Boring Billion

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https://en.wikipedia.org/wiki/Boring_Billion

Life timeline
−4500 — – — – −4000 — – — – −3500 — – — – −3000 — – — – −2500 — – — – −2000 — – — – −1500 — – — – −1000 — – — – −500 — – — – 0 —	Water   Single-celled life   Photosynthesis   Eukaryotes   Multicellular life   P l a n t s   Arthropods Molluscs Flowers Dinosaurs   Mammals Birds Primates H a d e a n A r c h e a n P r o t e r o z o i c P h a n e r o z o i c	← Earth formed ← Earliest water ← LUCA ← Earliest fossils ← LHB meteorites ← Earliest oxygen ← Pongola glaciation* ← Atmospheric oxygen ← Huronian glaciation* ← Sexual reproduction ← Earliest multicellular life ← Earliest fungi ← Earliest plants ← Earliest animals ← Cryogenian ice age* ← Ediacaran biota ← Cambrian explosion ← Andean glaciation* ← Earliest tetrapods ← Karoo ice age* ← Earliest apes / humans ← Quaternary ice age*
(million years ago) *Ice Ages

The Boring Billion, otherwise known as the Mid Proterozoic and Earth's Middle Ages, is an informal geological time period between 1.8 and 0.8 billion years ago (Ga) during the middle Proterozoic eon spanning from the Statherian to the Tonian periods, characterized by more or less tectonic stability, climatic stasis and slow biological evolution. Although it is bordered by two different oxygenation events (the Great Oxygenation Event and Neoproterozoic Oxygenation Event) and two global glacial events (the Huronian and Cryogenian glaciations), the Boring Billion period itself actually had very low oxygen levels and no geological evidence of glaciations.

The oceans during the Boring Billion may have been oxygen-poor, nutrient-poor and sulfidic (euxinia), populated by mainly anoxygenic purple bacteria, a type of bacteriochlorophyll-based photosynthetic bacteria which uses hydrogen sulfide (H₂S) for carbon fixation instead of water and produces sulfur as a byproduct instead of oxygen. This is known as a Canfield ocean, and such composition may have caused the oceans to be colored black- and milky-turquoise instead of blue or green as later . (By contrast, during the much earlier Purple Earth phase during the Archean, photosynthesis was performed mostly by archaeal colonies using retinal-based proton pumps that absorb green light, and the oceans would be magenta-purple.)

Despite such adverse conditions, eukaryotes may have evolved around the beginning of the Boring Billion, and adopted several novel adaptations, such as various organelles, multicellularity and possibly sexual reproduction, and diversified into algae, fungi and early animals at the end of this time interval. Such advances may have been important precursors to the evolution of large, complex life later in the Ediacaran Avalon Explosion and the subsequent Phanerozoic Cambrian Explosion. Nonetheless, prokaryotic cyanobacteria were the dominant autotrophic lifeforms during this time, and likely supported an energy-poor food-web with a small number of protists at the apex level. The land was likely inhabited by prokaryotic cyanobacteria and eukaryotic proto-lichens, the latter more successful here probably due to the greater availability of nutrients than in offshore ocean waters.

Description

In 1995, geologists Roger Buick, Davis Des Marais, and Andrew Knoll reviewed the apparent lack of major biological, geological, and climatic events during the Mesoproterozoic era 1.6 to 1 billion years ago (Ga), and, thus, described it as "the dullest time in Earth's history". The term "Boring Billion" was coined by paleontologist Martin Brasier to refer to the time between about 2 and 1 Ga, which was characterized by geochemical stasis and glacial stagnation. In 2013, geochemist Grant Young used the term "Barren Billion" to refer to a period of apparent glacial stagnation and lack of carbon isotope excursions from 1.8 to 0.8 Ga. In 2014, geologists Peter Cawood and Chris Hawkesworth called the time between 1.7 and 0.75 Ga "Earth's Middle Ages" due to a lack of evidence of tectonic movement.

The Boring Billion is now largely cited as spanning about 1.8 to 0.8 Ga, contained within the Proterozoic eon, mainly the Mesoproterozoic. The Boring Billion is characterized by geological, climatic, and by-and-large evolutionary stasis, with low nutrient abundance.

In the time leading up to the Boring Billion, Earth experienced the Great Oxygenation Event due to the evolution of oxygenic photosynthetic cyanobacteria, and the resultant Huronian glaciation (Snowball Earth), formation of the UV-blocking ozone layer, and oxidation of several metals. Oxygen levels during the Boring Billion are thought to have been markedly lower than during the Great Oxidation Event, perhaps 0.1% to 10% of modern levels. It was ended by the breakup of the supercontinent Rodinia during the Tonian (1000–720 Ma) period, a second oxygenation event, and another Snowball Earth in the Cryogenian period.

Tectonic stasis

Reconstruction of Columbia (1.6 Gya)

The evolution of Earth's biosphere, atmosphere, and hydrosphere has long been linked to the supercontinent cycle, where the continents aggregate and then drift apart. The Boring Billion saw the evolution of two supercontinents: Columbia (or Nuna) and Rodinia.

The supercontinent Columbia formed between 2.0 and 1.7 Ga and remained intact until at least 1.3 Ga. Geological and paleomagnetic evidence suggest that Columbia underwent only minor changes to form the supercontinent Rodinia from 1.1 to 0.9 Ga. Paleogeographic reconstructions suggest that the supercontinent assemblage was located in equatorial and temperate climate zones, and there is little or no evidence for continental fragments in polar regions.

Due to the lack of evidence of sediment build-up (on passive margins) which would occur as a result of rifting, the supercontinent probably did not break up, and rather was simply an assemblage of juxtaposed proto-continents and cratons. There is no evidence of rifting until the formation of Rodinia, 1.25 Ga in North Laurentia, and 1 Ga in East Baltica and South Siberia. Breakup did not occur until 0.75 Ga, marking the end of the Boring Billion. This tectonic stasis may have been related in ocean and atmospheric chemistry.

It is possible the asthenosphere—the molten layer of Earth's mantle that tectonic plates essentially float and move around upon—was too hot to sustain modern plate tectonics at this time. Instead of vigorous plate recycling at subduction zones, plates were linked together for billions of years until the mantle cooled off sufficiently. The onset of this component of plate tectonics may have been aided by the cooling and thickening of the crust that, once initiated, made plate subduction anomalously strong, occurring at the end of the Boring Billion.

Nonetheless, major magmatic events still occurred, such as the formation (via magma plume) of the 220,000 km² (85,000 sq mi) central Australian Musgrave Province from 1.22 to 1.12 Ga, and the 2,700,000 km² (1,000,000 sq mi) Canadian Mackenzie Large Igneous Province 1.27 Ga. Plate tectonics were still active enough to build mountains, with several orogenies, including the Grenville orogeny, occurring at the time.

