# Iron-rich microband formation in marine sediments by hydrothermal iron cycling bacteria at Lucky Strike

Jérémie Aubineau, Ernest Chi Fru, Christine Destrigneville, Thierry Decrausaz, Fleurice Parat, Fabien Baron, Alexandra Nederbragt, Manuel Henry, Alain Castillo, Céline Rommevaux, et al.

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Jérémie Aubineau, Ernest Chi Fru, Christine Destrigneville, Thierry Decrausaz, Fleurice Parat, et al. Iron-rich microband formation in marine sediments by hydrothermal iron cycling bacteria at Lucky Strike. Communications Earth & Environment, 2025, 10.1038/s43247-025-02223-2. hal-05053696

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# Iron-rich microband formation in marine sediments by hydrothermal iron cycling bacteria at Lucky Strike—

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Jérémie Aubineau Ⓓ, Ernest Chi Fru Ⓔ, Christine Destrigneville<fnref n="1" />, Thierry Decrausaz³, Fleurice Parat³, Fabien Baron⁴, Alexandra Nederbragt<fnref n="2" />, Manuel Henry<fnref n="1" />,⁵, Alain Castillo<fnref n="1" />, Céline Rommevaux Ⓔ⁶ & Valérie Chavagnac Ⓓ<fnref n="1" />

Models describing banded iron formation deposition beneath oxygenated Paleoproterozoic marine shelf waters assume that microaerophilic iron-oxidizing microorganisms were significant contributors. However, direct evidence for this mechanism is lacking, and it remains unclear how associated primary organic biomass and biogenic iron oxyhydroxides transformed into organic carbon-poor iron oxide/ carbonate-rich bands in banded iron formations. Here we explore in situ modern Zetaproteobacterial iron-oxidizing mats thriving in the Lucky Strike hydrothermal vent field to develop an empirical model that explains the potential contribution of iron-based metabolisms to deposition of rocks resembling banded iron formations. Petrographic, mineralogical, and geochemical data suggest that abundant production of Zetaproteobacterial stalks nucleates and drives spontaneous biotic/abiotic ferrihydrite precipitation in the mats. Subsequent early diagenetic anaerobic microbial oxidation of dead biomass as the mats aged, coupled to microbial reduction of iron oxyhydroxides/oxides, generates ferrous carbonates and mixed valence magnetite. Collectively, this forms crystalline iron-rich organic carbon-lean ¹³C-depleted layers reminiscent of Precambrian banded iron formation microbands.

Banded Iron Formations (BIFs) are ancient marine chemical sedimentary deposits rich in iron (Fe) and silica (Si), with peak deposition 4000–1850 Million years (Ma) ago and briefly during the 720–635 Ma Neoproterozoic Snowball Earth glaciations<fnref n="1" />⁻³. They provide key information about the chemistry of the ancient seawater from which they were formed and about the iron cycling in the early oceans. Algoma-type BIFs, often interlayered with Archean-Paleoproterozoic-aged volcano-sedimentary sequences, are interpreted to have been deposited near submarine volcanic centres via exhalative-hydrothermal processes<fnref n="1" />,<fnref n="2" />. Their distinct sedimentary and geochemical features suggest an open or partially restricted deep sea depositional setting. Algoma-type BIFs tend to be widespread in Archean cratons and sporadically appear in Proterozoic marine sedimentary environments. The deposition of Quaternary Algoma-type BIF analogue rocks on the coast of Milos, Greece⁴,⁵ suggests Algoma-type BIFs may also be associated with near continental margin shallow hydrothermal vent fields, driven by microbial iron cycling⁶,⁷. In contrast, Superior-type BIFs are typically

interbedded with Paleoproterozoic carbonates and black shales, reflecting deposition in nearshore shelf environments<fnref n="1" />,<fnref n="2" />. Algoma-type BIFs are much thinner and smaller in lateral extent than Superior-type BIFs, although the former can be more abundant in individual occurrences⁸. Rapitan-type BIFs, linked to the Neoproterozoic Snowball Earth glaciations, were formed when thick ice sheets resulted in the deoxygenation of the seawater, enabling ferruginous conditions to develop in the global ocean through hydrothermal exhalation<fnref n="2" />.

Generally, BIFs contain at least 15 wt.% bulk Fe and often show spectacular alternating millimetric (microbands) to metric silica and Fe(III) oxide or Fe(II) silicate layers<fnref n="1" />⁻³. According to a consensual model, the large-scale accumulation of Fe required that dissolved Fe(II) sourced from hydrothermal vents was oxidized by various mechanisms to insoluble Fe(III) in association with abiotic precipitation of amorphous silica<fnref n="1" />. Existing geochemical models propose that the microbands in Algoma-type BIFs likely formed through seasonal variations or microbial activity⁹. Similar

<footnote n="1">¹Géosciences Environnement Toulouse, CNRS UMR 5563 (CNRS/UPS/IRD/CNES), Université de Toulouse, Observatoire Midi-Pyrénées, Toulouse, France.</footnote>

<footnote n="2">²Center of Geobiology and Geochemistry, School of Earth and Environmental Sciences, College of Physical Sciences and Engineering, Cardiff University, Cardiff, Wales, UK. ³Géosciences Montpellier, Université de Montpellier, CNRS, Montpellier, France. ⁴Université de Poitiers, IC2MP CNRS, Poitiers, France. ⁵CEFREM, Université de Perpignan Via Domitia, UMR 5110, CEDEX, 52 Avenue Paul Alduy, Perpignan, France. ⁶Aix-Marseille University, Université de Toulon, CNRS, IRD, MIO, Marseille, France. ✉e-mail: jeremaubineau@aol.com</footnote>
microbial mechanisms have been suggested for the deposition of Qua-
ternary shallow submarine hydrothermal vent field analogues⁴,⁵. However,
there is a notable lack of in situ studies that detail the pathways of Fe(II)
oxidation leading to the formation of primary amorphous Fe(III) minerals
in BIFs and the processes involved in the sequential transformation of these
minerals into highly crystalline Fe(III) phases with low organic carbon
content, remain understudied. Here, we provide a mechanism from a
modern environment that could support a potential biological origin for
Algoma-type BIF-like microbands in a modern deep-sea hydrothermal
setting.

Planktonic (i) O₂-producing photosynthetic cyanobacteria in the Proterozoic Ocean, coupled with abiotic Fe oxidation, or chemolithoautotrophic Fe(II)-oxidizers, and (ii) Archean anoxygenic photoautotrophic Fe(II)-oxidizing bacteria (photoferrotrophy)⁶,¹⁰⁻¹³, are among some of the most prominent mechanisms proposed for Precambrian BIF deposition. For example, microaerophilic Fe(II)-oxidizing microorganisms (FeOM), including *Zetaproteobacteria*, that likely evolved to dominate pelagic microaerobic conditions following the Great Oxidation Event (GOE), ca. 2500–2100 Ga¹⁴,¹⁵, are suggested to have produced a “ferrihydrite rain” to the Precambrian seafloor². In contrast, a consortium of benthic microbial communities may have formed discrete microbands of iron oxides in stromatolitic facies¹⁶. However, it remains unclear how these processes and benthic microbial communities contributed to the generation of BIF bands.

Although present in varying amounts, iron oxides in Archean and Proterozoic BIFs are dominated by magnetite (Fe(III)₂Fe(II)O₄) and haematite (Fe(III)₂O₃), respectively¹⁷. Both are interpreted as having formed from an initial poorly crystalline Fe(III)-oxyhydroxide precursor phase, such as ferrihydrite¹,³ (Fe – Fe(III)(OH)₃), suggesting the iron oxide mineralogy of Precambrian BIFs is likely secondary rather than primary. Alternatively, greenalite (Fe(II)₃Si₂O₅(OH)₄) formed under oxidant-depleted environmental conditions, has been suggested as an alternate primary Fe precursor in BIFs¹⁸⁻²⁰ prior to the GOE. Based on this theory, iron oxides in Archean BIFs are secondary products formed following post-depositional oxidation of the primary ferrous silicates²¹,²². In addition, green rust, a ferrous-ferric hydroxy salt, which would have precipitated near the Fe redoxcline and transformed into stable mineral phases (e.g., magnetite) within the water column and sediments, has been suggested as a precursor to BIF mineral assemblages²³. Consequently, there is no consensus on the dominant Fe mineral precursors to BIFs. One reason for these disagreements arises from the scarcity of modern analogues of banded Fe-Si-rich precipitates from diverse marine settings mimicking the Precambrian Ocean to enable the appropriate testing and validation of these various hypotheses.

