8.7.2 Forced Abrupt Climate Change
18.104.22.168 Meridional Overturning Circulation Changes
As the radiative forcing of the planet changes, the climate system responds on many different time scales. For the physical climate system typically simulated in coupled models (atmosphere, ocean, land, sea ice), the longest response time scales are found in the ocean (Stouffer, 2004). In terms of thresholds and abrupt climate changes on decadal and longer time scales, the ocean has also been a focus of attention. In particular, the ocean’s Atlantic MOC (see Box 5.1 for definition and description) is a main area of study.
The MOC transports large amounts of heat (order of 1015 Watts) and salt into high latitudes of the North Atlantic. There, the heat is released to the atmosphere, cooling the surface waters. The cold, relatively salty waters sink to depth and flow southward out of the Atlantic Basin. The complete set of climatic drivers of this circulation remains unclear but it is likely that both density (e.g., Stommel 1961; Rooth 1982) and wind stress forcings (e.g., Wunsch, 2002; Timmermann and Goosse, 2004) are important. Both palaeoclimate studies (e.g., Broecker, 1997; Clark et al., 2002) and modelling studies (e.g., Manabe and Stouffer, 1988, 1997; Vellinga and Wood, 2002) suggest that disruptions in the MOC can produce abrupt climate changes. A systematic model intercomparison study (Rahmstorf et al., 2005) found that all 11 participating EMICs had a threshold where the MOC shuts down (see Section 8.8.3). Due to the high computational cost, such a search for thresholds has not yet been performed with AOGCMs.
It is important to note the distinction between the equilibrium and transient or time-dependent responses of the MOC to changes in forcing. Due to the long response time scales found in the ocean (some longer than 1 kyr), it is possible that the short-term response to a given forcing change may be very different from the equilibrium response. Such behaviour of the coupled system has been documented in at least one AOGCM (Stouffer and Manabe, 2003) and suggested in the results of a few other AOGCM studies (e.g., Hirst, 1999; Senior and Mitchell, 2000; Bryan et al., 2006). In these AOGCM experiments, the MOC weakens as the greenhouse gases increase in the atmosphere. When the CO2 concentration is stabilised, the MOC slowly returns to its unperturbed value.
As discussed in section 10.3.4, the MOC typically weakens as greenhouse gases increase due to the changes in surface heat and freshwater fluxes at high latitudes (Manabe et al., 1991). The surface flux changes reduce the surface density, hindering the vertical movement of water and slowing the MOC. As the MOC slows, it could approach a threshold where the circulation can no longer sustain itself. Once the MOC crosses this threshold, it could rapidly change states, causing abrupt climate change where the North Atlantic and surrounding land areas would cool relative to the case where the MOC is active. This cooling is the result of the loss of heat transport from low latitudes in the Atlantic and the feedbacks associated with the reduction in the vertical mixing of high-latitude waters.
A common misunderstanding is that the MOC weakening could cause the onset of an ice age. However, no model has supported this speculation when forced with realistic estimates of future climate forcings (see Section 10.3.4). In addition, in idealised modelling studies where the MOC was forced to shut down through very large sources of freshwater (not changes in greenhouse gases), the surface temperature changes do not support the idea that an ice age could result from a MOC shut down, although the impacts on climate would be large (Manabe and Stouffer, 1988, 1997; Schiller et al., 1997; Vellinga and Wood, 2002; Stouffer et al., 2006). In a recent intercomparison involving 11 coupled atmosphere-ocean models (Gregory et al., 2005), the MOC decreases by only 10 to 50% during a 140-year period (as atmospheric CO2 quadruples), and in no model is there a land cooling anywhere (as the global-scale heating due to increasing CO2 overwhelms the local cooling effect due to reduced MOC).
Because of the large amount of heat and salt transported northward and its sensitivity to surface fluxes, the changes in the MOC are able to produce abrupt climate change on decadal to centennial time scales (e.g., Manabe and Stouffer, 1995; Stouffer et al., 2006). Idealised studies using present-day simulations have shown that models can simulate many of the variations seen in the palaeoclimate record on decadal to centennial time scales when forced by fluxes of freshwater water at the ocean surface. However, the quantitative response to freshwater inputs varies widely among models (Stouffer et al., 2006), which led the CMIP and Paleoclimate Modelling Intercomparison Project (PMIP) panels to design and support a set of coordinated experiments to study this issue (http://www.gfdl.noaa.gov/~kd/CMIP.html and http://www.pmip2.cnrs-gif.fr/pmip2/design/experiments/waterhosing.shtml).
In addition to the amount of the freshwater input, the exact location of that input may also be important (Rahmstorf 1996, Manabe and Stouffer, 1997; Rind et al., 2001). Designing experiments and determining the realistic past forcings needed to test the models’ response on decadal to centennial time scales remains to be accomplished.
The processes determining MOC response to increasing greenhouse gases have been studied in a number of models. In many models, initial MOC response to increasing greenhouse gases is dominated by thermal effects. In most models, this is enhanced by changes in salinity driven by, among other things, the expected strengthening of the hydrological cycle (Gregory et al., 2005; Chapter 10). Melt water runoff from a melting of the Greenland Ice Sheet is a potentially major source of freshening not yet included in the models found in the MMD (see Section 22.214.171.124). More complex feedbacks, associated with wind and hydrological changes, are also important in many models. These include local surface flux anomalies in deep-water formation regions (Gent, 2001) and oceanic teleconnections driven by changes to the freshwater budget of the tropical and South Atlantic (e.g., Latif et al., 2000; Thorpe et al., 2001; Vellinga et al., 2002; Hu et al., 2004). The magnitudes of the climate factors causing the MOC to weaken, along with the feedbacks and the associated restoring factors, are all uncertain at this time. Evaluation of these processes in AOGCMs is mainly restricted by lack of observations, but some early progress has been made in individual studies (e.g., Schmittner et al., 2000; Pardaens et al., 2003; Wu et al., 2005; Chapter 9). Model intercomparison studies (e.g., Gregory et al., 2005; Rahmstorf et al., 2005; Stouffer et al., 2006) were developed to identify and understand the causes for the wide range of MOC responses in the coupled models used here (see Chapters 4, 6 and 10).