Biogenic pyrite and metastable iron sulfides: Emerging formation pathways and geological and societal relevance

Abstract

Iron sulfide (Fe-S) minerals such as mackinawite (FeS), greigite (Fe₃S₄) and pyrite (FeS₂) are widespread on Earth, where their formation and dissolution are strongly linked to the biogeochemical cycles of iron, sulfur, carbon, oxygen, nutrients and trace metals. Recent studies have shed light on how microorganisms mediate their formation, with breakthroughs linked to biogenic pyrite. In this review, we highlight the formation pathways of Fe-S minerals, starting with the increasingly recognized roles of Fe(III) and intermediate sulfur species (e.g. S⁰ and polysulfides) during the initial steps. The mechanisms by which microorganisms affect Fe-S mineral formation are compiled and discussed for low (25–35°C) and high (≥ 80°C) temperatures, with specific examples from experimental studies. The morphology and precipitation rates obtained from experiments are compared to natural environments, and their similarities and differences are critically discussed. We then review the current state of the art for Fe-S minerals in the context of the origin of life and as environmental proxies and biosignatures in the geological record using their texture and chemical and isotopic compositions. We end by highlighting the importance of Fe-S minerals for current societal issues, such as the sequestration of organic carbon, the formation of acid drainages, metal recovery and nitrate removal, and their potential use as technological bio-materials in the future.

Introduction

Sedimentary iron sulfide (Fe-S) minerals constitute the most abundant type of sulfide minerals at the surface of the Earth (Fig. 1). Most sulfide in Fe-S minerals originates from a biological process: microbial sulfate reduction (MSR) (Rickard et al., 2017). The most stable of Fe-S minerals, pyrite (FeS₂), is the focus of this review. We also consider the metastable phases, such as mackinawite (FeS) and greigite (Fe₃S₄), as their importance is being increasingly revealed. Fe-S mineral formation was likely widespread on early Earth due to the early onset of MSR, the availability of Fe, mostly in its reduced form, and the absence of oxygen in the atmosphere (Rickard et al., 2017). Although anoxic environments have retreated to subsurface environments after the rise of oxygen in the atmosphere, the formation of Fe-S minerals in sediments is significant for the global geochemical cycles of sulfur, iron, carbon and other nutrients and trace metals. The burial of Fe-S minerals, together with organic carbon, controls the oxygenation of the atmosphere on geological timescales (Berner, 1989; Berner and Raiswell, 1983; Rickard, 2012a).

Figure 1. Images of iron sulfide minerals produced experimentally or naturally, representative of common sedimentary sulfides on Earth: (a) Characteristic colours of biogenic Fe sulfides formed in microbial cultures. Mackinawite (FeS) and greigite (Fe₃S₄) tend to form fine black nanoparticles. Initially-formed pyrite can also be black but transforms over time to dense shiny grey particles with increasing crystallinity and size; (b) sulfate-reducing bacteria encrusted in mackinawite, imaged using transmission electron microscopy; (c) false-colour image of pyrite spherules (blue) associated with cells of Desulfocapsa sulfexigens (yellow) and residual Fe(III) (oxyhydr)oxides or other Fe sulfides (orange); (d) pyrite spherules (blue arrow) formed together with euhedral vivianite (green arrow) in sulfur/sulfate-reducing enrichment cultures from Lake Pavin; (e) a cluster of pyrite spherules (blue arrow) together with greigite nanocrystals (red arrow) produced by the hyperthermophilic archaeon Thermococcus prieurii isolated from hydrothermal deep-sea vents; (f) diversity of the size and shape of pyrite framboids associated with smaller nanocrystals found in the modern Gulf of Lion (PRGL 1-4 borehole); (g) nanocrystals of pyrite in the process of recrystallizing to form a larger euhedral crystal in shelf sediments of the Gulf of Lion (PRGL 1-4 borehole); and (h) recrystallization with time and burial eventually leads to larger-sized euhedral pyrite commonly observed in the geological record, such as in the Mendon sedimentary Formation (3.2 Ga, South Africa). Pittings on the grain originate from in situ spot analysis such as secondary ion mass spectrometry. Images c-h were obtained using scanning electron microscopy.

In modern low-temperature, anoxic environments, interactions involving microorganisms and Fe-S minerals are common. While Fe-S mineral formation pathways have been extensively studied in abiotic conditions (Rickard, 2012b; Rickard 2012c), the impact of microorganisms on the properties of Fe-S minerals is still relatively poorly characterized, and the emphasis is often on controlled intracellular precipitation of Fe-sulfides in magnetotactic bacteria (MTB) (Pósfai and Dunin-Borkowski, 2006; Picard et al., 2016; Park and Faivre, 2022). However, advances in analytical methods have allowed a better understanding of the interplays between biogeochemical cycles; for example, with the description of cryptic cycles (e.g. Canfield et al., 2010; Holmkvist et al., 2011). Additionally, well-constrained experimental studies in recent years have allowed significant progress in deciphering biological from abiotic controls on the formation of Fe-S minerals, notably on mackinawite, greigite and pyrite formation.

This review first inventories the sources of Fe and S available for Fe-S mineral formation in natural environments, before delving into the most recent knowledge on how microorganisms affect Fe-S mineral formation, and how Fe-S minerals might have played a role in the origin of life. It is intended to update the recent review by Picard et al. (2016) and focuses on extracellular Fe-S mineral formation as a biologically induced mineralization process (Pósfai and Dunin-Borkowski, 2006). For a comprehensive overview of the chemistry of Fe-S minerals, the reader is referred to the book by Rickard (2012a,b,c) and the references contained therein. Furthermore, an excellent layman’s introduction to the history of pyrite is available in Rickard (2015). In this review, we focus on linking new knowledge from experimental studies to natural environments, with an emphasis on pyrite, which helps reconstruct ancient environments and could also record the involvement of microbial life in early biogeochemical cycles. Finally, we discuss the societal impacts of Fe-S minerals, such as the importance of Fe-S minerals for the sequestration of organic carbon in anoxic environments, the reactivity and stability of Fe-S minerals in (sub)oxic environments and the formation of acid drainages and the potential production of biogenic Fe-S minerals for future industrial applications.

Sources of Fe and S for the formation of iron sulfides

Sources of iron

Iron (Fe) and sulfur (S) are the 4^th and 16^th most abundant elements in the Earth’s crust, respectively. Iron is cycled between its 2+ and 3+ redox states by various abiotic and biotic processes (Kappler et al., 2021). Fe-S mineral formation requires Fe(II), which can come from multiple sources. Direct sources of Fe²⁺ originate from reduced water bodies such as groundwaters, porewaters and hydrothermal vents, as well as from the dissolution of Fe(II)-containing minerals such as siderite (FeCO₃), vivianite (Fe₃(PO₄)₂·8H₂O) and mackinawite (FeS). The reaction between dissolved Fe²⁺ and sulfide (S(-II); the sum of H₂S and HS^-) in microbial cultures typically results in the precipitation of mackinawite without further transformation to pyrite (Picard et al., 2016). Therefore, recent studies on biogenic pyrite formation have shifted focus onto indirect sources of Fe(II) coming from microbial or abiotic reduction of Fe(III)-bearing minerals such as Fe(III) (oxyhydr)oxides, silicates and phosphates, which also produce intermediate sulfur species in the system.

Iron(III) (oxyhydr)oxides such as ferrihydrite (Fe(OH)₃), lepidocrocite (γ-FeOOH), goethite (α-FeOOH) and hematite (Fe₂O₃) are common constituents of soils and sediments worldwide (Table 1). In marine sediments and soils that have 3–4 wt.% Fe on average, 20–50% of the Fe exists as Fe(III) (oxyhydr)oxides that are soluble in oxalate, dithionite and/or HCl (Canfield, 1997; Raiswell and Canfield, 1998; Johnson et al., 2020; Pasquier et al., 2022). These minerals exist in various sizes, shapes and crystallinities, which affect their surface areas and reactivity towards sulfide (Poulton et al., 2004b). This in turn affects the supply rate of Fe(II) and the amount of residual reactive surfaces that enable interfacial reactions for pyrite formation (Peiffer et al., 2015; Wan et al., 2017; Hockmann et al., 2020). For example, while ferrihydrite is highly reactive towards sulfide, its surfaces may be dissolved or coated too quickly by FeS or other minerals for pyritization to occur. By contrast, a less reactive mineral such as hematite may react too slowly with sulfide for pyritization to occur at an appreciable rate.

Table 1. Iron-bearing minerals and their reactivity towards sulfide.

Information compiled from:

Fe(III) (oxyhydr)oxides: Raiswell and Canfield (1998); Schwertmann and Cornell (2000); Poulton et al. (2004b); Sklute et al. (2018); Caraballo et al. (2022); Jiang et al. (2022).

Fe-clays: Langmuir (1997); Raiswell and Canfield (1998); Fan et al. (2023).

Magnetite: Amor et al. (2020); Kappler et al. (2023).

Vivianite & ferric phosphates: Huffman et al. (1960); Eynard et al. (1992); Hyacinthe and Van Cappellen (2004); Kandori et al. (2006); Cosmidis et al. (2014); Rothe et al. (2016); Schütze et al. (2020); Metz et al. (2023).

Clays are the most abundant mineral host of Fe in nature, making up around 40–50% of total Fe in marine sediments (Raiswell and Canfield, 1998). Iron-bearing clays are thought to be poorly reactive towards sulfide, with slow pyritization possible but occurring in the timescale of hundreds to thousands of years (Raiswell and Canfield, 1996). This view is, however, slowly changing. Layered clay minerals such as illite, smectite and kaolinite contain Fe(III) in either structural, basal or edge sites with different reactivities (Fan et al., 2023). Some of these Fe(III) are reducible by sulfide produced by sulfate-reducing microorganisms (SRM) within a timescale of days (Li et al., 2004; Liu et al., 2012), making them potentially relevant for rapid pyrite formation (Pasquier et al., 2022). The Fe-S minerals produced from these experiments are poorly characterized. Empirical observations have always noted close associations between diagenetic pyrite (including framboids) and clays (Canfield et al., 1992; Marin-Carbonne et al., 2022; Sanz-Montero et al., 2009; Wang et al., 2020). It is possible that pore spaces with low diffusivity within clays could lead to microenvironments with high supersaturation that encourage pyrite formation.

Besides (oxyhydr)oxides and clays, Fe phosphates are also potential Fe sources for Fe sulfide formation. Recent studies have shown that biogenic pyrite could precipitate via sulfidation of ferric phosphates (Berg et al., 2020; Duverger et al., 2020). Ferric phosphates are key minerals involved in phosphorus cycling, especially in ferruginous lakes and estuarine sediments (Hyacinthe and Van Cappellen, 2004; Cosmidis et al., 2014). Another phosphate mineral, the Fe(II)-containing vivianite, is also commonly formed when dissolved Fe²⁺ and phosphate are released into the solution. Such conditions are found in Fe(III) reducing zones of water columns and sediments or in microbial cultures (Rothe et al., 2016; Bronner et al., 2023). Vivianite could be sulfidized to pyrite as the mineral is buried in deeper sulfate reduction zones. Interestingly, vivianite can persist under sulfidic conditions for months in microbial cultures as observed empirically (Picard et al., 2018; Nabeh et al., 2022; Bronner et al., 2023). Hence, it is unclear if vivianite is an important Fe source for pyrite formation. Nonetheless, the trends in the last few years show that the research field is shifting towards investigating alternative sources of Fe, with different interfacial chemistry, templating effects and microenvironments that promote a multitude of pathways for pyrite formation.

It is important to keep in mind that minerals with the same chemical formula and crystal structure may also exhibit different reactivities depending on their formation pathways, association with organics, trace metal contents and sizes. Natural minerals will probably have associated organics, especially when they are formed by widespread Fe-metabolizing microorganisms (Kappler et al., 2021). Organics and trace metals (e.g. Ni) have been shown to retard the extent of sulfidation or pyritization (ThomasArrigo et al., 2020; Duverger et al., 2021; Wang et al., 2023; Wu et al., 2023). Particle sizes also greatly affect pyritization rates. For example, micron-sized magnetite (mixed-valent Fe₃O₄) grains exhibit low reactivity towards sulfide with estimated half-lives in the order of 100 years (Canfield and Berner, 1987). By contrast, freshly precipitated nano-magnetite that are more similar to biogenic magnetite are rapidly sulfidized within days, with pyrite formation accelerated in the presence of elemental sulfur (S⁰) and polysulfides (Poulton et al., 2004b; Runge et al., 2023), while being retarded in the presence of organics (Runge et al., 2024).

Sources of sulfur

Mackinawite and greigite are Fe-S minerals in which both sulfur atoms have -2 redox states, while pyrite is an Fe polysulfide in which the sulfur atom has a -1 redox state. The availability of reduced sulfur species is tied to complex sulfur cycling driven by microbial and abiotic processes. Below, we discuss the sources of sulfide, polysulfides and S⁰ as the main sulfur-bearing species involved in sedimentary Fe-S mineral formation.

