The atmosphere of the Archean eon—one-third of Earth’s history—is important for understanding the evolution of our planet and Earth-like exoplanets. New geological proxies combined with models constrain atmospheric composition. They imply surface O 2 levels <10 −6 times present, N 2 levels that were similar to today or possibly a few times lower, and CO 2 and CH 4 levels ranging ~10 to 2500 and 10 2 to 10 4 times modern amounts, respectively. The greenhouse gas concentrations were sufficient to offset a fainter Sun. Climate moderation by the carbon cycle suggests average surface temperatures between 0° and 40°C, consistent with occasional glaciations. Isotopic mass fractionation of atmospheric xenon through the Archean until atmospheric oxygenation is best explained by drag of xenon ions by hydrogen escaping rapidly into space. These data imply that substantial loss of hydrogen oxidized the Earth. Despite these advances, detailed understanding of the coevolving solid Earth, biosphere, and atmosphere remains elusive, however.

We will not attempt to resolve controversies over how much solid Earth evolution drove atmospheric evolution. Consequently, we omit the large topic of how Earth’s outgassing history depended on debated tectonic and geological evolution models.

In addition to a gradual increase in solar luminosity, slow changes in the solid Earth over time provided boundary conditions for atmospheric evolution. On geological time scales, volcanic and metamorphic gases replenish atmospheric volatiles that escape to space or are chemically sequestered into solid materials.

Atmospheric composition, in turn, affected Archean climate. At 4 Ga ago, solar luminosity was 25 to 30% lower than today ( 12 , 13 ), but Archean Earth was not persistently frozen because abundant evidence shows an active hydrological cycle. Liquid water under a fainter Sun likely implies more abundant greenhouse gases than today ( 14 , 15 ). We review what the gases were and their levels.

Oxygenic photosynthesis produced the most impactful waste gas. The O 2 from cyanobacterial ancestors flooded the atmosphere rapidly at a time between 2.4 and 2.3 Ga, with the transition marked in the rocks by the sudden disappearance of mass-independent fractionation (MIF) of sulfur isotopes (discussed in detail later) ( 7 , 8 ). The Great Oxidation Event (GOE) thus began, which ended ~2.1 to 2.0 Ga ago ( 9 , 10 ). Although this switch to an oxygenated atmosphere and shallow ocean occurred in the Paleoproterozoic era (2.5 to 1.6 Ga ago), the weakly reducing atmosphere that was eliminated typified the Archean [e.g., ( 11 )]. Here, “weakly reducing” means minor levels of reducing gases, such as CO, H 2 , and CH 4 , in an anoxic atmosphere of bulk oxidized gases, CO 2 and N 2 . A major outstanding question concerns how trends of biological and geological evolution relate to the GOE.

Our discussion also considers evidence for early life and its possible global influence. We assume that metabolically useful gases would have been consumed, while waste gases would have been excreted, as they are today.

Data about the Archean atmosphere come from how individual gases, or the air as a whole, affected chemical and physical phenomena (e.g., the composition of aerosols, chemical reactions in soils, raindrop terminal velocity, isotopic fractionations, etc.) that were recorded in rocks. So, after a brief discussion of the Hadean, we review what the Archean atmosphere was made of. However, because of limited proxy data, major uncertainties remain about the exact levels of atmospheric gases over time.

The Archean was originally conceived to span the time from after the origin of life to the advent of free O 2 ( 5 ). While the origin of life dates back to before 3.5 to 3.8 Ga ago or earlier [e.g., ( 6 )], newer information puts atmospheric oxygenation after ~2.4 Ga ago, inside the Proterozoic. Here, considering the Archean in the older sense, the origin of life falls outside this Review, while the story of oxygen’s rise falls within.

The earliest well-preserved sedimentary and volcanic rocks are Archean and provide insights into atmospheric composition, climate, and life. These perspectives are unavailable for the Hadean eon from ~4.6 to 4 Ga ago, which generally lacks these rocks. For context, the Archean precedes the Proterozoic eon of 2.5 Ga to 541 ± 1 million years (Ma) ago, and Archean eras provide a timeline for our discussion: the Eoarchean (4 to 3.6 Ga ago), Paleoarchean (3.6 to 3.2 Ga ago), Mesoarchean (3.2 to 2.8 Ga ago), and Neoarchean (2.8 to 2.5 Ga ago).

The environment of the Archean eon from 4 to 2.5 billion years (Ga) ago has to be understood to appreciate biological, geological, and atmospheric evolution on our planet and Earth-like exoplanets ( Fig. 1 ) [e.g., ( 1 , 2 )]. Its most distinguishing characteristic was negligible O 2 , unlike today’s air, which contains, by dry volume, 21% O 2 , 78% N 2 , 0.9% Ar, and 0.1% other gases. With its radically different atmosphere and lack of macroscopic, multicellular life, the Archean world was alien. However, at that time, the beginnings of modern Earth emerged. For example, cyanobacteria probably evolved [e.g., ( 3 )] during this period, and these oxygenic photoautotrophs eventually oxygenated the air, setting the stage for later, complex life, including us ( 4 ).

In summary, at the end of the Hadean, Earth had oceans, continents ( 16 ), an anoxic atmosphere likely rich in CO 2 and N 2 , and probably life ( Fig. 1 ).

