Paleocene–Eocene Thermal Maximum

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Climate change during the last 65 million years. The Paleocene-Eocene Thermal Maximum is labeled PETM and is likely to be understated by a factor of 2 or more due to coarse sampling and averaging in this data set.
Climate change during the last 65 million years. The Paleocene-Eocene Thermal Maximum is labeled PETM and is likely to be understated by a factor of 2 or more due to coarse sampling and averaging in this data set.

The Paleocene/Eocene boundary, 55.8 million years ago, was marked by the most rapid and significant climatic perturbation of the Cenozoic Era. A sudden global warming event, leading to the Paleocene-Eocene Thermal Maximum (PETM, alternatively "Eocene thermal maximum 1" (ETM1), and formerly known as the "Initial Eocene" or "Late Paleocene Thermal Maximum",[1] (IETM/LPTM)), is associated with changes in oceanic and atmospheric circulation, the extinction of numerous deep-sea benthic foraminifera, and a major turnover in mammalian life on land which is coincident with the emergence of many of today's major mammalian orders.

The event saw global temperatures rise by around 6 °C over 20,000 years, with a corresponding rise in sea level as the whole of the oceans warmed.[2] Atmospheric carbon dioxide (CO2) concentrations rose, causing a shallowing of the lysocline. Regional deep water anoxia may have played a part in marine extinctions. The event is linked to a negative excursion in the δ13C isotope record, which occurs in two short (~1,000 year) pulses. These probably represent degassing of clathrates ("methane ice" deposits), which accentuated a pre-existing warming trend. The release of these clathrates, and ultimately the event itself, may have been triggered by a range of causes. Evidence currently seems to favour an increase in volcanic activity as the main perpetrator.


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Holocene begins 11.5 ka ago
An approximate timescale of key Cenozoic events.
Axis scale: Ma before present.

Contents

[edit] Setting

The Paleocene-Eocene Thermal Maximum lasted around 20,000 years, and was superimposed on a 6 million year period of more gradual global warming,[6] peaking later in the Eocene at the "Eocene climatic optimum". Other "hyperthermal" events can be recognised during this period of cooling, including the Elmo event (ETM2). During these events, of which the PETM was by far the most severe, around 1,500 to 2,000 gigatons of carbon were released into the ocean/atmosphere system over the course of 1,000 years. This rate of carbon addition almost equals the rate at which carbon is being released into the atmosphere today through anthropogenic activity.

The globe was subtly different during the Eocene. The Panama Isthmus did not yet connect North and South America, allowing circulation between the Pacific and Atlantic oceans. Further, the Drake Passage was shut, preventing the thermal isolation of Antarctica. This, combined with higher CO2 levels, meant that there were no significant ice sheets - the globe was essentially ice free.[6]

[edit] Evidence

Our strongest evidence for climate change comes from a global, synchronous and uniform excursion in the δ13C record, of −2-3 ‰. Its magnitude is larger in terrestrial environments.[7] This excursion implies the release of large amounts of 12C into the ocean and atmosphere, and implies the release of at least 6,800 Pg C.[8]

The timing of the PETM δ13C excursion has been calculated in two complementary ways. The iconic core covering this time period is the ODP's Core 690, and the timing is based exclusively on this core's record. The original timing was calculated assuming a constant sedimentation rate.[9] This model was improved using the assumption that 3He flux is constant; this cosmogenic nuclide is produced at a (roughly) constant rate by the sun, and there is little reason to assume large fluctuations in the solar wind across this short time period.[10] Both models have their failings, but agree on a few points. Importantly, they both detect two steps in the drop of δ13C, each lasting about 1000 years, and separated by about 20,000 years. The models diverge most in their estimate of the recovery time, which ranges from 150,000[9] to 30,000[10] years. There is other evidence to suggest that warming predated the δ13C excursion by some 3,000 years.[11]

[edit] Effects

[edit] Climate

Average global temperatures increased by ~6 °C in the space of 20,000 years. This is based on Mg/Ca and δ18O values of forams. δ18O is a more useful proxy for palæotemperature during the Eocene, as the lack of ice makes it safe to assume that the oceans' δ18O signature is constant.[12] Due to the positive feedback effect of melting ice reducing albedo, temperature increases would have been greatest at the poles, which reached an average annual temperature of 10-20 °C;[13] the surface waters of the northernmost[14] Arctic ocean warmed, seasonally at least, enough to support tropical lifeforms[15] requiring surface temperatures of over 22°C.[16]

