Carbon has two stable isotopes: ¹²C accounting for 98.9% of the total abundance and ¹³C, making up the balance of 1.1%. The international standard for δ¹³C measurements is the PDB. Autotrophic organisms employ various pathways for inorganic C assimilation (i.e. carboxylation) with specific isotope fractionation (see Table 1), which always favor the lighter isotope. The resulting C_org is thus ¹²C-enriched relative to the inorganic C compounds from which it feeds from [83] and [85]. In contrast, heterotrophic organisms have the same isotopic composition as the organic matter they feed from. Thus, organic matter derived from heterotrophic organisms records the isotopic composition of the primary producers in a food chain.

Fractionation during C assimilation is the result of a kinetic isotope effect inherent to irreversible enzymatic CO₂-fixing reactions. In case of the quantitatively dominant carboxylation pathway by ribulose-1.5-bisphosphate (RuBP) carboxylase and the subsequent Calvin–Benson cycle, the magnitude of the isotope effect mostly ranges between 17 and 30‰ (Table 1). This large range derives from the facts that: •

the fractionation in enzymatic reactions varies widely as a function of pH, metal cofactor, temperature and number of other variables [129];

RuBisCO IB and IA give a maximum fractionation of ∼30‰ [111] and ∼24‰ [102], respectively and RuBisco II has a smaller fractionation ∼19.5‰ in photoautotrophic purple bacteria [94]. Fractionations are lower for anoxygenic photoautotrophic bacteria that employ either the 3-hydroxypropionate pathway or the reductive tricarboxylic acid cycle (Table 1; [106]). Anaerobes using the acetyl-CoA pathway can produce acetate and/or methane with a fractionation factor comprised between 15 and 58‰. Methane produced by methanogens bacteria using this pathway is probably the most fractionated with extremely ¹³C-depleted CH₄ relative to CO₂ (∼58‰, [116]). Incorporation of this methane into biomass is performed during methanotrophic assimilation with an associated fractionation of 10 to 34‰ [131]. It produces the isotopically lightest organic carbon observed in the terrestrial biosphere with δ¹³C values < −80‰ [39].

Two distinct families of C-bearing compounds are found in sedimentary rocks, namely organic carbon (C_org) and carbonate (C_carb). Carbonates are mainly composed of the skeletal remains of calcite or aragonite secreting organisms, and of inorganically precipitated carbonate crystals in the water column, on the sea floor or as intergrain cements. In any case, carbonates record the isotopic composition of Dissolved Inorganic Carbon (DIC) from which they precipitate, with an offset of the order of 1‰ due to thermodynamic isotopic equilibrium [82]. We can therefore use Archean δ¹³C_carb values as reliable indicators of the DIC isotope ratios in early Earth environments, i.e., from surface waters to sediment porosity depending on the carbonate type. Sedimentary organic matter originates from residues of living organisms, which have survived remineralization. Several parameters control its δ¹³C_org ([123] for a review), including source effects (isotopic composition of the DIC), kinetic effects associated to the C assimilation and to a minor extend, the kinetic effects associated to the recycling into biosphere and remineralization in the water column and sediments. When the isotopic composition of the C source is recorded in the δ¹³C_carb, the difference between δ¹³C_carb and δ¹³C_org (Δ¹³C_org-carb = δ¹³C_org − δ¹³C_carb) can be used as a good approximation of the kinetic effects associated to the carbon assimilation pathways and can thus be used as a proxy for life in sedimentary rocks.

The most important concern that could invalidate the use of C isotope as a proxy for life would be the existence of inorganic processes able to mimic, in direction and magnitude, metabolic-induced isotopic fractionations. One serious candidate is the Fischer-Tropsch Type reaction (FTT), which produces reduced C from oxidized C (e.g. CO₂) in hydrothermal environments. Hydrothermal alteration of ultramafic rocks (serpentinization) leads to the formation of H₂ and minerals including serpentine, magnetite and brucite. Produced H₂ can reduce CO₂ in hydrothermal fluids to form CH₄ and complex hydrocarbons, including lipids [78]. A recent study has shown that hydrocarbons formed during FTT-synthesis have low δ¹³C (ca. −36‰ relative to CO₂-source), which is in the range of most C isotopic signatures observed in Early Archean rocks including those related to hydrothermal precipitation of iron and silica [79]. Therefore, while the C isotope signal is commonly used as a biosignature in sedimentary formation, it could be ambiguous in an Early Archean hydrothermal environment.

