System/Period Series/Epoch Stage/Age Age (Ma)PaleogenePaleoceneDanianyounger Cretaceous Upper/Late Maastrichtian 66.0 72.1 Campanian 72.1 83.6 Santonian 83.6 86.3 Coniacian 86.3 89.8 Turonian 89.8 93.9 Cenomanian 93.9 100.5 Lower/Early Albian 100.5 113.2 Aptian 113.2 121.4 Barremian 121.4 125.77 Hauterivian 125.77 132.6 Valanginian 132.6 137.05 Berriasian 137.05 ≈143.0JurassicUpper/LateTithonianolderSubdivision of the Cretaceous systemaccording to the ICS, as of 2024.Cohen, K.M., Finney, S.C., Gibbard, P.L. & Fan, J.-X. (2013; updated) The ICS International Chronostratigraphic Chart. Episodes 36: 199–204.
The Cenomanian–Turonian boundary event, also known as the Cenomanian–Turonian extinction, Cenomanian–Turonian Oceanic Anoxic Event (OAE 2), and referred to also as the Bonarelli Event or Level, was an anoxic extinction event in the Cretaceous period. The Cenomanian–Turonian oceanic anoxic event is considered to be the most recent truly global oceanic anoxic event in Earth's geologic history. There was a large carbon cycle disturbance during this time period, signified by a large positive carbon isotope excursion. However, apart from the carbon cycle disturbance, there were also large disturbances in the ocean's nitrogen, oxygen, phosphorus, sulfur, and iron cycles.
Background The Cenomanian and Turonian stages were first noted by D'Orbigny between 1843 and 1852. The global type section for this boundary is located in the Bridge Creek Limestone Member of the Greenhorn Formation near Pueblo, Colorado, which are bedded with the Milankovitch orbital signature. Here, a positive carbon-isotope event is clearly shown, although none of the characteristic, organic-rich black shale is present. It has been estimated that the isotope shift lasted approximately 850,000 years longer than the black shale event, which may be the cause of this anomaly in the Colorado type section. A significantly expanded OAE2 interval from southern Tibet documents a complete, more detailed, and finer-scale structures of the positive carbon isotope excursion that contains multiple shorter-term carbon isotope stages amounting to a total duration of 820 ±25 ka.
The level is also known as the Bonarelli Event because of layer of thick, black shale that marks the boundary and was first studied by in 1891.G. Bonarelli, Il territorio di Gubbio - Notizie geologiche, Roma 1891 It is characterized by interbedded black shales, chert and radiolarian sands and is estimated to span a 400,000-year interval. Planktonic foraminifera do not exist in this Bonarelli Level, and the presence of radiolarians in this section indicates relatively high productivity and an availability of nutrients.G. Parisi, F. Piergiovanni e M. Marcucci, Il livello Bonarelli nell'area umbro-marchigiana, in Stratigrafia del Mesozoico e Cenozoico nell'area Umbro-Marchigiana, Roma, 1989 In the Western Interior Seaway, the Cenomanian-Turonian boundary event is associated with the Benthonic Zone, characterised by a higher density of benthic foraminifera relative to planktonic foraminifera, although the timing of the appearance of the Benthonic Zone is not uniformly synchronous with the onset of the oceanic anoxic event and is thus cannot be used to consistently demarcate its beginning.
Timeline Selby et al. in 2009 concluded the OAE 2 occurred approximately 91.5 ± 8.6 Ma, though estimates published by Leckie et al. (2002) are given as 93–94 Ma. The Cenomanian-Turonian boundary has been refined in 2012 to 93.9 ± 0.15 Ma. The total duration of OAE2 has been estimated at 0.9 Myr, 0.82 ± 0.025 Myr, or 0.71 ± 0.17 Myr. At high latitudes, the event lasted for a shorter time: only ~600 kyr.
Biodiversity patterns of planktic foraminifera indicate that the Cenomanian–Turonian extinction occurred in five phases. Phase I, which took place from 313,000 to 55,000 years before the onset of the anoxic event, witnessed a stratified water column and high planktonic foraminiferal diversity, suggesting a stable marine environment. Phase II, characterised by significant environmental perturbations, lasted from 55,000 years before OAE2 until its onset and witnessed a decline in rotaliporids and heterohelicids, a zenith of schackoinids and hedbergellids, a 'large form eclipse' during which foraminifera exceeding 150 microns disappeared, and the start of a trend of dwarfism among many foraminifera. This phase also saw an enhanced oxygen minimum zone and increased productivity in surface waters. Phase III lasted for 100,000 to 900,000 years and was coincident with the Bonarelli Level's deposition and exhibited extensive proliferation of radiolarians, indicative of extremely eutrophic conditions. Phase IV lasted for around 35,000 years and was most notable for the increase in the abundance of hedbergellids and schackoinids, being extremely similar to Phase II, with the main difference being that rotaliporids were absent from Phase IV. Phase V was a recovery interval lasting 118,000 years and marked the end of the 'large form eclipse' that began in Phase II; heterohelicids and hedbergellids remained in abundance during this phase, pointing to continued environmental disturbance during this phase.