The climate would also have become much wetter, with the increase in evaporation rates peaking in the tropics. Deuterium isotopes reveal that much more of this moisture was transported polewards than normal.[17] This would have resulted in the largely isolated Arctic ocean taking a more freshwater character as northern hemisphere rainfall was channelled towards it.[17]

[edit] Sea level

Despite the global lack of ice, the sea level would have risen due to thermal expansion.[16] Evidence for this can be found in the shifting palynomorph assemblages of the Arctic ocean, which reflect a relative decrease in terrestrial organic material compared to marine organic matter.[16]

[edit] Circulation

At the start of the PETM, the ocean circulation patterns changed radically in the course of under 5,000 years.[18] Global-scale current directions reversed; for example, deep water in the Atlantic flowed from north to south instead of the usual south to north.[18] This "backwards" flow persisted for 40,000 years.[18] Such a change would transport warm water to the deep oceans, enhancing further warming.[18]

[edit] Lysocline

The lysocline marks the depth at which carbonate spontaneously dissolves in the oceans: today, this is at about 4km, comparable to the median depth of the oceans. This depth depends on (among other things) temperature and the amount of CO2 dissolved in the ocean. Adding CO2 initially shallows the lysocline,[19] resulting in the dissolution of deep water carbonates. This deep-water acidification can be observed in ocean cores, which show (where bioturbation has not destroyed the signal) an abrupt change from grey carbonate ooze to red clays (followed by a gradual grading back to grey).[20] It is far more pronounced in north Atlantic cores than elsewhere, suggesting that acidification was more concentrated here, related to a greater rise in the level of the lysocline.[4] In parts of the south east Atlantic, the lysocline rose by 2 km in just a few thousand years.[4]

[edit] Anoxia?

In parts of the oceans, especially the north Atlantic Ocean, bioturbation is absent. This may be due to bottom-water anoxia, or by changing ocean circulation patterns changing the temperatures of the bottom water. However, many ocean basins remain bioturbated through the PETM.[4]

[edit] Life

The PETM is accompanied by a mass extinction of 35-50% of benthic foramanifera (especially in deeper waters) over the course of ~1000 years - the group suffering more than during the dinosaur-slaying K-T extinction. Contrarily, planktonic foramanifera diversified, and dinoflagellates bloomed. Success was also enjoyed by the mammals, who radiated profusely around this time.

The deep sea extinctions are difficult to explain, as many were regional in extent (mainly affecting the north Atlantic): this means that we cannot appeal to general hypotheses such as a temperature-related reduction in oxygen availability, or increased corrosiveness due to carbonate-undersaturated deep waters. The only factor which was global in extent was an increase in temperature, and it appears that the majority of the blame must rest upon its shoulders. Regional extinctions in the North Atlantic can be attributed to increased deep-sea anoxia, which could be due to the slowdown of overturning ocean currents,[8] or the release and rapid oxidation of large amounts of methane.[21][verification needed]

In shallower waters, it's undeniable that increased CO2 levels result in a decreased oceanic pH, which has a profound negative effect on corals.[22] Experiments suggest it is also very harmful to calcifying plankton.[23] However, the strong acids used to simulate the natural increase in acidity which would result from elevated CO2 concentrations may have given misleading results, and the most recent evidence is that coccolithophores (E. huxleyi at least) become more, not less, calcified and abundant in acidic waters.[24] Interestingly, no change in the distribution of calcareous nannoplankton such as the coccolithophores can be attributed to acidification during the PETM.[24] Acidification did lead to an abundance of heavily calcified algae[25] and weakly calcified forams.[26]

The increase in mammalian abundance is intriguing. There is no evidence of any increased extinction rate among the terrestrial biota. Increased CO2 levels may have promoted dwarfing[27] - which may (perhaps?) have encouraged speciation. Many major mammalian orders, including the Artiodactyla, horses and primates, appeared as if from nowhere, and spread across the globe, 13,000 to 22,000 years after the initiation of the PETM.[27]

[edit] Possible causes

Discriminating between different causes of the PETM is difficult. Temperatures were rising globally at a steady pace, and a mechanism must be invoked to produce a sudden spike - which may have been accentuated by positive feedbacks. Our biggest aid in disentangling these factors comes from a consideration of the carbon isotope mass balance. We know the entire exogenic carbon cycle (i.e. the carbon contained within the oceans and atmosphere, which can change on short timescales) underwent a −2-3 ‰ perturbation in δ13C, and by considering the isotopic signatures of other carbon reserves, can consider what mass of the reserve would be necessary to produce this effect. The assumption underpinning this approach is that the mass of exogenic carbon was the same in the Palæogene as it is today - something which is very hard to confirm.

