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~30 km. (The gap around 10 km is an artifact of the calculations.) However, increased 03 above ~30 km causes cooling due to increased thermal emission to space and to increased absorption of solar radiation before it can reach the earth's surface (e.g., see Ra-manathan et al., 1985).

Addition of a constant, absolute amount of 03 may be unrealistic, however, in that 10 DU is a large percentage change in the existing 03 concentration at lower altitudes but a smaller percentage at the altitudes at which the ozone concentration peaks. Figure 14.21b, for example, shows one global mean vertical profile for 03 (in Dobson units). Adding fO DU at the tropopause corresponds to an increase in 03 of ~400%, whereas at 24 km, adding fO DU only increases 03 by ~60% (Forster and Shine, 1997). Figure 14.21c shows a calculated change in surface temperature due to a systematic change in 03 of f0% in each layer. While 03 at the tropopause is still important, that at lower and higher altitudes is relatively more important compared to the case in Fig. 14.21a where an absolute increase in 03 in each layer is assumed.

A number of model studies have explored the climate implications of changes in the vertical distribution of ozone (e.g., see Schwarzkopf and Ramaswamy, f 993; Wang et al., 1993; Molnar et al., 1994; and Chalita et al., 1996). For example, Fig. 14.22a shows a model-calculated percentage change in tropospheric 03 in July as a function of altitude and latitude from prein-dustrial times to the present (Chalita et al., 1996). The largest increase in concentration is predicted in the boundary layer, with smaller increases at higher altitudes. However, as seen in Fig. 14.22b, the major contribution to radiative forcing (see Section B.3) comes from the relatively small ozone increase predicted for the 6- to f2-km region.

Changes in stratospheric ozone also impact atmospheric temperatures through three effects. As discussed in Chapter 3, UV absorption by stratospheric

03 and the energy released in the O + 02 reaction warm the stratosphere. As a result, destruction of stratospheric ozone due to chlorofluorocarbons results in cooling of the stratosphere. From the Boltzmann relationship (Eq. (A)), there is then less downward radiation across the tropopause from ozone at these lower temperatures, in addition, there is the direct effect of a smaller ozone concentration to emit infrared radiation, part of which is in a downward direction and which normally contributes to heating of the troposphere. These two effects of changes in stratospheric ozone lead to cooling of the troposphere. Counterbalancing these effects is the increased solar UV reaching the troposphere due to less absorption by stratospheric 03, which is expected to cause surface heating (e.g., see Ramaswamy et al., 1992, 1996; Zhong et al., 1996; and Shine et al., 1998).

It should be noted that while changes in stratospheric ozone can impact tropospheric heating and cooling, greenhouse gases may also impact stratospheric ozone destruction. As described earlier in this chapter, while C02 causes warming in the troposphere, it causes cooling in the stratosphere through the efficient emission of infrared to space. Some model calculations suggest that additional stratospheric cooling in the polar regions due to increased greenhouse gases may increase the formation of polar stratospheric clouds. Since these play such a key role in ozone destruction in those regions (see Chapter 12.C.5), increased ozone destruction is predicted, particularly in the Arctic, where temperatures are not as routinely cold as in the Antarctic (Austin et al., 1992, 1994; Shindell et al., 1998). (ft should be noted, however, that warming of the troposphere by trapping of outgoing terrestrial radiation by PSCs during winter has also been proposed as being important historically at times of high methane concentrations that oxidized to form water; Sloan and Pollard, 1998.) In addition, it has been suggested that the recovery of these regions as the

Latitude Latitude

FIGURE 14.22 Model-calculated percentage (a) increase in 03 (zonal average) in July and (b) the corresponding contributions to instantaneous radiative forcing calculated as a function of latitude and altitude from preindustrial times to the present (adapted from Chalita et al., 1996).

Latitude Latitude

FIGURE 14.22 Model-calculated percentage (a) increase in 03 (zonal average) in July and (b) the corresponding contributions to instantaneous radiative forcing calculated as a function of latitude and altitude from preindustrial times to the present (adapted from Chalita et al., 1996).

emissions of ozone-destroying chlorine and bromine compounds decline may be delayed by about a decade due to this effect of greenhouse gases (Shindell et al., 1998).

There are a number of factors that affect the ultimate climate response to changes in tropospheric and stratospheric ozone. These include the altitude dependence of the forcing previously discussed, its role in absorbing solar UV in both the stratosphere and troposphere, its depletion through chain reactions of CFCs in the stratosphere, and finally, the large variability in its concentrations geographically, vertically, and temporally. Because of these complexities, the net effect is expected to also vary from one location to another, as well as temporally. The spatial and temporal effects due to ozone formed from emissions from biomass burning over large areas of the tropics are one example (Portmann et al., 1997). A number of studies have addressed the net effect of changes in 03 on climate, and the reader is referred to them for more detailed information (e.g., see Hauglustaine et al., 1994; Marenco et al., 1994; Lelieveld and van Dorland, 1995; Forster et al., 1996; Chalita et al., 1996; Portmann et al., 1997; Berntsen et al., 1997; van Dorland et al., 1997; Graf et al., 1998; Haywood et al., 1998c; Brasseur et al., 1998; Stevenson et al., 1998; and Wang and Jacob, 1998).