Moreover, the low organic matter content of BIFs has been used to
dismiss the potential direct involvement of biological activity in the pre-
cipitation of BIF Fe(III)-rich minerals. However, near-complete respiration/
oxidation of organic matter coupled to Fe(III) reduction by various bacteria
and archaea are frequently hypothesized to explain the low organic C
content (<0.8 wt.%) in BIFs¹,¹²,²⁴. This view is anchored on intimately
associated magnetite and ¹³C-depleted Fe(II)-bearing carbonates (e.g.,
siderite), and the general absence of definitive microfossils preserved in the
non-silicified Fe-rich bands¹,²⁴. However, direct (i.e., textural) evidence for
biological activity provided by microfossils and stromatolitic oncoidal
structures associated with granular IF deposition has been recently
documented²⁵⁻²⁷. Therefore, understanding the complex interplay between
BIFs, ocean chemistry, and the biosphere to unravel the mechanisms
responsible for the extensive generation of organic C-lean BIFs, remains an
ongoing endeavour as are the pathways involved in the transformation of
primary Fe-bearing mineral precursors into crystalline phases.

To provide new insights into these primary and secondary BIF for-
mation pathways, we focus on present-day active submarine hydrothermal
ecosystems displaying high dissolved Fe fluxes associated with chemo-
lithoautotrophic microaerophilic FeOMs, typical of the mid-ocean ridges,
back-arc spreading centres, and submarine arc volcanoes²⁸. There, hydro-
thermal fluid discharge produces chemical and redox gradients that enable

the development of a variety of microbial mats along cracks and fractures on
various seafloor lithologies²⁹⁻³⁵. Although deep-sea hydrothermal environ-
ments have been widely studied for their microbial composition and
environmental conditions, there is a dearth of knowledge on their potential
association with the biological development of Precambrian BIFs³⁵⁻³⁹. We
therefore collected in situ modern Fe-rich microbial mats from the Lucky
Strike Hydrothermal Field (LSHF; 37°17′N & 32°16′W; Mid-Atlantic
Ridge⁴⁰) (Fig. 1; Table 1; Supplementary note 1), to investigate a potential
linkage and association. At the seafloor, white and orange microbial mats,
linked to sulphur oxidation³⁰ and Fe oxidation²⁹,⁴¹, respectively, flourish on a
heterogeneous substratum in close vicinity of diffuse vents at temperatures
<100 °C⁴²,⁴³ from different well-known sites, including Capelinhos (CAP),
North Tour Eiffel (NTE), South Isabel (SI), West Sintra (WS), and Y3.

Given the relatively high Fe(II) concentrations in Archean oceans (µM then versus nM today)¹⁴, it is commonly accepted that deep-marine hydrothermal vents in the Archean delivered tremendous amounts of Fe(II) and other reduced compounds that modulated nutrient cycling (e.g., C, N, S, Fe) and contributed to BIF deposition¹,⁴⁵. The LSHF mats host abundant microbial communities, including autotrophic Fe(II)-oxidizing *Zetaproteobacteria* coexisting with N, S, and methane-cycling microorganisms²⁹,⁴¹ (see Supplementary note 1). *Zetaproteobacteria* are the presumed primary producers in iron-rich mats flourishing in diverse hydrothermal environments⁴¹,⁴⁶ as they are obligate chemoautotrophs, oxidizing Fe(II) to gain energy. Thus, they are essential for the development of ecological niches within the mats produced in Fe(II)-rich hydrothermal vent fields.

Considering that metabolic pathways within such nutrient cycles (i.e., C, N, S, Fe) emerged and diversified in the Archean⁴⁵, including microaerophilic microbial Fe oxidation⁴⁷, we couple petrographic and mineralogical observations with C-N systematics in the LSFH mats to validate primary and secondary microbial processes potentially involved in BIF microbands. We discuss the possibility that similar mat-forming FeOMs could have contributed to the generation of microbands in BIFs that formed close to hydrothermal vent field settings, for example, Algoma-type BIFs, under constrained microaerobic environmental depositional conditions and hydrothermal exhalation of Fe(II)-rich fluids.

Results

Texture and mineralogy of microbial mats

Except for the Lava Lake (LL) site where no diffuse venting was observed, two types of microbial mats, dominated by crumbly and flocculent structures, flourished in the vicinity of diffuse hydrothermal fluid venting at CAP, NTE, SI, WS, and Y3 (Fig. 1)⁴⁸⁻⁵⁰. Coherent and porous deposits with variable thickness of <20 cm, together with occasional dark-coloured laminated structures typify the crumbly microbial mats located at the CAP, NTE, and Y3 sites (Figs. 1, 2a–i, 3a–g, and 4a–i; Table 1). The mats at the SI, WS, and LL sites are composed of a mixture of flocculent material associated with microorganisms and inorganic particles (Fig. 1; Table 1). All the sampled Fe-rich mats have been previously studied for their *Zetaproteobacterial diversity and microbial filament morphology*²⁹.

Based on petrographic observations, the CAP mats dominantly exhibit mm-scale outward-convex structures with internal concentric bands consisting of a more or less porous network of filamentous structures dominated by stalks (Fig. 2b–h), a typical feature linked to FeOMs⁵¹. The twisted stalks, among other morphologies of filaments in the CAP mats²⁹, are thought to reflect dynamic microbial mat growth features¹³,⁵². X-ray diffraction (XRD) patterns further show two broad peaks centred at ~2.70–2.90 Å and ~1.51 Å (Fig. 5; Supplementary Fig. 1), which deviate from the pure 2-line Fh peaks characterized by d-spacings at 2.54 and 1.47 Å⁵³,⁵⁴. This shift, together with the poorly resolved peak at ~1.51 Å, the energy-dispersive spectra (EDS) analysis (Fig. 2i), and infrared-based spectroscopy (Supplementary Figs. 2, 3 & Supplementary note 2), point to poorly crystalline Si-bearing Fh⁵⁵,⁵⁶. The combined datasets highlight that the Fe-rich microstructures are heavily mineralized by Si-bound Fh.

Contrary to the CAP site, mat growth patterns are not visible at the
NTE site. Instead, micronodule structures of <120 µm in diameter and
angular minerals from the altered substratum compose the mats (Fig. 3b–f). Results from XRD, Mössbauer and infrared-based spectroscopies, and EDS analyses highlight the widespread occurrence of poorly crystalline Mn oxides (todorokite, birnessite, and pyrolusite), anorthite, and amphibole (edenite), and Si-bound Fh as the main Fe(III)-rich mineral (Figs. 3g, 5; Supplementary Figs. 1–3; Supplementary Table 1).

The Y3 mats comprise two Fe-rich microbands composed of multiple laminated and upward-convex layers (Fig. 4a, b), probably reflecting successive steps of mat growth⁵². Petrographic observation of the bottom most part of the mats is distinctively represented by a mm-thick black band consisting mainly of siderite and magnetite that appear as isotropic octahedral crystals showing a higher reflectance relative to other Fe(III)-rich minerals (iron oxides and well-crystallized Fe(III)-oxyhydroxides are referred to as FeOx) (Figs. 4c–e, 5; Supplementary Fig. 1). Towards the top of the mats, the black laminae are succeeded by alternating dark and light brown laminated layers composed of well-crystallized FeOx of various crystal shapes, including goethite [α-Fe(III)OOH], lepidocrocite [γ-Fe(III)OOH], and akageneite [β-Fe(III)OOH-Cl] (Figs. 4f–h, 5; Supplementary Figs. 1–3 & Supplementary note 2). In contrast to the other LSHF mats, little Si is bound to these FeOx in all laminae as suggested by the lack of an EDS Si peak, high EDS Fe/Si ratios, and the absence of Fe-O-Si stretching (Fig. 4i; Supplementary Figs. 3, 4). In addition, Mössbauer spectroscopy was employed to confirm the presence of Fh in the Y3 mats (Fig. 5; Supplementary Table 1 & Supplementary note 2). Similar to CAP mats and other natural biofilms⁵⁷, microbial filaments predominantly composed of stalks, coral-like, and Y-shaped structures at Y3 associated with *Zetaproteobacteria*²⁹, are thickly coated with poorly crystalline Fh. From top to bottom, the Y3 mats contain Fh-encrusted microbial filaments that disappear at the expense of increasing enrichment of well-crystallized FeOx minerals and ferrous carbonates (Fig. 6).

The Fe mineralogy of the flocculent mats at the WS, LL, and SI sites is dominated by both Si-Fh and Fh phases (Fig. 5; Supplementary Figs. 2, 3; Supplementary Table 1). Neither Si-rich bands nor amorphous silica phases like opal-A, common in modern marine sediments and defined by a broad peak centred at ~4.10 Å (~25.2° 2θ CoKα)³¹,³²,³⁴,⁵⁸,⁵⁹ were evidenced by XRD and petrographic observations in the LSHF mats. Finally, unidentified organic molecules are associated with the Fh-dominated mats (Supplementary Fig. 3).