In metamorphosed Archean sedimentary rocks, thermal devolatilization of organic matter (production of CH₄ and CO₂) and isotopic exchange between organic C and carbonates may have modified the initial δ¹³C_org. Devolatilization processes can shift C isotope composition of the residual carbonaceous matter towards higher values due to preferential loss of ¹²C. The isotopic shift is usually estimated from a relation between δ¹³C_org and H/C ratio [36] and [54]. However, in some cases, no isotopic shift is observed (see for example [101]) and corrections should be carefully employed. When both carbonate and organic C co-exist in a rock, metamorphic-driven isotopic exchange between these two phases may occur and decreases the isotope difference between C_carb and C_org with increasing temperature [119]. The 3.8 Ga sedimentary rocks of Isua Supracrustal belt (Southwest Greenland) offer a good example of extremely ambiguous δ¹³C_org data in metamorphosed rocks. These rocks experienced amphibolite facies metamorphism (500 °C; 3–5 kbar), so that their carbonaceous matter is present as graphite. In metacarbonates, bulk carbonate and graphite carbon isotope compositions range from −12 to −10‰ and −9 to +1‰, respectively. The low values and the variability of the isotope fractionations between graphite and carbonate, are more compatible with various extent of isotopic exchange during metamorphism than with values inherited from the deposition time [118]. However, graphite globules armored in garnet crystals present lower δ¹³C values of ∼ −19‰ [95]. The garnets being supposed to have sheltered to some extent the graphite from isotope exchange, their δ¹³C values are the closest known so far to the initial organic matter δ¹³C, which must have been lower and thus arguably of biological origin.

A systematic difference in the isotope composition of organic and carbonate C reservoirs (Fig. 1) – with Δ¹³C_org-carb ∼ −30‰ – has been observed in most organic-rich sediments of different ages from 3.5 to 2.0 Ga (Fig. 1). This carbon isotope record is an important, uninterrupted evidence for the presence of a substantial biosphere over geological time down to ∼3.5 Ga [54] and [99]. The δ¹³C_org record shows a negative excursion between ∼2.7 and 2.6 Ga (Fig. 1), with δ¹³C values as low as −60‰, while carbonate δ¹³C values remain close to 0‰ [35], [52] and [113]. This transitory Δ¹³C_org-carb ∼ −60‰ suggests that methanotrophy, and hence methanogenesis, was contributing significantly to the C cycling during this period.

The first step in N cycling is biological fixation of atmospheric N₂ by aerobic or anaerobic autothrophs, such as cyanobacteria, which have the nitrogenase enzyme that combines gaseous N with hydrogen to produce ammonia (NH₃) with little isotopic fractionation (ɛ ≤ 2‰; [105]). Ammonia is quickly assimilated and incorporated into proteins and other organic N compounds, through the process of uptake. Organic-N is often converted back into inorganic N by mineralization. This process produces an enzymatic breakdown of organic matter to release amine groups (-NH₂) of amino acids and subsequently NH₄ ⁺ through ammonification with little isotopic fractionation (ɛ ≤ −5 to +0‰; [105]). Under aerobic conditions, NH₄ ⁺ is rapidly oxidized to nitrite (NO₂ ⁻) and subsequently to nitrate (NO₃ ⁻) in a two-step process called nitrification, where a strong isotopic fractionation is produced (ɛ = −17‰; [105]). Marine organisms can recycle nitrates by: •

nitrate uptake or assimilation with a relatively constant ɛ of +5‰ [105];

This process, called denitrification is characterized by a much larger isotope effect (ɛ = −25‰) [105], which is responsible for the enrichment in ¹⁵N of oceanic nitrate (relative to atmospheric N).

The δ¹⁵N of nitrate resulting from this multistep kinetic metabolic fractionation is ∼ +5‰ [105]. Because nitrate is utilized by primary producers, the δ¹⁵N of the sinking flux (and of organic matter prior to diagenesis) will be close to 5‰ [88]. Diagenesis should have little isotopic effect on the N of fixed ammonium as disclosed by the isotopic similarity of oceanic N_org (δ¹⁵N = +5‰) with the bulk sedimentary N (δ¹⁵N = +7 ± 1‰) [3] and [88].

