Introduction to climate dynamics and climate modelling

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5.3.3 Cenozoic climate

Over the last 65 million years, the CO₂ concentration has gradually decreased from more than 1000 ppmv (part per million in volume) during the Paleocene and the beginning of the Eocene epochs to less than 300 ppmv during the Pleistocene. This long-term decrease is partly due to volcanic emissions, which were particularly large during the Paleocene and Eocene but which have diminished since then, and to changes in the rate of weathering of silicate rocks. The decline in the CO₂ concentration is associated with a cooling from the warm conditions of the early Eocene climatic optimum between 52 and 50 million years ago (Fig. 5.13). This shift is often referred to as a transition from a greenhouse climate to an icehouse, in which ice sheets are present over Antarctica (starting around 35 million years) and over Greenland (starting around 3 million years ago).

The climate reconstructions for this epoch are often based on the oxygen isotopic composition of the shell of small marine organisms called foraminifera (Fig. 5.13). Temperature influences the ¹⁸O/¹⁶O isotopic fractionation between seawater and the carbonate ions that form the shell. For some species the temperature-fractionation relationship is well known and appears to remain stable with time. So measuring the isotopic composition of the shell remains in sediments provides estimates of past temperatures. This relationship is strictly valid only for ice free conditions since ice sheets are built from water precipitating at high latitudes which is characterised by a very low ¹⁸O relative abundance. The growth of ice sheets is thus associated with a global decrease in the amount of ¹⁶O available in the other reservoirs, in particular in the ocean. As a consequence, the signal recorded in the shell of foraminifera becomes related to a mixture of temperature and ice volume influences. In a similar way as for ¹³C (Eq. 5.4), the isotopic signal is described using the delta value $δ$ ¹⁸O defined as:

δ^{18} O = [\frac{{(^{18} O /^{16} O)}_{sample}}{{(^{18} O /^{16} O)}_{standard}} - 1] .1000

(5.5)

**Figure 5.13:** The development of the global climate over the past 65 million years based on deep-sea oxygen-isotope measurements in the shell of benthic foraminifera (i.e. foraminifera living at the bottom of the ocean). The $δ$ ¹⁸O temperature scale, on the right axis, is only valid for an ice-free ocean. It therefore applies only to the time preceding the onset of large-scale glaciation in Antarctica (about 35 million years ago, see inset in the upper left corner). Figure from Zachos et al. (2008). Reprinted by permission from Macmillan Publishers Ltd: Nature, copyright 2008.

60 million years ago, the location of the continents was quite close to that of the present-day (Fig. 5.14). However, a relatively large seaway was present between North and South America while Antarctica was still connected to South America. The uplift of Panama and the closure of the Central America seaway likely modified the circulation in the Atlantic Ocean, possibly influencing the glaciation over Greenland. More importantly, the opening, deepening and widening of the Drake Passage (between South America and Antarctica) and the Tasmanian Passage (between Australia and Antarctica) allowed the formation of an intense Antarctic Circumpolar Current that isolates Antarctica from the influence of milder mid-latitudes and increased the cooling there. Finally, the uplift of the Himalayas and the Tibetan Plateau strongly modified the monsoon circulation in these regions. Those few examples illustrate the strength of the driving force associated with the changes in boundary conditions due to plate tectonics. This role should not be underestimated.

**Figure 5.14:** Land configuration about 60 million years ago. Source http://jan.ucc.nau.edu/~rcb7/global_history.html. Reproduced with permission.

In addition to the low frequency changes described above, relatively brief events are also recorded in the geological archives. One of the most spectacular is the large meteorite impact that occurred 65 million years ago at the boundary between the Cretaceous and Tertiary periods (or K-T boundary). This cataclysm has been hypothesised to have caused the extinction of many plant and animal species, including the dinosaurs, but its climatic impact is not well known and its long-term influence is not clear. The warming during the Paleocene Eocene Thermal Maximum (PETM, 55 million years ago, see Fig. 5.13), which also had a major impact on life on Earth, is better documented. During this event which lasted less than 170 000 years, the global temperature increased by more than 5^oC in less than 10 000 years. This period is also characterised by a massive injection of carbon into the atmosphere-ocean system as recorded by variations in the $δ$ ¹³C measured in sediments. The source of these massive inputs of carbon remains uncertain. It may be related to volcanism, or to the release of the methane stored in the sediments of the continental margins. Alternatively, the methane in these regions may have been destabilised by the initial warming, resulting in a strong positive feedback.

Closer to the present, large climate fluctuations have occurred over the last 5 million years. This is not clear at the scale of Fig. 5.13, but a higher resolution plot shows fluctuations with a dominant period of 100 000 years for the last million years and 41 000 years before that (Fig. 5.15). Those periodicities are very likely related to variations in the insolation, as discussed below.

**Figure 5.15:** Benthic $δ$ ¹⁸O, which measures global ice volume and deep ocean temperature, over the last 5.3 million years as given by the average of globally distributed records. Data from Lisiecki and Raymo (2005). Source http://www.lorraine-lisiecki.com/stack.html. Reproduced with permission.