Global Tropospheric Chemistry and Climate Change

Global Tropospheric Chemistry and Climate Change

C H A P T E R 14 Global Tropospheric Chemistry and Climate Change Over the past several decades, there has been increasing recognition in a number of...

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14 Global Tropospheric Chemistry and Climate Change Over the past several decades, there has been increasing recognition in a number of areas of the environmental impacts, both realized and potential, of human activities not only on local and regional scales but also globally. This is particularly true of changes to the composition and chemistry of the atmosphere caused by such anthropogenic activities. One example, for which there is irrefutable evidence, is stratospheric ozone depletion by chlorofluorocarbons, discussed in detail in Chapters 12 and 13. Another area of global-scale dimensions that is commanding increased attention is the potential impact of atmospheric trace gases and aerosol particles on climate, the subject of this chapter. Climate is the longterm statistical characterization of parameters describing what we commonly term ‘‘weather,’’ such as surface temperature. For example, the mean surface temperature with its associated variability over some time period, typically taken as 30 years, is one measure of climate. Thus, climate is distinguished from short-term, e.g., day-to-day, variations, which are typically referred to as ‘‘weather.’’ The recognition that atmospheric gases play a central role in determining the earth’s climate goes back more than a century to Joseph Fourier, who proposed in 1827 that heat is trapped by the atmosphere. In 1861, John Tyndall showed that O 2 , N2 , and H 2 do not absorb infrared radiation but that CO 2 and N2 O Žas well as a number of organic compounds. do. Subsequently, Arrhenius Ž1896. considered the role of changes in atmospheric CO 2 on the earth’s temperature due to the absorption of infrared radiation. He estimated, for example, that the temperature in the Arctic would rise by 8᎐9⬚C from an increase in atmospheric CO 2 by a factor of 2.5᎐3 over that present in 1896. Beginning in the late 1800s, Chamberlin explored the relationship between climate and atmospheric composition. A brief history of this area is given by Weart Ž1997. and by Fleming Ž1992..

As we shall see, the interrelationships between atmospheric composition, chemistry, and climate are very complex. For example, as discussed in more detail herein, it is clear that CO 2 emissions, primarily from fossil fuel combustion, have increased dramatically over the past century, leading to substantial increases in its atmospheric concentrations. The concentrations of a number of other greenhouse gases have been increasing as well ŽRamanathan et al., 1985.. In the simplest approach, these increases are expected to lead to a significant increase in the surface temperature, and indeed, there is general agreement that an increase of about 0.3᎐0.6⬚C over the past century has occurred ŽIPCC, 1996.. Thus, there is a sound scientific basis for anticipating that chemical changes in the atmosphere will impact climate. However, the interplay between all of the contributing factors and hence the ultimate quantitati¨ e impacts are, at present, not well understood and the subject of intense research activity. For example, the increase in surface temperatures over the past century has not been continuous, occurring primarily from approximately 1910 to 1940 and from 1975 to the present, with recent years being some of the warmest since extensive record keeping began about 1860. Ice core studies have documented changes over the past approximately 100,000 years in the earth’s climate, prior to extensive fossil fuel use, with some of the changes occurring quite rapidly Žover time scales of a decade or less.. Accompanying these have been changes in the concentrations of a number of atmospheric trace gases such as CH 4 , indicative of complex feedbacks occurring between the atmosphere, land masses, and oceans. Such natural variability complicates assessment of the anthropogenic influences on climate, particularly when the time scales for the effects of emissions due to human activity to be manifested can be a century or more for some gases.



Given the breadth and complexity of the scientific issues involved in global climate, we shall focus in this chapter primarily on the current state of understanding of the role of atmospheric composition and chemistry in determining the radiation balance of the atmosphere. However, some of the variables affecting global climate and potential feedback mechanisms with associated implications for climate change as well as the global climate record are briefly discussed as well; the reader should consult the references cited in those sections for further details. An excellent summary of research in this expanding and dynamic area through 1995 is found in Climate Change 1995: The Science of Climate Change by the Intergovernmental Panel on Climate Change ŽIPCC, 1996.. Radiative transfer in the atmosphere is not treated here in detail. The reader should consult books by Liou Ž1980., Goody and Yung Ž1989., and Lenoble Ž1993. for excellent detailed treatments of this subject.