Climatic stability

Graph showing average temperatures during the Mesoproterozoic. Blue line uses CO₂ concentration 5 times modern day, red line 10 times, and red dot shows average temperature range at the tropics

There is little evidence of significant climatic variability during this time period. Climate was likely not primarily dictated by solar luminosity because the Sun was 5–18% less luminous than it is today, but there is no evidence that Earth's climate was significantly cooler. In fact, the Boring Billion seems to lack any evidence of prolonged glaciations, which can be observed at regular periodicity in other parts of Earth's geologic history. High CO₂ could not have been a main driver for warming because levels would have needed to be 30 to 100 times greater than pre-industrial levels and produced substantial ocean acidification to prevent ice formation, which also did not occur. Mesoproterozoic CO₂ levels may have been comparable to those of the Phanerozoic eon, perhaps 7 to 10 times higher than modern levels. The first record of ice from this time period was reported in 2020 from the 1 Ga Scottish Diabaig Formation in the Torridon Group, where dropstone formations were likely formed by debris from ice rafting; the area, then situated between 35–50°S, was a (possibly upland) lake which is thought to have frozen over in the winter and melted in the summer, rafting occurring in the spring melt.

A higher abundance of other greenhouse gases, namely methane produced by prokaryotes, may have compensated for the low CO₂ levels; a largely ice-free world achieved by an atmospheric methane concentration of 140 parts per million (ppm). Methanogenic prokaryotes could not have produced so much methane, implying some other greenhouse gas, probably nitrous oxide, was elevated, perhaps to 3 ppm (10 times today's levels). Based on presumed greenhouse gas concentrations, equatorial temperatures during the Mesoproterozoic may have been about 295–300 K (22–27 °C; 71–80 °F), in the tropics 290 K (17 °C; 62 °F), at 60° 265–280 K (−8–7 °C; 17–44 °F), and the poles 250–275 K (−23–2 °C; −10–35 °F); and the global average temperature about 19 °C (66 °F), which is 4 °C (7.2 °F) warmer than today. Temperatures at the poles dropped below freezing in winter, allowing for temporary sea ice formation and snowfall, but there were likely no permanent ice sheets.

It has also been proposed that, because the intensity of cosmic rays has been shown to be positively correlated to cloud cover, and cloud cover reflects light into space and reduces global temperatures, lower rates of bombardment during this time due to reduced star formation in the galaxy caused less cloud cover and prevented glaciation events, maintaining a warm climate. Also, some combination of weathering intensity which would have reduced CO₂ levels by oxidation of exposed metals, cooling of the mantle and reduced geothermal heat and volcanism, and increasing solar intensity and solar heat may have reached an equilibrium, barring ice formation.

Conversely, glacial movements over a billion years ago may not have left many remnants today, and an apparent lack of evidence could be due to the incompleteness of the fossil record rather than absence. Further, the low oxygen and solar intensity levels may have prevented the formation of the ozone layer, preventing greenhouse gasses from being trapped in the atmosphere and heating the Earth via the greenhouse effect, which would have caused glaciation. Though not much oxygen is necessary to sustain the ozone layer, and levels during the Boring Billion may have been high enough for it, the Earth may have been more heavily bombarded by UV radiation than today.

Oceanic composition

The oceans seem to have had low concentrations of key nutrients thought to be necessary for complex life, namely molybdenum, iron, nitrogen, and phosphorus, in large part due to a lack of oxygen and resultant oxidation necessary for these geochemical cycles. Nutrients could have been more abundant in terrestrial environments, such as lakes or nearshore environments closer to continental runoff.

In general, the oceans may have had an oxygenated surface layer, a sulfidic middle layer, and suboxic bottom layer. The predominantly sulfidic composition may have caused the oceans to be a black- and milky-turquoise color instead of blue.

Oxygen

Earth's geologic record indicates two events associated with significant increases in oxygen levels on Earth, with one occurring between 2.4 and 2.1 Ga, known as the Great Oxidation Event (GOE), and the second occurring an approximate 0.8 Ga, known as the Neoproterozoic Oxygenation Event (NOE).^[39] The intermediary period, during the Boring Billion, is thought have had low oxygen levels (with minor fluctuations), leading to widespread anoxic waters.

The oceans may have been distinctly stratified, with surface water being oxygenated and deep water being suboxic (less than 1 μM oxygen), the latter possibly maintained by lower levels of hydrogen (H₂) and H₂S output by deep sea hydrothermal vents which otherwise would have been chemically reduced by the oxygen, i.e., euxinic waters. Even in the shallowest waters, significant quantities of oxygen may have been restricted mainly to areas near the coast. The decomposition of sinking organic matter would have also leached oxygen from deep waters.

The sudden drop in O₂ after the Great Oxygenation Event—indicated by δ13C levels to have been a loss of 10 to 20 times the current volume of atmospheric oxygen—is known as the Lomagundi-Jatuli Event, and is the most prominent carbon isotope event in Earth's history. Oxygen levels may have been less than 0.1 to 1% of modern-day levels, which would have effectively stalled the evolution of complex life during the Boring Billion. However, a Mesoproterozoic Oxygenation Event (MOE), during which oxygen rose transiently to about 4% PAL at various points in time, is proposed to have occurred from 1.59 to 1.36 Ga. In particular, some evidence from the Gaoyuzhuang Formation suggests a rise in oxygen around 1.57 Ga, while the Velkerri Formation in the Roper Group of the Northern Territory of Australia, the Kaltasy Formation (Russian: Калтасинская свита) of Volgo-Uralia, Russia, and the Xiamaling Formation in the northern North China Craton indicate noticeable oxygenation around 1.4 Ga, although the degree to which this represents global oxygen levels is unclear. Oxic conditions would have become dominant at the NOE causing the proliferation of aerobic activity over anaerobic, but widespread suboxic and anoxic conditions likely lasted until about 0.55 Ga corresponding with Ediacaran biota and the Cambrian explosion.

