The evolving redox state of the solid Earth in deep time

The ƒO₂ of Earth’s present-day upper mantle is QFM ± 2 (Fig. 1, see Box 1), but it decreases with depth (e.g. Haggerty, Reference Haggerty1978; Christie et al., Reference Christie, Carmichael and Langmuir1986; O’Neill and Wall, Reference St.C and Wall1987; Wood and Virgo, Reference Wood and Virgo1989; Wood et al., Reference Wood, Bryndzia and Johnson1990; Ballhaus et al., Reference Ballhaus, Berry and Green1991; O’Neill, Reference St.C1991; Holloway et al., Reference Holloway, Pan and Gudmundsson1992; Kasting, Reference Kasting1993; McCammon, Reference McCammon2005; Frost and McCammon, Reference Frost and McCammon2008; Cottrell and Kelley, Reference Cottrell and Kelley2011; Trail et al., Reference Trail, Watson and Tailby2011; Ardia et al., Reference Ardia, Hirschmann, Withers and Stanley2013; Gaillard et al., Reference Gaillard, Bouhifd, Füri, Malavergne, Marrocchi, Noack, Ortenzi, Roskosz and Vulpius2021; Yang et al., Reference Yang, Liu and Zhang2022) (Fig. 2). The transition zone is assumed to have an ƒO₂ of about QFM – 4 (McCammon, Reference McCammon2005; Frost and McCammon, Reference Frost and McCammon2008; Ardia et al., Reference Ardia, Hirschmann, Withers and Stanley2013; Yang et al., Reference Yang, Liu and Zhang2022) and the lower mantle is supposed to have an ƒO₂ below QFM – 5 (Frost et al., Reference Frost, Liebske, Langenhorst, McCammon, Trønnes and Rubie2004; McCammon, Reference McCammon2005; Ardia et al., Reference Ardia, Hirschmann, Withers and Stanley2013; Yang et al., Reference Yang, Liu and Zhang2022).

Figure 1. Commonly used mineral redox buffers and their relationship to ƒO₂ plotted over temperature. Quartz-fayalite-magnetite QFM) and nickel-nickel-oxide (NiNiO) depict oxidised conditions, while iron-wustite (IW) and quartz-iron-fayalite (QIF) represent reduced conditions.

Figure 2. Evolution of the Earth´s redox state for different formation stages. The changing ƒO₂ is indicated by the deviation in log units from the quartz-fayalite-magnetite (QFM) buffer and is explained in the text. The colours range from blue (reduced) to red (oxidised). The Earth is assumed to become more oxidised with time, with the most reduced values during the accretion period before core formation. It is thought that during the magma ocean period, ƒO₂ evolved towards more oxidised values. The modern Earth is comparatively oxidised, with a decreasing redox state with depth (after McCammon, Reference McCammon2005). See the text for references on the redox state of the early Earth.

In contrast to the present day, it is commonly thought that the Earth was initially reduced, with an ƒO₂ of IW – 2 (ca. QFM – 5.7) as a maximum value during core formation (before that, the ƒO₂ could have been even as low as IW – 5 or about QFM – 8.7; Wade and Wood, Reference Wade and Wood2001, Reference Wade and Wood2005; Rubie et al., Reference Rubie, Frost, Mann, Asahara, Nimmo, Tsuno, Kegler, Holzheid and Palme2011, Reference Rubie, Jacobson, Morbidelli, O’Brien, Young, De Vries, Nimmo, Palme and Frost2015; Scaillet and Gaillard, Reference Scaillet and Gaillard2011; Cartier et al., Reference Cartier, Hammouda, Boyet, Bouhifd and Devidal2014; Fischer et al., Reference Fischer, Nakajima, Campbell, Frost, Harries, Langenhorst, Miyajima, Pollok and Rubie2015; Schaefer and Elkins-Tanton, Reference Schaefer and Elkins-Tanton2018; Gaillard et al., Reference Gaillard, Bouhifd, Füri, Malavergne, Marrocchi, Noack, Ortenzi, Roskosz and Vulpius2021) (Fig. 2). One argument for a reducing start of the Earth is the assumption that it initially accreted from highly reduced, volatile-depleted material (like enstatite chondrites). Another argument is that the metal–silicate equilibrium required for core formation suggests a considerably low ƒO₂. A low ƒO₂ enhances the siderophile behaviour of certain elements like nickel, cobalt, manganese, chromium, vanadium and silicon (Gessmann et al., Reference Gessmann, Rubie and McCammon1999; Wade and Wood, Reference Wade and Wood2001, Reference Wade and Wood2005; Rubie et al., Reference Rubie, Frost, Mann, Asahara, Nimmo, Tsuno, Kegler, Holzheid and Palme2011, Reference Rubie, Jacobson, Morbidelli, O’Brien, Young, De Vries, Nimmo, Palme and Frost2015; Scaillet and Gaillard, Reference Scaillet and Gaillard2011; Siebert et al., Reference Siebert, Badro, Antonangeli and Ryerson2013; Cartier et al., Reference Cartier, Hammouda, Boyet, Bouhifd and Devidal2014; Fischer et al., Reference Fischer, Nakajima, Campbell, Frost, Harries, Langenhorst, Miyajima, Pollok and Rubie2015; Schaefer and Elkins-Tanton, Reference Schaefer and Elkins-Tanton2018; Gaillard et al., Reference Gaillard, Bouhifd, Füri, Malavergne, Marrocchi, Noack, Ortenzi, Roskosz and Vulpius2021; but see Badro et al. (Reference Badro, Brodholt, Piet, Siebert and Ryerson2015) for an alternative viewpoint).

Assuming reduced conditions early in Earth’s history, its relatively oxidised state today requires oxidation over time (e.g. Wade and Wood, Reference Wade and Wood2005; Cartier et al., Reference Cartier, Hammouda, Boyet, Bouhifd and Devidal2014; Schaefer and Elkins-Tanton, Reference Schaefer and Elkins-Tanton2018; Gaillard et al., Reference Gaillard, Bouhifd, Füri, Malavergne, Marrocchi, Noack, Ortenzi, Roskosz and Vulpius2021). Pahlevan et al. (Reference Pahlevan, Schaefer and Hirschmann2019) found that if the magma ocean was initially reduced, it must have evolved to a higher ƒO₂ (>IW + 1 or ~QFM – 2.7) during its final stages based on the D/H ratio. Moreover, Deng et al. (Reference Deng, Du, Karki, Ghosh and Lee2020) suggested that the magma ocean had a vertical gradient in ƒO₂, with the upper layer reaching IW + 2 (~QFM – 1.7). However, it was also proposed that more oxidised, volatile-rich material (e.g. CI chondrites) was delivered during the last stages of accretion and core formation (e.g. Wänke et al., Reference Wänke, Dreibus, Jagoutz, Kröner, Hanson and Goodwin1984; Javoy, Reference Javoy1995; Wade and Wood, Reference Wade and Wood2005; Schönbächler et al., Reference Schönbächler, Carlson, Horan, Mock and Hauri2010; Rubie et al., Reference Rubie, Frost, Mann, Asahara, Nimmo, Tsuno, Kegler, Holzheid and Palme2011; Scaillet and Gaillard, Reference Scaillet and Gaillard2011; Marty, Reference Marty2012; Cartier et al., Reference Cartier, Hammouda, Boyet, Bouhifd and Devidal2014; Fischer et al., Reference Fischer, Nakajima, Campbell, Frost, Harries, Langenhorst, Miyajima, Pollok and Rubie2015; Dauphas, Reference Dauphas2017; Fischer-Gödde and Kleine, Reference Fischer-Gödde and Kleine2017; Lammer et al., Reference Lammer2018; Grewal et al., Reference Grewal, Dasgupta, Sun, Tsuno and Costin2019; Budde et al., Reference Budde, Burkhardt and Kleine2019; Fischer-Gödde et al., Reference Fischer-Gödde, Elfers, Münker, Szilas, Maier, Messling, Morishita, Van Kranendonk and Smithies2020; Gaillard et al., Reference Gaillard, Bouhifd, Füri, Malavergne, Marrocchi, Noack, Ortenzi, Roskosz and Vulpius2021). Rubie et al. (Reference Rubie, Frost, Mann, Asahara, Nimmo, Tsuno, Kegler, Holzheid and Palme2011) concluded that 30–40% of the final mass accreted was rather oxidised, therefore the evolution towards a more oxidised planet probably occurred already during the formation of the Earth.

Besides the variation in the delivered material, the change in the Earth´s redox state during accretion can also be explained by the increasing size of the Earth (Wade and Wood, Reference Wade and Wood2005). It was proposed that due to the higher pressures associated with the growth of the Earth, perovskite (Mg,Fe,Al)(Al,Si)O₃) becomes the dominant phase of the lower mantle (stable below 660 km depth in present-day Earth). Perovskite formation drives Fe(II) disproportionation to Fe(III) and Fe(0) via Eq. (1):

(1) $$ 3{\mathrm{Fe}}^{2+}\mathrm{O}+{\mathrm{Al}}_2{\mathrm{O}}_3=2{\mathrm{Fe}}^{3+}{\mathrm{Al}\mathrm{O}}_3+{\mathrm{Fe}}^0 $$

Because Fe(0) has been sequestered into the core, the lower mantle became relatively enriched in Fe(III). It was suggested that the upper mantle became enriched over time due to convection (Mao and Bell, Reference Mao, Bell, Saxena, Bhattacharji, Annersten and Stephansson1977; Frost et al., Reference Frost, Liebske, Langenhorst, McCammon, Trønnes and Rubie2004, Reference Frost, Mann, Asahara and Rubie2008; Wade and Wood, Reference Wade and Wood2005). This transfer of Fe(III) from the lower to the upper mantle is also known as the ‘oxygen pump’ (Frost et al., Reference Frost, Liebske, Langenhorst, McCammon, Trønnes and Rubie2004, Reference Frost, Mann, Asahara and Rubie2008; Wade and Wood, Reference Wade and Wood2005). This process would also explain why the Martian mantle is more reduced than Earth (Righter and Drake, Reference Righter and Drake1996; Herd et al., Reference Herd, Papike and Brearley2001, Reference Herd, Borg, Jones and Papike2002; Wadhwa, Reference Wadhwa2001, Reference Wadhwa2008; Wade and Wood, Reference Wade and Wood2005; Righter et al., Reference Righter, Yang, Costin and Downs2008), even though Mars is enriched in volatiles and FeO. Perovskite is unstable in the Martian mantle due to its smaller size, resulting in lower lithostatic pressures, thus the self-oxidation of the mantle via perovskite formation could not occur on Mars (Wade and Wood, Reference Wade and Wood2005). It was proposed that the preferential partitioning of Fe(III) into the liquid phase enhances the equilibration of the redox state between the lower and upper mantle by mixing processes (Carmichael, Reference Carmichael1991; Scaillet and Gaillard, Reference Scaillet and Gaillard2011). In addition, FeO and FeO_1.5 have different molar volumes and densities, further favouring a more oxidised upper mantle and a more reduced lower mantle (Deng et al., Reference Deng, Du, Karki, Ghosh and Lee2020).

Alternatively, it was suggested that the oxidation of the mantle occurred directly during the magma ocean state (Schaefer and Elkins-Tanton, Reference Schaefer and Elkins-Tanton2018; Pahlevan et al., Reference Pahlevan, Schaefer and Hirschmann2019). Accordingly, the sink and sequestration of iron metal into the core would leave behind an oxidised mantle without requiring crystallisation and whole-scale mantle mixing (Schaefer and Elkins-Tanton, Reference Schaefer and Elkins-Tanton2018; Pahlevan et al., Reference Pahlevan, Schaefer and Hirschmann2019). In particular, the crystallising magma ocean would become progressively oxidised over time (Scaillet and Gaillard, Reference Scaillet and Gaillard2011). Similar arguments have been brought forward for a carbon pump leading to the formation of diamonds in the lower mantle (causing oxidation) in case of a deep (potentially giant-impact-induced) magma ocean, which may explain the thick CO₂ atmosphere of Venus in the absence of a late giant impact (Armstrong et al., Reference Armstrong, Frost, McCammon, Rubie and Boffa Ballaran2019). Moreover, H₂ loss from the mantle by outgassing is also discussed as a mechanism for oxidising the upper mantle (Sharp et al., Reference Sharp, McCubbin and Shearer2013).

