(($^{56}$Fe/$^{54}$Fe)$_{meas}$/($^{56}$Fe/$^{54}$Fe)$_{true}$, defined as the ratio between iron isotopic ratio measured by SIMS and the true ratio determined by MC-ICP-MS (for Multi Collector Inductively Coupled Plasma Mass Spectrometry). The reproducibility was ± 0.24‰ for Balmat pyrite (2SD, standard deviation, April 2018 session; Electronic Annex (EA); Fig. EA-3) and ± 0.28‰ for Spain pyrite (July 2018 session; Electronic Annex (EA); Fig. EA-4; Table EA-1).

### 3. RESULTS

#### 3.1. Petrographic investigations

The four studied samples are part of a well-developed stromatolite facies and belong to the stromatolite-rich interval described between ~66 to ~71.5 m depth of the

PDP1 drill core (Fig. 1, Philippot et al., 2009). The four investigated samples (GIS67.8, GIS68.1, GIS68.9, GIS70.8) show an internal variability in the stromatolite fabric, leading to the recognition of four main types of laminae (Fig. 2):

* Type 1: Micritic laminae, where a cement of micrite represents the major mineral phase, associated with scarce automorphic chlorites and micropyrites. Few micrometric (10 µm to 20 µm) detrital zircons were also observed. This stromatolitic fabric is characterized by well-marked laminated domal structures.

* Type 2: Organic-rich laminae, which are composed of nano to micropyrites and automorphic chlorites. Accessory mineral phases such as sphalerite, zircon and analate are rare. This type of lamina is often laterally
discontinuous and is observed between different types of
laminae.

* Type 3: Fenestral laminae, which are composed of micritic and sparitic cements, and micropyrites. This laminoid-fenestral fabric is characterized by spar-filled cavities that are arranged along the lamination. These cavities can be either by irregular or elongated horizontal fenestrae, which are embedded in a micritic cement.

* Type 4: coarse-grained laminae, composed of various sedimentary grains such as peloids and ooids, with a sparse and non-continuous internal lamination. These do not contain any pyrite and therefore are not considered in the following geochemical ($\delta^{56}\text{Fe}$) comparisons.

Type 1 and 3 laminae have a pluri-millimeter to centimeter thickness whereas Type 2 are sub-millimetric to millimetric. Microcrystalline quartz cement is identified in all samples, independently of the type of laminae, which indicates a post-depositional silicification process (Lepot et al., 2009; Philippot et al., 2009). Pyrites are identified in Type 1, 2 and 3 laminae (Fig. 3) in various proportions (more abundant in organic-rich laminae relative to the others) and exhibit euhedral morphologies, as single grains or aggregates. The size of the pyrites ranges from ~100 nm-20 µm, except for GIS68.1, which present some larger pyrites up to 60 µm. Pyrites associated with organic-rich laminae (Type 2) are mainly between 5 to 10 µm in size and in average ~10 µm in micritic and fenestral laminae (Types 1 and 3). The measured micropyrites have similar size ranges (5 to 15 µm) whatever the lamina type, but the smallest pyrites (<5µm to hundreds nm) are preferentially located in organic-rich (Type 2) laminae (Fig. 3).

## 3.2. Iron species

Iron speciation was used to evaluate the relative contribution of carbonate, oxide, silicate and sulfide as iron-bearing minerals compared to the total iron content. Iron speciation has been measured in 41 samples from 42.7 m to 90.4 m depth of the PDP1 drill core. Total iron (FeT) content in three different lithologies (stromatolites, mudstones and siltstones) vary between 0.64 and 5.69wt.% with a mean value of 2.99wt.%. Results for Fe speciation in the stromatolite facies are shown in Table 1.

$Fe_T$ in stromatolites range between 1.03 and 4.67wt.% with a mean value of 2.22wt.%. Among the different Fe-bearing phases, Fe content was subdivided into three main categories. These include extractable iron ($Fe_{ext}$) clumped iron in pyrites, oxides, magnetites and carbonates ($Fe_{py}$, $Fe_{ox}$, $Fe_{mag}$ and $Fe_{carb}$); iron in poorly reactive sheet silicates (mainly chlorite, $Fe_{PRS}$); and unreactive iron ($Fe_U$), defined as the difference between extractable iron and the total iron contents. The extraction of Fe from poorly reactive sheet silicates is however often incomplete and part of Fe contained in chlorites may participate to the fraction of unreactive iron. Iron oxides display low Fe content between 12 to 325 ppm with an average value of 165 ppm. Fe content in magnetites ranges between 149 to 1620 ppm with an average value of 645 ppm. Except for GIS68.9 and GIS69.2, which display Fe concentrations in pyrites close to 0.37wt.%, the average pyrite Fe content is around 0.10wt.%. Fe content in carbonates varies between 0.14 and 0.49wt.% with a mean value of 0.36wt.%. Fe content associated with poorly reactive silicates displays a range between 0.02 and 0.41wt.%, with an average of 0.15wt.%. We observe that Fe is mostly concentrated in carbonate and in sheet silicates as chlorites. Moreover, the difference between $Fe_T$ and $Fe_{ext}$ shows that a large fraction (~25%) of iron is concentrated in unreactive phases.

## 3.3. SIMS analyses

### 3.3.1. Iron isotope compositions

Iron isotope compositions of the 142 micropyrites analyzed in GIS67.8, GIS68.1, GIS68.9 and GIS70.8 are summarized in Table 2. These micropyrites show a large range in iron isotope compositions, between -2.20‰ and +4.44‰ ±0.24‰ (2σ,SD) (Fig. 4a), encompassing the entire terrestrial iron isotope range of values reported in the literature (Dauphas et al., 2017). Stromatolites GIS67.8 and GIS68.9 show a similar large range of $\delta^{56}\text{Fe}$ between -2.20‰ and +4.15‰ (n = 70) and between -1.44‰ and +4.44‰ (n = 60) respectively, while GIS68.1 and GIS70.8 display mostly positive values and smaller ranges of $\delta^{56}\text{Fe}$ values from -0.01‰ to +3.79‰ (n = 8) and from +0.49‰ to +2.41‰ (n = 4) respectively (Fig. 4a). In addition, we observe that iron isotope compositions of analyzed micropyrites might be dependent on the

<figure><img src="image_4.png" /><figcaption>Fig. 3. SEM images of the different types of laminae in which micropyrites were measured. (a) Micropyrites in a micritic lamina (Type 1); (b) Micropyrites in an organic-rich lamina (Type 2); (c) Micropyrites in a fenestra (Type 3). Chl: Chlorite; Qtz: Quartz; Py: Pyrite; Calc: Calcite; Ap: Apatite.</figcaption></figure>
<caption>Table 2</caption>

Iron isotope compositions of 142 micropyrites from GIS67.8, GIS68.1, GIS68.9 and GIS70.8 samples of the Tumbiana Formation. The type of lamina refers to micritic laminae (Type 1), organic-rich laminae (Type 2) and fenestral laminae (Type 3).

