18.104.22.168 What Do the Last Glacial Maximum and the Last Deglaciation Show?
Past glacial cold periods, sometimes referred to as ‘ice ages’, provide a means for evaluating the understanding and modelling of the response of the climate system to large radiative perturbations. The most recent glacial period started about 116 ka, in response to orbital forcing (Box 6.1), with the growth of ice sheets and fall of sea level culminating in the Last Glacial Maximum (LGM), around 21 ka. The LGM and the subsequent deglaciation have been widely studied because the radiative forcings, boundary conditions and climate response are relatively well known.
The response of the climate system at the LGM included feedbacks in the atmosphere and on land amplifying the orbital forcing. Concentrations of well-mixed greenhouse gases at the LGM were reduced relative to pre-industrial values (Figures 6.3 and 6.4), amounting to a global radiative perturbation of –2.8 W m–2 – approximately equal to, but opposite from, the radiative forcing of these gases for the year 2000 (see Section 2.3). Land ice covered large parts of North America and Europe at the LGM, lowering sea level and exposing new land. The radiative perturbation of the ice sheets and lowered sea level, specified as a boundary condition for some LGM simulations, has been estimated to be about –3.2 W m–2, but with uncertainties associated with the coverage and height of LGM continental ice (Mangerud et al., 2002; Peltier, 2004; Toracinta et al., 2004; Masson-Delmotte et al., 2006) and the parametrization of ice albedo in climate models (Taylor et al., 2000). The distribution of vegetation was altered, with tundra expanded over the northern continents and tropical rain forest reduced (Prentice et al., 2000), and atmospheric aerosols (primarily dust) were increased (Kohfeld and Harrison, 2001), partly as a consequence of reduced vegetation cover (Mahowald et al., 1999). Vegetation and atmospheric aerosols are treated as specified conditions in some LGM simulations, each contributing about –1 W m–2 of radiative perturbation, but with very low scientific understanding of their radiative influence at the LGM (Claquin et al., 2003; Crucifix and Hewitt, 2005). Changes in biogeochemical cycles thus played an important role and contributed, through changes in greenhouse gas concentration, dust loading and vegetation cover, more than half of the known radiative perturbation during the LGM. Overall, the radiative perturbation for the changed greenhouse gas and aerosol concentrations and land surface was approximately –8 W m–2 for the LGM, although with significant uncertainty in the estimates for the contributions of aerosol and land surface changes (Figure 6.5).
Understanding of the magnitude of tropical cooling over land at the LGM has improved since the TAR with more records, as well as better dating and interpretation of the climate signal associated with snow line elevation and vegetation change. Reconstructions of terrestrial climate show strong spatial differentiation, regionally and with elevation. Pollen records with their extensive spatial coverage indicate that tropical lowlands were on average 2°C to 3°C cooler than present, with strong cooling (5°C–6°C) in Central America and northern South America and weak cooling (<2°C) in the western Pacific Rim (Farrera et al., 1999). Tropical highland cooling estimates derived from snow-line and pollen-based inferences show similar spatial variations in cooling although involving substantial uncertainties from dating and mapping, multiple climatic causes of treeline and snow line changes during glacial periods (Porter, 2001; Kageyama et al., 2004), and temporal asynchroneity between different regions of the tropics (Smith et al., 2005). These new studies give a much richer regional picture of tropical land cooling, and stress the need to use more than a few widely scattered proxy records as a measure of low-latitude climate sensitivity (Harrison, 2005).
The Climate: Long-range Investigation, Mapping, and Prediction (CLIMAP) reconstruction of ocean surface temperatures produced in the early 1980s indicated about 3°C cooling in the tropical Atlantic, and little or no cooling in the tropical Pacific. More pronounced tropical cooling for the LGM tropical oceans has since been proposed, including 4°C to 5°C based on coral skeleton records from off Barbados (Guilderson et al., 1994) and up to 6°C in the cold tongue off western South America based on foraminiferal assemblages (Mix et al., 1999). New data syntheses from multiple proxy types using carefully defined chronostratigraphies and new calibration data sets are now available from the Glacial Ocean Mapping (GLAMAP) and Multiproxy Approach for the Reconstruction of the Glacial Ocean surface (MARGO) projects, although with caveats including selective species dissolution, dating precision, non-analogue situations, and environmental preferences of the organisms (Sarnthein et al., 2003b; Kucera et al., 2005; and references therein). These recent reconstructions confirm moderate cooling, generally 0°C to 3.5°C, of tropical SST at the LGM, although with significant regional variation, as well as greater cooling in eastern boundary currents and equatorial upwelling regions. Estimates of cooling show notable differences among the different proxies. Faunal-based proxies argue for an intensification of the eastern equatorial Pacific cold tongue in contrast to Mg/Ca-based SST estimates that suggest a relaxation of SST gradients within the cold tongue (Mix et al., 1999; Koutavas et al., 2002; Rosenthal and Broccoli, 2004). Using a Bayesian approach to combine different proxies, Ballantyne et al. (2005) estimated a LGM cooling of tropical SSTs of 2.7°C ± 0.5°C (1 standard deviation).
These ocean proxy synthesis projects also indicate a colder glacial winter North Atlantic with more extensive sea ice than present, whereas summer sea ice only covered the glacial Arctic Ocean and Fram Strait with the northern North Atlantic and Nordic Seas largely ice free and more meridional ocean surface circulation in the eastern parts of the Nordic Seas (Sarnthein et al., 2003a; Meland et al., 2005; de Vernal et al., 2006). Sea ice around Antarctica at the LGM also responded with a large expansion of winter sea ice and substantial seasonal variation (Gersonde et al., 2005). Over mid- and high-latitude northern continents, strong reductions in temperatures produced southward displacement and major reductions in forest area (Bigelow et al., 2003), expansion of permafrost limits over northwest Europe (Renssen and Vandenberghe, 2003), fragmentation of temperate forests (Prentice et al., 2000; Williams et al., 2000) and predominance of steppe-tundra in Western Europe (Peyron et al., 2005). Temperature reconstructions from polar ice cores indicate strong cooling at high latitudes of about 9°C in Antarctica (Stenni et al., 2001) and about 21°C in Greenland (Dahl-Jensen et al., 1998).
The strength and depth extent of the LGM Atlantic overturning circulation have been examined through the application of a variety of new marine proxy indicators (Rutberg et al., 2000; Duplessy et al., 2002; Marchitto et al., 2002; McManus et al., 2004). These tracers indicate that the boundary between North Atlantic Deep Water (NADW) and Antarctic Bottom Water was much shallower during the LGM, with a reinforced pycnocline between intermediate and particularly cold and salty deep water (Adkins et al., 2002). Most of the deglaciation occurred over the period about 17 to 10 ka, the same period of maximum deglacial atmospheric CO2 increase (Figure 6.4). It is thus very likely that the global warming of 4°C to 7°C since the LGM occurred at an average rate about 10 times slower than the warming of the 20th century.
In summary, significant progress has been made in the understanding of regional changes at the LGM with the development of new proxies, many new records, improved understanding of the relationship of the various proxies to climate variables and syntheses of proxy records into reconstructions with stricter dating and common calibrations.