It has also been suggested that recycling of surface material could have oxidised the upper part of the mantle (Arculus, Reference Arculus1985; Kasting, Reference Kasting1993; Kasting et al., Reference Kasting1993; Kump et al., Reference Kump, Kasting and Barley2001; Smart et al., Reference Smart, Tappe, Stern, Webb and Ashwal2016; Nicklas et al., Reference Nicklas, Puchtel, Ash, Piccoli, Hanski, Nisbet, Waterton, Pearson and Anbar2019; Stagno and Aulbach, Reference Stagno, Aulbach, Moretti and Neuville2021). At least today, the material transported with the subducting slab is more oxidised than the surrounding mantle (e.g. Wood et al., Reference Wood, Bryndzia and Johnson1990; Ballhaus et al., Reference Ballhaus, Berry and Green1991; Blundy et al., Reference Blundy, Brodholt and Wood1991). Mikhail and Sverjensky (Reference Mikhail and Sverjensky2014) found that under oxidising conditions, N₂ is the dominant nitrogen species over NH₄⁺. They argue that, during subduction, the increased ƒO₂ of the mantle wedges, compared to the surrounding upper mantle, results in N₂-rich fluids. The ascent and outgassing of such fluids allow an enhanced N₂ outgassing. Plate tectonics would, therefore, not only favour oxidised mantles and atmospheres but would also be needed for nitrogen-rich atmospheres like the Earth’s.

Duncan and Dasgupta (Reference Duncan and Dasgupta2017) turned the argument around: if reduced material (like organic carbon) was subducted, then this may have led to a transient increase of biological O₂ in the atmosphere by removing reducing power from the surface reservoir. At the same time, it would lead to a reducing effect on the mantle (unless permanently sequestered into a hidden reservoir) and result in releasing reducing gases into the atmosphere on melting.

The estimates on when the upper mantle was oxidised to near-modern values range from 4.4 to 2.7 Ga (e.g. Canil, Reference Canil1997; Delano, Reference Delano2001; Lee et al., Reference Lee, Brandon and Norman2003; Li and Lee, Reference Li and Lee2004; Foley, Reference Foley2011; Scaillet and Gaillard, Reference Scaillet and Gaillard2011; Trail et al., Reference Trail, Watson and Tailby2011; Aulbach and Stagno, Reference Aulbach and Stagno2016; Rollinson et al., Reference Rollinson, Adetunji, Lenaz and Szilas2017; Nicklas et al., Reference Nicklas, Puchtel and Ash2018, Reference Nicklas, Puchtel, Ash, Piccoli, Hanski, Nisbet, Waterton, Pearson and Anbar2019) (Figs 3, 4a). An important archive for understanding the redox state of the early Earth’s mantle is the cerium concentration in zircons (ZrSiO₄) (Loucks et al., Reference Loucks, Fiorentini and Henríquez2020). Cerium exists in both tri- and quadrivalent states in silicate melts. Zircons crystallising from these melts preferentially incorporate Ce⁴⁺ over Ce³⁺, substituting for Zr⁴⁺ in the zircon crystal structure. The ratio of Ce⁴⁺ to Ce³⁺ is influenced by the ƒO₂ of the melt. As a result, the cerium concentration in magmatic zircons can indicate the oxygen content in the magma (Trail et al., Reference Trail, Watson and Tailby2011). Trail et al. (Reference Trail, Watson and Tailby2011) calibrated the relationship between the zircon/melt partitioning coefficient of cerium and the ƒO₂ of the melt. Using their different oxygen isotopic compositions, Trail et al. (Reference Trail, Watson and Tailby2011) distinguished zircons derived from the mantle (δ¹⁸O = +5.3‰) and those from the crust. The cerium concentration data of both populations indicated that the host magmas had similar ƒO₂ to the modern mantle, which has QFM ± 2 (Yang et al., Reference Yang, Liu and Zhang2022). It was observed that the primary mantle melts were not saturated in zircon, but rather, the ‘mantle’ zircons crystallised in melt residues. The δ¹⁸O values of these residues would still closely resemble the composition of the host mantle. At such ƒO₂ values, the outgassing of CO₂, N₂, H₂O, and SO₂ would dominate over more reduced species like CO, H₂, NH₃ and H₂O (Frost and McCammon, Reference Frost and McCammon2008).

Figure 3. Literature estimates regarding the ƒO₂ of the upper mantle on early Earth (after Aulbach and Stagno, Reference Aulbach and Stagno2016; Schaefer and Elkins-Tanton, Reference Schaefer and Elkins-Tanton2018; Stagno and Aulbach, Reference Stagno, Aulbach, Moretti and Neuville2021). The shaded rectangles represent the results of the ƒO₂ estimated from individual studies (normalised to the QFM buffer) over the respective ages of the samples examined. The references for the horizontal rectangles are Canil (Reference Canil1997), Delano (Reference Delano2001), Li and Lee (Reference Li and Lee2004), Nicklas et al. (Reference Nicklas, Puchtel and Ash2018, Reference Nicklas, Puchtel, Ash, Piccoli, Hanski, Nisbet, Waterton, Pearson and Anbar2019), Aulbach and Stagno (Reference Aulbach and Stagno2016) (AS), Aulbach et al. (Reference Aulbach, Woodland, Vasilyev, Galvez and Viljoen2017). The crosses are data points from Nicklas et al. (Reference Nicklas, Puchtel and Ash2018) and Nicklas et al. (Reference Nicklas, Puchtel, Ash, Piccoli, Hanski, Nisbet, Waterton, Pearson and Anbar2019) that represent an estimated ƒO₂ based on the redox-dependent partitioning of vanadium between liquidus olivine and melt. The squares are orogenic eclogites, the circles are mantle eclogites and the diamond is a mid-ocean ridge ophiolite from Aulbach and Stagno (Reference Aulbach and Stagno2016). The symbols display the ƒO₂ (corrected to 1 GPa) calculated from V/Sc ratios. The vertical error bars are predicted 1σ errors of the V/Sc ratios (representing 1σ of the mean per sample suite) and the horizontal error bars show age ranges or 1σ errors for isochron ages from the literature. The red star shows the calculated ƒO₂ of the modern MORB and the arrow points toward the estimated ƒO₂ of the uppermost mantle according to Trail et al. (Reference Trail, Watson and Tailby2011) of samples from 4.4 Ga. QFM, quartz-fayalite-magnetite.

The oxygen isotope ratios in these up to 4.4 Ga zircons have provided valuable insights into the history of early Earth’s water cycle. Studies by Peck et al. (Reference Peck, Valley, Wilde and Graham2001), Valley et al. (Reference Valley, Peck, King and Wilde2002) and Cavosie et al. (Reference Cavosie, Valley and Wilde2005) also examined oxygen isotope ratios in Hadean zircon, some of which yield elevated δ¹⁸O values consistent with the assimilation of sediments or crustal material derived from low-temperature water–rock interactions. The isotopic signatures found in these zircons thus suggest the presence of liquid water on the Earth’s surface during the Hadean as early as 4.4 Ga. These findings support the presence of a hydrosphere on Earth at that time. If correct, the δ¹⁸O of the hydrosphere needs to be considered because the δ¹⁸O of the early oceans may have been lower than the present oceans (Wallmann, Reference Wallmann2001; Sengupta and Pack, Reference Sengupta and Pack2018; Herwartz et al., Reference Herwartz, Pack and Nagel2021; Tatzel et al., Reference Tatzel, Frings, Oelze, Herwartz, Lünsdorf and Wiedenbeck2022; Isson and Rauzi, Reference Isson and Rauzi2024) and meteoric water generally comprises lower δ¹⁸O than seawater. Due to the variability of δ¹⁸O in the hydrosphere, water–rock interaction at high and low temperatures can result in a large range of silicate δ¹⁸O, and assimilation of such altered material is also known to generate low δ¹⁸O magmas (Bindeman et al., Reference Bindeman, Schmitt and Evans2010; Herwartz et al., Reference Herwartz, Pack, Krylov, Xiao, Muehlenbachs, Sengupta and Di Rocco2015; Zakharov et al., Reference Zakharov, Bindeman, Tanaka, Friðleifsson, Reed and Hampton2019). In general, assimilation of the altered mafic crust may not lead to elevated δ¹⁸O of magmas from which the zircons crystallised, therefore a mantle-like δ¹⁸O of Hadean zircons may be taken cautiously as an argument for the mantle origin of the zircons. An origin from the crust, however, would not allow any conclusions to be drawn about the redox state of the mantle.

Indeed, Hopkins et al. (Reference Hopkins, Harrison and Manning2008) suggested that the Hadean zircons they studied (which contained mineral inclusions) formed in a crustal setting. Based on the derived formation pressure and temperature (700°C, 7 kbar), they concluded that the surface heat flow from 4.2 to 4.0 Ga was only 75 mW/m². Additionally, they proposed that the crustal zircon host melts may have formed above a subduction-like setting, where the subducting slab cools the underlying lithosphere. A similar conclusion was reached by Harrison et al. (Reference Harrison, Schmitt, McCulloch and Lovera2008), who suggested that Hadean zircons formed through crystallisation from crustal magmas. The negative ε(Hf,T) values observed in the set of zircons studied by Harrison et al. (Reference Harrison, Schmitt, McCulloch and Lovera2008) imply formation in a reservoir with sub-chondritic Lu/Hf (i.e. felsic crust), which may have formed as early as 4.5 Ga. Zircons that crystallise from a residual mantle melt should have positive ε(Hf,T) values.

Others used redox-sensitive elements like vanadium and chromium (e.g. Canil, Reference Canil1997; Delano, Reference Delano2001; Lee et al., Reference Lee, Brandon and Norman2003; Li and Lee, Reference Li and Lee2004; Aulbach and Viljoen, Reference Aulbach and Viljoen2015; Nicklas et al., Reference Nicklas, Puchtel and Ash2016, Reference Nicklas, Puchtel and Ash2018, Reference Nicklas, Puchtel, Ash, Piccoli, Hanski, Nisbet, Waterton, Pearson and Anbar2019; Aulbach and Stagno, Reference Aulbach and Stagno2016) or the Fe³⁺/(Fe³⁺ + Fe²⁺) ratio (Rollinson et al., Reference Rollinson, Adetunji, Lenaz and Szilas2017; Aulbach et al., Reference Aulbach, Woodland, Vasilyev, Galvez and Viljoen2017) to determine the redox state of early Earth’s mantle. Applying these methods, it was claimed that oxidation occurred early because since 3.9 to 3.5 Ga samples exhibit ƒO₂ similar to modern mid-ocean ridge basalts (MORBs; Canil, Reference Canil1997; Delano, Reference Delano2001; Li and Lee, Reference Li and Lee2004; Rollinson et al., Reference Rollinson, Adetunji, Lenaz and Szilas2017). In contrast, more recent studies observed that Archean samples are still relatively reduced (QFM – 1.19 ± 0.33) compared to the post-Archean samples (including MORB: QFM – 0.26 ± 0.44). This observation hints at a transition from a relatively reduced towards an oxidised upper mantle during the mid or late Archean (Lee et al., Reference Lee, Brandon and Norman2003; Aulbach and Viljoen, Reference Aulbach and Viljoen2015; Aulbach and Stagno, Reference Aulbach and Stagno2016; Aulbach et al., Reference Aulbach, Woodland, Vasilyev, Galvez and Viljoen2017; Stagno and Fei, Reference Stagno and Fei2020) (Figs 3, 4a), which would have direct consequences on the volcanic outgassing efficiency and atmospheric evolution (Guimond et al., Reference Guimond, Noack, Ortenzi and Sohl2021).

The data shown in Fig. 3 suggest a gradual increase in the ƒO₂ of the upper mantle observed from 3.0 to 2.0 Ga, which contrasts with a sudden increase proposed by previous studies (Canil, Reference Canil1997; Delano, Reference Delano2001; Li and Lee, Reference Li and Lee2004; Scaillet and Gaillard, Reference Scaillet and Gaillard2011; Rollinson et al., Reference Rollinson, Adetunji, Lenaz and Szilas2017). This discrepancy between a gradual and sudden increase in the redox sate is explained by heterogeneity of the early upper mantle due to incomplete mixing with the lower mantle, the addition of reduced meteoritic material or inherited from magma ocean processes (Ringwood, Reference Ringwood1979; Arculus, Reference Arculus1985; Nicklas et al., Reference Nicklas, Puchtel, Ash, Piccoli, Hanski, Nisbet, Waterton, Pearson and Anbar2019; Stagno and Fei, Reference Stagno and Fei2020; Stagno and Aulbach, Reference Stagno, Aulbach, Moretti and Neuville2021). Gu et al. (Reference Gu, Li, McCammon and Lee2016) experimentally demonstrated that oxidised lower mantle material is less dense than reduced lower mantle material. This enhances the ascent probability, leading to an efficient mixing between the lower and upper mantle. According to Gu et al. (Reference Gu, Li, McCammon and Lee2016), the upper mantle could have been oxidised within 800 m.y. via this mixing process. However, the process was probably prolonged due to the effect of the strength of bridgmanite (which is about three orders of magnitude higher compared to ferropericlase) on the mantle viscosity and, thus, on the mixing behaviour (Girard et al., Reference Girard, Amulele, Farla, Mohiuddin and Karato2016; Ballmer et al., Reference Ballmer, Houser, Hernlund, Wentzcovitch and Hirose2017; O’Neill and Aulbach, Reference O’Neill and Aulbach2022). Another reason for a delayed mantle mixing could have been a larger grain size resulting from hotter early Earth conditions. This larger grain size could have led to stronger plate boundaries, decreasing convective motion (Foley and Rizo, Reference Foley and Rizo2017). An inefficient mixing of the material from the lower mantle with the upper mantle would explain the preservation of primordial reservoirs suggested to explain observed isotope anomalies (e.g. Mukhopadhyay, Reference Mukhopadhyay2012; Debaille et al., Reference Debaille, O’Neill, Brandon, Haenecour, Yin, Mattielli and Treiman2013; Rizo et al., Reference Rizo, Boyet, Blichert-Toft and Rosing2013, Reference Rizo, Walker, Carlson, Touboul, Horan, Puchtel, Boyet and Rosing2016b, Reference Rizo, Walker, Carlson, Horan, Mukhopadhyay, Manthos, Francis and Jackson2016a; Girard et al., Reference Girard, Amulele, Farla, Mohiuddin and Karato2016; Ballmer et al., Reference Ballmer, Houser, Hernlund, Wentzcovitch and Hirose2017; Mundl et al., Reference Mundl, Touboul, Jackson, Day, Kurz, Lekic, Helz and Walker2017; Horan et al., Reference Horan, Carlson, Walker, Jackson, Garçon and Norman2018; Tusch et al., Reference Tusch, Münker, Hasenstab, Jansen, Marien, Kurzweil, Van Kranendonk, Smithies, Maier and Garbe-Schönberg2021, Reference Tusch, Hoffmann, Hasenstab, Fischer-Gödde, Marien, Wilson and Münker2022).