<table><thead><tr><th># Sample</th><th>Analyses</th><th><sup>56</sup>Fe<sup>+</sup> intensity (cps)</th><th>δ<sup>56</sup>Fe (‰)</th><th>2σ (‰)</th><th>Type of lamina</th></tr></thead><tbody><tr><td rowspan="25">GIS67.8</td><td>PDP1-G1Z8a</td><td>3.63E+07</td><td>-1.86</td><td>0.21</td><td>3</td></tr><tr><td>PDP1-G1Z16b</td><td>1.58E+07</td><td>1.39</td><td>0.30</td><td>3</td></tr><tr><td>PDP1-G1Z14c</td><td>7.34E+07</td><td>1.13</td><td>0.18</td><td>2</td></tr><tr><td>PDP1-G1Z13d</td><td>8.23E+07</td><td>-0.87</td><td>0.16</td><td>2</td></tr><tr><td>PDP1-G1Z11k</td><td>3.37E+07</td><td>-0.93</td><td>0.19</td><td>3</td></tr><tr><td>PDP1-G1Z11j</td><td>3.40E+07</td><td>-1.17</td><td>0.19</td><td>3</td></tr><tr><td>PDP1-G1Z11i</td><td>3.98E+07</td><td>-0.54</td><td>0.19</td><td>3</td></tr><tr><td>PDP1-G1Z11f</td><td>4.40E+07</td><td>0.90</td><td>0.19</td><td>3</td></tr><tr><td>PDP1-G1Z11b</td><td>5.60E+07</td><td>-0.77</td><td>0.16</td><td>3</td></tr><tr><td>PDP1-G1-Z33i</td><td>3.07E+07</td><td>0.81</td><td>0.19</td><td>1</td></tr><tr><td>PDP1-G1-Z33h</td><td>2.15E+07</td><td>-0.36</td><td>0.22</td><td>1</td></tr><tr><td>PDP1-G1-Z33e</td><td>2.67E+07</td><td>0.53</td><td>0.20</td><td>1</td></tr><tr><td>PDP1-G1-Z33c</td><td>1.50E+07</td><td>-0.30</td><td>0.32</td><td>1</td></tr><tr><td>PDP1-G1-Z33d</td><td>1.55E+07</td><td>-0.22</td><td>0.31</td><td>1</td></tr><tr><td>PDP1-G1-Z25e</td><td>2.63E+07</td><td>2.67</td><td>0.23</td><td>1</td></tr><tr><td>PDP1-G1-Z33b</td><td>2.00E+07</td><td>0.52</td><td>0.25</td><td>1</td></tr><tr><td>PDP1-Z18e</td><td>3.51E+07</td><td>-0.54</td><td>0.21</td><td>3</td></tr><tr><td>PDP1-Z19a</td><td>4.62E+07</td><td>-1.14</td><td>0.19</td><td>3</td></tr><tr><td>PDP1-Z18b</td><td>2.94E+07</td><td>-1.00</td><td>0.21</td><td>3</td></tr><tr><td>PDP1-Z18c</td><td>2.75E+07</td><td>0.32</td><td>0.18</td><td>3</td></tr><tr><td>PDP1-Z19b</td><td>2.72E+07</td><td>-2.20</td><td>0.21</td><td>3</td></tr><tr><td>PDP1-Z19c</td><td>5.68E+07</td><td>-0.57</td><td>0.16</td><td>3</td></tr><tr><td>PDP1-Z19d</td><td>2.19E+07</td><td>-1.96</td><td>0.23</td><td>3</td></tr><tr><td>PDP1-Z19e</td><td>1.54E+07</td><td>1.73</td><td>0.34</td><td>3</td></tr><tr><td>PDP1-Z19f</td><td>4.69E+07</td><td>-1.35</td><td>0.17</td><td>3</td></tr><tr><td>PDP1-Z19g</td><td>2.92E+07</td><td>2.65</td><td>0.22</td><td>3</td></tr><tr><td>PDP1-Z20d</td><td>3.05E+07</td><td>-0.11</td><td>0.21</td><td>3</td></tr><tr><td>PDP1-Z20e</td><td>4.64E+07</td><td>-0.81</td><td>0.17</td><td>3</td></tr><tr><td>PDP1-Z20g</td><td>4.77E+07</td><td>0.24</td><td>0.18</td><td>3</td></tr><tr><td>PDP1-Z27a</td><td>4.32E+07</td><td>-0.54</td><td>0.16</td><td>2</td></tr><tr><td>PDP1-Z27b</td><td>1.45E+07</td><td>0.74</td><td>0.31</td><td>2</td></tr><tr><td>PDP1-Z22a</td><td>5.63E+07</td><td>0.45</td><td>0.17</td><td>3</td></tr><tr><td>PDP1-Z31a</td><td>2.90E+07</td><td>0.09</td><td>0.17</td><td>2</td></tr><tr><td>PDP1-Z31b</td><td>2.39E+07</td><td>-0.28</td><td>0.26</td><td>2</td></tr><tr><td>PDP1-Z31c</td><td>5.43E+07</td><td>0.63</td><td>0.14</td><td>2</td></tr><tr><td>PDP1-Z31d</td><td>2.34E+07</td><td>1.78</td><td>0.22</td><td>2</td></tr><tr><td>PDP1-Z31e</td><td>3.27E+07</td><td>-1.54</td><td>0.19</td><td>2</td></tr><tr><td>PDP1-Z31f</td><td>3.59E+07</td><td>1.79</td><td>0.16</td><td>2</td></tr><tr><td>PDP1-Z36a</td><td>3.82E+07</td><td>-0.38</td><td>0.16</td><td>2</td></tr><tr><td>PDP1-Z37a</td><td>5.87E+07</td><td>1.91</td><td>0.13</td><td>1</td></tr><tr><td>PDP1-Z37b</td><td>3.43E+07</td><td>0.81</td><td>0.22</td><td>1</td></tr><tr><td>PDP1-Z1a</td><td>2.14E+07</td><td>2.36</td><td>0.18</td><td>1</td></tr><tr><td>PDP1-Z1b</td><td>5.49E+07</td><td>2.77</td><td>0.11</td><td>1</td></tr><tr><td>PDP1-Z1c</td><td>5.08E+07</td><td>1.92</td><td>0.14</td><td>1</td></tr><tr><td>PDP1-Z1d</td><td>2.10E+07</td><td>1.37</td><td>0.22</td><td>1</td></tr><tr><td>PDP1-Z2a</td><td>2.67E+07</td><td>0.35</td><td>0.23</td><td>3</td></tr><tr><td>PDP1-Z2b</td><td>2.61E+07</td><td>0.98</td><td>0.18</td><td>3</td></tr><tr><td>PDP1-Z2d</td><td>2.42E+07</td><td>-0.63</td><td>0.19</td><td>3</td></tr><tr><td>PDP1-Z2e</td><td>1.91E+07</td><td>0.18</td><td>0.26</td><td>3</td></tr><tr><td>PDP1-Z2f</td><td>5.23E+07</td><td>1.54</td><td>0.16</td><td>3</td></tr><tr><td>PDP1-Z2g</td><td>2.94E+07</td><td>1.04</td><td>0.16</td><td>3</td></tr><tr><td>PDP1-Z2h</td><td>3.35E+07</td><td>1.21</td><td>0.14</td><td>3</td></tr><tr><td>PDP1-IV-Z9a</td><td>2.48E+07</td><td>4.15</td><td>0.26</td><td>1</td></tr><tr><td>PDP1-IV-Z21</td><td>2.65E+07</td><td>-1.03</td><td>0.20</td><td>3</td></tr><tr><td>PDP1-IV-Z19a</td><td>2.52E+07</td><td>-1.27</td><td>0.20</td><td>3</td></tr><tr><td>PDP1-IV-Z19c</td><td>1.76E+07</td><td>-0.88</td><td>0.24</td><td>3</td></tr><tr><td>PDP1-IV-Z19d</td><td>1.99E+07</td><td>-0.26</td><td>0.25</td><td>3</td></tr><tr><td>PDP1-IV-Z19e</td><td>1.50E+07</td><td>2.15</td><td>0.33</td><td>3</td></tr><tr><td>PDP1-IV-Z19f</td><td>1.94E+07</td><td>-0.92</td><td>0.29</td><td>3</td></tr><tr><td>PDP1-IV-Z20</td><td>2.68E+07</td><td>-0.15</td><td>0.22</td><td>1</td></tr></tbody></table>
<table><caption>Table 2 (continued)</caption><thead><tr><th># Sample</th><th>Analyses</th><th><sup>56</sup>Fe<sup>+</sup> intensity (cps)</th><th>δ<sup>56</sup>Fe (‰)</th><th>2σ (‰)</th><th>Type of lamina</th></tr></thead><tbody><tr><td rowspan="20">GIS68.1</td><td>PDP1-IV-Z18a</td><td>3.32E+07</td><td>1.17</td><td>0.18</td><td>3</td></tr><tr><td>PDP1-IV-Z18b</td><td>2.11E+07</td><td>-2.06</td><td>0.25</td><td>3</td></tr><tr><td>PDP1-IV-Z12a</td><td>3.66E+07</td><td>0.84</td><td>0.17</td><td>1</td></tr><tr><td>PDP1-IV-Z11a</td><td>2.46E+07</td><td>2.33</td><td>0.22</td><td>1</td></tr><tr><td>PDP1-IV-Z11b</td><td>2.77E+07</td><td>-1.40</td><td>0.21</td><td>1</td></tr><tr><td>PDP1-IV-Z6a</td><td>2.09E+07</td><td>3.43</td><td>0.36</td><td>1</td></tr><tr><td>PDP1-IV-Z6b</td><td>2.08E+07</td><td>0.71</td><td>0.23</td><td>1</td></tr><tr><td>PDP1-IV-Z4a</td><td>2.68E+07</td><td>-0.27</td><td>0.18</td><td>1</td></tr><tr><td>PDP1-IV-Z5a</td><td>1.24E+07</td><td>-1.76</td><td>0.37</td><td>1</td></tr><tr><td>PDP1-IV-Z5b</td><td>1.54E+07</td><td>2.43</td><td>0.28</td><td>2</td></tr><tr><td>PDP1-III-Z1b</td><td>3.49E+07</td><td>3.48</td><td>0.16</td><td>1</td></tr><tr><td>PDP1-III-Z15e</td><td>2.70E+07</td><td>3.30</td><td>0.19</td><td>1</td></tr><tr><td>PDP1-III-Z15d</td><td>2.41E+07</td><td>3.79</td><td>0.29</td><td>1</td></tr><tr><td>PDP1-III-Z20c</td><td>3.62E+07</td><td>2.20</td><td>0.17</td><td>1</td></tr><tr><td>PDP1-III-Z23a</td><td>3.42E+07</td><td>-0.01</td><td>0.20</td><td>1</td></tr><tr><td>PDP1-III-Z13a</td><td>3.07E+07</td><td>0.44</td><td>0.17</td><td>1</td></tr><tr><td>PDP1-III-Z13b</td><td>2.62E+07</td><td>1.20</td><td>0.18</td><td>1</td></tr><tr><td>PDP1-III-Z10b</td><td>3.17E+07</td><td>2.42</td><td>0.23</td><td>1</td></tr><tr><td