Estimates of nitrogen bound in today’s solid Earth range a few to ~40 bar equivalent ( 37 , 38 ), which allows for considerable N 2 in the Hadean atmosphere unless N was incorporated into a reducing, deep layer of magma—a magma ocean—formed after the Moon-forming impact ~4.5-Ga ago [( 39 ) and references therein]. So, although N 2 was one of the bulk atmospheric gases, its Hadean level—higher than today or lower—remains unclear.

Because a liquid ocean likely existed by ~4.4 Ga ago, feedbacks in the geologic carbon cycle (discussed later) probably stabilized the long-term climate ( 34 ). However, consumption of CO 2 in the weathering of impact ejecta by carbonic acid suggests a cool early Hadean surface near 0°C under the faint Sun ( 35 , 36 ).

The mantle contains excess highly siderophile elements (HSEs) relative to concentrations expected after Earth’s iron core formed, which removed HSEs. Similarities of isotopes and relative proportions of these HSEs to those in enstatite chondrite and achondrite meteorites suggest that this highly reducing meteoritic material was delivered late in Earth’s accretion ( 28 ). Thus, carbon and nitrogen were supplied in graphite and nitrides. Therefore, the Hadean atmosphere and mantle were probably initially highly reducing, before subsequent oxidation either by hydrogen escape ( 29 ) or disproportionation of mantle FeO accompanied by Fe loss to the core ( 30 , 31 ). In any case, iron-cored impactors would reduce seawater to hydrogen and create transient, highly reducing atmospheres that may have been important for the origin of life ( 32 , 33 ).

Impact bombardment would have affected Hadean and subsequent Archean environments. The lunar record implies a decay of terrestrial impact bombardment extending into the Archean ( 20 , 21 ). The estimated median age of the last impact big enough to vaporize the entire ocean is ~4.3 Ga ago ( 22 ), which provides a crude upper age limit on the origin of life. An origin of life during the period ~4.3 to 4.0 Ga ago is consistent with phylogenetic inferences [e.g., ( 23 )]. Later, the Late Heavy Bombardment (LHB) is a hypothesized interval of enhanced bombardment superposed on the general decline, which occurred between 4.2 and 4.0 to ~3.5 Ga ago, based on the ages of lunar rocks and meteorite shocks ( 24 ), although many dispute that the LHB was a discrete event ( 25 ). Regardless, an LHB would likely not sterilize Earth ( 26 ); any microbial life would have rebounded ( 27 ).

The composition of the Hadean atmosphere is obscured by a lack of well-preserved rocks, but analysis of zircons—crystals of zirconium silicate (ZrSiO 4 )—suggests that continents, oceans, and perhaps life all originated in the Hadean ( 16 ). Zircons are tiny (<0.5 mm) durable pieces of continental crust. Elevated 18 O/ 16 O in 4.3-Ga-old zircons was possibly inherited from 18 O-enriched, weathered surface rocks that were later buried and melted, which implies the presence of surficial liquid water and even land ( 17 , 18 ). In addition, graphite inside a 4.1-Ga-old zircon has a biogenic-like δ 13 C of −24 per mil (‰) ( 19 ), although lack of context means that an abiotic origin cannot be eliminated. [Here, δ 13 C is the 13 C/ 12 C ratio of a sample in parts per thousand (‰) relative to a standard reference material: δ 13 C = 1000 × [( 13 C/ 12 C) sample /( 13 C/ 12 C) standard − 1]].

WHAT WAS THE ARCHEAN ATMOSPHERE MADE OF?

Proxies constrain Archean atmospheric composition. Gases reacted with the seafloor or land, leaving chemical traces in seafloor minerals (40, 41) or in soils that became paleosols (42, 43). In addition, atmospheric particles carried isotopic signatures into sediments that were diagnostic of atmospheric composition (44–47). Occasionally, fluid inclusions in rocks trapped seawater with dissolved air (48–50) or even microbial gases (51).

Sometimes the physical environment affected rocks and minerals. Their preservation allows estimates of environmental temperature (52–55) and barometric pressure (56, 57).

Table 1 summarizes inferences about the composition of the Archean atmosphere and ocean. Here, gas concentrations are generally for the base of the troposphere. Although the Archean lacked a stratospheric ozone layer, it remains valid to refer to a troposphere and stratosphere as vertical regions where, respectively, convection and radiation dominated the energy transfer.

Table 1 Archean environmental constraints. Constraints on atmospheric gases at ground level (unless stated otherwise) and some bulk marine species. Gas level constraints are given in the same units as in the cited papers: partial pressure in bar or atm, where 1 bar = 0.9869 atm, or as mixing ratios (ppmv = parts per million by volume; S-MIF = sulfur isotope mass-independent fractionation). View this table:

