IPCC (1992, 1994) and the SAR have dealt with the climatic effects of episodic, explosive volcanic events which result in significant enhancements of aerosol concentrations in the stratosphere. The most dramatic of these in recent times has been the 1991 eruption of Mt. Pinatubo. The radiative, chemical, dynamical and climatic consequences accompanying the transient duration of sulphuric acid aerosols in the stratosphere have been discussed in previous IPCC assessments. The eruption of Mt. Pinatubo reached a peak forcing of about -3 Wm-2 (uncertainty of 20%) in 1991 (Hansen et al., 1998; Stenchikov et al., 1998), and was plausibly the largest volcanic aerosol forcing of this century, perturbing the stratospheric and surface climate significantly (SAR). However, aerosol levels in the stratosphere have now fallen to well below the peak values seen in 1991 to 1993, and are comparable to the very low value seen in about 1979, which was a quiescent time as far as volcanic activity was concerned (WMO, 1999). It is likely that even the heterogeneous chemical effects initiated by aerosols upon the O3 chemistry and its destruction (Solomon et al., 1996) have diminished. One consequence of the aerosol-induced ozone depletion may have been a prolonged cooling of the lower stratosphere to abnormally low values through the mid-1990s, as estimated from satellite and radiosonde observations (WMO, 1999, Chapter 5). Although episodic in nature and transient in duration, volcanic events can exert a significant influence on the time history of the radiative forcing evolution and thereby on the long-term (interannual to decadal scale) temperature record (see Chapter 9).
As noted in previous IPCC Assessments, there are difficulties in compiling a good quantitative record of the episodic volcanic events (see also Rowntree, 1998), in particular the intensity of their forcings prior to the 1960s. Efforts have been made to compile the optical depths of the past volcanoes (SAR; Robock and Free, 1995, 1996; Andronova et al., 1999); however, the estimated global forcings probably have an uncertainty of a factor of two or more. Several major volcanic eruptions occurred between 1880 and 1920, and between 1960 and 1991.
Changes in land surface albedo can result from land-use changes (Henderson-Sellers, 1995) and thus be tied to an anthropogenic cause. Hansen et al. (1997b) estimate that a forcing of -0.4 Wm-2 has resulted, about half of which is estimated to have occurred in the Industrial Era. The largest effect is estimated to be at the high latitudes where snow-covered forests that have a lower albedo have been replaced by snow-covered deforested areas. Hansen et al. (1998) point out that the albedo of a cultivated field is affected more by a given snowfall than the albedo of an evergreen forest. They performed a simulation with pre-industrial vegetation replaced by current land-use patterns and found the global mean forcing to be -0.21 Wm-2, with the largest contributions coming from deforested areas in Eurasia and North America. In a similar study, Betts (2001) estimates an instantaneous radiative forcing of -0.20 Wm-2 by surface albedo change due to present day land use relative to natural vegetation. In agreement with Hansen et al. (1997b), the greatest effect is seen in the high latitude agricultural regions. If, as above, half of the land clearance is assumed to have taken place since the industrial revolution, this suggests a forcing of -0.10 Wm-2 by land use over this period. In a parallel simulation with the climate free to respond to the change in albedo and other vegetation characteristics, lower temperatures are simulated in the northern mid-latitudes. These are mainly attributed to the increased surface albedo, although increases in cloud cover cause further localised reductions in the net surface short-wave radiation in some regions. However, some areas exhibit higher temperatures in their dry season, consistent with a decrease in evapotranspiration due to reduced access of soil moisture by the shallower roots of the crops compared with forest.
Following Hansen et al. (1997b), Shine and Forster (1999) recommended in their review a value of -0.2 Wm-2 with at least a 0.2 Wm-2 uncertainty. We adopt those values here for the best estimate and range, respectively; however, in view of the small number of investigations and uncertainty in historical land cover changes, there is very low confidence in these values at present.
Changes in land use can also exert other kinds of climatic impacts, e.g., changes in roughness length, turbulent fluxes, and soil moisture and heat budgets (see also Chapters 7 and 8). Further, there are a host of factors that are potentially affected by land-use change and that could have an impact on the atmospheric concentrations of radiatively active trace gases and aerosols. For instance, the dry deposition rates of species could be affected owing to the surface roughness change. Precipitation changes induced by deforestation etc. could affect the wet deposition of species and thereby bring about biogeochemical changes, leading to changes in lifetimes. The impacts due to such changes have not been comprehensively investigated.
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