rowspan="20">GIS68.9</td><td>PDP1-G3Z4a</td><td>6.32E+07</td><td>-0.21</td><td>0.15</td><td>1</td></tr><tr><td>PDP1-G3Z4b</td><td>4.33E+07</td><td>3.81</td><td>0.50</td><td>1</td></tr><tr><td>PDP1-G3Z4e</td><td>5.57E+07</td><td>1.34</td><td>0.15</td><td>1</td></tr><tr><td>PDP1-G3Z4f</td><td>6.63E+07</td><td>0.81</td><td>0.16</td><td>1</td></tr><tr><td>PDP1-G3Z4g</td><td>3.72E+07</td><td>3.41</td><td>0.18</td><td>1</td></tr><tr><td>PDP1-G3Z4h</td><td>6.43E+07</td><td>0.10</td><td>0.18</td><td>1</td></tr><tr><td>PDP1-G3Z4i</td><td>5.61E+07</td><td>0.45</td><td>0.16</td><td>1</td></tr><tr><td>TumbianaZ25b</td><td>4.62E+07</td><td>2.18</td><td>0.29</td><td>2</td></tr><tr><td>TumbianaZ30a</td><td>5.28E+07</td><td>-0.33</td><td>0.14</td><td>1</td></tr><tr><td>TumbianaZ31a</td><td>5.98E+07</td><td>-0.58</td><td>0.16</td><td>1</td></tr><tr><td>TumbianaZ31b</td><td>6.30E+07</td><td>0.17</td><td>0.15</td><td>1</td></tr><tr><td>TumbianaZ17a</td><td>7.06E+07</td><td>-0.75</td><td>0.14</td><td>1</td></tr><tr><td>TumbianaZ17b</td><td>7.10E+07</td><td>0.07</td><td>0.13</td><td>1</td></tr><tr><td>TumbianaZ18a</td><td>7.65E+07</td><td>-1.44</td><td>0.13</td><td>1</td></tr><tr><td>TumbianaZ16a</td><td>7.35E+07</td><td>-0.68</td><td>0.12</td><td>1</td></tr><tr><td>Tumbiana-Z19</td><td>7.37E+07</td><td>0.45</td><td>0.14</td><td>1</td></tr><tr><td>Tumbiana-Z15</td><td>7.99E+07</td><td>-0.36</td><td>0.15</td><td>1</td></tr><tr><td>Tumbiana-Z23a</td><td>2.92E+07</td><td>0.02</td><td>0.17</td><td>2</td></tr><tr><td>Tumbiana-Z23b</td><td>6.56E+07</td><td>-0.27</td><td>0.15</td><td>2</td></tr><tr><td>Tumbiana-Z23c</td><td>1.86E+07</td><td>-1.12</td><td>0.45</td><td>2</td></tr><tr><td>Tumbiana-Z23f</td><td>5.37E+07</td><td>1.73</td><td>0.17</td><td>2</td></tr><tr><td>Tumbiana-Z24b</td><td>5.11E+07</td><td>2.23</td><td>0.24</td><td>2</td></tr><tr><td>LE_Z1</td><td>6.48E+07</td><td>0.91</td><td>0.15</td><td>2</td></tr><tr><td>LE-Z2e</td><td>5.89E+07</td><td>4.44</td><td>0.16</td><td>2</td></tr><tr><td>LE-Z2c</td><td>4.42E+07</td><td>3.80</td><td>0.21</td><td>2</td></tr><tr><td>LE-Z4b</td><td>5.47E+07</td><td>4.02</td><td>0.22</td><td>2</td></tr><tr><td>LE-Z4a</td><td>5.50E+07</td><td>0.08</td><td>0.17</td><td>2</td></tr><tr><td>LE-Z8a</td><td>5.72E+07</td><td>2.09</td><td>0.23</td><td>2</td></tr><tr><td>LE-Z12b</td><td>5.97E+07</td><td>1.42</td><td>0.15</td><td>2</td></tr><tr><td>LE-Z12a</td><td>4.77E+07</td><td>2.80</td><td>0.17</td><td>2</td></tr><tr><td>LE-Z13b</td><td>3.20E+07</td><td>3.62</td><td>0.21</td><td>2</td></tr><tr><td>L1_Z2a</td><td>4.68E+07</td><td>2.05</td><td>0.20</td><td>2</td></tr><tr><td>L3_Z3</td><td>5.48E+07</td><td>1.59</td><td>0.16</td><td>2</td></tr><tr><td>Matrix_Z1b</td><td>2.78E+07</td><td>1.94</td><td>0.19</td><td>1</td></tr><tr><td>Matrix_Z1c</td><td>5.06E+07</td><td>3.06</td><td>0.13</td><td>1</td></tr><tr><td>Matrix_Z1f</td><td>4.73E+07</td><td>1.92</td><td>0.14</td><td>1</td></tr><tr><td>Matrix_Z1e</td><td>4.77E+07</td><td>3.34</td><td>0.15</td><td>1</td></tr><tr><td>Matrix_Z1d</td><td>5.39E+07</td><td>0.55</td><td>0.15</td><td>1</td></tr><tr><td>Matrix_Z2a</td><td>4.60E+07</td><td>1.75</td><td>0.16</td><td>1</td></tr><tr><td>Matrix_Z2b</td><td>3.64E+07</td><td>1.33</td><td>0.15</td><td>1</td></tr><tr><td>Matrix_Z2c</td><td>2.72E+07</td><td>1.65</td><td>0.17</td><td>1</td></tr><tr><td>Matrix_Z2d</td><td>3.03E+07</td><td>1.87</td><td>0.19</td><td>1</td></tr><tr><td>Matrix_Z2e</td><td>3.86E+07</td><td>0.19</td><td>0.14</td><td>1</td></tr><tr><td>Matrix_Z3a</td><td>4.57E+07</td><td>1.02</td><td>0.12</td><td>1</td></tr></tbody></table>
<table><caption>Table 2 (continued)</caption><thead><tr><th># Sample</th><th>Analyses</th><th><sup>56</sup>Fe<sup>+</sup> intensity (cps)</th><th>δ<sup>56</sup>Fe (‰)</th><th>2σ (‰)</th><th>Type of lamina</th></tr></thead><tbody><tr><td rowspan="25">GIS70.8</td><td>Matrix_Z3d</td><td>1.68E+07</td><td>-1.34</td><td>0.27</td><td>1</td></tr><tr><td>Matrix_Z3e</td><td>4.05E+07</td><td>0.78</td><td>0.14</td><td>1</td></tr><tr><td>Matrix_Z3f</td><td>2.78E+07</td><td>2.93</td><td>0.24</td><td>1</td></tr><tr><td>Matrix_Z3g</td><td>5.56E+07</td><td>1.06</td><td>0.13</td><td>1</td></tr><tr><td>Matrix_Z3h</td><td>4.39E+07</td><td>1.25</td><td>0.13</td><td>1</td></tr><tr><td>Matrix_Z3i</td><td>2.21E+07</td><td>1.27</td><td>0.24</td><td>1</td></tr><tr><td>Matrix_Z4a</td><td>3.46E+07</td><td>1.08</td><td>0.18</td><td>1</td></tr><tr><td>Matrix_Z4b</td><td>1.95E+07</td><td>-0.61</td><td>0.23</td><td>1</td></tr><tr><td>Matrix_Z4c</td><td>3.41E+07</td><td>1.36</td><td>0.15</td><td>1</td></tr><tr><td>Matrix_Z4d</td><td>4.71E+07</td><td>0.45</td><td>0.12</td><td>1</td></tr><tr><td>Matrix_Z5a</td><td>6.67E+07</td><td>0.41</td><td>0.11</td><td>1</td></tr><tr><td>Matrix_Z5b</td><td>4.64E+07</td><td>1.34</td><td>0.13</td><td>1</td></tr><tr><td>Matrix_Z5c</td><td>3.42E+07</td><td>1.09</td><td>0.18</td><td>1</td></tr><tr><td>Matrix_Z5d</td><td>2.98E+07</td><td>1.01</td><td>0.18</td><td>1</td></tr><tr><td>Matrix_Z6b</td><td>5.75E+07</td><td>0.72</td><td>0.12</td><td>1</td></tr><tr><td>Matrix_Z6c</td><td>5.81E+07</td><td>0.96</td><td>0.12</td><td>1</td></tr><tr><td>PDP1-G2-Z9a</td><td>4.28E+07</td><td>2.04</td><td>0.29</td><td>1</td></tr><tr><td>PDP1-G2-Z9b</td><td>5.60E+07</td><td>0.49</td><td>0.14</td><td>1</td></tr><tr><td>PDP1-G2-Z9d</td><td>5.24E+07</td><td>1.72</td><td>0.15</td><td>1</td></tr><tr><td>PDP1-G2-Z9e</td><td>5.69E+07</td><td>2.41</td><td>0.15</td><td>1</td></tr></tbody></table>

