A robust test for the presence of ROS in the 3.7 Ga ocean atmosphere system rests on the fidelity with which our Cr and U data record primary seawater signatures as opposed to tracking subsequent modifications through hydrothermal processes during BIF deposition, post-formational BIF alteration, metamorphic overprinting, or recent oxidative weathering. Hydrothermal alteration (i.e. serpentinization, hydration), indeed, shifts the Cr isotope compositions of oxidatively weathered ultramafic rocks to heavier values19. Such isotopically heavy δ53Cr values in altered ultramafic rocks are associated with various secondary and metamorphic Cr-rich minerals and thereby support previous results by Schoenberg and co-workers15. These isotopically heavy secondary Cr minerals likely incorporate Cr(III) as the product of Cr(VI) reduction in mineral-forming fluids. Such a process should be accompanied then by an associated shift (Rayleigh distillation) of the residual Cr in the fluid towards heavier δ53Cr values24. Redistribution of Cr resulting in isotopically heavy Cr signatures in granitic25, basaltic26,27 and ultramafic28 modern and ancient rocks, point to dynamic and complex redox cycling of Cr in some of the Earth’s crustal and near-surface environments. A key point here, however, is the requirement of redox processes and therefore ROS, in both imparting Cr isotope fractionation and controlling the magnitude of Cr redistribution and the expression of isotopic fractionation at the scale of bulk rock samples. We can evaluate the extent of such alteration for the Isua BIFs we analyzed by combining diverse geochemical and petrological data available from Isua rocks with our new data.

Zinc-isotopes in a variety of rocks from Isua reveal a pronounced depletion in isotopically heavy Zn with respect to the igneous average pointing to a scenario whereby Isua rocks were permeated by carbonate-rich, high-pH reducing hydrothermal solutions at temperatures between 100–300 °C29. Unlike in the studies mentioned above, the BIFs and associated volcanic and sedimentary rocks in the least altered and least deformed central tectonic domain of the eastern ISB do not exhibit features of secondary hydrothermal overprinting (see Supplementary information for details). Our detailed petrographic inspection did not reveal secondary carbonates, which are the likely hosts for Zn species (and potentially also Cr(VI) species) under the low-temperature, highly alkaline hydrothermal conditions indicated previously29. Likewise, the absence of Ce-anomalies and the preservation of LREE-depleted, seawater-type REY patterns with strong positive Eu anomalies, which we report for the BIFs studied herein (Supplementary information) argue against large-scale permeation of BIF by post-formational hydrothermal solutions in this part of the ISB. Such hydrothermal fluids, if they had been present, would have likely caused an enrichment of the LREE signatures of the BIF samples, a feature we do not observe in our samples (see Supplementary information). Our analyses did not identify (despite intensive optical and electron microscopic investigations) secondary U-bearing phases (such as apatite, monazite etc.) which would be expected following U redistribution by hydrothermal fluids. This is corroborated by Pb isotope results of BIFs from within the same tectonic domain, which indicate a closed U-Pb system since BIF deposition30. Likewise, Sm–Nd isotopic relationships between individual layers reveal two distinct REE sources; seafloor-vented hydrothermal fluids (εNd (3.7 Ga)∼+3.1) and ambient surface seawater. The latter attained its composition by erosion of parts of the protocrustal landmass (εNd(3.7 Ga) ∼ +1.6)30. A third (i.e. post-depositional) REE source is not apparent. Finally, other sedimentary and volcanic rocks within the same tectonic domain as the BIFs studied herein preserve igneous inventory δ53Cr values and typical crustal U/Th ratios, a feature that speaks strongly against pervasive secondary alteration or metamorphic redistribution of these elements.