Sulfur

Diagram of how euxinic conditions form

In 1998, geologist Donald Canfield proposed what is now known as the Canfield ocean hypothesis. Canfield claimed that increasing levels of oxygen in the atmosphere at the Great Oxygenation Event would have reacted with and oxidized continental iron pyrite (FeS₂) deposits, with sulfate (SO₄²⁻) as a byproduct, which was transported into the sea. Sulfate-reducing microorganisms converted this to hydrogen sulfide (H₂S), dividing the ocean into a somewhat oxic surface layer, and a sulfidic layer beneath, with anoxygenic bacteria living at the border, metabolizing the H₂S and creating sulfur as a waste product. This created widespread euxinic conditions in middle-waters, an anoxic state with a high sulfur concentration, which was maintained by the bacteria. Many deposits from the Boring Billion contain mercury isotopic ratios characteristic of photic zone euxinia  More systematic geochemical study of the Mid-Proterozoic indicates that the oceans were largely ferruginous with a thin surface layer of weakly oxygenated waters, and euxinia may have occurred over relatively small areas, perhaps less than 7% of the seafloor. The very low concentrations of molybdenum in the Mesoproterozoic could sufficiently be explained even with such a relatively low percentage of the seafloor being euxinic. Euxinia expanded and contracted, sometimes reaching the photic zone and at other times being relegated to deeper waters. Evidence from the McArthur Basin of northern Australia reveals that intracontinental settings in particular were low in sulphide intermittently.

Iron

Among rocks dating to the Boring Billion, there is a conspicuous lack of banded iron formations, which form from iron in the upper water column (sourced from the deep ocean) reacting with oxygen and precipitating out of the water. They seemingly cease around the world after 1.85 Ga. Canfield argued that oceanic SO2−4 reduced all the iron in the anoxic deep sea. Iron could have been metabolized by anoxygenic bacteria. It has also been proposed that the 1.85 Ga Sudbury meteor impact mixed the previously stratified ocean via tsunamis, interaction between vaporized seawater and the oxygenated atmosphere, oceanic cavitation, and massive runoff of destroyed continental margins into the sea. Resultant suboxic deep waters (due to oxygenated surface water mixing with previously anoxic deep water) would have oxidized deep-water iron, preventing it from being transported and deposited on continental margins.

Nonetheless, iron-rich waters did exist, such as the 1.4 Ga Xiamaling Formation of Northern China, which perhaps was fed by deep water hydrothermal vents. Iron-rich conditions also indicate anoxic bottom water in this area, as oxic conditions would have oxidized all the iron.

Lifeforms

Low nutrient abundance may have facilitated photosymbiosis—where one organism is capable of photosynthesis and the other metabolizes the waste product—among prokaryotes (bacteria and archaea), and the emergence of eukaryotes. Bacteria, Archaea, and Eukaryota are the three domains, the highest taxonomic ranking. Eukaryotes are distinguished from prokaryotes by a nucleus and membrane-bound organelles, and almost all multicellular organisms are eukaryotes.

See also: Volyn biota

Prokaryotes

1.44 Ga stromatolite from Glacier National Park, Montana

Prokaryotes were the dominant lifeforms throughout the Boring Billion. Microfossils indicate the presence of cyanobacteria, green and purple sulfur bacteria, methane-producing archaea, sulfate-metabolizing bacteria, methane-metabolizing archaea or bacteria, iron-metabolizing bacteria, nitrogen-metabolizing bacteria, and anoxygenic photosynthetic bacteria.

Anoxygenic cyanobacteria are thought to have been the dominant photosynthesizers, metabolizing the abundant H₂S in the oceans. In iron-rich waters, cyanobacteria may have suffered from iron poisoning, especially in offshore waters where iron-rich deep water mixed with surface waters, and thus were outcompeted by other bacteria which could metabolize both iron and H₂S. However, iron poisoning could have been abated by silica-rich waters or biomineralization of iron within the cell.

Unicellular planktonic lineages of cyanobacteria evolved in freshwater during the middle of the Mesoproterozoic, and during the Neoproterozoic both benthic marine and some freshwater ancestors gave rise to marine planktonic cyanobacteria (both nitrogen-fixing and non-nitrogen fixing), contributing to the oxygenation of the Pre-Cambrian oceans.

Research on cyanobacteria in the laboratory has shown that the enzyme nitrogenase, which is used to fix atmospheric nitrogen, stops working when oxygen levels are higher than 10% of current atmospheric levels. The absence of nitrogen due to an increased amount of oxygen would have created a negative feedback loop where atmospheric oxygen levels stabilised at 2%, which began to change about 600 million years ago when landplants started releasing oxygen. By 408 million years ago, nitrogen fixating cyanobacteria had evolved heterocysts to protect their nitrogenase from oxygen.

Eukaryotes

Eukaryotes may have arisen around the beginning of the Boring Billion, coinciding with the accretion of Columbia, which could have somehow increased oceanic oxygen levels. Although there have been claimed records of eukaryotes as early as 2.1 billion years ago, these have been considered questionable, with the oldest unambiguous eukaryote remains dating to around 1.8-1.6 billion years ago in China. Following this, eukaryotic evolution was rather slow, possibly due to the euxinic conditions of the Canfield ocean and a lack of key nutrients and metals which prevented large, complex life with high energy requirements from evolving. Euxinic conditions would have also decreased the solubility of iron and molybdenum, essential metals in nitrogen fixation. A lack of dissolved nitrogen would have favored prokaryotes over eukaryotes, as the former can metabolize gaseous nitrogen. An alternative hypothesis for the lack of diversification among eukaryotes implicates high temperatures during the Boring Billion rather than low oxygen levels, positing that the fact that oxygenation events prior to the Late Neoproterozoic did not kickstart eukaryotic evolution suggests it was not the main limiting factor inhibiting it.

1.6 Ga Ramathallus fossil, the earliest known red alga

Nonetheless, the diversification of crown group eukaryotic macroorganisms seems to have started about 1.6–1 Ga, seemingly coinciding with an increase in key nutrient concentrations. According to molecular clock analysis, plants diverged from animals and fungi about 1.6 Ga; animals and fungi about 1.5 Ga; Bilaterians and cnidarians (animals respectively with and without bilateral symmetry) about 1.3 Ga; sponges 1.35 Ga; and Ascomycota and Basidiomycota (the two divisions of the fungus subkingdom Dikarya) 0.97 Ga. The paper's authors state that their time estimates disagree with the scientific consensus.

Fossils from the late Palaeoproterozoic and early Mesoproterozoic of the Vindhyan sedimentary basin of India, the Ruyang Group of North China, and the Kotuikan Formation of the Anabar Shield of Siberia, among other places, indicate high rates (by pre-Ediacaran standards) of eukaryotic diversification between 1.7 and 1.4 Ga, although much of this diversity is represented by previously unknown, no longer existing clades of eukaryotes. The earliest known red algae mats date to 1.6 Ga. The earliest known fungus dates to 1.01–0.89 Ga from Northern Canada. Multicellular eukaryotes, thought to be the descendants of colonial unicellular aggregates, had probably evolved about 2–1.4 Ga. Likewise, early multicellular eukaryotes likely mainly aggregated into stromatolite mats.