type of their hosting lamina (Fig. 4b). Pyrites measured in micritic laminae show large isotopic variations (δ<sup>56</sup>Fe from -1.76‰ to +4.15‰, n = 77), comparable with those recorded in pyrites hosted in organic-rich laminae (δ<sup>56</sup>Fe from -1.54‰ to +4.44‰, n = 29). In comparison, fenestral laminae yield a smaller Fe isotope signal from -2.20‰ to +2.65‰ (n = 36). The median δ<sup>56</sup>Fe value is +1.01‰ in micritic type laminae, +1.42‰ in organic-rich laminae and -0.54‰ in fenestral laminae (Fig. 4c).

Iron isotopic compositions can also be correlated with the pyrite sizes (Fig. 5). Fe isotopic compositions range between -0.54‰ and +2.67‰ for pyrites below 5 µm in size (n = 9, mean δ<sup>56</sup>Fe = 0.69‰), between -1.76‰ and +3.94‰ for the 5-10 µm category (n = 64, mean δ<sup>56</sup>Fe = 0.68‰), -2.20‰ and +4.44‰ for pyrites of [10-15] µm (n = 53, mean δ<sup>56</sup>Fe = 0.92‰) and between -1.44‰ and +3.48‰ for pyrites bigger than 15 µm (n = 16, mean δ<sup>56</sup>Fe = 1.07‰). Pyrites > 10 µm in size seem to record <sup>56</sup>Fe-enriched isotopic compositions. However, due to the difference in the population size of each categories, we cannot conclude on a clear influence of the size of the pyrites on their respective Fe isotopic compositions. Furthermore, pyrites with similar size can record a large Fe isotopic range (Fig. 6a and b; pyrites of 8-9 µm size have δ<sup>56</sup>Fe values between -0.61‰ and +4.02‰). Finally, individual pyrite grain observations after SIMS analysis support that there is no correlation between the δ<sup>56</sup>Fe values and the pyrite shapes (Fig. 6c; Electronic Annex (EA); Fig. EA-5).

### 3.3.2. Probability density and comparison with previous in situ studies

Iron isotopic compositions are presented using probability densities, which calculate a Gaussian function for

each analysis considering the associated error (Marin-Carbonne et al., 2012). Because probability density distribution strongly depends on bin size and bin boundary, we chose a bin size of 0.1 in order to reduce the bin boundary effects, i.e. the distribution has the same shape whatever the starting δ<sup>56</sup>Fe value of the probability density, even if artificial secondary peaks are produced (Fig. 7a). We also report cumulative functions which allow us to clearly identify bimodal and modal distributions (Fig. 7). In previous works from the Redmont area, Nishizawa et al. (2010) and Yoshiya et al. (2012) reported a bimodal distribution (Fig. 7b) with a positive mode around +1‰ and a negative mode around -1.8‰. In this study we observe a monomodal δ<sup>56</sup>Fe distribution with mode around +1‰ (Fig. 7a). The cumulative probability functions provide an alternative dataset representation (Fig. 7c and d) and show a stair shape distribution, which highlights the two modes observed in the previous studies (Fig. 7d). In this study, probability density and cumulative functions are in good agreement with a theoretical gaussian curve, showing a progressive increasing ramp for the cumulative representation (Fig. 7c). The whole dataset thus follows a gaussian distribution centered at +1‰ (Fig. 7a). The difference between our distribution and previously published distributions is likely due to different grade of metasomatism that affected variously the Redmont and Meentheena localities (White et al., 2014). Moreover, most of the pyrite grains in this study range from 5 to 10 µm when pyrites previously measured by Nishizawa et al. (2010) and Yoshiya et al. (2012) were tens to hundreds of µm. Micropyrites from this study likely recorded more discreet processes that are not evident in larger pyrite grains from the Redmont locality.
<figure><img src="image_5.png" /></figure>

Fig. 4. Iron isotopic compositions measured (a) in the different samples GIS67.8 (orange dots, n = 70), GIS68.1 (red dots, n = 8), GIS68.9 (yellow dots, n = 60) and GIS70.8 (green dots, n = 4); (b) in the different laminae: micritic (Type 1, n = 77), organic-rich (Type 2, n = 29) and fenestrae (Type 3, n = 36). GIS67.8 displays large δ⁵⁶Fe isotopic range between -2.20‰ and +4.15‰, similar to GIS68.9 isotopic range. Amplitude of isotopic ranges for GIS68.1 and GIS70.8 are less wide and probably un-representative of the whole isotopic variability. In general, more δ⁵⁶Fe-depleted values are observed in fenestrae whereas organic-rich laminae are characterized by δ⁵⁶Fe-enriched isotopic values. Error bars yield the 2SE (standard error); (c) Box plots using δ⁵⁶Fe values of the different laminae types. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

<figure><img src="image_6.png" /></figure>

Fig. 5. Iron isotopic compositions as a function of the size of micropyrites. We considered 4 categories of size: less than 5 µm, 5 µm to 10 µm, 10 µm to 15 µm and > 15 µm. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

# 4. DISCUSSION

The large range of δ⁵⁶Fe values measured in micropyrites of the Tumbiana Formation could reflect several biotic or abiotic processes. In the following section, we discuss the possible iron sources involved in pyrite precipitation and processes able to explain the range of δ⁵⁶Fe value recorded in the studied stromatolitic laminae.