A. RADIATION BALANCE OF THE ATMOSPHERE: THE GREENHOUSE EFFECT 1. Global Absorption and Emission of Radiation In Chapter 3 we examined the interaction of incoming solar radiation in the UV and visible regions of the spectrum with atmospheric gases, which drives atmospheric photochemistry. This incoming solar radiation


also determines the temperature of the earth’s surface through its absorption and reradiation. Although, as chemists, we tend to think of the absorption and emission of radiation in molecular terms, the greenhouse effect is best thought of in terms of the energy balance of the earth᎐atmosphere system taken as a whole Že.g., see Ramanathan et al., 1987; Ramanathan, 1976, 1988a,b; and Wang et al., 1995.. Figure 14.1 shows the solar flux outside the atmosphere, which is approximated by blackbody emission at 6000 K, and at sea level, respectively. Absorption of incoming solar radiation by O 3 , O 2 , H 2 O, and CO 2 as the light passes through the atmosphere to the earth’s surface is evident. Recall in Chapter 3 that the average total incoming light intensity per unit area normal to the direction of propagation outside the earth’s atmosphere, i.e., the solar constant, is 1368 W my2 . As seen in Fig. 14.2a, this is the energy density that would strike a planar disk of area ␲ r 2 Žwhere r is the radius of the earth. centered along the earth’s axis. However, this incoming solar energy is spread over the entire 4␲ r 2 surface area of the earth. The effective incoming solar radiation per unit area of the earth’s surface is therefore 1368r4 s 342 W my2 . wIt should be noted that, as discussed in Chapter 4.A, collision complexes of O 2 with a second O 2 molecule or with N2 also may contribute an additional small amount to the absorption of incoming solar radiation; for example, O 2 ⭈ O 2 and O 2 ⭈ N2 may contribute an additional 0.57᎐3 W my2 to this total ŽPfeilsticker et al., 1997; Solomon et al., 1998; Mlawer et al., 1998..x

FIGURE 14.1 Solar flux outside the atmosphere and at sea level, respectively. The emission of a blackbody at 6000 K is also shown for comparison. The species responsible for light absorption in the various regions ŽO 3 , H 2 O, etc.. are also shown Žadapted from Howard et al., 1960..



FIGURE 14.2 Global average mean radiation and energy balance per unit area of earth’s

surface. The numbers in parentheses are the energy in units of W my2 typically involved in each path Žadapted with permission from IPCC Ž1996., with numbers from Kiehl and Trenberth Ž1997...

Figure 14.2b summarizes typical fates of this radiation. Of the incoming solar radiation, about 31% is reflected back to space either at the surface Ž30 W my2 . or by the atmosphere itself Ž77 W my2 .. The remaining 235 W my2 is absorbed, about 168 W my2 by the earth’s surface and 67 W my2 by O 3 , CO 2 , H 2 O, and O 2 Žsee Fig. 14.1. and by particles and clouds in the atmosphere. The net absorption of 235 W my2 by the earth’s surface and atmosphere leads to heating and hence to

the thermal emission of radiation. The Stefan᎐Boltzmann law can be applied to the combination of surface and atmosphere as a system ŽRamanathan et al., 1987. approximated by a blackbody. Recall that this law says that the energy radiated by a blackbody at temperature T per unit time is given by E s ␴ T 4 , where ␴ is the Stefan᎐Boltzmann constant, equal to 5.67 = 10y8 W my2 Ky4 . If this absorbed solar energy is radiated in accordance with the Stefan᎐Boltzmann law, the effective temperature, Te , of the surface᎐atmosphere system


can be estimated using E s 235 W my2 s Ž5.67 = 10y8 W my2 Ky4 .Te4 , giving Te s 254 K, or y19⬚C. This assumes no interaction with the atmosphere, which, we shall see, is certainly not the case. Figure 14.3 is a schematic illustration of the wavelength distribution of the direct, incoming solar radiation and the outgoing, lower-energy, terrestrial radiation emitted by the earth’s surface for a temperature of ;254 K. The global climate issues discussed in this chapter focus on the interactions of both the longer wavelength, terrestrial radiation and the shortwave, solar radiation with atmospheric gases and particles. Clearly, 254 K is much colder than the typical temperatures around 288 K Ž15⬚C. found at the earth’s surface. This difference between the calculated effective temperature and the true surface temperature is dramatically illustrated in Fig. 14.4, which shows the spectra of infrared radiation from earth measured from the Nimbus 4 satellite in three different locations, North Africa, Greenland, and Antarctica ŽHanel et al., 1972.. Also shown by the dotted lines are the calculated emissions from blackbodies at various temperatures. Over North Africa ŽFig. 14.3a., in the window between 850 and 950 cmy1 , where CO 2 , O 3 , H 2 O, and other gases are not absorbing significantly, the temperature corresponds to blackbody emission at 320 K due to the infrared emissions from hot soil and vegetation. The negative peaks that appear to be absorption bands superimposed on the continuous emission curve from the earth’s surface are actually due to a combination of two processes involving the atmospheric greenhouse gases: Ž1. absorption of outgoing terrestrial infrared radiation by the gases, causing vibration᎐ rotation transitions Žand in the case of H 2 O, pure rotational transitions ., and Ž2. emission of infrared radiation by the greenhouse gases due to excited states populated by collisions. The population of these excited states is determined by the Boltzmann distribution Žsee Section A.2b. and hence the emission intensity is char-

