The Icehouse world - the last 34 million years
- This page is part of the topic Deep time
The first continental-scale ice sheets formed on Antarctica close to the Eocene-Oligocene boundary around 34 Ma), and are physically recorded in strata from Prydz Bay, McMurdo Sound and Seymour Island (Francis et al., 2008b). Deep-sea isotope data suggest they were similar in size to that of today (Miller et al., 2008). According to DeConto and Pollard (2003) the development of these first Antarctic ice sheets was triggered in an interval when the Earth’s orbit of the Sun favoured cool summers as atmospheric CO2 levels declined below a critical threshold (~2.8 times pre-industrial). This decline has been documented (Pagani et al., 2005; Siegert et al., 2008) and has been ascribed to reduced CO2 outgassing from ocean ridges, volcanoes and metamorphic belts and increased carbon burial (Pearson and Palmer, 2000), dropping global temperatures from at least 6 to around 4ºC higher than today (Figure 3.5 and Figure 3.6). The fall in CO2 levels at this time is also reflected in a 1-km drop in the calcium compensation depth in the tropical Pacific Ocean (Coxall et al., 2005).
The pulsating style of Antarctic glaciation in Oligocene-early Miocene times is best recorded from 1,500 m of nearshore marine sediments drilled off Cape Roberts in the southwest Ross Sea (Figure 3.1 and Figure 3.7) and in the middle to late Oligocene Polonez Cove Formation, exposed on south-eastern King George Island, South Shetland Islands (Troedson and Smellie, 2002). At Cape Roberts, sediments resulting from 55 glacial-interglacial cycles have accumulated close to sea level on the subsiding margin of the Victoria Land Basin, spanning the period from 33 to 17 Ma (Barrett, 2007). The site was close to the edge of the continental ice sheet and recorded its cyclic expansion and contraction on Milankovitch frequencies (41 ka and 100 ka, Naish et al., 2001; Naish et al., 2008b; Huybrechts, 2009; Naish et al., 2009; Pollard and DeConto, 2009). The changes in sediment type that characterize the cycles - from glacial deposits through nearshore sand and offshore mud to sand again - indicate sea level changes on a scale of tens of metres (Dunbar et al., 2008). Despite episodes of extensive ice cover palynological studies indicate a coastal vegetation ranging from low woodland (Nothofagus) to tundra persisting through the Oligocene with a slight cooling in early Miocene times (Prebble et al., 2006).
Recent studies of an ancient glacial landscape in the Olympus Range on the inland edge of the McMurdo Dry Valleys have revealed warm-based glacial deposits overlain by ridges of cold-based gravelly debris, each bearing volcanic ash beds ~14 and ~13.6 Ma respectively. The landscape has changed little since that time, the lack of alteration being ascribed to a persistent frozen state from that time on, recording a sharp Middle Miocene cooling (Lewis et al., 2007). Ash-bearing proglacial lake beds, also dated at ~14 Ma, include an ostracod fauna, Nothofagus pollen and beds of moss, and are regarded as possibly the last vestiges of this fauna and flora in the region (Ashworth et al., 2007; Lewis et al.; 2008). A well preserved flora of in situ Nothofagus dwarf shrubs, mosses and cushion plants, along with the remains of beetles, molluscs, fish and parts of flies, from the Sirius Group, Oliver Bluffs in the Dominion Range, Transantarctic Mountains also provide good evidence for tundra conditions in this region only 300 miles from the South Pole. The fossils are preserved in a glacio-fluvial-lacustrine-palaeosol layer that represents a warmer interval that prompted glacial retreat between colder intervals during which glaciers were present at that site (Francis and Hill, 1996; Ashworth and Cantrill, 2004). Unfortunately an undisputed age for these deposits is not available.
The sharp cooling in the Middle Miocene has long been known from deep-sea isotopic studies (Shackleton and Kennett, 1975), and was probably caused by the growing thermal isolation of Antarctica and related intensification of the Antarctic Circumpolar Current described above, accompanied by a drop in atmospheric CO2 (Shevenell et al., 1996). This thickened the ice sheet to more or less its modern configuration, which is thought to have persisted through the early Pliocene warming from 5 Ma to 3 Ma (Kennett and Hodell, 1993; Barrett, 1996; McKay et al., 2008).
During the Pliocene, mean global temperatures were 2-3ºC above pre-industrial values (Figure 3.6) with CO2 values less than 400 ppm and sea levels 15-25m above modern levels (Raymo et al., 1996; Jansen et al., 2007, and references therein). The Antarctic margin also records Pliocene temperatures several degrees warmer than today in diatom-bearing coastal sediments (Harwood et al., 2000) and offshore cores (Whitehead et al., 2005).
A particularly instructive record comes from the ANDRILL McMurdo Ice Shelf site, where over 1,200 m of strata dating back to 13 Ma were cored from a deep-water basin south of Ross Island (Naish et al., 2007; 2008a,b, 2009). The record comprises many cycles of sedimentation, alternating between deposition beneath grounded ice (diamictite) and ice shelf ice (mudstone) in the last million years. However, from 1 to ~5 Ma depositional environments ranging from grounded ice (diamictite) to open water (diatomite) (Figure 3.8). The diatomite beds show the drill site to have been essentially ice-free at this time, i.e. no McMurdo Ice Shelf and hence no Ross Ice Shelf. The lack of buttressing from the loss of the Ross Ice Shelf in turn implies a much reduced West Antarctic Ice Sheet (Mercer, 1978; Dupont and Alley, 2005a). However geomorphological evidence, and the antiquity of high level surfaces in the Transantarctic Mountains (Sugden et al., 1993), along with modelling Pliocene ice sheets (Hill et al., In Press), supports the persistence of an ice sheet in the East Antarctic interior through this period. A new ice sheet model by Pollard and DeConto (2009) confirms the persistence of East Antarctic ice and the disappearance of the West Antarctic ice sheet during warm Pliocene times and even as recently as the MIS 31 interglacial stage a little over a million years ago (Figure 3.8).
Global cooling from around 3 Ma onwards (Ravelo et al., 2004) led to the first big ice sheets on North America and NW Europe around 2.6 Ma (Shackleton et al., 1984), enhancing the Earth’s climate response to orbital forcing with a 40,000 year cyclicity, and taking us to the Earth’s present intense “ice house” state. For the last million years (Figure 3.8) this has alternated between (i) longer (100,000 years) glacial cycles, when much of the Northern Hemisphere was ice-covered, global average temperature was around 5ºC colder, and sea level was approximately 120 m lower than today, and (ii) much shorter warm interglacial cycles like that of the last ~10,000 years, with sea levels near or slightly above those of the present.
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