[edit] Volcanic activity

In order to balance the mass of carbon and produce the observed δ13C value, at least 1,500 Gt of carbon would have to be degassed from the mantle via volcanoes over the course of the two 1,000 year steps. To put this in perspective, this is about 200 times the background rate of degassing for the rest of the Palæogene. There is no indication that such a burst of volcanic activity has occurred at any point in Earth's history. However, substantial volcanism had been active in East Greenland for around the preceding million years or so, but this struggles to explain the rapidity of the PETM. Even if the bulk of the 1,500 Gt of carbon was released in a single pulse, further feedbacks would be necessary to produce the observed isotopic excursion.

On the other hand, there are suggestions that surges of activity occurred in the later stages of the volcanism and associated continental rifting; intrusions of hot magma into carbon-rich sediments may have triggered the degassing of methane.[28] Further phases of volcanic activity could have triggered the release of more methane, and caused other early Eocene warm events such as the ETM2.[8]

[edit] Comet impact

A briefly popular theory held that a 12C-rich comet struck the earth and initiated the warming event.[29] Even allowing for feedback processes, this would require at least 100 Gt of extra-terrestrial carbon[29] - such a catastrophic impact should have left its mark on the globe. Unfortunately, the evidence put forwards does not stand up to scrutiny. An unusual 9 m thick clay layer supposedly formed soon after the impact, containing unusual amounts of magnetism. But it formed too slowly for these magnetic particles to be a result of the comet's impact - [11] it turns out they were created by bacteria.[30] Further, an iridium anomaly - often an indicator of extra-terrestrial impact - observed in Spain is far too small to denote a comet impact.

[edit] Burning of peat

This combustion of prodigal quantities of peat was once postulated, but in order to produce the δ13C excursion observed, over 90% of the Earth's biomass would have to be combusted. Since plants in fact grew more voraciously during the period of the PETM, this theory has been discounted.

[edit] Orbital forcing

The presence of later (smaller) warming events of a global scale, such as the Elmo horizon (aka ETM2), has led the the hypothesis that the events repeat on a regular basis, driven by maxima in the 400,000 and 100,000 year eccentricity cycles in the Earth's orbit. The orbital increase in insolation (and thus temperature) would force the system over a threshold and unleash positive feedbacks.[31]

[edit] Methane release

None of the above causes are alone sufficient to cause the carbon isotope excursion or warming observed at the PETM. The most obvious feedback mechanism that could amplify the initial perturbation is that of clathrates. At certain temperature and pressure conditions, methane - which is being produced continually by decomposing microbes in sea bottom sediments - is stable in a complex with water, which forms ice-like cages trapping the methane in solid form. As temperature rises, so the pressure at which this clathrate configuration is stable falls - so shallow clathrates dissociate, releasing methane gas to make its way into the atmosphere. Since biogenic clathrates have a δ13C signature of −60 ‰ (inorganic clathrates are the still rather large −40 ‰), relatively small masses can produce large δ13C excursions. Further, methane is a potent greenhouse gas - as it is released into the atmosphere, so it causes warming, and as the ocean transports this to the bottom sediments, it destabilises more clathrates. It would take around 2,300 years for an increased temperature to diffuse warm the sea bed to a depth sufficient to cause clathrates' release - although the exact time frame is highly dependant on a number of poorly-constrained assumptions.[32]