FIGURE 14.3 Schematic of wavelength dependence of energy emitted by the sun and hence entering the earth’s atmosphere and the energy emitted by the earth’s surface at a temperature of ;254 K. The absorptions of various atmospheric constituents have been omitted for clarity.


FIGURE 14.4 Infrared emission from earth measured from the Nimbus 4 satellite Ža. over the Niger Valley, North Africa Ž14.8⬚N, 4.7⬚W. at 12:00 GMT; Žb. over Greenland Ž72.9⬚LN, 41.1⬚W. at 12:18 GMT, and Žc. over Antarctica Ž74.6⬚S, 44.4⬚E. at 11:32 GMT. Emissions from blackbodies at various temperatures are shown by the dotted lines for comparison Žadapted from Hanel et al., 1972..

acteristic of the particular temperature of the atmosphere where the emitting molecule is located. Contributions from CO 2 Ž600᎐750 cmy1 . and O 3 Ž1000᎐1070 cmy1 . as well as from H 2 O wrotation bands below ;625 cmy1 Že.g., see Clough et al., 1992. and vibration᎐rotation bands in the region from ;1200 to 2000 cmy1 x are evident. There are also smaller contributions from other greenhouse gases such as CH 4 and N2 O Žvide infra.. Figures 14.4b and 14.4c show that the surface temperatures of Greenland and Antarctica are much colder, ;240 and 200 K, respectively, at the times these spectra were recorded. The emissions from other greenhouse gases such as CFC-11 and CFC-12 have also been observed from the earth’s surface in downward radiation Že.g., see Walden et al., 1998.. The fact that both absorption and reemission contribute to the spectral features assigned to CO 2 and O 3 is illustrated by the spectra in Fig. 14.4 as well. Over North Africa ŽFig. 14.4a., the air temperature is less than that at the surface, so that the emission intensity around 1040 cmy1 due to atmospheric O 3 is less than its absorption, leading to a distinct ‘‘negative’’ band superimposed on the continuous emission curve. The



1040-cmy1 ozone band is very weak in the Greenland spectrum ŽFig. 14.4b. not because O 3 is absent, but because the average temperature at which the atmospheric ozone is emitting at this location is about the same as the surface temperature, so that absorption and emission balance out. In Antarctica ŽFig. 14.4c., the atmosphere is warmer than the surface so that infrared emissions due to CO 2 and O 3 more than counterbalance the absorption of terrestrial radiation and their bands actually appear as ‘‘positive’’ peaks on top of the colder surface emissions. While there are obviously extreme variations in surface temperature, the measured emission profile in the ‘‘window’’ between the strong gas absorptions corresponds to a temperature of 288 K as an average over the earth’s surface. From the Stefan᎐Boltzmann relationship, the measured average temperature of 288 K corresponds to an energy of emission of E s ␴ T 4 s 390 W my2 . This temperature and the corresponding energy are clearly much greater than the effective temperature of 254 K calculated earlier assuming the earth is a blackbody emitting the absorbed solar radiation of 235 W my2 with no interactions with the atmosphere above it. Figure 14.2c schematically summarizes the transfer of thermal radiation and heat in the troposphere. Since there is a net absorption of 235 W my2 of incoming solar radiation ŽFig. 14.2b., a net 235 W my2 in outgoing radiation is needed to balance this. Of the 390 W my2 emitted as thermal infrared radiation by the earth Žcorresponding to the satellite-derived temperature of 288 K., approximately 40 W my2 is radiated directly to space in the atmospheric ‘‘window’’ region from 7 to 13 ␮ m where absorptions by CO 2 , H 2 O, and O 3 are relatively weak. It is this radiation that is detected as the ‘‘background’’ in Fig. 14.4 upon which the greenhouse gas bands are superimposed. The remaining 350 W my2 is absorbed by the greenhouse gases and clouds. Water vapor is by far the most important greenhouse gas Že.g., Wang et al., 1976.. For example, Kiehl and Trenberth Ž1997. calculate that in a standard atmosphere containing 353 ppm CO 2 , 1.72 ppm CH 4 , and 0.31 ppm N2 O as well as ozone and water vapor, water vapor contributes ;60% of the total radiative forcing Ždefined later.. CO 2 is the next larger contributor, at ;26%, followed by O 3 , at ;8%. Water vapor in the stratosphere, although present in small concentrations Žsee Chapter 12., is particularly important ŽWang et al., 1976.. While we shall focus on the global view in this chapter, it is important to recognize that the processes shown in Fig. 14.2 are not homogeneous on a global scale. Thus, Fig. 14.5 shows the absorbed short-wavelength energy and the emitted long-wavelength energy