Beaumont and Robert [8] showed an evolution of the N isotopic signature of kerogens with δ¹⁵N_kerogen from −6.2‰ in the Early Archean to +10‰ in the Late Proterozoic. This evolution was interpreted as changes in the redox potential of the Earth towards more oxidizing condition around the Great Oxidation Event [55]. The increase of atmospheric O₂ at the end of the Archean would have encouraged the biological production of NO₃ ⁻ and its use as a source for organic ¹⁵N-enriched nitrogen.

Earlier negative δ¹⁵N values during Archean might reflect a metabolic isotopic fractionation in anoxic conditions, when the oceanic nitrate pool was scarce. Bacteria might have used reduced, inorganic forms of nitrogen through NH₄ ⁺ uptake [43] and NH₃ assimilation [7], which are able to produce some negative isotopic shift. Alternatively, Beaumont and Robert [8] argued that atmospheric N₂ could have been isotopically lighter than today (δ¹⁵N = −5‰) if derived from an Enstatite Chondrite source [60] and thus, simple biological fixation of that N could have been reflected in the organic pool.

Pinti and Hashizume [90] and Pinti et al. [91] challenged the interpretation of Beaumont and Robert [8], suggesting that their kerogens represented an environmentally biased sample of the Archean biota, in that many were extracted from silica precipitated from hydrothermal fluids. In such anoxic setting, metabolic processes are regulated by chemosynthetic biota [26]. Chemoautolithotrophy describes the synthesis of organic C compounds from CO₂ using energy and reducing power derived from the oxidation of inorganic compounds, such as NH₄ ⁺, which support microbial chemosynthesis [59]. The N isotopic signature of the biomass produced by chemosynthesis, via N fixation [70] and [80] is significantly depleted in ¹⁵N, with δ¹⁵N values ranging from −9.6 to +0.9‰ [25], [26], [70], [86] and [117].

N isotopic ratios measured in Late Archean (2.7 Ga) shales and kerogens by Jia and Kerrich [62] and [63] and Kerrich et al. [71] showed exclusively positive δ¹⁵N values from +5 to +24‰. They interpreted the high δ¹⁵N values as a fossil residual signature of a veneer atmosphere having an initial CI-chondritic composition with δ¹⁵N values from +30 to +42‰. The isotopic shift from +24 at 2.7 Ga to 0‰ at present was interpreted as the combination of three processes: •

degassing of ¹⁵N-depleted mantle N (δ¹⁵N = −5 ± 2‰);

However, the occurrence of an isotopically heavier atmosphere during Archean can be dismissed by the presence of very variable negative δ¹⁵N values in Archean diamonds from −5 to −10‰ and down to −25‰ [22] which possibly reflect the mixing of a crustal component and an atmospheric component recycled into the mantle with isotopically lighter (rather than heavier) primeval N inherited from prototerrestrial material [114].

In Fig. 1, we reported all published (and unpublished data from D. Pinti internal database) N isotopic data measured in bulk rocks and kerogens, plotted against time since Earth formation. Paleoarchean (3.8–3.2 Ga) N isotopic signatures exhibit a bimodal distribution, with δ¹⁵N values from −7 to +7‰ [92]. Nitrogen from Neoarchean organic matter (3.2–2.5 Ga) is much more ¹⁵N-enriched, with a δ¹⁵N_mean value around +11‰. In a restricted range of age clustered around 2.7 Ga, kerogens, cherts and particularly BIF show an extreme enrichment in ¹⁵N (from +24‰ up to +35‰; [8] and [62]). Proterozoic kerogens and cherts show lower δ¹⁵N, with an average of +5.6‰, close to the value of modern ocean nitrates. It is worth noting that the large positive shift in the N isotopic composition measured in kerogens and rocks is reached before the Great Oxidation Event (2.35 Ga; [9]), during the largest deposition of Banded Iron Formation (BIF) ever on Earth, between 2.7 and 2.6 Ga [58]. The deposition of BIF requires the presence of little oxygen, possibly 10^-2 Present Atmospheric Level (PAL) or less [57]. If these conditions were sufficient to have limited the amount of nitrates in the ocean, we could speculate that the observed elevated δ¹⁵N values were obtained by enhanced denitrification [2] and [3]. The degree of N kinetic isotope fractionation depends on the nitrate availability as a reactant: the lower the reactant availability, the higher the isotopic effect that can be produced by Rayleigh distillation [105]. In a nitrate-poor Archean ocean, the nitrification-denitrification-assimilation processes could have easily produced a higher nitrate δ¹⁵N than in the modern ocean. With the increase in oceanic nitrate content following the oxygenation of the ocean at the end of Archean, the nitrate isotopic composition would have decreased to the modern average value of +5‰.

