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The Effects on the Atmosphere of a Major Nuclear Exchange (1985)

Chapter:7 Atmospheric Effects and Interactions

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Suggested Citation:"7 Atmospheric Effects and Interactions." National Research Council. 1985. The Effects on the Atmosphere of a Major Nuclear Exchange. Washington, DC: The National Academies Press. doi: 10.17226/540.
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Suggested Citation:"7 Atmospheric Effects and Interactions." National Research Council. 1985. The Effects on the Atmosphere of a Major Nuclear Exchange. Washington, DC: The National Academies Press. doi: 10.17226/540.
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Suggested Citation:"7 Atmospheric Effects and Interactions." National Research Council. 1985. The Effects on the Atmosphere of a Major Nuclear Exchange. Washington, DC: The National Academies Press. doi: 10.17226/540.
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Suggested Citation:"7 Atmospheric Effects and Interactions." National Research Council. 1985. The Effects on the Atmosphere of a Major Nuclear Exchange. Washington, DC: The National Academies Press. doi: 10.17226/540.
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7 Almo spheric Effects and Interactions OVERVIEW The dispersion, evolution, and effects of dust and smoke injected into the atmosphere from a major nuclear conflict involve a large set of interacting processes whose complexity precludes detailed quantitative prediction at the present time. The available tools include a variety of models, of which the most advanced are the general circulation models (GCMs) developed for application to studies of weather prediction and climate dynamics. In these models, pressure, temperature, wind, moisture, and cloudiness fields are represented with a horizontal resolution of a few hundred kilometers and at a number of tropospheric and stratospheric levels (see, for example, Gates and Schlesinger, 1977; Mahlman and Moxim, 1978; Washington, 1982~. Smaller scale processes such as microscale and mesoscale turbulence, convection, gravity waves, local topography, and land-sea circulations can only be treated parametrically. Nevertheless, several of these models provide realistic simulations of the present climate. For applications to the problem of atmospheric effects of dust and smoke from nuclear war, however, GCMs are deficient in several respects. Transport of trace gases and diurnal variations have been simulated in some GCM studies (Levy et al., 1980; Cess et al., 1984; MacCracken and Walton, 1984~. However, no existing GCM simulates the full physics of a radiatively active trace material where net heating effects drive the circulation while the distribution of material is itself continuously varying in response to the flow and to complex flow-dependent removal processes. Formulations of boundary layer processes in these models are necessarily somewhat crude because of the low spatial resolution. Some recent model calculations have included particulate transport and diurnally varying absorption of solar radiation by the particulates, but these calculations have thus far had very limited vertical resolution (Cess et al., 1984; MacCracken and Walton, 1984~. Perhaps most serious for the nuclear war particulate problem, the cloud microphysical processes that are primarily responsible for the removal of per ticulates from the atmosphere cannot now be included in detail in these models. Other more specialized models can be applied to aspects of the problem, for example: cumulus-scale and mesoscale circulation models, 127

128 some with crude treatments of cloud microphysics, could be used to investigate specific processes that occur at scales smaller than that of the GCM grids. One-dimensional {vertical) radiative-convective models coupled to particle microphysical models have been used for detailed investigations of these critical processes, and, because of their computational efficiency, such models are extremely useful for sensitivity studies. Two-dimensional circulation models, though far less realistic than GCMs, can simulate the zonally symmetric components of the flow and the corresponding transport and radiative heating effects of nuclear per ticulates. Because they are relatively convenient computationally, they can be used for sensitivity studies, and therefore provide a valuable complement to Gems. Energy balance climate models (EBCMs) make up another class of relatively simple model that can be used to investigate radiative perturbations of surface energy balance and surface temperature (e.g., Sellers, 1973; Robock, 19831. Most such models deal only with the energy balance at the surface, and horizontal heat transport is modeled as a diffusive process with diffusion coefficients chosen to provide reasonable simulations of the present climate. Consequently, results from such models must be interpreted judiciously. The advantage of EBCMs is that, because of their computational efficiency and modeling of horizontal variations, they can be used to provide an indicator of the feedback effects of such relatively persistent climate factors as snow and ice albedo, sea ice cover, and sea surface temperature. Some of the principal results that are now available from one-, two-, and three-dimensional models are displayed in Tables 7.3 and 7.4. Of necessity, the results of simulations using models constitute the core of our knowledge of the likely atmospheric effects of smoke and dust from a nuclear war. In discussing these results, it is convenient to divide the problem into several subdivisions: early spread and evolution of the particulate clouds, direct optical effects, thermal effects as calculated by one-dimensional (vertical) models, thermal and circulation effects calculated by multidimensional models, and modification of circulation, cloudiness, and precipitation fields by the radiation perturbations induced by these particulate clouds. Several of these are rapidly evolving areas of research, and it should be clear that parts of this chapter may be superseded TV new developments in the near future. In the absence of observational analogs of the atmosphere as perturbed by nuclear war, observations of related, though inevitably very different atmospheric situations must be used. Several such partial analogs are discussed near the end of this chapter. The global-scale atmospheric perturbations associated with major volcanic eruptions and with plausible meteor impact events and their relationship to nuclear war scenarios are considered in the following chapter. EARLY SPREAD AND EVOLUTION OF PARTICULATE CLOUDS The area initially covered by the smoke plumes would depend on the number of fires, the cross-wind width of each fire, the average wind

129 speed, the directional variability of the wind near the level of plume stabilization, the duration of the fires, and the overlap among fire zones. If urban fire plumes extended into the middle troposphere, they would be transported by winds whose average speeds are of order 20 m/s, so that fires of several hours duration would produce plumes several hundred kilometers in length. For this reason alone, it is reasonable that substantial fractions of Eurasia, North America, and the North Atlantic, would be covered initially by smoke plumes. Crutzen et al. (1984) have estimated that the initial area covered by smoke plumes would be between 1 x 107 km2 and 2 x 107 km2 for a scenario similar to the Ambio scenario (Ambio, 1982~. For the committee's 6500-Mt scenario with about 1000 urban mass fires, an initial coverage area (immediately following the phase of rapid burning and plume rise) of about 107 km2 seems to be reasonable. In a statically stable atmosphere subject to solar heating, local wind systems would develop in response to the differential heating associated with nonuniformities in smoke distribution, and these winds would tend to smooth out both the thermal perturbations and the smoke nonuniformities. Such forced circulation systems were found to be effective smoothing agents in a cumulus-scale circulation model with an initially nonuniform distribution of carbon black (Chen and Orville, 1977~. The committee is not aware of similar numerical experiments at larger scales, but there is good reason to believe that such wind systems would be effective at scales out to several hundred kilometers. This is the typical scale of the Rossby radius of deformation.* At this and larger scales, the effect of earth's rotation becomes important and would impose a structure that could partially restrict such thermally forced lateral spreading of the smoke. Nevertheless, many of the smoke-free holes originating over the North American and Eurasian continents between 30°N and 70°N latitude would be filled within the first 2 days. After about 3 days, under typical meteorological conditions, the major gap over the Atlantic in the 30°N to 60°N latitude belt would be largely filled and very likely would have drifted over Western Europe. Portions of the mid-latitude Pacific would also be covered. The speed of this further spreading would depend somewhat on season, being greater in winter and smaller in summer. Figure 7.1 shows a specific winter season realization of the smoke and dust distribution after 3 days, based on winds derived from the Oregon State GCM (Gates and Schlesinger, 1977), and a nuclear war smoke injection scenario somewhat similar to the baseline case (MacCracken, 19831. The initial injections for this case were 207 Tg soot and 118 Tg dust. The feedback between the radiatively induced perturbation to circulation and particulate transport was not included; it was, however, included in a more recent *The Rossby radius of deformation for mid-latitude disturbances driven by heating in the mid-troposphere is (N/f)H, where N ~ 10-2 s- is the frequency of buoyancy oscillations, f ~ 10-4 s-1 is the Coriolis frequency, and H ~ 0.7 x 104 m is the scale height (e.g., Holton, 1979~. Hence the Rossby radius is about 700 km.

- 130 _ \ FIGURE 7.1 Hemispheric distribution of smoke-induced optical depth 3 days after a hypothetical nuclear exchange. (From MacCracken, 1983.) calculation using the Oregon State GCM, which produced quite similar results (MacCracken and Walton, 1984~. The spreading of smoke is probably underestimated in the calculation shown in Figure 7.1 because vertical wind shear has been neglected and the thermally forced smoothing and spreading of the smoke have not been taken into account. Nevertheless, smoke and dust cover much of the northern mid-latitude region. According to this calculation, there are large patches tof order 106 km2 in area) in which optical depth exceeds 20 at 3 days after the start of fires, but approximately 20 percent of the area of the hemisphere (about 40 x 106 km2) is already covered by smoke and dust with optical depth of 5 or more. The initial area covered by stratospheric dust, corresponding to the area occupied initially by stabilized nuclear clouds, would be much smaller, about 0.4 x 106 km2 for the baseline case. Although dust absorbs solar radiation far less efficiently than smoke, the heating per unit mass of air would still be significant at the lower densities of the stratosphere. Thus these clouds would also tend to spread laterally in response to their self-induced thermal circulation. Calculations of stratospheric dust cloud dispersion for a nuclear war scenario involving counterforce strikes show distributions qualitatively similar to that in Figure 7.1 when climatological mean midwinter winds are used (B. Yoon, private communication, 1983~.