FIGURE 14.5 Annual mean radiation measured by satellite in 1988 at the top of the atmosphere as a function of latitude for incoming absorbed short-wavelength radiation and emitted outgoing long-wavelength radiation Žadapted from Trenberth and Solomon, 1994..

at the top of the atmosphere, derived from satellite measurements in 1988, as a function of latitude ŽTrenberth and Solomon, 1994.. Around the equator, a great deal more incoming solar energy is absorbed than is emitted at the longer wavelengths, whereas the opposite is true at high latitudes. As a result, low latitudes are warmed and high latitudes cooled, causing heat transport from the equator toward the poles by the atmosphere and oceans.

2. Radiative Transfer Processes in the Atmosphere a. Macroscopic View The spectral distribution of the radiation emitted by the earth’s surface is determined by its temperature; i.e., the emission is that of a blackbody at an average temperature of 288 K. Thermal emission at the earth’s surface produces a net upward flux of energy. Let us denote upward energy fluxes by Fq and downward fluxes by Fy. The transfer of radiant energy in the troposphere can be thought of as occurring between vertical layers in the atmosphere as shown in Fig. 14.6a. The total energy flux Ž F net . crossing a plane at a given altitude z is the difference between the upward and downward fluxes; i.e., F net s Fqy Fy. As shown in Fig. 14.6b, the flux of light energy through a volume of air is determined by a combination of transmission, absorption, scattering both in and out of the ‘‘beam,’’ and thermal emission from the gas molecules Žshown on a molecular level in Fig. 14.6c and described in more detail later..



FIGURE 14.6 Schematic diagram of radiative transfer in the atmosphere: Ža. between vertical layers in the atmosphere; Žb. within a volume element; Žc. on a molecular level. The filled circles on the energy levels portray qualitatively the much larger populations in the ground state.

In a particular layer of air in the atmosphere, the net change in energy of the layer is a result of the total fluxes crossing its boundaries from above and below. For example, in Fig. 14.6a for the layer between altitudes z1 and z 2 , if the net flux into the layer F1net is larger in magnitude than the net flux out, F2net , the difference Ž F1net y F2net . must go into heating the layer. There are a variety of radiative transfer models for the atmosphere that incorporate all of the emission, absorption, and scattering processes as a function of altitude. These are used to predict the total radiance as

a function of wavelength as well as the effect of changes in atmospheric gases, aerosol particles, and clouds on it. For details, see the books by Liou Ž1980., Goody and Yung Ž1989., and Lenoble Ž1993. as well as the article by Clough et al. Ž1992. for a typical line-by-line radiative transer model. b. Molecular-Le© el View On a molecular level, absorption of terrestrial infrared radiation of the appropriate wavelength corresponding to the energy-level splittings in the molecule