Furthermore, Aulbach and Stagno (Reference Aulbach and Stagno2016) propose that, in contrast to their suite of samples, the rocks measured by previous studies were not derived from the convective mantle. They argue that the latter intruded into a cratonic setting and thus experienced mixing with the sublithospheric mantle. An oxidised mantle at the end of the Archean has also been suggested due to an increase of mantle mixing gradually over time (O’Neill and Aulbach, Reference O’Neill and Aulbach2022), by a change in interior convection patterns from two-layered to one-layered mantle convection (Breuer and Spohn, Reference Breuer and Spohn1995) or by the onset of plate tectonics (Debaille et al., Reference Debaille, O’Neill, Brandon, Haenecour, Yin, Mattielli and Treiman2013; Andrault et al., Reference Andrault, Muñoz, Pesce, Cerantola, Chumakov, Kantor, Pascarelli, Rüffer and Hennet2018). The main argument for mantle mixing due to plate tectonics is to allow the more oxidised, bridgmanite-rich lower mantle (Mao and Bell, Reference Mao, Bell, Saxena, Bhattacharji, Annersten and Stephansson1977; Frost et al., Reference Frost, Liebske, Langenhorst, McCammon, Trønnes and Rubie2004, Reference Frost, Mann, Asahara and Rubie2008; Wade and Wood, Reference Wade and Wood2005) to efficiently mix with the more reducing upper mantle material due to slabs penetrating and stirring up the lower mantle. The mechanisms resulting in mantle mixing could explain the observed rise in upper mantle ƒO₂ between 3.0 and 2.0 Ga (Figs 3, 4a; Aulbach and Viljoen, Reference Aulbach and Viljoen2015; Aulbach and Stagno, Reference Aulbach and Stagno2016; Aulbach et al., Reference Aulbach, Woodland, Vasilyev, Galvez and Viljoen2017; Stagno and Fei, Reference Stagno and Fei2020; O’Neill and Aulbach, Reference O’Neill and Aulbach2022).

First hints of locally oxidised surface environments around 3.0 Ga: implications from stable isotopes

The first geochemical evidence for locally oxidised conditions in marginal marine basins comes from measurements of stable isotopes (e.g. chromium, molybdenum, uranium) of marine black shales (e.g. Anbar et al., Reference Anbar2007; Scott et al., Reference Scott, Lyons, Bekker, Shen, Poulton, Chu and Anbar2008; Lyons et al., Reference Lyons, Reinhard and Planavsky2014; Planavsky et al., Reference Planavsky2014; Kendall et al., Reference Kendall, Creaser, Reinhard, Lyons and Anbar2015; Ossa Ossa et al., Reference Ossa Ossa, Hofmann, Vidal, Kramers, Belyanin and Cavalazzi2016, Reference Ossa Ossa, Hofmann, Wille, Spangenberg, Bekker, Poulton, Eickmann and Schoenberg2018; Wang et al., Reference Wang2018, Reference Wang, Ossa Ossa, Hofmann, Agangi, Paprika and Planavsky2020; Brüske et al., Reference Brüske2020; Kendall, Reference Kendall2021). Earth’s mantle and crustal rocks contain chromium in a trivalent state. In modern surface environments, Cr(III) is oxidised to soluble Cr(VI), which is preferentially enriched in heavy isotopes (δ⁵³Cr > 0). The oxidation occurs by the reaction of Cr(III) with Mn(IV) oxides, which require free O₂ exceeding 0.1–1% of the present atmospheric level (PAL) (Planavsky et al., Reference Planavsky2014); therefore heavy chromium isotopes are a proxy for the presence of free O₂ in the surface environment. Molybdenum isotopes are another tracer for the presence of free O₂. Molybdenum adsorbs on Mn(IV) oxide surfaces, a reaction with strong mass-dependent fractionation toward lighter isotopes, therefore low δ^98/95Mo values hint towards the existence of Mn(IV) oxides, which require free O₂ to form.

No chromium isotope fractionation has been observed in 3.8 Ga banded iron formations (BIFs) from Isua (Frei et al., Reference Frei, Gaucher, Poulton and Canfield2009), which is taken as evidence for atmospheric O₂ pressures below 0.02–0.2 bar (i.e. 0.1–1% PAL) (Fig. 4b). This can be regarded as an indication that oxygenation of the surface reservoir had not yet initiated at 3.8 Ga. The earliest hints of locally oxidised conditions are currently recorded in the 3.0 Ga Sinqeni Formation of the Mozaan Group in South Africa (Planavsky et al., Reference Planavsky2014; Ossa Ossa et al., Reference Ossa Ossa, Hofmann, Vidal, Kramers, Belyanin and Cavalazzi2016, Reference Ossa Ossa, Hofmann, Wille, Spangenberg, Bekker, Poulton, Eickmann and Schoenberg2018; Smith and Beukes, Reference Smith and Beukes2023) (Fig. 4c). Contemporaneous oxidative weathering in soils was suggested based on the extensive mobilisation of redox-sensitive elements and fractionation of the redox-sensitive δ⁵³Cr value. Crowe et al. (Reference Crowe, Døssing, Beukes, Bau, Kruger, Frei and Canfield2013) reported marked negative δ⁵³Cr from the 3.0 Ga Nsuze paleosol and small positive δ⁵³Cr from contemporaneous Ijzermyn iron formation (both from the Pongola Supergroup, South Africa). They concluded that free O₂ exceeding 0.1% PAL existed in the Mesoarchean, some 600 m.y. before the GOE. However, modern weathering was identified at this site and may have altered the chromium isotope ratios (Albut et al., Reference Albut, Babechuk, Kleinhanns, Benger, Beukes, Steinhilber, Smith, Kruger and Schoenberg2018, Reference Albut, Kamber, Brüske, Beukes, Smith and Schoenberg2019). Post-depositional alteration as the cause for the measured chromium isotope fractionation was supported by Heard et al. (Reference Heard, Aarons, Hofmann, He, Ireland, Bekker, Qin and Dauphas2021). They could not confirm the fractionation of chromium isotopes in the Pongola Supergroup paleosol and concluded that the Mesoarchean was anoxic. Irrespective of these arguments, Smith and Beukes (Reference Smith and Beukes2023) combined evidence from detailed stratigraphy, mineralogy, petrography and carbonate mineral chemistry with isotopic evidence from δ¹³C to conclude that the local surface ocean within this basin was oxidised supporting previous δ⁵⁶Fe and δ⁹⁸Mo data. They suggest microaerophilic chemolithoautotrophs were responsible for iron and manganese oxidation, which would require the presence of free oxygen in the water column, but not the atmosphere. Thus, at least concerning O₂, the chemical exchange between the hydrosphere and atmosphere can be suppressed. In the following, evidence for a persistently anoxic Archean atmosphere is summarised.

Atmospheric O₂ content remains low between 3.25 and 2.75 Ga: implications from mineral archives

In addition to stable isotopes, indirect proxies such as certain mineral deposits can be used as oxygen barometers. Many minerals that are stable in the subsurface environment become oxidised when exposed to the O₂-rich modern atmosphere. Notable among these minerals are sulfides like pyrite (FeS₂), uraninite (UO₂]) or siderite (FeCO₃). In the presence of O₂, pyrite is oxidised to Fe(III) (oxyhydr)oxides (rust), uraninite to soluble hexavalent species and siderite to Fe(III) (oxyhydr)oxides. Fluvial uraninite and pyrite detritus were described, e.g. by Ramdohr (Reference Ramdohr1958) and Schidlowski (Reference Schidlowski1981), in Archean sedimentary rocks from the Witwatersrand basin (South Africa). The rounded shape of the mineral grains and absence of oxidation rims suggest that they once occurred as river sand in an O₂-free Archean environment. Detrital pyrite, gersdorffite [NiAsS], uraninite and siderite were described by Rasmussen and Buick (Reference Rasmussen and Buick1999) from Archean (3.25–2.75 Ga) fluvial sediments from Pilbara (Australia) and later by Hofmann et al., (Reference Hofmann, Bekker, Rouxel, Rumble and Master2009) from South Africa (3.2–2.7 Ga; Fig. 4c). These minerals can be used as oxygen barometers. For instance, the stability of uraninite in the surface environment is limited to atmospheric O₂ levels below 10^-2 times the PAL (Grandstaff, Reference Grandstaff1980). Detailed thermodynamic modelling resulted in an upper p(O₂) limit of 3.2 × 10^-5 bar (1.4 × 10^-4 times the PAL) (Johnson et al., Reference Johnson, Gerpheide, Lamb and Fischer2014). The presence of detrital siderite puts an upper limit not only on free O₂ but also on H₂S. Abundant H₂S would lead to the pyritisation of siderite, which is not observed in the Archean sediments studied by Rasmussen and Buick (Reference Rasmussen and Buick1999). They concluded that the Archean atmosphere was poor in H₂S, with levels below 10^–5 bar.

Hexavalent sulfur S(VI), as present in sulfate (SO₄^2–), should not exist in the reduced Archean environment. Instead, S(IV), S(0), or S(–II) should be the dominating sulfur oxidation states in equilibrium with the lithosphere and atmosphere. However, sedimentary and hydrothermal barite (BaSO₄) exists in Paleo- and Mesoarchean rocks from Australia (e.g. Dresser Formation) and South Africa (Barberton Greenstone Belt) (Heinrichs and Reimer, Reference Heinrichs and Reimer1977; Thorpe, Reference Thorpe1979; Walter et al., Reference Walter, Buick and Dunlop1980; Lowe et al., Reference Lowe, Drabon and Byerly2019) (Fig. 4c). The presence of oxidised sulfate within at least some surface waters is regarded as disequilibrium sulfate, i.e. it is produced by local processes but is not in thermodynamic equilibrium with the entire reduced environment (Olson et al., Reference Olson, Drabon and Johnston2022). Thus, the presence of sulfate minerals (such as barite) in the geological record is generally not regarded as representative of the redox state of the Archean ocean (Huston and Logan, Reference Huston and Logan2004).

One process to obtain the S(VI) to form barite is the UV-induced photodissociation and disproportionation of SO₂ from volcanic degassing into reduced elemental sulfur S(0) and oxidised sulfate S(VI). Indeed, the Paleo- and Mesoarchean sulfate comprises sulfur isotope signatures revealing at least a partial origin from the atmosphere (Bao et al., Reference Bao, Rumble and Lowe2007; Ueno et al., Reference Ueno, Ono, Rumble and Maruyama2008). Triple oxygen isotope data reveal at least two distinct sources of oxygen in sulfate. Apart from an atmospheric endmember, photooxidation of dissolved Fe²⁺ to Fe³⁺ could have acted as a sulfur oxidiser and microbial sulfur cycling may also have been significant (Olson et al., Reference Olson, Drabon and Johnston2022). Further suggestions include sulfate formation through the reaction between reduced S(IV, 0, –II) components and water at high temperatures of igneous systems and the disproportionation of SO₂ in hydrothermal systems (Halevy, Reference Halevy2013). Thus, sulfate does not require an oxidised environment but may result from very particular reactions involving SO₂ from volcanic emissions. All these processes form sulfate, which is in thermodynamic disequilibrium with the atmo-, hydro- and lithosphere and hence contains little information about the redox state of the Archean Earth.