131 Dispersion would be faster with actual time-dependent winter winds, but during summer, spring, and autumn, zonal winds in the extratropical lower stratosphere are weaker, and dispersion would be correspondingly slower. Material injected above 18 km in midsummer would drift westward (e.g., Holton, 1975~. As discussed on pages 77 to 80, nuclear smoke clouds would be subject to early rainout and coagulation of particles during the initial plume rise phase, but the effectiveness of these processes would rapidly decrease after the clouds have stabilized and begun to spread out in horizontally stratified plumes. Crutzen et al. (1984), using a simplified model, found less than a factor of 2 increase of particle mode radius during the 30 days following the initial rapid rise phase of the fire plumes. Coagulation in slowly dispersing smoke clouds was also evaluated by Turco et al. (1983b). In a case intended to maximize the Brownian coagulation rate, they assumed initial plume coverage equal to that of the stabilized nuclear clouds (about 106 km2 for their baseline case); they also assumed slow horizontal diffusive growth such that coverage increased linearly with time, reaching 20 x 10~ km2 only after 20 days. For the reasons cited above, this spreading rate is unrealistically slow, but even with these extreme assumptions, average smoke particle radius was found to increase by only about 65 percent after 1 week. For spherical particles whose initial radii are <0.4 um having an imaginary refractive index (the absorption component of refractive index) of <0.1, such size increases cause a decrease in the absorption coefficient per unit mass of less than a factor of 2 (Bergstrom, 1973; Lee, 1983~. For smaller or more weakly absorbing particles and for infrared radiation, the effect of such a size change is smaller. As will be shown below, early temperature changes near the surface are not very sensitive to variations of a factor of 2 or less from the absorption coefficient per unit mass of the baseline smoke injections. This is because the baseline injection initially contains more than enough smoke to absorb almost all sunlight in the areas affected by the smoke cloud. The duration of direct thermal effects of the particulates is more sensitive to the absorption coefficient, however. In addition, factor of 2 reductions below the baseline in several quantities (e.g., initial injected mass and absorption coefficient) would affect even the early temperature changes. Thus coagulation and early rainout are very important and complex issues requiring additional research. Longer term chemical and physical modification, or "aging," of aerosols in the atmosphere is another area on which additional basic information is needed. Because elemental carbon is hydrophobic and unreactive in the atmospheric temperature range, this may be a slow process for soot, depending on coalescence with preexisting hydroscopic particles. When coalescence occurs, the resulting particles behave as hydroscopic particles and can grow further by adsorption of water (Ogren, 1982; Ogren and Charlson, 1983~. Because of the internally mixed elemental carbon, these composite particles would still be efficient absorbers of sunlight (Ackerman and Toon, 1981), but an increase in composite particle size due to aging could have an important effect on the ratio of absorption efficiencies at visible and infrared

132 wavelengths. Since this ratio is an important factor controlling the influence of per ticulates on net radiation, the aging issue requires careful additional scrutiny. DIRECT OPTICAL EFFECTS Figure 7.2 displays the transmission of visible sunlight, including diffuse as well as direct radiation, as a function of smoke and dust opacities for particulates having the size and refractive index properties specified in the baseline case. Dust and smoke properties for the injections of the baseline case have been described and presented in Chapters 4 and 5 (readers are referred particularly to pages 27 to 32 and Table 5 .7 ~ . For convenience , the baseline in jection parameters are summarized in Table 7.1.* For these optical properties, light levels decrease very rapidly for smoke optical depths greater than one. When these light level reductions are combined with the extinction optical depths calculated by MacCracken (1983), and illustrated in Figure 7.1, the result is that light levels for much of the continental area north of 30°N would be reduced below the limit of photosynthesis during the first week, and widespread dense patches of smoke would make seeing impossible for several days after the nuclear exchange. For the NRC baseline case, with smoke and dust assumed to be instantaneously dispersed to a uniform distribution over the 30°N to 70°N latitude belt, average light levels for the belt would be below those for a very cloudy day (about 10 percent of the normal clear sky illumination) for about 2 weeks after the exchange. This can be seen by comparing the total downward solar flux versus time for this case (Figure 7.3) with the transmission levels shown in Figure 7.2. The values shown in Figure 7.3 were calculated using the one-dimensional model of Turco et al. (1983a,b). As explained by these authors, this model combines a detailed radiative transfer model with a detailed particle microphysics model (Pollack et al., 1976; Toon et al., 1979; Turco et al., 1979; Ackerman and Toon, 1981; Pollack et al., 19831. There is an approximately exponential dependence of the total downward solar flux on smoke opacity when full allowance is made for multiple scattering, as shown in Figure 7.2. This is largely because of the high absorptivity of the smoke. As a consequence, a saturation effect occurs such that most of the solar flux is removed by a smoke optical depth as small as 2; further increases in smoke optical depth *Readers unfamiliar with radiative transfer theory may wish to consult Liou (1980), which describes the theory and computational approaches in detail. tThe abbreviation NRC is used to denote the committee's baseline and excursions; LLNL denotes Lawrence Livermore Laboratory (e.g., MacCracken, 1983), and TTAPS denotes Turco et al. (1983a,b).

O Very Cloudy Day \ En .10 _ ~_ 6 \ \ 10-2 133 .75 .50 _ 25 \ Smoke \ _ Dust 10-4 10-6 10-8 1 o~ 1 0 Limit of Photosynthesis _ \ F ul I Moonl ight \ _ Limit of Human Vision \ - , , , , 1 ~, , it, 1 1 2 4 6 8 10 20 40 60 80100 OPTICAL DEPTH FIGURE 7.2 Fraction of incident solar radiation reaching the surface as a function of extinction optical depth for smoke and dust particulates with optical properties as in the NRC baseline case (Table 7.1~. Solar zenith angle of 60° is assumed. Diurnally averaged illumination depletions would be somewhat smaller at latitudes and seasons with smaller minimum zenith angles. These calculations use the radiative transfer algorithm of Pollack et al. (1976, 1983), in which full account is taken of multiply scattered radiation (cf. Pollack et al., 1983, and references therein for a fuller description). Note that the vertical scale is logarithmic. have relatively little additional effect on solar flux received at the surface. This saturation effect carries through to the temperature changes computed by one-dimensional radiative-convective models (Turco et al., 1983a) and energy balance climate models (Robock, 1984), since the degree of cooling at the surface predicted by these models is not very sensitive to variations in illumination at very low light levels, and these models do not allow for gaps and nonuniformity in the smoke. Because of the high absorptivity, smoke clouds produce much larger depletions of solar radiation than water clouds or dust clouds of comparable extinction optical depth. However, even for a relatively moderate depletion in surface illumination comparable to that produced by dense water clouds, smoke clouds would have a larger effect on the surface thermal balance than water clouds. This is because water clouds have a high ratio of infrared to visible opacity so that increased

200 150 NRC Baseline '~ NRC Baseline. Hemisoheric 100 l 1 / 1/ 50 o / , 1~1 1 1 1 1 1 1 1 0 25 50 75 100 125 150 175 200 - / 134 , Hemispheric, Fast Rainout _ _ _ __ NRC Baseline, 30° -70° N TIME (days) FIGURE 7.3 Time evolution of the total downward solar flux at the surface for the NRC baseline (30°N to 70°N), the NRC baseline injections spread over 0° to 90°N, and the NRC fast-rainout variant with injections spread over 0° to 90°N. downward flux of infrared radiation can equal, or even exceed (on a 24-h basis), the depletion of solar flux. Such compensation between solar radiation depletion and infrared radiation enhancement would not occur for smoke clouds because of their low ratio of infrared to visible opacity, except in regions where the normal daily total of solar radiation is already very low, such as is the case very close to the polar twilight boundary during winter, or would be the case at very early times following a nuclear exchange in dense patches in which the optical depth reaches values of 20 or more. The corresponding saturation regime is not reached for dust until the optical depth of dust alone reaches a value of about 12 (see Figure 7.21. For this reason, among others, the thermal effect of dust is far more sensitive than that of smoke to the nuclear war scenario. Smoke opacity is initially well within the saturation regime for the baseline smoke emission given in Table 7.1--180 Tg spread over the 30°N-70°N latitude belt--and approaches the margin of the saturation regime only as this value is decreased by about a factor of 4 to ~40 to 50 Tg. For smoke injections below this level, saturation no longer applies, and the light reduction and temperature effects would decrease rapidly with

135 TABLE 7.1 Properties of Injected Aerosols, NRC Baseline Case Dust (see pages 27 to 32) Smoke (see Table 5.7) Total injected mass (Tg) 15 180 Median particle radius rm rumba 0.25 0.10 Log normal dispersion ha 2.0 2.0 Refractive index (real part, 0.5 nm) 1.5 1.55 Refractive index (imaginary part, 0.5 um) 0.001 0.10 Extinction coefficient at 0.5 um (m2/g) 2.8 5.5 Absorption coefficient at 0.5 um (m2/g) 0.1 2.0 Infrared optical Wavelength-dependent Absorption only, properties basaltic glass (cf. cross section Pollack et al., 1973) 0.5 m2/g Vertical distribution 37% stratosphere Uniform mass per of injection 63% troposphere unit volume (see Table 4.1J between 0 and 9 km (see pages 73 to 76 and 83) Horizontal distribution Uniform in the Uniform in the of injection latitude belt latitude belt 30°N-70°N; none 30°N-70°N; outside none outside aParameters of the log normal size distribution; see page 62. decreasing injected mass. On the other hand, dust opacity approaches saturation only for rather extreme excursions that involve large numbers of surface bursts. For the dust optical properties and quantities of the baseline case, the extinction cross section of dust at 0.5 Am is 2.8 m2/g, and the corresponding extinction optical depths [lower limit (best estimate) upper limit for submicron dust in both troposphere and stratosphere] are [0.6 (0.9) 1.51 for dust uniformly spread around the 30°N to 70°N latitude belt. The extinction optical depths in the same