In order for the clathrate hypothesis to work, the oceans must show signs of being warmer slightly before the carbon isotope excursion - because it would take some time for the methane to become mixed into the system and δ13C-reduced carbon to be returned to the deep ocean sedimentary record. Until recently, the evidence suggested that the two peaks were in fact simultaneous, weakening the support for the methane theory. But recent work has managed to detect a short gap between the initial warming and the δ13C excursion.[33] Chemical markers of surface temperature (TEX86) also indicate that warming occurred around 3,000 years before the carbon isotope excursion, but this does not seem to hold true for all cores.[11] Notably, deeper (non-surface) waters do not appear to display evidence of this time gap.[34]

Analysis of these records reveals another interesting fact: plantktonic (floating) forams record the shift to lighter isotope values earlier than benthic (bottom dwelling) forams. The lighter (lower δ13C) methanogenic carbon can only be incorporated into the forams' shells after it has been oxidised. A gradual release of the gas would allow it to be oxidised in the deep ocean, which would make benthic forams' tests lighter earlier. The fact that the planktonic forams are the first to show the signal suggests that the methane was released so rapidly that its oxidation used up all the oxygen at depth in the water column, allowing some methane to reach the atmosphere unoxidised, where atmospheric oxygen would react with it. This observation also allows us to constrain the duration of methane release to under around 10,000 years.[33]

[edit] Ocean circulation

The large scale patterns of ocean circulation are important when considering how heat was transported through the oceans. Our understanding of these patterns is still in a preliminary stage. Models show that there are possible mechanisms to quickly transport heat to the shallow, clathrate-containing ocean shelves, given the right bathymetric profile, but the models cannot yet match the distribution of data we observe.[35]

[edit] Recovery

The δ13C record records a recovery time of around 150,000[9] to 30,000[10] years, relatively rapid compared to the residence time of carbon in the modern atmosphere (100-200 thousand years). A satisfactory explanation of this rapid recovery must incorporate a feedback system.[36]

The most likely method of recovery invokes an increase in biological productivity, transporting carbon to the deep ocean. This would be assisted by higher global temperatures and CO2 levels, as well as an increased nutrient supply (which would result from higher continental weathering due to higher temperatures and rainfall; volcanics may have provided further nutrients). Evidence for higher biological productivity comes in the form of biogenic Barium.[36] However, this proxy may instead reflect the addition of Barium dissolved in methane.[37] However, diversifications suggest that productivity increased in near-shore environments, which would have been warm and fertilised by run-off - outweighing the reduction in productivity in the deep oceans.[26]