Negligible Archean O 2 but oxygen oases after oxygenic photosynthesis evolved The strongest constraint on Archean atmospheric composition is that the ground-level mixing ratio of O 2 was <10−6 PAL (present atmospheric level) or <0.2–parts per million by volume (ppmv) O 2 for air of 1 bar, indicated by the presence of sulfur isotope MIF (S-MIF) in Archean sedimentary minerals (Fig. 2A) (44, 47). An oft-quoted limit of 10−5 PAL O 2 derives from an earlier photochemical model that could not address O 2 levels in the range of 10−5 to 10−15 PAL (45). Usually, isotope fractionation is proportional to the mass difference between isotopes; e.g., in diffusive separation of sulfur-containing gases, 34S becomes about half as abundant, relative to 32S, as 33S. However, some particular photochemical reactions produce MIF that, by definition, deviates from proportionality to mass. Fig. 2 Schematic histories of atmospheric O 2 and surface barometric pressure or N 2 . (A) Colored arrows faithfully represent known O 2 constraints, but the black line is speculative. An Archean upper bound of <0.2-μbar O 2 (blue) is for photochemistry that generates S 8 aerosols, preserving observed mass-independent isotope fractionation in sulfur compounds (44). The size and shape of an O 2 overshoot during the GOE are highly uncertain; a lower bound (red arrow) comes from iodine incorporation into carbonates (251). In the Proterozoic, a lower bound (light green) of 6 × 10−4 bar is required for an O 2 -rich atmosphere to be photochemically stable (44). However, O 2 levels likely remained low for most of the Proterozoic (252). Neoproterozoic oxygenation began around ~800 Ma ago. From ~600 Ma ago, a lower bound of >0.02-bar O 2 (dark green) is from plausible O 2 demands of macroscopic Ediacaran and Cambrian biota (120). Charcoal since 0.4 Ga ago implies a lower bound of >0.15 bar (purple) (253). The post-Devonian black line for O 2 evolution approximately represents curves from calculations of C and S isotopic mass balance (254, 255). (B) Constraints on surface atmospheric pressure (red) (56, 57) and the partial pressure of nitrogen, pN 2 (blue) (48, 49, 140). Blue shading shows a schematic and speculative pN 2 range in different time intervals consistent with very sparse proxy data. Archean S-MIF is tied to the production of elemental sulfur, S 8 , from photochemistry in anoxic air; in contrast, in an oxic atmosphere, S-MIF nearly disappears. Photochemistry imparts S-MIF and produces elemental sulfur, starting with reactions that photolyze volcanic SO 2 such as SO 2 + hν (<217 nm) = SO + O and SO + hν (< 231 nm) = S + O (58, 59). This photolysis occurs when short-wavelength ultraviolet (UV) penetrates Earth’s troposphere in the absence of a stratospheric ozone (O 3 ) layer. Scarce ozone implies negligible O 2 from which O 3 derives. In anoxic air, sulfur ends up in insoluble S 8 aerosols and water-soluble sulfate and SO 2 , unlike today’s atmospheric sulfur, which almost entirely oxidizes to sulfate (44, 45). In the Archean, as well as anoxia, gases such as CH 4 or H 2 produced sufficiently reducing conditions that S, S 2 , S 3 , etc. gases persisted and polymerized into S 8 (44, 60). Reactions that dominantly imparted S-MIF are debated: They include polysulfur formation (61), SO 2 photolysis, and other reactions (46, 59). In any case, when S 8 particles fell to Earth’s surface, their S-MIF isotopic composition complemented that of sulfate particles, allowing preservation of different S-MIF signs in these phases as sedimentary pyrite (FeS 2 ) and barite (BaSO 4 ) (62). Sulfate, of course, could later be microbially transformed to pyrite. Numerous redox-sensitive tracers corroborate negligible Archean O 2 [reviewed, e.g., in (10, 63, 64)]. Anoxia proxies that have long been recognized include the lack of pre-GOE continental sediments stained by red ferric oxides (redbeds), detrital grains from well-aerated rivers of siderite (FeCO 3 ), uraninite (UO 2 ), or pyrite that would oxidize and dissolve or rust at high pO 2 (65, 66), and paleosols with iron washed out by anoxic rainwater [e.g., (67)]. Furthermore, Archean marine sediments have low concentrations of elements that enter rivers during oxidative continental weathering (68–70). Conversely, glacial sediments contain continental materials lacking oxidative weathering loss of molybdenum (71). Iron formations (IFs), which are marine chemical sedimentary rocks rich in iron and silica [15 to 40 weight % (wt %) Fe and 40 to 60 wt % SiO 2 ], indicate that the deep Archean ocean contained Fe2+(aq) and so was anoxic [e.g., (72)]. In “Superior-type IFs” that formed near-shore, rare earth elements show that dissolved iron was partly sourced from seafloor hydrothermal vents and upwelled onto continental shelves where the iron precipitated. In today’s deeply oxygenated oceans, oxidized iron instead precipitates locally around vents. Archean shallow-water IFs constrain atmospheric O 2 to <2.4 × 10−4 bar (73). The absence or presence of mass-dependent fractionation of various isotopes can also indicate anoxic versus oxic conditions. In the case of sulfur, today’s oxidative weathering of continental sulfides produces soluble sulfate, which rivers carry to the ocean. When bacteria reduce sulfate to pyrite in seafloor sediments, they impart mass-dependent S isotope fractionation if sulfate is present at sufficient concentrations as in the modern oceans, but usually not in Archean seawater (74, 75). The absence of this isotope fractionation indicates little Archean seawater sulfate and implies anoxic air. Using similar arguments of O 2 -sensitive weathering, transport, and fractionation, isotopes of Cu (76), Cr (77), Fe (78), U (79), Mo [e.g., (80, 81)], and Se (82) indicates an anoxic Archean atmosphere. Atmospheric oxidation of 2.7-Ga-old iron-nickel micrometeorites has been used to argue for O 2 near-modern levels above ~75-km altitude (83, 84). However, given copious evidence for an anoxic Archean atmosphere, an alternative explanation is that high CO 2 levels (perhaps >70%) oxidized the micrometeorites (85). Nonetheless, even under a globally anoxic atmosphere, lakes and shallow seawater inhabited by oxygenic photosynthesizers could have become “oxygen oases”—local or regional areas with elevated O 2 . Modern surface seawater dissolves 0.25 mM O 2 at 15°C, while