136 belt with the added opacity due 32) are [1.3 (2.1) 3.31. - well below the saturation threshold for climatologically significant since most of this dust is in the stratosphere and has a long residence time. In the baseline case, only about 40 percent of the dust is initially injected into the stratosphere, but the remainder may have an anomalously long residence time in the upper troposphere if precipitation is suppressed because of the smoke (see below). Crutzen et al. (1984) have also estimated transmission versus time for the smoke cloud. They consider models with rainout removal times of 15 days and 30 days. For the 15-day rainout time, calculated solar illumination reductions to the 10 percent level persist for 10 to 14 days (the exact value depending on the assumed extent of forest fires) by which time a uniform cloud has dispersed to cover 60 percent of the northern hemisphere, whereas for the 30-day rainout time the reduction to 10 percent persists for about 14 to 24 days. According to their estimates, about one-half of the northern hemisphere will have been covered by the smoke cloud in 10 days, and about two-thirds of the hemisphere in 20 days. These reductions correspond quite well to the NRC baseline case despite differences in the scenarios and in the treatment of cloud dispersion and evolution. to the 8500-Mt dust excursion (see page The values for the 8500-Mt excursion, though dust, may nevertheless be THERMAL EFFECTS IN ONE-DIMENSIONAL MODELS General circulation models can provide the most detailed and reliable assessments of temperature changes associated with nuclear war; however, because of their complexity and computational requirements, they are not suitable for sensitivity studies in which parameters such as input scenarios and particulate removal rates are varied over wide ranges. Turco et al. (1983a,b) have carried out such sensitivity studies using the TTAPS one-dimensional model. In order to relate the results of the TTAPS studies to the current baseline and to the results of multidimensional modeling studies using the NRC baseline, the TTAPS model has been applied to the NRC baseline case and to two variations: a rast-ra~nout removal case, and a case in which the baseline smoke injection is uniformly distributed over the entire northern hemisphere. The TTAPS one-dimensional model (Turco et al., 1983b) calculates the microphysical evolution of particulates subject to coagulation, agglomeration, sedimentation, vertical eddy diffusion, surface deposition, and removal by parameterized rainout processes (see Turco et al., 1983b, and references therein--particularly Turco et al., 1979, 1981; Toon et al., 1979; Hamill et al., 1982--for details). In the NRC baseline case the smoke and dust clouds have been assumed to be uniformly distributed around the 30°N to 70°N latitude belt. The microphysical implications of this simplifying assumption have been discussed in previous sections of this chanter and in Chanter 5. Results are most sensitive to the particle process, which is parameterized as a linear loss mechanism with a height-dependent exponential lifetime. Since the particle lifetime increases rapidly ~. · ~

137 with altitude, there is a strong interaction between the altitude of initial smoke plume injection and the assumed vertical profile of rainout rate. AS described on pages 73 to 76 the committee has assumed that the smoke is distributed uniformly with altitude over the O to 9 km range, partly for simplicity in the absence of better information, and partly because it is the committee's judgment that the intensity of urban fires would tend to drive the plumes into the upper troposphere. If vertical mixing in the plumes is very rapid, it would tend to produce a uniform smoke mixing ratio rather than uniform smoke concentration. However, as will be seen below, the tendency to develop a uniform mixing ratio would probably decrease rapidly with time and would be strongly opposed by the increase in the rainout rate near the ground. The rainout removal rate profile assumed for the NRC baseline case is given in Table 7.2, where it is compared with the profile used in the TTAPS study. The TTAPS group chose baseline values designed to represent the rainout characteristics of the unperturbed atmosphere; for the NRC baseline case, these values have been modified so that faster rainout occurs in the lower troposphere (O to 5 km) and no rainout at all occurs above 5 km. These changes have been made in order to simulate possible effects of changes in static stability and cloudiness expected in the perturbed atmosphere (see pp. 156 to 158 below), and they are of course highly uncertain. Even in the absence of rainout, however, eddy diffusion acts in the model as an effective mechanism for removing particulates from the upper troposphere. Following Massie and Hunten (19811, the vertical eddy diffusion coefficient value 10 m2/s has been assumed for the NRC baseline case, as it was for the TTAPS calculations. This value gives a characteristic lifetime against dry deposition for particulates in the upper troposphere of about 40 days. Because the rainout time and its interaction with the initial vertical smoke distribution are so critical to the evaluation of climatic effects, a fast-rainout excursion has been considered, with rainout times given in the last column of Table 7.2. These high values of rainout rate are believed to provide a reasonable case bounding the smoke lifetime on the low side for the NRC baseline smoke injection. In this case, the smoke has been assumed to be dispersed over the entire hemisphere rather than over the 30°N to 70°N latitude band. The initial opacity for this case is nearly equivalent to that for smoke and dust spread over the 30°N to 70°N latitude band with half of the initial smoke and dust injections of the NRC baseline case. The TTAPS ~slow-rainout" case, with an effective removal rate about one-third as fast as their baseline case, represents a plausible bound to the smoke lifetime on the high side. This case is also compared with the NRC baseline. Figure 7.4 shows vertical profiles of the contributions to optical depth from smoke and dust in each 2-km layer for (a) the NRC baseline case, and (b) the fast-rainout excursion. For comparison, the TTAPS baseline is also shown (Figure 7.4c). In the TTAPS baseline case, the relatively rapid rainout removal assumed for the upper troposphere causes the center of mass of smoke to lower over time, while in both the NRC baseline case and the fast-rainout excursion, the rapid downward

138 TABLE 7 . 2 Smoke Removal Rates (s-l) NRC "Fast- Altitude (km) TTAPS (1983, 1984) NRC Baseline Rainout" Case 0-1 1.0 x 10-6 4 .0 x 10-6 4 .0 x 10-6 1-3 8 .6 x 10-7 2 .7 x 10-6 4 .0 x 10-6 3-5 7 .1 x 10-7 1.3 x 10-6 2 .0 x 10-6 5-7 5.7 x 10-7 0 1.0 x 10-6 7-9 4 .3 x 10-7 0 1.0 x 10-6 9-11 2 .9 x 10-7 0 1.0 x 10-6 1 1-13 1. 4 x 10-7 0 0 NOTE: Values represent [(l/m)(dm/dt)l, where m is the mass of particulates per unit volume. Zeroes indicate removal controlled by eddy diffusion with K = 10 m2/s. Smoke from above 5 km diffuses downward to 5 km, where it is removed by the rates shown. Effective removal rates for the atmosphere above 5 km can be estimated from the data shown in Figure 7.4a, and amount to about 3 x 10-7 s-1 for the layer between 5 and 8 km for the first 30 days following smoke injection. increase in rainout rate quickly removes smoke in the lower troposphere, leaving an elevated smoke cloud. Note that the low initial values of opacity in the TTAPS baseline and in the NRC fast-rainout varient are due to the assumption of initial dispersion over the hemisphere for these cases. The NRC baseline also differs from the TTAPS baseline in that the latter has smoke in the lower stratosphere and a larger mass of stratospheric dust. These features appear as bulges near 15 and 25 km in the initial profiles of Figure 7.4c. The differences in shape below 10 km between the TTAPS and NRC cases are due to differences in assumed initial injections and removal rate profiles. Figure 7.5 shows the time evolution of the optical depth for the NRC baseline, the fast-rainout excursion, and the TTAPS baseline. The total optical depth with contributions from dust as well as smoke is shown. These results reflect the removal rates shown in Table 7.2. The radiative-convective component of the TTAPS model has been described elsewhere, and readers are referred to these sources for details (Pollack et al., 1976; Ackerman and Toon, 1981; Toon and Ackerman, 1981; Pollack et al., 19831. The model takes into account the size distributions and optical properties of smoke and dust, and explicitly calculates solar and infrared radiation including the effects of multiple scattering, absorption and emission by atmospheric gases, and the radiative effects of a prescribed representative distribution of cloudiness. Heat capacity of the underlying surface is neglected for land areas, but is included when the model is applied to ocean areas as it was in one of the cases considered by Turco et al. (1983a). The

139 30 25 [ 20 10 O / / / ~ r 0.0 0.5 1.0 (a) · = Day O = After 5 Days = After 30 Days = After 60 Days · = After 130 Days 0 = After 250 Days _~ JO 1.5 2.0 OPTICAL DEPTH (per 2-km layer) 2.5 3.0 FIGURE 7.4 Contr ibutions to total extinction optical depth for 2-km-thick layers at various times after hypothesized nuclear war smoke and dust injections. (a) NRC baseline, 30°N to 70°N; (b) NRC baseline, O ° to 90°N' fast rainout; (c) TTAPS baseline.

140 30 25 20 - UJ 15 o 0.0 FIGURE 7 . 4 (continued ) 9 10 to E 1 ~ .~ (b) o . 0 = After 253 Days = After O Days = After 4 Days = After 29 Days = After 61 Days = After 1 28 Days ~ _ t ~ ' l 0.5 1.0 OPTICAL DEPTH (per 2-km layer) 1.5

141 35t~ 30 25 20 10 5 - 1 0.0 FIGURE 7 . 4 (continued) (c) ~ r - 0.5 · = After O Days 0 = After 4 Days · = After 31 Days = After 62 Days = After 125 Days = After 255 Days 1/ V' - .~- -. 1 . ~ 1 .0 OPTICAL DEPTH (per 2-km layer) 1.5

142 9.0 80 70 60 50 40 cat o 30 20 1.0 O _ 0 25 50 75 100 125 150 175 200 \ \ -am\ N RC Basel ine, 30° -70° N NRC Baseline, Hemispheric NRC Baseline, Hemispheric, Fast Rainout TIME (days) FIGURE 7.5 Time evolution of the total column extinction optical depth for the NRC baseline (30°N to 70°N), the NRC baseline injections (0° to BOON), and the NRC fast-rainout variant (0° to 90°N). Contributions from both dust and smoke are included in the total. "surfaces temperatures given by the TTAPS model actually correspond to temperatures of a 2-km-deep lowest atmospheric layer. Because this has a significant thermal inertia, the neglect of land surface thermal inertia is probably not serious. Vertical mixing of heat occurs in the model only by convective adjustment. Figure 7.6 shows the time evolution of near-surface temperature (actually the mean temperature of the lowest 2-km-thick atmospheric layer) for the NRC baseline, the excursion from the NRC baseline in which the smoke and dust clouds are distributed over the entire hemisphere rather than the 30°N to 70°N latitude strip, and the fast-rainout excursion from the NRC baseline. These temperatures drop rapidly to minima that are reached at times ranging from 10 to 25 days after the initial nuclear conflagrations. In each case there is a slower recovery, with the point of 50 percent temperature recovery reached at times ranging from 20 to 75 days. These results are summarized in Table 7.3. For comparison, Table 7.3 also shows the results of several other calculations. The "no-fires case from TTAPS (Turco et al., 1983b) isolates the effect of