The red alga Bangiomorpha is the earliest known sexually reproducing and meiotic lifeform, and evolved by 1.047 Ga. Based on this, these adaptations evolved between ca. 2–1.4 Ga. Alternatively, these may have evolved well before the last common ancestor of eukaryotes given that meiosis is performed using the same proteins in all eukaryotes, perhaps stretching to as far back as the hypothesized RNA world.

Cell organelles probably originated from free-living cyanobacteria (symbiogenesis) possibly after the evolution of phagocytosis (engulfing other cells) with the removal of the rigid cell wall which was only necessary for asexual reproduction. Mitochondria had already evolved in the Great Oxygenation Event, but plastids used in primoplants for photosynthesis are thought to have appeared about 1.6–1.5 Ga. Histones likely appeared during the Boring Billion to help organize and package the increasing amount of DNA in eukaryotic cells into nucleosomes. Hydrogenosomes used in anaerobic activity may have originated in this time from an archaeon.

Given the evolutionary landmarks achieved by eukaryotes, this time period could be considered an important precursor to the Cambrian explosion about 0.54 Ga, and the evolution of relatively large, complex life.

Ecology

Due to the marginalization of large food particles, such as algae, in favor of cyanobacteria and prokaryotes which do not transmit as much energy to higher trophic levels, a complex food web likely did not form, and large lifeforms with high energy demands could not evolve. Such a food web probably only sustained a small number of protists as, in a sense, apex predators.

The presumably oxygenic photosynthetic eukaryotic acritarchs, perhaps a type of microalga, inhabited the Mesoproterozoic surface waters. Their population may have been largely limited by nutrient availability rather than predation because species have been reported to have survived for hundreds of millions of years, but after 1 Ga, species duration dropped to about 100 Ma, perhaps due to increased herbivory by early protists. This is consistent with species survival dropping to 10 Ma just after the Cambrian explosion and the expansion of herbivorous animals.

The relatively low concentrations of molybdenum in the ocean throughout the Boring Billion have been suggested as a major limiting factor that kept populations of open ocean nitrogen fixing microorganisms, which require molybdenum to produce nitrogenases, low, although freshwater and coastal environments close to riverine sources of dissolved molybdenum may have still hosted significant communities of nitrogen fixers. The low rate of nitrogen fixation, which only ended during the Cryogenian with the evolution of planktonic nitrogen fixers, meant that free ammonium was in short supply across this time interval, severely constraining the evolution and diversification of multicellular biota.

Life on land

Some of the earliest evidence of the prokaryotic colonization of land dates to before 3 Ga, possibly as early as 3.5 Ga. During the Boring Billion, land may have been inhabited mainly by cyanobacterial mats. Dust would have supplied an abundance of nutrients and a means of dispersal for surface-dwelling microbes, though microbial communities could have also formed in caves and freshwater lakes and rivers. By 1.2 Ga, microbial communities may have been abundant enough to have affected weathering, erosion, sedimentation, and various geochemical cycles, and expansive microbial mats could indicate biological soil crust was abundant.

The earliest terrestrial eukaryotes may have been lichen fungi about 1.3 Ga, which grazed on the microbial mats. Abundant eukaryotic microfossils from the freshwater Scottish Torridon Group seems to indicate eukaryotic dominance in non-marine habitats by 1 Ga, probably due to increased nutrient availability in areas closer to the continents and continental runoff. These lichen may have later facilitated plant colonization 0.75 Ga in some manner. A massive increase in terrestrial photosynthetic biomass seems to have occurred about 0.85 Ga, indicated by a flux in terrestrially-sourced carbon, which may have increased oxygen levels enough to support an expansion of multicellular eukaryotes.

Great Oxidation Event

From Wikipedia, the free encyclopedia
https://en.wikipedia.org/wiki/Great_Oxidation_Event

Timescale

O₂ build-up in the Earth's atmosphere. Red and green lines represent the range of the estimates while time is measured in billions of years ago (Ga).

• Stage 1 (3.85–2.45 Ga): Practically no O₂ in the atmosphere. The oceans were also largely anoxic – with the possible exception of O₂ in the shallow oceans.
• Stage 2 (2.45–1.85 Ga): O₂ produced, rising to values of 0.02 and 0.04 atm, but absorbed in oceans and seabed rock.
• Stage 3 (1.85–0.85 Ga): O₂ starts to gas out of the oceans, but is absorbed by land surfaces. No significant change in oxygen level.
• Stages 4 and 5 (0.85 Ga – present): Other O₂ reservoirs filled; gas accumulates in atmosphere.

The Great Oxidation Event (GOE) or Great Oxygenation Event, also called the Oxygen Catastrophe, Oxygen Revolution, Oxygen Crisis or Oxygen Holocaust, was a time interval during the Early Earth's Paleoproterozoic era when the Earth's atmosphere and the shallow ocean first experienced a rise in the concentration of oxygen. This began approximately 2.460–2.426 Ga (billion years) ago during the Siderian period and ended approximately 2.060 Ga ago during the Rhyacian. Geological, isotopic, and chemical evidence suggests that biologically produced molecular oxygen (dioxygen or O₂) started to accumulate in Earth's atmosphere and changed it from a weakly reducing atmosphere practically devoid of oxygen into an oxidizing one containing abundant free oxygen, with oxygen levels being as high as 10% of their present atmospheric level by the end of the GOE.

The sudden injection of highly reactive free oxygen, toxic to the then-mostly anaerobic biosphere, may have caused the extinction of many organisms on Earth – mostly archaeal colonies that used retinal to utilize green-spectrum light energy and power a form of anoxygenic photosynthesis (see Purple Earth hypothesis). Although the event is inferred to have constituted a mass extinction, due in part to the great difficulty in surveying microscopic organisms' abundances, and in part to the extreme age of fossil remains from that time, the Great Oxidation Event is typically not counted among conventional lists of "great extinctions", which are implicitly limited to the Phanerozoic eon. In any case, isotope geochemistry data from sulfate minerals have been interpreted to indicate a decrease in the size of the biosphere of >80% associated with changes in nutrient supplies at the end of the GOE.

The GOE is inferred to have been caused by cyanobacteria that evolved porphyrin-based photosynthesis, which produces dioxygen as a byproduct. The increasing oxygen level eventually depleted the reducing capacity of ferrous compounds, hydrogen sulfide and atmospheric methane, and compounded by a global glaciation, devastated the microbial mats around the Earth's surface. The subsequent adaptation of surviving archaea via symbiogenesis with aerobic proteobacteria (which went endosymbiont and became mitochondria) may have led to the rise of eukaryotic organisms and the subsequent evolution of multicellular life-forms.

