Cenomanian-Turonian boundary event

Source: Wikipedia, the free encyclopedia.
Period
Epoch
Age
Ma
)
Paleogene Paleocene Danian younger
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 ≈113.0
Aptian ≈113.0 ≈125.0
Barremian ≈125.0 ≈129.4
Hauterivian ≈129.4 ≈132.9
Valanginian ≈132.9 ≈139.8
Berriasian ≈139.8 ≈145.0
Jurassic Upper/
Late 		Tithonian older
Subdivision of the Cretaceous system
according to the ICS, as of 2017.[1]

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,[2] 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.[3] There was a large carbon cycle disturbance during this time period,[4][5] signified by a large positive carbon isotope excursion.[6][7][8] However, apart from the carbon cycle disturbance, there were also large disturbances in the ocean's nitrogen,[9] oxygen,[10] phosphorus,[11][12][13] sulphur,[14] and iron cycles.[15]

Background

The Cenomanian and Turonian stages were first noted by

type section.[16] 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.[17]

The level is also known as the Bonarelli Event because of 1-to-2-metre (3 ft 3 in to 6 ft 7 in) layer of thick, black shale that marks the boundary and was first studied by Guido Bonarelli [it] in 1891.[18] 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.[19] 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.[20]

Timeline

Selby et al. in 2009 concluded the OAE 2 occurred approximately 91.5 ± 8.6 Ma,[21] though estimates published by Leckie et al. (2002) are given as 93–94 Ma.[22] The Cenomanian-Turonian boundary has been refined in 2012 to 93.9 ± 0.15 Ma.[23] The total duration of OAE2 has been estimated at 0.82 ± 0.025 Myr[17] or 0.71 ± 0.17 Myr.[24] At high latitudes, the event lasted for a shorter time: only ~600 kyr.[25]

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.[26]

Causes

Climate change

Earth pronouncedly warmed just before the beginning of OAE2.

sea surface temperatures (SSTs) were very warm, about 27-29 °C.[30] The onset of OAE2 was concurrent with a 4-5 °C rise in shelf sea temperatures.[31] 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.[32] Minimum SSTs in mid-latitude oceans were >20 °C.[33] This exceptional warmth persisted until the Turonian-Coniacian boundary.[34]

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.

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.[36] 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,[37][38] Caribbean,[39][40][41] Gorgona,[42] Ontong Java,[37] and High Arctic LIPs.[43][44][45] The abundance of LIPs at this time reflects a major overturning in mantle convection.[46] 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.[47] 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.[48] An osmium isotope excursion coeval with OAE2 strongly suggests submarine volcanism as its cause;[49] 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;[50] 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.[51] Osmium data also reveal that three distinct pulses of intense volcanism occurred ~60, ~270, and ~400 kyr after OAE2's onset, prolonging it.[52] Positive neodymium isotope excursions provide additional indications of pervasive volcanism as a cause of OAE2.[53] Enrichments in zinc further bolster and reinforce the existence of extensive hydrothermal volcanism,[54] as do extreme negative δ53Cr excursions.[55] 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.[56] 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.[57]

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.[58]

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.[59] Global average temperatures fell to around 4 °C lower than they were pre-OAE2.[30] Equatorial SSTs dropped by 2.5–5.5 °C.[60] 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,[59] although this theory has been criticised and the warming after the Plenus Cool Event attributed to decreased silicate weathering instead.[61]

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,[62][63][64] as well as coccolith malformation and dwarfism.[65] 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.[66]

Anoxia and euxinia

When the volcanic activity declined, this run-away

mass extinction.[67] An acceleration of the hydrological cycle induced by warmer global temperatures drove greater fluxes of nutrient runoff into the oceans, fuelling primary productivity.[68][69][70] The global environmental disturbance that resulted in these conditions increased atmospheric and oceanic temperatures. Extreme hothouse conditions encouraged ocean stratification.[71] Boundary sediments show an enrichment of trace elements, and contain elevated δ13C values.[72][73] The positive δ13C excursion found at the Cenomanian-Turonian boundary is one of the main carbon isotope events of the Mesozoic. It represents one of the largest disturbances in the global carbon cycle from the past 110 million years. This δ13C excursion indicates a significant increase in the burial rate of organic carbon, indicating the widespread deposition and preservation of organic carbon-rich sediments and that the ocean was depleted of oxygen at the time.[74][75][76] Depletion of manganese in sediments corresponding to OAE2 provides additional strong evidence of severe bottom water oxygen depletion.[54] An increase in the abundance of the planktonic foraminifer Heterohelix provides further evidence still of anoxia.[77][52] The resulting elevated levels of carbon burial would account for the black shale deposition in the ocean basins.[72][78] The proto-North Atlantic in particular was a hotbed of carbon burial during OAE2 as it was in later, less severe anoxic events.[79] Though anoxia was prevalent throughout the interval, there were transient periods of reoxygenation during OAE2.[6]