Large-scale oxidation begins around 2.7 Ga: insights from iron formations

Iron formations (IFs) are iron- and silica-rich marine chemical sediments that commonly display a distinct banding (i.e. banded iron formations, BIFs) (e.g. Bekker et al., Reference Bekker, Slack, Planavsky, Krapez, Hofmann, Konhauser and Rouxel2010; Konhauser et al., Reference Konhauser2017; Mänd et al., Reference Mänd, Robbins, Planavsky, Bekke and Konhauser2021; Aftabi et al., Reference Aftabi, Atapour, Mohseni and Babaki2021; Dreher et al., Reference Dreher, Schad, Robbins, Konhauser, Kappler and Joshi2021). Two main endmember types are distinguished. Algoma-type IFs are generally associated with volcanic provinces and comprise large positive europium anomalies inherited from anoxic vent fluids. These comparably small-scale deposits appear throughout the Archean and early Proterozoic (Barrett et al., Reference Barrett, Fralick and Jarvis1988; Bolhar et al., Reference Bolhar, Van Kranendonk and Kamber2005; Ohmoto et al., Reference Ohmoto, Watanabe, Yamaguchi, Naraoka, Haruna, Kakegawa, Hayashi and Kato2006b; Bekker et al., Reference Bekker, Slack, Planavsky, Krapez, Hofmann, Konhauser and Rouxel2010; Pirajno and Yu, Reference Pirajno and Yu2021). Superior-type IFs form on continental shelves covering extensive areas between 2.7 and 1.8 Ga (Fig. 4c), with a few occurrences already around 3.0 Ga (Huston and Logan, Reference Huston and Logan2004; Smith and Beukes, Reference Smith and Beukes2023). Especially after 2.4 Ga, some of these formed above the storm wave base, destroying the banding and generating granular iron formations (GIFs). The depositional depth seems to be related to the depth of the photic zone (Herwartz and Viehmann, Reference Herwartz and Viehmann2024). Superior-type IFs exhibit smaller europium anomalies, pointing to dominant contributions of rare earth elements derived from continental weathering or low-temperature alteration of oceanic crust rather than hydrothermal vents. The direct precipitation from open seawater makes superior-type IFs the prime target for reconstructing ambient seawater conditions (Bekker et al., Reference Bekker, Slack, Planavsky, Krapez, Hofmann, Konhauser and Rouxel2010; Konhauser et al., Reference Konhauser2017; Mänd et al., Reference Mänd, Robbins, Planavsky, Bekke and Konhauser2021). Today, IFs comprise iron-rich phases, including hematite, magnetite, siderite and iron silicates with variable redox states (mean oxidation state of ~Fe^2.4+) and low (<<0.5 wt.%) organic carbon content (Klein and Beukes, Reference Klein and Beukes1992; Trendall, Reference Trendall, Altermann and Corcoran2002). However, the mineralogy observed today does not represent the primary precipitates from an ancient ocean (Konhauser et al., Reference Konhauser2017; Muhling and Rasmussen, Reference Muhling and Rasmussen2020). Most candidates for primary precipitates comprise Fe(III) (but see Muhling and Rasmussen, Reference Muhling and Rasmussen2020). Hence, large-scale oxidation of soluble Fe(II) to insoluble Fe(III) is required to form IFs. Several abiotic and biotic mechanisms have been suggested, most of which are proposed to occur within the photic zone of ocean water.

In the absence of an ozone layer, UV irradiation reaches the Earth’s surface, which induces photochemical oxidation of dissolved Fe²⁺ to Fe³⁺ (Cairns-Smith, Reference Cairns-Smith1978; Braterman et al., Reference Braterman, Cairns-Smith and Sloper1983; Anbar and Holland, Reference Anbar and Holland1992). It is suggested that this process occurs at a sufficient rate to form IF deposits (François, Reference François1986). In contrast, Konhauser et al. (Reference Konhauser, Amskold, Lalonde, Posth, Kappler and Anbar2007a) argue that the photochemical contribution to solid-phase precipitation is negligible, as most of the Fe²⁺ quickly forms poorly crystalline precursor phases to Fe(II) silicates and/or Fe(II) carbonates. The rate of indirect photochemical oxidation via atmospheric H₂O₂ is found to be too low to account for depositional rates of IF (Pecoits et al., Reference Pecoits, Smith, Catling, Philippot, Kappler and Konhauser2015). Another source of H₂O₂ is the decay of primordial radioactive isotopes dissolved in seawater. Ershov (Reference Ershov2021) estimates that the decay of highly soluble ⁴⁰K alone may account for the oxidation of 10²¹ g of iron within a period between 4.3 and 2.5 Ga. The aqueous oxidation of Fe²⁺ to Fe³⁺ is favourable at high pH because this reaction generates protons (2Fe²⁺ + 4H₂O → 2FeOOH + H₂ + 4H⁺). Shibuya et al. (Reference Shibuya, Komiya, Nakamura, Takai and Maruyama2010) argue that high-temperature hydrothermal vent fluids, which are acidic today, had elevated pH in the Archean and comprised Fe³⁺. Experimental results by Dodd et al. (Reference Dodd2022) show that the decomposition of Fe(OH)₂ in Archean seawater analogues produces Fe³⁺ species. The Fe(OH)₂ compound is stable at elevated pH.

The spontaneous conversion of green rust (Fe₄²⁺Fe₂³⁺(OH)₁₂SO₄•8H₂O) to magnetite (Fe²⁺Fe³⁺₂O₄) goes along with a net increase in Fe³⁺ (Tamaura et al., Reference Tamaura, Yoshida and Katsura1984; Li et al., Reference Li, Konhauser and Zhai2017). Green rust is commonly considered a primary iron precipitate in Archean oceans (e.g. Sun et al., Reference Sun, Lechte, Shi, Zhou, Zhou, Feng, Xie, Wu and Tang2022) which are, however, considered to be sulfate-poor, at least between 3.2 and 2.4 Ga (Huston and Logan, Reference Huston and Logan2004). Archean seawater chemistry (including pH and ion concentrations) considerably affects the efficiency of abiotic iron oxidation pathways (e.g. Konhauser et al., Reference Konhauser, Amskold, Lalonde, Posth, Kappler and Anbar2007a; Shibuya et al., Reference Shibuya, Komiya, Nakamura, Takai and Maruyama2010). Therefore, the respective net contribution to individual BIF deposits remains unclear and may vary spatially and over time for each abiotic oxidation mechanism.

The proposed biotic iron oxidation mechanisms can be subdivided into indirect oxidation by free O₂ from oxygenic photosynthesis (Cloud, Reference Cloud1973; Klein and Beukes, Reference Klein and Beukes1992) and direct oxidation either by chemolithoautotrophic or anoxygenic photoautotrophic iron-oxidising bacteria (Konhauser et al., Reference Konhauser, Hamade, Raiswell, Morris, Ferris, Southam and Canfield2002; Kappler and Newman, Reference Kappler and Newman2004; Kappler et al., Reference Kappler, Pasquero, Konhauser and Newman2005). The relative proportions of these pathways can be approximated from reactive transport modelling (Ozaki et al., Reference Ozaki, Thompson, Simister, Crowe and Reinhard2019). This approach shows how variations between individual settings with variable nutrient and Fe²⁺ supply and the available light intensity within a given water mass control the dominating oxidation pathway (Ozaki et al., Reference Ozaki, Thompson, Simister, Crowe and Reinhard2019; Herwartz and Viehmann, Reference Herwartz and Viehmann2024).

Manganese in IFs is a main tracer for the oxygenation of Earth’s hydrosphere (Robbins et al., Reference Robbins2023). Tetravalent Mn(IV) oxides form at redoxclines via consumption of dissolved molecular O₂ and are thus direct evidence for oxygenic photosynthesis (see Robbins et al., Reference Robbins2023 for a review). Oxidised Mn⁴⁺, Fe³⁺ and organic matter form particles that sink towards the seafloor. This process is observed in anoxic basins today and is known as the Fe–Mn shuttle (Dellwig et al., Reference Dellwig, Leipe, März, Glockzin, Pollehne, Schnetger, Yakushev, Böttcher and Brumsack2010). While respective particles slowly sink below the chemocline into the anoxic water body, the Mn⁴⁺ is reduced again by dissolved Fe²⁺ (Dellwig et al., Reference Dellwig, Leipe, März, Glockzin, Pollehne, Schnetger, Yakushev, Böttcher and Brumsack2010; Kurzweil et al., Reference Kurzweil, Wille, Gantert, Beukes and Schoenberg2016; Ossa Ossa et al., Reference Ossa Ossa, Hofmann, Wille, Spangenberg, Bekker, Poulton, Eickmann and Schoenberg2018). Deposition of such Mn⁴⁺ particles in the sediment is only viable at low Fe²⁺, e.g. distal to the iron source (Smith and Beukes, Reference Smith and Beukes2023), or when oxygenic photosynthesis is so active that the flux of sinking Mn⁴⁺ particles outcompetes the upwelling flux of Fe²⁺. During and in the aftermath of the GOE, enormous amounts of Mn⁴⁺ particles have been deposited on the seafloor, forming the world’s largest manganese deposits (Gutzmer and Beukes, Reference Gutzmer and Beukes1996; Tsikos et al., Reference Tsikos, Beukes, Moore and Harris2003; Sekine et al., Reference Sekine2011), reflecting the high productivity around that time. Elevated manganese contents are a prime indicator for “whiffs of oxygen” (Planavsky et al., Reference Planavsky2014; Ossa Ossa et al., Reference Ossa Ossa, Hofmann, Vidal, Kramers, Belyanin and Cavalazzi2016; Smith and Beukes, Reference Smith and Beukes2023) and a general increase in manganese contents in IFs is observed at the onset of the GOE, e.g. in the Transvaal Supergroup of South Africa (Tsikos et al., Reference Tsikos, Beukes, Moore and Harris2003; Schröder et al., Reference Schröder, Bedorf, Beukes and Gutzmer2011; Kurzweil et al., Reference Kurzweil, Wille, Gantert, Beukes and Schoenberg2016; Smith, Reference Smith, Siegesmund, Basei, Oyhantçabal and Oriolo2018).

Subsequent oxidation of organic matter in the sediment partially reduced Fe³⁺ and Mn⁴⁺ back to soluble Fe²⁺ and Mn²⁺. Hence, diagenetic processes can be responsible for the variable mineralogy observed in IFs today. For instance, the Fe and Mn in siderite and rhodochrosite can be derived from the oxidation of organic matter by Fe³⁺ and Mn⁴⁺, which precludes the use of IF mineralogy to reconstruct paleo-atmospheric gas concentrations (Reinhard and Planavsky, Reference Reinhard and Planavsky2011). Identifying primary mineral phases and other features, such as the banding of BIFs, has been the main challenge in using these rocks as reliable archives (Mänd et al., Reference Mänd, Robbins, Planavsky, Bekke and Konhauser2021; Mundl-Petermeier et al., Reference Mundl-Petermeier, Viehmann, Tusch, Bau, Kurzweil and Münker2022; Bau et al., Reference Bau, Frei, Garbe-Schönberg and Viehmann2022).

Abundant whiffs of oxygen between 2.6 and 2.5 Ga: implications from stable isotopes and black shales

Whiffs of oxygen in marine sediments become more abundant in the Neoarchean (2.6–2.5 Ga) towards the GOE (Anbar et al., Reference Anbar2007; Scott et al., Reference Scott, Lyons, Bekker, Shen, Poulton, Chu and Anbar2008; Lyons et al., Reference Lyons, Reinhard and Planavsky2014; Kendall et al., Reference Kendall, Creaser, Reinhard, Lyons and Anbar2015; Ostrander et al., Reference Ostrander, Nielsen, Owens, Kendall, Gordon, Romaniello and Anbar2019; Brüske et al., Reference Brüske2020) (Fig. 4c). Frei et al. (Reference Frei, Gaucher, Poulton and Canfield2009) reported on sedimentary rocks with marked positive δ⁵³Cr, suggesting that O₂ rich oases existed before the GOE. These oases probably occurred near the shore, and rivers washed heavy Cr(VI) into the oceans, where chemical sediments preserved the isotope signature.

Significant volumes of black shales started forming at 2.7 Ga (Fig. 4c), indicating a substantial burial of organic carbon that was probably a response to increasing primary productivity via oxygenic photosynthesis (Condie, Reference Condie2001; Lyons et al., Reference Lyons, Reinhard and Planavsky2014). Oxygenic photosynthesis is assumed to be one of the primary mechanisms leading to the significant accumulation of O₂ in the oceans and the atmosphere. Additionally, O₂ accumulation is favoured by organic carbon burial (Lee et al., Reference Lee, Yeung, McKenzie, Yokoyama, Ozaki and Lenardic2016). Thus, black shales indirectly record the enhanced oxidation of the hydrosphere-atmosphere system.