[edit] See also

[edit] Notes

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  2. ^ Kennett, J.P. & Stott, L.D. 1991. Abrupt deep-sea warming, palaeoceanographic changes and benthic extinctions at the end of the Palaeocene. Nature, 353: 225-229
  3. ^ Retallack, G.J. (1997). "Neogene Expansion of the North American Prairie". PALAIOS 12 (4): 380-390. 
  4. ^ a b c d Zachos, J.C.; Kump, L.R. (2005). "Carbon cycle feedbacks and the initiation of Antarctic glaciation in the earliest Oligocene". Global and Planetary Change 47 (1): 51-66. 
  5. ^ Krijgsman, W.; Garcés, M.; Langereis, C.G.; Daams, R.; Van Dam, J.; Van Der Meulen, A.J.; Agustí, J.; Cabrera, L. (1996). "A new chronology for the middle to late Miocene continental record in Spain". Earth and Planetary Science Letters 142 (3-4): 367-380. 
  6. ^ a b Zachos, J.C.; Dickens, G.R.; Zeebe, R.E. (2008). "An early Cenozoic perspective on greenhouse warming and carbon-cycle dynamics". Nature 451 (7176): 279-83. doi:10.1038/nature06588. 
  7. ^ Norris, R.D.; Röhl, U. (1999). "Carbon cycling and chronology of climate warming during the Palaeocene/Eocene transition". Nature 401 (6755): 775-778. doi:10.1038/44545. 
  8. ^ a b c Panchuk, K.; Ridgwell, A.; Kump, L.R. (2008). "Sedimentary response to Paleocene-Eocene Thermal Maximum carbon release: A model-data comparison". Geology 36 (4): 315-318. doi:10.1130/G24474A.1. 
  9. ^ a b c Rohl, U.; Bralower, T.J.; Norris, R.D.; Wefer, G. (2000). "New chronology for the late Paleocene thermal maximum and its environmental implications". Geology 28 (10): 927-930. doi:10.1130/0091-7613(2000)28<927:NCFTLP>2.0.CO;2. 
  10. ^ a b c Farley, K.A.; Eltgroth, S.F. (2003). "An alternative age model for the Paleocene--Eocene thermal maximum using extraterrestrial 3He". Earth and Planetary Science Letters 208 (3-4): 135-148. doi:10.1016/S0012-821X(03)00017-7. 
  11. ^ a b c Sluijs, A.; Brinkhuis, H.; Schouten, S.; Bohaty, S.M.; John, C.M.; Zachos, J.C.; Reichart, G.J.; Sinninghe Damste, J.S.; Crouch, E.M.; Dickens, G.R. (2007). "Environmental precursors to rapid light carbon injection at the Palaeocene/Eocene boundary.". Nature 450 (7173): 1218-21. doi:10.1038/nature06400. 
  12. ^ Thomas, E.; Shackleton, N.J. (1996). "The Paleocene-Eocene benthic foraminiferal extinction and stable isotope anomalies". Geological Society London Special Publications 101 (1): 401. doi:10.1144/GSL.SP.1996.101.01.20. 
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  14. ^ Drill cores were recovered from the Lomonosov ridge, presently at 87°N
  15. ^ the dinoflagellates Apectodinium augustum
  16. ^ a b c Sluijs, A.; Schouten, S.; Pagani, M.; Woltering, M.; Brinkhuis, H.; Damsté, J.S.S.; Dickens, G.R.; Huber, M.; Reichart, G.J.; Stein, R.; Others, (2006). "Subtropical Arctic Ocean temperatures during the Palaeocene/Eocene thermal maximum". Nature 441 (7093): 610-613. doi:10.1038/nature04668. 
  17. ^ a b Pagani, M.; Pedentchouk, N.; Huber, M.; Sluijs, A.; Schouten, S.; Brinkhuis, H.; Sinninghe Damsté, J.S.; Dickens, G.R.; Others, (2006). "Arctic hydrology during global warming at the Palaeocene/Eocene thermal maximum". Nature 442 (7103): 671-675. doi:10.1038/nature05043. 
  18. ^ a b c d Nunes, F.; Norris, R.D. (2006). "Abrupt reversal in ocean overturning during the Palaeocene/Eocene warm period.". Nature 439 (7072): 60-3. doi:10.1038/nature04386. 
  19. ^ Dickens, G.R.; Castillo, M.M.; Walker, J.C.G. (1997). "A blast of gas in the latest Paleocene; simulating first-order effects of massive dissociation of oceanic methane hydrate". Geology 25 (3): 259-262. doi:10.1130/0091-7613(1997)025<0259:ABOGIT>2.3.CO;2. 
  20. ^ Zachos, J.C.; Röhl, U.; Schellenberg, S.A.; Sluijs, A.; Hodell, D.A.; Kelly, D.C.; Thomas, E.; Nicolo, M.; Raffi, I.; Lourens, L.J.; et al. (2005). "Rapid Acidification of the Ocean During the Paleocene-Eocene Thermal Maximum". Science 308 (5728): 1611-1615. doi:10.1126/science.1109004. 
  21. ^ Zachos, J.C.; Dickens, G.R. (1999). "An assessment of the biogeochemical feedback response to the climatic and chemical perturbations of the LPTM". GFF 122: 188-189. 