e. CFCs, HCFCs, and HFCs

The atmospheric concentrations of chlorofluoro-carbons (CFCs), hydrochlorofluorocarbons (HCFCs), and hydrofluorocarbons (HFCs) and the trends in these concentrations are discussed in detail in Chapter 13. In brief, the atmospheric concentrations of CFCs increased as they came into increasing use in the f950s. As the phase-outs specified in the Montreal Protocol and its amendments have come into play (Figs. 12.13 and 13.4), the growth rates have fallen dramatically. For example, that for CFC-12 was ~18 ppt per year prior to the Montreal Protocol but in mid-1993 had fallen to 6-7 ppt per year (e.g., see Cunnold et al., 1997) and has since become slightly negative (e.g., Derwent et al., 1998a). However, the concentrations of their alternatives, the HCFCs and HFCs, are increasing as expected (see Fig. 13.6). Although they do not contribute significantly to radiative forcing at present, they could do so if their emissions approach those of the compounds they are replacing (Derwent et al., 1998b).

/. Other Gases

Other anthropogenically emitted gases such as CO have also been suggested to contribute to the greenhouse effect (e.g., see Evans and Puckrin, 1995). CO concentrations also increased during the 1980s but then decreased from 1988 to 1992 (e.g., see Khalil and Rasmussen, 1984, 1994a; Novelli et al., 1994; and Yurganov et al., 1997). CO is not believed to directly contribute significantly to the greenhouse effect (IPCC, 1996). However, increasing CO emissions may decrease the OH concentration, which would then increase the concentrations of other greenhouse gases that react with OH, such as CH4. For example, Daniel and Solomon (1998) estimate that this indirect effect associated with anthropogenic emissions may be as or more significant over the next f5 years than that due to anthropogenic emissions of N20.

3. Radiative Forcing by Greenhouse Gases and Global Warming Potentials

In the simplest of worlds, the greenhouse gases would exert their influences independent of each other and of other factors such as aerosols and clouds, feedback mechanisms, and ozone depletion. This, of course, is not the case. However, it is useful before examining these "real-world" considerations to consider the direct effects on the radiation balance of the atmosphere of the greenhouse gases. These are commonly expressed in terms of the radiative forcing. Another tool used for examining the relative effects of various gases on the radiation balance of the atmosphere is the global warming potential.

a. Instantaneous and Adjusted Radiative Forcing

As discussed at the beginning of this chapter, changes in the radiation balance of the atmosphere can occur due to changes either in incoming solar radiation or in the outgoing infrared radiation. Radiative forcing is defined as a change in the average net radiation at the tropopause due to a particular perturbation of interest. This change (usually expressed in W m~2) could be in either the incoming or outgoing radiation.

Using the flux at the tropopause to define radiative forcing is believed to be appropriate because of the rapid vertical mixing by convection and large-scale processes within the troposphere which closely couples the troposphere and the earth's surface (Wang et al., 1995). As a result, energy absorbed in the troposphere is assumed to be effective in warming the earth's surface (Lacis et al., 1990) and the change in the flux at the tropopause can be used to calculate the change in the surface temperature (Ramanathan, 1976).

Two approaches to calculating radiative forcing due to greenhouse gases have been taken. In the first, the immediate forcing due to increases in the greenhouse gas is calculated without allowing for a change in the stratospheric temperature. This is what is known as the instantaneous radiative forcing.

The second approach is to calculate the radiative forcing after allowing stratospheric temperatures to readjust to radiative equilibrium, but with the temperatures of the earth's surface and troposphere, as well as atmospheric moisture, fixed. This is known as the adjusted radiative forcing. The reason for allowing a stratospheric readjustment is that an increase in a greenhouse gas increases the net radiation absorbed in the troposphere. As a result, there is less upwelling radiation reaching the stratosphere (Fig. 14.2c). This causes cooling of the stratosphere, which decreases the net downward radiative flux from the stratosphere at the tropopause, contributing a negative component to the net radiative forcing attributable to an increase in a greenhouse gas (Wang et al., 1995). The time for the stratosphere to adjust is of the order of months (Manabe and Strickler, 1964), so that for the longer term perturbations of interest, this adjustment will occur and decrease the net radiative forcing from the instantaneous value.

This is illustrated by the data in Table 14.3, which shows the calculated instantaneous and adjusted radiative forcing attributed to the increase in tropospheric 03 from preindustrial times to the present (Berntsen et al., 1997). Table 14.3 also illustrates the larger relative importance of absorption of long-wavelength IR by ozone compared to short-wavelength UV. [Note that these calculations represent the contribution due to changes in tropospheric ozone only; as discussed earlier and in the following text, the destruction of stratospheric ozone leads to a negative forcing, i.e., cooling. Indeed, it appears that this cooling effect is likely dominant at the present time (Hansen et al., 1997b).]

Radiative forcing can be calculated for greenhouse gases in a fairly straightforward manner, particularly in the simplest case where there are no feedbacks or indirect effects on the chemistry of the atmosphere. However, translating these radiative forcings into real temperature changes at the earth's surface is much more uncertain due to the complex feedbacks involving

TABLE 14.3 Radiative Forcing" Due To Changes in Tropospheric Ozone from Preindustrial Times to the Present Time for Clear Skies Calculated Using the Oslo Modelb

Radiative

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