Early atmosphere

Main article: Paleoatmosphere

See also: Paleoclimatology and Atmosphere of Earth

The composition of the Earth's earliest atmosphere is not known with certainty. However, the bulk was likely nitrogen, N₂, and carbon dioxide, CO₂, which are also the predominant nitrogen-and-carbon-bearing gases produced by volcanism today. These are relatively inert gases. Oxygen, O₂, meanwhile, was present in the atmosphere at just 0.001% of its present atmospheric level. The Sun shone at about 70% of its current brightness 4 billion years ago, but there is strong evidence that liquid water existed on Earth at the time. A warm Earth, in spite of a faint Sun, is known as the faint young Sun paradox. Either CO₂ levels were much higher at the time, providing enough of a greenhouse effect to warm the Earth, or other greenhouse gases were present. The most likely such gas is methane, CH
₄, which is a powerful greenhouse gas and was produced by early forms of life known as methanogens. Scientists continue to research how the Earth was warmed before life arose.

An atmosphere of N₂ and CO₂ with trace amounts of H₂O, CH₄, carbon monoxide (CO), and hydrogen (H₂) is described as a weakly reducing atmosphere. Such an atmosphere contains practically no oxygen. The modern atmosphere contains abundant oxygen (nearly 21%), making it an oxidizing atmosphere. The rise in oxygen is attributed to photosynthesis by cyanobacteria, which are thought to have evolved as early as 3.5 billion years ago.

The current scientific understanding of when and how the Earth's atmosphere changed from a weakly reducing to a strongly oxidizing atmosphere largely began with the work of the American geologist Preston Cloud in the 1970s. Cloud observed that detrital sediments older than about 2 billion years contained grains of pyrite, uraninite, and siderite, all minerals containing reduced forms of iron or uranium that are not found in younger sediments because they are rapidly oxidized in an oxidizing atmosphere. He further observed that continental red beds, which get their color from the oxidized (ferric) mineral hematite, began to appear in the geological record at about this time. Banded iron formation largely disappears from the geological record at 1.85 Ga, after peaking at about 2.5 Ga. Banded iron formation can form only when abundant dissolved ferrous iron is transported into depositional basins, and an oxygenated ocean blocks such transport by oxidizing the iron to form insoluble ferric iron compounds. The end of the deposition of banded iron formation at 1.85 Ga is therefore interpreted as marking the oxygenation of the deep ocean. Heinrich Holland further elaborated these ideas through the 1980s, placing the main time interval of oxygenation between 2.2 and 1.9 Ga.

Constraining the onset of atmospheric oxygenation has proven particularly challenging for geologists and geochemists. While there is a widespread consensus that initial oxygenation of the atmosphere happened sometime during the first half of the Paleoproterozoic, there is disagreement on the exact timing of this event. Scientific publications between 2016–2022 have differed in the inferred timing of the onset of atmospheric oxygenation by approximately 500 million years; estimates of 2.7 Ga, 2.501–2.434 Ga 2.501–2.225 Ga, 2.460–2.426 Ga, 2.430 Ga, and 2.33 Ga have been given. Factors limiting calculations include an incomplete sedimentary record for the Paleoproterozoic (e.g., because of subduction and metamorphism), uncertainties in depositional ages for many ancient sedimentary units, and uncertainties related to the interpretation of different geological/geochemical proxies. While the effects of an incomplete geological record have been discussed and quantified in the field of paleontology for several decades, particularly with respect to the evolution and extinction of organisms (the Signor–Lipps effect), this is rarely quantified when considering geochemical records and may therefore lead to uncertainties for scientists studying the timing of atmospheric oxygenation.

Geological evidence

Evidence for the Great Oxidation Event is provided by a variety of petrological and geochemical markers that define this geological event.

Continental indicators

Paleosols, detrital grains, and red beds are evidence of low oxygen levels. Paleosols (fossil soils) older than 2.4 billion years old have low iron concentrations that suggest anoxic weathering. Detrital grains composed of pyrite, siderite, and uraninite (redox-sensitive detrital minerals) are found in sediments older than ca. 2.4 Ga. These minerals are only stable under low oxygen conditions, and so their occurrence as detrital minerals in fluvial and deltaic sediments are widely interpreted as evidence of an anoxic atmosphere. In contrast to redox-sensitive detrital minerals are red beds, red-colored sandstones that are coated with hematite. The occurrence of red beds indicates that there was sufficient oxygen to oxidize iron to its ferric state, and these represent a marked contrast to sandstones deposited under anoxic conditions which are often beige, white, grey, or green.

Banded iron formation

Banded iron formations are composed of thin alternating layers of chert (a fine-grained form of silica) and iron oxides (magnetite and hematite). Extensive deposits of this rock type are found around the world, almost all of which are more than 1.85 billion years old and most of which were deposited around 2.5 Ga. The iron in banded iron formations is partially oxidized, with roughly equal amounts of ferrous and ferric iron. Deposition of a banded iron formation requires both an anoxic deep ocean capable of transporting iron in soluble ferrous form, and an oxidized shallow ocean where the ferrous iron is oxidized to insoluble ferric iron and precipitates onto the ocean floor. The deposition of banded iron formations before 1.8 Ga suggests the ocean was in a persistent ferruginous state, but deposition was episodic and there may have been significant intervals of euxinia. The transition from deposition of banded iron formations to manganese oxides in some strata has been considered a key tipping point in the timing of the GOE because it is believed to indicate the escape of significant molecular oxygen into the atmosphere in the absence of ferrous iron as a reducing agent.

Iron speciation

Black laminated shales, rich in organic matter, are often regarded as a marker for anoxic conditions. However, the deposition of abundant organic matter is not a sure indication of anoxia, and burrowing organisms that destroy lamination had not yet evolved during the time frame of the Great Oxygenation Event. Thus laminated black shale by itself is a poor indicator of oxygen levels. Scientists must look instead for geochemical evidence of anoxic conditions. These include ferruginous anoxia, in which dissolved ferrous iron is abundant, and euxinia, in which hydrogen sulfide is present in the water.

Examples of such indicators of anoxic conditions include the degree of pyritization (DOP), which is the ratio of iron present as pyrite to the total reactive iron. Reactive iron, in turn, is defined as iron found in oxides and oxyhydroxides, carbonates, and reduced sulfur minerals such as pyrites, in contrast with iron tightly bound in silicate minerals. A DOP near zero indicates oxidizing conditions, while a DOP near 1 indicates euxinic conditions. Values of 0.3 to 0.5 are transitional, suggesting anoxic bottom mud under an oxygenated ocean. Studies of the Black Sea, which is considered a modern model for ancient anoxic ocean basins, indicate that high DOP, a high ratio of reactive iron to total iron, and a high ratio of total iron to aluminum are all indicators of transport of iron into a euxinic environment. Ferruginous anoxic conditions can be distinguished from euxenic conditions by a DOP less than about 0.7.