Significant enrichment of free O₂ in the surface reservoir recorded by sediments and sulfur isotopes starting at 2.5 Ga

Widespread release of O₂ from an oxygenated surface ocean is evident after around 2.3 Ga when red beds emerge (Konhauser et al., Reference Konhauser2017) (Fig. 4c). Iron loss is observed during paleosol weathering (Rye and Holland, Reference Rye and Holland1998) and detrital pyrite and uraninite disappearance, just like the MIF-S (mass-independent sulfur isotope fractionation) anomaly disappears with the rise of O₂ in the atmosphere (Bekker, Reference Bekker2001). The MIF-S signal is assumed to occur due to photolysis and/or photoexcitation of volcanogenic SO₂ by ultraviolet light in a reduced, anoxic atmosphere. Reduced sulfur species displaying a positive ∆³³S signature result from this photolytic process (e.g. Farquhar, Reference Farquhar2000; Farquhar et al., Reference Farquhar, Savarino, Airieau and Thiemens2001; Ono et al., Reference Ono, Eigenbrode, Pavlov, Kharecha, Rumble, Kasting and Freeman2003; Whitehill and Ono, Reference Whitehill and Ono2012; Whitehill et al., Reference Whitehill, Xie, Hu, Xie, Guo and Ono2013). The sudden disappearance of the MIF-S signal at ca. 2.3 Ga is one of the most well-known and solid evidence for the accumulation of free O₂ above 10^–5 PAL and thus marks the onset of the GOE (Fig. 4c) (e.g. Pavlov and Kasting, Reference Pavlov and Kasting2002; Ono et al., Reference Ono, Eigenbrode, Pavlov, Kharecha, Rumble, Kasting and Freeman2003; Bekker et al., Reference Bekker, Holland, Wang, Rumble, Stein, Hannah, Coetzee and Beukes2004; Zahnle et al., Reference Zahnle, Claire and Catling2006; Domagal-Goldman et al., Reference Domagal-Goldman, Kasting, Johnston and Farquhar2008; Guo et al., Reference Guo2009; Luo et al., Reference Luo, Ono, Beukes, Wang, Xie and Summons2016; Warke et al., Reference Warke, Di Rocco, Zerkle, Lepland, Prave, Martin, Ueno, Condon and Claire2020; Poulton et al., Reference Poulton, Bekker, Cumming, Zerkle, Canfield and Johnston2021). The positive ∆³³S signal in Archaean sedimentary sulfides is recognisable but comparatively low between 3.9 and 2.7 Ga (except for a peak between ca. 3.2 and 3.2 Ga). Between about 2.7 and 2.5 Ga a pronounced MIF-S spike is observed coinciding with the GOE (Farquhar, Reference Farquhar2000; Mojzsis et al., Reference Mojzsis, Coath, Greenwood, McKeegan and Harrison2003; Ono et al., Reference Ono, Eigenbrode, Pavlov, Kharecha, Rumble, Kasting and Freeman2003; Hu et al., Reference Hu, Rumble and Wang2003; Bekker et al., Reference Bekker, Holland, Wang, Rumble, Stein, Hannah, Coetzee and Beukes2004; Whitehouse et al., Reference Whitehouse, Kamber, Fedo and Lepland2005; Papineau et al., Reference Papineau, Mojzsis, Coath, Karhu and McKeegan2005; Jamieson et al., Reference Jamieson, Wing, Hannington and Farquhar2006; Ohmoto et al., Reference Ohmoto, Watanabe, Ikemi, Poulson and Taylor2006a; Ono et al., Reference Ono, Beukes, Rumble and Fogel2006; Cates and Mojzsis, Reference Cates and Mojzsis2006; Johnston et al., Reference Johnston, Poulton, Fralick, Wing, Canfield and Farquhar2006; Kamber and Whitehouse, Reference Kamber and Whitehouse2007; Papineau et al., Reference Papineau, Mojzsis and Schmitt2007; Philippot et al., Reference Philippot, Van Zuilen, Lepot, Thomazo, Farquhar and Van Kranendonk2007; Kaufman et al., Reference Kaufman, Johnston, Farquhar, Masterson, Lyons, Bates, Anbar, Arnold, Garvin and Buick2007; Bao et al., Reference Bao, Rumble and Lowe2007; Farquhar et al., Reference Farquhar, Johnston and Wing2007; Domagal-Goldman et al., Reference Domagal-Goldman, Kasting, Johnston and Farquhar2008; Partridge et al., Reference Partridge, Golding, Baublys and Young2008; Johnston et al., Reference Johnston, Farquhar, Summons, Shen, Kaufman, Masterson and Canfield2008; Ueno et al., Reference Ueno, Ono, Rumble and Maruyama2008; Ono et al., Reference Ono, Beukes and Rumble2009; Thomazo et al., Reference Thomazo, Ader, Farquhar and Philippot2009a; Shen et al., Reference Shen, Farquhar, Masterson, Kaufman and Buick2009; Guo et al., Reference Guo2009; Gaillard et al., Reference Gaillard, Scaillet and Arndt2011; Lyons et al., Reference Lyons, Reinhard and Planavsky2014; Ono, Reference Ono2017; Kendall, Reference Kendall2021).

While the atmosphere, surface oceans and marginal basins are oxidised after the GOE (Lyons et al., Reference Lyons, Reinhard and Planavsky2014), the deep ocean remains anoxic until 1.8 Ga (Huston and Logan, Reference Huston and Logan2004). Deep ocean oxygenation requires Phanerozoic-like atmospheric O₂ levels and deep ocean convection (Reinhard and Planavsky, Reference Reinhard and Planavsky2022). Therefore, fully oxidised oceans as we know them today do not appear until atmospheric O₂ levels approach modern levels in the late Proterozoic oxidation event (Reinhard and Planavsky, Reference Reinhard and Planavsky2022; but see Xu et al., Reference Xu, Qin, Wang, Li, Shi, Tang and Liu2023), and even then bottom-water anoxia seem to have been the rule rather than the exception until the mid-Paleozoic era (Stockey et al. Reference Stockey, Cole, Farrell, Agić, Boag, Brocks, Canfield, Cheng, Crockford, Cui, Dahl, Del Mouro, Dewing, Dornbos, Emmings, Gaines, Gibson, Gill, Gilleaudeau and Sperling2024).

Figure 4. Oxidation and oxygenation of the upper mantle and the surface environment over time. (a) Calculated ƒO₂ of samples derived from the upper mantle after Aulbach and Stagno (Reference Aulbach and Stagno2016) and Stagno and Aulbach (Reference Stagno, Aulbach, Moretti and Neuville2021) (see Fig. 3 for the legend). The rectangles at the top of the figure display some important geodynamic events: extensive formation of continental crust ca. 3.5–2.4 Ga (Collerson and Kamber, Reference Collerson and Kamber1999; Huston and Logan, Reference Huston and Logan2004), onset of modern style plate tectonics (PT) ca. 3.2–3.0 Ga (Smithies et al., Reference Smithies, Champion, Van Kranendonk, Howard and Hickman2005; Van Kranendonk et al., Reference Van Kranendonk, Smithies, Hickman and Champion2007; Van Kranendonk, Reference Van Kranendonk2011; Duncan and Dasgupta, Reference Duncan and Dasgupta2017; Kuang et al., Reference Kuang, Morra, Yuen, Kusky, Jiang, Yao and Qi2023), major crustal growth ca. 2.7 Ga (Gaillard et al., Reference Gaillard, Scaillet and Arndt2011), intrusion of large igneous provinces (LIPs) 2.5–2.4 Ga (Ernst and Bleeker, Reference Ernst and Bleeker2010; Gumsley et al., Reference Gumsley, Chamberlain, Bleeker, Söderlund, de Kock, Larsson and Bekker2017), first glaciations (Ice) ca. 2.4 Ga (Kirschvink et al., Reference Kirschvink, Gaidos, Bertani, Beukes, Gutzmer, Maepa and Steinberger2000; Gumsley et al., Reference Gumsley, Chamberlain, Bleeker, Söderlund, de Kock, Larsson and Bekker2017). (b) Partial pressure of specific gas species over time after Catling and Zahnle (Reference Catling and Zahnle2020). The two grey dashed lines indicate the partial pressure of CO₂. The upper line (K) is after Kasting (Reference Kasting1987) and Herwartz et al. (Reference Herwartz, Pack and Nagel2021), and the lower line (CZ) is after Catling and Zahnle (Reference Catling and Zahnle2020). (c) Geological evidence for oxygenation of the surface environment is explained in detail in the text. The red star indicates early whiffs of oxygen at 3.0 Ga. WO, more abundant later whiffs of oxygen; BS, black shales; Red, red beds (see the main text for references). Mass-independent sulfur isotope fractionation (MIF-S) marks the positive and negative excursions of Δ³³S in ‰ (after Ono, Reference Ono2017; see also the main text). The data for banded iron formations (BIFs) (S, = superior-type; A, algoma-type) and sulfate deposits (barite) is from Huston and Logan (Reference Huston and Logan2004). Reddish colours display evidence of oxygenation, while blueish colours indicate reduced conditions. The boxes without colours are deposits discussed in the literature as possible hints for redox conditions, even though the general opinion is that they cannot be used as redox proxies. (d) Timetable for the emergence of the microbial metabolic processes discussed in this review. Solid lines represent well-established timeframes. Dashed lines represent tentative timeframes. Question marks signify highly uncertain periods. Adapted from Lepot (Reference Lepot2020) and modified based on references in the text. The red bar across the whole figure indicates the timing of the GOE (2.5–2.3 Ga; see text for references). GOE, Great Oxygenation Event; QFM, quartz-fayalite-magnetite

Constraining the O₂ source-limited oxygenic photosynthesis?

One explanation for this delayed oxygenation of the Earth’s surface environments is that the productivity of oxygenic photosynthesis in early cyanobacteria was limited, decreasing the microbial O₂ flux (see Dick et al., Reference Dick, Grim and Klatt2018 for a detailed review). Cyanobacteria depend on bioavailable nitrogen and are major agents for nitrogen fixation in today’s surface oceans (Field et al., Reference Field, Behrenfeld, Randerson and Falkowski1998; Zehr and Kudela, Reference Zehr and Kudela2011). Nitrogen fixation is catalysed by the enzyme nitrogenase, which contains molybdenum (Postgate, Reference Postgate1998). Molybdenum may have been scarce in the reducing environments of the early Archean Earth, where it was poorly soluble (Williams and Fraústo Da Silva, Reference Williams and Fraústo Da Silva2003). Thus, it was suggested that nitrogen fixation in cyanobacteria was inhibited (Zerkle et al., Reference Zerkle, House, Cox and Canfield2006). At the same time, O₂ output by oxygenic photosynthesis could have inhibited other nitrogen-fixing microorganisms, ultimately starving cyanobacteria of bioavailable nitrogen (Shi and Falkowski, Reference Shi and Falkowski2008; Kasting and Canfield, Reference Kasting, Canfield, Knoll, Canfield and Konhauser2012). Nitrogen fixation could have occurred via lightning-driven atmospheric reactions (Navarro-González et al., Reference Navarro-González, Molina and Molina1998; Wong et al., Reference Wong, Charnay, Gao, Yung and Russell2017). Still, the isotopic composition of most Archean sedimentary nitrogen isotope records suggests this process was not quantitatively important for sustaining primary production (Barth et al., Reference Barth, Stüeken, Helling, Rossmanith, Peng, Walters and Claire2023). Instead, nitrogen isotope evidence is consistent with biological nitrogen fixation by at least 3.2 Ga (Stüeken et al., Reference Stüeken, Buick, Guy and Koehler2015a, Reference Stüeken, Kipp, Koehler and Buick2016). This is consistent with phylogenetic studies suggesting an early emergence of nitrogen fixation in cyanobacteria (Latysheva et al., Reference Latysheva, Junker, Palmer, Codd and Barker2012). Hydrothermal sources may have sufficiently compensated the low supply of molybdenum for nitrogenase from oxidative weathering in the Archean (Evans et al., Reference Evans, Coogan, Kaçar and Seyfried2023). Moreover, hydrothermal systems probably played a role in recycling sedimentary ammonium (Stüeken et al., Reference Stüeken, Boocock, Robinson, Mikhail and Johnson2021; Martin et al., Reference Martin, Stüeken, Michaud, Münker, Weyer, Van Hees and Gehringer2024). Therefore, nitrogen may not have been a limiting factor for cyanobacterial productivity in the late Archean.