-10 - LL ~ -20 us -15 -25 -30 -35 o 143 NRC Baseline, Hemispheric, Fast Rainout, 0-90°N / / / / ~ /' ~ / NRC Baseline, 30° -70° N /: ~NRCBaseline, / ,/ Hemispheric, 0-90° N / 1 1 1 1 25 50 75 100 Tl M E (days) 1 1 1 125 150 175 200 FIGURE 7.6 Time evolution of the surface temperature for the NRC baseline (30°N to 70°N), NRC baseline (0° to 90°N), and NRC fast-rainout variant (0° to 90°N). The temperature shown is actually the mean for the lowest 2-km atmospheric layer in the TTAPS model. stratospheric dust alone, and has an initial dust optical depth of 1.4, so that it is within the bounds of the NRC 8500-Mt excursion for stratospheric dust discussed in Chapter 4. The LLNL one-dimensional model closely parallels the TTAPS one-dimensional model calculation (MacCracken, 1983~. Initial injections include 207 Tg of soot from urban and forest fires, and 118 Tg stratospheric dust. Also included are the optical effects from the injection of nitrogen oxides into the stratosphere, taken as 8.3 Tg N. together with the corresponding ozone reduction. The net effect is a slight addition to the absorption of solar radiation in the lower stratosphere. Results from this calculation are similar to those of Turco et al. (1983b). Differences arise in part from differences in the treatments of precipitation scavenging and in the assumed optical properties of the per ticulates. The one-dimensional models do not provide reliable estimates of mean, or even typical, temperature changes over continental areas. Better estimates are provided by multidimensional models, also summarized in Table 7.3, though these too are subject to large uncertainties. Results from these models will be discussed in detail below.

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148 40 _ 30 - I I U]' o CL Am: O _ - 0 10 20 30 40 50 60 70 80 90 100 TIME (days) FIGURE 7.7 Time-height section of the temperature perturbation (°C) for the NRC baseline case (30°N to 70°N). The stippled area represents negative values. The time evolution of the vertical temperature change profile for the NRC baseline is shown in Figure 7.7. As with the cases considered by the TTAPS group, the strongly absorbing particulate clouds produce rapid increases in temperature aloft, leading to a radical modification of the initial temperature profile, and replacing the normal tropospheric lapse rate with a deep and intense inversion. As a consequence of the optical saturation effect mentioned above, one-dimensional model responses are not sensitive to the precise amount of smoke injected, provided the optical depth is in the neighborhood of 2 or greater--nor is the magnitude of the temperature increase sensitive to the details of the vertical distribution of smoke provided that substantial amounts of smoke remain above about 4 km. Except for the NRC hemispheric fast-rainout excursion, in which the optical depth drops below 2 within 7 days, each of these cases shows a maximum temperature increase near 80 K. To illustrate this insensitivity, magnitudes, times, and altitudes of maximum temperature increases for several cases are given in Table 7.4. The essence of the temperature change tendencies indicated by the numbers in Tables 7.3 and 7.4 derives from the efficiency of the smoke particles as absorbers of solar radiation coupled with their inefficiency for absorption and emission of infrared radiation. The

149 TABLE 7.4 Elevated Layer Temperature Changes Maximum Temperature Time of Height of Increase Maximum Maximum Case Region Season {°C) (days) (km) NRC 30°-70°N Annual +85 43 11 baseline average NRC 0°-90°N Annual +65 47 11 baseline average hemispheric NRC 0°-90°N Annual +35 30 11 hemispheric average fast rainout TTAPS 0°-90°N Annual +95 55 17 baseline average ratio of the extinction efficiency of the smoke at a wavelength of 0.55 um to that at 10 um is about 10 for the NRC baseline case. This ratio follows from the model of the properties of the injected smoke in the baseline model discussed in Chapter 5. The ratio of absorption efficiencies is about 4. As long as the absorptivity of the cloud for solar radiation is sufficiently large that the altitude at which most of the solar radiation is absorbed is higher than the altitude from which most of the thermal infrared radiation is emitted, the normal warming of the surface due to the "greenhouse effect" will be cut off, and the "antigreenhouse. tendencies depicted in Figures 7.6 and 7.7 will appear in radiative convective calculations. This conclusion remains valid even for optical depths at visible wavelengths that are sufficiently large that the infrared opacity is also large, so long as the altitude of solar energy absorption is above the level of thermal infrared emission (T.P. Ackerman, private communication, 1984; Golitsyn and Ginsburg, 19843. In the limiting case of a globally uniform absorbing cloud in thermal equilibrium, with solar radiation absorbed above the altitude of thermal emission to space, the mean surface temperature would correspond approximately to the effective emission temperature of a planet in thermal equilibrium with the incident sunlight (about 255 K for the present earth) while the cloud temperature would be much warmer, the exact temperature depending on cloud thermal emissivity. Because of the optical depth saturation effect for smoke absorption discussed on pages 132 and 133, the one-dimensional model temperature changes are insensitive to changes from the NRC baseline in smoke

150 particle size distribution as optical properties. For the same reason, the results are not very sensitive to the initial mass of smoke input, provided the initial mass is large enough. Lower injected smoke mass (or smoke absorptivity) by a factor of about 4 would lie near the edge of the saturation regime and climatic effects would decrease rapidly for smoke injections in this range and below. In addition, the maximum temperature depression is sensitive to smoke lifetime if rainout rates are unexpectedly large, as in the NRC fast-rainout case. In this case, the smoke opacity may not remain above the strong absorbing threshold (optical depth of about 2) long enough for even approximate thermal equilibration of perturbed continental surface temperatures to take place e THERMAL AND CIRCULATION EFFECTS CALCULATED BY MULTIDIMENSIONAL MODELS The one-dimensional models discussed in the previous sections should not be expected to predict accurate temperature changes, especially near the surface. They are useful for evaluating the possible magnitude of the problem, and for efficiently examining the sensitivity of the thermal tendencies to various assumptions. AS mentioned above, the temperature changes predicted by these models should not be interpreted as mean or even typical changes for the regions given in the third column of Table 7.3. These models might provide upper limit values for the given smoke and dust inputs and removal rate assumptions, applicable perhaps to the deep interiors of continents where the ameliorating influences of the oceans do not apply. Even for this limited quantitative purpose, such results should be considered with some reservations, however, because patches with much larger optical depth might prevail for some time over localized regions of the continental interiors, and might lead to larger temperature drops. There are three serious limitations to the use of one-dimensional models to deduce representative temperature changes. The first problem is that they neglect the ameliorating effects of the nearly constant sea-surface temperatures on the climate over coastal areas and continental areas exposed to fresh flows of maritime air. This limitation would be most serious for the near-surface temperature changes, because the temperature increases due to absorption of solar radiation aloft would be likely to develop and persist over the oceans as well as over continents. The second problem is somewhat more subtle. As pointed out by MacCracken (1983), the nonuniform distribution of dust implies that some regions would have very high Optical depths while others would be nearly clear (note the patchiness in Figure 7.1, for example). Thus fewer smoke particles would be exposed to sunlight than would be the case if the same amounts of smoke and dust were uniformly distributed, and as a consequence the absorptive power of the per ticulates would be less efficiently utilized; the uniform optical depth assumption tends to maximize the overall mean thermal tendencies. A third problem is that these models do not predict or account for changes in water cloud distributions, which could in turn

151 have an additional effect on the radiation balance. That this effect could be important is illustrated by a one-dimensional model calculation of MacCracken (1983), in which a 33 percent higher induced surface temperature drop occurs for a case without water clouds than for the corresponding case with normal clouds. In one attempt to account for these effects, MacCracken has used a two-dimensional meridional plane climate model to calculate temperature changes for particulate optical depths specified as meridional averages obtained from the Oregon State GCM (Gates and Schlesinger, 1977~. Transport and removal processes in this calculation were accounted for using the LLNL two-dimensional model (this model used the LLNL "GRANTOUR" code; cf. Walton and MacCracken (1984~. The initial particulate injections were, as mentioned above, 118 Tg of dust and 207 Tg of smoke. The circulation responsible for particulate transport was represented by only a single mid-troposphere wind, but the climate model included algorithms for predicting meridional circulation, cloudiness, precipitation, and the hydrological cycle based on weighted means of land and ocean over each latitude belt (MacCracken, 1983~. The result is summarized in Table 7.3 for the normal rainout parameterization of the LLNL two-dimensional model, calibrated to represent the unperturbed atmosphere. The maximum surface temperature decrease for this model was much less than for the one-dimensional LLNL model. This difference is primarily due to the treatment of the exchange of heat between continental and oceanic regions and to the more realistic nonuniform and evolving spatial distribution of smoke in the two-dimensional case. The calculated 30°N value should not be taken to represent the extreme temperature perturbation that might occur at mid-latitude continental locations with this scenario; the extreme may be substantially larger. MacCracken has also carried out a two-dimensional calculation with the rainout removal rate suppressed to simulate the possible effects of the nuclear war smoke and dust injections on precipitation, and in this case somewhat larger and much more persistent temperature perturbations developed. A GCM calculation has been carried out by Covey et al. (1984) using the NCAR Community Climate Model (CCM; Washington, 1982~. The horizontal resolution of this model corresponds to approximately 4.5° latitude and 7.5° longitude, and the model has nine layers in the vertical. Other details are given in Covey et al. and references therein. Smoke was introduced as a perturbation to initial atmospheric states corresponding to the undisturbed climate for winter, spring, and summer seasons. The total amount of smoke, its distribution, and its absorptivity for solar radiation corresponded closely to the NRC baseline case: nearly uniform concentration contained in the 27°N to 71°N latitude and 0 to 10 km altitude band, with an absorption optical depth of 3.0. Dust injection was neglected. Multiple scattering and infrared absorption and emission by the smoke particles was also neglected, and the smoke distribution was not allowed to change with time. The NCAR CCM accounts for absorption and emission by atmospheric gases at all wavelengths and scattering and absorption by water clouds. The distribution of water clouds (liquid or ice) is calculated based on