Sulphate reduction increased during OAE2,[15] causing euxinia, a type of anoxia defined by sulphate reduction and hydrogen sulphide production, to occur during OAE2, as revealed by negative δ53Cr excursions,[80] positive δ98Mo excursions,[81] a low seawater molybdenum inventory,[82] and molecular biomarkers of green sulfur bacteria.[83][84][85] Although euxinia was not uncommon in the latter part of the Cenomanian, it only expanded into the photic zone during OAE2 itself.[86]

OAE2 began on the southern margins of the proto-North Atlantic, from where anoxia spread across the rest of the proto-North Atlantic and then into the Western Interior Seaway (WIS) and the epicontinental seas of the Western Tethys.[87] Anoxic waters spread rapidly throughout the WIS due to marine transgression and a powerful cyclonic circulation resulting from an imbalance between precipitation in the north and evaporation in the south.[88] Anoxia was especially intense in the eastern North Sea, evidenced by its very positive δ13C values.[89] Thanks to persistent upwelling, some marine regions, such as the South Atlantic, were able to remain partially oxygenated at least intermittently.[90] Indeed, redox states of oceans vary geographically, bathymetrically and temporally during OAE2.[91]

Milankovitch cycles

It has been hypothesised that the Cenomanian-Turonian boundary event occurred during a period of very low variability in Earth's insolation, which has been theorised to be the result of coincident nodes in all orbital parameters. Barring chaotic perturbations in Earth's and Mars' orbits, the simultaneous occurrence of nodes of

obliquity on Earth occurs approximately every 2.45 million years.[92] Numerous other oceanic anoxic events occurred throughout the extremely warm greenhouse conditions of the Middle Cretaceous,[93] and it has been suggested that these Middle Cretaceous ocean anoxic events occurred cyclically in accordance with orbital cycle patterns.[92] The mid-Cenomanian Event (MCE), which occurred in the Rotalipora cushmani planktonic foraminifer biozone, has been argued to be another example supporting this hypothesis of regular oceanic anoxic events governed by Milankovitch cycles.[93] The MCE took place approximately 2.4 million years before the Cenomanian-Turonian oceanic anoxic event, roughly at the time when an anoxic event would be expected to occur given such a cycle.[92] Geochemical evidence from a sediment core in the Tarfaya Basin is indicative of the main positive carbon isotope excursion occurring during a prolonged eccentricity minimum. Carbon isotope shifts smaller in scale observed in this core likely reflected variability in obliquity.[94] Ocean Drilling Program Site 1138 in the Kerguelen Plateau yields evidence of a 20,000 to 70,000 year periodicity in changes in sedimentation, suggesting that either obliquity or precession governed the large-scale burial of organic carbon.[95] Within the OAE2 positive δ13C excursion, short eccentricity scale carbon isotope variability is documented in a significantly expanded OAE2 interval from southern Tibet;[17] periodic negative δ13C excursions paced by the short eccentricity cycle are easily detectable in southwestern Utah too.[96]

Enhanced phosphorus recycling

The phosphorus retention ability of seafloor sediments declined during OAE2,

nitrogen fixing bacteria, increasing the availability of yet another limiting nutrient and supercharging primary productivity through nitrogen fixation.[100] The ratio of bioavailable nitrogen to bioavailable phosphorus, which is 16:1 in the present, fell precipitously as the ocean transitioned from being oxic and nitrate-dominated to anoxic and ammonium-dominated.[58] A potent feedback loop of nitrogen fixation, productivity, deoxygenation, nitrogen removal, and phosphorus recycling was created.[9] Bacterial hopanoids indicate populations of nitrogen fixing cyanobacteria were high during OAE2, providing a rich supply of nitrates and nitrites.[101] Negative δ15N values reveal the dominance of ammonium through regenerative nutrient loops in the proto-North Atlantic.[102]