Sulfide (S(-II)) is produced in porewater as the main product of microbial sulfate reduction (MSR), the dominant mode of organic matter oxidation in marine sediments (Jørgensen et al., 2019), where it is fuelled by sulfate ions present at relatively high concentrations in seawater (~28 mM) and diffusing downwards into the sediment. Sulfide speciates into two major species: H₂S and bisulfide (HS^-). The pKa of H₂S is about 7, meaning that at normal seawater pH (around 8), it represents less than 10% of the total dissolved S(-II) pool. The majority of the pool will then be in the form of HS^-. The link between MSR and pyrite formation is potentially weaker in low-sulfate lake sediments, where S(-II) may be derived from the breakdown of biomass containing reduced organic sulfur (Wei et al., 2023). The breakdown of reduced organic sulfur, as well as the reduction of organic sulfite, is also thought to have been a prevalent source of reduced sulfur for pyrite formation in low-sulfate Archean oceans (Fakhraee and Katsev, 2019). Microbial disproportionation or reduction of elemental sulfur (S⁰) and thiosulfate may also constitute a source of S(-II) for pyrite formation, as suggested by both microbial culture experiments (Canfield et al., 1998) and observations from natural marine sediments (Zopfi et al., 2008).

Polysulfides are soluble ionic species consisting of relatively short chains of sulfur atoms terminated by negative charges (n in S_n^2- is typically smaller than 9, with n = 5, 6, and 4 being the more abundant forms under certain experimental conditions; Kamyshny et al., 2004). At high concentrations, they exhibit a distinct yellow-green colouration with absorbances in the 250–350 nm region (e.g. Domingos et al., 2023). At pH values representative of seawater and marine pore waters, polysulfides form spontaneously through the chemical reaction of S⁰ with S(-II), as depicted in the equation below:

(1) $$ {\mathrm{H}\mathrm{S}}^{{\textstyle \hbox{-} \hbox{-} }}+( n{\textstyle \hbox{--}}1)\hskip2pt {\mathrm{S}}^0={\mathrm{H}}^{+}+{{\mathrm{S}}_{\mathrm{n}}}^{2{\textstyle \hbox{--} }} $$

Polysulfides can also form through the oxidation of hydrogen sulfide by O₂, Fe(III) and manganese (oxyhydr)oxides, with kinetic reaction rates in the following order: MnO₂ > O₂ > Fe(OH)₃ > S⁰, corresponding to characteristic reaction times ranging from a few minutes to about a year at conditions relevant for sulfidic marine sediments (Chen and Morris, 1972; Poulton et al., 2004a; Avetisyan et al., 2019, 2021). They can furthermore be formed biologically, as a by-product of both phototrophic and chemotrophic S-oxidative microbial metabolisms (Findlay, 2016). The environmental prevalence and relative abundance of polysulfide species in the environment have remained relatively elusive until the development of robust analytical methods based on derivatization almost twenty years ago (Kamyshny et al., 2006, 2009), and even today reliably measuring polysulfide concentrations in sediments is a difficult endeavour. However, it has been established that under many conditions, polysulfides can be expected to occur in concentrations approaching calculated equilibrium with elemental sulfur based on reaction (1) (Kamyshny and Ferdelman, 2010). For this reason, we focus our discussion on factors controlling the abundance of sedimentary S⁰ in the paragraph below.

In low-temperature environments, S⁰ is formed by oxidation of more reduced sulfur species. This oxidation process can occur abiotically in the presence of oxygen or oxidized Fe or Mn phases (e.g. Rickard and Luther, 2007) or it can be mediated by S-oxidizing bacteria and archaea (Dahl et al., 2008). Since measured rates of prokaryotic S-oxidation are several orders of magnitudes faster than rates of chemical sulfide oxidation to S⁰ by molecular oxygen (Luther et al., 2011), it is often assumed that S⁰ formed in low-temperature environments is mostly formed as a result of microbial activity. A diversity of phototrophs and chemotrophs are able to biomineralize S⁰ in the form of intra- or extra-cellular S⁰ globules (Dahl and Prange, 2006; Cron et al., 2019; Marnocha et al., 2019) or extracellular S⁰ filaments (Wirsen et al., 2002; Sievert et al., 2007). Sulfur rarely accumulates in sedimentary environments, due to its chemical and biological instability. Chemically, S⁰ is only stable in a very narrow range of Eh-pH conditions, and not at all above neutral pH values (Rickard and Luther, 2007). It has actually been found to be thermodynamically unstable in a range of natural sulfidic waters (Helz, 2014). Biologically, S⁰ is used as a source of energy for a diverse range of S-oxidizers, S-reducers and microorganisms that perform S⁰ disproportionation (Dahl, 2020a). Some microbes, such as the thermoacidophile Acidianus, are particularly efficient at recycling S⁰ as they can grow from all three reactions (Amenabar and Boyd, 2018). Abundances of S⁰ in sediments typically range from 11 μmol/kg to 60 mmol/kg (see compilations in Ye and Jing (2022), their Table S-1 and Zopfi et al. (2004).

Not all forms of S⁰ are created equal, as a number of factors can influence the chemical and biological reactivity of S⁰ in the environment. Obviously, S⁰ biominerals stored intracellularly (Dahl, 2020b) are unavailable for extracellular chemical reactions and consumption by other microorganisms. Some microbes that produce S⁰ biominerals as an extracellular energy storage resource have also evolved strategies to increase their environmental stability and/or avoid piracy by other S⁰-consuming cells (Cosmidis and Benzerara, 2022). As an example, some S-oxidizers encapsulate their extracellular S⁰ globules in an organic membrane, allowing S⁰ to be formed and persist in the extracellular medium in a thermodynamically unstable state (Cron et al., 2019, 2021; Marnocha et al., 2019). Sulfur minerals produced chemically by oxidation of sulfide in the presence of organics (a process called S⁰ organomineralization) can also exist outside of their thermodynamic stability domain (Lau et al., 2017; Cosmidis et al., 2019). The size of S⁰ particles affects their chemical reactivity (Steudel, 2003) and bioavailability (Franz et al., 2007), with smaller particles being overall more unstable. In sediments, S⁰ is thought to exist mainly as colloidal sols, which are more reactive than crystalline S⁰ (Zopfi et al., 2004). Due to its reactivity, colloidal or nanoparticulate S⁰ is likely to be the main source of polysulfides in sulfidic sediments (Kleinjan et al., 2005; Mol et al., 2022), but abundances of S⁰ particles in the micron or sub-micron size range are rarely reported (Findlay et al., 2014). Such considerations on S⁰ reactivity should be taken into account in experimental studies investigating the role of this mineral in Fe-S mineral formation.

Biogenic iron sulfide mineral formation at low temperature

Mackinawite and greigite formation at low temperature

Mackinawite (FeS) and greigite (Fe₃S₄) are described as metastable Fe-S minerals with respect to pyrite (cubic FeS₂). For that reason, they are generally assumed to be present in modern sedimentary environments but absent from ancient rocks and sediments (Rickard, 2012b). Recent modelling studies have nonetheless found that greigite could be much more stable than originally expected (Subramani et al., 2020; Shumway et al., 2022; Son et al., 2022). In natural environments, metastable Fe-S minerals are assumed to be the main constituent of acid-volatile sulfides (AVS), which are the solid and aqueous phases that produce sulfide after treatment of samples with HCl. The contribution of mackinawite and greigite to the AVS fraction is likely to vary depending on the environment considered and it is possible that the AVS fraction does not capture all of the mackinawite and greigite content of an environment, as these two minerals might not completely dissolve in HCl (Rickard and Morse, 2005). While direct identification of mackinawite and greigite is most easily done using X-ray diffraction (XRD) (Berner, 1962; Evans Jr et al., 1964; Skinner et al., 1964; Lennie, 1995), combined studies including AVS analysis and XRD in sedimentary environments are not common. Biogenic precipitates of mackinawite and greigite are certainly present in anoxic sediments, and understanding their physical properties is of crucial importance because they could react differently from abiotic precipitates to analytical procedures, such as leaching procedures for AVS quantification.

Abiotic mackinawite has a tetragonal layer structure with cell parameters a=b=3.6735 Å and c=5.0329 Å (Lennie, 1995). Greigite has a cubic unit cell, with a=9.876 Å (Skinner et al., 1964). Mackinawite is stoichiometric FeS and precipitates rapidly from the reaction between aqueous Fe²⁺ and dissolved sulfide (H₂S/HS^-) (Rickard, 1995, 2024; Rickard et al., 2006). The chemical formula for the greigite formula is averaged to Fe₃S₄, although its exact composition has not been determined (Rickard, 2012b). Greigite forms through the solid-state transformation of mackinawite, which is driven by the oxidation of Fe atoms and their rearrangement (Rickard, 2012b, Lennie et al., 1997). As greigite does not precipitate directly from the solution, its characterization is difficult because residual mackinawite is always present (Rickard, 2012b). Owing to its magnetic properties, greigite can be detected in the sedimentary record but could be difficult to differentiate from magnetite (Fe₃O₄) (Roberts et al., 2011). In experimental studies and microbial cultures, a strong neodymium magnet can be used to check for its presence.

In low-temperature environments, the main source of sulfide for mackinawite and greigite formation is microbial sulfate reduction (MSR) (Rickard et al., 2017). As discussed in the ‘Sources of sulfur’ section above, other microbial processes can provide sulfide for metastable Fe-S mineral formation, but their significance varies depending on the environment (e.g. Jørgensen et al., 2019; Wu et al., 2021). Mesophilic sulfate-reducing bacteria (SRB) have been used in most experimental studies to decipher the role of microorganisms in the formation of extracellular Fe-S minerals at low temperatures, e.g. through biologically induced mineralization (Pósfai and Dunin-Borkowski, 2006; Picard et al., 2016; Park and Faivre, 2022). The few sulfate-reducing archaea (SRA) available in cultures are (hyper)thermophilic and have not been used for Fe-S mineral formation investigations. Experimental studies consist of precipitating biogenic Fe-S minerals by adding a source of Fe to cultures of SRB and (ideally) comparing them to abiotic Fe-S minerals precipitated by adding sulfide to the culture medium containing the same source of Fe. The following strains have been used in a range of temperatures between room temperature and 35°C: Desulfovibrio capillatus (Ikogou et al., 2017), Desulfovibrio desulfuricans (Rickard, 1969b; Neal et al., 2001; Li et al., 2006; Stanley and Southam, 2018; Duverger et al., 2020), Desulfovibrio hydrothermalis (Picard et al., 2018; Nabeh et al., 2022), Desulfovibrio magneticus (Nabeh et al., 2022), Desulfovibrio vulgaris (Zhou et al., 2014; Mansor et al., 2019), Desulfovibrio spp. (Li et al., 2004), Desulfosporosinus orientis (Stanley and Southam, 2018), Desulfotomaculum sp. (Fortin et al., 1994). Additionally, uncharacterized enrichments of SRB from various environments have also been used in experimental studies (Herbert et al., 1998; Donald and Southam, 1999; Gramp et al., 2010). The redox state of the initial source of Fe in cultures of SRB appears to determine the mineralogy of the final Fe-S mineral products (Rickard, 1969b; Duverger et al., 2020). In experimental studies using Fe²⁺ as a unique source of Fe, mackinawite precipitates first and transforms into greigite over time (Picard et al., 2018). In experimental studies using Fe(III), the reduction of the latter by sulfide produces Fe(II) and intermediate sulfur species, which appear necessary for the formation of pyrite (see the section ‘Biogenic pyrite formation at low temperatures’ below).

Owing to the semi-conducting properties of Fe-S minerals, the role of mackinawite in improving the efficiency of extracellular electron transfer in microbial cultures has been investigated (Nakamura et al., 2010; Jiang et al., 2014; Kondo et al., 2015; Zhu et al., 2022). These studies have explored the production of biogenic Fe-S minerals by Fe(III)-reducing bacteria, e.g. Shewanella or Geobacter, which also have the ability to reduce intermediate sulfur species. In those experimental systems, microbial reduction of thiosulfate or elemental sulfur produces sulfide, while microbial and/or chemical reduction of Fe(III) by sulfide is the source of Fe(II) that promotes the precipitation of Fe-S minerals. Although many studies of interest also considered biogenic mackinawite in the context of bioremediation studies (e.g. Sharma et al., 2024; Yang et al., 2017), we restrict this section to studies that have attempted to characterize and quantify the differences between abiotic and biogenic mackinawite and greigite experimentally.

Microbial influence on the physical characteristics of mackinawite

The presence of microorganisms in experimental systems does not prevent the formation of mackinawite. However, the availability of cell surfaces as templates for mineral nucleation and growth can impact its physical properties, such as the size of crystallite domains, crystallinity and propensity to aggregate, and can lead to cell encrustation (Picard et al., 2016, 2018, 2021; Mansor et al., 2019; Nabeh et al., 2022). Templating occurs when Fe²⁺ first interacts with negatively charged bacterial cell surfaces before precipitating with sulfide (Beveridge, 1989); i.e. when SRBs are grown in a culture medium that contains millimolar concentrations of Fe²⁺ (Picard et al., 2018). Physical and chemical characteristics are important when considering reactivity and transformation of solid phases in natural environments.