152 relative humidity distributions and convection predicted within the model. Zonally averaged temperatures for the summer period from 10 to 20 days after smoke injections case averaged over the are shown in Figure 7.8. There is a large elevated temperature maximum that is much like that predicted by the one-dimensional models, except that it spreads in an attenuated form over a far wider latitude range than that of the smoke. The continental average surface temperature change 10 days after smoke injection averaged over the 30°N to 60°N latitude band was -26°C for the summer case and -17°C for spring (Table 7.3; Thompson et al., 1984~. Temperature changes at the surface averaged over days 6 to 10 following the nuclear attacks are shown in Figure 7.9. By day 10, in this case, subfreezing temperatures extend over much of the Eurasian and North American continents, although the western part of Eurasia and coastal strips of North America are not so seriously affected. As Covey et al. point out, this calculation also shows the rapidity with which subfreezing temperatures might develop under large patches at relatively low latitude. This model predicts a number of meteorological effects in addition to temperature change, as discussed in the next section. Another GCM calculation has been carried out at the Computing Center of the USSR Academy of Sciences by Alekeandrov and Stenchikov (1983), using a modified version of the Oregon State two-level model. The calculation starts from simulated normal annual mean conditions with clouds of absorbing particulates injected into both the troposhpere and the stratosphere. In this calculation, the particulates are assumed to be distributed uniformly over the entire northern hemisphere poleward of 12°N. Only absorption of solar radiation is taken into account, and the initial optical depth for absorption is about 7, so that, if one allows for the areal extent of the cloud, the initial mass of material corresponds to about 4 times that of the NRC baseline case. The smoke lifetime is also assumed to be rather large with the optical depth remaining above 3 until day 100. This calculation produces a decrease in zonal mean surface temperature of as much as 22°C at 65°N on day 40, and a strong temperature inversion throughout the northern hemisphere. Extreme surface temperature changes produced by the model are quite localized, with extreme departures exceeding -40°C by day 40 in patches throughout the land-covered area of the northern hemisphere. Based on the one- and two-dimensional model results discussed previously, these results on day 40 are reasonable, given the very large initial injections and slow removal rates of this calculation. The initial zonal mean state and the temperature distribution after disappearance of the smoke and dust clouds at 360 days show unrealistic features, however, possibly due to the low horizontal and vertical resolutions of the model used for this early investigation of the problem. Discussion, interpretations, and comparisons of the results of the CCM and the USSR Academy of Sciences models are given in Thompson et al. (1984~. In comparing the one-, two-, and three-dimensional model simulations for similar scenarios and taking seasonal effects into account, the multidimensional models give average continental surface Toolings that are smaller than those given by the one-dimensional models by factors of

153 1 5 10 20 30 50 70 95 1 5 10 J in en ~ 20 cr 30 50 70 95 1 - 5 UJ cr 10 20 30 50 70 95 , ~ ' ' 1 ' ' 1 ' '1 ' ' 1 '/ ' Pertu rbed Case230 ~ I ~ ' ~ (a) `~;Ji 2~` _ / 220 - Hi\ 21'0 if= ' ' 1 ' ' 1 ' Control | /~ -2l0~JC - :~ L: -1 07: (,~\~,~,~: (Perturbed) (c) :~: -(Control) :40'33 :20 ~11,, 90 60 30 0 -30 LATITUDE (dog) -60 -90 FIGURE 7.8 Meridional temperature cross-section for the perturbed case, the control case, and the perturbed minus the control at t = 10 days from the CCM calculations for summer. The vertical scale is pressure, with 10 kPa corresponding approximately to 15 km and 7 kPa to 30 km. (From Covey et al., 1984.)

154 Perturbation Minus Control (Days 6-~0) i-:-: ~ <-~0°C ~ <-30°C 60° 3oo no 9oo 60° 3oo Go If:;' APRI L A, JULY ~ x_ ~ ~ , ,,, , . '.. . . . . . a. . . .~ . ~ ~ .. A.... :~, d ~ %~ >~53 \ a, a, ~o _~8OO _90o 0° 90° 180° FIGURE 7.9 Surface temperature change from the control after 10 days for a summer run of the CCM. (From Thompson et al., 19841. 2 to 3. This difference, which was anticipated by the TTAPS group, can be understood primarily as a consequence of the exchange of heat between ocean and continent as it is represented in the two- and three- dimensional models; in addition, the evolving smoke distribution plays a role in MacCracken's two-dimensional calculation. The extremes of mid-continent surface temperature decrease found in the NCAR CCM calculations are comparable to those occurring in the one-dimensional calculations, and this is consistent with what one would expect given the limitations of the one-dimensional models, for scenarios in which dense clouds of smoke remain over the continental interiors for the order of 10 days or more. It is worthwhile to note the possible effects of some of the improvements in this type of calculation that future research should address. The inclusion of scattering of sunlight and infrared absorption and emission, including the optical effects of injections of dust, would produce different predicted temperature changes. Multiple scattering and inclusion of dust would probably lead to somewhat lower near-surface temperatures, since the particulate cloud would have a higher albedo than the cloud assumed in the CCM simulation. The initial cooling of continental areas is faster in the CCM calculation of Covey et al. than in the calculations using the TTAPS one-dimensional model. This difference is almost certainly due to differences in the treatment of near-surface thermal inertia and vertical resolution. The "surface" temperature in the TTAPS model corresponds to the lowest 2 km of the atmosphere (see page 142~. As a

155 result, this model does not allow for the possibility of very shallow cold layers over dry continental surfaces that form at early times in the CCM calculation. Accurate treatment of the near-surface boundary layer is a prerequisite for reliable estimation of surface temperature changes. The neglect of inferred absorption and emission due to the smoke in the CCM calculations could lead to exaggerated cooling rates, but detailed calculations show that this effect is not important at early times for the NRC baseline injections (Ramaswamy and Kiehl, 19841. On the other hand, it could be significant during winter at middle and high latitudes, since the contribution of thermal infrared radiation to the surface energy balance is more important relative to solar radiation during winter. Infrared emission could also significantly affect cooling rates if smoke particle sizes are much larger than those of the NRC baseline. Simulations that permit the transport of particulates should have a high priority in future research. Such models allow calculation of the early spread of smoke and dust clouds and explicit calculation of the feedback loop between the particulates and circulation; not only would transport spread the cloud, but the enhanced circulation driven by the heating produced by the cloud could accelerate the spread. Rainout and aging of smoke particles could have significant ameliorative effects on the thermal perturbation after 10 to 20 days. Too little is currently known about fundamental aspects of these processes to permit convincing modeling at present, although parameterized treatments of rainout could be incorporated with relative ease. In addition, diurnal variations should be explicitly included in future studies, not only because the diurnal near-surface temperature extremes are of practical importance, but also because diurnal heating and temperature variations would be associated with diurnal variations of wind and vertical mixing that could be quite important for particulate transport and removal. Diurnal variations should even be included in future studies involving one-dimensional radiative- convective model calculations. The climatic impact of smoke and dust injections is slightly smaller when diurnal variations are explicitly simulated in these models (Cess, 1984~. Finally, seasonal variations need to be considered in more detail. The one-dimensional models have, at the time of writing, tended to focus on annual mean conditions, as did Alexabdrov and Stenchikov. Covey et al. carried out a calculation for winter using the CCM, but did not report in detail on the results. Robock (1984) has used an EBCM to simulate nuclear exchanges in autumn and winter as well as spring and summer. However, in this calculation, as in the CCM calculation, infrared radiation absorbed and emitted by the particulates was neglected. Because of the relatively important role played by thermal infrared radiation during winter, the significance of these results is not clear. For this reason, as well as because winter is the dormant season in middle and high latitudes, the ecological effects of the smoke and dust clouds could be less serious during winter. Winter cases deserve serious investigation, however, because a postwar population already exposed to the rigors of winter could be particularly sensitive

156 to the additional darkness and persistent cold that could be produced by nuclear smoke and dust clouds. Moreover, transport and dispersion rates are normally greater during winter, so that tropical and subtropical regions whose ecologies are particularly sensitive to meteorological excursions could be at greater risk (but see the discussion in the next section of the possibility of greatly enhanced meridional transport during summer and spring). MODIFICATION OF CLOUDINESS, PRECIPITATION , AND WINDS The strong thermal effects indicated by studies reviewed in preceding sections would certainly produce large changes in other climatically significant quantities. The thermal effects are themselves uncertain, and deductions about any consequent effects are necessarily even less certain, and must be considered somewhat speculative. Nevertheless, in this section an attempt is made to assess such effects, drawing on analogies to known meteorological phenomena and, where possible, on available model results. The following phenomena are considered: fog, cloudiness and precipitation distributions, zonal mean winds, other large-scale wind systems, and ultra-high clouds. Ground Fog Under the influence of large-scale dust and smoke clouds, radiation fogs would form over land areas as the surface temperatures dropped below the dewpoint. Initially at least, these fogs could provide some protection against further temperature decreases, particularly in affected tropical or subtropical regions where dewpoints are normally high. The lifetimes of such induced radiation fogs and the amount of thermal protection they would provide are uncertain. Normally, when radiation fog is persistent over a period of days, there is a tendency for thermal balance, in the diurnal mean, between the competing effects of cooling of the entire foggy layer by emission of infrared radiation, heating due to that portion of the incident sunlight not reflected by the fog and therefore absorbed at the ground or in the layer, and entrainment of warm air from above.* Such fogs are usually most persistent under otherwise clear skies during winter with strong subsidence above the foggy layer and near-stagnation of the low-level winds. If some mixing occurs in the surface layer, the fog can lift to form a low stratus layer. Such fog and stratus layers have been known to persist for as long as a month in California's Central Valley. Under postnuclear war smoke and dust clouds, sunlight would be virtually absent, and ventilation conditions would probably be at least as variable as in the unperturbed *While fogs are forming, they release latent heat, but this effect would be small in comparison with the radiative perturbations of the nuclear war scenar iOS.