In an approach similar to that used by Li and co-workers31 for the 3.46 Ga Marble Bar chert (MBC) from the Pilbara Craton, NW Australia, we estimated the concentration of U in the Eoarchean Isua seawater basin, based on the U concentrations measured in BIF samples from Isua. Uranium concentrations in the Isua BIFs are, on average, around four times higher (~27 ppm) than in the MBC (~6 ppm; excluding three samples with anomalously high concentrations31), but measured U/Th values agree well (Isua: 0.70 ± 0.29; MBC: 0.67 ± 0.1731). The higher U concentrations in the Isua BIF samples may reflect the affinity of U(VI) (and Th(IV)) for sorption onto Fe(III) oxyhydroxides, i.e. they reflect the higher Fe 2 O 3 concentrations in the Isua BIFs (5.6–77.6 wt%) relative to the cherts of MBC (0.06–9.7 wt%)31 shown in Fig. 2b. The Isua BIFs (and likewise the MBC) are not 100% pure detritus-free chemical sediments and minor detrital contamination, for example, is indicated by Al 2 O 3 concentrations between 0.02 to 1.02 wt% (Supplementary Table 1) for the individual BIF mesobands. This suggests that in both sample sets, the pure marine chemical sediment-endmember showed an even higher U/Th ratio than those measured in the Isua BIF and Marble Bar chert samples. The U/Th ratios above the average crustal value of ~0.26 in the chemical sediments imply a relative enrichment of dissolved U relative to Th in the seawater from which the chemical sediments precipitated. The required decoupling of U and Th can only result from oxidative mobilization of U in the Earth’s surface system.

Based on authigenic (detrital and decay-corrected U (U*; Supplementary Table 1; details in Li et al.31) and Fe 2 O 3 concentration relationships in the MBC samples and a conservative U distribution coefficient (K d ) value of 104 between Fe(III) hydroxide and aqueous solution32,33 (taking the potentially lower seawater pH and higher atmospheric CO 2 contents during the Archean relative to present conditions into consideration), Li and co-workers31 showed that the U concentration of 3.46 Ga seawater was at least two orders of magnitude lower (20–750 ppt) than that of modern seawater (3 ppb34), a result which according to Li and co-workers31 attests to anoxic atmospheric and ocean conditions at 3.46 Ga. Using the same parameters and 3.7 Ga for the decay correction, we calculated even lower U concentrations in the Isua seawater, with values ranging between 1.2 and 32 ppt (Fig. 3). The low redox potential of U(VI)/U(IV) couple makes U an element that is very sensitive to continental oxidation. We here emphasize, in conjunction with the positively fractionated Cr isotope data measured for the Isua BIF, that oxidative removal of U and Cr from the continental landmasses was apparently possible under very low oxygen levels of the 3.7 Ga atmosphere. The similarly elevated U/Th ratios in the MBC31 indicate that mobilization of U under low atmospheric oxygen levels may have persisted at least until 3.47 Ga.

Figure 3 U versus Fe 2 O 3 diagram with BIF samples from Isua. Data from Isua are compared with U and Fe 2 O 3 contents of the 3.46 Ga Marble Bar Chert (MBC)31 samples and with Phanerozoic cherts from three different tectonic settings (grey filled area; data sources in Li et al.31. U* depict authigenic U concentrations in the BIFs (measured U corrected for 3.7 Ga decay and for detrital share using Th concentrations and an average crustal U/Th of 0.25). Using a conservative U K d value of 104 between Fe(III) hydroxides and aqueous solution32,33, very low seawater U (U SW ) concentrations of between 1.2 to 32 ppt (dashed red lines) are estimated for the 3.7 Ga seawater. The dotted black line is the reference MBC 3.42 Ga seawater U = 20 ppt line31. In conjunction with the Cr isotope data presented herein (Fig. 1a) and the positive U-Cr correlation (Supplementary Fig. 7) defined by the Isua BIF samples, we hypothesize that very low atmospheric oxgen levels were sufficient to mobilize small amounts of redox sensitive U and Cr through oxidative weathering conditions prevailing on land at 3.7 Ga. Full size image

Collectively, our data imply that oxidative processes were active under a very low oxygen atmosphere during the deposition of Isua sediments. Such oxidative processes, nevertheless, require the presence and action of ROS. Czaja and co-authors23 concluded that the δ56Fe distributions in Isua BIF reflect concentrations of <0.001% modern seawater O 2 in the photic zone. Such low oxygen levels are well in line with independent proxies for atmospheric oxygen levels including the photochemical mass independent fractionation of S-isotopes35, which implies less than about 10−3 to 10−5 present atmospheric level (PAL) O 2 .