causes vibrational᎐rotational excitation or, in the case of H 2 O, pure rotational transitions as well. Vibration᎐rotation transitions occur if there is an oscillating dipole moment in the molecule, a requirement not met by homonuclear diatomics such as N2 and O 2 . Recall from Chapter 3.A.1 that the vibrational energy spacing is typically sufficiently large that at room temperature Ž;298 K., most molecules are in the lowest vibrational energy level, and hence absorption of infrared radiation occurs from this ¨ ⬘ s 0 state. The selection rules dictate that ⌬¨ s 1 transitions are by far the most likely, giving vibrationally excited molecules in ¨ ⬙ s 1 upon absorption of infrared radiation emitted by the earth’s surface. ŽOf course, associated rotational transitions occur simultaneously, and overtone and combination bands can also be important.. Take the simplified case of transitions between the ground state Ženergy E0 . and one higher energy level Ž E1 . shown in Fig. 14.6c. Most molecules reside in the ground vibrational state at atmospheric temperatures and can therefore absorb energy corresponding to E1 y E0 . Once a molecule has made the transition to the upper state, it can undergo spontaneous emission of spont . light Žshown as h␯out , induced emission of light ind Žshown as h␯out ., or deactivation back to E0 through collisions with other molecules. Induced emission is not important in the atmosphere due to the low light levels. In the troposphere, the total gas densities are sufficiently high that the collision frequency is large. In addition, the radiative lifetime for vibration᎐rotation transitions is typically quite long, of the order of 1᎐100 ms for most molecules ŽLambert, 1977.. As a result, collisional deactivation of the excited molecules is their major fate Žsee Problem 1.. Another consequence of the high gas concentrations in the troposphere and most of the stratosphere is that it is collisions, rather than radiative processes, that control the population of molecules in various vibrational and rotational energy levels. As a result, excited molecules in E1 are formed primarily by collisions Žshown as ‘‘collisional activation’’ in Fig. 14.6c., not by absorption of radiation. Under these conditions, in the simplest Žhypothetical. case of a molecule with ground-state energy E0 of degeneracy g 0 , and one in excited-state energy E1 of degeneracy g 1 , the ratio of the number of molecules in E1 to that in the ground Ž E0 . state is given by the Boltzmann distribution: N1 N0


g1 g0

eyŽ E1yE 0 .r kT .

Ž A.

Thus, gas collisions lead to a small equilibrium population of excited states. A small fraction of molecules in the excited states emit radiation rather than being

collisionally deactivated. Based on the Boltzmann distribution, the population of the emitting states and hence the intensity of such radiation would be expected to decrease with decreasing temperature. This situation with thermal equilibrium, where the population of the excited states and hence emission intensity is determined by collisions, is known as ‘‘local thermodynamic equilibrium’’ ŽLTE. and holds in the atmosphere up to altitudes of ;50᎐60 km ŽLenoble, 1993.. Above this altitude, non-LTE models must be used Že.g., see Lopez-Puertas et al., 1998a, 1998b.. ´ It is this emission from the Boltzmann population of excited states that is the thermal emission shown in Fig. 14.6b as well as the upward and downward emission shown in Fig. 14.2c. It is also responsible for the positive bands to to CO 2 and O 3 observed in Fig. 14.4c. For a detailed discussion of emission Žthe source function., see Goody and Yung Ž1989., Liou Ž1980, 1992., and Lenoble Ž1993.. c. From Molecules to the Global Atmosphere Overall, then, there is absorption of infrared terrestrial radiation by the greenhouse gases, collisional deactivation to convert this energy to heat, and emission of infrared radiation but at the lower temperatures characteristic of higher altitudes. As a result, the energy input into the troposphere is increased. This is clearly going to be a function of the concentrations of absorbing gases, their infrared absorption cross sections, the flux of terrestrial radiation, and the total gas pressure, which determines the rate of collisional deactivation. However, simultaneously there is emission of infrared radiation from the Boltzmann distribution of molecules in excited states, which leads to a negative energy component. This emission process depends not only on the concentration of the gas but very sensitively on the temperature since this determines the population of the excited states that emit ŽEq. ŽA... As we shall see in some specific cases below, it is the balance between these two at any given altitude that determines the changes in fluxes and the ultimate impact of a change in a greenhouse gas concentration. Because of efficient trapping of specific wavelengths of infrared radiation at lower altitudes by gases such as H 2 O, CO 2 , and O 3 , radiation emitted to space by such infrared-active species generally originates from molecules at higher altitudes, where the temperature is lower. Because of the Boltzmann population temperature dependence, the relative proportion of excited states that are the source of the emission is lower. This leads to smaller total energy emission out to space compared to what would be the case for a higher temperature.