157 atmosphere. Consequently, one could expect that fog or low stratus layers would persist in some places, but would be rapidly removed in others; and that, where persistent fog did occur, it would be somewhat less effective in ameliorating surface temperatures than similar fogs for the unperturbed atmosphere. There is little hard evidence bearing on this question. The CCM calculations of Covey et al. and the two-dimensional model results of MacCracken do show increases in low-level cloudiness in the perturbed regime, but improved boundary layer formulations are needed for a more precise assessment of the role of fog. Because such fogs may be persistent and have ameliorative effects, the sensitivity of surface temperature changes over extensive continental areas to particulate lifetime could be more pronounced than indicated by the considerations of the preceding sections. Cloudiness and Precipitation The heated elevated layer produced by absorption of sunlight in a widespread smoke cloud would suppress convection and prevent the formation of clouds within the layer because of the increase in static stability and lowering of relative humidity associated with high temperatures in the layer. Since precipitation normally forms only in deep clouds, it would be suppressed as well, at least over the continents. The situation over the oceanic regions covered by particulate clouds would be quite different from that over the continents. While the continental boundary layer would be statically stable, the oceanic boundary layer would generally be unstable, at least in regions affected by flow of air from the cold continents over the still-warm surface waters. As a result, enhanced convection would be likely over large areas of the ocean, especially adjacent to continents. Because of the high-temperature layer aloft, this convection would probably be shallow, however, and would perhaps be similar to that observed when very stable cold polar air from the continents flows over adjacent oceans. These convective boundary layers are rarely more than 2 km deep (Walter, 1980~. Clouds of this depth can produce showers, however, especially if temperatures are subfreezing below the cloud top level. Thus such oceanic convective boundary layers could be regions of frequent precipitation and effective particulate removal. Coastal regions might also be regions of enhanced cyclonic storm activity because of the increased temperature gradient between land and sea. This effect could add to the effectiveness of these areas as sites of particulate removal. However, if there were a well-developed elevated heated layer, it is unlikely that even the cloudiness associated with such coastal storms could penetrate to very great heights. Convective cloudiness and precipitation could be enhanced near the edges of major smoke cloud bands, and in the vicinity of outlying streamers. _ conditions, small rain showers could be produced by seeding the atmosphere with carbon black particles. The effectiveness of Chen and Orville (1977) have shown that, under suitable

158 precipitation developing along smoke cloud edges is extremely difficult to estimate. This process is likely to be most effective at early times when the cloud edge-to-area ratio is largest and before a widespread hot elevated layer has had a chance to develop. The possibility of enhanced precipitation during this stage of smoke cloud evolution constitutes one of the major uncertainties in the analysis. AS Figure 7.8 shows, full development of the stabilizing warm layer extends far beyond the edges of the smoke cloud. This effect, which is due to the tendency of subsiding motion in the cloud environment to compensate rising motion in the cloud region, is likely to suppress deep convection near the cloud edges after a week or two. Zonal Mean Winds Changes in the zonally averaged west-to-east winds are relatively easy to calculate, given any set of temperature field changes, since wind changes will be in thermal wind balance* with the temperature changes. For the summer GCM simulation of a nuclear war scenario by Covey et al., the predicted changes in the mean west-to-east winds are shown in the bottom panel of Figure 7~10. These changes correspond to the temperature changes shown in the bottom panel of Figure 7.8. The main features are strongly enhanced west-to-east winds in the region normally occupied by the stratosphere at middle and high latitudes of the southern hemisphere and high latitudes of the northern hemisphere, and very strongly enhanced easterly winds above the subtropical {northern hemisphere) edge of the smoke. Each of these features can be understood in the light of the thermal wind relation as it applies to the temperature perturbation field. Their effects would be to greatly increase the rate of zonal transport of particulates, especially if some of the particles were to spread upward and equatorward over time. There is more uncertainty associated with predictions of the change in the Tonally averaged meridional component of the flow, but, given the altered heating field, predictions can be made with reasonable confidence (see Held and Hou, 1980, for a discussion of the theory of the zonally averaged circulation). For the spring and summer simulations of Covey et al., the nuclear-war-induced changes are dramatic. Figure 7.11 shows the stream function for the meridional circulation mass flow for the spring control and perturbed cases. Intense heating along the tropical edge of the cloud has caused the normal two-cell circulation to be replaced by a single cell with rising motion near the southern edge of the cloud. The intensified heating has also intensified and elevated the meridional flow toward the southern *That is, conditions that are completely determined by the temperature and surface pressure distributions. This relationship is a consequence of the close balance between Coriolis and pressure gradient forces in large-scale terrestrial wind systems (geostrophic wind balance). It should be noted that the true temperature change field inevitably includes the effects of dynamical as well as radiative processes.

159 - - llJ on en LL ~ 20 Cal 10 30 50 70 _ ~ I 95 ~F-~-r-r- 920 H 25 -~3 ~ I (a) | 40 Perturbed 2~0W{t'i_ , 1, 1 5 10 20 30 ~0 70 95 - - cr oh 20 UJ 30 50 70 95 1 :1 1 1 ~ 1 ~ I 40 , (b) Control / l H 11 \L] :> ~\~: l 20 i _ / / ~ 36 ~ .1 , , 1 , ,1 l ~ I ~ F 1 (Perturbed) ' -(Control)' 0; ll ~ _: _ H 20 1, _ 20 35 A--\, I ,~ - --~-w~ , 90 60 30 0 -30 LATITUDE (deg) FIGURE 7.10 Zonal wind cross sections for the perturbed case, the control case, and the perturbed minus the control at t = 10 days from the CCM calculations for summer. Isolines are labeled in meters per second with positive values eastward; westward winds are shaded. (From Covey et al., 1984.)

160 APRIL CONTROL (Days 16-20) - A - LL ~10 on en 20 50 100 1 - - UJ ~10 on cc 20 5 50 MAT ' ' ' ' I i' ~0 30 25 I \ o o \ \ 20 ~ I 15 ~ I 10 ~/ I I 11 I, ~ 1 , , 1 , , 1 , , 1 1 , 1 , , 1 APR I L PE RTU REED (Days 16-20) ' ' 1 ' ' 1 ' ' 1 ' ' 1 ' ' 1 ' ' :0 j:.-05~ 100 _ 90 60 30 0 -30 o 30 25 / -15 :' J 1~ :)~. ~°~ ,, 1,, 20 ~ - 15 I 111 0 I 5 O -60 -90 LATITUDE (deg) FIGURE 7.11 Meridional circulation mass flow stream functions from the NCAR CCM for the April (spring) control case and for the perturbed case. Between any two contours, the mass flow is 101° kg/s. Averages over the period t = 16 to 20 days are shown. (From Covey et al., 1984.) hemisphere. The mean meridional velocity in this branch is about 4 m/s. Note that the descending branch of the meridional cell spreads southward about 10° latitude as well. This intensification of the mean meridional circulation did not occur in the case of a perturbed circulation calculated by Covey et al. for winter because the differential heating along the southern edge of the cloud was not

161 intense enough during winter to counter the normal circulation driven by heating centered farther south. A strong enhancement and widening of the mean meridional circulation also developed in the simulation of Alexandrov and Stenchikov for annual mean conditions. In some additional model studies, S.H. Schneider and S.L. Thompson (private communication, 1984) have found that the onset of this enhanced cross-equatorial circulation is sensitive to the reflectivity of the particulate cloud. For a sufficiently reflective cloud,* this qualitative change in the mean meridional circulation does not occur in the CCM. The altered meridional circulation of Figure 7.11 would spread particulates rapidly upward and equatorward from the southern edge of the cloud. AS the particulates spread, the heating would spread upward and equatorward with them. Since the particulate heating becomes more effective at the lower air densities of higher altitudes, the intensity of the thermally driven meridional circulation would increase as the particles rose. This positive feedback arising from the coupling between transport and heating has been observed in the GCM calculation of MacCraken and Walton (1984) and in two-dimensional model calculations (Ha~erle et al., 1983; M.C. MacCracken, private communication, 19847. In this simulation of a summer case in which an initial smoke cloud was located below 4 km, smoke from the equatorward edge of the cloud rapidly rose to 30 km, where its further evolution was influenced by the top of the model domain. Thus it is likely that this positive feedback mechanism could propel smoke to high levels, where its lifetime could be greatly lengthened. Even without the operation of the feedback mechanism, an enhanced meridional circulation like that produced by the CCM could transport smoke into the southern subtropics in 1 or 2 weeks. The large transport rates (up to 400 km/day} illustrated in Figure 7.11 developed without the feedback arising from transport of particulates, however. Other Large-Scale Wind Systems Large temperature differences between land and sea at low levels would produce effects that might be analogous to those occurring in midwinter at fairly high latitudes: prevalence of anticyclones with low-level outflow over the continents, surface cyclones over the high-latitude oceans, and development of frequent intense coastal storms, especially along eastern coasts of continents. At present, these effects must be regarded as speculative, although additional calculations with GCMs could narrow the uncertainties. Since many of the observed storms occurring in high-latitude coastal waters during winter and spring are quite small, model studies intended to simulate their behavior require higher horizontal spatial resolution than was used in the GCMs of either Covey et al. or Alexandrov and Stenchikov. *Such a reflectivity could be produced by high-altitude dust injections and would itself imply a major long-term climate perturbation.