X-ray diffractograms of ‘templated’ minerals display fine peaks with high intensity, while those of ‘non-templated’ minerals display broad peaks with low intensity (Picard et al., 2018; Mansor et al., 2019; Duverger et al., 2020, 2021, Nabeh et al., 2022). For a quantitative approach, XRD data can be fitted to estimate the average size of the crystalline domains using the Scherrer equation (Wolthers et al., 2003; Unruh and Forbes, 2019). Abiotic mackinawite precipitated in water is nanocrystalline and has an average particle size of 7.4 nm (a/b axis) x 2.9 nm (c axis) (Wolthers et al., 2003; Ohfuji and Rickard, 2006). While there is variability among the few studies that reported crystallite size data, biogenic mackinawite grows along both the a/b axis and the c axis more than the abiotic controls (Zhou et al., 2014; Picard et al., 2018; Mansor et al., 2019). Specifically, crystallite domains of biogenic mackinawite precipitated with SRB grown with Fe²⁺, and of mackinawite precipitated with dead SRB incubated with Fe²⁺ then with sulfide, are on average significantly larger than those of abiotic mackinawite precipitated in water and the SRB medium (with or without simple or complex organic molecules), and those of biogenic mackinawite precipitated in SRB cultures to which Fe²⁺ has been added after growth and sulfide production (Picard et al., 2018, 2021). When SRBs are grown with Fe(III)-citrate, biogenic mackinawite displays smaller crystallite domains than when Fe²⁺ is directly available (Ikogou et al., 2017). In all conditions in which crystallite domains are small, templating cannot occur for the two following reasons: (1) salts and organic molecules in the medium do not provide scaffolds for mineral nucleation and growth and (2) Fe²⁺ does not have the opportunity to bind to cell surfaces if sulfide is already present in the environment. A number of other studies have investigated the formation of biogenic Fe-S sulfide minerals using Fe(III) minerals as starting Fe source (e.g. Fe(III) (oxyhydr)oxides, Fe(III)-containing clays, Fe(III) phosphate minerals); however, they did not provide information about the size of crystallite domains (Duverger et al., 2020).

High-resolution studies of biogenic mackinawite indicated that it becomes crystalline very rapidly. Selected-area electron diffraction (SAED) patterns showed a shift from a polycrystalline material after one week to a single-crystal pattern after one month in mackinawite produced in cultures of Desulfovibrio desulfuricans grown with Fe²⁺. As early as one week of incubation, and in longer experiments, the 5 Å d-spacing of mackinawite can be clearly observed using high-resolution transmission electron microscopy (HR-TEM) (Duverger et al., 2020). It is usually assumed in abiotic studies that mackinawite is transient and short-lived in anoxic sedimentary environments. The crystallinity of biogenic mackinawite will probably play an important role in its stability and further transformations and should be considered.

Aggregation of biogenic mackinawite particles in cultures of SRB grown with Fe²⁺ is visible when observing cultures by eye. Minerals observed in unshaken cultures of SRB grown with Fe(II) form sticky clumps, while abiotic mackinawite sediments homogeneously at the bottom of serum vials. When biogenic minerals are resuspended in solutions, they also appear less opaque than abiotic minerals (Picard et al., 2018). Observations using scanning and/or transmission electron microscopy (SEM and/or TEM) reveal that biogenic mackinawite has a flaky texture and can reach mm-range sizes, much larger than aggregates of abiotic mackinawite (Herbert et al., 1998; Gramp et al., 2010; Picard et al., 2018; Duverger et al., 2020). Biogenic particles precipitated in cultures of SRB grown with Fe²⁺ form larger aggregates than abiotic particles precipitated in the culture medium, as determined using dynamic light scattering (Picard et al., 2018).

Cell encrustation by Fe-S minerals has been reported in cultures of SRB grown with Fe²⁺. This observation is consistent with the ability of cells to serve as templates, e.g. when Fe²⁺ is available to bind to the cell surface before interacting with sulfide. Cells from cultures of SRB grown with and without Fe(II) have a similar smooth aspect when imaged with SEM (Picard et al., 2018). To detect the presence of Fe-S crusts at the surface of cells, it is best to prepare thin sections of resin-embedded pellets that contain cells and minerals and to image them using TEM (Shuster et al., 2019; Picard et al., 2018). Fe-S minerals have been directly observed at the surface of both gram-negative and gram-positive SRB using TEM imaging of thin sections prepared after short incubations (Fortin et al., 1994; Donald and Southam, 1999; Picard et al., 2018; Stanley and Southam, 2018). Although no direct mineral characterization of the crusts has been performed in these studies using SAED, encrusted cells were imaged after short periods of time (one week or less), when mackinawite is the only mineral phase detected by powder X-ray diffraction (XRD) (Picard et al., 2018). It is unknown if cell encrustation in Fe-S minerals is a common situation for microbial cells in natural environments. There have been only two reports of microorganisms encrusted with Fe-S minerals in mine tailing sediments, where SRB are ubiquitous (Ferris et al., 1987; Fortin and Beveridge, 1997). The impact of Fe-S mineral encrustation on microbial metabolic activity is also unknown. The assimilation of carbon substrates by bacteria encrusted in Fe(III) oxyhydroxides appears to be inhibited (Miot et al., 2015). Duverger et al. (2020) suggested that sulfate reduction in cultures of Desulfovibrio desulfuricans grown with Fe²⁺ is hindered by cell encrustation in Fe-S minerals. It is possible that the high concentration of soluble Fe²⁺ (20 mM) used in their study, rather than encrustation, inhibited metabolic activity. Indeed, other studies that described cell crusts around cells used lower concentrations of Fe²⁺ (e.g. 3.5–4.0 mM) and did not report inhibition of sulfate reduction (Fortin et al., 1994; Picard et al., 2018). Although Fe(II) is required by SRB for growth, high concentrations of soluble Fe²⁺ might trigger stress or other metabolic responses. SRB released seven unidentified organic molecules in response to 4 mM Fe²⁺ in the growth medium (Picard et al., 2019). It is possible that (1) these extracellular compounds are inhibitory at higher concentrations, and/or that (2) high concentrations of Fe²⁺ are directly inhibitory for SRB. As Fe-S minerals are semi-conducting materials, some studies have evaluated the potential role that biogenic Fe-S minerals could play in enhancing extracellular electron transfer (EET) between microorganisms and solid phases. Direct contact between SRB and Fe-S minerals might be beneficial, and electrochemical studies indicated an increase in electron transfer from cells to electrodes in the presence of biogenic Fe-S minerals (Deng et al., 2020). This observation has also been reported in studies in which Shewanella strains produced Fe-S minerals when grown with thiosulfate and a source of Fe (Nakamura et al., 2010; Jiang et al., 2014; Kondo et al., 2015). The facilitated electron transfer through biogenic Fe-S minerals could be of use in energy-depleted environments. As discussed later in this review, Fe-S minerals can store significant amounts of organic carbon that could be potentially accessed as energy sources by SRB and other microorganisms (Picard et al., 2019; Nabeh et al., 2022).

Microbial influence on the transformation of mackinawite to greigite under anoxic conditions

In abiotic experiments, greigite formation from mackinawite requires an oxidant, which can be traces of oxygen at the surface of mackinawite (Benning et al., 2000), aldehydes (with the intention of sterilizing the experimental system) (Rickard et al., 2001), or polysulfides (Benning et al., 2000). High temperature accelerates the abiotic transformation of mackinawite to greigite (Lennie, 1995). A recent modelling study indicated that greigite formation is favourable in anoxic, alkaline and low-temperature environments where Fe is enriched and sulfide limited (Turney et al., 2023). There is experimental evidence that SRM could be driving and/or accelerating the transformation of mackinawite to greigite under strictly anoxic conditions. As noted above, greigite can form in the absence of microorganisms. However, in experimental studies under strict anoxic conditions, greigite forms in cultures of SRM grown with Fe²⁺ at their optimal pH and temperature after several months of incubation, while it does not form in abiotic experiments (Rickard, 1969b; Picard et al., 2018, 2021; Mansor et al., 2019; Nabeh et al., 2022). In old cultures, greigite is stable and does not transform further, and mackinawite is still detectable after several years of incubation, suggesting that the full transformation of mackinawite to greigite in these experimental conditions is slow (Picard et al., 2018; Nabeh et al., 2022; Picard, unpublished). In the study by Mansor et al. (2019), greigite could be detected by SAED in the TEM after six months of incubation but not by XRD, suggesting that transformation yield can vary with experimental conditions and microbial strains. In abiotic studies that maintain strict anoxic conditions, mackinawite remains stable and does not transform into greigite (Benning et al., 2000; Picard et al., 2018, 2021).

Biogenic greigite produced in cultures of SRB displays an average crystalline domain size of 19.2 nm, which is in the size range of biogenic mackinawite precipitated in cultures of SRB grown with Fe²⁺, supporting the hypothesis of solid-state transformation from biogenic mackinawite (Picard et al., 2018). The ‘small’ mackinawite produced in cultures to which Fe²⁺ has been added after sulfide production does not transform to greigite, nor does abiotic mackinawite precipitated with organic molecules or mixtures of complex organics (Picard et al., 2018, 2021). Interestingly, the ‘large’ mackinawite that is produced at the surface of dead cells also does not transform into greigite (Picard et al., 2018). This indicates that under strict anoxic conditions, crystalline mackinawite and metabolically active cells are required to form greigite. The oxidant necessary to oxidize Fe(II) in mackinawite could originate from H⁺ produced by residual activity of SRB in old cultures (Mansor et al., 2019). However, it is unknown why the transformation only takes place after several months, when greigite precipitation is actually favourable in anoxic and alkaline conditions at low temperatures (Turney et al., 2023). The experiments cited above have been performed in anoxic conditions, around neutral pH and at temperatures between RT and 35°C (Rickard, 1969b; Zhou et al., 2014; Picard et al., 2018, 2021; Mansor et al., 2019).

There are still many unknowns into what controls greigite formation in low-temperature environments. Changes in one or several parameters in culture media can promote or suppress the formation of greigite. For example, varying the amount of sulfide produced in cultures of Desulfovibrio hydrothermalis, and of Desulfovibrio magneticus RS-1 grown with Fe(II) resulted in different patterns of greigite production (Nabeh et al., 2022). A recent report of the formation of greigite in cultures of Geobacter sulfurreducens grown with Fe(III) and S(0) suggests that other conditions than the ones described above might exist to promote greigite formation in the presence of microorganisms (Bronner et al., 2023). The encrustation of cells by greigite has not been reported. As noted in the previous section, encrustation has been reported in relatively short-term cultures when mackinawite is the only phase detected by XRD. It is thus unknown how the interactions between SRM and Fe-S minerals evolve and if the transformation of mackinawite to greigite affects the localization of minerals on microbial cells.

While its significance in terms of amounts precipitated is unknown, biogenic greigite produced intracellularly by magnetotactic bacteria (MTB) through biologically controlled mineralization (as opposed to biologically induced mineralization described above) has attracted attention for their potential to participate in the magnetic properties of sediments and sedimentary rocks and to produce biosignatures (Amor et al., 2020). Two morphological types of intracellular greigite producers have been described: magnetotactic multicellular prokaryotes (MMPs) and large rod-shaped bacteria. Intracellular greigite producers are found in reducing and sulfidic environments and use sulfate as an electron acceptor (Farina et al., 1990; Mann et al., 1990; Lefèvre et al., 2011; Descamps et al., 2017). Intracellular greigite formation is favoured at low Fe and high sulfide concentrations (Descamps et al., 2017). Similar to extracellular greigite, intracellular crystals in MTB form from the solid-state transformation of mackinawite (Pósfai et al., 1998). Mackinawite can be detected at the end of magnetosome chains during short periods of time before transforming to greigite in a matter of days to weeks, suggesting that the intracellular formation of greigite is faster than that of extracellular greigite (Pósfai et al., 1998).

Biogenic pyrite formation at low temperatures

Pyrite is the most widespread iron sulfide on the Earth’s surface and is often ascribed a biological origin, as microbial sulfate reduction (MSR) is the main pathway to produce sulfide precursors from the particularly stable molecule sulfate. Chemically speaking, pyrite is one of two iron disulfide (FeS₂) polymorphs, along with marcasite. Pyrite’s ubiquitous presence in the geological record is due to its incredible thermodynamic stability in comparison to other reduced sulfur minerals (Schoonen, 2004). It is thus tempting to assume that its formation results from the simple precipitation of its constituent cation (Fe²⁺) and anion (S₂^2-), as is the case for most non-silicate authigenic sedimentary minerals (e.g. halides, sulfates, carbonates, phosphates). It was therefore assumed that pyrite is relatively easy to obtain in laboratory experiments. Indeed, a long list of recipes for pyrite formation can be found in the literature, as compiled in Rickard (2012).

However, the conditions used to precipitate pyrite in the laboratory are often far removed from those found in natural sediments, particularly those supporting microbial life. Indeed, many experiments have been carried out at so-called ‘low temperatures’ of <100°C, which are lower than hydrothermal temperatures but not so low in a biological sense. In fact, temperatures above 45°C and 80°C are in the ecological niche ranges of thermophiles and hyperthermophiles, respectively, which are not responsible for most MSR globally. Similarly, experiments carried out at pH levels far from neutral raise questions about the parallels that can be drawn with the natural environment. All these abiotic syntheses have been crucial to understanding the mechanisms of pyrite formation and the different environmental factors at play. However, they do not allow us to correctly decipher the role played by microorganisms in pyrite formation.

