large d13C shift occurred slightly later in surface waters (Figure 5.3).
No such deep oceanic mass extinction and isotopic excursion has yet been discovered at any other time in the Cenozoic. This was an unusual, if not unique, event. Why did the event occur at about 55 Ma, early in the Cenozoic, rather than later? During the early Paleogene, different global geography and climate (Kennett, 1977; Haq, 1981; Hay, 1989) combined to make ocean circulation distinct from that of modern and, indeed, Neogene oceans (Kennett, 1977; Benson, 1979; Kennett and Stott, 1990a). Much evidence exists for relatively warm climates in the Antarctic region during the early Cenozoic (Kennett and Barker, 1990). Oxygen isotopic data suggest average Antarctic Ocean surface water temperatures of ~14°C during the Late Paleocene. Decreased meridional thermal gradients led to a decrease in global zonal wind intensity (Janecek and Rea, 1983; Hovan and Rea, 1992). Clay mineral assemblages in offshore sequences derived from the Antarctic continent were formed predominantly by chemical weathering under conditions of relative continental warmth and humidity (Robert and Kennett, 1992). Extensive coastal cool temperate rain forests dominated by Nothofagus indicate high continental rainfall and a lack of perglacial conditions at sea-level (Case, 1988). Ice-rafted sediments are absent, as is other evidence for continental cryosphere of any extent (Kennett and Barker, 1990), although montane glaciation seems probable. The Antarctic Ocean was dominated by calcareous planktonic microfossil assemblages of high diversity rather than siliceous forms (Kennett, 1977). Faunas were cool to warm temperate in character. Deep waters in the global ocean were warm, averaging 10 to 12°C compared with ~2°C in the modern ocean (Shackleton and Kennett, 1975; Stott et al., 1990). The Earth was clearly in a ''greenhouse" mode—a condition that appears to have been much exaggerated during the terminal Paleocene isotopic excursion. Relatively high precipitation in the Antarctic region at this time is inferred to have contributed to a large reduction in deep water production at high latitudes (Kennett and Stott, 1991).
At the same time, the extensive mid-latitude Tethys Seaway north of Africa was a likely location for the production of large volumes of warm saline deep waters (Kennett and Stott, 1990). Tectonic reconstructions (Dercourt et al., 1986) show that the Tethys Seaway in the early Cenozoic contained extensive shallow carbonate platforms with dolomites and evaporitic sediments. During the excursion, these various factors combined to form, through positive feedback responses, an extreme case of the Proteus Ocean, an ocean dominated by middle-latitude deep water production (Kennett and Stott, 1990, 1991). Climate model studies (Barron, 1987; Covey and Barron, 1988) suggest that large-scale meridional heat transport became more effective via the deep oceans relative to the atmosphere (Hovan and Rea, 1992).
The forcing mechanism of ocean warming and associated faunal turnover at the end of the Paleocene is not yet known. Rea et al. (1990) have suggested that the abruptness of environmental changes and the associated mass extinction were possibly triggered by rapid input of CO2 into the atmosphere from volcanism and/or hydrothermal activity that was extensive over the Paleocene-Eocene transition. The warming of the oceans would have been the most obvious effect of enhanced greenhouse forcing resulting from this. Whether or not this would have been sufficiently rapid and large to explain the rapid rise in temperatures associated with the extinctions at 55 Ma remains to be tested. The triggering mechanism for the rapid climate change at the end of the Paleocene remains unknown.
In one attempt to test whether CO2 might be implicated in the oceanic warming, Stott (1992) presented evidence that the Paleocene ocean-atmosphere system was indeed associated with higher levels of CO2 compared to the present time. However, on the basis of the same data it appears that the extinction interval was actually associated with lower oceanic CO2, not higher. How could an abrupt warming at the end of the Paleocene be associated with lower oceanic CO2? The answer may lie in the way the ocean and atmosphere cycle CO2. The problem is that the solubility of CO2 in seawater decreases with increasing temperature. The exchange of CO2 between the ocean and atmosphere was further complicated by changes in atmospheric circulation occurring at that time, which would have affected turbulence of the mixed layer ocean. This, together with changes in the biological pump (photosynthesis), constitute factors that are not yet well constrained for the Paleocene-Eocene. However, it is evident that with the high sea surface temperatures at the end of the Paleocene, particularly in regions of normal deep water advection (e.g., high latitudes), the oceans were probably less efficient in taking up CO2 from the atmosphere.
A conceptual model of the possible chain of environmental events related to the terminal Paleocene mass extinction in the deep-sea is shown in Figure 5.5. The extinction occurred in less than ~3000 yr and was associated with global deep-sea warming of similar rapidity (Figure 5.3). Both benthic foraminifera and ostracoda were se-