Bioavailable phosphorous, in the form of phosphate, is widely considered another limiting factor for primary productivity in the Archean and early Proterozoic ocean (Derry, Reference Derry2015; Reinhard et al., Reference Reinhard, Planavsky, Gill, Ozaki, Robbins, Lyons, Fischer, Wang, Cole and Konhauser2017; Ossa Ossa et al., Reference Ossa Ossa, Hofmann, Spangenberg, Poulton, Stüeken, Schoenberg, Eickmann, Wille, Butler and Bekker2019; Walton et al., Reference Walton, Ewens, Coates, Blake, Planavsky, Reinhard, Ju, Hao and Pasek2023). This is despite its supply from continental weathering (Hao et al., Reference Hao, Knoll, Huang, Hazen and Daniel2020; Watanabe and Tajika, Reference Watanabe and Tajika2021) and possible hydrothermal sources (Rasmussen et al., Reference Rasmussen, Muhling, Suvorova and Fischer2021, Reference Rasmussen, Muhling, Tosca and Fischer2023). For example, Ozaki et al. (Reference Ozaki, Thompson, Simister, Crowe and Reinhard2019) provide a model for open ocean settings and investigate the competition between O₂-producing cyanobacteria and photoferrotrophs, the latter being adapted to lower light levels, allowing them to thrive deeper in the water column. Accordingly, nutrients such as phosphate and Fe²⁺ from upwelling water masses are consumed by photoferrotrophs, leaving surface water starved in either phosphate or Fe²⁺ (Kappler et al., Reference Kappler, Pasquero, Konhauser and Newman2005; Ozaki et al., Reference Ozaki, Thompson, Simister, Crowe and Reinhard2019). If oceanic iron/phosphate ratios are high, oxygenic photosynthesis in the upper water column is efficiently suppressed (Ozaki et al., Reference Ozaki, Thompson, Simister, Crowe and Reinhard2019). In coastal settings, where the water column is shallow and more nutrients are supplied from the continent, benthic microbial mats are observed that probably produce O₂ (Homann et al., Reference Homann, Heubeck, Airo and Tice2015; Homann, Reference Homann2019), therefore the productivity of cyanobacteria probably varied spatially and over time (Konhauser et al., Reference Konhauser2018). Moreover, the bioavailability of phosphorous may have been limited due to inefficient remineralisation of organic matter (Kipp and Stüeken, Reference Kipp and Stüeken2017) or phosphate scavenging by Fe²⁺ and adsorption on or co-precipitation with (biogenic) Fe(III) minerals in ferruginous oceans (Bjerrum and Canfield, Reference Bjerrum and Canfield2002; Laakso and Schrag, Reference Laakso and Schrag2014; Derry, Reference Derry2015). The efficiency of phosphate scavenging by Fe (III) minerals is debated (Konhauser et al., Reference Konhauser, Lalonde, Amskold and Holland2007b; Jones et al., Reference Jones, Nomosatryo, Crowe, Bjerrum and Canfield2015). Recent analyses of carbonate-associated phosphate in Archean rocks also challenge severe phosphate limitation in coeval waters (Ingalls et al., Reference Ingalls, Grotzinger, Presen, Rasmussen and Fischer2022; Crockford and Halevy, Reference Crockford and Halevy2022). Hence, the possible phosphate limitation of Archean cyanobacteria remains an open question.

Phylogenetic evidence suggests the earliest cyanobacteria were benthic freshwater strains that only diversified into brackish and marine habitats in the late Archean (Blank and Sánchez-Baracaldo, Reference Blank and Sánchez-Baracaldo2010; Schirrmeister et al., Reference Schirrmeister, Sanchez-Baracaldo and Wacey2016; Sánchez-Baracaldo et al., Reference Sánchez-Baracaldo, Raven, Pisani and Knoll2017; Grettenberger et al., Reference Grettenberger, Gold and Brown2025). Planktonic cyanobacteria may have only appeared in the Neoproterozoic, expanding their habitat to the open ocean (Sánchez-Baracaldo et al., Reference Sánchez-Baracaldo, Ridgwell and Raven2014, Reference Sánchez-Baracaldo, Bianchini, Di Cesare, Callieri and Chrismas2019; Sánchez-Baracaldo, Reference Sánchez-Baracaldo2015; Schirrmeister et al., Reference Schirrmeister, Sanchez-Baracaldo and Wacey2016). A benthic lifestyle, on the other hand, would have constrained the spatial extent of Archean cyanobacterial habitats to terrestrial or coastal areas, limiting their overall O₂ production (Sánchez-Baracaldo et al., Reference Sánchez-Baracaldo, Ridgwell and Raven2014, Reference Sánchez-Baracaldo, Raven, Pisani and Knoll2017; Lalonde and Konhauser, Reference Lalonde and Konhauser2015; Sánchez-Baracaldo, Reference Sánchez-Baracaldo2015). Once cyanobacteria expanded to marine environments, they may have suffered from iron toxicity in the Archean oceans (Swanner et al., Reference Swanner, Mloszewska, Cirpka, Schoenberg, Konhauser and Kappler2015a, Reference Swanner, Wu, Hao, Wüstner, Obst, Moran, McIlvin, Saito and Kappler2015b; Dreher et al., Reference Dreher, Schad, Robbins, Konhauser, Kappler and Joshi2021). If true, this was most probably due to reactive oxygen species produced during Fe(II) oxidation by photosynthetic O₂ (Rush and Bielski, Reference Rush and Bielski1985). However, more recent experiments did not observe such effects in open bottle cultures that allowed for gas exchange, limiting the accumulation of photosynthetic O₂ to concentrations assumed for Archean oxygen oases (≤10 μM; Herrmann et al., Reference Herrmann, Sorwat, Byrne, Frankenberg-Dinkel and Gehringer2021). In these sunlit environments, early cyanobacteria would have also been exposed to high levels of UV radiation due to the absence of an ozone shield (Mloszewska et al., Reference Mloszewska, Cole, Planavsky, Kappler, Whitford, Owttrim and Konhauser2018). Recently, it has also been suggested that the net O₂ production in Archean cyanobacterial mats was lower than previously thought due to shorter day lengths (Klatt et al., Reference Klatt, Chennu, Arbic, Biddanda and Dick2021) or an inefficient photosystem in early cyanobacteria (Grettenberger and Sumner, Reference Grettenberger and Sumner2024). One or more of these factors could have limited the productivity of cyanobacteria in the late Archean. The increasing abundance of continental crust through the Archean (Fig. 4a) (Kemp and Hawkesworth, Reference Kemp and Hawkesworth2014; Smit and Mezger, Reference Smit and Mezger2017; Korenaga, Reference Korenaga2018) may have helped overcome some of these limitations by supplying weathering-derived phosphate and creating shallow marine habitats, increasing the biological O₂ source. However, reliable primary productivity estimates depend on the magnitudes of these effects, which remain to be determined.

Critical O₂ sinks: reduced gases and solutes in Archean surface environments

Apart from limited biological O₂ production, low ambient O₂ levels could also be due to large fluxes into O₂ sinks. The most prominent O₂ sinks are reduced gases and aqueous solutes in the Archean atmosphere and oceans (e.g. Fe²⁺, Mn²⁺, H₂S, C_org) (Holland, Reference Holland2002; Claire et al., Reference Claire, Catling and Zahnle2006; Gaillard et al., Reference Gaillard, Scaillet and Arndt2011; Lyons et al., Reference Lyons, Reinhard and Planavsky2014, Reference Lyons, Tino, Fournier, Anderson, Leavitt, Konhauser and Stüeken2024; Lee et al., Reference Lee, Yeung, McKenzie, Yokoyama, Ozaki and Lenardic2016; Catling and Zahnle, Reference Catling and Zahnle2020). The sizes and capacities of these sinks are partly controlled by fluid–rock interactions and volcanic outgassing in chemical equilibrium with the redox state of the Earth’s mantle and crust (Gaillard et al., Reference Gaillard, Bouhifd, Füri, Malavergne, Marrocchi, Noack, Ortenzi, Roskosz and Vulpius2021). As discussed above, Earth supposedly accreted from relatively reduced material. However, the mantle subsequently experienced oxidation between 3.0 and 2.0 Ga (Fig. 4), shifting the redox state of volcanic gases toward more oxidised species (Claire et al., Reference Claire, Catling and Zahnle2006; Aulbach and Stagno, Reference Aulbach and Stagno2016; O’Neill and Aulbach, Reference O’Neill and Aulbach2022). When the abundance of continental crust rose, magmatic outgassing was increasingly shallow and subaerial and, thus, more oxidised due to the pressure dependence of volatile speciation in magmatic systems (Holland, Reference Holland2002; Kump and Barley, Reference Kump and Barley2007; Gaillard et al., Reference Gaillard, Scaillet and Arndt2011). Less mafic and ultramafic rocks at the surface meant fewer reduced solutes (e.g. Fe²⁺, Mn²⁺) in seawater (Kump et al., Reference Kump, Kasting and Barley2001; Lee et al., Reference Lee, Yeung, McKenzie, Yokoyama, Ozaki and Lenardic2016) and fewer reduced species from serpentinisation (e.g. H₂) (Hoffmann, Reference Hoffmann2017; Smit and Mezger, Reference Smit and Mezger2017). More continental landmass increased the accommodation space for the burial of reduced sediments, removing these critical O₂ sinks from the surface environment (Canfield, Reference Canfield2005; Lee et al., Reference Lee, Yeung, McKenzie, Yokoyama, Ozaki and Lenardic2016; Zhao et al., Reference Zhao, Mills, Homoky and Peacock2023).

Biotic processes like microbial methanogenesis also control the abundance of gaseous sinks in surface environments. Methanogens are strictly anaerobic archaea that form CH₄ either by reduction of a carbon substrate (e.g. hydrogenotrophic methanogenesis, CO₂ reduction with H₂ as the electron donor) or by disproportionation (i.e. fermentation) of organic compounds (e.g. acetoclastic methanogenesis, acetate disproportionation) (Head, Reference Head, Ehrlich, Newman and Kappler2016) (Table 1). Methane can contribute to O₂ consumption via photochemically generated CH₃ and OH radicals in the atmosphere (Pavlov et al., Reference Pavlov, Brown and Kasting2001; Kasting and Siefert, Reference Kasting and Siefert2002; Claire et al., Reference Claire, Catling and Zahnle2006; Daines and Lenton, Reference Daines and Lenton2016). A study on the ferruginous and sulfate-poor Lake Matano, an Archean ocean analogue site, showed a limitation of dissimilatory sulfate reduction in favour of methanogenesis (Crowe et al., Reference Crowe2011). Methanogens could, therefore, also have been important agents for the remineralisation of organic matter from Archean primary production (e.g. Thompson et al., Reference Thompson2019). Highly ¹³C depleted CH₄ (δ¹³C down to –56‰) from fluid inclusions in hydrothermal quartz of the 3.48 Ga Dresser Formation may represent the oldest direct evidence for methanogenesis (Ueno et al., Reference Ueno, Yamada, Yoshida, Maruyama and Isozaki2006). However, this CH₄ may originate from abiotic organic synthesis (Sherwood Lollar and McCollom, Reference Sherwood Lollar and McCollom2006). Carbon isotopic evidence for methanogenesis is also found in ca. 3.0 Ga fluvio-lacustrine Lalla Rokh Sandstone (δ¹³C_org –30 to –38‰; Stüeken and Buick, Reference Stüeken and Buick2018), and in the shallow marine or lacustrine 2.72 Ga Tumbiana Formation in Western Australia (δ¹³C_org down to –56‰; Thomazo et al., Reference Thomazo, Ader and Philippot2011), which are probably unaffected by hydrothermal CH₄. Indeed, recent molecular clock studies place the emergence of methanogenesis even ≥3.5 Ga (Wolfe and Fournier, Reference Wolfe and Fournier2018; but see Roger and Susko, Reference Roger and Susko2018 for an alternative viewpoint). Assuming biogenic CH₄ fluxes similar to today, the Archean atmosphere could have maintained CH₄ concentrations of thousands of ppm (Pavlov et al., Reference Pavlov, Brown and Kasting2001; Kasting and Siefert, Reference Kasting and Siefert2002; Kharecha et al., Reference Kharecha, Kasting and Siefert2005). The CH₄ flux probably decreased when sufficient seawater sulfate was available and sulfate-reducing bacteria overtook methanogens in organic carbon mineralisation (Zahnle et al., Reference Zahnle, Claire and Catling2006). Declining oceanic nickel concentrations may have further inhibited methanogens (Konhauser et al., Reference Konhauser, Pecoits, Lalonde, Papineau, Nisbet, Barley, Arndt, Zahnle and Kamber2009). It seems plausible that the late Archean decline of the atmospheric CH₄ pool was crucial for the subsequent rise of atmospheric O₂, rendering biogenic CH₄ an important O₂ sink in the Archean. Together with an increasing flux of biogenic O₂, these processes exhausted the capacity of Earth’s gaseous and aqueous O₂ sinks throughout the Archean, paving the way for the GOE.

