Unraveling the Role of Aerosols in
Climate Change Significant progress has been made in understanding how aerosols affect climate change, but big unknowns must still be addressed.
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© 2001 American Chemical Society
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lthough it has been recognized that aerosols of human origin could have a large impact on climate change, there is great uncertainty about how to model their effects. Aerosols—suspensions of small particles in air—affect climate in two important ways. They scatter and absorb radiation (the direct effect), and they change the microphysical structure and possibly the lifetime and amount of clouds (the indirect effect). This article examines what has been learned so far about aerosols and climate, and what still needs to be investigated. The direct and indirect effects of aerosols can be understood by considering their separate effects on solar and thermal radiation in the atmosphere. Solar radiation comes from the sun and has wavelengths ranging from 0 to ~4 µm. Thermal radiation, on the other hand, is emitted by the earth’s surface and atmosphere and has wavelengths ranging from 4 to 20 µm. Because most aerosols are smaller than ~4 µm and aerosol particles scatter radiation most strongly at wavelengths near their size, most radiation scattering by atmospheric aerosols takes place in the solar spectrum. For selected compounds of appropriate molecular structure (e.g., elemental carbon), appreciable absorption of solar radiation can also occur. Aerosols having diameters >2 µm, such as sea salt and dust particles, can also absorb at infrared wavelengths and interact strongly with thermal radiation. The scattering of solar radiation cools the planet, and absorption of solar radiation warms the air directly instead of allowing sunlight to be absorbed by earth’s surface. When sunlight is absorbed by the earth’s surface, the heat is then transferred back to the atmosphere either in the form of a flux of thermal radiation—by sensible heat flux (radiative flow from the surface to the atmosphere through advection, conduction, and convection processes)—or by latent heat flux associated with the evaporation of water (energy flow from the surface to the atmosphere through evaporation and condensation processes). (Readers less familiar with some of the terminology used in this feature article can refer to climatology glossaries at www.globalchange.org/glossall and www.shsu.edu/ ~chemistry/Glossary/glos.html.) The direct effect of aerosols arises from increases in aerosol concentrations, but it is difficult to calculate because aerosols may both absorb and scatter solar radiation. The amount of total scattering relative to the total scattering and absorption by aerosols is measured by the single-scattering albedo (the fraction of radiation reflected by a surface).
JOYCE E. PENNER, DEAN HEGG, AND RICHARD LEAITCH
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Although the exact effect of the aerosol single-scattering albedo depends on the zenith angle (the angle between the local zenith and the line of sight to the sun) and the reflectivity of the underlying atmosphere and surface, under average conditions, aerosols with single-scattering albedo less than ~0.85 tend to lead to a net warming, and aerosols with higher albedoes tend to lead to a net cooling. The single-scattering albedo is mainly determined by the amount of black carbon in the aerosol. Moreover, the specific configuration of black carbon within the aerosol is important. If it is internally mixed, its specific absorption is roughly twice that when it is externally mixed (1). Internal mixing (where black carbon is mixed with other aerosol components within each particle) is thought to be more important. Another issue for calculating the direct effect of aerosols is the amount of water vapor that condenses on the aerosol. At high relative humidities, acid aerosols like sulfate can absorb a significant amount of water. This generally moves them into a size range that causes a very large increase in the amount of scattering, especially at very high relative humidity (>90%). Because of this feature, aerosols of different chemical characteristics can be more or less effective at scattering radiation, even if they have the same dry size distribution. Quantifying this effect requires an understanding of both an aerosol’s chemistry and dry size distribution and the temperature and water vapor distribution in the atmosphere. The indirect effects of aerosols result from changes in cloud properties. Because of their abundance and size, cloud droplets both absorb thermal radiation and scatter solar radiation. The total radiative effect of clouds (called the cloud forcing) is of the order of −20 W m−2 (which is composed of 30 W m−2 thermal and −50 W m−2 solar). (“Radiative forcing” generally refers to perturbations of the climate system that change its radiative balance.) Clearly, compared to the average forcing by anthropogenic greenhouse gases today (~2.5 W m−2), even small changes in clouds could be associated with significant climate forcing. Changes in liquid water clouds at low altitudes tend to mainly change the amount of scattered solar radiation. In contrast, the radiative effect of changes in