  22. ^ Langdon, C.; Takahashi, T.; Sweeney, C.; Chipman, D.; Goddard, J.; Marubini, F.; Aceves, H.; Barnett, H.; Atkinson, M.J. (2000). "Effect of calcium carbonate saturation state on the calcification rate of an experimental coral reef". Global Biogeochemical Cycles 14 (2): 639-654. doi:10.1029/1999GB001195. 
  23. ^ Riebesell, U.; Zondervan, I.; Rost, B.; Tortell, P.D.; Zeebe, R.E.; Morel, F.M.M. (2000). "Reduced calcification of marine plankton in response to increased atmospheric CO2". Nature 407 (6802): 364-367. 
  24. ^ a b "Phytoplankton Calcification in a High-CO2 World M. Debora Iglesias-Rodriguez,1* Paul R. Halloran,2* Rosalind E. M. Rickaby,2 Ian R. Hall,3 Elena Colmenero-Hidalgo,3{dagger} John R. Gittins,1 Darryl R. H. Green,1 Toby Tyrrell,1 Samantha J. Gibbs,1 Peter von Dassow,4 Eric Rehm,5 E. Virginia Armbrust,5 Karin P. Boessenkool" . doi:10.1126/science.1154122. 
  25. ^ Bralower, T.J. (2002). "Evidence of surface water oligotrophy during the Paleocene-Eocene thermal maximum: Nannofossil assemblage data from Ocean Drilling Program Site 690, Maud Rise, Weddell Sea". Paleoceanography 17 (2): 1023. 
  26. ^ a b Kelly, D.C.; Bralower, T.J.; Zachos, J.C. (1998). "Evolutionary consequences of the latest Paleocene thermal maximum for tropical planktonic foraminifera". Palaeogeography, Palaeoclimatology, Palaeoecology 141 (1): 139-161. doi:10.1016/S0031-0182(98)00017-0. 
  27. ^ a b Gingerich, P.D. (2003). "Mammalian responses to climate change at the Paleocene-Eocene boundary: Polecat Bench record in the northern Bighorn Basin, Wyoming". Causes and Consequences of Globally Warm Climates in the Early Paleogene 369: 463. doi:10.1130/0-8137-2369-8.463. 
  28. ^ Storey, M.; Duncan, R.A.; Swisher III, C.C. (2007). "Paleocene-Eocene Thermal Maximum and the Opening of the Northeast Atlantic". Science 316 (5824): 587. doi:10.1126/science.1135274. 
  29. ^ a b Kent, D.V.; Cramer, B.S.; Lanci, L.; Wang, D.; Wright, J.D.; Van Der Voo, R. (2003). "A case for a comet impact trigger for the Paleocene/Eocene thermal maximum and carbon isotope excursion". Earth and Planetary Science Letters 211 (1-2): 13-26. doi:10.1016/S0012-821X(03)00188-2. 
  30. ^ Kopp, R.E.; Raub, T.; Schumann, D.; Vali, H.; Smirnov, A.V.; Kirschvink, J.L. (2007). "Magnetofossil Spike During The Paleocene-eocene Thermal Maximum: Ferromagnetic Resonance, Rock Magnetic, And Electron Microscopy Evidence From The Atlantic Coastal Plain Of New Jersey". Palaeoceanography 22: PA4103. doi:10.1029/2007PA001473. 
  31. ^ Lourens, L.J.; Sluijs, A.; Kroon, D.; Zachos, J.C.; Thomas, E.; Röhl, U.; Bowles, J.; Raffi, I. (2005). "Astronomical pacing of late Palaeocene to early Eocene global warming events". Nature 435 (7045): 1083-1087. doi:10.1038/nature03814. 
  32. ^ Katz, M.E.; Cramer, B.S.; Mountain, G.S.; Katz, S.; Miller, K.G. (2001). "Uncorking the bottle: What triggered the Paleocene/Eocene thermal maximum methane release". Paleoceanography 16 (6): 667. doi:10.1029/2000PA000615. 
  33. ^ a b Thomas, D.J.; Zachos, J.C.; Bralower, T.J.; Thomas, E.; Bohaty, S. (2002). "Warming the fuel for the fire: Evidence for the thermal dissociation of methane hydrate during the Paleocene-Eocene thermal maximum". Geology 30 (12): 1067-1070. doi:10.1130/0091-7613(2002)030. 
  34. ^ Tripati, A.; Elderfield, H. (2005). "Deep-Sea Temperature and Circulation Changes at the Paleocene-Eocene Thermal Maximum". Science 308 (5730): 1894-1898. doi:10.1126/science.1109202. 
  35. ^ Bice, K.L.; Marotzke, J. (2002). "Could changing ocean circulation have destabilized methane hydrate at the Paleocene/Eocene boundary". Paleoceanography 17 (2): 1018. doi:10.1029/2001PA000678. 
  36. ^ a b Bains, S.; Norris, R.D.; Corfield, R.M.; Faul, K.L. (2000). "Termination of global warmth at the Palaeocene/Eocene boundary through productivity feedback.". Nature 407 (6801): 171-4. doi:10.1038/35025035. 
  37. ^ Dickens, Fewless, thomas, brawoler (2003). "Excess barite accumulation during the PETM.....". GSA spec 369. 

[edit] External links