estimates for Archean oxygen oases range from 0.001 to 0.017 mM (0.4 to 7% of present) (86, 87). When exactly oxygenic photosynthesis began and dominated over anoxygenic photosynthesis is debated, but signs of biological carbon fixation appear early. Graphite in a ~3.7-Ga-old outcrop of sedimentary rock is 12C-enriched in amounts typical of photosynthetic microbes (88, 89). Then, at 3.52 Ga ago, δ13C in kerogen of −24‰ and associated marine carbonate of −2‰ found in Australia is similar to biological isotope fractionation in modern oceans (90). Fossil evidence for Archean cyanobacteria is reported. Light organic carbon isotopes and structures like those made by filamentous cyanobacteria found within stromatolites or other microbially induced sedimentary structures are consistent with cyanobacteria by 3.2 to 2.7 Ga ago (91–94). Cyanobacteria could be corroborated by biomarkers, which are remnant organic molecules from particular organisms. However, putative Archean biomarkers have been plagued by younger contamination [e.g., (95)]. Instead, cyanobacterial interpretations are strengthened by geochemical data suggesting O 2 oases at 3.2 to 3.0 Ga ago. Fractionated iron and molybdenum isotopes and levels of redox-sensitive metals suggest marine photic zone O 2 (96–99). [Chromium isotope data have been used to argue for the existence of ~3 Ga-old terrestrial oxygen (100), but they were probably caused by modern oxidative weathering (101).] Then, by 2.8 to 2.6 Ga ago, increasing concentrations and isotopic fractionation of Mo and S in marine shales suggest that O 2 proximal to cyanobacterial mats and stromatolites on land oxidized sulfides and boosted sulfate and Mo riverine fluxes to the oceans (70, 80, 102–104). These trends are consistent with isotopic evidence for Neoarchean methanotrophy and oxidative nitrogen cycling (105–109). Concentration spikes at 2.5 to 2.66 Ga ago in Mo, Se, and Re and isotopic excursions of Mo, Se, U, and N have been interpreted as arising from O 2 transients or “whiffs” of O 2 (110–113). Critics argue that the data derive from post-GOE alteration (114, 115) [but see (116)]. Alternatively, oxygen oases of cyanobacteria within soils or lakebeds may have mobilized these elements into rivers and then the sea (117). Phylogenetic analyses mostly suggest that cyanobacteria originated by 2.8 Ga. Molecular clocks must be calibrated by physical evidence, and phylogenetic methods are themselves debated. While some argue that oxygenic photosynthesis evolved only 50 to 100 Ma before the GOE (115), most studies suggest an earlier Paleoarchean or Mesoarchean age [e.g., (3, 118, 119)]. A major unresolved issue is how the GOE was related to underlying geological or biological trends (Fig. 2A). Life on its own cannot change the net redox state of the global environment because each biological oxidant is complemented by a reductant. In oxygenic photosynthesis, a mole of organic carbon accompanies every mole of O 2 CO 2 + H 2 O ⇄ respiration , decay photosynthesis CH 2 O + O 2 (1) Today, with ~4.4 × 105–Tmol surface organic matter and ~9000 Tmol/year of oxidative decay or respiration (120), reaction 1 reverses in ~50 years. Consequently, for long-term O 2 accumulation, some organic carbon must be segregated from the O 2 and buried. Alternatively, O 2 molecules are liberated if microbes use the organic carbon from reaction 1 to make other reductants, such as sulfides from sulfates, that are buried. However, Archean seawater sulfate concentrations were small (Table 1), so organic carbon burial is the O 2 flux that matters for the GOE. Atmospheric oxidation also occurs when hydrogen escapes to space after the photochemical breakdown of gases such as H 2 and CH 4 , which ultimately derive from water and are relatively abundant in anoxic air. Atmospheric O 2 is determined by net redox fluxes into and out of the atmosphere. This simple truth is deceptive because these redox fluxes are themselves controlled by less easily constrained oxidative weathering (both seafloor and continental) and volcanic and metamorphic degassing, as well as hydrogen escape to space. Permanent atmospheric oxygenation requires ~3 × 10−3 PAL O 2 or more to prevent a destabilizing positive feedback of photochemical destruction of tropospheric O 2 that otherwise occurs when an incipient ozone column is still transparent to far UV (44, 121, 122). These O 2 levels would have only been attained when the O 2 flux from the burial of organic carbon exceeded the kinetically efficient sink from O 2 -consuming gases (CO, H 2 , H 2 S, and SO 2 ) from volcanism and metamorphism plus fluxes of reducing cations such as Fe2+ from seafloor vents (9, 120, 121, 123, 124). A minority assume that such a flux imbalance applied as soon as oxygenic photosynthesis evolved, mandating a rapid rise of O 2 (125). In the consensus assessment that oxygenic photosynthesis evolved long before the GOE, efficient consumption of O 2 initially suppressed O 2 levels. Hypotheses about the GOE tipping point are reviewed elsewhere [(64), chap. 10]. Briefly, ideas favoring increased O 2 fluxes from organic burial appeal to more continental shelf area available for burial (126), more phosphorus to stimulate photosynthesis (127, 128), or subduction of organic carbon relative to ferric iron (129). Because organic carbon burial extracts 12C and leaves inorganic carbonates 12C depleted, it is difficult to reconcile these hypotheses with the remarkable constancy of the carbon isotope record, which indicates little change in average organic burial rates between the Archean and Proterozoic (130). However, an increase in organic burial might have occurred if negligible oxidative weathering of 12C-rich organics on land or a sink of seafloor carbonate meant that the operation of the carbon cycle or the isotopic composition of carbon input into the surface environment differed from today (130–132). Hypotheses for a slowly decreasing O 2 sink to the GOE tipping point rely on a decline in the ratio of reduced-to-oxidized species from volcanic, metamorphic, and hydrothermal sources (9, 123, 124, 133, 134). Some emphasize the role of hydrogen escape to space in oxidizing solid Earth, lowering Earth’s capacity to release O 2 -consuming reductants (121, 135, 136). New evidence from xenon isotopes supports rapid Archean hydrogen escape, as discussed below.