ice clouds at high altitude is dominated by the change in outgoing thermal radiation. Increases in low-altitude clouds act to mainly cool the planet, whereas increases in high-altitude clouds tend to warm it. Aerosols can affect clouds by acting as cloud condensation nuclei (CCN), the particles on which water vapor condenses to form cloud droplets). At constant liquid water content, increases in aerosols will tend to increase the number of cloud droplets (and decrease their size), a process that is called the “Twomey effect” (2). The effect of increasing aerosol concentrations on cloud droplet number concentrations at constant liquid water content is called the “first indirect effect”. Increases in the number concentration of cloud droplets in low-altitude, liquid water clouds may greatly increase scattering of solar radiation (while having only a small effect on their interaction with thermal radiation as long as 334 A
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the total amount of water in the cloud remains constant), and thereby cool the planet. Although aerosols are also likely to significantly affect the radiative properties of ice clouds, much less is understood about such processes. A second effect of increases in aerosols occurs if changes in droplet concentrations alter the precipitation efficiency of clouds (3). Decreases in precipitation efficiency might occur because smaller droplets may require longer times to develop into precipitating droplets. This may increase cloud lifetime and cloud cover and lead to clouds having higher liquid water content on average. These changes are referred to as the “second indirect effect”. Calculation of the indirect effect of anthropogenic aerosols is highly sensitive to the number of natural aerosols, because at high aerosol concentrations, the indirect effect “saturates”—the addition of more aerosols does not increase the total number of droplets as much as it does when aerosol concentrations are small. Therefore, an accurate calculation of the indirect forcing requires an understanding of both natural and anthropogenic aerosols.
Sources of aerosols Particles in the atmosphere are most often composed of sulfate, nitrate, ammonium, organics, silicates, and other compounds associated with soils and dust; sea salt and compounds associated with the sea salt aerosol; soot or black carbon; and trace metals (the total trace metal mass is small and therefore not described here because of its very small climatic influence). The sources of these aerosol components vary significantly with region and time of year. In addition, aerosol lifetime is relatively short (~5 days), implying that the component concentrations also vary significantly with location and time. The main anthropogenic sources of aerosols are derived from fossil fuel emissions and other industrial activities, biomass burning, and agricultural emissions (see Table 1 and Figure 1 (4)). Fossil fuel emissions are mainly emitted into the Northern Hemisphere, and emissions associated with biomass burning are mainly emitted into the tropics, with a seasonal variation that depends on the timing of the dry season and agricultural activities. Using estimated total source strengths of submicrometer anthropogenic aerosols together with typical conversion efficiencies of aerosol precursors to condensed products, we estimate a total anthropogenic and natural source strength of submicrometer aerosols of 434 Tg yr −1 and 385 Tg yr −1, respectively (1 teragram (Tg) is 1012 g). These estimates are subject to considerable uncertainty (perhaps a range of 244–667 Tg yr −1 and 178–681 Tg yr −1, respectively), but demonstrate that anthropogenic sources are of considerable importance to the global aerosol budget. Specific regions will have far larger anthropogenic influences (Figure 1). In addition to the anthropogenic sources estimated for their effect on today’s climate, climate change itself can lead to additional emissions of some aerosol types that are normally considered natural. For example, climate change is associated with increased temperature, winds, and arid areas, leading to enhanced
TA B L E 1
Aerosol sources and source strengths Main sources and source strengths (4) of anthropogenic and natural aerosols (r < 1 µm). The source strengths are representative of the mid-1980s except for the changes in natural aerosols (for which the values are estimates for 2100). Uncertainty ranges are given in parentheses. Anthropogenic Aerosol type
Source strength (Tg yr −1)
Sulfates (as HSO4−)
104 (59–182)
Natural Source strength (Tg yr −1)
Main source activities
Fossil fuels and smelting
49 (24–101)
Main sources
Dimethylsulfide and H2S from oceans, land biota, and soils
3.2 (1.5−9)
Enhanced emissions of dimethylsulfide associated with stronger winds and higher temperature from climate change
18 (8–41)
Volcanic SO2
20 (10–30)
Fossil fuels, outdoor cooking
14 (8–40)
6 (3–17)
Enhanced emissions of terpenes from higher temperature due to climate change
Photochemical conversion of terpenes to condensable products and primary biogenics
Black carbon
7 (4–11)
Fossil fuels, outdoor cooking
Smoke
70 (50–90)
Nitrates (as NO3−)
Organic carbon
0
No sources
Biomass burning; smoke is largely composed of organic and black carbon and is therefore often included in those categories.