Pyrite formation in the presence of microbes has more recently been reported in environmental studies of sediments (Quevedo et al., 2021; Tribovillard et al., 2022), thermal springs (Tsyrenova et al., 2018) and deep geological repositories (Boylan et al., 2019), as well as in numerous recent experimental studies such as in metal corrosion (Etim et al., 2018; Jia et al., 2018, 2019; Rasheed et al., 2019), arsenic bioremediation in contaminated soils (Pi et al., 2017; Saunders et al., 2018; Lee et al., 2019), various microbial enrichment cultures (Gao et al., 2019, 2021; Berg et al., 2020; Allen et al., 2021; Ikkert et al., 2021; Karnachuk et al., 2021; Wei et al., 2023; Yang et al., 2023) and (co-)cultures (Thiel et al., 2019; Zhou et al., 2019; Duverger et al., 2020; Sun et al., 2023; Ke et al., 2024). It should be pointed out that the role of microorganisms in the formation of biogenic pyrite was not the main objective of some of these studies and, therefore, although anoxic conditions were applied during the microbial culture, they were not maintained throughout the experiments, particularly during the analysis, which may result in the transformation of metastable iron sulfide phases into more stable pyrite that may not have been present initially. Moreover, the identification of pyrite in some of these studies is not infallible: SEM images rely solely on morphology, semi-quantitative SEM-EDS analyses may be subject to large errors due to surface effects, chemical extractions are not very selective to pyrite (e.g. chromium-reducible sulfur can also contain zero-valent sulfur species if not initially separated with methanol and HNO₃ can also target iron silicates) and there are still poorly indexed X-ray diffractograms in the literature that over-interpret the presence of pyrite, as they are made on multi-phase systems (containing elemental sulfur or vivianite, for example) that generate numerous peaks, many of which coincide with those of pyrite. Despite the fact that it is widely accepted that microorganisms play a key role in the production of pyrite, it is interesting to note that most attempts to produce pyrite in the presence of microorganisms have failed (Picard et al., 2016, 2018; Ikogou et al., 2017; Zhou et al., 2017; Stanley and Southam, 2018; Mansor et al., 2019; Zhang et al., 2021; Nabeh et al., 2022) with only a few known exceptions (Rickard, 1969b; Ivarson and Hallberg, 1976; Donald and Southam, 1999; Thiel et al., 2019; Zhou et al., 2019; Berg et al., 2020; Duverger et al., 2020).

For a quarter of a century, there has been a consensus on two reaction pathways for pyrite formation: the polysulfide pathway (Rickard, 1969a, 1975) and the H₂S pathway (Rickard, 1997; Rickard and Luther, 1997). Others have fallen into disuse because they can be interpreted as variants of the two main pathways (Table 2). Recently a third reaction pathway has been proposed under the name ferric-hydroxide-surface pathway (FHS; Wan et al., 2017).

Table 2. Example of proposed reactions for pyrite formation.

The role of polysulfides in pyrite formation has long been known. Apart from the synthesis of pyrite by metallurgical processes requiring very high temperatures, two almost concurrent works attesting to the formation of pyrite from polysulfides at relatively low temperatures go as far back as the mid-19th century. During an expedition to Iceland following the 1845–1846 Hekla eruption, the formation of pyrite was observed in fumarole systems (Bunsen, 1847). Based on the contemporaneous chemical knowledge about alkaline polysulfide dissolution of ferrous sulfide, which then reprecipitates, it was postulated that hot hydrogen sulfide vapours emitted by fumaroles transform the iron minerals in weathered basalt into pyrite through the transient formation of iron sulfide and polysulfides. Several years later, in pioneering hydrothermal synthesis experiments, several metal sulfide minerals including pyrite were successfully formed from a mixture of iron salts and ‘persulfide’ (obsolete word for polysulfide S_n^2-) in a sealed tube heated to 165–180°C (Senarmont, 1851). Later investigations on hydrothermal synthesis of iron sulfide showed that pyrite could be formed by heating various mixtures of H₂S and iron salts, elemental sulfur and iron sulfide or dissolved polysulfide and ferrous salts, reproducing acidic to alkaline environments (Allen et al., 1912). The authors generalized these three different methods by “the action of sulphur on ferrous sulphide”, laying the foundations for a common pathway to pyrite formation regardless of the reagents used. The idea that elemental sulfur is directly involved in the formation of sedimentary pyrite subsequently gained ground, not least because of its common presence in the sediments (GW Harmsen, 1954; Kaplan et al., 1963; Berner, 1970; Sweeney and Kaplan, 1973).

Because the solid-state reaction between elemental sulfur and iron sulfide at low temperatures is mechanically impossible, it is now generally accepted that pyrite formation resulting from the addition of a sulfur compound to iron monosulfide occurs via the dissolved polysulfide reagent. This reaction is known as the polysulfide pathway (Rickard, 1975; Luther, 1991). The reaction of HS^- with elemental sulfur (Reaction 1, Table 2) is a ubiquitous source of dissolved polysulfide which helps explain how pyrite can easily be formed anywhere from modern sediments to early experiments with elemental sulfur and iron sulfide. Isotopic measurements have revealed that pyrite formed by the polysulfide pathway inherits the polysulfide isotopic signature, meaning that the sulfur atom from the initial FeS is replaced by two sulfur atoms of the polysulfide via a cyclic reaction (Butler et al., 2004).

The polysulfide pathway is probably of relevance in transitional redox environments with limited molecular oxygen content (Rickard, 1997). In addition to oxygen, other oxidants could promote this pathway in the environment. For instance, H₂S released by sulfate-reducing microorganisms (SRM) can be oxidized to polysulfide during abiotic ferric iron reduction (Wei and Osseo-Asare, 1997; Morin et al., 2017; Baya et al., 2022). It is interesting to note that under strictly anoxic conditions, the few studies reporting pyrite formation in microbial cultures of SRM have mostly been carried out in the presence of ferric iron: goethite (FeO(OH), Rickard, 1969a), jarosite (KFe₃(SO₄)₂(OH)₆, Ivarson and Hallberg, 1976) or ferric phosphate (FePO₄.2H₂O, Berg et al., 2020; Duverger et al., 2020). In these various works, amorphous black precipitates were initially observed and pyrite was only detected after longer incubation times. It would therefore seem that the H₂S produced by microbial sulfate reduction (MSR) reacted with the ferric iron minerals, producing not only ferrous iron and subsequently iron sulfide but also intermediate sulfur compounds paving the way for pyrite formation via the polysulfide pathway. A similar mechanism is probably at play in sulfur-disproportionating bacterial cultures that grow only with the addition of Fe(III) (oxyhydr)oxides as a hydrogen sulfide scavenger (Thamdrup et al., 1993; Canfield et al., 1998; Finster et al., 1998). Polysulfides can also be directly produced by microorganisms and eventually released into the environment via polyvalent metabolisms such as incomplete sulfide oxidation (Berg et al., 2014; Findlay, 2016). The rate law for pyrite formation via the polysulfide pathway was derived by Rickard (1975), with a correction later published due to incorrect unit conversion in the original publication (Wan et al., 2017).

The thermodynamic feasibility of pyrite formation via the H₂S pathway, even at low-temperature, biological conditions, was known long before being experimentally demonstrated (Berner, 1970; Rickard, 1997; Rickard and Luther, 1997). The first observations of pyrite formation coupled with dihydrogen production came from hydrothermal experiments (Wikjord et al., 1976, 1980; Taylor et al., 1979). In highly reducing environments where stronger oxidants are absent, H₂S can oxidize FeS to pyrite (Rickard, 1997; Rickard and Luther, 1997). This direct reaction leads to the production of H₂, which can then reduce carbon dioxide to organic molecules in the presence of a mineral or enzyme catalyst. In fact, the formation of pyrite via the H₂S pathway helped fuel the iron-sulfur world hypothesis for the origin of life (Wächtershäuser, 1988, 1992; Drobner et al., 1990). Although another mechanism, known as the iron-loss pathway, was proposed to explain the production of dihydrogen by the oxidation of two FeS molecules to pyrite and dissolved Fe²⁺ (Schoonen and Barnes, 1991; Wilkin and Barnes, 1996), it is now considered to be derived from the H₂S pathway (Butler et al., 2004). There is evidence from the sulfur isotopic composition of pyrite formed via the H₂S route that it is a product of the equimolar mixing of the FeS and H₂S pools (Butler et al., 2004).

Although HS^- is simply the deprotonated form of H₂S, differences in electron orbital energy levels make HS^- an incompatible electron acceptor in the H₂S pathway. For this reason, the H₂S pathway was initially suggested to be dominant in neutral to acidic environments, while the polysulfide pathway is dominant under alkaline conditions (Rickard and Luther, 1997; Thiel et al., 2019). The favourability of the reaction is also influenced by the microbial scavenging of H₂. For example, hydrogenotrophic methanogens growing together with sulfate-reducing bacteria probably maintain this reaction in pyrite-precipitating enrichment cultures (Thiel et al., 2019). The H₂S pathway is assumed to be more widespread in strictly anoxic subsurface sediments (Rickard, 1997) and could have been the dominant pyrite formation mechanism in the ferruginous Archean Ocean before the Great Oxygenation Event (Lyons et al., 2014).

Nonetheless, studies have shown that mackinawite and H₂S can remain incredibly stable over months in strict anoxic conditions even during hydrothermal processing (Benning et al., 2000; Cahill et al., 2000; Swanner et al., 2019), in contrast to the fast rates of the H₂S pathway initially reported in Rickard (1997). These studies demonstrated that slightly oxidized iron sulfide precursors actually drove pyrite formation. Anaerobic cultivation techniques and sensing technologies used to measure oxygen have evolved a long way since early experiments (Wikjord et al., 1976; Taylor et al., 1979; Drobner et al., 1990; Rickard, 1997; Rickard and Luther, 1997), so one cannot exclude the possible influence of very low oxygen concentrations in previous studies. In these cases, the surface oxidation of iron sulfide precursors may have been sufficient to drive pyrite formation under not strictly anoxic conditions, perhaps via a combination of different pathways.

Since FeS formation can be directly linked to the sulfidation of Fe(III) (hydr)oxides, it eventually became clear that surface-mediated reactions could play a role in pyrite formation. Nevertheless, it required a combination of bulk chemical and modern micro- to nano-scale imaging methods to finally elucidate the steps following the electron transfer reaction between sulfide and Fe(III) (hydr)oxides. In experiments mixing dissolved sulfide and lepidocrocite, TEM, XRD and Mössbauer were used to observe a very rapid (<2 h) surface reaction with the formation of a mackinawite rim on lepidocrocite, which eventually dissolved at the expense of pyrite precipitation delocalized from the lepidocrocite surface (Hellige et al., 2012). Further investigations with iron minerals of different crystallinity (ferrihydrite < lepidocrocite < goethite) revealed that sulfidation proceeds at different rates and that the formation of FeS (hours) and pyrite (weeks) is decoupled in time (Peiffer et al., 2015). This led to the proposal of a third pathway for pyrite formation catalysed by Fe(III) minerals. The so-called ferric hydroxide surface pathway (FHS) begins with the formation of surface-bound, non-sulfur-associated Fe(II) (>Fe(II)OH₂⁺) formed through surface complexation reactions (Table 2). These iron-hydroxy groups react with sulfide radicals to form surface-bound Fe(II)S₂⁻ species and induce pyrite nucleation in conditions below supersaturation by creating a new equilibrium with the aqueous phase. The FHS pathway is predicted to be prominent in aquatic systems with abundant Fe(III) minerals, i.e. with terrestrial influence, (Wan et al., 2017) because the reaction only proceeds if the precursor species >Fe^IIS₂^– incompletely covers the ferric hydroxide surface. Other compounds, like organic matter, have been found to mask ferric hydroxide reactive sites, slowing the formation of pyrite via the FHS pathway from 120 days (Wan et al., 2017) to 12 months (ThomasArrigo et al., 2020). Interestingly, the addition of dissolved (complexed) Fe³⁺ has commonly produced pyrite in abiotic experiments (Wei and Osseo-Asare, 1997; Morin et al., 2017; Baya et al., 2022) but only metastable iron sulfide phases in microbial experiments (Bertel et al., 2012; Ikogou et al., 2017). The few successes in obtaining biogenic pyrite in laboratory experiments have used solid Fe(III) minerals (Rickard, 1969b; Ivarson and Hallberg, 1976; Berg et al., 2020; Duverger et al., 2020). Together, these results suggest that further studies on sulfidization pathways on common iron mineral phases should include organic matter and microorganisms in order to better represent environmental conditions.

Biogenic iron sulfide mineral formation at high temperatures

The black smokers as Fe- and S-rich systems

The most extreme hyperthermophilic microorganisms have been isolated from hydrothermal vents (Huber et al., 1991; Blöchl et al., 1997; Takai et al., 2001), which represent the most biologically active sites in the deep ocean found along mid-ocean ridges and discovered in 1977 (Corliss et al., 1979; Hannington et al., 1995). These environments are characterized by unique physical and chemical properties such as high hydrostatic pressures, high temperatures and often highly dissolved metal contents. Among those, black smokers are iron- and sulfur-rich anaerobic systems (Holden et al., 2012) that form chimneys up to 45 metres tall (Hannington et al., 1995). Cold seawater infiltrates down through the oceanic crust, is heated in contact with the magma chamber, becomes less dense, and then rises to the seafloor, dissolving metals and sulfides from the surrounding basaltic rocks (Humphris and Mccollom, 1998; Fouquet et al., 2010). Sulfur species are abundant and present in several redox states in hydrothermal environments (from –2 to +6), both in inorganic and organic forms. Inorganic sulfur appears in various forms such as metal sulfides (chalcopyrite CuFeS₂, pyrite FeS₂, sphalerite ZnS) (Tivey and Delaney, 1986; Peng and Zhou, 2005), as well as nanoparticles of elemental sulfur (S⁰), hydrogen sulfide (H₂S), hydrosulfide (HS^-), hydrogen polysulfides (HS_n^-) and polysulfides (S_n^2-) (Schwarzenbach and Fischer, 1960; Rickard and Luther, 2007; Gartman et al., 2011). These are predominantly found in plumes and in the reducing parts of hydrothermal chimneys (Findlay et al., 2014; Findlay, 2016). The presence of both oxidized and reduced sulfur compounds in the hydrothermal ecosystem supports the development of metabolically diverse microorganisms (Orcutt et al., 2011; Dick, 2019). Sulfides and hydrogen sulfide can also result from MSR and/or reduction of elemental sulfur. Consequently, the sulfide content of black smokers originates from a mixture of abiotically produced sulfide through high-temperature chemical reactions and biotically produced sulfide via microbial activities, complicating the interpretation of sulfur-based biosignatures in such ecosystems.