The Archean aerobic O₂ sink: insights from recent environments, ancient rocks and modern genomes

Another potential O₂ sink in the Archean is its reduction coupled to the oxidation of various electron donors (organic matter, CH₄, NH₄⁺, NO₂, Mn(II), Fe(II), sulfide) for conserving energy (i.e. aerobic respiration) by microorganisms (Table 1). Aerobic respiration is associated with a greater energy yield compared to anaerobic (i.e. O₂-free) respiration, making it highly competitive in environments where O₂ is available. Indeed, it is the most competitive pathway for organic carbon remineralisation to CO₂ on modern Earth and a critical buffer against further atmospheric O₂ accumulation (Berner, Reference Berner1989). Microbial oxidation of CH₄ to CO₂ using O₂ as the terminal electron acceptor (i.e. aerobic methanotrophy) is a critical CH₄ sink (Table 1). Ammonium oxidation to nitrite and nitrate by O₂ (i.e. nitrification; Table 1) is dominantly controlled by aerobic microorganisms like Nitrosomonas and Nitrobacter sp. in modern oceans (Falkowski, Reference Falkowski1997; Stüeken et al., Reference Stüeken, Pellerin, Thomazo, Johnson, Duncanson and Schoepfer2024). Microbial manganese oxidation by O₂ (Table 1) is the main mechanism for Mn(IV) oxide production in modern oceans (Tebo et al., Reference Tebo, Johnson, McCarthy and Templeton2005). This is because the microbially mediated oxidation of Mn(II) by O₂ is up to five orders of magnitude faster than its abiotic counterpart in seawater-like conditions (Nealson et al., Reference Nealson, Tebo and Rosson1988; Tebo, Reference Tebo1991; Hansel, Reference Hansel and Poole2017; Yu and Leadbetter, Reference Yu and Leadbetter2020). Therefore, Mn(II) oxidation in the presence of O₂ is generally mediated by Mn-oxidising microorganisms, even at sub-micromolar O₂ concentrations (Tebo et al., Reference Tebo, Bargar, Clement, Dick, Murray, Parker, Verity and Webb2004, Reference Tebo, Johnson, McCarthy and Templeton2005; Schippers et al., Reference Schippers, Neretin, Lavik, Leipe and Pollehne2005; Clement et al., Reference Clement, Luther and Tebo2009; Learman et al., Reference Learman, Voelker, Vazquez-Rodriguez and Hansel2011). Microaerophilic Fe(II)-oxidising bacteria (e.g. Gallionella, Leptothrix, Mariprofundus sp) couple the oxidation of Fe(II) to the reduction of O₂ (Table 1). At neutral pH, the oxidation of reduced sulfur compounds (e.g. H₂S) to S⁰ or sulfate can occur at micromolar O₂ concentrations using O₂ or nitrate as an electron acceptor (e.g. Beggiatoa; Hentschel and Felbeck, Reference Hentschel and Felbeck1993; Jørgensen and Gallardo, Reference Jørgensen and Gallardo1999; Girguis et al., Reference Girguis, Lee, Desaulniers, Childress, Pospesel, Felbeck and Zal2000; Dahl et al., Reference Dahl, Friedrich and Kletzin2008) (Table 1). In acidic environments (e.g. hydrothermal sulfide systems, acid rock drainage sites), microorganisms can exploit the aerobic oxidation of S⁰ or sulfide minerals, such as pyrite (e.g. Acidothiobacillus; Segerer et al., Reference Segerer, Neuner, Kristjansson and Stetter1986; Dahl et al., Reference Dahl, Friedrich and Kletzin2008). Microbial sulfide oxidation is orders of magnitude faster than abiotic sulfide oxidation in most sedimentary environments, particularly at low O₂ concentrations, highlighting its role in consuming O₂ in the environment (Luther et al., Reference Luther, Findlay, MacDonald, Owings, Hanson, Beinart and Girguis2011).

Despite the significance and diversity of microbial O₂ sinks in recent environments and previous suggestions for the antiquity of aerobic respiration (‘respiration early hypothesis’; Castresana and Saraste, Reference Castresana and Saraste1995), the role of biological O₂ consumption in the Archean is poorly explored. The proliferation of aerobic microorganisms is traditionally assumed to postdate the GOE due to the canonical lower limit for aerobic respiration (the ‘Pasteur Point’) of 2.2 μM O₂ in seawater at 25°C (Devol, Reference Devol1978). However, the discovery of ‘nanaerobic’ life (Baughn and Malamy, Reference Baughn and Malamy2004), respiring aerobically at nanomolar O₂ concentrations, challenges this paradigm. For instance, Escherichia coli, a well-studied model organism, grows aerobically at O₂ concentrations as low as 3 nM (Stolper et al., Reference Stolper, Revsbech and Canfield2010). Moreover, it is increasingly recognised that many microorganisms typically considered strict anaerobes can also respire aerobically at low O₂ concentrations (Cypionka, Reference Cypionka2000; Lee et al., Reference Lee, Youn, Kang and Lee2019; Berg et al., Reference Berg2019). It was cautioned that aerobic growth rates under nanomolar O₂ concentrations are strongly muted, suggesting that anaerobic respiration (e.g. Fe(III) reducers) could outcompete aerobic microorganisms in the Archean (Ducluzeau et al., Reference Ducluzeau, Schoepp-Cothenet, Van Lis, Baymann, Russell and Nitschke2014). However, advances in O₂ microsensing and metatranscriptomic analysis demonstrate that aerobic respiration is widespread in recent environments, even under apparent anoxia, where aerobic microorganisms may consume O₂ faster than it can accumulate (Berg et al., Reference Berg, Ahmerkamp, Pjevac, Hausmann, Milucka and Kuypers2022 and references therein).

Recent oxygen minimum zones and stratified lakes parallel late Archean environments by their low O₂ concentrations within or below the photic zone, shedding light on the possible role of microbial communities and biogeochemical processes before the GOE. Aerobic methanotrophy (Table 1) has been identified as an efficient CH₄ sink coupled to cryptic O₂ from oxygenic photosynthesis in the photic zone of lakes (Oswald et al., Reference Oswald, Milucka, Brand, Littmann, Wehrli, Kuypers and Schubert2015; Milucka et al., Reference Milucka, Kirf, Lu, Krupke, Lam, Littmann, Kuypers and Schubert2015). These conditions could have been widespread in Archean lakes and oceans, supporting the suggested role of aerobic methanotrophy in buffering atmospheric O₂ accumulation (Daines and Lenton, Reference Daines and Lenton2016). Nitrification is efficient in oxygen minimum zones, even at nanomolar O₂ concentrations (Kalvelage et al., Reference Kalvelage, Jensen, Contreras, Revsbech, Lam, Günter, LaRoche, Lavik and Kuypers2011, Reference Kalvelage2015; Füssel et al., Reference Füssel, Lam, Lavik, Jensen, Holtappels, Günter and Kuypers2012; Thamdrup et al., Reference Thamdrup, Dalsgaard and Revsbech2012; Beman et al., Reference Beman, Leilei Shih and Popp2013; Bristow et al., Reference Bristow2016). Optimal rates of microaerophilic Fe(II) oxidation by Sideroxydans were observed at 5–20 μM O₂ (Maisch et al., Reference Maisch, Lueder, Laufer, Scholze, Kappler and Schmidt2019), but the marine strain Mariprofundus can still grow at sub-micromolar O₂ concentrations (Chiu et al., Reference Chiu, Kato, McAllister, Field and Chan2017; McAllister et al., Reference McAllister, Moore, Gartman, Luther, Emerson and Chan2019). Most importantly, however, because the abiotic oxidation of Fe(II) is slow at neutral pH and low O₂ (Søgaard et al., Reference Søgaard, Medenwaldt and Abraham-Peskir2000), microaerophilic Fe(II) oxidation outcompetes abiotic oxidation at or below 50 μM O₂ (Druschel et al., Reference Druschel, Emerson, Sutka, Suchecki and Luther2008). Aerobic sulfide oxidation also occurs in apparently anoxic environments of modern lakes and oxygen minimum zones when the influx and microbial consumption of O₂ are balanced, resulting in a cryptic O₂ cycle (Sommer et al., Reference Sommer2017; Callbeck et al., Reference Callbeck2018; Berg et al., Reference Berg2019). The micromolar O₂ concentrations inferred for late Archean oxygen oases satisfy even conservative lower limits for aerobic respiration, demonstrating that aerobic respiration was viable in Neoarchean and, perhaps transiently, in Mesoarchean surface waters. Cryptic O₂ consumption in recent environments, resulting in O₂ concentrations below the detection limits of modern microsensors, suggests the downwelling and downward diffusion of oxygenated surface waters may have even allowed for aerobic metabolism in apparently anoxic deeper Archean settings.

Early studies suggested that aerobic respiration is required to mass-balance the preserved organic carbon in Archean black shales (Towe, Reference Towe1990). It is difficult to verify this by investigating the rock record because isotope fractionation involved in heterotrophy is much less than during the initial autotrophic carbon fixation (Hayes, Reference Hayes2001). The resulting carbon species of aerobic methanotrophy are strongly depleted in ¹³C and can be bound in carbonates or assimilated in microbial biomass, enabling the reconstruction of CH₄ oxidation in the geological record (Hayes, Reference Hayes2001; Eigenbrode and Freeman, Reference Eigenbrode and Freeman2006). The δ¹³C values in sedimentary carbonaceous matter of down to –60‰ were interpreted as evidence for aerobic methanotrophs in the late Archean (Hayes, Reference Hayes and Schopf1983, Reference Hayes and Bengtson1994; Hayes and Waldbauer, Reference Hayes and Waldbauer2006). Eigenbrode and Freeman (Reference Eigenbrode and Freeman2006) present indirect evidence for aerobic respiration based on δ¹³C_org analysis of ≤2.7 Ga sedimentary rocks. They conclude that the more consistently ¹³C-depleted deep versus shallow facies (δ¹³C_org = –40 to –45‰ and –57 to –28‰, respectively) demonstrate a more prominent role of CH₄ cycling in anoxic deep water versus (aerobic) respiration of photosynthetic organic matter in oxic surface water. The presence of 3β-methylhopane biomarkers in the 2.7–2.5 Ga Transvaal Supergroup and Hamersley Group supports this (Brocks et al., Reference Brocks, Buick, Logan and Summons2003; Eigenbrode et al., Reference Eigenbrode, Freeman and Summons2008; Waldbauer et al., Reference Waldbauer, Sherman, Sumner and Summons2009), but the syngeneity of biomarkers in these localities to their host rocks was contested (Brocks, Reference Brocks2011). Highly depleted δ¹³C signatures in Archean carbonaceous matter could also be explained by anaerobic oxidation of methane (AOM) (Hinrichs, Reference Hinrichs2002; Thomazo et al., Reference Thomazo, Pinti, Busigny, Ader, Hashizume and Philippot2009b; Guy et al., Reference Guy, Ono, Gutzmer, Kaufman, Lin, Fogel and Beukes2012; Stüeken et al., Reference Stüeken, Buick, Anderson, Baross, Planavsky and Lyons2017; Flannery et al., Reference Flannery, Allwood, Summons, Williford, Abbey, Matys and Ferralis2018; Lepot et al., Reference Lepot, Williford, Philippo, Thomazo, Ushikubo, Kitajima, Mostefaoui and Valley2019), which is a major CH₄ sink on modern Earth (Knittel and Boetius, Reference Knittel and Boetius2009). The quantitative importance of AOM may have been limited in the Archean due to low marine sulfate levels before the GOE (Catling et al., Reference Catling, Claire and Zahnle2007) but alternative electron acceptors like Fe(III) seem plausible (Knittel and Boetius, Reference Knittel and Boetius2009; Stüeken and Buick, Reference Stüeken and Buick2018).

