IPCC Fourth Assessment Report: Climate Change 2007
Climate Change 2007: Working Group I: The Physical Science Basis

6.4.2 Abrupt Climatic Changes in the Glacial-Interglacial Record What Is the Evidence for Past Abrupt Climate Changes?

Abrupt climate changes have been variously defined either simply as large changes within less than 30 years (Clark et al., 2002), or in a physical sense, as a threshold transition or a response that is rapid compared to forcing (Rahmstorf, 2001; Alley et al., 2003). Overpeck and Trenberth (2004) noted that not all abrupt changes need to be externally forced. Numerous terrestrial, ice and oceanic climatic records show that large, widespread, abrupt climate changes have occurred repeatedly throughout the past glacial interval (see review by Rahmstorf, 2002). High-latitude records show that ice age abrupt temperature events were larger and more widespread than were those of the Holocene. The most dramatic of these abrupt climate changes were the Dansgaard-Oeschger (D-O) events, characterised by a warming in Greenland of 8°C to 16°C within a few decades (see Severinghaus and Brook, 1999; Masson-Delmotte et al., 2005a for a review) followed by much slower cooling over centuries. Another type of abrupt change were the Heinrich events; characterised by large discharges of icebergs into the northern Atlantic leaving diagnostic drop-stones in the ocean sediments (Hemming, 2004). In the North Atlantic, Heinrich events were accompanied by a strong reduction in sea surface salinity (Bond et al., 1993), as well as a sea surface cooling on a centennial time scale. Such ice age cold periods lasted hundreds to thousands of years, and the warming that ended them took place within decades (Figure 6.7; Cortijo et al., 1997; Voelker, 2002). At the end of the last glacial, as the climate warmed and ice sheets melted, climate went through a number of abrupt cold phases, notably the Younger Dryas and the 8.2 ka event.

Figure 6.7

Figure 6.7. The evolution of climate indicators from the NH (panels a to d), and from Antarctica (panels e to g), over the period 64 to 30 ka. (a) Anhysteretic remanent magnetisation (ARM), here a proxy of the northward extent of Atlantic MOC, from an ocean sediment core from the Nordic Seas (Dokken and Jansen, 1999); (b) CH4 as recorded in Greenland ice cores at the Greenland Ice Core Project (GRIP), Greenland Ice Sheet Project (GISP) and North GRIP (NGRIP) sites (Blunier and Brook, 2001; Flückiger et al., 2004; Huber et al., 2006); CH4 data for the period 40 to 30 ka were selected for the GRIP site and for 64 to 40 ka for the GISP site when sample resolution is highest in the cores; (c) surface temperature estimated from nitrogen isotope ratios that are influenced by thermal diffusion (Huber et al., 2006); (d) d18O, a proxy for surface temperature, from NGRIP (2004) with the D-O NH warm events 8, 12, 14 and 17 indicated; (e) d18O from Byrd, Antarctica (Blunier and Brook, 2001) with A1 to A4 denoting antarctic warm events; (f) nss-Ca2+, a proxy of dust and iron deposition, from Dome C, Antarctica (Röthlisberger et al., 2004); and (g) CO2 as recorded in ice from Taylor Dome, Antarctica (Indermühle et al., 2000). The Heinrich events (periods of massive ice-rafted debris recorded in marine sediments) H3, H4, H5, H5.2, and H6, are shown. All data are plotted on the Greenland SS09sea time scale (Johnsen et al., 2001). CO2 and CH4 are well mixed in the atmosphere. CH4 variations are synchronous within the resolution of ±50 years with variations in Greenland temperature, but a detailed analysis suggests that CH4 rises lag temperature increases at the onset of the D-O events by 25 to 70 years (Huber et al., 2006). CO2 co-varied with the antarctic temperature, but the exact synchronisation between Taylor Dome and Byrd is uncertain, thus making the determination of leads or lags between temperature and CO2 elusive. The evolution of Greenland and antarctic temperature is consistent with a reorganisation of the heat transport and the MOC in the Atlantic (Knutti et al., 2004).

The effects of these abrupt climate changes were global, although out-of-phase responses in the two hemispheres (Blunier et al., 1998; Landais et al., 2006) suggest that they were not primarily changes in global mean temperature. The highest amplitude of the changes, in terms of temperature, appears centred around the North Atlantic. Strong and fast changes are found in the global CH4 concentration (of the order of 100 to 150 ppb within decades), which may point to changes in the extent or productivity of tropical wetlands (see Chappellaz et al., 1993; Brook et al., 2000 for a review; Masson-Delmotte et al., 2005a), and in the Asian monsoon (Wang et al., 2001). The NH cold phases were linked with a reduced northward flow of warm waters in the Nordic Seas (Figure 6.7), southward shift of the Inter-Tropical Convergence Zone (ITCZ) and thus the location of the tropical rainfall belts (Peterson et al., 2000; Lea et al., 2003). Cold, dry and windy conditions with low CH4 and high dust aerosol concentrations generally occurred together in the NH cold events. The accompanying changes in atmospheric CO2 content were relatively small (less than 25 ppm; Figure 6.7) and parallel to the antarctic counterparts of Greenland D-O events. The record in N2O is less complete and shows an increase of about 50 ppb and a decrease of about 30 ppb during warm and cold periods, respectively (Flückiger et al., 2004).

A southward shift of the boreal treeline and other rapid vegetation responses were associated with past cold events (Peteet, 1995; Shuman et al., 2002; Williams et al., 2002). Decadal-scale changes in vegetation have been recorded in annually laminated sequences at the beginning and the end of the Younger Dryas and the 8.2 ka event (Birks and Ammann, 2000; Tinner and Lotter, 2001; Veski et al., 2004). Marine pollen records with a typical sampling resolution of 200 years provide unequivocal evidence of the immediate response of vegetation in Southern Europe to the climate fluctuations during glacial times (Sánchez Goñi et al., 2002; Tzedakis, 2005). The same holds true for the vegetation response in northern South America during the last deglaciation (Hughen et al., 2004).