Iron is found at very high concentrations, up to 25 mM (Holden and Adams, 2003; Tivey, 2007; Toner et al., 2016). In the anaerobic, reducing and high-temperature hydrothermal fluid, iron is mainly present as the soluble ferrous form Fe²⁺, while in the surrounding oxygenated seawater, iron is rapidly oxidized into the ferric form Fe(III) and precipitates as iron (oxyhydr)oxide minerals (Rickard and Luther, 2007; Scholten et al., 2019). In the fluid (>250°C), Fe²⁺ reacts with sulfide to form inorganic massive iron sulfide deposits, beginning with FeS mackinawite/ pyrrhotite, which is thermodynamically unstable, and then rapidly evolving into FeS₂ pyrite (Rickard and Luther, 2007), the major sulfide component of the interior of black smokers (Fouquet et al., 1997). Whereas Fe₃S₄ greigite is considered an intermediate phase in the process of pyrite formation (Hunger and Benning, 2007), no occurrence of greigite has been reported in hydrothermal chimneys. Pyritization is thus the main process in black smokers but the mechanism of pyrite formation is still being debated, especially for pyrite forming at relatively low temperatures (<150°C) in the external parts of chimneys, which may involve the living compartment (Juniper and Martineu, 1995; McCollom, 2007).

To elucidate the process of microbial synthesis of pyrite, several experiments have aimed to mimic environmental conditions in laboratory experiments using mesophilic microorganisms to produce pyrite (see section ‘Biogenic pyrite formation at low temperature’). But pyrite formation in direct connection to the activity of hyperthermophilic microorganisms has rarely been reported in the literature (Stetter et al., 1983; Gorlas et al., 2018, 2022; Truong et al., 2023).

Pyrite formation by hyperthermophiles

More than two decades ago, Stetter and colleagues published the first observation of pyrite formation resulting from a biogenic process at a high temperature (Stetter et al., 1983). Pyrite was formed during the growth of the sulfur-reducing archaeon Pyrodictium occultum in a medium containing a final concentration of 10 μM FeSO₄, in a fermenter at 95°C. This biogenic pyrite formation could be explained by an interaction between the metabolic end-product H₂S, formed from sulfur reduction, with dissolved Fe²⁺ according to the H₂S pathway (Table 2).

Unfortunately, there have been no further studies to confirm or refute this following these initial observations. Recent experiments were specifically designed to investigate the formation of Fe-S minerals by hyperthermophilic Archaea belonging to the Thermococcales order (Gorlas et al., 2018), which are predominant inhabitants of the hottest parts of hydrothermal environments (Takai et al., 2001). Under laboratory conditions, the sulfur-reducers Thermococcales are able to generate significant amounts of pyrite within a few hours and could be an important contributor to pyrite formation in their ecosystem (Gorlas et al., 2022). Cells grown in a modified rich medium supplemented with S⁰ (1 g/L or 31 mM) were incubated with a solution of ferrous sulfate (5 mM of FeSO₄), under strict anoxia at 85°C (Gorlas et al., 2018, 2022). The amount of pyrite production increases as the mineralization duration lengthens (Truong et al., 2023). Pyrite nanocrystals are consistently found to have a close association with the cells and with sulfur vesicles produced by Thermococcales during their growth (Gorlas et al., 2022). This leads to the production of pyrite spherules with ultra-smooth surfaces having diameters ranging between 200 nm to 1 μm (with a marked abundance maximum around 1 μm) made up of an assemblage of very numerous domains ranging between 5 and 15 nm (Truong et al., 2023). This assembly of small domains explains the ultra-smooth appearance of pyrite spherules. Interestingly, the presence of pyrite was detected only when Thermococcales were cultivated in a growth medium initially containing zero-valent sulfur S⁰ (Gorlas et al., 2022). In these growth conditions, Thermococcales cells internalized high sulfur concentrations leading to the production of numerous sulfur vesicles probably derived from polysulfides. This process has been interpreted as a polysulfide detoxification mechanism (Gorlas et al., 2015). Gorlas et al. (2022) proposed that in the presence of abundant Fe²⁺, those sulfur vesicles allowing reactive sulfur to be exposed at cell surfaces could act as precursors for pyrite formation. Conversely, under growth conditions in which Thermococcales do not produce such sulfur vesicles but sulfide instead, no formation of pyrite occurs, suggesting that pyrite is produced through Thermococcales preferentially via the polysulfide pathway rather than the H₂S pathway (Gorlas et al., 2022).

The initial occurrence of the FeS nano-mackinawite, observed in short-term experiments (i.e. after 5 hours of mineralization) (Truong et al., 2023), could be attributed to the interaction between Fe²⁺ and the H₂S produced during Thermococcales growth. Thus Truong et al. (2023) proposed that the formation of pyrite particles induced by the presence of Thermococcales and their sulfur vesicles occurs due to a redox comproportionation of S⁰ (from elemental sulfur) and sulfide (S-II) (from FeS) to yield S(-I) in pyrite particles. This suggests a high degree of metabolic adaptability by Thermococcales. This capability allows Thermococcales to adjust their metabolic activities in response to variations in sulfur and iron availability, which may fluctuate in hydrothermal vent environments.

Greigite formation by Thermococcales

Thermococcales have also been recognized for producing cuboidal extracellular nanocrystals of Fe₃S₄ greigite within a few days, with sizes ranging from 40 to 60 nm, regardless of the presence or absence of sulfur vesicles (Gorlas et al., 2018, 2022; Truong et al., 2023). Greigite formation occurs in close proximity to cells and vesicles that yield pyrite spherules. While conventional models of greigite formation from FeS typically involve an excess of sulfur or an iron loss pathway (Wilkin and Barnes, 1996), Gorlas et al. (2018) proposed that those greigites related to the biological activity of Thermococcales were formed by sulfurization of amorphous Fe(III)-bearing phosphates loaded onto cellular debris. Production of greigite was also observed during the growth of the hyperthermophile methanogen Methanocaldoccus jannaschii at 80°C when hematite was added to the growth medium (Igarashi et al., 2016). Following the reduction of sulfur by the methanogen, hematite was reduced to form amorphous FeS, subsequently reacting with residual Fe(III) into greigite-like nanoflakes. The presence of Fe(III), either because it is added or because it is intrinsically produced, therefore seems to be a sufficient condition to form greigites in these strictly anoxic environments. In the case of Thermococcales, the mechanism by which Fe(III) is generated has yet to be elucidated (Kish et al., 2016). Finally, it is likely that over long periods of time, some of the greigites formed by these processes subsequently evolve into pyrite, but it is clear that in the case of Thermococcales, greigite formation is not just an intermediate towards pyrite but constitutes a pathway on its own.

The efficiency with which Thermococcales produce pyrite and greigite in laboratory settings suggests that they are able to replicate this biomineralization process in their native environment. Consequently, the hyperthermophiles Thermococcales may play a significant role in the formation of ‘low temperature’ pyrite in their ecosystem. Furthermore, this biologically induced mineralization mechanism of Fe-S minerals by Thermococcales could be a key component of their survival strategy to thrive in highly mineralized, high-temperature environments (Gorlas et al., 2022).

Integrating various pyrite formation pathways: Connecting experiments to the environment

Recent experiments have shed light on how microorganisms contribute to pyrite formation beyond the simple role of SRM as providers of sulfide. In this section, we synthesize the relevance of these experiments to natural environments.

As discussed in previous sections, there is a consensus that biogenic pyrite can form via three distinct pathways. In nature, while each of these pathways may be more important under certain conditions, it is likely that all three will operate at the same time. This is in fact what is observed in recent cultivation work (see previous two sections), in which microorganisms (sometimes as a community) affect both the Fe and S cycles, generating intermediates that promote pyrite formation. Therefore, it is important to consider the interactions and processes associated with all the pathways rather than viewing them separately (Fig. 2).

Figure 2. Summary of interactions between Fe and S cycles driven by abiotic and microbial processes to generate Fe sulfide minerals. In this figure, the H₂S, polysulfide and FHS pathways are considered together rather than separately. SRM (blue): sulfate-reducing microorganisms, SOM (green): sulfur/sulfide oxidizing microorganisms, IRM (red): iron-reducing microorganisms.

Biogenic pyrite formed in experiments have adopted either a micrometric spherulitic (Berg et al., 2020; Duverger et al., 2020; Truong et al., 2023) or euhedral morphology (Thiel et al., 2019; Allen et al., 2021), with no reported formation of framboids. It has long been suggested that the unique framboidal morphology is a strong indicator of biological activity, but evidence for that is lacking (Runge et al., 2024). The research focus is now shifting to trying to understand the origin and abundance of pyrite spherules. Reports of framboids and euhedral pyrite are common in nature, perhaps due to their easily recognizable morphology even in complex sediments. By contrast, micrometric pyrite spherules have only recently been reported (Truong et al., 2024) but could be more common once greater scrutiny is applied.

Lastly, as the morphology of pyrite is controlled by precipitation rate (among other factors; see Raiswell, 1982; Butler and Rickard, 2000; Runge et al., 2023, 2024), it is important that the formation rates of biogenic pyrite in experiments match those observed in nature. To this end, we have expanded the pyrite precipitation rate dataset from Mansor and Fantle (2019) to include older data from salt marshes (based on³⁵ S incorporation into pyrite; Howarth and Giblin, 1983; Howarth and Merkel, 1984) and recent data from microbial (Berg et al., 2020; Gorlas et al., 2022; Truong et al., 2023) and abiotic experiments (Hockmann et al., 2020; ThomasArrigo et al., 2020; Baya et al., 2021; Domingos et al., 2023; Runge et al., 2023, 2024). As can be seen in Fig. 3, most microbial pyrite (0.0009–0.35 mM/day) is precipitated at the same rates as in salt marshes (0.0009–0.35 mM/day) and the upper estimates of marine sediments (10^-5 to 0.09 mM/day). The only exception is the extremely fast pyrite formation by S⁰ disproportionaters (0.2–4.3 mM/day; Canfield et al., 1998). This comparison generates confidence that lab studies are indeed relevant to nature. Rates in natural hydrothermal systems were unfortunately not readily available from the literature for comparison.

Figure 3. Compiled pyrite precipitation rates in the environment and in biological and abiotic experiments. The figure was updated from Mansor and Fantle (2019) with additional data from: salt marshes – Howarth and Giblin (1983); Howarth and Merkel (1984); framboids – Rickard (2019); microbial 24–35°C – Thiel et al. (2019); Berg et al. (2020); microbial 85°C – Gorlas et al. (2022; Truong et al. (2023); abiotic at 25°C with Fe(III) minerals – Hockmann et al. (2020); ThomasArrigo et al. (2020); abiotic at 40–100°C with wet FeS or magnetite - Domingos et al. (2023); Runge et al. (2023, 2024) and abiotic at pH 5–6 at 25°C – Baya et al. (2021).

It is interesting to note that pyrite precipitates faster in salt marshes than in marine sediments, perhaps due to its more dynamic oxic-anoxic cycles that generate redox-active intermediates (e.g. S⁰ or polysulfides). It is also interesting to note that rates derived for pyrite precipitation within plant cells (Rickard et al., 2007) and the theoretical estimates for framboids (Guilbaud et al., 2011; Rickard, 2019; based on the burst nucleation model) are around 800–17,000 mM/day, which are orders of magnitude higher than most of the dataset. Whether these anomalously fast rates are valid or not remains to be determined. One explanation could be that these rates reflect precipitation in microenvironments with locally enhanced supersaturation, while the rest of the dataset primarily reflects bulk rates averaged over at least centimetre scales. In sediments, the site of pyrite formation is known to be heterogenous and to be enhanced around organic matter, clays or shell remains (e.g. Marin-Carbonne et al., 2022). Hence, the elevated rates could in theory be also achievable in cultivation experiments and at the same time remain relevant to nature.

Fe-S minerals and the origin of life

Fe-S minerals probably played a pivotal role in the genesis of life by actively participating in the generation of prebiotic molecules (Picard et al., 2021) and/or by preserving ancient traces of life (Wacey et al., 2011; Baumgartner et al., 2019, 2020). Ferredoxin, one of the oldest biological catalysts, contains Fe-S clusters, which suggests an ancient origin. Examples of Fe-S clusters are widespread in biogeochemistry, where they serve as active centres in essential proteins, including NADH dehydrogenase, coenzyme Q – cytochrome C reductase, hydrogenases and nitrogenase (Beinert, 2000; Johnson et al., 2005; Huang et al., 2024).