Post-depositional changes of the pristine, metabolic-induced N isotopic fractionation are difficult to evaluate. Studies on thermal metamorphism showed that the δ 15 N N H 4 + $ {\mathrm{\delta }}^{15}{\text{N}}_{\text{N}{\text{H}}_{4}+}$ measured in mica increased progressively with the metamorphic grade up to +15‰. This isotope shift was interpreted as a progressive devolatilization of the rock (with preferential loss of lighter ¹⁴N compared to ¹⁵N) by increasing temperature [49]. Dauphas and Marty [30] showed that devolatilization could have caused large isotopic shift (≥ 15‰) in some Archean hydrothermal mica. However, recent studies have shown that during regional metamorphism (high pressure, high temperature) the δ 15 N N H 4 + $ {\mathrm{\delta }}^{15}{\text{N}}_{\text{N}{\text{H}}_{4}+}$ shift is limited to +2 − 3‰ (e.g., [14] and [61]), while the δ¹⁵N_kerogen shift is null [1]. Moreover, as for carbon, the most important concern that could invalidate the use of N isotopes as biosignature would be abiotic processes mimicking metabolic-induced isotopic fractionations. The few experimental data existing on N isotopic fractionation during FTT synthesis of N-based polymers and amino acids gave contradictory results. Miller–Urey reactions of CH₄–NH₃–H₂ mixtures have produced N-containing nonvolatile soluble organics (amino acids, organic acids) and polymers with δ¹⁵N values 8 to 11‰ greater than the starting NH₃ [75]. Plasma-discharge of CO–N₂–H₂ produced carbonaceous materials with δ¹⁵N values of −3 to −17‰ relative to the reactant N₂ [72]. These few experimental evidences do not allow estimating whether FTT reactions could have an important role in N isotopic fractionation in Precambrian rocks.

The Earth surface cycle of S is dominated by biological processes, which utilize the eight-electron S redox gradient and impart large S isotopic fractionations.

Organisms generally use S according to three main processes: •

bacterial sulfate reduction (BSR), where sulfate acts as an electron acceptor during organic C oxidation (i.e. respiration) or H₂ oxidation;

We describe below the S isotope fractionations associated with these three processes. The fractionation (ɛ³⁴) imposed by BSR between initial sulfate and produced sulfides varies from 4 to 46‰ [23], [50] and [68]. It depends on: •

the organism involved [34];

At high sulfate concentrations, the biological sulfur isotope fractionation of natural populations of sulfate reducers is typically comprised between 20 and 40‰, independently of temperature and reduction rate [16] and [46]. At low sulfate concentrations (less than 200 μM), it is greatly reduced [16] and [50].

The biological pathways of sulfides oxidation in nature are diverse and poorly known. They include phototrophic and non-phototrophic oxidation of various sulfide species such as H₂S, S⁰, S₂O₃ ²⁻ and SO₃ ²⁻. The fractionations produced during phototrophic oxidation are small or negligible (between −2 and 0‰ [41] and [42]). Small fractionation also accompanies the non-phototrophic oxidation between −1 and 1‰ [68].

Three different disproportionation reactions involving intermediate S compounds are known: S⁰ (4S⁰ + 4H₂O → 3H₂S⁺ + SO₄ ²⁻ + 2H⁺), S₂O₃ ²⁻ (S₂O₃ ²⁻ + H₂O → H₂S + SO₄ ^2-), and SO₃ ^2- (4SO₃ ²⁻ + 2H⁺ → H₂S + 3SO₄ ²⁻). The fractionation associated with the disproportionation of elemental S shows a narrow range, where sulfide is depleted in ³⁴S by 6.1 ± 0.4‰, and sulfate is enriched in ³⁴S by 18.3 ± 1.3‰ [21]. ³⁴S depletions higher than 70‰ recorded in natural sulfides derive from multistep sulfide oxidation of sulfur intermediate subsequently disproportionated, and so on [20]. Inorganic disproportionation involves only a minor < 3‰ kinetic isotope fractionation [108].