## Bulk geochemistry

Bulk mat geochemistry at LSHF shows heterogeneous chemical major element composition, with Fe₂O₃ content varying between 10.8 and 99.6 % (Supplementary Table 2 & Supplementary note 2). At the NTE site, elevated concentrations of Al₂O₃ and CaO ranging from 9.4–13.3 % and 4.2–6.2 %, respectively, compared to other sites with <3.3% Al₂O₃ and <1.6% CaO, are related to basaltic-derived alteration fraction composed of anorthite and amphibole. P concentrations vary considerably in function of Fh/FeOx mineralogy, where very low (<0.04 %) P content is associated with crystalline FeOx phases and high (0.15–0.72 %) P concentrations with poorly crystalline Fh (Supplementary Fig. 5), reflecting a 3- to 18-fold decrease from crystalline FeOx to poorly crystalline Fh.

X-ray fluorescence (XRF) analysis of the LSHF mats reveals SiO₂ concentrations in the range of 3.9-38.1% (Supplementary Table 3), the lowest in the Y3 mats, moderate in the CAP, WS, and LL mats, and highest in the SI and NTE mats. XRF findings agree with the above mineralogical and petrographic results and support the potential occurrence of opal-A at the SI site. Moreover, the observations indicate a limited presence of Si-bound FeOx minerals at Y3, the presence of silicate minerals at NTE, and the general lack of amorphous silica phases at LSHF.

<figure><img src="image_4.png" /><figcaption>**Fig. 1 |** Bathymetric map of the Lucky Strike hydrothermal field and images of the sampled iron-rich microbial mats, modified from Chavagnac et al.⁶⁰ and Astorch-Cardona et al.²⁹. The studied mats and their associated active</figcaption></figure>

hydrothermal vents are shown. No hydrothermal vent was detected at LL site. Chlorinity increases from CAP to LL sites. Scale bars: 5 cm. [Capelinhos (CAP); North Tour Eiffel (NTE); South Isabel (SI); West Sintra (WS); Lava Lake (LL)].
<table><caption>Table 1 | Specific features of the studied Fe(II)-oxidizing microbial mats from Lucky Strike Hydrothermal Field. For further details, see Supplementary discussion concerning the high positive δ¹³C~carb~ values</caption><thead><tr><th>Site</th><th>Year</th><th>Dive</th><th>Depth (mbsl)</th><th>Distance from diffuse hydrothermal vent (m)</th><th>Thickness (cm)</th><th>Mat type</th><th>Colours</th><th>Abundance of Zetaproteobacteria (%)†</th><th>Mineralogy of the mat</th><th>Substratum</th></tr></thead><tbody><tr><td>South Isabel</td><td>2021</td><td>1998-10</td><td>1699</td><td>No diffuse flow</td><td>-</td><td>Flocculent</td><td>orange-<br>dark brown</td><td>28.68</td><td>Ferrihydrite, barite, opal-A?</td><td>Hydrothermal slab#</td></tr><tr><td>N37° 17.336<br>W32° 16.628</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>West Sintra</td><td>2021</td><td>1997-9</td><td>1669</td><td>No diffuse flow</td><td>-</td><td>Flocculent</td><td>reddish brown</td><td>44.95</td><td>Si-ferrihydrite, ferrihydrite</td><td>Massive hydrothermal deposits dominated by pyrite and marcassite</td></tr><tr><td>N37° 17.500<br>W32° 16.570</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Lava Lake</td><td>2021</td><td>2000-12</td><td>1741</td><td>No vent</td><td>-</td><td>Flocculent</td><td>orange-<br>yellowish brown</td><td>33.21</td><td>Si-ferrihydrite</td><td>Basalt</td></tr><tr><td>N37° 17.549<br>W32° 16.790</td><td>2023</td><td>844-6</td><td>1738</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Capelinhos</td><td>2022</td><td>2035-8</td><td>1669</td><td>Immediately below</td><td>20</td><td>Crumbly</td><td>orange-yellowish<br>brown to<br>reddish brown</td><td>21.55</td><td>Si-ferrihydrite</td><td>Hydrothermal slab#</td></tr><tr><td>N37° 17.360<br>W32° 15.829</td><td>2023</td><td>844-6</td><td>1669</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>North Tour<br>Eiffel</td><td>2022</td><td>2033-6</td><td>1689</td><td>&lt;1</td><td>&lt;10</td><td>Crumbly</td><td>orange-yellowish<br>brown, brown,<br>and black</td><td>25.07</td><td>Si-ferrihydrite, Mn oxides,<br>anorthite, amphibole</td><td>Hydrothermal slab#<br>composed of sphalerite,<br>marcassite, and barite</td></tr><tr><td>N37° 17.353<br>W32° 16.535</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Y3</td><td>2022</td><td>2034-7</td><td>1727</td><td>&lt;1</td><td>10-15</td><td>Crumbly</td><td>light brown to<br>orange-brown</td><td>11.69</td><td>Ferrihydrite, goethite,<br>lepidocrocite, akaganite,<br>magnetite, siderite, sulphurs</td><td>Hydrothermal gravitational<br>deposits; aggregates of iron<br>sulfide minerals and barite</td></tr><tr><td>N37° 17.518<br>W32° 16.663</td><td>2023</td><td>846-8</td><td>1727</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr></tbody></table>

All data with the exception of thickness and colours of the mats are from Matabos & Sarrazin⁴⁸, Sarradin & Matabos⁴⁹; Matabos¹⁰. The mineralogy of the LSHF mats is based on our integrated study (petrographic, XRD, and Mössbauer analyses). Salts are not included. †, Zetaproteobacterial relative abundance within the microbial mats⁵⁰, #, layered silicified volcaniclastic deposits; mbsl, metres below sea level.
<figure><img src="image_5.png" /><figcaption>**Fig. 2** | Morphological, textural, and elemental image of Capelinhos (CAP A) mat. **a** Brown to orange Fe(II)-oxidizing microbial mats. **b** Thin section. The Fe mineralogy exclusively consists of ferrihydrite (Fh), as demonstrated by XRD results of our study. **c**, **d** Transmitted light photomicrographs (plane-polarized light).</figcaption></figure>

Filamentous microbial structures. **e–h** SEM images in BSE mode showing the aggregation of a large number of filaments (mainly stalks here). Filaments crossing with each other to form a mesh structure. **i** EDS analysis of stalks displaying a predominantly Fe- and Si-rich composition.

## Spatial chemical concentration of Fe-bearing mineral phases

Electron microprobe analyses (EPMA) align with the XRD mineralogy (Supplementary Table 4 & Supplementary note 2). Moreover, EPMA chemical compositions of Fh-encrusted filamentous microstructures and crystalline FeOx show three distinct groups in the crumbly mats (Supplementary Table 5). Considering the Si concentrations of end-member hydrothermal fluids are of the same range of concentration (12.7 to 18 mM) at LSHF$^{41,60}$, Fh in CAP and NTE sites show an average $Fe_2O_3$/SiO$_2$ ratio of $3.5 \pm 1.2$ ($n = 16$; $1\sigma$), and $6.7–7.2$ ($n = 2$) in the filaments of the Y3 mats, whereas a considerably higher ratio of $90.6 \pm 46.6$ ($n = 17$; $1\sigma$) is recorded in the crystalline FeOx phases at Y3. The latter observation clearly aligns with electron distribution maps and infrared-based spectroscopy, illustrating stalks enriched with Si-Fh phases in the CAP mats containing quantitatively higher $SiO_2$ content and to a lesser extent $P_2O_5$ compared to the filamentous microfossils and the more crystalline FeOx minerals in Y3 mats (Fig. 7).