## 4.1. Context of micropyrites precipitation

### 4.1.1. Pyrite as a minor sink of iron in Tumbiana stromatolites

Although iron speciation proxy is not calibrated for lacustrine environment, important features can be drawn from these analyses. In Tumbiana stromatolites, Fe speciation data show that FePRS and FeU, mostly represented by
<figure><img src="image_7.png" /><figcaption>Fig. 6. SEM pictures of micropyrites in GIS68.9. (a) and (b) Fe isotopic compositions measured in pyrites of 8 to 9 µm size. This size category shows an isotopic range up to 4‰, showing the absence of correlation between δ$^{56}$Fe and the size of the pyrite; c) Fe isotopic compositions of two hexagonal pyrites of 16 µm and 8 µm in size. These pyrites recorded an isotopic difference of 1‰, located at 30 µm from each other, meaning that iron isotopic compositions are not controlled by the shape of the pyrite.</figcaption></figure>

<figure><img src="image_8.png" /><figcaption>Fig. 7. Probability density distributions calculated for a) the entire δ$^{56}$Fe dataset (black curve, this study) showing a continuous, gaussian shape distribution centered on +1‰. The red curve represents the theoretical Gaussian function given by $f(x) = \frac{1}{\sigma*\sqrt{2\pi}} * e^{-\frac{(x-\mu)^2}{2\sigma^2}}$ with μ the dataset mean value and σ the standard deviation; b) Previous published δ$^{56}$Fe values for stromatolite samples describing a bimodal distribution (grey curve, data from Nishizawa et al., 2010; Yoshiya et al., 2012); c) Cumulative function of the entire dataset (black curve, this study) and Gaussian fit (red curve) and d) Cumulative function associated with Nishizawa et al. (2010) and Yoshiya et al. (2012) data. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>

chlorites, represent in average 73% of the total iron content (FeT). The occurrence of few detrital minerals like zircon and anatase suggests that continental weathering and riverine transport were active during deposition of the Tumbiana sediments. Moreover, in situ weathering and

devitrification of the mafic substrate, including volcanic ashes, likely drove the formation of chlorites that are abundant in the samples. As aluminium is a conservative element during weathering and is not remobilized by low temperature hydrothermal fluids, the positive correlation between
Fe~T~ and Al contents within stromatolite layers (Electronic Annex (EA); Fig. EA-6) suggests that Fe is mostly inherited from Fe-rich mafic lithologies (Sugisaki, 1978; Wintsch and Kvale, 1994; Sugitani et al., 1996). The input of Fe-rich detritus can explain the low abundance of highly reactive iron (Fe~HR~) compare to Fe~T~ (Table 1). The low pyrite content (4% relative to Fe~T~) coupled with low Fe~pyr~/Fe~HR~ ratio (~0.11 on average) argue for non-sulfidic conditions in the Tumbiana water column. However, as organic-matter is sulfidized in the Tumbiana stromatolites (Lepot et al., 2009), variable local sulfidic conditions could have occurred in the sediment pore water. Finally, the micropyrites studied here only reflect a small fraction of the Fe budget in the Tumbiana environment and therefore cannot be considered to account for the whole Tumbiana basin Fe cycling.

### 4.1.2. *Origin of the laminae diversity*

The observed fabric variability in our samples underlines the existence of four different types of laminae reflecting diverse mineralization processes leading to the formation of stromatolites (Fig. 2a). The micritic laminae (Type 1), which are mainly made of microcrystalline calcite, reflect a microbial cement precipitation commonly observed in stromatolites through Earth history (Visscher et al., 1998). The organic-rich laminae (Type 2) are chlorite-rich suggesting a terrigenous input in the system. The spatial distribution of these laminae along the domal structure of the stromatolite, their morphology and the lateral variability in thickness argue against a stylolitic origin. Process of trapping of detrital material is highlighted and illustrated by the presence of chlorite, zircon and anatase. The formation of organic matter is associated with the activity of microbial mats. Fenestral laminae (Type 3) are laterally more or less continuous. These sparitic fenestrae indicate a syngeneic or early diagenetic filling of a primary porosity that can be formed by gas released after organic matter or mat breakdown (Tebbutt et al., 1965; Choquette and Pray, 1970; Mata et al., 2012). Finally, the coarse-grained laminae (Type 4), that are devoid of pyrite, highlight an intense trapping and binding of peloids and ooids from adjacent sedimentary bodies. The depositional model of Awramik and Buchheim (2009) for the Meentheena Member positioned the studied stromatolites between shallow and high-energy ooidal grainstones and deeper siltstones and shales. In such a depositional environment, micritic laminae (Type 1) reflect optimal conditions for stromatolite development associated with increased carbonate precipitation. However, the thickness of these micritic laminae does not exceed few millimeters, indicating that their growth was frequently disrupted. Indeed, the Meentheena sedimentary system was relatively dynamic, notably disturbed by rapid lateral facies changes due to frequent storm events (Awramik and Buchheim, 2009). Stromatolite growth was thereby interrupted by recurrent ooidal sand body migration leading to the formation of the coarse-grained laminae (Type 4), or by periods of increased terrigenous runoff to form organic-rich laminae (Type 2). The presence of nutrients associated with river discharges is also consistent with higher rates of organic matter production recorded in these Type 2 laminae. The presence of chlorites also promoted a

best post-deposit preservation of organic compounds in this lamina type (Jardine et al., 1989).

Differences in stromatolite growth rates are generally assumed to be controlled by an interplay between internal (microbial communities) and external (i.e. environmental) factors (Reid et al., 2000; Dupraz et al., 2009; Bowlin et al., 2012; Bouton et al., 2016). Intrinsic parameters relate to the metabolic activities in the microbial mat or the biofilm, that is the dynamism and diversity of the microbial community. These parameters strongly act on the mineralization of laminae by controlling the availability of HCO₃⁻ (the alkalinity) and Ca²⁺ ions (Aloisi, 2008; Dupraz et al., 2009). Extrinsic factors relate to variations in environmental factors, which can influence both the microbial activity and ions availability. These variables include for example, temperature, chemical composition of water (e.g., salinity, pH, alkalinity), solar irradiance, supply (i.e. source material and flux) of sedimentary material (Reid et al., 2000; Arp et al., 2003; Andres and Pamela Reid, 2006; Planavsky and Grey, 2008; Jahnert et al., 2013; Bouton et al., 2016; Bouton et al., 2020). Among the different factors that can influence the stromatolite growth and the microbial mineralization potential, physical parameters such as trapping and binding also directly influence stromatolite growth and carbonate precipitation rates (Awramik et al., 1976; Reid et al., 2000; Awramik and Grey, 2005). The diversity of laminae recorded in the Tumbiana stromatolites is a clear illustration of the different processes (mineralization, trapping, terrigenous influence) that can be involved in stromatolite fabric and growth. It also illustrates how the depositional environment and early diagenesis may have influenced the micropyrite distribution and Fe isotopic compositions.

### 4.1.3. *A synsedimentary origin of the micropyrites*

In the four analyzed samples, the absence of rounded grains argues against a detrital origin of the pyrites. Similarly, the absence of hydrothermal veins and secondary veins of pyrites argue against a late stage hydrothermal overprint. A limited metamorphic imprint is further supported by δ¹³C analyses in carbonate and organic matter, where the range of equilibrium fractionation (Δ¹³Cₚcarb-org) expected for an isotopic re-equilibration during high-grade metamorphism is inconsistent with the one determined in the Tumbiana samples (Thomazo et al., 2009). The absence of major secondary overprint and preservation of primary features have been likely favored by the pervasive silicification affecting the different samples studied here. Petrographic investigations show that the size, morphology and distribution of the pyrites recorded in the different laminae cannot be used to identify a relative chronology of formation. The smallest pyrites (<5µm to hundreds nm) occur preferentially in intimate association with organic matter in Type 2 laminae and are more concentrated in this lamina type (Fig. 3; Electronic Annex (EA); Fig. EA-7). This indicates a causal relationship between organic matter availability on the abundance of pyrite formed (Rickard et al., 2017). Because micropyrites are enclosed in subparallel stromatolitic laminae, associated with pristine organic matter (Lepot et al., 2009; Marin-Carbonne et al., 2018), and
display a unimodal Fe isotopic distribution, we interpreted
these micropyrites to be precipitated during stromatolite
growth or during early diagenesis.

## 4.2. Large δ⁵⁶Fe range of micropyrites reflects a kinetic effect

Different processes can account for the large Fe isotope
range of values recorded in the studied micropyrites. These
include: i) δ⁵⁶Fe values inherited from Fe sources, ii) Ray-
leigh distillation process during precipitation of pyrites
from monosulfides, iii) variable pyrite precipitation rate
(Mansor and Fantle, 2019), iv) H₂S reduction of Fe(III)-
oxides and v) Dissimilatory Iron Reduction (DIR)
(Yoshiya et al., 2012). All Fe isotope fractionations consid-
ered in the following discussion are summarized in Fig. 8.