High temperature hydrothermal processes (> 100 °C) promote the abiological thermo sulphate reduction (TSR) to sulfides [44]. The maximum isotopic fractionation associated to TSR is ∼20‰ at 100 °C. Therefore, while the S isotope signal is commonly used as a biosignature of BSR and disproportionation in sedimentary formations, it could be ambiguous in Early Archean hydrothermal formations [77]. Sulfides could also derive from H₂S released from NSO-compounds, with kinetic sulfur isotope fractionation lower than −2‰ [77]. Hence, sulfides produced by organic matter thermal cracking should present much heavier δ³⁴S values than those formed by BSR.

Sulfur isotope measurements are performed essentially on three families of sulfur-bearing compounds which can be found in the rock record. Evaporite sulfate minerals (gypsum, anhydrite, barite), carbonate-associated sulfates (CAS) and sulfur minerals (essentially pyrite). Both evaporites and CAS record faithfully the isotopic composition of the dissolved sulfate they originate from, and pyrite precipitates from dissolved sulfide with a small isotope fractionation (∼1‰, [93]). Thus, the difference between sedimentary sulfate (evaporite or CAS) and pyrite isotope compositions can be used as an indicator of BSR and S disproportionation pathways.

Between 3.5 and 2.4 Ga, sulfate δ³⁴S secular evolution seems to be centered on 0‰. It is however poorly constrained so far because of the lack of evaporite sulfates minerals in the rock record, and the scarcity of available δ³⁴S data on CAS (Fig. 1). Pyrite δ³⁴S Archean secular evolution is better constrained and is also centered around 0‰, except for the Dresser Formation (western Australia) at ∼3.49 Ga, where δ³⁴S_pyrite are lower than −20‰ (Fig. 1) and interpreted as a biological signature in spite of the fact that they belong to a hydrothermal system [89] and [104] (see § 5.4). The range of δ³⁴S_pyrite around 0 ± 5‰ observed in most Archean pyrites can thus be interpreted either as an absence of evidence for the operation of BSR/disproportionation [109], or, assuming that these sulfur metabolisms were operating [46] and [104], as the evidence for a sulfate-poor ocean [17]. At ∼2.4 Ga, the increasing range of δ³⁴S_pyrite provides indisputable evidences both for the involvement of BSR and/or disproportionation in the sulfur cycling and for an increasing ocean sulfate concentration. Assuming that the sulfate concentration in the oceans depends on the rate of pyrite oxidative weathering [18] and [110], this increase in sulfate concentration is interpreted as a response to a rise in the atmospheric O₂ content known as the Great Oxidation Event (see [55] for a review).

More recently, combined isotopic measurements of δ³³S, δ³⁴S and δ³⁶S from Archean sulfides and sulfates have provided new insights into the Archean sulfur metabolisms and oxygenation history. The multiple S isotope studies revealed S anomalous mass-independent fractionation (MIF-S) [9], [37], [81] and [87]. MIF-S is expressed by the Δ³³S (i.e. Δ³³S = δ³³S-1000*[1 + δ³⁴S/1000]^0.515 − 1) and refers to any chemical or physical process that acts to separate isotopes, where the amount of separation does not scale in proportion with the difference in the masses of the isotopes (see review in [112]). Most isotopic fractionations (including typical kinetic fractionations and equilibrium fractionations) are mass dependent since they are caused by the effects of the mass of an isotope on atomic or molecular velocities, diffusivities or bond strengths. MIF processes are less common, and for sulfur isotopes are known to occur during UV photolysis of SO₂ and SO. In its most basic form, photolysis produces two isotopically-distinct types of sulfur bearing aerosols that are likely to rain out to Earth's surface environments and transfer positive Δ³³S via reduced sulfur species (S⁰) and negative Δ³³S via oxidized sulfur species (H₂SO₄) to the sedimentary record. MIF-S are thought to be produced and preserved in the sediments only when atmospheric partial pressure of O₂ is lower than 10^-5 PAL, thus preventing both oxygen UV shielding effect and re-homogenization of atmospheric sulfur species through oxidation [38], [69] and [98]. MIF-S are significant (i.e. different from 0‰) from 3.8 to 2.4 Ga (Fig. 1) and therefore, provide evidence for a low atmospheric O₂ content before 2.4 Ga [9], and for a very different atmospheric S chemistry in the Archean. Moreover, the sign of Δ³³S can be used to trace the oxidation state of the source of the sedimentary S component. This approach was recently used to re-evaluate the significance of the strongly negative δ³⁴S_pyrite observed in the 3.49 Ga Dresser Formation (Pilbara, western Australia), which was suggested to reflect the first evidence of BSR in a locally sulphate-rich Early Archean environment [104]. Combined observation of negative δ³⁴S and positive Δ³³S of sulfides was interpreted as evidence for reduced S conversion to sulfides by S disproportionating microorganisms [89].

