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etate were not well correlated with Rn, suggesting marine biogenic sources for these species. A major contributor to the aerosol mass (~80% of the total mass) was unspecified water-soluble organics.

As discussed in Section C.la, sea salt particles in the marine boundary layer have been shown to likely play a major role in backscattering of solar radiation (Murphy et al., 1998), i.e., to the direct effect of aerosol particles. However, they also contribute to the indirect effect involving cloud formation, since they can also act as CCN. Since such particles are a natural component of the marine atmosphere, their contribution will not play a role in climate change, unless their concentration were somehow to be changed by anthropogenic activities, e.g., through changes in wind speed over the oceans, which largely determines the concentration of sea salt particles (Gong et al., 1997a, 1997b). However, the presence of sea salt particles can still have an impact on the effects of anthropogenically derived species. Thus, activation of sea salt particles results in lowering of the peak supersaturation in the cloud. From the Köhler curves (Fig. 14.38), this means that the size of other particles such as non-sea salt sulfate (nss) must be larger in order to activate into cloud droplets. If fewer anthropogenic particles grow into this larger size, the number of cloud drops formed is reduced and the anthropogenic contribution to cloud droplet formation is proportionately smaller (O'Dowd et al., 1997a, 1997b).

Sea salt particles also provide an aqueous medium for the oxidation of S02 to sulfate (e.g., see Sievering et al., 1992; Chameides and Stelson, 1992; and O'Dowd et al., 1997a, 1997b) and hence play a role in both the direct and indirect forcing by sulfate particles. As discussed in Chapter 8.C.3, such aqueous-phase processes, which usually dominate the overall conversion of SOz to sulfate, are pH dependent. This is in part due to the decreasing concentrations of dissolved S(IV) in the aqueous phase as the pH falls and in part due to the dependence of the reaction kinetics on pH. Seawater is basic (pH ~8) so that newly formed sea salt particles are likely basic when initially formed. Under these conditions, oxidation by 03 is important (see Chapter 8.C.3d). This oxidation dominates until the alkalinity of the droplet has been consumed by the acid formed. As the pH of the droplet falls, oxidation by H202 and the gas-phase oxidation by OH become relatively more important. While the pH of atmospheric sea salt particles has not been well established, experimental studies in Bermuda under moderately polluted conditions suggest that it can be in the range 3.5-4.5 (Keene and Savoie, 1998, 1999). Oxidation by HOC1, believed to be an important intermediate in halogen chemistry in the marine boundary layer, has been proposed to be important as well (e.g., see Vogt et al., 1996; Keene et al., 1998; and Chapter 8.C.3).

The significance of this oxidation of S(IV) in sea salt particles is that if it occurs in existing aerosol particles, sulfate formation will not result in new particles and hence potentially new CCN, but rather contribute to the mass of existing particles (e.g., O'Dowd et al., 1997b). A significant fraction of all particulate nss is believed to be generated by this oxidation in existing sea salt particles.

It has also been proposed that the uptake of gases such as HN03 and HCl onto particles may alter their ability to act as CCN (e.g., see Kulmala et al., 1993, 1995, 1998; and Laaksonen et al., 1997). Clearly, these are areas that need much further investigation.

In short, it is becoming clear that although the focus to date has been mainly on sulfate, the effects of other components, including both natural and anthropogenic species, need to be taken into account in both the direct and indirect effects of particles on global climate.

Given the evidence for a relationship between anthropogenic emissions and CCN, the next link to global climate is the assumption that increased CCN lead to increased cloud droplet concentrations (iV). As seen in Eqs. (JJ) and (KK), increased concentrations affect both cloud albedo and its sensitivity to changes in the cloud droplet number. There is a great deal of evidence gathered over decades for a relationship between increased CCN and increased concentration of droplets in clouds. For example, some 30 years ago Warner and Twomey (1967) measured cloud droplet number concentrations and condensation nuclei at 0.5% supersaturation below the base of clouds upwind (over the ocean) and downwind of a region in which sugar cane was burning. The average concentration of CCN was 280 cm~3 over the ocean but 750 cm-3 downwind of the fire; the cloud droplet number concentration similarly increased from 300 to 920 cm"3. The relationship between the average cloud droplet number and that expected from transport of the below-cloud CCN into the cloud was approximately linear in that particular case, as well as in other sets of measurements carried out at other locations under more normal conditions (e.g., Twomey and Warner, 1967).