## Carbon and nitrogen systematics

The C isotopic compositions of carbonates ($\delta^{13}C_{carb}$) and organic matter ($\delta^{13}C_{org}$) and N isotope ($\delta^{15}N$) values in the LSHF Fe-rich mats differ considerably between and within sites and samples (Table 2). These mats contain low 0.18-0.83 wt.% total organic C (TOC) and 0.02-0.07 wt.% total N (Table 2). $\delta^{13}C_{carb}$, $\delta^{13}C_{org}$, and $\delta^{15}N$ values vary from -6.0 to +11.5‰ ($n = 11$), -35.1 to -18.9‰ ($n = 11$), and -1.4 to +3.1‰ ($n = 4$), respectively. The siderite-bearing Y3 mats show negative to slightly positive $\delta^{13}C_{carb}$ values, while siderite-free mats (SI, LL, and CAP) display both negative and highly positive $\delta^{13}C_{carb}$ values. Besides, TOC contents, $\delta^{13}C_{carb}$, $\delta^{13}C_{org}$, and $\delta^{15}N$ values decrease from top to bottom with increasing siderite abundance in the Y3 B mat (Table 2). Finally, P is moderately correlated with TOC ($R^2 = 0.50, n = 6, p < 0.12$) (Supplementary Fig. 6a), whereas a weak covariation between $\delta^{13}C_{org}$ and $\delta^{13}C_{carb}$ ($R^2 = 0.26, n = 11, p < 0.109$) and a strong one between $\delta^{13}C_{org}$ and $\delta^{15}N$ ($R^2 = 0.74, n = 5, p < 0.054$) are observed (Supplementary Fig. 6b, c).
<figure><img src="image_6.png" /><figcaption>Fig. 3 | Morphological, textural, and elemental image of NTE (NTE B) mat. a Brown to yellowish Fe(II)-oxidizing microbial mats with mm-thick dark laminae. The sample is extremely crumbly. b Thin section. Note the absence of specific structures. c Transmitted light photomicrograph (plane-polarized light). d–f SEM images in BSE mode showing the presence of numerous minerals of various shapes</figcaption></figure>

and sizes. The Fe mineralogy exclusively consists of ferrihydrite (Fh), as suggested by the Mössbauer spectroscopy. The identification of Mn oxides and silicates was carried out with EPMA. g EDS analysis of Fh displaying an Fe- and Si-rich composition.

## Discussion

### Microbial activity

*Carbon fixation and oxidation:* Microaerophilic Fe(II)-oxidizing *Zetaproteobacteria*, grow at low O₂ concentrations (5–20 µM⁶¹) where they compete with abiotic oxidation of Fe(II) due to the slower oxidation rate in such conditions⁶²,⁶³. They use the Calvin–Benson–Bassham (CBB) pathway to autotrophically build biomass⁶⁴ according to the equation⁶⁵:

$$6\text{Fe}^{2+} + 0.5\text{O}_2 + \text{CO}_2 + 16\text{H}_2\text{O} \rightarrow [\text{CH}_2\text{O}] + 6\text{Fe}(\text{OH})_3 + 12\text{H}^+$$

The variable δ¹³Corg values of -35.1 to -18.9‰ in the LSHF mats likely point to multiple autotrophic C fixation pathways⁶⁶,⁶⁷, and potential respiration of generated organic matter by Fe(III)-reducing bacteria within the mats (Fig. 8a), connecting the Fe cycle with carbon and nitrogen cycling (see Supplementary discussion). In this regard, most of our δ¹³Corg values are consistent with C fixation through the CBB cycle pathway that fractionates Corg isotopes by 20 to 30‰⁶⁶, considering a δ¹³C value of the dissolved inorganic C close to 0‰. In support of this observation, the Lō'ihi seamount mats in the Pacific Ocean, containing the RubisCO gene (key enzyme in the CBB cycle)⁶⁴ and *Zetaproteobacteria*, have shown δ¹³Corg values ranging from -28 to -23‰⁶⁸, which further strengthens the suggestion that

*Zetaproteobacteria* contributed significantly to the δ¹³Corg values bracketing between -20 and -30‰ in our samples and that they are foundational to the mat building process. In contrast, the reductive citric acid and 3-hydroxypropionate C fixing cycles produce δ¹³Corg values > -20‰⁶⁶. Except for one sample, our δ¹³Corg values < -20‰ suggest that the latter autotrophic mechanisms are not dominant in the LSHF mats. Further, the anaerobic C fixation via the reductive acetyl-coenzyme A or Wood–Ljungdahl pathways produces δ¹³Corg values < -30‰⁴⁶,⁶⁶, similar to those observed in some Y3 mats. Although the extent of methanotrophic microorganisms is limited in the LSHF mats²⁹,⁴¹, they could have also contributed to the lowest δ¹³Corg values due to the ¹³C-depleted methane hydrothermal source⁶⁹ and large isotopic fractionation of up ~30‰ during methanotrophy⁷⁰.

### Mechanism of microbial Fh precipitation.
*Zetaproteobacteria* oxidize Fe(II) on mat surfaces, leading to even greater O₂ depletion resulting in the development of anaerobic heterotrophic microbial communities deeper in the mats. This is consistent with increasing oxidation of organic C by dissimilatory Fe(III)-reducing bacteria, resulting in a much lower residual biomass content and formation of siderite and magnetite deeper in the Y3 mat, together with more negative δ¹³Corg and δ¹³CCarb values⁷¹ (Fig. 8a).
<figure><img src="image_7.png" /><figcaption>**Fig. 4** | Morphological, textural, and elemental image of Y3 (Y3 B) mat. a Brown to orange Fe(II)-oxidizing microbial mats with darker bands (arrows). b Thin section. Note the micrometric laminated structures with darker layers towards the bottom of the mat. c, d Reflected light photomicrograph (plane-polarized light) of the dark layers and related EDS analysis. Close association of magnetite, a grey-coloured mineral with a higher reflectance than other Fe(III)-rich minerals, and</figcaption></figure>

siderite. e-i SEM images in BSE mode and related EDS analysis showing crystal shapes of crystalline Fe oxides and oxyhydroxides (here referred to as Fe oxides for simplicity). Highly aggregated octahedral crystals in (e) colloform texture with radiating fibrous morphologies in (f) and (g) and lath-like crystals in (h). Crystalline Fe(III)-bearing minerals mainly consist of goethite and lepidocrocite in (f-h), as suggested by EPMA.

Based on mineralogical and textural observations, the ubiquitous occurrence of Fh, together with abundant biogenic filaments in the studied mats, indicate initiation of poorly crystalline Fh precipitation by *Zetaproteobacteria*, with the filaments acting as nucleation sites<sup>72</sup>. Moreover, the stalk-/sheath-forming *Zetaproterobacteria* have been specifically linked with the LSHF mats<sup>29</sup>, similar to many Fe-Si microbial deposits reported near modern hydrothermal vents<sup>31,32,34,58,59,73</sup>, which are believed to reflect a response to changing Fe(II) and O<sub>2</sub> redox gradients<sup>52</sup>. Besides, the widespread association of filamentous morphologies in the LSHF mats has been suggested to represent a growth strategy that enables *Zetaproteobacteria* to escape permanent entombment and asphyxiation by Fh/FeOx precipitates<sup>29,52</sup>. Consequently, the resulting encrustation of abandoned filaments by Fh leads to exquisite preservation of mat growth patterns (except for the NTE site where substratum alteration occurs in this area<sup>43</sup>). Moreover, nucleated Fh triggers additional autocatalytic oxidation of Fe(II) in the absence of biomass, further templating and scavenging abiotic Fh from seawater<sup>3,74</sup>.

## Sequential deposition of iron oxyhydroxide phases

**Fh reactivity and environmental conditions.** Because of their high specific surface area and sorption capacity, poorly crystalline nm-size

Fe(III)-oxyhydroxide (Fh) particles are highly reactive compared to other crystalline FeOx (e.g., goethite and magnetite)<sup>53</sup>. However, surface coating of Fh by organic molecules, silicate, phosphate, and trace metals may slow down or inhibit transformation into crystalline FeOx<sup>3,54,74,75</sup>, thereby enhancing its preservation<sup>54,76</sup>. In this regard, our data show an abundance of organic molecules (infrared-based analyses; Supplementary Fig. 3), SiO<sub>2</sub>, and P<sub>2</sub>O<sub>5</sub> in Fh-dominated mats compared to the crystalline FeOx phases in the Y3 mats. In agreement with the accepted consensus that dissolved aqueous Si favours direct adsorption to Fh even at extremely low Si concentrations<sup>55,77</sup>, the indurated Fh-encrusted microbial filaments at LSHF host high Si concentrations sourced from active hydrothermal vents. Further, our mineralogical observations indicate that Si tends to associate intimately with poorly crystalline Fh rather than form separate Fh and Si minerals at the surface of the mats. Similar to Si-Fh mineral associations reported in other deep-sea hydrothermal vent settings<sup>e,g,34</sup>, environmental factors such as fluid temperature and pH likely influenced the binding capacity of Fh for Si and various elements<sup>78</sup>. For example, an increase in Fh crystallite size lowers the proportion of atoms present at the mineral surface, resulting in a proportionate reduction in the mineral's specific surface area available for adsorption and/or scavenging of dissolved trace elements<sup>54,79</sup>. No
<figure><img src="image_8.png" /><figcaption>**Fig. 5** | Representative XRD patterns of randomly oriented powders and Mössbauer spectra from the LSHF iron-rich microbial mats. Diagnostic XRD spacings (Å) are shown. Although the presence of crystalline minerals makes amorphous to poorly crystalline phases difficult to detect using XRD, the isomer shift (IS) and quadrupole splitting (QS) are consistent with reported values for Fh at SI, NTE, and</figcaption></figure>