4.2.1. Iron isotope compositions inherited from Fe sources

Iron isotope compositions of sedimentary micropyrites
could be inherited from various sources of Fe to the system.
Fe²⁺ may have been delivered from mafic mineral dissolu-
tion, the incorporation into authigenic carbonates or mag-
matic magnetites. Fe isotope fractionation in magmatic
rocks are partially attributed to magma evolution pro-
cesses, such as fractional crystallization or fluid exsolution.
The largest range of Fe isotope compositions of ~1.6‰ has
been measured in olivine and assigned to Fe-Mg exchange
diffusion between olivine and melt (Teng et al., 2011).
Except this large isotope range, small fractionations in the
order of 0.1‰ to 0.3‰ are typical for magmatic processes
(Dauphas et al., 2017 for a review). For example, an

equilibrium fractionation of +0.2 ± 0.016‰ has been
described between fayalite and magnetite at temperatures
ranging between 600 and 800 °C (Shahar et al., 2008). Fur-
thermore, several studies have assessed Fe isotope fraction-
ations that occur during abiotic mafic rock and mineral
dissolution. A large Fe isotope fractionation of ~-1.5‰
has been measured for the weathering of a tholeiitic basalt
between aqueous Fe released after few hours of experiment
and the basalt. Nonetheless, the extend of this fractionation
tends to decrease (Δ⁵⁶Fe solution-rock = -0.5‰) as a function
of leaching time (Chapman et al., 2009). Only few studies
have investigated Fe isotope fractionation during dissolu-
tion of silicates. The dissolution of hornblende in the pres-
ence of different organic ligands produces dissolved iron
slightly enriched in light isotopes, with a maximum frac-
tionation of -0.6‰ between Fe in solution and the mineral
(Brantley et al., 2001; Brantley et al., 2004). Similar frac-
tionations were determined during the dissolution of biotite
and chlorite from granite (Kiczka et al., 2010). Experiments
measuring Fe isotope fractionations between aqueous Fe
and siderite have shown Δ⁵⁶Fe Fe(II)aq-siderite of 0‰ and
+0.5‰, respectively under biotic and abiotic influences
(Wiesli et al., 2004; Johnson et al., 2005). An equilibrium
Fe fractionation factor of +0.9‰ has been reported for
ankerite formation during microbial DIR (Johnson et al.,
2005). Finally, Sossi et al. (2012) reported Fe isotope frac-
tionation of ~+0.1‰ between magnetite and melt at tem-
perature of 900–1000 °C. Thus the Fe isotope range
inherited from magmatic processes and/or mineral mineral
precipitation is limited and it seems difficult to explain

<figure><img src="image_9.png" /><figcaption>Fig. 8. Theoretical Fe isotopic fractionations between oxidized Fe species, ferrous Fe, monosulfide and pyrite. Data from Welch et al. (2003); Bullen et al. (2001); Croal et al. (2004); Guilbaud et al. (2011); Polyakov et al. (2007); Butler et al. (2005); Crosby et al. (2007). Blue and red colors indicate respectively Fe fractionations due to biological and abiotic reactions. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>
the ~6‰ range in Fe isotope compositions measured in micropyrites by one or several of the fractionation processes described above.

### 4.2.2. Rayleigh distillation during precipitation of pyrites from monosulfides

Considering that the pyrites studied here have a synsedimentary origin, micropyrites Fe isotope signatures can result from a Rayleigh distillation process during conversion of pyrite precursors attending stromatolite growth or during early stages of diagenesis. Assuming Fe isotope composition of 0‰ for the overlying water column (Sharma et al., 2001; Beard et al., 2003; Severmann et al., 2004; Johnson et al., 2008), monosulfide pyrite precursor, most likely mackinawite, should in average carry a δ⁵⁶Fe value of about -0.85‰ (Butler et al., 2005). Using a fractionation factor between pyrite and mackinawite of ΔFeS₂-FeS = -2.2‰ (Guilbaud et al., 2011), we can model the theoretical evolution of precipitated pyrite iron isotopes signature during a Rayleigh distillation process. Results of this isotopic evolution reported as a cumulative probability plot in Fig. 9 show a large isotopic range with values as high as 8‰. The observed range of values obtained in this study between -2.20‰ and +4.44‰ is completed for 98% of the Fe fraction precipitated as pyrites in the Rayleigh process. However, in the case of Rayleigh fractionation during the conversion of mackinawite to pyrite, a much higher abundance of isotopically light values would be expected, in contrast to what is observed in the micropyrites (Fig. 9). Additionally, a gradual increase of Fe isotopic compositions due to Rayleigh distillation may generate a specific spatial distribution, from low to progressively higher δ⁵⁶Fe values from the lamina to the sample scale. The random distribution of δ⁵⁶Fe values of the micropyrites

within a single lamina (e.g., GIS68.9; δ⁵⁶Fe from +0.08 ± 0.17‰ to +4.44 ± 0.16‰), which is observed in the whole sample, does not support the spatial distribution that might be expected during a Rayleigh distillation (Electronic Annex (EA); Fig. EA-8). Hence, we suggest that a single stage of pyrite precipitation associated with Rayleigh distillation is not supported here.

### 4.2.3. Variable pyrite precipitation rate

A recent study proposed that a combination of both kinetic and equilibrium isotopic effects could explain the whole Fe isotopic range in sedimentary pyrites recorded through geological history (Mansor and Fantle, 2019). Pyrites with positive δ⁵⁶Fe values can be formed under equilibrium considering a fractionation factor between pyrite and the initial Fe(II)aq of about +4.5‰ (Polyakov et al., 2007). Alternatively, pyrites with negative δ⁵⁶Fe values reflect kinetic effect between pyrite and mackinawite (ΔFeS₂-FeS = -2.2‰, Guilbaud et al., 2011). Hence, to account for the large Fe isotopic range recorded in sedimentary pyrite, Mansor and Fantle (2019) proposed a variable expression of the kinetic (KIE) and equilibrium isotope effect (EIE) as a function of the rate of pyrite precipitation (RPP) and isotopic exchange between pyrite and Fe(II)aq, the latter being controlled by the particle size and the concentrations of Fe(II) and H₂S in the system. For example, heavy Fe isotope pyrites are expected when EIE is expressed in association with low RPP. Consequently, pyrite showing δ⁵⁶Fe signatures close to 0‰ can reflect a growing influence of the KIE and an increase in RPP (Guilbaud et al., 2011; Mansor and Fantle, 2019). This hypothesis of a first order control on the δ⁵⁶Fe signal of pyrite precipitation on the long-time range (Guilbaud et al., 2011; Rolison et al., 2018) is further supported by the positive correlation

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

Fig. 9. Rayleigh distillation model. Cumulative function calculated for a near complete (98%) Rayleigh distillation (in red) using a fractionation factor of -2.2‰ between FeS₂ and FeS, and an initial δ⁵⁶FeFeS of -0.85‰ (Guilbaud et al., 2011; Butler et al., 2005). The black curve represents our dataset. δ⁵⁶Fe values for Rayleigh Distillation is given by $\delta^{56}Fe_{prod} = \alpha_{FeS2-FeS} * (\delta^{56}Fe_{res} + 1000) - 1000$ and $\delta^{56}Fe_{res} = (\delta^{56}Fe_{init} + 1000) * f^{(\alpha_{FeS2-FeS} - 1)} - 1000$, where “prod” refers to the product, “res” to the residue and f is the remaining fraction of reactant. The comparison suggests that a process of pyrite recrystallisation from monosulfides controlled by a Rayleigh distillation do not fit our measurements. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)
revealed between $\delta^{56}$Fe and $\delta^{34}$S values recorded in sedi-
mentary pyrites (Heard et al., 2020). At the scale of a single
pyrite grain, we can assume that RPP is faster in the initial
stage relative to the final stage of precipitation due to either
Fe(II)aq and S pool availability or surface area to volume
ratio variations. Accordingly, a single pyrite grain should
show evolving isotope signal toward more positive $\delta^{56}$Fe
values from core to rim. Because of the size of the micropy-
rites and the low image resolution of the SIMS, profiles
along pyrite grain were not completed here. Nonetheless,
the exhaustive comparison between $\delta^{56}$Fe values and the
size of micropyrites does not show a specific trend
(Fig. 5). The different concentrations of pyrites between
the Type 1, 2 and 3 laminae may reflect variable Fe/S ratios
or variable organic matter availability among these lami-
nae, which should influence the RPP. Although micritic
and organic-rich laminae include variable concentration
of pyrites, comparison of Fe isotopic compositions between
these two lamina types shows similar ranges, with $\delta^{56}$Fe
between -1.76‰ and +4.15‰ in the micritic laminae (Type
1) compared to -1.54‰ and +4.44‰ in the organic-rich
laminae (Type 2). Comparable Fe isotopic ranges in both
types of lamina suggests that the micropyrite Fe isotopic
compositions are not controlled by the availability of
organic matter between the lamina types. In a previous
study focused on S isotope compositions of nanoprites
preserved in Type 2 organic-rich laminae, Marin-
Carbonne et al. (2018) concluded that sulfate was a limiting
reactant but may have been locally concentrated in micro-

metric scale niches. In our samples, the different laminae
record variable amounts of detrital material that can
account for varying Fe/S ratios at the micro-scale. The
higher terrigenous fraction recognized in organic-rich lam-
inae (see Section 4.1.2) should reflect a different Fe/H₂S
ratio compare to micritic and fenestral laminae. However,
micropyrites from organic-rich and micritic laminae seem
to record random $\delta^{56}$Fe ranges over a distance of one to
several hundreds of microns (~4.2‰ in Fig. 10a, ~1.4‰ in
Fig. 10b, ~2.7‰ in Fig. 10d and ~4‰ in Fig. 10e).