3 (2−4)
Natural fires
14 (10–20)
NOx from biomass burning, fossil fuel, and aircraft; agricultural soil NOx
4 (2–8)
Lightning, natural soil, and stratospheric NOx
Ammonium (as NH4+)
19 (11–34)
Enhanced soil emissions from application of nitrogen fertilizer; domestic animals; human emissions; biomass burning; fossil fuel and industry
12 (6–26)
Natural soils, wild animals, and oceans
Sea salt
67 (23–126)
Enhanced wind injections associated with climate change in 2100
88 (30–165)
Formation of jets and bubbles from wind
Dust r < 1 µm
200 (100–300)
Agriculturally disturbed lands and increased desertification
200 (100–300)
Wind-blown dust in deserts and other arid, susceptible areas
20 (10–30)
Dust associated with enhanced winds and arid areas due to climate change in 2100
emissions of sea salt, terpenes, ocean dimethylsulfide, and dust. Table 1 indicates estimates of the increases in these sources associated with the climate in 2100.
Aerosol direct effects A simple box model (see sidebar on page 338A (5–7)) of the overall change in planetary albedo can be used to assess how the uncertainty associated with each of the factors determining aerosol direct forcing contributes to the overall uncertainty in forcing. The ap-
proach (and some of the parameter values) follows one previously described (6), but it is updated here. On the basis of the model analysis, the resulting global mean aerosol forcing along with the overall uncertainty by aerosols associated with fossil fuels and other industrial activities is −0.6 ± 0.42 W m−2 (Table 2). Thus, an estimate of the 2/3 uncertainty range of the global mean forcing is −0.2 to −1.0 W m−2. This range encompasses values for recent evaluations of the forcing for the individual components of industrial aerosols (8–15). The main uncertainties AUGUST 1, 2001 / ENVIRONMENTAL SCIENCE & TECHNOLOGY
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TA B L E 2
Direct forcing variables: Fossil fuels and other industrial processes Several factors contribute to direct forcing by aerosols associated with fossil fuels and other industrial processes and its uncertainty. Note that optical parameters are for a wavelength of 550 nm and dry aerosols. Central value
Quantity
Total emission of anthropogenic organic carbon (OC) from fossil fuel burninga (Tg yr −1) 20 Atmospheric burden of OC from fossil fuelsb (Tg) 0.48 Total emission of anthropogenic black carbon (BC) from fossil fuel burning (Tg yr −1) 7 Atmospheric burden of BC from fossil fuel burningb (Tg) 0.133 Total emission of anthropogenic sulfate from fossil fuel burning (Tg yr −1) 69 Atmospheric burden of sulfate from fossil fuel burning (Tg sulfur) 0.525 Fraction of light scattered into upward hemispherec, 僓 0.23 2 −1 Aerosol mass scattering efficiencyc (m g ), αs 3.5 0.92 Aerosol single-scattering albedo (dry)c, ω0 Atmospheric transmittance above aerosol layerd, Ta 0.87 Fractional increase in aerosol scattering efficiency due to hygroscopic growth at 80% RHe 2.0 Fraction of earth not covered by cloudsd 0.39 Mean surface albedod 0.15
2/3 Uncertainty range
10–30 0.33–0.70 4.67–10.5 0.11–0.16 57.5–82.8 0.35–0.79 0.17–0.29 2.3–4.7 0.85–0.97 0.72–1.00 1.7–2.3 0.35–0.43 0.08–0.22
aThe central value estimated here was taken from Table 1, whereas the value used in the IPCC model comparison was 29. bThe burden was that estimated from the IPCC model comparison for total anthropogenic carbon, taking the fraction associated
with fossil fuels from the fraction of emissions associated with fossil fuels. central estimate and uncertainty range were calculated assuming a size distribution for polluted continental aerosols (i.e., a log normal size distribution with geometric mean radius of 0.05 µm and geometric standard deviation of 1.9). dCentral estimate and uncertainty range adapted from values used in reference (5). eRH is relative humidity. cThe