Y3 sites (see Supplementary Note 2). The light and dark red shaded regions indicate the center of broad peak areas of Si-Fh and Fh, respectively. Note that a diffractometer using MoKα radiation was used to analyse Fe- and Mn-containing mats at the NTE site. [Ferrihydrite (Fh); room temperature (RT); diffraction peak assigned to the sample holder (*)].

correlations were observed between the pH of the diffuse hydrothermal fluids and Fh/FeOx mineralogy at LSHF (pH ranges from 5.1 to 6.1 at CAP and 5.6 to 5.8 at Y3)⁴². In contrast, persistent 90 °C diffuse fluids at Y3 compared to temperatures of ~50 °C at the CAP and NTE sites⁴² would have promoted nucleation of slightly larger Fh particles at the former⁷⁸ (Fig. 8b). This phenomenon likely decreased Fh surface area reactivity, adsorption capacity for Si, P, and other elements, and likely enhanced transformation to secondary FeOx mineral phases solely in the Y3 mats.

**Transformation of primary Fh to secondary Fe(III) and Fe(II) mineral phases.** Different environmental conditions (e.g., temperature of diffuse fluids) at LSHF would be expected to produce different Fh transformation products. For example at the Y3 site, stalk-forming *Zetaproteobacteria* and organic molecules facilitate the nucleation of biogenic Fh⁵⁷. As the mats aged, thermodynamically unstable Fh eventually transformed into more crystalline FeOx phases associated with residual filamentous microbial structures, as evidenced by petrographic textures (Fig. 6). Haematite formation from Fh occurs rapidly at temperatures >50 °C or after extended reaction times⁸⁰, while goethite/lepidocrocite is an intermediate mineral formed as a function of temperature, pH, and Fe(II) concentrations during the transformation of Fh to haematite⁷⁴,⁸⁰. As a consequence, the LSHF elevated Fe(II) concentrations⁴¹,⁶⁰ would have preferentially transformed Si-free Fh into

goethite and lepidocrocite³,⁷⁴ (Fig. 8b), which explains the lack of haematite. This is supported by the fact that goethite may require several years/centuries (or even more) to form haematite under suitable environmental conditions, as observed in Quaternary shallow marine hydrothermal deposits associated with microbial filaments⁴⁻⁶.

We observe a unique transformation of Fh/FeOx to siderite and magnetite in the bottommost part of the Y3 mat, which could result from both biotic and abiotic processes⁵. Siderite and magnetite do not appear to have replaced or overgrown an earlier Fe-rich phase. Magnetite could have formed from the conversion of Fe(III)-oxyhydroxides with the addition of high amounts of hydrothermally derived Fe(II)⁷⁴. However, petrographic observations and all isotopic evidence suggest that they are a product of early diagenetic Fh/FeOx transformation rather than primary growth (Fig. 8b). We use C isotope systematics to decipher the origin of both minerals²⁴, considering the variable δ¹³Ccarb values could have originated from dissolved inorganic C in hydrothermal-seawater fluid mixtures because δ¹³C values of hydrothermal vent fluids are systematically < −1‰⁶⁹. However, the lack of correlation between Eu/Eu* anomaly as a tracer of hydrothermal contribution⁸¹,⁸² and δ¹³Ccarb values (R² = 0.03, n = 6, p < 0.75) (Supplementary Fig. 7; Supplementary Table 2) indicates that alternative mechanisms controlled the δ¹³Ccarb values in the mats. Furthermore, hydrothermal fluid influence cannot explain the unique presence of siderite in the bottom layer of the Y3 mats, seeing that our samples covered several sites associated with hydrothermal activity (except for the LL site).
<figure><img src="image_9.png" /><figcaption>**Fig. 6** | **Micro textures of Y3 (Y3 B) mat.** SEM-BSE images of microbial filaments showing an in situ model of increasing enrichment and crystallization of sheaths and stalks with FeOx from the upper (a, b), middle (c–g), to the bottom (h), parts of the Y3 B mat. As the mat aged, the microbial filaments are usually encrusted by poorly crystalline ferrihydrite (Fh)⁵⁷ first act as nucleation sites, followed by gradual</figcaption></figure>

transformation into well-crystallized FeOx that cover them with successive steps of FeOx crystallization (b–f), and eventually disappear (g–h). The orange arrow reported in (d) shows the decreasing proportion of filamentous structures with increasing depth.

Instead, the transition to lower δ¹³Ccarb, with δ¹³Corg depletion and decreasing TOC content deeper in the Y3 mat (Supplementary Fig. 6b; Table 2), as shown by the low δ¹³Ccarb and δ¹³Corg values of +2.0‰ and -21.4‰ in the bottom part, respectively, compared to +11.5‰ and -19.0‰ at the mat surface, respectively, point to a biotic origin⁷¹ (Fig. 8). This largely supports a coupling of early diagenetic microbial reduction of Fh/FeOx to the oxidation of residual organic C¹,¹²,²⁴ to produce ¹³C-depleted biomass

and siderite, and double valence magnetite. This process is also consistent with the 3- to 18-fold reduction of P concentration in the FeOx-rich layer compared to the Fh-rich mats via the reductive release of organic-bound P into seawater and sediment pore water to sustain local primary productivity. Furthermore, the weak correlation between δ¹³Ccarb and δ¹³Corg values in the LSHF mats implies that biologically mediated organic C oxidation promoted the release of CO₂ into pore waters, during which some combined
**Fig. 7** | SEM-BSE images showing features of the Fe(II)-oxidizing mats from CAP (CAP A; left panels) and Y3 (top of the Y3 B mat; right panels) sites, and corresponding EPMA element mapping (viridis colour map). Fe₂O₃, SiO₂, and P₂O₅ compositional variations revealed by EMPA maps. SiO₂ and P₂O₅ concentrations in Fh-dominated CAP mat show higher proportions than those in crystalline FeOx from the Y3 mat. Red arrows denote Si-poor lath-like minerals, likely goethite/lepidocrocite, compared to Fh-encrusted microbial filaments (white arrows). Scale bars are 30 µm.

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

with Fe(II) to form siderite and the remainder was lost to the overlying water column. The above mechanism for siderite and magnetite formation also aligns with the presence of *Geospsychobacter* Fe(III)-reducing bacteria and other archaeal Fe(III) reducers in the LSHF mats²⁹.

## Implication for microband formation in BIFs

**Role of benthic microbial contribution to Fe-rich microband formation.** Our approach provides a detailed modern understanding of interactions between benthic microbial activity, environmental conditions, and Fe-rich minerals where Fe-rich microbands reminiscent of BIFs formed solely in the Y3 mats. It has been suggested that