Progressive distillation of the ambient sulfate reservoir
seems consistent with measurements performed in fenestral
laminae, where we observed micropyrites in GIS68.1 dis-
tributed along a profile of ~400 µm displaying $\delta^{56}$Fe values
of +2.15‰, -0.92‰, -0.26‰, -0.88‰ and -1.27‰,
respectively from the bottom to the top of the profile
(Fig. 10c). This type of distribution is not systematic in fen-
estral laminae however, as shown by the randomly dis-
tributed $\delta^{56}$Fe values (Fig. 10f). Therefore, this study does
not fully support nor exclude that intra-facies $\delta^{56}$Fe vari-
ability, recorded in micropyrites of the Tumbiana Forma-
tion, can reflect variation of the RPP during pyrite
formation.

4.2.4. Sulfidation of highly-reactive Fe phases

The subsequent formation of monosulfide precursors
(FeS) and pyrites can be promoted by sulfide mediated
reductive dissolution of Fe(III)-oxides (Berner, 1984; Dos
Santos Afonso and Stumm, 1992). The experimental

<figure><img src="image_11.png" /><figcaption>Fig. 10. SEM images of micropyrites and their corresponding $\delta^{56}$Fe values in the different lamina types. (a) and (d) Fe isotopic compositions measured in micropyrites from micritic laminae (Type 1) in GIS68.1; b) and e) Fe isotopic compositions measured in micropyrites from organic-rich laminae (Type 2) in GIS68.9; c) and f) Fe isotopic compositions measured in micropyrites from fenestral laminae (Type 3) in GIS68.1 and GIS68.9.</figcaption></figure>
sulfidation of a variety of Fe(III)-bearing phases, such as ferrihydrite, goethite, hematite and lepidocrocite produces a significant Fe isotope fractionation of –0.85‰ (Δ⁵⁶FeFe₂+aq-Fe(III)solid) and should be considered for cycling Fe in the sediment porewater (McAnena, 2011). If we cannot fully exclude abiotic sulfidation as an efficient reduction pathway of Fe(III)-oxides, this process alone is not likely to explain δ⁵⁶Fe values as low as –2.2‰.

### 4.2.5. Dissimilatory Iron Reduction (DIR)

A previous study has proposed that the Fe isotope composition of pyrites from the Tumbiana Formation can account for Dissimilatory Iron Reduction process (Yoshiya et al., 2012). Positive δ⁵⁶Fe values of pyrites can reflect the total reduction of preexisting Fe-oxides, while negative δ⁵⁶Fe values can be explained by partial reduction (Yoshiya et al., 2012). However, the recent discovery of negative Fe isotopic fractionation during abiotic Fe-oxides reduction that mimics DIR isotopic signatures (Frierdich et al., 2019) has changed the interpretation of the δ⁵⁶Fe signal. Knowing that abiotic Fe reduction can occur at similar rate than DIR in the presence of Fe-oxides and amorphous FeS (Mortimer et al., 2011), it seems difficult to firmly infer which one controlled pyrite precipitation and ensuing iron isotope signal. In the four analyzed samples, no Fe-oxide has been observed but the Fe speciation data support the presence of a non-negligible amount of Fe-oxides (between 90 ppm and 204 ppm, Table 1). The difference between partial and total iron reduction should therefore be due to organic matter availability. Indeed, the preservation of organic-matter in Type 2 laminae could be the result of partial reduction, whereas in Type 1 micritic laminae, the absence or discrete occurrence of organic matter, could reflect near-complete reduction. However, the rather comparable Fe isotopic ranges recorded in pyrites do not support the hypothesis of varying Fe reduction rate between Type 1 and 2 laminae. Thus, the preservation of organic-matter in Type 2 laminae is likely due to the ability of clays to adsorb organic-matter compared to micrite in Type 1 laminae.

Covariation between Fe and S isotopes could provide evidence for paired microbial Fe and sulfate reduction (Archer and Vance, 2006). A large S isotopic range measured in nanopyrites from organic-rich laminae (Type 2) has been previously interpreted as diagnostic for MSR (Marin-Carbonne et al., 2018). However, the Fe-S covariation proposed by Archer and Vance (2006) has not been tested at the micropyrite scale, although large isotopic ranges have been evidenced for both Fe (δ⁵⁶Fe from –1.5‰ to +4.4‰) and S (δ³⁴S from –33.7‰ to +50.4‰) for organic-rich laminae. The presence of iron enriched micrite (Fe > 0.5wt.%) in Type 1 laminae and magnetite, as indicated by iron speciation, can be either interpreted as a direct incorporation of Fe(II) provided by the dissolution of mafic minerals, or as an end-product of DIR. Hence, if a microbial iron reduction seems plausible, it cannot solely explain the whole Fe isotopic distribution recorded in the four samples.

## 4.3. An emerging view of the biogeochemical Fe cycling recorded in Archean stromatolites

### 4.3.1. Pyrite Fe isotopes and microbial activity

In situ analyses applied to stromatolites from the Tumbiana Formation allowed the identification of S, C and Fe isotopic variations (Nishizawa et al., 2010; Yoshiya et al., 2012; Marin-Carbonne et al., 2018; Lepot et al., 2019) significantly different than those obtained so far by bulk-rock analyses (Thomazo et al., 2009; Czaja et al., 2010). These micro-scale heterogeneities in both organics and pyrites promote a microenvironmental, i.e. local, control on δ³⁴Spy, δ¹³Corg and δ⁵⁶Fepy (Thomazo et al., 2009; Fike et al., 2015; Marin-Carbonne et al., 2018; Lepot et al., 2019). MSR has been previously called upon to explain S isotope heterogeneities and to be a major driver for micro- to nanometric pyrite precipitation on a local scale (Marin-Carbonne et al., 2018). Methanotroph metabolic activity, that can be aerobically mediated using oxygen (Hayes, 1994) or anaerobically (AOM for Anaerobic oxidation of Methane) using sulfates (Hinrichs, 2002), nitrates (Haroon et al., 2013) or even Fe-oxides (Beal et al., 2009), was likely active to create the large negative δ¹³C excursion recorded in organic matter (Hayes, 1994; Thomazo et al., 2009; Lepot et al., 2019). C-isotope analyses of organic matter from another Tumbiana stromatolite samples (Lepot et al., 2019), equivalent to our Type 1 (micritic) laminae and Type 2 (organic-rich) laminae, revealed highly δ¹³C-depleted signatures as low as –58.8 ± 3.6‰ (SD) in micrite and –53.5 ± 1.8‰ (SD) in organic rich horizons. Extremely δ¹³C-depleted values of organic matter from micritic laminae associated with high abundance of organic sulfur have been interpreted as the result of sulfate-thriving AOM. The less negative δ¹³Corg values recorded in organic matter from Type 2 laminae is interpreted to reflect a competition between sulfate and Fe-oxides during AOM (Lepot et al., 2019).

