Great Oxidation Event
The Great Oxidation Event or Great Oxygenation Event, also called the Oxygen Catastrophe, Oxygen Revolution, Oxygen Crisis, or Oxygen Holocaust, was a time interval during the Earth's Paleoproterozoic era when the Earth's atmosphere and shallow seas first experienced a rise in the concentration of free oxygen. This began approximately 2.460–2.426 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 started to accumulate in the Archean prebiotic atmosphere by microbial photosynthesis and eventually 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 modern atmospheric level by the end of the GOE.
Overview
The appearance of highly reactive free oxygen, which can oxidize organic compounds and thus is toxic to the then-mostly anaerobic biosphere, may have caused the extinction/extirpation of many early organisms on Earth—mostly archaeal colonies that used retinal to use green-spectrum light energy and power a form of anoxygenic photosynthesis. 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, which evolved chlorophyll-based photosynthesis that releases dioxygen as a byproduct of water photolysis. The continually produced oxygen eventually depleted all the surface reducing capacity from ferrous iron, sulfur, hydrogen sulfide and atmospheric methane over nearly a billion years. The oxidative environmental change, 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 may have led to the rise of eukaryotic organisms and the subsequent evolution of multicellular life-forms.
Early atmosphere
The composition of the Earth's earliest atmosphere is not known with certainty. However, the bulk was likely nitrogen, and carbon dioxide, which are also the predominant nitrogen- and carbon-bearing gases produced by modern volcanism. These are relatively inert gases. Oxygen in its form, meanwhile, was present in the atmosphere at just 0.001% of recent atmospheric levels.The Sun shone at about 70% of its modern 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 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,, 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 and with trace amounts of,, carbon monoxide, and hydrogen is described as a weakly reducing atmosphere. Such an atmosphere contains practically no oxygen. The modern atmosphere contains abundant oxygen, 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 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 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, 2.33 Ga, and 2.3 Ga have been given. Factors limiting calculations include an incomplete sedimentary record for the Paleoproterozoic, 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, 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
s, detrital grains, and red beds are evidence of low oxygen levels. Paleosols older than 2.4 billion years old have low iron concentrations that suggest anoxic weathering. Detrital grains composed of pyrite, siderite, and uraninite 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
are composed of thin alternating layers of chert and iron oxides. 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, 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 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.