Wächterhauser introduced the ‘iron sulfur world’ prebiotic model in 1988, proposing that iron sulfides in hydrothermal conditions could facilitate the formation of prebiotic molecules such as amino acids. This theory suggests that mineral surfaces provide the catalytic properties necessary for the formation of simple organic molecules from inorganic compounds. In high temperature, high-pressure conditions, H₂ and CO₂ can interact with metal sulfides leading to the synthesis of amino acids, peptides (Bonfio et al., 2017) and eventually nucleotides, the building blocks of life (Table 3). Pyrite precipitation provides the energy required for CO₂ reduction, leading to the creation of reduced sulfur organic molecules (Huber and Wächterhäuser, 1997). Negatively charged organic molecules can then bind to the positively charged pyrite surface, catalysing their transformation into more complex molecules and enhancing molecular diversity. Experimental successes include the production of thiolated formic acid and CH₃COSH, a precursor for the citrate cycle acetyl coenzyme A (Huber and Wächterhäuser, 1997). Thermodynamic calculations have also supported the formation of amino acids in hydrothermal conditions (Amend and Shock, 1998 and references therein). Notably, experiments on other sulfides, such as chalcopyrite or sphalerite, have not yielded amino acids, emphasizing the unique role of Fe-S minerals (and sometimes nickel sulfides) in these prebiotic processes (Schreiner et al., 2011).

Table 3. Production of reduced organic molecules from CO₂ reactions with metal.

An alternative theory suggests that mackinawite or greigite can serve as catalytic centres, with FeS bubbles acting as proto-cell membranes, instead of lipids, that promote the formation of organic molecules (Russell et al., 2010). In this model, FeS bubbles can precipitate when alkaline fluids are injected into acidic Fe²⁺-rich solutions, mimicking conditions in hydrothermal regions of the Hadean Ocean. The geochemical pH gradient then provides energy through the FeS membrane (Fig. 4). Nucleic acids can bind to mackinawite nanoparticles, as demonstrated by Hatton and Rickard (2008). This hypothesis suggests that the compartmentalization of membranous FeS precipitates is crucial for maintaining a chemical gradient needed for carbon fixation metabolism (Fig. 4).

Figure 4. Various examples of reactions relevant to the origin of life that involve Fe-S minerals, (a) CO₂ fixation schematic modified from De Graaf et al. (2023), based on prior experiments (Herschy et al., 2014; Sojo et al., 2016; Hudson et al., 2020), (b) prebiotic metabolic reaction in a protocell modified from Alpermann et al. (2011) and (c) RNA-peptide co-evolution around hydrothermal vents. Bubbles from vents could form a membrane associated with iron sulfides. RNA bound to the minerals could act as a template for peptide formation (Russell and Hall, 1997).

Moreover, iron sulfide surfaces can also support an autocatalytic chemolithotrophic metabolism driven by the exergonic formation of pyrite, leading to a chemoautotrophic origin of primordial metabolisms (Wächtershäuser, 1988; Cody, 2004). The hypothesis of protometabolism catalysed by ancient iron sulfur active centres has been successfully tested (Bonfio et al., 2017). Recent studies have highlighted that greigite can catalyse the fixation of CO₂ under hydrothermal alkaline conditions, thus potentially preceding the enzyme route for the acetyl-CoA pathway (Preiner et al., 2020). Similarities between the spatial organization of enzymes and Fe-S minerals have been recognized, suggesting an abiotic origin of these catalysts (Russell et al., 2014). Other experimental studies have indicated that Fe-S minerals (Fig. 4) possess the capability to support a proto-metabolic process (Lazcano and Miller, 1999; Cody, 2004; Goldford et al., 2019).

An ancient origin for Fe and S microbial metabolisms is supported by phylogenetic studies and other geochemical indicators (Lepot, 2020; Lyons et al., 2024). Indeed, microorganisms that are closely related to the last common ancestor are mainly anaerobic and sulfur-reducing hyperthermophiles. Isotopic studies on ancient pyrite sediments have constrained the antiquity of microbial sulfate reduction at 3.5 Ga (Shen and Buick, 2004), with further evidence at 2.7 Ga (Archer and Vance, 2006; Marin-Carbonne et al., 2018). Experimental studies have also shown that both archaea and bacteria closely related to the last common ancestor can reduce Fe (III) (Vargas et al., 1998; Lovley et al., 2022), suggesting an early origin of dissimilatory iron reduction. Geochemical studies on ancient pyrite have thus identified isotopic fingerprints of this metabolism in the geological record, from 3.2 Ga (Marin-Carbonne et al., 2020), to the late Archean (Archer and Vance, 2006; Craddock and Dauphas, 2011; Czaja et al., 2016). These findings highlight the important role that microbial metabolism played in shaping the early Earth’s environment and lay the groundwork for the identification of biosignatures in the sedimentary record.

Sedimentary pyrite as environmental proxies and biosignatures

Background

In contrast to metastable iron sulfides, pyrite is prevalent in sedimentary rocks, with its presence dating back to approximately 3.8 billion years ago (Ga) (Smith et al., 2005). The ubiquity of pyrite in sedimentary rocks, as shown in Fig. 1, has led to the use of pyrite abundance, isotopic characteristics, and trace metal concentrations to reconstruct past environmental conditions, aiding in the understanding of historical atmospheric oxygen levels, as well as global sulfur and iron geochemical cycles over Earth’s geological history. For instance, the recognition of detrital pyrite grains apparently eroded by flowing water serves as a robust indicator of low oxygen (O₂) levels before approximately 2.4 Ga (Johnson et al., 2014). This timing aligns with the loss of sulfur mass-independent fractionation (S-MIF, see box), which is commonly attributed to the shift from an anoxic atmosphere (pO₂ < 10^-15 PAL) to an oxic atmosphere (Farquhar et al., 2000).

Furthermore, the primary mechanism for the long-term storage of reduced sulfur species on geological timescales is the formation and subsequent burial of pyrite. This process, in conjunction with the burial of organic carbon, contributes significantly to maintaining the oxidized surface conditions of Earth (Canfield and Farquhar, 2009). Consequently, methods have been developed to gauge past oxidation states and water chemistry by examining the iron (-sulfur) mineralogy in sediment and sedimentary rocks (Raiswell and Canfield, 2012; Raiswell et al., 2018). These methods are grounded in the differential reactivity of iron phases towards sulfide, and thus, rely on pyrite precipitation. The presently utilized framework (Poulton and Canfield, 2005), known as the ‘Fe speciation’ proxy, involves quantifying the ratio of highly reactive iron to total iron (Fe_HR/Fe_T) and the ratio of pyrite to highly reactive iron (Fe_PYR/Fe_HR). The application of Fe speciation data, whether derived from stratigraphic variations or collected across various locations over time or space, has significantly contributed to our understanding of the evolution of ocean redox conditions (Poulton, 2021 and references therein). It has revealed that during the Archean and Phanerozoic eras, oceans were predominantly ferruginous (characterized by anoxic conditions and high Fe²⁺ levels) and oxygenated, respectively. In the Proterozoic era, rising oxygen levels facilitated the development of euxinic conditions at mid-depth, particularly in restricted basins while the deep ocean maintained predominantly ferruginous conditions (Poulton et al., 2004a; Planavsky et al., 2011; Ostrander et al., 2019). It is important to note that recent concerns have been raised regarding the effectiveness of the extraction protocol in retrieving the targeted mineral phases (Hepburn et al., 2020; Slotznick et al., 2020) and the influence of early diagenesis (Eroglu et al., 2021; Hutchings and Turchyn, 2021; Pasquier et al., 2022). These concerns question the capacity of the ‘Fe speciation’ approach to accurately constrain the chemistry and oxidation state of water columns, both in the present and in the past (Pasquier et al., 2022).

Texture and composition

Pyrite grains display a large range of sizes, spanning from centimetres to nanometres, and a variety of textures, e.g. euhedral/anhedral, nodule and framboidal are among the most common (see Fig. 1). Framboids, in particular, are characterized as spherical to subspherical clusters comprised of numerous microcrystals of pyrite, predominantly found in sedimentary environments (Wang and Morse, 1996; Wilkin and Barnes, 1997; Rickard, 2019). Irrespective of its shape, during its crystallization, pyrite has the capability to integrate various trace elements (TE), with chalcophile and siderophile elements being the most frequently encountered. These TE can be integrated into pyrite through two distinct mechanisms: either by substituting for Fe or S within the pyrite structure, or by existing as inclusions of distinct mineral phases or amorphous masses enclosed within the pyrite matrix (Gregory, 2020 and references therein). Of particular interest in enhancing our understanding of past environmental conditions is the observation that the TE content in pyrite is directly linked to the TE content of the water from which it precipitates (Gregory et al., 2014). Consequently, TE within pyrite has been utilized to decipher the formation of ore deposits (Kusebauch et al., 2019) and to unravel the evolution of ocean and atmospheric chemistry (e.g. Large et al., 2014; Gregory et al., 2017). Nevertheless, recent investigations have revealed that pyrite’s TE concentration varies with depth within the sediment, ultimately reaching a plateau in composition (Gregory et al., 2022a). This suggests that the relationship between the water column and pyrite’s TE concentration is more intricate than previously assumed. Hence, TE contents in pyrite not only reflect the composition of the water column but also the composition of pore water, maybe allowing for tracking the release of trace elements from organic matter and/or iron (oxyhydr)oxides during diagenesis, particularly when examined at a finer spatial resolution (Tribovillard et al., 2006; Gregory et al., 2022b; Atienza et al., 2023). Biogenic pyrite experiments in the presence of TE have not been performed and it is unclear how biological activities can affect the incorporation of TE into pyrite. A recent study has also highlighted the importance of pyrite growth via particle attachment under certain conditions, which might impact the morphology and distribution of trace metals and isotopes within pyrite (Domingos et al., 2023).

Box: Isotope notation

Most chemical reactions, including those involved in the formation of pyrite, distribute isotopes proportionally to their mass. Here, we illustrate this concept using the four stable sulfur isotopes (³²S, ³³S, ³⁴S and ³⁶S) as an example. A similar analysis can be conducted with the iron isotopic system (Dauphas et al., 2017).

In Mass-Dependent Fractionation (MDF), it is anticipated that the enrichment or depletion in ³⁴S is roughly twice and half of that in ³³S and ³⁶S, respectively. The theoretical slopes of the ³³S-³⁴S and ³⁶S-³⁴S MDF, denoted as ³³λ and ³⁶λ, are 0.515 and 1.89, respectively (Young et al., 2002; Farquhar et al., 2003). However, various physical and (bio)chemical processes exhibit mass dependencies that subtly deviate from these values (Farquhar et al., 2003; Ono et al., 2006; Johnston et al., 2007). When these slight variations appear in the isotopic composition of pyrite (and other materials), they are commonly represented as Δ^3xS, where the superscript 3x denotes one of the rare isotopes of sulfur, ³³S, or ³⁶S. In its linear definition, Δ^3xS = δ^3xS – ^3xλ x δ³⁴S, where δ^3xS=[(^3xS/³²S)_sample / (^3xS/³²S)_standard-1]*1000. The values of δ^3xS are reported in permil (‰) with respect to the sulfur isotope international standard Vienna Canyon Diablo Troilite (VCDT). The observation that different processes exhibit distinct mass dependencies (e.g. Young et al., 2002; Johnston et al., 2007; Zerkle et al., 2009; Johnston, 2011; Wing and Halevy, 2014; Eldridge and Farquhar, 2018) suggests that studying MDF may offer a more comprehensive understanding of how biological processes and/or local sedimentological conditions influence sulfur cycling in marine sediments.

A notable characteristic of early Earth sulfur-bearing material is the preservation of ³³S-³²S and ³⁶S-³²S ratios that significantly deviate from MDF, known as Mass-Independent Fractionation (MIF, or S-MIF; (Farquhar et al., 2000). Based on experimental SO₂ photolysis and an atmospheric chemistry model, the conservation of substantial and variable S-MIF signals in Archean and early Paleoproterozoic sedimentary rocks necessitates an extremely low partial pressure of O₂ (Farquhar et al., 2001; Catling and Zahnle, 2020). Therefore, the S-MIF signal is considered one of the strongest pieces of evidence for an anoxic atmosphere before ≈2.4 Ga. Although the atmospheric production of S-MIF is well established, the exact underlying mechanism and the identity of its carriers to the surface remain unclear and subject to intense debate (Halevy et al., 2010; Halevy, 2013; Endo et al., 2019; Reed et al., 2022; Oduro et al., 2023).

Sulfur isotopes

Sedimentary pyrite retains a distinct isotopic signature that reflects a combination of microbial metabolic processes and physical transport and mineralization. A broad taxonomic spectrum of microbes can change the oxidation state of sulfur to gain the energy required for cellular function and growth. Three main metabolic pathways are particularly important for the sulfur isotopic composition of pyrite: microbial sulfate reduction (MSR), sulfur disproportionation, and sulfide oxidation (Fike et al., 2015; Jørgensen et al., 2019). From each of them arises a metabolic-specific microbial fractionation of sulfur isotopes which discriminates heavy isotopes, reported afterwards as ³⁴ε-³³λ (see isotope notation box).