Nitrogen gases in the Archean Molecular nitrogen dominates today’s atmosphere, and three lines of evidence have begun to constrain Archean N 2 levels (Fig. 2B). First, the largest size of 2.7-Ga-old fossil raindrop imprints provides a conservative limit of paleopressure of <2.1 bar and a probable limit of <0.52 to 1.2 bar (57). [Criticism of the raindrop paleopressure constraint (137) is refuted in (138).] Second, the N 2 /36Ar ratio in fluid inclusions indicates pN 2 <1.0 bar [2σ] at 3.3 Ga ago and <1.1 bar at 3.5 to 3.0 Ga ago (48, 49). Third, vesicle volumes in 2.7-Ga-old basaltic lava flows erupted at sea level imply a 0.23 ± 0.23–bar [2σ] paleopressure (56). The inferred history of pN 2 depends on how the geologic nitrogen cycle has changed over time. Nitrogen can only greatly accumulate as atmospheric N 2 or in rocks as ammonium, amide, nitride, or organic nitrogen. Under typical mantle temperatures and redox conditions, volcanic gases contain N 2 , not ammonia [e.g., (139), p. 49], and because N 2 is unreactive, it enters the air. Today, N 2 is also produced when oxidative weathering of organic matter on the continents makes nitrate (NO 3 −) that undergoes rapid biological denitrification into N 2 (140). Within large uncertainties, volcanic and oxidative weathering inputs of N 2 are comparable [(64), p. 204]. Using sedimentary C/N data, Berner (140) argues that the sum of these N 2 sources was balanced over the Phanerozoic primarily by N burial in organic matter, so that the Phanerozoic partial pressure of nitrogen, pN 2 , varied little. Som et al. (56) proposed that low Archean paleopressure arose because today’s long-term N 2 atmospheric input from oxidative weathering and denitrification was absent. If so, pN 2 would have risen at the GOE, and nitrogen in today’s air must have been in solid phases previously. Certainly on modern Earth, nitrification, by humankind’s addition of nitrate to land and sea, has enhanced denitrification, indicated by increased atmospheric nitrous oxide (N 2 O) [e.g., (141), chap. 6]. On the other hand, a model can be constructed where pN 2 declines after the GOE if burial of organic nitrogen increased (142). In the Hadean, pN 2 either started high and then diminished (143) or was initially low if nitrogen partitioned into a very reducing magma ocean (39). However, low N/C in today’s mid-ocean ridge source basalts (144) suggests considerable N 2 degassing once the upper mantle became oxidized because then nitrogen became insoluble in magmas and upper mantle fluids (145). Even today, some of the upper mantle lies within the stability field for ammonium, so that increased oxidation of the early mantle and mantle wedge could have caused more subducted nitrogen to outgas as N 2 (145). Marine phyllosilicates at 3.8 Ga ago are ammonium enriched (146, 147), which probably came from porewater ammonium ( NH 4 + ) derived from degraded organics (Table 1) (148, 149), and these data have been used to argue that Archean N 2 was sequestered into solid phases after an early advent of biological nitrogen fixation (56, 150). In the early ocean, NH 4 + would have been the stable form of dissolved nitrogen unlike today’s nitrate. Consequently, a postulated drawdown of Archean N 2 involves biological fixation, organic burial, and subduction of ammonium in refractory minerals. The rate of organic burial must have been relatively high for a time-integrated loss to affect pN 2 significantly (150), which is not necessarily inconsistent with carbon isotopic constraints because early high degassing of carbon required more carbon to be buried (130). Nitrogen isotopes appear to confirm that biologically fixed nitrogen entered the Archean mantle. Sedimentary organics have δ15N = 7 ± 1‰ compared to 0‰ in N 2 in modern and ancient air and − 5 ± 2‰ in the mantle (151, 152). Fractionation mostly arises when denitrification preferentially converts nitrate or nitrite 14N into N 2 . Thus, heavy δ15N in 3.1 to 3.5 Ga-old mantle-derived diamonds may be a sedimentary component (153). Ammonium substitutes for potassium, and breakdown of previously subducted ammonium-containing minerals in magmas at oceanic islands releases N 2 and radiogenic 40Ar derived from 40K. The ratio 40Ar/N 2 in plume-related lavas scatters by a factor of ~4 to 5, and higher values (older from more 40Ar) correlate with smaller, Archean-like values of δ15N, consistent with a history of ammonium subduction. Because N 2 is uncorrelated with nonradiogenic 38Ar or 36Ar, nitrogen in the current mantle is not primordial but recycled (154). Some use N ingassing versus outgassing fluxes to infer past pN 2 . Mallik et al. (155) estimate modern subduction of 6.4 ± 1.4 × 1010 mol N year−1 and, with a global degassing flux of 2 × 1010 mol N year−1 from (156), argue that net N ingassing today means higher past pN 2 . However, other outgassing estimates are 7 × 1010 mol N year−1 from arcs (157) or 9 ± 4 × 1010 mol N year−1 globally [(64), p. 204], such that current source and sink N fluxes balance within uncertainties. Unlike N 2 , other nitrogen-bearing Archean gases would have been trace quantities. With only tiny atmospheric fluxes of nitrate or nitrite, N 2 O from denitrification would have been negligible, perhaps restricted to lakes (158). Lightning production of NO, followed by H addition and HNO dissolution and decomposition, might maintain ground-level N 2 O to a few parts per billion by volume (ppbv), compared to a preindustrial ~270-ppbv N 2 O (159). The other nitrogen oxides (NO and NO 2 ) would have been at trace levels because their lightning production in CO 2 -N 2 air is inefficient (160). Ammonia (NH 3 ) levels of 10 to 100 ppmv would provide a greenhouse effect to counteract the faint young Sun (FYS) (14), but these levels of NH 3 cannot be sustained against UV photolysis (161). Possibly, a stratospheric organic haze, such as that on Saturn’s moon, Titan, shielded tropospheric NH 3 from UV (162). However, whether this shielding actually occurred depends on whether hazes actually existed and on the size and radiative properties of haze particles, which remain uncertain (1, 163). Another nitrogen-bearing gas, hydrogen cyanide (HCN) is more stable photochemically than NH 3 and made in reducing atmospheres by lightning, impacts (164, 165), or UV-driven photochemistry. In particular, N atoms from N 2 photolysis in the upper atmosphere can mix to lower levels and react with CH 4 photolysis products to make HCN (166, 167). At Archean biogenic methane levels of ~103 ppmv, HCN concentrations reach ~102 ppmv.