Nitrification produces ¹⁵N-enriched residual ammonium (δ¹⁵N up to +35‰) that can be assimilated and recorded in sedimentary organic matter (Mariotti et al., Reference Mariotti, Germon, Hubert, Kaiser, Letolle, Tardieux and Tardieux1981; Casciotti, Reference Casciotti2009; Mandernack et al., Reference Mandernack, Mills, Johnson, Rahn and Kinney2009). Nitrite can be used to oxidise ammonium in the absence of excess O₂ (anaerobic ammonium oxidation, ‘anammox’; Mulder et al., Reference Mulder, Graaf, Robertson and Kuenen1995; Van De Graaf et al., Reference Van De Graaf, Mulder, De Bruijn, Jetten, Robertson and Kuenen1995; Lam et al., Reference Lam, Lavik, Jensen, Van De Vossenberg, Schmid, Woebken, Gutiérrez, Amann, Jetten and Kuypers2009) (Table 1). However, the anammox reaction also requires O₂ because nitrite cannot be produced anaerobically (Stüeken et al., Reference Stüeken, Kipp, Koehler and Buick2016). Ammonium oxidation may also be coupled to the reduction of sulfate (i.e. sulfammox) or Fe(III) (i.e. feammox) (Clement et al., Reference Clement, Shrestha, Ehrenfeld and Jaffe2005; Yang et al., Reference Yang, Weber and Silver2012; Rios-Del Toro et al., Reference Rios-Del Toro, Valenzuela, López-Lozano, Cortés-Martínez, Sánchez-Rodríguez, Calvario-Martínez, Sánchez-Carrillo and Cervantes2018). The latter was suggested as the possibly dominant ammonium oxidation pathway in the early Archean when O₂ scarcity prevented nitrite and sulfate accumulation (Pellerin et al., Reference Pellerin, Thomazo, Ader, Marin-Carbonne, Alleon, Vennin and Hofmann2023). Both denitrification to N₂O or N₂ and anammox produce ¹⁵N-enriched residual nitrate (Stüeken et al., Reference Stüeken, Pellerin, Thomazo, Johnson, Duncanson and Schoepfer2024). Notably, the δ¹⁵N record in metasedimentary rocks can be further shifted to more positive values with increasing metamorphic grade due to the release of isotopically light ammonium or N₂ (Ader et al., Reference Ader, Thomazo, Sansjofre, Busigny, Papineau, Laffont, Cartigny and Halverson2016). Ample nitrogen isotope evidence from low-grade metasedimentary rocks (greenschist facies and below) shows highly variable δ¹⁵N signatures (ca. –11 to +50‰) and a secular increase of δ¹⁵N of approximately 2‰ at ca. 2.8–2.6 Ga (Garvin et al., Reference Garvin, Buick, Anbar, Arnold and Kaufman2009; Godfrey and Falkowski, Reference Godfrey and Falkowski2009; Thomazo et al., Reference Thomazo, Ader and Philippot2011; Busigny et al., Reference Busigny, Lebeau, Ader, Krapež and Bekker2013; Stüeken et al., Reference Stüeken, Buick and Schauer2015b, Reference Stüeken, Kipp, Koehler and Buick2016; Koehler et al., Reference Koehler, Buick, Kipp, Stüeken and Zaloumis2018; Pellerin et al., Reference Pellerin, Thomazo, Ader, Rossignol, Rego, Busigny and Philippot2024). The overall increase in sedimentary δ¹⁵N values strongly suggests an increasing role of ammonium oxidation in marine environments. Sulfate was virtually unavailable as an electron acceptor during this time (e.g. Crowe et al., Reference Crowe2014) and δ¹⁵N values of up to +37.5‰ from marine sediments are inconsistent with feammox (Pellerin et al., Reference Pellerin, Thomazo, Ader, Rossignol, Rego, Busigny and Philippot2024). Therefore, this trend is best explained by the rise of nitrification, denitrification and/or anammox in the Meso- to Neoarchean (Garvin et al., Reference Garvin, Buick, Anbar, Arnold and Kaufman2009; Godfrey and Falkowski, Reference Godfrey and Falkowski2009; Thomazo et al., Reference Thomazo, Ader and Philippot2011; Busigny et al., Reference Busigny, Lebeau, Ader, Krapež and Bekker2013; Stüeken et al., Reference Stüeken, Buick and Schauer2015b, Reference Stüeken, Kipp, Koehler and Buick2016; Koehler et al., Reference Koehler, Buick, Kipp, Stüeken and Zaloumis2018; Pellerin et al., Reference Pellerin, Thomazo, Ader, Rossignol, Rego, Busigny and Philippot2024), consistent with phylogenetic reconstructions (Parsons et al., Reference Parsons, Stüeken, Rosen, Mateos and Anderson2021). These processes, including the presence of nitrate, were probably transient and limited to settings with oxygenated surface waters.

Microaerophilic Fe(II)-oxidising bacteria commonly form characteristic mineral-organic structures that consist of Fe(III) (oxydr)oxide-encrusted stalks (Chan et al., Reference Chan, Emerson and Luther2016), which show good potential for fossilisation (Picard et al., Reference Picard, Kappler, Schmid, Quaroni and Obst2015). The fossil record of these stalks in hydrothermal jaspers, where preservation of such delicate structures is most probable, extends back to at least 1.74 Ga (Little et al., Reference Little2021). The oldest putative findings of such fossils are reported in 3.77 Ga hydrothermal vent deposits (Dodd et al., Reference Dodd, Papineau, Grenne, Slack, Rittner, Pirajno, O’Neil and Little2017; Papineau et al., Reference Papineau, She, Dodd, Iacoviello, Slack, Hauri, Shearing and Little2022), although the biogenicity of these features would require free O₂ in Eoarchean seafloor hydrothermal systems or a different metabolic affinity. On modern Earth, microaerophilic Fe(II) oxidisers inhabit a limited niche at opposing gradients of O₂ and Fe²⁺ in marine and terrestrial environments like hydrothermal vents, sediment–water interfaces and soils (Kappler et al., Reference Kappler, Bryce, Mansor, Lueder, Byrne and Swanner2021). However, in Archean oxygen oases where O₂ was produced in the photic zone and Fe²⁺ was supplied from below, this niche could have been much larger (Holm, Reference Holm1989; Konhauser et al., Reference Konhauser, Hamade, Raiswell, Morris, Ferris, Southam and Canfield2002; Dreher et al., Reference Dreher, Schad, Robbins, Konhauser, Kappler and Joshi2021).

The presence of abundant Mn oxides in late Archean sedimentary rocks is another signature of aerobic metabolism. In the absence of O₂, Mn(II) may also be oxidised by a range of abiotic and other biotic mechanisms, including photooxidation by UV radiation, Mn-dependent anoxygenic photosynthesis and Mn oxidation coupled to alternative electron acceptors like nitrate (Luther et al., Reference Luther, Sundby, Lewis, Brendel and Silverberg1997; Johnson et al., Reference Johnson, Webb, Thomas, Ono, Kirschvink and Fischer2013; Daye et al., Reference Daye, Klepac-Ceraj, Pajusalu, Rowland, Farrell-Sherman, Beukes, Tamura, Fournier and Bosak2019; Liu et al., Reference Liu, Hao, Elzinga, Piotrowiak, Nanda, Yee and Falkowski2020). However, the transfer of Mn oxides from the water column to the sediments and their preservation required an oxygenated depositional environment devoid of Fe(II) and H₂S, in which aerobic microorganisms dominate Mn(II) oxidation today (Jones et al., Reference Jones, Crowe, Sturm, Leslie, MacLean, Katsev, Henny, Fowle and Canfield2011; Smith and Beukes, Reference Smith and Beukes2023; Robbins et al., Reference Robbins2023; Mhlanga et al., Reference Mhlanga, Tsikos, Lee, Rouxel, Boyce, Harris and Lyons2023).

Fossil evidence for the antiquity of aerobic sulfur oxidation is scarce, probably due to the metastable nature of the product S⁰, the poor preservation potential of sulfur oxidiser cells upon silicification and a lack of distinct sulfur isotope signatures (Canfield, Reference Canfield, Valley and Cole2001; Cosmidis et al., Reference Cosmidis, Nims, Diercks and Templeton2019; Nims et al., Reference Nims, Lafond, Alleon, Templeton and Cosmidis2021). Nevertheless, morphological characteristics (size, cell wall structure) combined with paleoecological considerations were used to interpret carbonaceous microstructures in the 2.52 Ga Gamohaan Formation (South Africa) as sulfur-oxidising bacteria similar to Thiomargarita (Czaja et al., Reference Czaja, Beukes and Osterhout2016). Microbial sulfur oxidation is also supported by multiple sulfur isotope compositions of 3.22 Ga sulfates (Nabhan et al., Reference Nabhan, Marin-Carbonne, Mason and Heubeck2020) and possible microbially induced corrosion features on ca. 3.4 Ga detrital pyrite (Wacey et al., Reference Wacey, Saunders, Brasier and Kilburn2011). If true, this does not necessarily indicate an aerobic metabolism due to the potential coupling of sulfur oxidation with nitrate reduction. Chromium isotopes in Archean BIFs trace chromium mobilisation during the oxidative weathering of pyrite exposed on continents by 2.48 Ga (Konhauser et al., Reference Konhauser2011). It was suggested that this was due to the activity of acidophilic sulfide-oxidising bacteria (Konhauser et al., Reference Konhauser2011) or pyrite oxidation by photochemically generated Fe³⁺ (Hao et al., Reference Hao, Liu, Goff, Steadman, Large, Falkowski and Yee2022). In marine environments, aerobic sulfur oxidisers may have been limited to microbial mats, oxidising sulfide generated in underlying sediments (Konhauser, Reference Konhauser2007). In the open ocean, sulfide was dominantly sourced from hydrothermal systems and probably scavenged by Fe²⁺ before it reached oxygenated surface waters (Canfield et al., Reference Canfield, Rosing and Bjerrum2006), although euxinic conditions possibly prevailed locally in the late Archean (Reinhard et al., Reference Reinhard, Raiswell, Scott, Anbar and Lyons2009; Scott et al., Reference Scott, Bekker, Reinhard, Schnetger, Krapež, Rumble and Lyons2011). These examples highlight the fragmented nature of direct evidence for aerobic metabolism from the geological record.

Such biosignatures can be used as calibration points in phylogenetic studies exploring the emergence and diversification of aerobic microorganisms on early Earth. A Mesoarchean origin of aerobic respiration is supported by the emergence of enzymes involved in oxygen cycling to ≥2.9 Ga, i.e. at least ca. 500 m.y. before the GOE (David and Alm, Reference David and Alm2011; Wang et al., Reference Wang, Jiang, Kim, Qu, Ji, Mittenthal, Zhang and Caetano-Anollés2011; Kim et al., Reference Kim, Qin, Jiang, Chen, Xiong, Caetano-Anollés, Zhang and Caetano-Anollés2012; Jabłońska and Tawfik, Reference Jabłońska and Tawfik2021; Boden et al., Reference Boden, Konhauser, Robbins and Sánchez-Baracaldo2021). Possibly, early enzymes catalysing O₂ reduction might be a detoxification mechanism for coping with oxidative stress (‘aerotolerance’) rather than aerobic respiration (Brochier-Armanet et al., Reference Brochier-Armanet, Talla and Gribaldo2009; Gribaldo et al., Reference Gribaldo, Talla and Brochier-Armanet2009; Jabłońska and Tawfik, Reference Jabłońska and Tawfik2021, Reference Jabłońska and Tawfik2022). Aerotolerance could be an adaptation to abiotic sources of reactive oxygen species and O₂ on early Earth (Haqq-Misra et al., Reference Haqq-Misra, Kasting and Lee2011; He et al., Reference He, Wu, Xian, Zhu, Yang, Lv, Li and Konhauser2021, Reference He2023; Stone et al., Reference Stone, Edgar, Gould and Telling2022). Nevertheless, the early origin of these enzymes is consistent with geochemical proxies suggesting locally or transiently sufficient O₂ for aerobic respiration (Anbar et al., Reference Anbar2007; Ostrander et al., Reference Ostrander, Johnson and Anbar2021), current reconstructions of the emergence of oxygenic photosynthesis at ca. 3.0 Ga (Schirrmeister et al., Reference Schirrmeister, Gugger and Donoghue2015; Sánchez-Baracaldo, Reference Sánchez-Baracaldo2015; Cardona et al., Reference Cardona, Sánchez-Baracaldo, Rutherford and Larkum2019; Garcia-Pichel et al., Reference Garcia-Pichel, Lombard, Soule, Dunaj, Wu and Wojciechowski2019; Jabłońska and Tawfik, Reference Jabłońska and Tawfik2021; Fournier et al., Reference Fournier, Moore, Rangel, Payette, Momper and Bosak2021; Boden et al., Reference Boden, Zhong, Anderson and Stüeken2024) and the recent discovery of ancestral high redox potential quinones predating the emergence of cyanobacteria (Elling et al., Reference Elling2025). Notably, for O₂ to leave a proxy record, it must degas into the atmosphere (in the case of MIF-S) or affect the solubility of redox-sensitive elements (e.g. chromium, molybdenum). However, released O₂ may not reach concentrations reflected in O₂ proxies due to its consumption by various O₂ sinks, including aerobic microorganisms. Conversely, oxidative weathering within benthic microbial mats can produce O₂ proxy signals, even if O₂ does not accumulate in the environment due to microbial consumption (Lalonde and Konhauser, Reference Lalonde and Konhauser2015). Therefore, aerobic microorganisms are a plausible O₂ sink after the emergence of oxygenic photosynthesis, even when seawater or atmospheric O₂ concentrations were too low to leave a proxy record. Aerobic microorganisms may thus have consumed O₂ produced by cyanobacteria, even if the O₂ flux was small due to the possible inhibition of oxygenic photosynthesis. Thus, as soon as oxygenic photosynthesis emerged ca. 3.0 Ga, an aerobic niche appeared for microorganisms to exploit. In concert, the current evidence renders it probable that the Archean aerobic biosphere represented an O₂ sink that helped delay the GOE.