The currently available evidence suggests that the deep ocean remained anoxic and ferruginous as late as 580 Ma, well after the Great Oxygenation Event, remaining just short of euxenic during much of this interval of time. Deposition of banded iron formation ceased when conditions of local euxenia on continental platforms and shelves began precipitating iron out of upwelling ferruginous water as pyrite.

Isotopes

Some of the most persuasive evidence for the Great Oxidation Event is provided by the mass-independent fractionation (MIF) of sulfur. The chemical signature of the MIF of sulfur is found prior to 2.4–2.3 Ga but disappears thereafter. The presence of this signature all but eliminates the possibility of an oxygenated atmosphere.

Different isotopes of a chemical element have slightly different atomic masses. Most of the differences in geochemistry between isotopes of the same element scale with this mass difference. These include small differences in molecular velocities and diffusion rates, which are described as mass-dependent fractionation processes. By contrast, MIF describes processes that are not proportional to the difference in mass between isotopes. The only such process likely to be significant in the geochemistry of sulfur is photodissociation. This is the process in which a molecule containing sulfur is broken up by solar ultraviolet (UV) radiation. The presence of a clear MIF signature for sulfur prior to 2.4 Ga shows that UV radiation was penetrating deep into the Earth's atmosphere. This in turn rules out an atmosphere containing more than traces of oxygen, which would have produced an ozone layer that would have shielded the lower atmosphere from UV radiation. The disappearance of the MIF signature for sulfur indicates the formation of such an ozone shield as oxygen began to accumulate in the atmosphere. MIF of sulphur also indicates the presence of oxygen in that oxygen is required to facilitate repeated redox cycling of sulphur.

MIF provides clues to the Great Oxygenation Event. For example, oxidation of manganese in surface rocks by atmospheric oxygen leads to further reactions that oxidize chromium. The heavier ⁵³Cr is oxidized preferentially over the lighter ⁵²Cr, and the soluble oxidized chromium carried into the ocean shows this enhancement of the heavier isotope. The chromium isotope ratio in banded iron formation suggests small but significant quantities of oxygen in the atmosphere before the Great Oxidation Event, and a brief return to low oxygen abundance 500 Ga after the GOE. However, the chromium data may conflict with the sulfur isotope data, which calls the reliability of the chromium data into question. It is also possible that oxygen was present earlier only in localized "oxygen oases". Since chromium is not easily dissolved, its release from rocks requires the presence of a powerful acid such as sulfuric acid (H₂SO₄) which may have formed through bacterial oxidation of pyrite. This could provide some of the earliest evidence of oxygen-breathing life on land surfaces.

Other elements whose MIF may provide clues to the GOE include carbon, nitrogen, transitional metals such as molybdenum and iron, and non-metal elements such as selenium.

Fossils and biomarkers

See also: Biomarker (petroleum)

While the GOE is generally thought to be a result of oxygenic photosynthesis by ancestral cyanobacteria, the presence of cyanobacteria in the Archaean before the GOE is a highly controversial topic. Structures that are claimed to be fossils of cyanobacteria exist in rock formed 3.5 Ga. These include microfossils of supposedly cyanobacterial cells and macrofossils called stromatolites, which are interpreted as colonies of microbes, including cyanobacteria, with characteristic layered structures. Modern stromatolites, which can only be seen in harsh environments such as Shark Bay in Western Australia, are associated with cyanobacteria, and thus fossil stromatolites had long been interpreted as the evidence for cyanobacteria. However, it has increasingly been inferred that at least some of these Archaean fossils were generated abiotically or produced by non-cyanobacterial phototrophic bacteria.

Additionally, Archaean sedimentary rocks were once found to contain biomarkers, also known as chemical fossils, interpreted as fossilized membrane lipids from cyanobacteria and eukaryotes. For example, traces of 2α-methylhopanes and steranes that are thought to be derived from cyanobacteria and eukaryotes, respectively, were found in the Pilbara of Western Australia. Steranes are diagenetic products of sterols, which are biosynthesized utilizing molecular oxygen. Thus, steranes can additionally serve as an indicator of oxygen in the atmosphere. However, these biomarker samples have since been shown to have been contaminated, and so the results are no longer accepted.

Carbonaceous microfossils from the Turee Creek Group of Western Australia, which date back to ~2.45–2.21 Ga, have been interpreted as iron-oxidising bacteria. Their presence suggests a minimum threshold of seawater oxygen content had been reached by this interval of time.

Other indicators

Some elements in marine sediments are sensitive to different levels of oxygen in the environment such as the transition metals molybdenum and rhenium. Non-metal elements such as selenium and iodine are also indicators of oxygen levels.

Hypotheses

See also: Oxygen cycle

Life timeline
−4500 — – — – −4000 — – — – −3500 — – — – −3000 — – — – −2500 — – — – −2000 — – — – −1500 — – — – −1000 — – — – −500 — – — – 0 —	Water   Single-celled life   Photosynthesis   Eukaryotes   Multicellular life   P l a n t s   Arthropods Molluscs Flowers Dinosaurs   Mammals Birds Primates H a d e a n A r c h e a n P r o t e r o z o i c P h a n e r o z o i c	← Earth formed ← Earliest water ← LUCA ← Earliest fossils ← LHB meteorites ← Earliest oxygen ← Pongola glaciation* ← Atmospheric oxygen ← Huronian glaciation* ← Sexual reproduction ← Earliest multicellular life ← Earliest fungi ← Earliest plants ← Earliest animals ← Cryogenian ice age* ← Ediacaran biota ← Cambrian explosion ← Andean glaciation* ← Earliest tetrapods ← Karoo ice age* ← Earliest apes / humans ← Quaternary ice age*
(million years ago) *Ice Ages

The ability to generate oxygen via photosynthesis likely first appeared in the ancestors of cyanobacteria. These organisms evolved at least 2.45–2.32 Gaand probably as early as 2.7 Ga or earlier. However, oxygen remained scarce in the atmosphere until around 2.0 Ga, and banded iron formation continued to be deposited until around 1.85 Ga. Given the rapid multiplication rate of cyanobacteria under ideal conditions, an explanation is needed for the delay of at least 400 million years between the evolution of oxygen-producing photosynthesis and the appearance of significant oxygen in the atmosphere.