microorganisms played a significant role in the formation of microbands in BIFs⁹, a proposition challenged by the low organic C content of <0.8 wt.% in BIFs¹,¹²,²⁴. While a mean TOC of ~5 wt.% has been extrapolated from C isotopes, magnetite and siderite mass balance studies of BIF deposits¹, and suggestions made that this organic matter could have been consumed by post-depositional microbial oxidation, direct evidence remains lacking. Our data suggest that microbial-driven siderite and magnetite formation can occur at even much lower organic C content in BIFs. While it is unclear to what extent benthic mats may have contributed to BIF deposition, our observations suggest that the low TOC content in such BIFs, particularly Algoma-type BIFs, may be explained by
<table><caption>**Table 2 | Carbon and nitrogen data**</caption><thead><tr><th>Site</th><th>Sample ID</th><th>Year</th><th>Dive</th><th>Depth (mbsl)</th><th>δ<sup>13</sup>C<sub>carb</sub> (‰, VPDB)</th><th>δ<sup>13</sup>C<sub>org</sub> (‰, VPDB)</th><th>δ<sup>15</sup>N (‰, air)</th><th>TOC (wt.%)</th><th>TN (wt.%)</th><th>Siderite§</th></tr></thead><tbody><tr><td>South Isabel</td><td>M21-SI bulk</td><td>2021</td><td>1998-10</td><td>1699</td><td>-5.0</td><td>-21.8</td><td>3.1</td><td>0.32</td><td>0.03</td><td>No</td></tr><tr><td>N37° 17.336<br>W32° 16.628</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Lava Lake</td><td>M23-LL bulk</td><td>2023</td><td>844-6</td><td>1738</td><td>-1.2</td><td>-25.2</td><td>n.d.</td><td>0.43</td><td>0.02</td><td>No</td></tr><tr><td>N37° 17.543<br>W32° 16.796</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Capelinhos</td><td>M22-CAP_C bulk</td><td>2022</td><td>2035-8</td><td>1669</td><td>-0.3</td><td>-24.6</td><td>n.d.</td><td>0.56</td><td>0.04</td><td>No</td></tr><tr><td>N37° 17.368<br>W32° 15.834</td><td>M23-CAP-sample 1 (upper<br>or bottom part)#</td><td>2023</td><td>844-6</td><td>1669</td><td>1.7</td><td>-26.7</td><td>n.d.</td><td>0.83</td><td>0.07</td><td>n.d.</td></tr><tr><td></td><td>M23-CAP-sample 1<br>(bottom or upper part)#</td><td>2023</td><td>844-6</td><td>1669</td><td>1.2</td><td>-22.0</td><td>n.d.</td><td>0.76</td><td>0.06</td><td>n.d.</td></tr><tr><td>Y3</td><td>M22-Y3_A bulk</td><td>2022</td><td>2034-7</td><td>1727</td><td>-6.0</td><td>-35.1</td><td>-1.4</td><td>0.22</td><td>0.03</td><td>Yes</td></tr><tr><td>N37° 17.518<br>W32° 16.663</td><td>M22-Y3_B (upper part)</td><td>2022</td><td>2034-7</td><td>1727</td><td>11.5</td><td>-18.9</td><td>2.1</td><td>0.38</td><td>0.03</td><td>Limited</td></tr><tr><td></td><td>M22-Y3_B (middle part)</td><td>2022</td><td>2034-7</td><td>1727</td><td>5.5</td><td>-20.9</td><td>n.d.</td><td>0.26</td><td>0.02</td><td>Yes</td></tr><tr><td></td><td>M22-Y3_B (bottom part)</td><td>2022</td><td>2034-7</td><td>1727</td><td>2.0</td><td>-21.4</td><td>0.9</td><td>0.29</td><td>0.02</td><td>Yes</td></tr><tr><td>N37° 17.509<br>W32° 16.670</td><td>M23-Y3 sample 1</td><td>2023</td><td>846-8</td><td>1727</td><td>2.0</td><td>-33.8</td><td>n.d.</td><td>0.18</td><td>0.02</td><td>n.d</td></tr><tr><td></td><td>M23-Y3 sample 2</td><td>2023</td><td>846-8</td><td>1727</td><td>10.2</td><td>-22.8</td><td>1.7</td><td>0.44</td><td>0.03</td><td>n.d</td></tr></tbody></table>

#, no orientation; §, mineralogy obtained by XRD and infrared-based and Mössbauer spectroscopies; n.d., not determined.

three main mechanisms from the perspective of biological Fe cycling (Fig. 8b):

(1) *Zetaproteobacteria* that dominate Fe(II) oxidation in benthic marine ecosystems consistently escape encrustation to the surface of the mats to continue the mat-building process, leaving residual mineralized structures to form the framework on which the mat layers are built. The paucity of C in stalks of mat-forming *Zetaproteobacteria* from the Jan Mayen Vent Fields corroborates our interpretation³⁵. Regardless of FeOM lifestyles, the residual organic C could originate from a smooth coating of polysaccharides bound to the stalks during cell growth⁵¹.

(2) The remaining residual organic matter is rapidly oxidized by a consortium of microorganisms, a mechanism supported by genome-resolved metagenomics and metatranscriptomics from the Lō'ihi mats⁴⁶. Besides, Fe(III)-reducing microorganisms usually rely on fermenting microbes to produce low molecular weight organic compounds that can be more readily assimilated⁸³ (Fig. 8a). Fermenters also produce H₂ to facilitate Fe(III) reduction and oxidation of organic molecules for Fe(III) reducers⁸⁴. These microbial communities tend to form a relatively little biomass compared to the large volume of Fe(III)-oxyhydroxides needed for growth because of the much lower amount of energy extractable from Fe(III)-oxyhydroxides compared to O₂ as the terminal electron acceptor. We note that such processes potentially operated in all BIF types if they were associated with enough organic matter to support the development of a similar microbial consortium.

(3) Finally, as evidenced by Fig. 6, the Fe biomineralized filaments provide a large adsorptive surface area for continuous scavenging of considerable amounts of hydrothermally-derived Fe(II) in the absence of direct microbial biomass, collectively forming organic C-poor Fh precursors to FeOx. The resulting substantial increase in Fe₂O₃/SiO₂ ratios from the filamentous microbial structures to the more crystalline FeOx, together with the unique absence of bulk Si-O-Fe elemental bonds, suggest that the gradual Fh transformation dislodges Si during the initial stage of diagenesis (Fig. 8b), offering new insights into formation mechanisms of Fe-rich/Si-poor microbands in BIFs.

**Role of environmental contribution to the formation of Si-rich microbands.** The exact timing and depositional mechanism of Si-rich layers in BIFs is uncertain. Si enrichment in Precambrian IFs may have originated through the precipitation of Fe-Si gels before reaching the

seafloor and their subsequent dewatering and Si release during diagenesis¹, while metamorphic reactions have been proposed for the BIF meso/macrobands⁹. It has also been suggested that the formation of Si-rich layers in Quaternary BIF analogue rocks⁴⁻⁶ was a syn-depositional feature and controlled mainly by abiotic processes involving for example, cycles of intense and paused hydrothermal/volcanic venting. It was concluded that their deposition with exquisite alternating Fe and Si bands was enabled by basin restriction from the open ocean, resulting in Si saturation and abiotic precipitation from the water column⁵,⁶. Other studies infer that biannual temperature cycles resulting in episodic abiotic silica saturation could have caused the prominent micro- to mesoband layering in BIFs⁸⁵,⁸⁶. Alternatively, Si precipitation could occur at, or immediately below, the sediment-water interface in mats along a saturation front with respect to amorphous silica due to the conductive cooling of the local hydrothermal environment³²,³⁸. Critical explanations about the timing of silica precipitation could come from Si isotope compositions of Si-rich minerals in mats¹,⁴,⁴,³².

As previously discussed, at LSHF, the Fe-rich mats host Si-containing Fh marked by the absence of amorphous Si mineral phases and Si-rich bands, contrasting with other Fe-Si mat-related deposits worldwide³⁹,⁵⁹. Si concentrations of the diffuse hydrothermal fluids are >2 mM at LSHF (Supplementary Table 6), similar to the concentrations in the global oceans that deposited BIFs but much higher than the 0.07 mM of modern seawater⁸⁷. However, diffuse fluid temperatures of ~50–90 °C likely inhibit massive abiotic Si precipitation in the mats, as suggested by PHREEQC thermodynamic calculations (Supplementary Table 6). It is also probable that rapid diffusive dispersal of dissolved Si to the low Si-containing modern ocean compared to Precambrian times, dilutes aqueous Si, thereby inhibiting seawater saturation with regard to silica minerals and precipitation of Si-rich bands. We, therefore, hypothesize that the subseafloor architecture beneath hydrothermal fields, which controls the fluid flow temperature toward the surface, is critical to facilitating abiotic silica precipitation and formation of Si-rich microbands in modern mats during the initial stage of diagenesis. Finally, we suggest that early diagenetic Si-rich microbands would have formed in parallel with the seawater Si precipitation responsible for BIF meso/macrobands.

**Contribution of benthic mats to Algoma-type BIFs.** We pose the question of whether Fe-metabolizing bacteria as part of microbial mats
<figure><img src="image_11.png" /><figcaption>**Fig. 8** | Proposed conceptual models of the microbial composition and evolution of the LSHF mats. a Mat trophic structure (Y3 site) and their effect on C isotope compositions. b Formation of the Fe-rich microbanding patterns and their related mineralogical composition and biotic/abiotic processes in the LSHF Fe-rich mats. (I-III) Successive steps of the generation of the microbands at Y3. Temperature of the diffuse fluids controls Fh particle size, which, in turn, affects surface area reactivity</figcaption></figure>

and adsorption capacity for chemical elements. Note that diffuse hydrothermal fluids were not observed during mat collection at WS, SI, and LL sites. For simplicity, stalk-formers are only represented in these mats. Vertical and horizontal scales are not represented. Not all microbial metabolisms described at LSHF<sup>29,41</sup> and seawater are shown. LMWOC, low molecular weight organic compounds.
were associated with the deposition of Algoma-type BIFs. Our observations point to a relatively rapid and progressive disappearance of Fh-based microbial filaments and possible generation of Fe-rich BIF microbands without phototrophic involvement, although carbon from the lysed planktonic phototroph biomass can fuel Fe(III) reducers and methanogens⁸⁴. Furthermore, our model highlights that magnetite crystallization predates haematite formation as the mats aged and that benthic microbial communities may be a key ingredient for (episodic) siderite-/magnetite-rich microbands in BIFs, which does not preclude existing models of magnetite precipitation in Archean BIFs (see Supplementary discussion).