Recall that 235 W my2 must be emitted to space to balance the net energy absorbed from incoming solar radiation ŽFig. 14.2.. While a small part of this Ž40 W my2 . comes from direct emissions from the earth’s surface in the atmospheric window where strong absorptions do not occur, the larger portion Ž195 W my2 . comes from the lower temperature emissions from the greenhouse gases and the tops of clouds Žsee Fig. 14.2c.. Since by the Stefan᎐Boltzmann law, 235 W my2 corresponds to a temperature of 254 K, the emission of infrared radiation to space can then be thought of as occurring from an altitude where the temperature is 254 K, which is approximately 5.5 km above the earth’s surface. In effect, then, infrared emission out to space by the greenhouse gases and clouds occurs at lower temperatures than the corresponding absorptions. The 235 W my2 of incoming solar radiation that is absorbed at the surface and in the atmosphere ŽFig. 14.2b. is ultimately balanced by the outgoing radiation from the upward emission Žat lower temperatures. of approximately 195 W my2 from atmospheric constituents, including the greenhouse gases and clouds, and of 40 W my2 from the surface that occurs in the atmospheric window between the strong absorptions due to CO 2 , H 2 O, and O 3 . The difference of 155 W my2 between the 390 W y2 m emitted by the earth’s surface ŽFig. 14.2c. and the 235 W my2 escaping from the atmosphere represents the amount of ‘‘trapped’’ radiation, the ‘‘greenhouse effect.’’ As seen in Fig. 14.2c, in addition to the 350 W my2 absorbed in the atmosphere by the greenhouse gases and clouds, energy is also deposited in the atmosphere by convective, vertical mixing of surface heat through thermals Ž24 W my2 ., through the release of the latent heat of evaporation of water when it condenses into liquid water Ž78 W my2 ., and by direct absorption of light ŽFig. 14.2b, 67 W my2 .. This total of 519 W my2 Ž350 q 24 q 78 q 67 s 519 W my2 . is balanced by emission of infrared radiation by gases and the tops of clouds upward into space Ž195 W my2 . as well as downward Ž324 W my2 ., where it is absorbed by the surface and heats it. In summary, in a hypothetical world unperturbed by anthropogenic emissions, the presence of H 2 O, CO 2 , and, to a lesser extent, O 3 , CH 4 , and N2 O in the atmosphere leads to a natural greenhouse effect that results in an average surface temperature of about 288 K, rather than 254 K, which is expected in the absence of these gases. It is important to emphasize that because the greenhouse effect originates in radiative transfer processes in the earth᎐atmosphere system, the net effect of a


FIGURE 14.7 Model-calculated atmospheric temperature changes as a function of altitude due to an increase in CO 2 from 315 ppm in 1960 to 370 ppm projected for 2000 Žno feedbacks taken into account. Žadapted from Rind and Lacis, 1993..

greenhouse gas such as CO 2 on the temperature of the atmosphere depends on the altitude and temperature. For example, once CO 2 reaches the stratosphere, its density is too small to trap radiation to a significant extent. In addition, the temperature is increasing with altitude in the stratosphere. Thus, from the Boltzmann relationship ŽEq. ŽA.., the relative concentrations of the excited emitting states are increasing, resulting in a greater net emission of energy to space. The result is that although CO 2 in the troposphere leads to warming, in the stratosphere it leads to cooling Že.g., see Roble and Dickinson, 1989; Cicerone, 1990; Rind et al., 1990; and Rind and Lacis, 1993.. Figure 14.7 shows one model calculation of the atmospheric temperature change due to increasing the CO 2 concentration from 315 to 370 ppm, corresponding to the change over the years from 1960 to 2000 ŽRind and Lacis, 1993.. Heating in the troposphere and cooling in the stratosphere are clearly evident. It is interesting that this cooling of the stratosphere due to CO 2 may have some interesting side effects. For example, Yung et al. Ž1997. estimate using model calculations that a doubling of CO 2 would increase the erythema-weighted UV by ;1% at the earth’s surface due to the temperature effect on the UV absorption cross sections for O 3 .