Over the past several decades, laboratory experiments using pure cultures of SRM have yielded variable apparent isotopic fractionation, from 65‰ to ≈5‰ (Detmers et al., 2001; Johnston et al., 2005b; Hoek et al., 2006; Johnston et al., 2007; Sim et al., 2011a; Sim et al., 2011b; Sim et al., 2012; Leavitt et al., 2013; Deusner et al., 2014; Pellerin et al., 2015; Bradley et al., 2016; Smith et al., 2020). However, it was not until we understood the enzymatic processes that drive MSR that the debate over the extent of microbial fractionation could be settled (Goldhaber and Kaplan, 1980; Ohmoto et al., 1990; Rudnicki et al., 2001; Wortmann et al., 2001; Brunner et al., 2005). Recent work shows that the microbial reduction of sulfate to sulfide is catalysed by a reaction network of four enzymatic steps, all of which are reversible (Johnston et al., 2005b; Johnston et al., 2007; Sim et al., 2011a; Sim et al., 2011b; Sim et al., 2012; Leavitt et al., 2013, 2024; Bradley et al., 2016; Smith et al., 2020). The emerging picture from laboratory cultures, bio-isotopic models, and modern environments is that a large ³⁴ε_DSR predominates in natural environments, ranging from 66 to 78‰ (Wing and Halevy, 2014; Halevy et al., 2023). This substantial fractionation, which closely resembles the thermodynamic equilibrium between sulfate and sulfide (i.e. ≈70‰ at 20°C; Eldridge et al., 2016), is attributed to the inherently low cell-specific sulfate reduction rates (csSRR) in marine sediments. In other words, all steps of the enzymatic pathway are fully reversible (i.e. equilibrium between reactant and product; Wing and Halevy, 2014). The observed inverse relationship between the ³⁴ε_DSR and the csSRR is responsible for the overall deviation from thermodynamic equilibrium observed in laboratory experiments (i.e. lower apparent fractionation). This relationship between ³⁴ε_DSR and csSRR can be extended to ³³S and, consequently, one can expect sulfide produced from MSR to be characterized by large, near-equilibrium ³⁴ε_DSR and ³³λ_DSR values. Culture experiments with green sulfur bacteria that oxidize sulfide showed smaller microbial S-isotope fractionations and small but nonzero ³³λ deviations from equilibrium (Zerkle et al., 2009), whereas larger microbial S-isotope fractionations and ³³λ deviations from equilibrium were observed in pure cultures of sulfur disproportionators ( Johnston et al., 2005a; Fig. 5).

Figure 5. Comparison of S-isotope measurements in culture experiments to assess isotopic microbial fractionation, with the S-isotopic composition of pyrite preserved in natural samples over the geological record (measured by bulk and microscale techniques). Data are from: MSR (Detmers et al., 2001; Johnston et al., 2005b; Hoek et al., 2006; Johnston et al., 2007; Sim et al., 2011a; 2011b; Sim et al., 2012; Leavitt et al., 2013, 2024; Deusner et al., 2014; Pellerin et al., 2015; Bradley et al., 2016; Smith et al., 2020); S-oxidizers (Zerkle et al., 2009); disproportionators (Johnston et al., 2005a); bulk and microscale pyrite (Halevy et al., 2023). Modelled DSR refers to bio-isotopic model outputs analysed under a wide range of environmental parameters (i.e. temperature, sulfate, organic matter and Fe availabilities) expected to reflect modern marine conditions.

Once generated, a portion of this ³⁴S-depleted sulfide can react with dissolved ferrous iron (Fe²⁺) or iron-bearing minerals (Fe(III)) to form FeS. The formation of FeS and its subsequent transformation into pyrite involves a relatively small isotopic fractionation, typically not exceeding a few ‰ (>5‰; Fry et al., 1986; Böttcher et al., 1998). Consequently, pyrite captures the isotopic composition of the products of microbial sulfur metabolisms (H₂S, HS^-), including small deviations from MDF, and thus may serve as a good recorder of past microbial activity. Interestingly, bulk pyrite S-isotope data do not meet the experimental expectations above. Several non-unique explanations that involve mixing between different sulfur pools or combinations of metabolic effects have been invoked to explain the apparent mismatch, clearly demonstrating that our current methodology is not suitable for uniquely distinguishing the various microbial pathways involved in pyrite formation (see Johnston, 2011 for a recent review).

The uncertainty in the interpretation of bulk δ³⁴S_PYR data can be resolved by studying the isotopic composition of individual pyrite grains, which collectively contribute to the bulk signal (Fike et al., 2015; Marin-Carbonne et al., 2022; Bryant et al., 2023). As illustrated in Fike et al. (2015), when MSR occurs in a transport-limited system, the concentration of residual sulfate in porewater decreases, as do δ³⁴S_SO4 and Δ³³S_SO4, due to Rayleigh distillation. This results in a parallel increase in the instantaneous δ³⁴S_H2S-Δ³³S_H2S that can be preserved in the accumulated pyrite pool. The magnitude of this enrichment depends on the fraction of sulfate consumed, the connectivity of the porewater sulfate pool with the overlying seawater column, and the fractionation associated with the redox transformation itself. Most of these factors depend on the local sedimentation regime, including the sedimentation rate, the iron and organic carbon loading and reactivity (Pasquier et al., 2017; Liu et al., 2021; Pasquier et al., 2021a; Pasquier et al., 2021b; Houghton et al., 2022; Bryant et al., 2023; Halevy et al., 2023). As pyrite grains grow within the sediment, they continuously sample the evolving isotopic composition of microbially-produced sulfide throughout the sample’s burial history, and when pyrite grains grow quickly, over short durations as for framboid morphology, it is expected that a time series of increasing δ³⁴S_PYR values will be recorded within the population of pyrite grains in each sample. In such instances, the minimum δ³⁴S_PYR can be used to assess the microbial fractionation specific to the depositional environment (Marin-Carbonne et al., 2022; Bryant et al., 2023) whereas the overall distribution of δ³⁴S is more likely to reflect the diagenetic evolution over the history of the sediment burial (Halevy et al., 2023).

Assuming that MSR was established approximately 3.5 billion years ago (Shen and Buick, 2004; Mateos et al., 2023), the sulfur isotopic compositions of pyrite (denoted as δ³⁴S_PYR) are commonly utilized in both contemporary and ancient contexts to gain insights into the processes and fluxes within the global sulfur cycle. Within the framework of a global steady-state S cycle, the lowest δ³⁴S values reflect the maximum microbial fractionation whereas the departure from those low microbial values toward higher values indicates the evolution of MSR during the burial history of the sediment, irrespectively of the local and/or global oxygenation (Fig. 5). This pattern is typically interpreted as indicative of rising oxygen levels over time. This increase in oxygen may have resulted in more elevated oceanic sulfate levels and/or a reduction of the pyrite burial flux in response to more efficient organic matter oxidation (i.e. more aerobic respiration; Habicht et al., 2002; Wu et al., 2010; Leavitt et al., 2013).

To delve further into this, Archean sedimentary rocks exhibit relatively low apparent fractionation between sulfate and pyrite (Δ_PYR) values, which are often interpreted as a consequence of the limited sulfate reservoir in the ancient oceans (Habicht et al., 2002), with sulfate estimates ranging from hundreds to tens of micromoles per litre (μM) (Crowe et al., 2014). While the rates and isotopic implications of organic sulfur breakdown remain uncertain, both modelling and observations suggest that organic sulfur probably played a role in extremely low-sulfate systems (Fakhraee and Katsev, 2019). This has potential implications for understanding the pathways and isotopic compositions involved in early Earth pyrite formation. The transition from the Archean to the Early Proterozoic era is marked by the gradual oxidation of the Earth’s surface environment. This led to an increase in seawater sulfate content, resulting in the disappearance of sulfur mass-independent fractionation (S-MIF) signals (see Box) and an overall rise in Δ_PYR values (Fig. 5). Throughout much of the Proterozoic era, δ³⁴S_PYR values exhibit significant variability, which is commonly attributed to fluctuations in the pyrite burial rate and shifts in ocean redox conditions, such as changes in the extent of ocean anoxia (Emmings et al., 2022). After the oxygenation of the ocean atmosphere, the range of δ³⁴S_PYR values remained relatively constant at around -50‰ (Fig. 5), and the observed decreases during the Paleozoic era are probably associated with marine sulfate isotopic secular evolution (Fig. 5; e.g. Owens et al., 2013; Gomes et al., 2016; Raven et al., 2018). An essential aspect of the Phanerozoic era is the emergence of episodic intervals of ocean anoxia, known as Oceanic Anoxic Events (OAEs), driven by large-scale disturbances in the carbon cycle. High-resolution analysis of pyrite and seawater sulfate sulfur isotopic compositions and their utilization in isotopic box models have unveiled brief intervals of severe anoxia which led to significant drawdown of the seawater sulfate during those brief episodes of Earth’s history (100s kyr to My, Song et al., 2014; Bauer et al., 2022).

More recently, an alternative reading of the first-order pattern preserved in the S-isotope geologic record has emerged due to a growing body of evidence highlighting the importance of local sedimentary processes in shaping the preserved isotopic composition of pyrite (Fike et al., 2015; Halevy et al., 2023). These findings are based on the study of stratigraphic δ³⁴S_PYR variations within 100 kyr Pleistocene glacial-interglacial cycles, which have been linked to sedimentary parameters such as sedimentation rates and the availability and reactivity of iron and/or organic carbon, rather than changes in microbial fractionation (Bryant et al., 2023). These findings challenge the conventional view that sedimentary archives merely passively record the global sulfur cycle, especially given the long residence time of sulfate in the ocean (13 Myr, Kah et al., 2004). The emerging understanding suggests that, in addition to the expansion of the marine sulfate reservoir, sedimentary parameters through their influence on the accessibility of sulfate within the sediment play a crucial role in shaping the wide spectrum of Δ_PYR preserved in sedimentary rock records, rather than large-scale temporal shifts in the global sulfur cycle.

Iron isotopes

The formation of pyrite necessitates the presence of Fe(II), which in natural environments mostly originates from either dissimilatory iron reduction (DIR) or from abiotic sulfidization of Fe(III) (oxyhydr)oxides. When Fe(III) reducers are grown in pure cultures, they produce Fe²⁺ that exhibits a ⁵⁶Fe-depletion of approximately 2.9±0.9‰ compared to the initial Fe(III) substrates (Crosby et al., 2005, 2007). This microbial fractionation remains consistent across different Fe(III) substrates and bacterial strains, and it closely matches the thermodynamic equilibrium fractionation between Fe(III) and Fe(II) (Welch et al., 2003) (Fig. 6). In contrast, laboratory experiments involving the abiotic sulfidization of Fe(III) (oxyhydr)oxides release Fe²⁺ with a ⁵⁶Fe-depletion of approximately 0.8±0.3‰ relative to the Fe(III) minerals (McAnena et al., 2024). Interestingly, this abiotic fractionation appears to be independent of the dissolution rate of the Fe(III) (oxyhydr)oxides, which itself seems to be influenced by the mineralogy of Fe(III) or the S(-II):Fe(III) ratio.

Figure 6. Compilation of Fe-isotopes measurements during the reduction of Fe(III) minerals, from abiotic processes involved during mineral precipitation and from pyrite preserved in modern environments over the geological record (measured by bulk and microscale techniques). Data are from: dissimilatory iron reduction DIR (Crosby et al., 2005, 2007); sulfidization (McAnena et al., 2024); bulk and microscale pyrite (Dupeyron et al., 2023).

Furthermore, apart from the isotopic fractionation associated with the reduction step, the conversion of Fe²⁺ to FeS species also leads to a noticeable isotopic fractionation. Laboratory experiments have shown that this fractionation can vary between –0.8±0.2‰ (Butler et al., 2005) and an apparent equilibrium FeS-Fe²⁺ isotope fractionation of +0.4±0.2‰ (Wu et al., 2012). Near-equilibrium between Fe²⁺ to FeS in natural environments is consistent with the observed isotopic exchange between FeS and aqueous sulfide (Butler et al., 2004). Under laboratory conditions, the formation of pyrite from FeS results in pyrite that is ⁵⁶Fe-depleted by 0.5 to 2.2‰ compared to the FeS pool (Guilbaud et al., 2011; Mansor and Fantle, 2019). The overall Fe isotope fractionation between pyrite and Fe²⁺ (or Fe(III) (oxyhydr)oxides) depends on the relative importance of these reactions and the rate-dependent expression of kinetic and equilibrium isotope effects associated with these processes.

A limited number of studies (n=9) collectively provide valuable insights into δ⁵⁶Fe values associated with pyrite found in modern marine sedimentary environments, thereby enhancing our understanding of the influences of water column redox processes (e.g. Busigny et al., 2014; Rolison et al., 2018), offering a contemporary perspective on the benthic iron shuttle (Severmann et al., 2008; Scholz et al., 2014), and examining how early diagenesis can impact δ⁵⁶Fe_PYR in marine sediments (as observed in Fehr et al., 2008, 2010; Lin et al., 2017, 2018). However, it should be noted that the majority of bulk δ⁵⁶Fe_PYR data fail to align with the combined expectations derived from experimental and theoretical considerations mentioned earlier. One can expect that with recent advancements in in situ δ⁵⁶Fe_PYR analyses, they can be used to remove some of the uncertainty in interpreting bulk δ⁵⁶Fe_PYR data, particularly when conducted across a range of well-defined modern depositional settings.