Overall, we hypothesize that the diffusion efficiency of hydrothermal plumes and the relative location of hydrothermal/volcanic centres to sub-oxic surface waters would have strongly controlled the development of primitive benthic mats because of the FeOM need for microaerobic conditions. In the late Archean-Paleoproterozoic, these environmental conditions were to be expected where the Fe(II)/O₂ chemocline intersected with the shelf¹³,⁸⁸. Estimations for O₂ concentrations suggest that dissolved [O₂] of <1 µM in late Archean shallow seawater oases would have jumped to 5–30 µM in the surface ocean at the Archean-Proterozoic transition⁴⁵,⁸⁸. Considering that the [O₂] in the range of 5-20 µM corresponds to the optimum growth of microaerophilic FeOMs, their proliferation was likely plausible after the GOE. By contrast, the growth of modern FeOMs is limited at 1-5 µM [O₂], although cells are still active⁶¹. Evolving Archean-Proterozoic seawater [O₂] however likely helped shape the pace of FeOM evolution¹⁷, suggesting that the primitive FeOMs could have evolved and flourished in the presence of nanomolar [O₂] and, thus, contributed to the deposition of BIFs near shallower submarine volcanoes and hydrothermal vent fields to form Algoma-type BIFs. Finally, microbanding patterns formed in the LSHF mat material may be comparable to the Fe-rich microbands of the ~2.5 Ga-old Superior-type BIFs⁸⁹,⁹⁰. It is, thus, conceivable that similar microband-forming processes could have been operating on the continental shelf if microaerobic Fe(II)-oxidizing conditions were rife, coupled with pulsed delivery of Fe(II)-rich deep-sea hydrothermal vent fluids to shelf environments where Superior-type BIFs formed.

## Conclusion

We follow multiple lines of petrographic, mineralogical, and isotopic evidence to show that the development of C-poor Fe-rich BIF microbands could have been controlled by benthic mat-forming microbial processes. Our model shows that these processes begin with microbial oxidation of dissolved hydrothermal Fe(II) by *Zetaproteobacteria*, where microbial filaments provide a large surface of extensive nucleation and removal of dissolved Fe(II) from seawater in the presence or absence of microbial cells. This process is then followed by gradual biotic and abiotic transformation of poorly crystalline Fh into crystalline FeOx minerals as mats aged, during which the Fh-encrusted microbial filaments systematically disappear. Particularly, during early diagenesis, the coupling of sequential microbial reduction of Fh/FeOx to the oxidation of residual organic C by dissimilatory Fe(III)-reducing microorganisms leads to the transformation of Fh/FeOx to secondary crystalline siderite and magnetite. It is suggested that this process points to a key mechanism explaining the low organic C content of BIFs and helps to constrain similar isotopic signatures in BIFs as biomarkers. We suggest that these mechanisms may have been essential to the formation of BIF microbands, particularly in deposits formed in submarine hydrothermal vent fields and continental shelf environments where conducive Fe(II)/O₂ gradients developed. We further propose that the striking absence of Si-rich bands at LSHF can be related to the enduring dominance of high temperature ~50-90 °C diffuse fluids and dispersal of Si to the low Si-containing open ocean, which inhibits the supersaturation state required to drive Si band formation. Importantly, the present study sheds new light on the biotic and abiotic mechanisms by which Fe-rich marine rocks with organic C-poor and Fe-rich microbanding form in the modern ocean and helps explain deep-time observations in which biological processes are difficult to reconstruct.

## Methods

### Sampling and sample preparation

Mat morphology and their location are described in Table 1⁴⁸⁻⁵⁰. The Fe-bearing microbial mats were collected from Capelinhos (CAP), North Tour Eiffel (NTE), South Isabel (SI), West Sintra (WS), Y3, and Lava Lake (LL) sites during the annual MoMARSAT EMSO-Azores Observatory maintenance cruises. Two types of microbial mats were sampled, including flocculent mats at WS, SI, and LL and crumbly mats at CAP, NTE, and Y3 sites between 2021 and 2023⁴⁸⁻⁵⁰. The flocculent mats were easily disturbed and resuspended during sample collection. In situ sample collection and on-board sample processing were conducted, as previously described²⁹,⁴¹. A fraction of each Fe-rich mat was milled into a fine powder in agate mortar for whole-rock mineralogical and geochemical analyses. Thin sections of CAP, NTE, and Y3 mats were made using standard polishing procedures at GET laboratory, University of Toulouse (France). Thin sections were investigated under transmitted and reflected lights using an Olympus BX51 microscope combined with a Olympus SC30 microscope camera.

### Diffuse hydrothermal fluids sampling and analysis

Diffuse fluids are limited to the first few dozen metres away from the high temperature chimneys. The diffuse hydrothermal fluid samples of CAP, NTE, and Y3 sites were collected during the MoMARSAT cruise in June 2022⁴²,⁴⁹. Prior to fluid sampling using titanium gas-tight syringe with a fennel-like snorkel, the temperature of diffuse hydrothermal fluids was measured in situ using the temperature sensor of the HOV Nautilus submersible. The full procedure for diffuse fluid sampling is described in detail in Wheeler et al.⁴³. Aboard the research vessel, hydrothermal fluids were extracted from the fluid sampler, filtered with 0.22-µm Millipore filters and processed on board as described in Chavagnac et al.⁴³,⁶⁰. Aliquots were transferred in acid-cleaned HDPE bottles for a first set of analyses, i.e. density (refractometry), dissolved Fe concentrations (photometry), pH, Eh, salinity, total dissolved salts, dissolved H₂S concentrations, and conductivity (electrodes).

### Mineralogical analysis

X-ray diffraction analysis of Mn-poor bulk powder was undertaken with a Panalytical Empyrean diffractometer (CoKα radiation) operating at 35 kV and 45 mA. To determine the mineralogy of Fe- and Mn-rich samples, XRD patterns were acquired with a WAXS Panalytical Empyrean diffractometer (MoKα radiation) operating at 60 kV and 40 mA. In the latter case, a small fraction of bulk powders was mounted in a glass capillary tube of 0.5 mm diameter. Analysis of whole-rock powder samples was performed over angular ranges of 6–76° (CoKα) and 2–30° (MoKα) 2θ and a step size of 0.022° 2θ per 1 s at LPCNO, University of Toulouse III. Peak indexing was performed using Bruker Eva software and mineral identification was carried out by comparing XRD patterns with reference data⁵³,⁵⁴.

Short wavelength infrared (SWIR) and Fourier transformed infrared (FTIR) spectroscopies were carried out at IC2MP, University of Poitiers (France), to investigate further mineralogy, structure, and bonding of Fe oxides. SWIR reflectance spectra of bulk rock powders were obtained with an ASD TerraSpec 4 Standards-Resfield spectrometer within the 500–2500 nm range. The spectrometer is equipped with a contact probe including a white light source and three detectors (one Si detector for the 350–1000 nm range with 3 nm resolution and two InGaAs detector for the 1000-2500 nm range with 10 nm resolution). An average of 50 scans was measured on each sample. The Indico Pro software was used to collect the SWIR spectra. FTIR spectra were collected in transmission mode using a Nicolet iS50 FTIR spectrometer equipped with an Ever-Glo source, a potassium bromide (KBr) beamsplitter, and a DTGS KBr detector. Prior to FTIR analyses, KBr pellets, consisting of a mixture of 1 mg of sample and 149 mg of KBr, were prepared. The preparation was pressed for 5 min at 8 kbar to produce pellets and dried overnight in an oven at 120 °C to remove adsorbed water. The FTIR spectra were recorded in the 4000 to 400 cm⁻¹ region (middle infrared) and acquired from a mean of 100 scans at a resolution of 4 cm⁻¹. The
study of stretching and bending vibrations using FTIR spectroscopy is of
deep interest for characterizing adsorbed Si onto Fe oxide surface⁵⁵,⁵⁶.