Carbon gases in the Archean: CO 2 , CH 4 , and CO Let us now consider carbon gases, starting with carbon dioxide. Since the early Hadean, CO 2 has probably always been Earth’s most important noncondensable greenhouse gas. CO 2 also affects seawater pH and influences the carbon cycle through the formation of carbonates and organic matter. However, direct evidence for Archean CO 2 levels remains scanty. Paleosols provide some estimates. Acid leaching in Archean soils arose from CO 2 dissolved in rainwater. Mass-balance calculations give 10 to 50 PAL of CO 2 at 2.7 Ga ago and 23 ÷ 3 × 3 PAL of CO 2 at 2.2 Ga ago (42, 168). However, these analyses assume that all the CO 2 that entered the soils caused dissolution, so pCO 2 could have been higher if only a fraction of CO 2 had been used. Another study used an analysis with an estimate of the composition of temperature-dependent aqueous solutions during weathering and, because of a weaker dependence of weathering on pCO 2 , obtained higher pCO 2 of 85 to 510 PAL at 2.77 Ga ago, 78 to 2500 PAL at 2.75 Ga ago, and 160 to 490 PAL at 2.46 Ga ago (Fig. 3A) (43). Fig. 3 History of CO 2 and a CH 4 schematic since the Archean. (A) The black line is median CO 2 from a carbonate-silicate climate model, and yellow shading indicates its 95% confidence interval (34); this curve merges with a fit to CO 2 proxy estimates for 0.42 Ga ago to present from (256). Various Precambrian pCO 2 proxy estimates are shown (42, 43, 168, 257–259). (B) A very schematic history of CH 4 . Constraints include a lower limit (blue) required for Archean S-MIF (44) and a tentative lower limit of ~3.5 Ga ago from a preliminary interpretation of xenon isotopes (black) (187). The black curve is from a biogeochemical box model coupled to photochemistry (121). Orange shading is schematic but consistent with possible biological CH 4 fluxes into atmospheres of rising O 2 levels at the GOE and in the Neoproterozoic. Note that the suggestion that moderately high levels of methane may have contributed to greenhouse warming in the Proterozoic (260, 261) has been disputed (262, 263) and may depend on fluxes from sources on land (264). The curve for ~0.4 Ga ago to present is from (265). High Archean pCO 2 does not have to induce acidic seawater and dissolve marine carbonates. Instead, an increase in Ca2+ concentrations could maintain an ocean saturated in calcium carbonate at alkaline pH; alternatively, seawater pH could be slightly lower than today, but calcium carbonate would remain saturated. In fact, sedimentary marine carbonates appear from 3.52 Ga ago onward [(98) and references therein]. Calcium isotopes might provide insight into coupled Archean seawater pH and pCO 2 . These two variables and carbonate alkalinity (Alk = [ HCO 3 − ] + 2[ CO 3 − ]) define a system where any two variables imply the third. In evaporating seawater, Ca isotopes could undergo Rayleigh distillation if [Ca2+] << Alk, but limestones from 2.6-Ga-old Campbellrand marine evaporites show no spread in δ44/40Ca, which might imply [Ca2+] >> Alk and pH of 6.4 to 7.4 for a likely range of Neoarchean pCO 2 (Table 1) (169). Siderite (FeCO 3 ) in Archean IFs has been proposed as a pCO 2 proxy if it precipitated in equilibrium with the atmosphere (170, 171). However, data suggest that this siderite was diagenetic (172, 173), so we omit this proxy from Fig. 3A and Table 1. Potentially, two negative feedbacks control long-term pCO 2 . First, the net consumption of CO 2 in acid weathering of continents or the seafloor ends up making carbonates (174) X SiO 3 + CO 2 + H 2 O ⇌ X CO 3 + SiO 2 + H 2 O (2)where X is a cation. Seafloor weathering occurs when water in the permeable abyssal plains dissolves basaltic minerals, releasing calcium ions and precipitating calcium carbonate in veins and pores (175). Second, “reverse weathering” (RW) reactions have been proposed (176). Schematically, in net, we represent continental (or seafloor) weathering plus RW, as follows X SiO 3 + CO 2 + H 2 O + Al ( OH ) 3 ⇌ X AlSiO 2 ( OH ) 5 + CO 2 (3) This reaction consumes aqueous SiO 2 and uses cations to make aluminosilicate clays instead of carbonates that consume carbon, so that CO 2 stays in the atmosphere. High dissolved seawater silica and pH are hypothesized to promote RW and create clay minerals. If atmospheric pCO 2 is high, pH is low, and temperature is warm, reaction 2 is favored, and pCO 2 falls, in negative feedback on pCO 2 and climate. However, if pCO 2 is low and pH is high, RW, reaction 3, may be an alternative to reaction 2 and a negative feedback on low pCO 2 . Estimates of today’s RW flux (mol Si/year) vary an order of magnitude (177, 178), the rate coefficient for RW reactions varies by many orders of magnitude (179, 180), and the solubility of authigenic phyllosilicates (needed to calculate RW) varies by over an order of magnitude. Consequently, a self-consistent, coupled carbon-silica cycle model since 4 Ga ago shows that RW can be important or unimportant for the Proterozoic climate depending on parameter choice, while RW in the Archean is muted because of probably lower land fraction and sedimentation rate (181). Considering these factors, estimates of Archean pCO 2 and seawater pH that we give henceforth are based on carbonate-silicate cycle models without RW. In the Archean, with greater seafloor production than today, seafloor weathering could have been comparable to continental weathering (34, 35, 182), and negative feedback (Eq. 2) likely maintained average Archean surface temperatures between 0° and 40°C with seawater pH 6.4 to 7.4 (34). The corresponding pCO 2 would have been 0.006 to 0.6 bar 4 Ga ago, assuming that ~104-ppmv CH 4 also contributed to the greenhouse effect. Anoxic Archean air could hold 1000 s of parts per million by volume of CH 4 if a microbial flux of CH 4 was comparable to today’s (135, 183, 184). Phylogenetically, methanogens date back to >3.5 Ga ago (23). In contrast, the modern oxygenated atmosphere destroys reducing gases rapidly, limiting tropospheric CH 4 and H 2 abundances to 1.8 and 0.55 ppmv, respectively. Evidence points to high levels of Archean CH 4 (Fig. 3B). First, signs of methanogens and methanotrophs from light carbon isotopes in Archean organics imply methane’s presence [e.g., (109)]. Second, Archean S-MIF requires >20-ppmv CH 4 to generate particulate sulfur, S 8 (Table 1). Third, the deuterium-to-hydrogen (D/H) ratio of 3.7 Ga-old seawater estimated from serpentine minerals is 2.5% lighter than today, which could be explained by rapid Archean escape of hydrogen and isotopic fractionation (185, 186). This hydrogen was likely derived from UV photolysis of CH 4 in the upper atmosphere (135). Later, we discuss how fractionation of xenon isotopes in the Archean suggests that, ~3.5-Ga ago, CH 4 levels were >0.5%, i.e., >5000 ppmv (187). Fourth, globally extensive glaciations during the GOE [e.g., (188)] provide circumstantial evidence for high Archean CH 4 . At the tipping point, air flips from anoxic to oxic in only ~104 years, causing a ~10°C temperature drop by oxidizing CH 4 . This chemical transition is far faster than the ~105- to 106-year response of the carbonate-silicate thermostat. Another carbon-containing gas, carbon monoxide (CO), was probably not abundant in the presence of an Archean microbial biosphere. The Last Universal Common Ancestor was likely capable of anaerobic CO consumption (189), which involves water as a substrate and catalysts such as iron sulfide that were probably widespread H 2 O + CO → CO 2 + H 2 (4) With this reaction, microbes would draw CO down to 100 to 102 ppmv (183). However, episodic CO levels at percent levels may have occurred when large impacts delivered cometary CO ice or organic matter that was oxidized (190).

