9.2.2 Spatial and Temporal Patterns of the Response to Different Forcings and their Uncertainties
9.2.2.1 Spatial and Temporal Patterns of Response
The ability to distinguish between climate responses to different external forcing factors in observations depends on the extent to which those responses are distinct (see, e.g., Section 9.4.1.4 and Appendix 9.A). Figure 9.1 illustrates the zonal average temperature response in the PCM model (see Table 8.1 for model details) to several different forcing agents over the last 100 years, while Figure 9.2 illustrates the zonal average temperature response in the Commonwealth Scientific and Industrial Research Organisation (CSIRO) atmospheric model (when coupled to a simple mixed layer ocean model) to fossil fuel black carbon and organic matter, and to the combined effect of these forcings together with biomass burning aerosols (Penner et al., 2007). These figures indicate that the modelled vertical and zonal average signature of the temperature response should depend on the forcings. The major features shown in Figure 9.1 are robust to using different climate models. On the other hand, the response to black carbon forcing has not been widely examined and therefore the features in Figure 9.2 may be model dependent. Nevertheless, the response to black carbon forcings appears to be small.
Greenhouse gas forcing is expected to produce warming in the troposphere, cooling in the stratosphere, and, for transient simulations, somewhat more warming near the surface in the NH due to its larger land fraction, which has a shorter surface response time to the warming than do ocean regions (Figure 9.1c). The spatial pattern of the transient surface temperature response to greenhouse gas forcing also typically exhibits a land-sea pattern of stronger warming over land, for the same reason (e.g., Cubasch et al., 2001). Sulphate aerosol forcing results in cooling throughout most of the globe, with greater cooling in the NH due to its higher aerosol loading (Figure 9.1e; see Chapter 2), thereby partially offsetting the greater NH greenhouse-gas induced warming. The combined effect of tropospheric and stratospheric ozone forcing (Figure 9.1d) is expected to warm the troposphere, due to increases in tropospheric ozone, and cool the stratosphere, particularly at high latitudes where stratospheric ozone loss has been greatest. Greenhouse gas forcing is also expected to change the hydrological cycle worldwide, leading to disproportionately greater increases in heavy precipitation (Chapter 10 and Section 9.5.4), while aerosol forcing can influence rainfall regionally (Section 9.5.4).
The simulated responses to natural forcing are distinct from those due to the anthropogenic forcings described above. Solar forcing results in a general warming of the atmosphere (Figure 9.1a) with a pattern of surface warming that is similar to that expected from greenhouse gas warming, but in contrast to the response to greenhouse warming, the simulated solar-forced warming extends throughout the atmosphere (see, e.g., Cubasch et al., 1997). A number of independent analyses have identified tropospheric changes that appear to be associated with the solar cycle (van Loon and Shea, 2000; Gleisner and Thejll, 2003; Haigh, 2003; White et al., 2003; Coughlin and Tung, 2004; Labitzke, 2004; Crooks and Gray, 2005), suggesting an overall warmer and moister troposphere during solar maximum. The peak-to-trough amplitude of the response to the solar cycle globally is estimated to be approximately 0.1°C near the surface. Such variations over the 11-year solar cycle make it is necessary to use several decades of data in detection and attribution studies. The solar cycle also affects atmospheric ozone concentrations with possible impacts on temperatures and winds in the stratosphere, and has been hypothesised to influence clouds through cosmic rays (Section 2.7.1.3). Note that there is substantial uncertainty in the identification of climate response to solar cycle variations because the satellite period is short relative to the solar cycle length, and because the response is difficult to separate from internal climate variations and the response to volcanic eruptions (Gray et al., 2005).
Volcanic sulphur dioxide (SO2) emissions ejected into the stratosphere form sulphate aerosols and lead to a forcing that causes a surface and tropospheric cooling and a stratospheric warming that peak several months after a volcanic eruption and last for several years. Volcanic forcing also likely leads to a response in the atmospheric circulation in boreal winter (discussed below) and a reduction in land precipitation (Robock and Liu, 1994; Broccoli et al., 2003; Gillett et al., 2004b). The response to volcanic forcing causes a net cooling over the 20th century because of variations in the frequency and intensity of volcanic eruptions. This results in stronger volcanic forcing towards the end of the 20th century than early in the 20th century. In the PCM, this increase results in a small warming in the lower stratosphere and near the surface at high latitudes, with cooling elsewhere (Figure 9.1b).