Causes Climate change Earth pronouncedly warmed just before the beginning of OAE2. The Cenomanian–Turonian interval represents one of the hottest intervals of the entire Phanerozoic eon, and it boasted the highest carbon dioxide concentrations of the Cretaceous period. Even before OAE2, during the late Cenomanian, tropical sea surface temperatures (SSTs) were very warm, about 27-29 °C. The onset of OAE2 was concurrent with a 4-5 °C rise in shelf sea temperatures. Mean tropical SSTs during OAE2 have been conservatively estimated to have been at least 30 °C, but may have reached as high as 36 °C. Minimum SSTs in mid-latitude oceans were >20 °C. This exceptional warmth persisted until the Turonian–Coniacian boundary.
One possible cause of this hothouse was sub-oceanic volcanism. During the middle of the Cretaceous period, the rate of crustal production reached a peak, which may have been related to the rifting of the newly formed Atlantic Ocean. It was also caused by the widespread melting of hot mantle plumes under the ocean crust, at the base of the lithosphere, which may have resulted in the thickening of the oceanic crust in the Pacific and Indian Oceans. The resulting volcanism would have sent large quantities of carbon dioxide into the atmosphere, leading to an increase in global temperatures. Greenhouse gas release was further increased by the degassing of organic-rich sediments intruded into by volcanic sills. Several independent events related to large igneous provinces (LIPs) occurred around the time of OAE2. A multitude of LIPs were active during OAE2: the Madagascar, Caribbean, Gorgona, Ontong Java, and High Arctic LIPs. The abundance of LIPs at this time reflects a major overturning in mantle convection. Trace metals such as chromium (Cr), scandium (Sc), copper (Cu) and cobalt (Co) have been found at the Cenomanian–Turonian boundary, which suggests that an LIP could have been one of the main basic causes involved in the contribution of the event. The timing of the peak in trace metal concentration coincides with the middle of the anoxic event, suggesting that the effects of the LIPs may have occurred during the event, but may not have initiated the event. Other studies linked the lead (Pb) isotopes of OAE-2 to the Caribbean–Colombian and the Madagascar LIPs. An osmium isotope excursion coeval with OAE2 strongly suggests submarine volcanism as its cause; in the Pacific, an unradiogenic osmium spike began about 350 kyr before the onset of OAE2 and terminated around 240 kyr after OAE2's beginning; the osmium isotope data from a highly expanded OAE2 interval in southern Tibet show multiple osmium excursions with the most pronounced one lagging the onset of OAE2 by ≈50 kyr that was probably related to the ocean connectivity change at ~94.5 Ma. Osmium data also reveal that three distinct pulses of intense volcanism occurred ~60, ~270, and ~400 kyr after OAE2's onset, prolonging it. Positive neodymium isotope excursions provide additional indications of pervasive volcanism as a cause of OAE2. Enrichments in zinc further bolster and reinforce the existence of extensive hydrothermal volcanism, as do extreme negative δ53Cr excursions. The absence of geographically widespread mercury (Hg) anomalies resulting from OAE2 has been suggested to be because of the limited dispersal range of this heavy metal by submarine volcanism. A modeling study performed in 2011 confirmed that it is possible that a LIP may have initiated the event, as the model revealed that the peak amount of carbon dioxide degassing from volcanic LIP degassing could have resulted in more than 90 percent global deep-ocean anoxia.
Later on, when anoxia became widespread, the production of nitrous oxide, a greenhouse gas about 265 times more potent than carbon dioxide, drastically increased because of elevated nitrification and denitrification rates. This powerful positive feedback mechanism is what may have enabled extremely hot temperatures to persist in spite of the supercharged organic carbon burial associated with anoxic events.
Plenus Cool Event Large-scale organic carbon burial acted as a negative feedback loop that partially mitigated the warming effects of volcanic discharge of carbon dioxide, resulting in the Plenus Cool Event during the Metoicoceras geslinianum European ammonite biozone. Global average temperatures fell to around 4 °C lower than they were pre-OAE2. Equatorial SSTs dropped by 2.5–5.5 °C. This cooling event was insufficient at completely stopping the rise in global temperatures. This negative feedback was ultimately overridden, as global temperatures continued to shoot up in sync with continued volcanic release of carbon dioxide following the Plenus Cool Event, although this theory has been criticised and the warming after the Plenus Cool Event attributed to decreased silicate weathering instead.
Ocean acidification Within the oceans, the emission of SO2, H2S, CO2, and halogens would have increased the acidity of the water, causing the dissolution of carbonate, and a further release of carbon dioxide. Evidence of ocean acidification can be gleaned from δ44/40Ca increases coeval with the extinction event, as well as coccolith malformation and dwarfism. Lithologies characterised by low calcium carbonate concentrations predominated during intervals of carbonate compensation depth shoaling. Ocean acidification was exacerbated by a positive feedback loop of increased heterotrophic respiration in highly biologically productive waters, elevating seawater concentrations of carbon dioxide and further decreasing pH.
Anoxia and euxinia