By taking into account some of the effects mentioned above, researchers have employed the iron isotopic composition of pyrite, denoted as δ⁵⁶Fe_PYR, to explore how the biogeochemical iron (Fe) cycle in Earth’s oceans and the processes underlying pyrite formation may have evolved over geological history. In contrast to S-isotopes, the δ⁵⁶Fe_PYR record is fragmented, with a primary focus on significant shifts in Earth’s redox history, particularly associated with two major events of atmospheric oxygenation – the Great Oxidation Event (GOE) and the Neoproterozoic Oxidation Event (NOE). Consequently, it remains challenging to construct a comprehensive long-term narrative of the Fe cycle’s evolution.

Various environmental factors, whether in conjunction with the co-evolution of microbial life or not, have been suggested as potential explanations for the bulk negative isotopic variations that occurred prior to the GOE around 2.4 billion years ago. For instance, Rouxel et al., (2005) leveraged δ⁵⁶Fe_PYR values to glean insights into the oxygenation status of ancient oceans. Lighter δ⁵⁶Fe_PYR values were associated with partial oxidation of a vast Fe²⁺ oceanic reservoir indicative of more reducing conditions, while heavier values were indicative of more oxidizing conditions leading up to the GOE. Furthermore, the transition towards more positive δ⁵⁶Fe_PYR values has been suggested to attest to the onset of pyrite weathering (Heard et al., 2020). This would have significantly increased the availability of sulfate in the ocean, thereby affecting the interplay between kinetic and equilibrium processes during pyrite precipitation (Mansor and Fantle, 2019; Heard et al., 2020). Alternatively, the prevalence of low δ⁵⁶Fe_PYR values during the Neoarchean era (2.8–2.5 billion years ago) has been used to underscore the early emergence of DIR and the inference of a microbially driven Fe cycle (Johnson et al., 2008). These low values have also been interpreted as reflecting strong kinetic effects during the abiotic formation of pyrite from FeS precursors (Guilbaud et al., 2011). It’s important to highlight that a recent collection of microscale δ⁵⁶Fe_PYR values across the GOE presents distinctive patterns when compared with results obtained through conventional bulk analyses (Dupeyron et al., 2023). The discrepancy in the evolution of δ⁵⁶Fe_PYR between bulk and in situ measurements may be linked to a sampling bias, because the majority of traditional bulk investigations have been performed on millimetre-scale pyrite grains, often extracted from black shale matrices.

Societal impacts of Fe sulfides in modern and future environments

Sequestration of organic carbon by Fe-S minerals

Associations between Fe-S minerals and organic carbon have been hardly investigated, even though they might provide a mechanism for the long-term preservation of organic matter in rocks and sediments. Anoxic zones in the global ocean are expanding as a result of warming climate (Keeling et al., 2010; Levin and Bris, 2015; Keil, 2017; Breitburg et al., 2018; Ruvalcaba Baroni et al., 2020). Therefore, understanding interactions between Fe-S minerals and organic carbon might be of importance to predict how the biogeochemical cycles of Fe, S and C might evolve in marine environments. The importance of Fe(III) (oxyhydr)oxides for the preservation of organic carbon has been well explored and recognized because of their abundance in oxygenated surface environments (Lalonde et al., 2012; Longman et al., 2022; Moore et al., 2023). Recent experimental work highlighted the role of microbial cell surfaces in the growth and nucleation of Fe-S minerals, therefore suggesting that some type of association must form between microbial organic carbon and Fe-S minerals (Picard et al., 2018). It has been suggested that Fe-S minerals might play an important role in the preservation of organic carbon in anoxic marine sediments (Barber et al., 2017).

In laboratory experiments, Fe-S minerals can bind organic matter from various sources: live and dead microbial cells (Herbert et al., 1998; Picard et al., 2019, 2021; Nabeh et al., 2022; Truong et al., 2023), simple organic molecules (i.e. sugars, amino acids) and complex organic mixtures used in microbiological media (i.e. tryptone and yeast extract) (Picard et al., 2021; Nabeh et al., 2022), water-soluble extracts from aged compost soil, microalgae biomass and corn leaf (Tétrault and Gélinas, 2022). The capacity of mackinawite to sequester organic carbon is at least comparable to that of ferrihydrite (Wang et al., 2019; Ma et al., 2022). Light elements (e.g. C, O) can be detected in Fe-S minerals by energy dispersive X-ray spectroscopy (EDS) in the scanning electron microscope (SEM) (Picard et al., 2019). To characterize the speciation and redox state of light elements, the use of spectroscopic methods, such as X-ray photoelectron spectroscopy (XPS) and near-edge X-ray absorption fine structure (NEXAFS) spectroscopy, is required. These methods identified C and N in biogenic Fe-S minerals originating from organic molecules, and O originating from organic functional groups rather than from oxidation of Fe-S minerals (Herbert et al., 1998; Picard et al., 2019, 2021). The use of scanning transmission X-ray microscopy (STXM), coupled with NEXAFS spectroscopy, determined that organic carbon is homogenously distributed on biogenic Fe-S mineral aggregates precipitated with SRB, on biogenic pyrite and Fe-S minerals produced in cultures of sulfur-reducing archaea, and on abiotic minerals precipitated with simple and complex organic mixtures (Picard et al., 2019, 2021; Truong et al., 2023). Some organic carbon also occurs as ‘hot spots’ in mineral aggregates precipitated with SRB (whether Fe²⁺ was in the growth medium or added after growth), interpreted as the preservation of intact cells and contents (Picard et al., 2019, 2021).

In long-term laboratory experiments, Fe-S minerals could sequester significant amounts of organic carbon and nitrogen, as quantified by elemental analysis of solid phases (Nabeh et al., 2022). The sequestered amounts decreased as a function of time but stabilized rapidly within a few months (Nabeh et al., 2022). Semi-quantitative analysis from STXM/NEXAFS data also (Picard et al., 2021). The highest stabilized levels of C and N were 13.5 w/w% and 3.3 w/w%, respectively. In a marine medium, the organic carbon removal capacity of Fe-S minerals was higher when microbial biomass was present (42–51% organic carbon removed from medium) than when organic mixtures that did not contain cells (e.g. amino acid mixtures, tryptone, yeast extract) were used (1.2–5.2% organic carbon removed from medium). In freshwater medium, the removal capacity of microbial biomass was even higher than in marine medium (67–137%; values above 100% linked to uncertainties in elemental analysis of biomass) (Nabeh et al., 2022).

The nature of associations between Fe-S minerals and organic carbon has yet to be thoroughly investigated. Quantification of organic carbon in the studies reported above was done after several washes with anoxic ultrapure water, therefore removing loosely bound and water-soluble organics. Spectroscopic studies indicated that glucose and mannose bind less to Fe-S minerals than protein-rich organic mixtures (Picard et al., 2021). Synthetic mackinawite studied with Fourier transform infrared spectroscopy (FTIR) revealed that it mostly associates with polysaccharides when adsorbed to extracts from natural organic matter (corn leaf, marine algae and soil) (Tetrault and Gelinas 2022). The only study on natural pyrite (n = 5 framboids) determined fairly high organic carbon contents of 2.8–6.5 wt.% (Tribovillard et al., 2022). Overall, the studies indicated a high association between Fe-S minerals and organic carbon, which needs to be investigated further in the context of climate and long-term organic carbon preservation (Fig. 7).

Figure 7. Summary of the relevance of biogenic Fe-S minerals in various research fields and current societal issues.

Pyrite oxidation in the context of acidic drainages, metal recovery and nitrate removal

Pyrite oxidation in nature has been a subject of intense interest due to its key role in generating acid mine drainages (AMD) (Fig. 7). Decades of ore and metal mining have left pyrite-containing deposits exposed to the air and susceptible to oxidation mediated by aerobic Fe(II) and sulfur-oxidizing microorganisms, generating Fe³⁺ and sulfuric acid. Further oxidation of pyrite by Fe³⁺ generates a feedback loop that amplifies the oxidative reaction, resulting in AMD with low pH (< 5) and elevated toxic metals (e.g. Fe, Ni, Co, Cu, Zn) that negatively affect ecosystems worldwide (Baker and Banfield, 2003).

Despite its negative associations, research on acidic drainages presents some exciting opportunities to understand life in extreme environments. It is being increasingly recognized that some acidic drainages are natural and are better termed acid rock drainages (ARD) to distinguish them from anthropogenically-generated AMD. The most famous example is perhaps the Río Tinto, a 92 km long river that drains the Iberian Pyrite Belt in southwest Spain. Biogeochemical conditions in such acidic environments are similar to those that could have been found on ancient Mars as well as on early Earth (Amils and Fernández-Remolar, 2021). In particular, it was proposed that widespread ARDs occurred on the early Earth ~2.4 billion years ago, due to the GOE that accelerated terrestrial pyrite weathering (Konhauser et al., 2011). Thus, modern AMD and ARD act as analogues to provide valuable insights into ancient life and biogeochemical cycles that may have operated in the past and other habitable worlds.

Further downstream, the toxicity of acidic drainages is mitigated by several processes including hydrological dilution, pH buffering by bedrock (e.g. limestone) and sulfate reduction by SRM. Enhancing the activity of microbial sulfate reduction is of particular interest in bioremediation and resource recovery, as sulfate reduction increases pH and generates sulfide that reacts readily with metals, immobilizing them as various metal sulfides (e.g. FeS, CuS, ZnS, NiS). The precious metals can then be recovered and used for other applications (Fig. 7). Selective metal recovery has been demonstrated in some cases using precise pH control in bioreactors, although the efficiency highly depends on the specific composition of the feed solution (reviewed in Johnson and Sánchez-Andrea, 2019).

Besides aerobic pyrite oxidation, the environmental relevance of pyrite oxidation under anoxic conditions is being increasingly recognized. In fact, most of the pyrite oxidation that contributes to the aforementioned Río Tinto system is now thought to come primarily from subsurface pyrite oxidation coupled with nitrate reduction by anaerobic microorganisms (Amils et al., 2023). This process is also one of the major controls for the fate of agriculturally-sourced nitrate, a pollutant of ground- and drinking water in various parts of the world. It has been shown that pyrite of different sizes and shapes are associated with different degrees of pyrite oxidation and nitrate removal, and whether the final product is dissolved ammonium, harmless N₂ gas or the greenhouse gas N₂O (Bosch et al., 2011; Yan et al., 2019; Mansor and Xu, 2020; Pang and Wang, 2020; Jakus et al., 2021; Kappler et al., 2021). Subsurface pyrite oxidation is also an important factor to consider in engineering, as this process can lead to ground collapse and significant loss of economic and public life (Czerewko and Cripps, 2023).

While this section is focused primarily on pyrite oxidation, it is important to recognize that redox heterogeneity is expected in space and time (Peiffer et al., 2021). For example, sediment cores in Río Tinto are black within a few centimetres of depth, which can be attributed to sulfate reduction at depth (Sánchez-Andrea et al., 2012) and the formation of new metal sulfides that probably include pyrite. In the subsurface, hydrological fluctuations over time can shift redox conditions from being conducive to pyrite oxidation to being conducive to pyrite formation via the activity of sulfate/sulfur/iron-reducing microorganisms. It is likely that new pyrite is continuously being formed and redissolved in the environment. Thus, investigating biogenic pyrite formation will be key to understanding the reactivity of these dynamic phases.

Biogenic pyrite as future photovoltaics and semiconductors

As the world’s population continues to increase, there is a need for environmentally sustainable solutions to meet the future demands in energy. The optimization of energy gain from solar cells and output efficiencies of various electronics are considered high priorities in this regard. Silicon-based materials have traditionally dominated the field of photovoltaics and semiconductors. In recent years, it has been recognized that pyrite has the potential to replace silicon-based materials at a lower cost and with higher efficiency (Fig. 7). Pyrite has a suitable band gap (energy gap between two electron states) around 0.95 eV, which is neither too small (i.e. as a conductor) or too large (i.e. insulator) for it to function as a semiconductor. It further has a light absorption coefficient that is about two orders of magnitude higher than silicon. The relative abundance and non-toxicity of pyrite components (e.g. Fe, S) also make it more attractive compared to other explored materials (Wadia et al., 2009).

For pyrite to be feasible as an energy material of the future, a method to obtain high-quality pyrite in sufficient quantities needs to be developed. Thin films of pyrite have initially been synthesized through various methods such as hydro/solvothermal, hot injection and vapour deposition. Modifications of these procedures can lead to different sizes, shapes and trace metal contents, which allow for tunable band gap energy and electrical current production. These methods often require high temperatures and toxic solvents and tend to produce impurities such as the high-temperature phase marcasite (orthorhombic FeS₂) and pyrrhotite (Fe₇S₈) (Zaka et al., 2022). High-energy milling of pyrite ore can be used to obtain nanoparticles prior to thin film preparation, but there are concerns that this will exacerbate environmental problems associated with acid mine drainages. Furthermore, robust supplies of high-quality pyrite ore cannot be guaranteed as there are huge variations in the semiconductive properties of pyrite even within a deposit (Wang et al., 2021).

Could pyrite in the future be grown as a biomanufactured material? (Cosmidis, 2023). This approach takes its inspiration from nature, which has produced pyrite sustainably at a rate of ~200,000 tons per day for the last millions of years (Rickard, 2015). A mechanistic understanding of biogenic pyrite formation will be essential to allow for tunable bio-synthesis. The gap between laboratory experiments to large-scale production is vast. However, achieving this will be a worthwhile target.