(see sidebar on page 338A) are those for the up scatter fraction by a particle (the fraction of light scattered into the upward hemisphere), the burden or total aerosol amount in the atmosphere (which includes the propagated uncertainties in emissions), and the mass scattering efficiency (the fraction of incident light scattered by a particle per unit mass), in that order. One of the issues of uncertainty that is not included in this estimate is the mode of mixing of the black carbon associated with the aerosol. If the mode of mixing of black carbon were assumed to be external, instead of internal as used here, then the uncertainty range would be similar, but the central value for fossil fuel aerosols might be −0.7 W m−2 instead of −0.6 W m−2. Using similar methods (including the effect of black carbon on an internal mixture) leads to an estimate for the forcing by biomass aerosols of −0.3 W m−2 ± 0.24 W m−2 (Table 3); thus, an estimate of the 2/3 uncertainty range of global mean forcing for biomass aerosols is −0.1 to −0.5 W m−2 (10, 16 ). The main uncertainties (see sidebar on page 338A) are the singlescattering albedo, the fraction of radiation scattered into the upward hemisphere, and the combined uncertainty associated with emissions and the burden estimate. The actual uncertainty associated with estimating the direct effects of aerosols may be somewhat larger than quoted here. This is because the uncertainties associated with the vertical distribution of the aerosol and with potential correlations between clouds and aerosol abundances were not evaluated. Furthermore, other uncertainties are associated with the radiative transfer treatment, which might be of the order of 20% (17). A further caveat is that the 336 A
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uncertainty estimates (see Table 2) depend on the assumption that the data chosen for the size distribution are representative throughout those areas where these aerosol types contribute to forcing. Because we used observations from continental polluted regions for fossil fuel aerosols and from regional measurements for biomass aerosols, they may not encompass the full set of size distributions that may occur in other regions. One final direct effect deserves some mention. This is the “semidirect” effect due to the change in cloudiness associated with heating by absorbing aerosols. If cloud cover decreases because of heating by black carbon, the net effect would be an additional warming associated with a decrease of reflected radiation by clouds. Hansen and co-workers (18) argue that inclusion of this forcing would cause surface temperature changes associated with the total aerosol forcing to be very small, if the aerosol single-scattering albedo is in the 0.9–0.95 range. Ackerman and co-workers showed that this effect could be important locally (19). But Lohmann and Feichter found that their calculated indirect effect would dominate changes from the semidirect effect (20). Because the indirect effect is so uncertain, further research is needed to clarify the role of semidirect forcing.
Aerosol indirect effects Estimating climate forcing by aerosol indirect effects is far more complex than calculating direct effects. Calculation of the indirect effect of aerosols on liquid water clouds requires an understanding of CCN concentrations in the atmosphere and their change as a result of anthropogenic emissions. It also depends on the relationship between aerosol CCN concentration
FIGURE 1
Aerosol source strengths
90.0
90.0
60.0
60.0
30.0
30.0
Latitude
Latitude
Colors in the bars (bottom) and in the maps represent the annual average source strength (kg km–2 hr–1) for each of the major aerosol types and total aerosol optical depth: (a) column average H2SO4 production rate from anthropogenic sources; (b) column average H2SO4 production rate from natural sources (dimethylsulfide and SO2 from volcanoes); (c) anthropogenic sources of organic matter; (d) natural sources of organic matter; (e) anthropogenic sources of black carbon; (f) dust sources for dust with diameters