Because it was difficult to detect poorly crystalline ferrihydrite in the
presence of crystalline mineral phases by XRD, the SI, NTE, and Y3
microbial mats were investigated by ⁵⁷Fe Mössbauer spectroscopy at the
Charles Gerhardt Institute, University of Montpellier (France) to char-
acterize Fe-bearing minerals, but also the distribution and redox state of Fe.
Analyses were carried out on ca. 300 mg of rock powder loaded in a 30-mm
diameter plastic holder irradiated by a ⁵⁷Co source in rhodium metal. The
Mössbauer spectra were acquired using a spectrometer in the constant
acceleration mode over the velocity range of -4 to 4 mm/s at room tem-
perature (~ 22 °C). The velocity scale was calibrated using an α-Fe foil. The
spectrometer was operated using a triangular velocity waveform for trans-
mission geometry. Gamma rays were detected with a NaI scintillation
detector (detection limit of 2%). The Mössbauer spectra were fitted by
Lorentzian lines to determine the spectral parameters (IS: isomer shift; QS:
electric quadrupole splitting; LW: full line width at half maximum; A:
relative resonance areas).

The thin sections were carbon-coated and imaged in BSE mode with a Tescan Vega 4 LMU scanning electron microscope (SEM) equipped with an energy dispersive spectroscopy (EDS) system at GET laboratory. Detailed observations of filamentous structures were conducted on platinum-coated thin sections using a JEOL JSM-6700F field emission SEM at Centre Castaing, University of Toulouse. The SEM-EDS analyses were run at a 10-30 kV accelerating voltage and working distances of 7.9 and 14.7 mm.

Bulk elemental analysis

Whole-rock geochemical analyses of major and trace elements were per-
formed at the clean lab of the GET laboratory. Approximatly 100 mg of bulk
rock powders were dissolved using double-distilled reagent acids in closed
screw-top Teflon vessels (Savillex). First, samples were digested with a con-
centrated HNO₃-HCl mixture using a MARS 6 microwave digestion system
at 180 °C for 15 min. After evaporation of the solution to dryness on hot plate
at 70 °C, samples were dissolved in suprapur HF at 70 °C for 2 h to break
down silicates. Major (except Si due to the evaporation of Si-F complexes) and
trace elements were measured by ICP-MS using a Thermo Scientific iCAP
triple quadrupole instrument, with a precision better than 2%. Calibrations
were made with multielemental solutions diluted to different concentrations.
The reference material CRPG Mica-Fe was used as a control.

Furthermore, the Si concentration of bulk rock samples were deter-
mined by X-Ray Fluorescence (XRF) on glass fusion discs using a PANa-
lytical AXIOS Max spectrometer equipped with an SST-max X-ray tube at
RRXG platform, University of Montpellier. The glass fusion discs were
prepared by mixing 0.6 g of rock powder with 6 g of lithium tetraborate. We
analysed 9 reference geological standards for internal calibration and one
external reference standard (IF-G). Loss of ignition (LOI) was determined
by heating the rock powders in a muffle furnace at 900 °C for 4 h.

Electron probe micro-analyser

Electron probe micro-analyses (EPMA) were carried out on carbon-coated thin sections using a CAMECA SXFive electron microprobe equipped with five wavelength dispersive spectrometers at Centre Castaing, University of Toulouse III. Analytical conditions of EPMA were performed using a ~1 µm focused beam at 15 kV and 20 nA. The counting time was set to 10 s for all elements. Si, P, and Fe were measured with a focused electron beam and step size of 0.5 µm during quantitative mapping. For silicate EPMA, the instrument was calibrated using reference material standards, including albite for Na, MgO for Mg, Al₂O₃ for Al, wollastonite for Si and Ca, grafi-onite for P, sanidine for K, Cr₂O₃ for Cr, MnTiO₃ for Mn and Ti, haematite for Fe, and SrSO₄ for S. For sulfur EPMA, we used the following standards: chalcopyrite for S, Fe, and Cu, mispickel for As, and native Mo for Mo.

Stable isotope analysis

Representative microbial mat samples were analysed for their C isotope
ratios of carbonates (δ¹³C<sub>carb</sub>) and organic matter (δ¹³C<sub>org</sub>) and for their N

isotope compositions (δ¹⁵N) at Cardiff University. The δ¹³C<sub>carb</sub> analysis was conducted on a Thermo Delta V Advantage mass spectrometer, coupled to a Thermo gasbench II for head-space sampling of carbonate samples. The powdered samples were first weighed into septum vials, which were flushed with helium and then acidified manually with 99% orthophosphoric acid. Samples were then left to react for 24 h at 60 °C to ensure complete reaction of dolomite and siderite. The precision of an in-house Carrara marble standard calibrated against NBS19, expressed as the standard deviation of repeated analyses, resulted in a value 0.07‰ for δ¹³C<sub>carb</sub>. For total organic content (TOC), total N (TN), associated δ¹³C<sub>org</sub> and δ¹⁵N compositions, powdered samples acidified in 25% HCl were left to react for two days to eliminate inorganic C. Residues were rinsed three times with ultrapure deionize water to remove remaining acid, dried and weighed into tin capsules. Samples were then measured on a Thermo Delta V Advantage mass spectrometer with Thermo Flash EA and Conflo III. Samples were analysed in combination with three in-house standards composed of a lab-grade caffeine with δ¹³C = -33.30‰ and two commercial collagen food supplements (MarCol) with a δ¹³C of -16.20‰ and an MCF with a δ¹³C value of -22.36‰. The in-house standards were calibrated against the international standards IAEA-CH6 and IAEA-600. Data precision, expressed as the standard deviation for repeated analyses of the independent standard MCF, was 0.2‰ (1 SD) for δ¹³C<sub>org</sub> and 0.1‰ (1 SD) for δ¹⁵N. The results are reported in the standard δ-notation in ‰ on the Vienna-Pee Dee Belemnite (VPDB) scale for C and air for N.

PHREEQC geochemical modelling

Thermodynamic calculation of the amorphous silica solubility as a function of temperature was performed with the batch reaction option of PHREEQC software package developed by USGS<sup>91</sup>. In addition, the effect of pH was considered and modelled. The calculation of the maximum concentration of dissolved Si required to reach thermodynamic equilibrium between amorphous silica and a standard seawater solution<sup>92</sup> was calculated using the llnl.dat data base<sup>93</sup>. Temperatures were set between 25 °C and 90 °C to represent environmental temperature bathing the microbial mats and the effect of salinity on activity coefficients was taken into account. According to the thermodynamic calculation, the main factor controlling the solubility of Si is temperature rather than pH.

Authors' contributions

J.A., E.C.F., and V.C. designed the study. V.C., C.R., C.D., and A.C. collected the samples and, analysed diffuse hydrothermal fluids. J.A., T.D., and F.B. conducted mineralogical analyses. F.P., M.H, A.N., conducted geochemical analyses. C.D. performed PHREEQC modelling. C.R. and V.C. carried out funding acquisition and supervised the project. J.A. drafted the original manuscript and made the figures. The draft was reviewed before submission by E.C.F., C.D., T.D., F.P, F.B, C.R., and V.C.

Data availability

All data are included in present paper and related supplementary
information.

Received: 20 September 2024; Accepted: 18 March 2025;
Published online: 02 May 2025

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## Acknowledgements

The MoMARSAT cruises were mainly funded by the French Oceanographic Fleet (FOF). We acknowledge the captains and crews onboard French research vessels *L'Atalante* and *Pourquoi Pas?* as well as the ROV and HOV team during MoMARSAT 2021–2023 campaigns. This study was financially supported by the National Research Agency (ANR) of France as part of the IRONWOMAN project (ANR-21-CE02-0012). Sampling permissions are given by the Portuguese authorities. They are provided in the technical reports of each MoMARSAT cruise available online on the FOF website. The authors are also grateful to the European Union (ERDF) and Région Nouvelle Aquitaine for providing financial support for the study of infrared-based spectroscopy. For scientific discussion, we thank Romain Guilbaud. We are grateful to Fabienne de Parseval (GET, University of Toulouse – UT) for the preparation of high-quality thin sections. We deeply thank Hadrien Henry (GET, UT) for the generation of EPMA element mapping. In addition, Françoise Maubé (GET, UT), Philippe de Parseval (GET, UT), Stéphane Le Blond

Du Plouy (Centre Castaing, UT), Nicolas Ratel-Ramond (LPCNO, UT), Simon Cayez (LPCNO, UT), and Moulay Brahim (RRXG platform, University of Montpellier) are thanked for their technical assistance during analytical works. We thank Bertus Smith and 3 anonymous reviewers for their constructive reviews and valuable feedbacks.

## Competing interests

The authors declare no competing interests.

## Additional information

**Supplementary information** The online version contains supplementary material available at https://doi.org/10.1038/s43247-025-02223-2.

**Correspondence** and requests for materials should be addressed to Jérémie Aubineau.

**Peer review information** *Communications Earth & Environment* thanks Albertus Smith and the other, anonymous, reviewer(s) for their contribution to the peer review of this work. Primary Handling Editor: Carolina Ortiz Guerrero. A peer review file is available.

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