Iron has four stable isotopes, ⁵⁴Fe, ⁵⁶Fe, ⁵⁷Fe and ⁵⁸Fe, with natural abundances of 5.84, 91.75, 2.12 and 0.28%, respectively. The isotope ratio ⁵⁶Fe/⁵⁴Fe used by most authors is expressed as δ⁵⁶Fe, referred to the IRMM-014 standard. The most important parameters controlling Fe isotope fractionations in natural environment are oxidation state and bonding. In aqueous system, ferric iron (Fe[III]) is soluble under acidic and oxidizing conditions but is insoluble under near neutral-pH conditions. Ferrous iron (Fe[II]) is soluble over a large range of pH under anoxic conditions. Cycling of Fe is not only controlled by Eh-pH conditions but also by living organisms that derive energy through changes in Fe redox state [76] and [84]. Organisms generally use Fe according to three main processes: •

Fe(II) oxidation, where Fe(II) acts as an electron donor for energy generation;

We describe below Fe isotope fractionation associated with these three processes.

The oxidation of Fe(II) under anoxic condition can be performed by phototrophic or chemotrophic metabolisms. Fe(II) can also be oxidized indirectly by microorganisms when reacting with O₂ released from oxygenic photosynthesis. The oxidation of Fe(II)_aq into Fe(III)_aq leads to insoluble species such as Fe oxihydroxide that evolve ultimately into goethite, hematite and/or magnetite. Isotopic studies of Fe oxides and hydroxides stored in sedimentary rocks may thus reveal the signature of life. Unfortunately, experimental works showed that the oxidation of Fe(II)_aq to Fe(III)_aq enriches Fe(III)_aq in the heavy isotopes by ∼1 to 3.4‰, regardless on whether this oxidation is biologically mediated or not [4], [5], [64] and [122]. The enrichment in heavy Fe isotopes during oxidation is usually counterbalanced by a kinetic isotope fractionation during precipitation, enriching the precipitate in the light isotopes [107]. Overall, the net isotopic fractionation is modulated by the rate of precipitation but, in most cases, it produces a precipitate enriched in the heavy isotopes of Fe.

Bacterial DIR is a widespread process in anaerobic marine sediments and may be one of the oldest forms of respiration [120]. This process couples the oxidation of organic matter to the reduction of solid Fe(III). In addition to the production of Fe(II)_aq, the end-products of DIR may include Fe carbonates (siderite and ankerite) and magnetite [76]. Microbial DIR produces Fe isotope fractionation enriching Fe(II)_aq in the light isotopes relative to the initial Fe(III) substrate [6], [12], [28] and [65]. In experiments of bacterial DIR, the Fe isotope fractionation between Fe(II)_aq and a reactive Fe(III) component on a Fe(III) oxide surface was found constant at −2.95‰ over long timescales, for various Fe(III) substrates and bacterial species [28]. Bacterially-mediated Fe dissolution is thus associated with large isotope fractionation. Abiotic dissolution of minerals seems to be more complex. Some experiments as well as natural environment study illustrate that abiotic dissolution of minerals did not produce significant Fe isotope fractionation [11], [12], [15] and [107]. Abiotic goethite dissolution shows different results depending on the mechanism of dissolution. Proton-promoted dissolution produces no fractionation, while ligand-controlled and reductive dissolution enrich the produced Fe(II)_aq in the light isotopes by 1.7‰ relative to reactive surface [126]. These abiotic fractionations were significant for the first steps of substrate dissolution but were negligible after only ∼2% of the initial substrate dissolved. Accordingly, abiotic mineral dissolution produces either no or limited fractionation of Fe isotopes. While the direction of the fractionation is the same for biotic and abiotic dissolution, the amplitude is much larger (by at least ∼1‰) in the case of microbial dissolution. When microorganisms have intense metabolic activity, they can produce a large amount of Fe(II)_aq with very light Fe isotope signature under conditions out of equilibrium.

