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Setting

The configuration of oceans and continents was somewhat different during the Eocene. The Panama Isthmus did not yet connect North and South America, which allowed circulation between the Pacific and Atlantic oceans. Further, the Drake Passage was closed, perhaps preventing the thermal isolation of Antarctica. Although various proxies for past atmospheric CO₂ levels in the Eocene do not agree in absolute terms, all suggest that levels then were much higher than at present. In any case, there were no significant ice sheets during this time.^[5]

Earth surface temperatures increased by about 6°C from the late Paleocene through the early Eocene, culminating in the "Early Eocene Climatic Optimum" (EECO).^[5] Superimposed on this long-term, gradual warming were at least two (and likely more) "hyperthermals". These can be defined as geologically brief (<200,000 year) events characterized by rapid global warming, major changes in the environment, and massive carbon addition. Of these, the PETM was the most extreme and perhaps the first (at least within the Cenozoic). Another hyperthermal clearly occurred at approximately 53.7 Ma, and is now called ETM-2 (also referred to as H-1, or the Elmo event). However, additional hyperthermals likely occurred at about 53.6 Ma (H-2), 53.3 (I-1), 53.2 (I-2) and 52.8 Ma (informally called K, X or ETM-3). The number, nomenclature, absolute ages, and relative global impact of the Eocene hyperthermals are the source of considerable current research. Whether they only occurred during the long-term warming, and whether they are causally related to apparently similar events in older intervals of the geological record (e.g., the Toarcian turnover of the Jurassic) are open issues.

[edit] Evidence for global warming

Average global temperatures increased by ~6°C (11°F) within about 20,000 years. This is based on several lines of evidence. There is a prominent (>1‰) negative excursion in the δ¹⁸O of foraminifera shells, both those made in surface and deep ocean water. Because there was a paucity of continental ice in the early Paleogene, the shift in δ¹⁸O very likely signifies a rise in ocean temperature.^[6] The temperature rise is also supported by analyses of foraminifera Mg/Ca and ratios of certain organic compounds (TEX₈₆).

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 to 20 °C (50 to 68 °F);^[7] the surface waters of the northernmost^[8] Arctic ocean warmed, seasonally at least, enough to support tropical lifeforms^[9] requiring surface temperatures of over 22°C.^[10]

[edit] Evidence for carbon addition

Clear evidence for massive addition of ¹³C-depleted carbon at the onset of the PETM comes from two observations. First, a prominent negative excursion in the carbon isotope composition (δ¹³C) of carbon-bearing phases characterizes the PETM in numerous widespread locations from a range of environments. Second, carbonate dissolution marks the PETM in sections from the deep-sea.

The total mass of carbon injected to the ocean and atmosphere during the PETM remains the source of debate. In theory, it can be estimated from the magnitude of the δ¹³C excursion, the amount of carbonate dissolution on the seafloor, or ideally both. However, the shift in the δ¹³C across the PETM depends on the location and the carbon-bearing phase analyzed. In some records of bulk carbonate, it is about 2‰; in some records of terrestrial carbonate or organic matter it exceeds 6‰.^[11] Carbonate dissolution also varies throughout different ocean basins. It is extreme in parts of the north and central Atlantic Ocean but far less pronounced in the Pacific Ocean. With available information, estimates of the carbon addition range from about 2500 to over 6800 gigatons ^[12]

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.^[13] 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.^[14] 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 1,000 years, and separated by about 20,000 years. The models diverge most in their estimate of the recovery time, which ranges from 150,000^[13] to 30,000^[14] years. There is other evidence to suggest that warming predated the δ¹³C excursion by some 3,000 years.^[15]

[edit] Effects

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.^[16] This would have resulted in the largely isolated Arctic ocean's taking on a more freshwater character as northern hemisphere rainfall was channelled towards it.^[16]

[edit] Sea level

Despite the global lack of ice, the sea level would have risen due to thermal expansion.^[10] 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.^[10]

[edit] Circulation

At the start of the PETM, the ocean circulation patterns changed radically in the course of under 5,000 years.^[17] Global-scale current directions reversed; for example, deep water in the Atlantic flowed from north to south instead of the usual south to north.^[17] This "backwards" flow persisted for 40,000 years.^[17] Such a change would transport warm water to the deep oceans, enhancing further warming.^[17]

[edit] Lysocline

The lysocline marks the depth at which carbonate starts to dissolve (above the lysocline, carbonate is oversaturated): today, this is at about 4 km, 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,^[18] 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).^[19] 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.^[3] In parts of the southeast Atlantic, the lysocline rose by 2 km in just a few thousand years.^[3]