Hypotheses to explain this gap must take into consideration the balance between oxygen sources and oxygen sinks. Oxygenic photosynthesis produces organic carbon that must be segregated from oxygen to allow oxygen accumulation in the surface environment, otherwise the oxygen back-reacts with the organic carbon and does not accumulate. The burial of organic carbon, sulfide, and minerals containing ferrous iron (Fe²⁺) is a primary factor in oxygen accumulation. When organic carbon is buried without being oxidized, the oxygen is left in the atmosphere. In total, the burial of organic carbon and pyrite today creates 15.8±3.3 Tmol (1 Tmol = 10¹² moles) of O₂ per year. This creates a net O₂ flux from the global oxygen sources.

The rate of change of oxygen can be calculated from the difference between global sources and sinks. The oxygen sinks include reduced gases and minerals from volcanoes, metamorphism and weathering. The GOE started after these oxygen-sink fluxes and reduced-gas fluxes were exceeded by the flux of O₂ associated with the burial of reductants, such as organic carbon. About 12.0±3.3 Tmol of O₂ per year today goes to the sinks composed of reduced minerals and gases from volcanoes, metamorphism, percolating seawater and heat vents from the seafloor. On the other hand, 5.7±1.2 Tmol of O₂ per year today oxidizes reduced gases in the atmosphere through photochemical reaction. On the early Earth, there was visibly very little oxidative weathering of continents (e.g., a lack of red beds), and so the weathering sink on oxygen would have been negligible compared to that from reduced gases and dissolved iron in oceans.

Dissolved iron in oceans exemplifies O₂ sinks. Free oxygen produced during this time was chemically captured by dissolved iron, converting iron Fe and Fe²⁺ to magnetite (Fe²⁺Fe3+2O₄) that is insoluble in water, and sank to the bottom of the shallow seas to create banded iron formations. It took 50 million years or longer to deplete the oxygen sinks. The rate of photosynthesis and associated rate of organic burial also affect the rate of oxygen accumulation. When land plants spread over the continents in the Devonian, more organic carbon was buried and likely allowed higher O₂ levels to occur. Today, the average time that an O₂ molecule spends in the air before it is consumed by geological sinks is about 2 million years. That residence time is relatively short in geologic time; so in the Phanerozoic, there must have been feedback processes that kept the atmospheric O₂ level within bounds suitable for animal life.

Evolution by stages

Preston Cloud originally proposed that the first cyanobacteria had evolved the capacity to carry out oxygen-producing photosynthesis but had not yet evolved enzymes (such as superoxide dismutase) for living in an oxygenated environment. These cyanobacteria would have been protected from their own poisonous oxygen waste through its rapid removal via the high levels of reduced ferrous iron, Fe(II), in the early ocean. He suggested that the oxygen released by photosynthesis oxidized the Fe(II) to ferric iron, Fe(III), which precipitated out of the sea water to form banded iron formation. He interpreted the great peak in deposition of banded iron formation at the end of the Archean as the signature for the evolution of mechanisms for living with oxygen. This ended self-poisoning and produced a population explosion in the cyanobacteria that rapidly oxygenated the ocean and ended banded iron formation deposition. However, improved dating of Precambrian strata showed that the late Archean peak of deposition was spread out over tens of millions of years, rather than taking place in a very short interval of time following the evolution of oxygen-coping mechanisms. This made Cloud's hypothesis untenable.

Most modern interpretations describe the GOE as a long, protracted process that took place over hundreds of millions of years rather than a single abrupt event, with the quantity of atmospheric oxygen fluctuating in relation to the capacity of oxygen sinks and the productivity of oxygenic photosynthesisers over the course of the GOE. More recently, families of bacteria have been discovered that closely resemble cyanobacteria but show no indication of ever having possessed photosynthetic capability. These may be descended from the earliest ancestors of cyanobacteria, which only later acquired photosynthetic ability by lateral gene transfer. Based on molecular clock data, the evolution of oxygen-producing photosynthesis may have occurred much later than previously thought, at around 2.5 Ga. This reduces the gap between the evolution of oxygen photosynthesis and the appearance of significant atmospheric oxygen.

Nutrient famines

Another possibility is that early cyanobacteria were starved for vital nutrients, and this checked their growth. However, a lack of the scarcest nutrients, iron, nitrogen, and phosphorus, could have slowed but not prevented a cyanobacteria population explosion and rapid oxygenation. The explanation for the delay in the oxygenation of the atmosphere following the evolution of oxygen-producing photosynthesis likely lies in the presence of various oxygen sinks on the young Earth.

Nickel famine

Early chemosynthetic organisms likely produced methane, an important trap for molecular oxygen, since methane readily oxidizes to carbon dioxide (CO₂) and water in the presence of UV radiation. Modern methanogens require nickel as an enzyme cofactor. As the Earth's crust cooled and the supply of volcanic nickel dwindled, oxygen-producing algae began to outperform methane producers, and the oxygen percentage of the atmosphere steadily increased. From 2.7 to 2.4 Ga the rate of deposition of nickel declined steadily from a level 400 times that of today. This nickel famine was somewhat buffered by an uptick in sulfide weathering at the start of the GOE that brought some nickel to the oceans, without which methanogenic organisms would have declined in abundance more precipitously, plunging Earth into even more severe and long-lasting icehouse conditions than those seen during the Huronian glaciation.

Large igneous provinces

Another hypothesis posits that a number of large igneous provinces (LIPs) were emplaced during the GOE and fertilised the oceans with limiting nutrients, facilitating and sustaining cyanobacterial blooms.

Increasing flux

One hypothesis argues that the GOE was the immediate result of photosynthesis, although the majority of scientists suggest that a long-term increase of oxygen is more likely. Several model results show possibilities of long-term increase of carbon burial, but the conclusions are indeterminate.

Decreasing sink

In contrast to the increasing flux hypothesis, there are several hypotheses that attempt to use decrease of sinks to explain the GOE. One theory suggests that the composition of the volatiles from volcanic gases was more oxidized. Another theory suggests that the decrease of metamorphic gases and serpentinization is the main key of GOE. Hydrogen and methane released from metamorphic processes are also lost from Earth's atmosphere over time and leave the crust oxidized. Scientists realized that hydrogen would escape into space through a process called methane photolysis, in which methane decomposes under the action of ultraviolet light in the upper atmosphere and releases its hydrogen. The escape of hydrogen from the Earth into space must have oxidized the Earth because the process of hydrogen loss is chemical oxidation. This process of hydrogen escape required the generation of methane by methanogens, so that methanogens actually helped create the conditions necessary for the oxidation of the atmosphere.

