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 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 (Ga) 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 the Archean prebiotic atmosphere due to 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.

The appearance of highly reactive free oxygen, which can oxidize organic compounds (especially genetic materials) 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 (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, 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 (which went endosymbiont and became mitochondria) may have led to the rise of eukaryotic organisms and the subsequent evolution of multicellular life-forms.

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 modern volcanism. These are relatively inert gases. Oxygen, O₂, 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 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 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, 2.33 Ga, and 2.3 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.

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