The net effect of all forcings combined is a pattern of NH temperature change near the surface that is dominated by the positive forcings (primarily greenhouse gases), and cooling in the stratosphere that results predominantly from greenhouse gas and stratospheric ozone forcing (Figure 9.1f). Results obtained with the CSIRO model (Figure 9.2) suggest that black carbon, organic matter and biomass aerosols would slightly enhance the NH warming shown in Figure 9.1f. On the other hand, indirect aerosol forcing from fossil fuel aerosols may be larger than the direct effects that are represented in the CSIRO and PCM models, in which case the NH warming could be somewhat diminished. Also, while land use change may cause substantial forcing regionally and seasonally, its forcing and response are expected to have only a small impact at large spatial scales (Sections 9.3.3.3 and 7.2.2; Figures 2.20 and 2.23).
The spatial signature of a climate model’s response is seldom very similar to that of the forcing, due in part to the strength of the feedbacks relative to the initial forcing. This comes about because climate system feedbacks vary spatially and because the atmospheric and ocean circulation cause a redistribution of energy over the globe. For example, sea ice albedo feedbacks tend to enhance the high-latitude response of both a positive forcing, such as that of CO2, and a negative forcing such as that of sulphate aerosol (e.g., Mitchell et al., 2001; Rotstayn and Penner, 2001). Cloud feedbacks can affect both the spatial signature of the response to a given forcing and the sign of the change in temperature relative to the sign of the radiative forcing (Section 8.6). Heating by black carbon, for example, can decrease cloudiness (Ackerman et al., 2000). If the black carbon is near the surface, it may increase surface temperatures, while at higher altitudes it may reduce surface temperatures (Hansen et al., 1997; Penner et al., 2003). Feedbacks can also lead to differences in the response of different models to a given forcing agent, since the spatial response of a climate model to forcing depends on its representation of these feedbacks and processes. Additional factors that affect the spatial pattern of response include differences in thermal inertia between land and sea areas, and the lifetimes of the various forcing agents. Shorter-lived agents, such as aerosols, tend to have a more distinct spatial pattern of forcing, and can therefore be expected to have some locally distinct response features.
The pattern of response to a radiative forcing can also be altered quite substantially if the atmospheric circulation is affected by the forcing. Modelling studies and data comparisons suggest that volcanic aerosols (e.g., Kirchner et al., 1999; Shindell et al., 1999; Yang and Schlesinger, 2001; Stenchikov et al., 2006) and greenhouse gas changes (e.g., Fyfe et al., 1999; Shindell et al., 1999; Rauthe et al., 2004) can alter the North Atlantic Oscillation (NAO) or the Northern Annular Mode (NAM). For example, volcanic eruptions, with the exception of high-latitude eruptions, are often followed by a positive phase of the NAM or NAO (e.g., Stenchikov et al., 2006) leading to Eurasian winter warming that may reduce the overall cooling effect of volcanic eruptions on annual averages, particularly over Eurasia (Perlwitz and Graf, 2001; Stenchikov et al., 2002; Shindell et al., 2003; Stenchikov et al., 2004; Oman et al., 2005; Rind et al., 2005a; Miller et al., 2006; Stenchikov et al., 2006). In contrast, NAM or NAO responses to solar forcing vary between studies, some indicating a response, perhaps with dependence of the response on season or other conditions, and some finding no changes (Shindell et al., 2001a,b; Ruzmaikin and Feynman, 2002; Tourpali et al., 2003; Egorova et al., 2004; Palmer et al., 2004; Stendel et al., 2006; see also review in Gray et al., 2005).
In addition to the spatial pattern, the temporal evolution of the different forcings (Figure 2.23) generally helps to distinguish between the responses to different forcings. For example, Santer et al. (1996b,c) point out that a temporal pattern in the hemispheric temperature contrast would be expected in the second half of the 20th century with the SH warming more than the NH for the first two decades of this period and the NH subsequently warming more than the SH, as a result of changes in the relative strengths of the greenhouse gas and aerosol forcings. However, it should be noted that the integrating effect of the oceans (Hasselmann, 1976) results in climate responses that are more similar in time between different forcings than the forcings are to each other, and that there are substantial uncertainties in the evolution of the hemispheric temperature contrasts associated with sulphate aerosol forcing.