Decreased sulphide oxidation

In the present day, sulphidic waters are generally prevented from spreading throughout the water column by the oxidation of sulphide with nitrate. However, during OAE2, the inventory of seawater nitrate was lower, meaning that chemolithoautotrophic oxidation of sulphides with nitrates was inefficient at preventing the spread of euxinia.[103]

Sea level rise

A marine transgression in the latest Cenomanian resulted in an increase in average water depth, causing seawater to become less eutrophic in shallow, epicontinental seas. Turnovers in marine biota in such epicontinental seas have been suggested to be driven more so by changes in water depth rather than anoxia.[104] Sea level rise also contributed to anoxia by transporting terrestrial plant matter from inundated lands seaward, providing an abundant source of sustenance for eutrophicating microorganisms.[105]

Geological effects

Phosphate deposition

A phosphogenic event occurred in the Bohemian Cretaceous Basin during the peak of oceanic anoxia. Phosphorus liberation in the pore water environment, several centimetres below the interface between seafloor sediments and the water column, enabled the precipitation of phosphate through biological mediation by microorganisms.[106]

Increase in weathering

silicate weathering increased over the course of OAE2. Because of its effectiveness as a carbon sink on geologic timescales, the uptick in sequestration of carbon dioxide by the lithosphere may have helped to stabilise global temperatures after global temperatures soared.[107] Particularly so at high latitudes, where the increase in weatherability was very pronounced.[108]

Biotic effects

Changes in oceanic biodiversity and its implications

The event brought about the extinction of the

ichthyosaurs. Coracoids of Maastrichtian age were once interpreted by some authors as belonging to ichthyosaurs, but these have since been interpreted as plesiosaur elements instead.[109]

Although the cause is still uncertain, the result starved the Earth's oceans of oxygen for nearly half a million years, causing the extinction of approximately 27 percent of

oligotrophy and ocean warmth in an environment with short spikes of productivity followed by long periods of low fertility.[113] A study performed in the Cenomanian-Turonian boundary of Wunstorf, Germany, reveal the uncharacteristic dominance of a calcareous nannofossil species, Watznaueria, present during the event. Unlike the Biscutum species, which prefer mesotrophic conditions and were generally the dominant species before and after the C/T boundary event; Watznaueria species prefer warm, oligotrophic conditions.[114] In the Ohaba-Ponor section in Romania, the presence of Watznaueria barnesae indicates warm conditions, while the abundances of Biscutum constans, Zeugrhabdotus erectus, and Eprolithus floralis peak during cool intervals.[113] Sites in Colorado, England, France, and Sicily show an inverse relationship between atmospheric carbon dioxide levels and the size of calcareous nannoplankton.[115] Radiolarians also suffered heavy losses in OAE2, one of their highest diversity losses in the Cretaceous.[116] Rudist bivalves suffered high extinction rates combined with low origination rates during OAE2.[117]

The diversity of

trace fossils sharply plummeted during the beginning of the Cenomanian-Turonian boundary event. The recovery interval after the anoxic event's conclusion features an abundance of Planolites and is characterised overall by a high degree of bioturbation.[118]

At the time, there were also peak abundances of the green algal groups Botryococcus and prasinophytes, coincident with pelagic sedimentation. The abundances of these algal groups are strongly related to the increase of both the oxygen deficiency in the water column and the total content of organic carbon. The evidence from these algal groups suggest that there were episodes of halocline stratification of the water column during the time. A species of freshwater dinocyst—Bosedinia—was also found in the rocks dated to the time and these suggest that the oceans had reduced salinity.[119][120]

Changes in terrestrial biodiversity

No major change in terrestrial ecosystems is known to have been synchronous with the marine transgression associated with OAE2, although the loss of freshwater

extirpation of some metatherians and brackish water vertebrates is associated with the later marine regression following OAE2 in the Turonian. Whatever the nature and magnitude of terrestrial extinctions at or near the Cenomanian-Turonian boundary was, it was most likely caused mainly by other factors than eustatic sea level fluctuations.[121] The effect of the ecological crisis on terrestrial plants has been concluded to have been inconsequential, in contrast to extinction events driven by terrestrial large igneous provinces.[122]

See also

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Further reading
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