Similarly, Martin and co-workers (1994) measured aerosol particles in the size range from 0.05 to f.5 /¿m below the base of stratocumulus clouds, along with cloud droplet number concentrations in maritime and in continental air masses. Figure 14.46 shows the relationship between cloud droplet number concentration and the aerosol particle concentration for a set of flights carried out in the vicinity of the British Isles and in the South Atlantic (Martin et al., 1994). There is an almost linear relationship between the two for maritime air masses. Given that the cutoff for particle measurements was 0.05 /¿m, these concentrations may have been underestimated, so that the slope of the line for maritime air masses can be taken as unity. That is, essentially all of the maritime particles at the cloud base could act as CCN under the range of supersaturations in these studies.

However, this relationship did not hold true for continental air masses. The fraction of aerosol particles that lead to cloud droplet formation is clearly less than one, in agreement with the studies of Gillani et al. (1995) discussed earlier. In addition, the relationship is much more scattered, indicating that the chemical c 600 o

o cP 400

0 500 1000 1500

Aerosol particle concentration (number cm"3)

FIGURE 14.46 Average cloud droplet number concentration as a function of subcloud aerosol particle concentration (0.05-1.5 /im) in marine (•) and continental (□) air masses (adapted from Martin et al., 1994).

composition and hence ability to act as CCN are much more variable over the continents.

Number concentrations of ice crystals in cirrus clouds have also been observed to increase with aerosol particle concentrations (with diameters >0.018 /jlm) and, in particular, with the concentration of light-absorbing materials in the ice crystals (Strom and Ohlsson, f998).

Not only do CCN affect the number of cloud droplets formed, but they also affect the size distribution of these droplets. This also affects cloud albedo and its sensitivity to changes in the number concentration (see Eqs. (JJ) and (KK)). Figure 14.47, for example, shows the size distribution for cloud droplets measured in urban and nonurban air around Denver, Colorado (Al-kezweeny et al., 1993). The median volume diameter was 14 ¡xm for the urban air cloud, and this was only ~50% of that of the much larger droplets in the nonurban air cloud. As expected, the cloud number concentration in the urban air cloud was also larger, 226 cm-3 compared to 22 cm~3 for the nonurban cloud. Analogous results have been reported in other studies (e.g., Pueschel et al., 1986). For example, based on satellite data, the radii of cloud drops over the oceans were observed to be 2-3 /jlm larger than over continental regions; in addition, marine cloud droplet radii in the Southern Hemisphere are ~1 pm larger than in the Northern Hemisphere (Han et al., 1994).

Hudson and Li (1995) measured aerosol particle concentrations below clouds, as well as various cloud parameters, in polluted as well as clean air masses in aircraft measurements around the Azores. Table 14.8 summarizes some of these data. Consistent with the studies discussed above, the polluted air mass had increased aerosol particle concentrations, increased CCN, and larger numbers of, but smaller sized, cloud droplets.

Albrecht (1989) suggested that increasing CCN concentrations would lead to decreasing cloud drop size and decreased drizzle production in marine stratocumulus and fair-weather cumulus clouds, leading to an increase in the geographical extent of clouds as well as their lifetime. Modeling studies suggest that this could be a significant effect (Lohmann and Feichter, 1997). Clearly, this too could play a role in global climate change. The studies by Hudson and Li (1995) also reported evidence for this effect in that the number of "drizzle drops" with diameters >50 /xm was smaller (by an order of magnitude) in the cloud in the polluted air mass (Table 14.8). Related to this is the observation by Parungo et al. (1994) that there has been an increase in total oceanic clouds from 1930 to 1981, with the change in the Northern Hemisphere (2.3%) being about double that for the Southern Hemisphere (1.2%).

c 600 o

o cP 400

0 500 1000 1500

Aerosol particle concentration (number cm"3)

FIGURE 14.46 Average cloud droplet number concentration as a function of subcloud aerosol particle concentration (0.05-1.5 /im) in marine (•) and continental (□) air masses (adapted from Martin et al., 1994).

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