A high-altitude Archean organic haze? Because of relatively abundant CH 4 , a high-altitude Archean organic haze might have formed, as mentioned earlier. The idea was first suggested in the 1980s (191). Empirically, if the CH 4 :CO 2 ratio exceeds ~0.1 in a UV-irradiated CO 2 -N 2 -CH 4 mixture, radicals from methane photolysis polymerize into organic particles (192). When and whether an organic haze formed are uncertain. A haze could have affected tropospheric sulfur gases by blocking UV photons. Consequently, the structure of S-MIF variations in Archean sedimentary minerals and their correlation with light, organic δ13C have been attributed to episodic hazes driven by variable atmospheric CH 4 :CO 2 ratios (60, 193–196). However, given—dare we say—only a hazy understanding of which species and reactions are important for S-MIF (see earlier), interpretations of episodic hazes are permissive rather than definitive.

Hydrogen abundance The Archean lower atmosphere is unlikely to have been H 2 -rich given the antiquity of methanogens (23, 51, 197), some of which convert H 2 into methane through a net reaction 4 H 2 + CO 2 → CH 4 + 2 H 2 O (5) Anoxygenic photosynthesis also consumes H 2 [e.g., (198)]. In models with methanogens and H 2 -based photosynthesizers, atmospheric H 2 mixing ratios depend on assumed H 2 outgassing and biological productivity but generally are ≤10−4 (183, 184). These levels preclude H 2 as an important Archean greenhouse gas. Detrital magnetite carried in rivers 3.0- to 2.7-Ga ago would have dissolved at high pH 2 via microbial reduction of Fe3+ to soluble Fe2+ using H 2 (199), providing an upper limit of pH 2 ≤ 10−2 bar (200).

