Fig. 1. Schematic diagram showing the modern fluxes of O₂. The fluxes include metamorphic (F_metamorphic) and volcanic (F_volcanic) sources of reducing gases. F_source is the flux of burial of organic carbon, pyrite and ferrous iron that contribute equivalent oxygen into the atmosphere–ocean system. F_escape is the escape of hydrogen to interplanetary space and F_weathering is the flux of reductants uplifted on continental surfaces that react with O₂ dissolved in rainwater. Solid arrows indicate flow of reductants, while dashed arrows indicate an equivalent flux of oxygen.

Fig. 2. A schematic diagram showing reductant fluxes that govern the oxidation state of the atmosphere, ocean and lithosphere. Curved arrows between the upper and lower crust reservoirs indicate mixing due to tectonic activity. Summing fluxes into and out of the atmosphere–ocean box gives Eq. (3).

Fig. 3. Oxygen reservoirs and fluxes in the modern O₂ cycle. Primary production is from [114]. Fluxes of burial, weathering and reaction with volcanic and metamorphic gases are from [13].

Fig. 4. (a) Reducing volcanic gases (with significant H₂, CH₄ and NH₃) were introduced into the atmosphere before core formation, when the mantle was rich in metallic iron. (b) A weakly reducing mixture of volcanic gases has fed the atmosphere for most of Earth history, after the Earth differentiated into core, mantle and crust.

Fig. 5. Hydrogen in the Earth's prebiotic atmosphere. (a) Schematic of outgassing source and escape sink (b) how dynamic equilibrium sets the prebiotic H₂ concentration.

Fig. 6. Schematic showing the effect of organisms on the early atmosphere before the origin of oxygenic photosynthesis. For the sake of argument, this graph assumes methanogens evolved before anoxygenic photosynthesis, but this is uncertain.

Fig. 7. The history of O₂, where the thick dashed line shows a possible evolutionary path that satisfies biogeochemical data. Dotted horizontal lines show the duration of biogeochemical constraints, such as the occurrence of detrital siderite (FeCO₃) in ancient riverbeds. Downward-pointing arrows indicate upper bounds on the partial pressure of oxygen (pO₂), whereas upward-pointing arrows indicate lower bounds. Unlabelled solid horizontal lines indicate the occurrence of particular paleosols, with the length of each line showing the uncertainty in the age of each paleosol. Inferences of pO₂ from paleosols are taken from [58]. An upper bound on the level of pO₂ in the prebiotic atmosphere at c. 4.4 Ga (shortly after the Earth had differentiated into a core, mantle and crust) is based on photochemical calculations. MIF is “mass-independent isotope fractionation”, which in sulfur is caused by photochemistry in an O₂-poor atmosphere. The pO₂ level inferred from MIF observed in pre-2.4 Ga sulfur isotopes is based on the photochemical model results of [70]. Biological lower limits on pO₂ are based on the O₂ requirements of: (1) the marine sulfur-oxidizing bacterium, Beggiatoa [3]; (2) animals that appear after 0.59 Ga [115]; (3) charcoal production in the geologic record. A “bump” in the oxygen curve around 300 Ma, in the Carboniferous, is based on the interpretation of Phanerozoic carbon and sulfur isotope data by [116].

Fig. 8. The carbon isotope record of δ¹³C in marine carbonates. The horizontal line is the average of all values (neglecting those from banded iron formations (square symbols)) and the surrounding rectangular box extends to one standard deviation. The dashed line is a sketch indicating how the Paleoproterozoic and Neoproterozoic were times of unusually isotopically heavy carbonates. Isotopically light carbonates between 2.7–2.4 Ga are likely due to carbonates derived from respired organic matter, which is particularly true of banded iron formations [117]. The Precambrian carbon isotope data is from R. Buick (unpublished), modified from [118] by additional literature and improved radiometric dating. Phanerozoic carbon isotope data is from [87].

Fig. 9. A schematic diagram showing the dependence of the atmospheric redox state on changes in the reduction of sulfur, following Holland [13]. The dashed box indicates an average batch of volcanic and metamorphic gas. Hydrogen is ultimately used to reduce carbon gases to organic carbon and to reduce SO₂ to sedimentary sulfide. Geochemical data suggests that over the course of Earth history the fraction of carbon gas converted to organic matter has been about 1 / 5 (dashed arrows). The fraction of sulfur gas reduced to pyrite has been 2 / 3 in the Phanerozoic (dotted arrows). If enough H₂ had been present in Archean volatiles to convert all of the SO₂ to sulfide, excess H₂ would have accumulated in the Archean atmosphere making it anoxic. The consequence would be loss of hydrogen to space and oxidation of the lithosphere. As a result, volcanic and metamorphic H₂ fluxes would gradually dwindle. Eventually, not enough H₂ would be present to convert all the SO₂ to sulfide and instead SO₂ would be oxidized to create sulfate minerals (dash–dot arrows).

Fig. 10. The inventory of oxygen in the Earth's crust shows that there is excess oxygen. Data is from a tabulated compilation in [91].

Fig. 11. (a) A schematic diagram showing a plausible evolution of redox fluxes due to oxidative weathering, hydrogen escape, volcanic and metamorphic gases, and the burial of organic carbon (a net source of O₂). (b) A schematic diagram showing the evolution of atmospheric gases (CH₄, O₂) and oceanic SO₄²⁻from the late Archean to Proterozoic.

Glossary