162 In addressing the question of interhemispheric transport, Covey et al. noted that the CCM results for spring showed strong localized cross-equatorial flows extending as far south as 30°S in the upper troposphere and lower stratosphere. Even in the unperturbed atmosphere, large-scale upper tropospheric troughs occasionally extend from mid-latitudes to the equator and beyond (young and Hitchman, 1982; Vincent, 1982; Huang and Vincent, 1983~. These occurrences are more frequent during winter and over the oceans. Thus, if smoke reaches the upper troposphere, there is a good possibility that bands or streamers would be separated from the main cloud mass and stretched into the southern hemisphere even if the enhanced meridional circulation described above does not operate. The separation of bands of smoke from the southern edge of the main cloud would be associated with the complementary process: injection of streamers of clear air northward. In conjunction with spatial nonuniformity in the precipitation scavenging rate, this would ensure a degree of nonuniformity in optical depth even at long times after smoke injection, especially in the southern portion of the smoke-covered region.* More research is needed to understand transport near the southern edge of the clouds, both in the unperturbed and in the perturbed atmospheres. Ultra-High Clouds The intense heating in the upper portions of the particulate cloud should drive intense convection above the cloud. In the TTAPS model calculations, such convection does appear in the form of convective adjustment above the cloud. The temperature distribution shown in Figure 7.7 contains the effect of this convective adjustment; the temperature is shown to have increased in the region above the cloud as a consequence of upperward heat transport by convective adjustment. Mixing in the connectively active layer would stir fine particles upward, thereby raising the altitude of maximum heating and further raising the altitude of the convective layer. This effect, which is in addition to any systematic tendency for large-scale circulation to raise the particulate cloud, was recognized by the TTAPS group, but not explicitly accounted for in their calculations. Water vapor would also be mixed through the convective layer. At the top of the layer, a deep temperature minimum would develop, particularly as the convective elements are likely to overshoot the level of neutral buoyancy. The water vapor transported upward by the convection would be likely to condense to form a widespread cirriform cloud cover. The mass of material in this cloud would depend primarily on the water vapor concentrations at and above the base of the convective layer (near the temperature maximum). Such a cloud could *In a series of unpublished GCM calculations, J. Mahlman (of Geophysical Fluid Dynamics Laboratory) has recently shown how these processes conspire to maintain nonuniformities in a simulated unperturbed atmosphere.

163 have significant radiative effects. Because ice crystals found in normal cirrus clouds tend to be of moderate size twith radii of several microns to a few tens of microns), and because ice is strongly absorbing in the infrared and reflective in the visible, normal cirrus generally has a larger influence on infrared radiation than on solar radiation. However, even a small increase in albedo due to such clouds would reduce the energy received by the atmosphere, so it is difficult to estimate the net climatic impact without detailed calculations. As an example, suppose that water vapor from the base of the convective region is mixed upward uniformly through the convective layer with a mixing ratio of 100 ppmv, a representative value for air originating near the 200-mbar level. With adiabatic cooling of the rising air, condensation could begin near or slightly below the 50-mbar level. If the cloud extends 1 km above the condensation level and most of the water vapor in the cloud layer condenses, the resulting cloud mass would be about 7 g/m2. The absorption cross section at 10-pm wavelength for spherical ice particles whose radii are a few microns or less is about 0.1 to 0.2 m2/g (Bergstrom, 1973~. Thus, in this example, an absorption optical depth of 1 for 10-pm radiation could develop for such a cloud. More work is needed to assess the significance of such ultra-high clouds. For example, if the absorbing particulate cloud moves upward, as a result of self-induced circulation or mixing, the infrared opacity of such an elevated cirrus layer would be correspondingly smaller. Longer Term Effects on Climate If nuclear war injections of smoke were as large as those of the NRC baseline case, longer term meteorological effects, extending beyond the time at which most of the smoke is removed from the atmosphere, might occur. Such effects could arise from changes in the distribution of snow, sea ice, and vegetation cover, which would cause changes in surface albedo, thermal inertia, and evapotranspiration potential. It is also possible that persistent changes in ocean current systems leading to changes in sea surface temperature distributions would be produced. The upward mixing of water vapor by convection to altitudes above 10 km could also have significant long~term climatic implications. Such possibilities are extremely difficult to evaluate, particularly because shorter term effects themselves are highly uncertain. However, Robock (1984) has recently attempted to assess some of these effects using an EBCM with snow and ice albedo feedback and sea ice thermal inertia and meltwater feedbacks included in the model (Robock, 1983~. Applying this model to the TTAPS scenario, he found depressed surface temperatures persisting but gradually ameliorating over several years in northern, middle, and high latitudes, primarily as a result of an increase in the surface covered by sea ice with a corresponding reduction in thermal inertia of the northern high-latitude oceans. An effect that could be significant but would favor warming of high-latitude surface temperatures is the depression of snow and ice

164 albedo due to the fallout of smoke particles. If as little as 10 to 20 Tg of smoke particles was to fall out over the Arctic during the course of a few months and if the smoke particles were mixed with no more than the normal amount of snowfall, they could have a very significant effect on snow albedo (Warren and Wiscombe, 1984~. The actual importance of this effect is difficult to evaluate, however, since it depends on many detailed processes, such as the exact timing of smoke and snow fallout events, washout of smoke particles due to surface melting on snow or ice, and changes in the morphology of the snow or ice surfaces. Such longer term effects are difficult to investigate, but they should not be ignored. ANALOGS Of necessity the previous discussion relies heavily on model results, supplemented by occasional references to our understanding of how the undisturbed atmosphere behaves. Confidence in these results can be enhanced by examining natural situations where some of the key processes and their effects can be seen. Indeed, bare model results in the absence of such natural analog situations would be quite unconvincing to many observers. In this section several such natural analogs are examined. Arctic Haze Recent research has shown that there is a remarkable amount of aerosol pollution in the central Arctic, especially during spring (Patterson et al., 1982; Rosen and Novakov, 1983~. A major component of this pollution is a fine particle mode (particle mode diameter of about 0.4 um), which in turn is rich in soot carbon. This material has been detected near the surface and in layers at elevations as high as 5 km (Hansen and Rosen, 1984; Radke et al., 1984~. The particles in such elevated layers, following essentially quasi-isentropic trajectories,* must have originated at distant mid-latitude pollution sources, and they must in some cases have been in transit for many days. Thus the properties of these particles provide valuable information on the aging of carbonaceous particulates in the unperturbed atmosphere. Microscopic analysis and analysis of the optical properties of these particles indicate that the soot particles sometimes occur internally mixed in a nonabsorbing material, probably sulfate (A.D. Clarke, private communication, 1984~. The polluted layers also contain nonabsorbing *Heating can probably be neglected to first order in considering the transport of these particles, so that they would tend to move approximately on surfaces of constant specific entropy. Since these slope upward toward the pole, pollutants originating near the surface can reach the middle troposphere in the Arctic.

165 particles unmixed with carbonaceous material so that the mean single scattering albedo of all particles varies around 0.86 (Clarke et al., 19843. This value is considerably higher than that of the postulated nuclear war smoke clouds, though nevertheless the polluted layers are quite strongly absorbing. In relating these aerosols to the smoke that could be produced by burning cities, it is important to keep in mind that the former are probably produced in pollution plumes that are rich in sulfur and not particularly black at the source; the smoke from burning cities is likely to be much blacker initially and throughout its life in the atmosphere. Elemental carbon several days removed from its sources has also been found to be an important component of the fine particle mode in the marine boundary layer over the Atlantic (Andreas, 1983~. Although highly variable, typical soot fractions of the fine particle mass were about 40 percent. Further experimental studies of the fine particle mode in regions remote from pollution sources should provide valuable information on the mechanisms, rates, and consequences of the aging of carbonaceous particles in the undisturbed atmosphere. This information is a necessary prerequisite to understanding the implications of soot aging for the consequences of nuclear war. Plumes from Large Forest Fires There are a number of accounts of observations of forest fire plumes at large distances from their sources (see Chapter 5~. Lyman (1918), for example, documents a case in which smoke from large fires in Minnesota darkened the sky over much of the northeastern United States and southeastern Canada. Shostakovitch (1925) gives a dramatic account of the obscuration persisting for more than a month due to the Siberian forest fires of 1915. Wexler (1950) provides a well-documented account of the plume from a large number of forest fires burning within a 40,000 km2 area of northwest Alberta and northeast British Columbia (although the extent of the area that actually burned is unclear from Wexler's account). Wexler describes events during the period September 24 to 30, 1950. Within 2 days of the beginning of the most intense phase of burning, the plume had reached Washington, D.C. Within 5 days, it had been observed over all of Canada except the far northeast and far west, over almost the entire United States east of the Mississippi River plus Minnesota and the Dakotas, and had stretched across the North Atlantic and had been observed throughout Western Europe from Portugal to Norway (Figure 7.12). At Washington, D.C., the smoke occurred in a layer between the 2.5 and 5 km altitudes bounded above and below by inversions, and was estimated by Wexler to have reduced the total incident solar radiation by as much as 54 percent. Associated with this reduction was a decrease in maximum temperature that Wexler estimated to be an average of 4°C for 4 days. Smith (1950) quotes an estimate by Fritz that the maximum temperature was reduced by as much as 6°C, with no compensating rise in minimum temperature. By the time the plume had reached England, it

166 1 ~q 1 ~ of h In' FIGURE 7.12 The hatched area represents the region over which smoke was · ~ . . . _ ~ · · _ . . ~ ~ observed from the western Canada forest fires of September 1950 (exclusive of observations from Western Europe). The boundary of this area is dotted where it is tentative. The darkened areas in western Canada are the areas in which the fires occurred, and the curves mark calculated trajectories for smoke reaching the vicinity of Washington, D.C., by September 24, two days after the most intense burning episode. (From Smith, 1950.) appears to have risen to an altitude range of 10 to 12 km (Bull, 1951) These incidents illustrate the rapid spread of fire plumes from relatively small areas. They also show that such plumes can have dramatic optical effects and can influence surface temperatures thousands of kilometers from the source. Such forest fire plumes are not necessarily highly absorbing for solar radiation, however. The reduction in solar radiation and the surface temperature decreases observed at Washington were probably due largely to reflection rather than absorption of sunlight by the cloud. As discussed in Chapter 5, urban fires are likely to produce much blacker smoke, and to produce much larger optical depths and reductions in solar radiation at the surface. .