Iron assimilation in microorganisms is largely due to the two stable valencies that impart considerable range to the reactivity of Fe. When inserted into organic ligand, Fe can be used to catalyze a wide range of chemical reactions essential to the cell. Fe is also the active component of the O₂ carrier proteins. Only few studies have examined Fe isotope fractionation during uptake by microorganisms [11], [12] and [121]. Experimental work indicates that bacterial Fe assimilation favors uptake of heavy isotope into the cell, with a fractionation of ∼1.1‰ [121]. This fractionation cannot be explained by a simple kinetic process (since lighter isotopes should be favored) and requires at least one step in isotopic equilibrium. However, whether it reflects active microbial assimilation or passive adsorption of Fe on cell walls is still unknown.

The earliest photosynthetic metabolism on Earth was probably anoxigenic. Because Fe(II) was the most important electron donor available in the Archean ocean, anaerobic Fe(II) oxidation may have been the main metabolic pathway [19] and [125]. This biogenic oxidation of Fe(II)_aq into Fe(III)_aq produces insoluble Fe oxihydroxide and is often considered to explain the deposition of BIF in the Archean [51] and [73]. Other hypotheses to explain the formation of BIF include Fe(II) oxidation through increase in ambient O₂ by photosynthetic activity [24] and abiotic Fe photo-oxidation by UV photons [13]. As noted above, Fe isotopes alone cannot differentiate biotic from abiotic oxidation. However, a recent study coupling laboratory experiments with thermodynamic modeling seems to rule out the possibility of a photo-oxidation [74]. If this is true, the two other scenarios involve living organisms, either as oxygenic or anoxygenic photosynthetic metabolism.

Before deciphering biogeochemical index of life from Fe isotopes, it is crucial to understand the cycling of Fe in the Archean. Large variations in the Fe isotope record were identified during Archean and Early Proterozoic by analyzing sedimentary pyrites [96]. Fig. 1 plots a compilation of Fe isotope data available in the literature for Archean rocks. The most significant variation is a large negative excursion of δ⁵⁶Fe values between 2.7 and 2.3 Ga flanked by near zero or slightly positive δ⁵⁶Fe values before 2.7 Ga and after 2.3 Ga. It was primarily interpreted as reflecting changes in the Fe isotopes signature of seawater [29] and [96]. However, this interpretation is weakened by the observation of large Fe isotope heterogeneity (> 2‰) at the cm-scale in Archean rocks, which cannot be explained by secular variation in seawater, because of the high Fe residence time in the Archean ocean [65]. Although such a small scale variability has only been reported once and still needs some confirmation, an alternate hypothesis has been proposed which is that the δ⁵⁶Fe of sedimentary rocks might reflect mainly biogeochemical cycling during sediment diagenesis [66], [67] and [130]. The large negative excursion of δ⁵⁶Fe values between 2.7 and 2.3 Ga would result from an extensive activity of DIR metabolism in response to increase Fe(III) and organic carbon delivery to the ocean [65]. If true, this exciting alternative would indicate that Fe isotopes do record the signature of iron based metabolic pathways.

The drawback with studying old Precambrian rocks is that most of them experienced metamorphism. An important question is thus whether Fe isotope composition reflects processes in water column, sediment diagenesis or metamorphism. Several works on Eoarchean metamorphic rocks demonstrated that Fe isotopes can be used to distinguish protolith signature from metamorphic overprint [31], [32] and [33]. Whitehouse and Fedo [124] observed small-scale isotopic heterogeneity in magnetite from very small areas within single bands of Eoarchean BIF, indicating a range of fractionation within single depositional event. This is unlikely to have resulted from water column chemistry or later metamorphic disturbance, which should produce isotopic homogeneization. In contrast, Fe isotope heterogeneity likely derives from isotopic reservoir effect during diagenetic reactions in pore waters isolated from the ocean [124]. In Biwabik Iron Formation, Fe isotope fractionation between coexisting minerals decreases with increasing metamorphic grade, indicating that isotopic exchange occurred during metamorphism [115]. Isotopic exchange was limited to the mineral-scale but bulk rock behaved as closed-system and preserved primary Fe isotope composition [40] and [115]. Any potential isotopic re-equilibration could be avoided by analyzing metamorphic rocks or layers with single Fe-bearing mineralogy, which may be particularly useful for searching isotopic traces of early life.