Large-scale Cr isotope fractionation preserved in marine sediments appears to be restricted to Neoproterozoic and younger marine sediments, implying that relatively high oxygen levels, of more than ~10−3 PAL may be needed to record fractionations on the order of several per mil. in marine sediments12,14. On the other hand, concentrations as low as 10−5–10−4 PAL appear sufficient to induce large Cr isotope fractionations in paleosols and these continental signals can lead to mildly positive δ53Cr values in contemporaneous BIFs11. As an absolute lower limit, thermodynamic considerations suggest (Supplementary Figs 8,9) that Cr oxidation may be theoretically possible at O 2 concentrations as low as 10−20 PAL at circumneutral pH and that even if Cr oxidation were dependent on catalysis by Mn, it would have been possible at O 2 concentrations below 10−5 PAL with pH above 4. The kinetics of Cr dissolution and oxidation including its microbial catalysis, however, need also be considered and some models predict that extensive Cr oxidation and isotope fractionation take place at O 2 concentrations of less than 0.1% PAL36.

The precise threshold O 2 concentrations for the induction of Cr isotope fractionation remain uncertain, but we argue here that our data are consistent with the very low levels of oxygen or other ROS indicated by other proxies. If we estimate that weathering could generate a ~3% fractionation in Cr run-off28, then mass budget calculations using the average δ53Cr of 0.05% defined by the Isua BIFs would imply that less than only ~2% of the total Cr in the BIF would need to carry this signal. If 3.8–3.7 Ga rivers had total Cr concentrations on the order of 3 nmol l−1, comparable to Cr(III) concentrations in modern rivers26,37, only about 60 pmol l−1 of Cr would need to have seen a redox cycle. At these low concentrations, we emphasize that the redox active Cr pool need not necessarily be transported to the ocean as Cr(VI). In contrast, assuming that rivers during the early Archean were likely reducing in nature, such small amounts of redox active Cr could have been delivered to the oceans as soluble Cr(III) instead. Importantly, any trace of Cr that cycled through redox reactions on land would tend both to be heavy and to mobilize into the contemporaneous run-off more readily than Cr weathered directly as Cr(III). Having reached the oceans, this fractionated Cr would have been stripped from seawater by Fe (oxy)hydroxides formed during the deposition of BIFs from low oxygen oceans.

Czaja and co-authors23 argued that iron deposition to form Isua BIF was likely the product of photoferrotrophic iron oxidation through anoxygenic photosynthesis rather than by oxygen produced through oxygenic photosynthesis. They further argued that if oxygenic photosynthesis had evolved and was active at this time, that cyanobacteria had not proliferated and that the oxygen produced must have been rapidly consumed leading to little Fe(III)-deposition compared to times later in the Archean. Similar arguments would apply to molecular oxygen produced directly through CO 2 photodissociation38.

Hydrogen peroxide (H 2 O 2 ) is another ROS, which could have led to iron oxidation in the Eoarchean. It can be generated through several processes and has been measured in a variety of different environments39 (and references therein), but Pecoits and co-workers39 consider H 2 O 2 formed through atmospheric photochemical reactions as the only process that can generate significant fluxes. The formation of H 2 O 2 in the ancient atmosphere is linked to initial photolysis of water vapour, or to the reaction of water vapour with an excited atomic oxygen39. Both OH and HO 2 molecules, which are produced during these reactions, are precursors to H 2 O 2 formation and H 2 O 2 may then be removed from the atmosphere by photolysis, reaction with OH, or by rainout40. Using conservative Fe(III) sedimentation rates predicted for submarine hydrothermal settings in the Eoarchean, Pecoits and co-workers39 argue that the flux of H 2 O 2 produced in an Eoarchean atmosphere was likewise insufficient by several orders of magnitude to account for iron formation deposition.