Xenon isotopic constraints on oxygen, hydrogen, and methane levels, plus Earth’s oxidation Changes in atmospheric Xe isotopes through the Archean stop after the GOE (analogous to S-MIF) (49) and potentially tell us about O 2 , CH 4 , and H 2 levels (187), including in the otherwise hidden Hadean, as discussed below. The trend also relates to how much hydrogen escaped from Earth and hence Earth’s total oxidation over time. Xenon in fluid inclusions in Archean rocks becomes isotopically heavier through the Archean relative to an initial solar composition until the fractionation reaches that of modern air around 2.1 to 2.0 Ga ago (Fig. 4) (49). Xenon dragged out into space by escaping hydrogen during the Archean and Hadean best explains the progressive mass fractionation (49, 187). An alternative explanation from trapping of Xe+ in organic hazes (201) has the problem that weathering or microbial processing of buried organics would release xenon and modulate the Xe isotopes in post-Archean air, which is not observed (202). Fig. 4 Mass fractionation of nine atmospheric xenon isotopes over time relative to modern air per atomic mass unit showing relative enrichment in light isotopes in the past. Data from (49). The vertical axis shows the fractionation per atomic mass unit (amu) of atmospheric xenon relative to modern air. To compute this average fractionation across the nine isotopes, Avice and co-workers (49) normalized the isotopic compositions to 130Xe and to the isotopic composition of the modern atmosphere using the delta notation. For a Xe isotope of mass i, δiXe air = 1000 × ((iXe/130Xe) sample /(iXe/130Xe) air − 1). The slope of a straight line fit to the normalized data provides the average fractionation per atomic mass per unit and its uncertainty, i.e., plotted points. Inset: A diagram showing schematically how the slope of the fractionation of the nine isotopes changed over time relative to the initial solar composition, where the graph is normalized to atomic mass 130. Today, the nine atmospheric Xe isotopes [124 to 136 atomic mass units (amu)], have a huge fractionation of 4.2% per amu relative to solar or chondritic sources, whereas the six, lighter krypton isotopes (76 to 86 amu) are barely fractionated. Earth’s Xr/Kr ratio is also 4 to 20 times less than meteorites, implying selective Xe loss. Unlike lighter krypton, xenon is easily ionized by solar UV or charge exchange with H+ ions, so Xe+ can be dragged out to space by escaping H+ ions (187). Whereas Xe+ is unreactive with H, H 2 , or CO 2 (203), any Kr+ ions are neutralized via Kr+ + H 2 →KrH+ + H and KrH+ + e−→Kr + H, explaining the lack of Kr isotope fractionation. Ions are tethered to Earth’s magnetic field lines, but a “polar wind” of hydrogen ions escapes along open field lines at the poles, accounting for ~15% of all hydrogen escapes today. Xe+ ions could be dragged by a vigorous ancient polar wind. That requires copious hydrogen to be derived from relatively abundant CH 4 and/or H 2 in the lower atmosphere. UV decomposes CH 4 in an anoxic upper atmosphere, releasing hydrogen [e.g., (166)]. The hypothesized xenon escape works only in anoxic air, so, like S-MIF, Xe isotopes record the GOE (Fig. 4). Oxic air destroys H 2 and CH 4 , making their abundances too low to supply enough hydrogen to drag along xenon. In addition, O 2 would remove Xe ions in a resonant charge exchange reaction (203) Xe + + O 2 → Xe + O 2 + (6) Our preliminary model finds that Xe can escape when the total hydrogen mixing ratio exceeds ~1% for the solar extreme UV flux expected around ~3.5 Ga ago (187). The total hydrogen mixing ratio, f T (H 2 ), in equivalent H 2 molecules, is f T ( H 2 ) = 0.5 f H + f H 2 + f H 2 O + 2 f CH 4 + … (7) Thus, if nearly all Archean hydrogen was biologically converted into CH 4 (Eq. 5), we interpret the xenon escape constraint of >1% f T (H 2 ) to be >0.5% CH 4 , which is a rare empirical constraint on Archean CH 4 (Fig. 3B). The inferred escape of a strong reductant, hydrogen, would oxidize entire Earth. The total oxidation during the Archean is equivalent to oxygen from a tenth of more of an ocean (187).