The three lamina types illustrate various concentration of organic-matter, terrigenous material and the effect of early diagenesis that can account for pyrite precipitation. The lower δ⁵⁶Fe values (δ⁵⁶Fe from –2.2‰ to +2.65‰) measured in micropyrites from fenestral laminae (Type 3), which are formed later compared to micritic and organic-rich laminae, can be related to the sparite filling that occured during early diagenesis. We can thus conclude that diagenetic processes affect the mean Fe isotopic composition recorded in micropyrites but the heterogeneity of δ⁵⁶Fe values is still preserved. The Fe isotopic record shows similar δ⁵⁶Fe variations and ranges between micritic and organic-rich laminae. Based on these two lamina types, the absence of variability rules out either a control of organic-matter, sulfate and Fe-oxides availability or a control of the biomineralization (higher in micritic laminae compared to organic-rich laminae) of the stromatolite on the Fe isotopic signatures. Therefore, the heterogeneity of δ⁵⁶Fe signatures seems to be primarily inherited from metabolic activities.
### 4.3.2. A complex interplay of Fe redox reactions

Based on previous knowledge on microbial processes operating in the Tumbiana environment, we propose a multi-step model of iron cycling than can account for the $\delta^{56}$Fe distribution attributed to the micropyrites. This model includes metabolisms previously proposed to be active in these stromatolites, such as MSR and AOM. While consistent with $\delta^{56}$Fe measured in micropyrites and previous geochemical studies performed in the Tumbiana Formation, some uncertainties due to processes that could significantly contribute to Fe isotope signatures must be considered in the model. These include the influence of EIE/KIE, potential isotopic fractionation during the conversion of mackinawite into pyrites or isotopic variability inherited from the Fe source. The detrital $Fe_{aq}^{2+}$ supplied by weathering of the mafic substrate (pathway 1, Fig. 11) was partially oxidized into $Fe^{3+}$, and precipitated as ferric oxide/hydroxides (pathway 2, Fig. 11). Whereas the question of the oxidant used for the $Fe(II)_{aq}$ oxidation remains open, Fe isotopic fractionations during partial oxidation mediated by $O_2$, anoxygenic photosynthesis or UV photo-oxidation are similar ranging from +0.9‰ to +1.5‰

(Bullen et al., 2001; Croal et al., 2004; Nie et al., 2017). We thus considered in the following section a mean isotopic fractionation between dissolved $Fe^{2+}$ and Fe(III)-oxides/hydroxides $Fe(OH)_3$ of +1‰. Fe(III)-oxides/hydroxides are then used as electron acceptors during the microbial reduction mediated by DIR bacteria (1) (Czaja et al., 2010) or abiotically (2) (Mortimer et al., 2011) (pathway 3, Fig. 11).

$$4Fe(OH)_3 + CH_2O \rightarrow 4Fe^{2+} + HCO_3^- + 7OH^- + 3H_2O \quad (1)$$

$$2Fe(OH)_3 + FeS \rightarrow 3Fe^{2+} + S^0 + 6OH^- \quad (2)$$

On one hand, the biotic or abiotic Fe reduction is partial in order to generate $Fe^{2+}$ depleted in $^{56}$Fe with $\delta^{56}$Fe values between -2‰ to -0.5‰. This $^{56}$Fe-depleted iron source can be used in turn to react with reduced sulfur species (e.g., $H_2S$, $S_0$) produced by MSR and hence precipitate successively as iron monosulfides and sedimentary pyrites (pathway 3a, Fig. 11). On the other hand, the partial reduction pathway allows the buildup of a residual $Fe^{3+}$-oxides/hydroxides pool which are experimentally shown to be

<figure><img src="image_12.png" /><figcaption>Fig. 11. Schematic view of microbial reactions discussed at local scale. Iron from water column is cycled several times by oxidation and reduction processes in the sediment. DIR: Dissimilatory Iron Reduction. MSR: Microbial Sulfate reduction. AOM IR: Anaerobic oxidation of methane coupled to iron reduction. Black numbers correspond to iron isotopic compositions of the different species at the first cycling. Grey numbers referred to iron isotopic compositions obtained after the second cycling. Probability density functions of the $\delta^{56}$Fe distribution calculated for 4 cumulative Fe cycling (bin size of 0.1) and the comparison with probability density function of the $\delta^{56}$Fe distribution of the 142 micropyrites are available in Electronic annex (Fig. EA-9). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>
enriched in heavy iron isotope (Welch et al., 2003). The Fe isotopic compositions of the residual Fe(OH)₃ is estimated to range between +1‰ and +2.4‰ (pathway 3b, Fig. 11). This residual Fe(III)-oxides pool can later be fully converted back to Fe²⁺, coupled for example with the activity of anaerobic methanotroph bacteria following the equation (3) (Czaja et al., 2010):

$$CH_4 + 8Fe(OH)_3 + 15H^+ \rightarrow HCO_3^- + 8Fe^{2+} + 21H_2O \quad (3)$$

The involvement of this step of complete reduction, possibly conducted by AOM-IR, allows to the preservation of the Fe isotopic range recorded by the residual Fe(III)-oxides. The conjoint activity of methanotrophs and potentially Fe reducing bacteria, leading to the formation of bicarbonate and ferrous iron, could participate to the precipitation of micrite observed in Type 2 laminae slightly enriched in iron. After reduction of Fe(III)-oxides into Fe²⁺, possibly by methanotrophs, a re-cycling of this isotopically heavy dissolved Fe²⁺ could produced:

* δ⁵⁶Fe values of Fe-oxi/hydroxides from +2‰ to +5‰ after partial oxidation (pathway 2, 2nd cycling, Fig. 11).

* δ⁵⁶Fe values of Fe²⁺ from -1‰ to +3.5‰ after iron reduction (pathway 3a, 2nd cycling, Fig. 11). Isotopically heavier Fe²⁺ is available for pyrite precursor precipitations.

The model includes episodic supplies of detrital Fe²⁺ in the sediment pore water, likely generated by the weathering of mafic volcanoclastic substrate. The distribution of δ⁵⁶Fe values measured in Tumbiana micropyrites is best modeled after four successive cycles with a partial reduction rate of 30% (Electronic Annex (EA); Fig. EA-9), which is in the range of the reduction rate reported for Dissimilatory Iron Reduction in batch experiments (Benner et al., 2002; Hansel et al., 2004). Such an intensive Fe oxido-reduction iron cycling echoes Fe and S cryptic cycles, which can occur in modern stratified lacustrine environments (e.g., Pavin lake, Cadagno lake; Berg et al., 2016; Berg et al., 2019) and oxygen minimum zone (e.g., Chilean coast; Canfield et al., 2010; Teske, 2010). Independently of the Fe concentration of the water column, the subsequent Fe(II) oxidation and Fe-oxide reduction are the result of a prevalent microbial iron cycling in the water column chemocline (Berg et al., 2016; Berg et al., 2019). Although the origin of the oxidation is still questionable for Archean stromatolites such as those from the Tumbiana Formation, this study demonstrates the importance of local redox processes controlling the multiple step Fe cycling associated with coeval microbial metabolisms that promoted pyrite precipitation.

## 5. CONCLUSIONS

Iron isotope compositions of micropyrites from the 2.724 Ga Tumbiana Formation provide a unique view of the possible iron cycle during the Archean. Based on the study of stromatolite samples, four types of laminae have been evidenced: micritic (Type 1), organic-rich (Type 2),

fenestral (Type 3) and coarse-grained (Type 4) laminae. These laminae reflect the different processes typically recognized in a stromatolite, such as carbonate precipitation (Type 1), terrigenous influence (Type 2), trapping and binding (Type 4) and cementation during early diagenesis (Type 3). Micropyrites recorded a local microenvironment and a limited impact of diagenesis. In addition to reflect a combined activity of AOM, MSR and possibly DIR, micropyrites recorded multiple stages of Fe oxidation and reduction. This study helps to better understand the iron isotopic signatures that result from multiple redox fluctuations in microbial mats. Therefore, stromatolites of the Tumbiana Formation can be considered to reflect a fossil microbial ecosystem that has recorded an exceptional diversity of microbial metabolisms. Future investigations should focus on the coupled Fe, S and C isotope evolution at the local scale in order to assess the effect of metabolism interactions on Fe isotope signatures.

## 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

Johan Villeneuve and Nordine Bouden are thanked for their help during SIMS sessions. Colette Guilbaud is thanked for her help with sample preparation. Virgil Pasquier, Elias Bloch and Mathieu Gravey are thanked for fruitful discussions. Béatrice Luais is acknowledged for providing us standard material. This research was supported by the European Union's Horizon 2020 research and innovation program ERC (STROMATA, grant agreement 759289). Pascal Philippot acknowledges support from the São Paulo Research Foundation (FAPESP grant 2015/16235-2).

## APPENDIX A. SUPPLEMENTARY DATA

Supplementary data to this article can be found online at https://doi.org/10.1016/j.gca.2021.07.020.

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Associate editor: Aubrey Zerkle