167 Early Plume from the Mount St. Helens Eruption The paroxysmal eruption of Mount St. Helens on May 18, 1980, produced a large plume of ash that spread rapidly across eastern Washington and into Idaho and Montana during the day following the eruption. Rapid daytime temperature decreases were observed beneath the plume. By comparing observed and forecast temperatures under the plume with those in the surroundings, Mass and Robock (1982) argued that the plume produced a drop in the maximum temperature of up to 8°C. However, during the following night, as the plume drifted over Montana, increases in minimum temperature of about the same magnitude were observed. Evidently, the substantial reduction in solar radiation produced by the plume was compensated by a corresponding increase in the downward infrared radiation. The properties of the ash particles in this early volcanic plume were quite different from those of the smoke particles of the nuclear war scenarios. The Mount St. Helens ash particles had high single scattering albedos, and the size distributions had maximum diameters between 1 and 10 Em. The plume is estimated to have contained about 2 Tg of ash particles with diameters greater than 2 Em, but less than 10-2 Tg of particles with diameters less than 2 um (Hobbs et al., 1982), so it is not surprising that the plume was an effective emitter of infrared radiation at this stage of its evolution. These observations illustrate the rapidity with which such plumes can influence surface temperatures, and they serve to focus attention on the role of the ratio of infrared to visible absorptivity of particles in the nuclear war scenarios. Sahara Dust Plumes, the "Harmattan" Sahara dust carried over West Africa and the tropical Atlantic Ocean by northeasterly and easterly winds provides another natural analog for some facets of the nuclear war problem. Outbreaks of dust over the Atlantic can produce extinction optical depths of about 1 over areas of 106 km2 (Carlson and Caverly, 1977; Carlson and Benjamin, 1980). AS much as 8 Tg of dust may be involved in a large outbreak (Carlson, 1979), and strong heating occurs in the dusty layer. Suppression of convection has been noted when Sahara dust in the middle troposphere is transported over the tropical Atlantic. During the dry season in West Africa, the prevailing northeasterly wind, which is often laden with dust, is known as the "harmattan. n Brinkman and McGregor (1983) report harmattan events in Nigeria with optical depths up to 2 and associated reductions in daily mean total solar radiation of 28 percent. They also report temperature decreases of up to 6°C for these events, although this is representative of the depression of the maximum rather than the daily mean temperature. Although these dust particles are probably generally much larger than the stratospheric dust particles and are both larger and more reflective than the smoke particles of the nuclear war scenarios, these observations show that such aerosols do have a rapid effect on surface

168 temper atur es. They also show that such particles, even though less absorbing than smoke, produce elevated heated layers that can act to suppress convection. Martian Global Dust Storms It is now known that the planet Mars is subject to occasional global-scale dust storms in which dust spreads over most of the planet with mean optical depths of order 5. Martian dust is somewhat more absorbing at visible wavelengths than typical terrestrial dusts, so that the absorptivity for these situations is intermediate between values for nuclear war scenarios with dust only and those with both smoke and dust. Consequently, the scale of the associated optical perturbation is within the range of interest. These events produce temperature increases in the upper part of the dusty layer of order 80°C over much of the planet. Temperature decreases at both subtropical and mid-latitude sites have also been observed in connection with these events (Martin and Kieffer, 1979; Pollack et al., 1979; Ryan and Henry, 1979~. The vertical profile of temperature changes associated with these events resembles that of the nuclear war scenarios except that the decrease in surface temperature is less. This is partly because Martian dust is much less absorbing in the visible than smoke, but, probably more important, it is because the "greenhouses effect is at most very weak on Mars, so that the ~antigreenhouse" effect at the surface due to the absorbing cloud is not very pronounced (see page 149~. These dust storms do not occur every Martian year. When they do occur, it is during southern hemisphere summer, Mars perihelion season, when dust generated locally in the summer subtropics is swept upward to great heights in the rising branch of the mean meridional circulation and then is swept rapidly poleward, reaching high latitudes of the opposite hemisphere within a few days (Haberle et al., 1982~. Proper phasing between dust injection and the meridional circulation is an essential feature of this phenomenon; dust injected into the normally subsiding branch of the tropical mean meridional circulation remains close to the latitude of injection. The analogy to the nuclear war scenarios should not be pressed too far. The total amount of material involved in the Martian dust storms is larger (Toon et al., 1977), but the particle sizes are larger so they are less efficient optically; precipitation processes are not active on Mars; and the global dust storms are driven by heating per unit mass of atmosphere that is larger than the largest reasonable values for the nuclear war smoke clouds. Nevertheless, Mars does provide a natural example of the thermal structure of an "antigreenhouse" atmosphere and of rapid meridional spread of per ticulates by an enhanced thermally driven meridional circulation.

169 SU+ARY None of the natural situations described above bears a close resemblance to the atmospheric condition that is likely to prevail following a full-scale nuclear war. Nevertheless, each has elements that tend to support various conclusions drawn from the models. In sum then, the various model results in concert with a limited set of observations of related natural phenomena provide a basis for concluding that a nuclear war scenario like the NRC baseline case could produce large temperature decreases near the surface and temperature increases aloft for a period of weeks to months following the event (cf. the two- and three-dimensional model results summarized in Tables 7.3 and 7.4~. Moreover, rapid spreading of particulates into the tropics and even into the southern hemisphere is a real possibility. These conclusions are contingent upon the assumptions that a substantial fraction of the smoke particles produced by burning cities would survive early scavenging and coagulation, and that subsequent aging and scavenging processes would not remove submicron smoke particles distributed throughout the middle and upper troposphere at a removal rate* greater than about (2 weeks)~l. Because of optical saturation due to the high absorptivity of smoke, the climatic effects are likely to be insensitive to moderate changes in smoke or absorptivity about the baseline values. However, lower values of either of these quantities by a factor of about 4 would lie near the edge of the saturation regime, and climatic effects would decrease rapidly for large reductions. Climatic effects are also sensitive to the removal rate of smoke. If middle and upper tropospheric rates were as large as (1 week)~1 temperature perturbations would be considerably moderated although still significant (see the Fast rainout" used in Figure 7.6~. Improvements in the models are needed, particularly to investigate further the effects of realistic transport and dispersion of smoke and dust in the perturbed atmosphere, the infrared opacity of the smoke, diurnal and seasonal effects, and the possible roles of ground fog and stratus and of ultra-high clouds forming at the top of the convective layer that may be driven by absorption of solar radiation in smoke and dust clouds. Long-term effects arising from possible changes in the properties of the underlying surface also require further study. REFERENCES Ackerman, T.P., and O.B. Toon (1981) Absorption of visible radiation in atmosphere containing mixtures of absorbing and non-absorbing particles. Appl. Opt. 20:3661-3668. Aleksandrov, V.V., and G.L. Stench~kov {1983) On the modeling of the climatic consequences of the nuclear war. In Proceedings on Applied Mathematics. Moscow: Computing Center of the Academy of Sciences USSR. *Removal rate is defined in Table 7.2.

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172 Mass, C., and A. Robock (1982) The short-term influence of the Mount St. Helens volcanic eruption on surface temperature in the northwest United States. Mon. Weather Rev. 110:614-622. Massie, S.T., and D.M. Hunten (1981) Stratospheric eddy diffusion coefficients from tracer data. J. Geophys. Res. 86:9859-9868. Ogren, J.A. (1982) Deposition of particulate elemental carbon from the atmosphere. In Particulate Carton: Atmospheric Life Cycle, edited by G.T. Wolff and R.L. Klimisch. New York: Plenum. Ogren, J.A., and R.J. Charlson (1983) Elemental carbon in the atmosphere: Cycle and lifetime. Tellus 358:241-254. Patterson, E.M., B.T. Marshall, and K.A. Rahn (1982) Radiative properties of the Arctic aerosol. Atmos. Environ. 16:2967-2977. Pollack, J.B., O.B. Toon, and B.N. Khare (1973) Optical properties of some terrestrial rocks and glasses. Icarus 19:372-389. Pollack, J.B., O.B. Toon, C. Sagan, A. Suborners, B. Baldwin, and W. van Camp (1976) Volcanic explosions and climate change: A theoretical assessment. J. Geophys. Res. 81:1071-1083. Pollack, J.B., D.S. Colburn, F.M. Flasar, R. Kahn, C.E. Carlston, and D. Pidek (1979) Properties and effects of dust particles suspended in the Martian atmosphere. J. Geophys. Res. 84:2929-2945. Pollack, J.B., O.B. Toon, T.P. Ackerman, C.P. McKay, and R.P. Turco. (1983) Environmental effects of an impact-generated dust cloud: Implications for the Cretaceous-Tertiary extinctions. Science 219:287-289. Radke, L.F., J. Lyons, D. Hegg, P.V. Hobbs, and I. Bailey (1984) Airborne observations of Arctic aerosols. 1. Characteristics of Arctic haze. Geophys. Res. Lett. 11:393-396. Ramaswamy, V., and J. Kiehl (1984) Sensitivity of the radiative forcing due to large loadings of smoke and dust aerosols. Manuscript, National Center for Atmospheric Research, Boulder, Colo. (Submitted to J. Geophys. Res.) Robock, A. {1983) Ice and snow feedbacks and the latitudinal and seasonal distribution of climate sensitivity. J. Atmos. Scz. 40: 986-997.. Robock, A. (1984) Snow and ice feedbacks for prolonged effects of nuclear winter. Nature 310: 667-670. Rosen, H., and T. Novakov (1983) Combustion-generated carbon particles in the Arctic atmosphere. Nature 306:768-778. Ryan, J.A., and R.M. Henry (1979) Mars atmospheric phenomena during major dust storms as measured at the surface. J. Geophys. Res. 84:2821-2829. Schneider, E.K. (1983) Martian great dust storms: Interpretive axially symmetric models. Icarus 55:302-331. Sellers, W.D. {1973) A new global climate model. J. Appl. Meteorol. 12:241-254. Shostakovitch, V.B. (1925) Forest conflagrations in Siberia. J. Forestry 23:365-371. Smith, C.D., Jr. (19S0) The widespread smoke layer from the Canadian forest fires during late September 1950. Mon. Weather Rev. 78:180-184.

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Most of the earth's population would survive the immediate horrors of a nuclear holocaust, but what long-term climatological changes would affect their ability to secure food and shelter? This sobering book considers the effects of fine dust from ground-level detonations, of smoke from widespread fires, and of chemicals released into the atmosphere. The authors use mathematical models of atmospheric processes and data from natural situations—e.g., volcanic eruptions and arctic haze—to draw their conclusions. This is the most detailed and comprehensive probe of the scientific evidence published to date.

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