Tectonic trigger

2.1-billion-year-old rock showing banded iron formation

One hypothesis suggests that the oxygen increase had to await tectonically driven changes in the Earth, including the appearance of shelf seas, where reduced organic carbon could reach the sediments and be buried. The burial of reduced carbon as graphite or diamond around subduction zones released molecular oxygen into the atmosphere. The appearance of oxidised magmas enriched in sulphur formed around subduction zones confirms changes in tectonic regime played an important role in the oxygenation of Earth's atmosphere.

The newly produced oxygen was first consumed in various chemical reactions in the oceans, primarily with iron. Evidence is found in older rocks that contain massive banded iron formations apparently laid down as this iron and oxygen first combined; most present-day iron ore lies in these deposits. It was assumed oxygen released from cyanobacteria resulted in the chemical reactions that created rust, but it appears the iron formations were caused by anoxygenic phototrophic iron-oxidizing bacteria, which does not require oxygen. Evidence suggests oxygen levels spiked each time smaller land masses collided to form a super-continent. Tectonic pressure thrust up mountain chains, which eroded releasing nutrients into the ocean that fed photosynthetic cyanobacteria.

Bistability

Another hypothesis posits a model of the atmosphere that exhibits bistability: two steady states of oxygen concentration. The state of stable low oxygen concentration (0.02%) experiences a high rate of methane oxidation. If some event raises oxygen levels beyond a moderate threshold, the formation of an ozone layer shields UV rays and decreases methane oxidation, raising oxygen further to a stable state of 21% or more. The Great Oxygenation Event can then be understood as a transition from the lower to the upper steady states.

Increasing photoperiod

Cyanobacteria tend to consume nearly as much oxygen at night as they produce during the day. However, experiments demonstrate that cyanobacterial mats produce a greater excess of oxygen with longer photoperiods. The rotational period of the Earth was only about six hours shortly after its formation 4.5 Ga but increased to 21 hours by 2.4 Ga in the Paleoproterozoic. The rotational period increased again, starting 700 million years ago, to its present value of 24 hours. The total amount of oxygen produced by the cyanobacteria remained the same with longer days, but the longer the day, the more time oxygen has to diffuse into the water.

Consequences of oxygenation

Eventually, oxygen started to accumulate in the atmosphere, with two major consequences.

• Oxygen likely oxidized atmospheric methane (a strong greenhouse gas) to carbon dioxide (a weaker one) and water. This weakened the greenhouse effect of the Earth's atmosphere, causing planetary cooling, which has been proposed to have triggered a series of ice ages known as the Huronian glaciation, bracketing an age range of 2.45–2.22 Ga.

Timeline of glaciations, shown in blue.

• The increased oxygen concentrations provided a new opportunity for biological diversification, as well as tremendous changes in the nature of chemical interactions between rocks, sand, clay, and other geological substrates and the Earth's air, oceans, and other surface waters. Despite the natural recycling of organic matter, life had remained energetically limited until the widespread availability of oxygen. The availability of oxygen greatly increased the free energy available to living organisms, with global environmental impacts. For example, mitochondria evolved after the GOE, giving organisms the energy to exploit new, more complex morphologies interacting in increasingly complex ecosystems, although these did not appear until the late Proterozoic and Cambrian.

Mineral diversification

Main article: Mineral evolution

The Great Oxygenation Event triggered an explosive growth in the diversity of minerals, with many elements occurring in one or more oxidized forms near the Earth's surface. It is estimated that the GOE was directly responsible for deposition of more than 2,500 of the total of about 4,500 minerals found on Earth today. Most of these new minerals were formed as hydrated and oxidized forms due to dynamic mantle and crust processes.

↓GOE

↓End of Huronian glaciation

Ediacaran

Cambrian

Ordovician

Silurian

Devonian

Carboniferous

Permian

Triassic

Jurassic

Cretaceous

Paleogene

Neogene

Quaternary

Palæoproterozoic

Mesoproterozoic

Neoproterozoic

Palæozoic

Mesozoic

Cenozoic

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−1900

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Million years ago. Age of Earth = 4,560

Cyanobacteria evolution

In field studies done in Lake Fryxell, Antarctica, scientists found that mats of oxygen-producing cyanobacteria produced a thin layer, one to two millimeters thick, of oxygenated water in an otherwise anoxic environment, even under thick ice. By inference, these organisms could have adapted to oxygen even before oxygen accumulated in the atmosphere. The evolution of such oxygen-dependent organisms eventually established an equilibrium in the availability of oxygen, which became a major constituent of the atmosphere.

Origin of eukaryotes

It has been proposed that a local rise in oxygen levels due to cyanobacterial photosynthesis in ancient microenvironments was highly toxic to the surrounding biota, and that this selective pressure drove the evolutionary transformation of an archaeal lineage into the first eukaryotes. Oxidative stress involving production of reactive oxygen species (ROS) might have acted in synergy with other environmental stresses (such as ultraviolet radiation and/or desiccation) to drive selection in an early archaeal lineage towards eukaryosis. This archaeal ancestor may already have had DNA repair mechanisms based on DNA pairing and recombination, and possibly some kind of cell fusion mechanism. The detrimental effects of internal ROS (produced by endosymbiont proto-mitochondria) on the archaeal genome could have promoted the evolution of meiotic sex from these humble beginnings. Selective pressure for efficient DNA repair of oxidative DNA damage may have driven the evolution of eukaryotic sex involving such features as cell-cell fusions, cytoskeleton-mediated chromosome movements and emergence of the nuclear membrane. Thus the evolution of eukaryotic sex and eukaryogenesis were likely inseparable processes that evolved in large part to facilitate DNA repair. The evolution of mitochondria, which are well suited for oxygenated environments, may have occurred during the GOE.

However, other authors express scepticism that the GOE resulted in widespread eukaryotic diversification due to the lack of robust evidence, concluding that the oxygenation of the oceans and atmosphere does not necessarily lead to increases in ecological and physiological diversity.

Lomagundi-Jatuli event

The rise in oxygen content was not linear: instead, there was a rise in oxygen content around 2.3 Ga, followed by a drop around 2.1 Ga. This rise in oxygen is called the Lomagundi-Jatuli event or Lomagundi event, (named for a district of Southern Rhodesia) and the time period has been termed Jatulian; it is currently considered to be part of the Rhyacian period. During the Lomagundi-Jatuli event, oxygen amounts in the atmosphere reached similar heights to modern levels, before returning to low levels during the following stage, which caused the deposition of black shales (rocks that contain large amounts of organic matter that would otherwise have been burned away by oxygen). This drop in oxygen levels is called the Shunga-Francevillian event. Evidence for the event has been found globally in places such as Fennoscandia and the Wyoming Craton. Oceans seem to have stayed rich in oxygen for some time even after the event ended.

It has been hypothesized that eukaryotes first evolved during the Lomagundi-Jatuli event.