3. Dependence of Net Infrared Absorption on Atmospheric Concentrations Net infrared absorption is determined by the intrinsic strength of the absorption for that particular



molecule and transition Ži.e., the absorption cross section., the effective path length, and the concentration of the absorbing gas. CO 2 , H 2 O, and to a lesser extent O 3 , all absorb infrared radiation strongly in the atmosphere. As discussed in more detail shortly, other infrared-absorbing trace gases, particularly those that have strong absorptions in the relatively clean atmospheric window from 7 to 13 ␮ m where CO 2 , H 2 O, and O 3 do not absorb strongly, also contribute to the net absorption of this radiation. However, as discussed in detail by Shine Ž1991., even a gas that absorbs in the same regions as the major greenhouse gases can contribute to trapping of infrared radiation; the contribution due to an increase in a particular trace gas depends on the combination of absorption region, initial concentration of the gas, and its absorption coefficients. The dependence of absorption on concentration is linear only for weak absorption lines in the atmospheric window; this is the case, for example, for the chlorofluorocarbons. For stronger absorptions such as those due to O 3 , CH 4 , and N2 O, absorption at the peak of the absorption bands approaches saturation; in these case, the net absorption varies with the square root of the absorber concentration. For very strongly absorbing peaks such as those due to CO 2 and H 2 O, absorption only occurs at the fringes of the band and the net absorption varies with the logarithm of the absorber concentration ŽDickinson and Cicerone, 1986; Mitchell, 1989.. The reasons for this are discussed in books devoted to the subject of atmospheric radiation, which should be consulted for details Že.g., see Liou, 1980; Goody and Yung, 1989; and Lenoble, 1993.. A brief account is given in Box 14.1. The absorption cross sections of a variety of gases of atmospheric relevance that are needed for these calculations are available in the literature. See, for example, the HITRAN database ŽRothman et al., 1992..

B. CONTRIBUTION OF TRACE GASES TO THE GREENHOUSE EFFECT Increased atmospheric concentrations of CO 2 , O 3 , and other greenhouse gases over the past century have now been well documented Žvide infra.. All other things being equal, increasing the tropospheric concentrations of infrared-absorbing greenhouse gases will increase the difference between the amount of long-wavelength radiation absorbed and that emitted to space, leading

to a net increase in the energy deposited in, and hence temperature of, the troposphere. One way of thinking of this ŽMitchell, 1989. is shown schematically in Fig. 14.9. Recall from Section A.2c that the emission of radiation to space can be thought of as occurring from a temperature of 254 K, corresponding to an altitude of ;5.5 km. An increased tropospheric concentration of a greenhouse gas means that trapping of the infrared radiation from the earth’s surface will continue to occur to higher altitudes. The molecules that are emitting infrared radiation to space will therefore be at higher altitudes where the tropospheric temperatures are lower, T1 in Fig. 14.9. At these lower temperatures, the populations of the emitting excited states will be lower ŽBoltzmann Eq. ŽA.. and hence the net emission of infrared energy smaller than that needed to balance the net absorbed incoming solar radiation of 235 W my2 . To restore the energy balance between the incoming absorbed solar radiation and the outgoing infrared radiation, the atmosphere at the effective infrared emitting altitude must increase back to ;254 K. If the lapse rate remains the same Žwhich, as discussed later, may not be the case., temperatures then must increase throughout the troposphere.

1. Infrared Absorption by Trace Gases As seen in Figs. 14.1 and 14.2, CO 2 and H 2 O, and to a lesser extent O 3 , are major absorbers of infrared radiation in the troposphere. These gases, along with contributions from CH 4 and N2 O, are responsible for a greenhouse effect that exists without any emissions from anthropogenic activities. However, as seen earlier, any infrared-active atmospheric species can also act as a greenhouse gas. Thus, anthropogenic emissions of new infrared-absorbing species, or increased emissions of the traditional greenhouse gases, provide additional trapping of infrared radiation to higher altitudes, altering the energy balance of the troposphere. For a new compound to be potentially important as a greenhouse gas in the atmosphere, it must have a sufficiently large infrared absorption cross section and be present in large enough concentrations to lead to significant absorption of infrared radiation. In addition, it will be most effective if it absorbs in the infrared window from approximately 7 to 13 ␮ m between the CO 2 and H 2 O absorptions. Finally, the atmospheric lifetime of the gas is important, in that a long-lived species such as the chlorofluorocarbons can make a larger contribution when integrated over time than a short-lived species Žsee Section B.3..



BOX 14.1

MOLECULAR BASIS FOR DIFFERING RELATIONSHIPS BETWEEN INFRARED ABSORPTION AND CONCENTRATION The reason for the different dependencies of absorption on the concentration can be seen starting with the Beer᎐Lambert Law. If LT␯ is the radiance of frequency ¨ incident on an absorber with absorption coefficient ␴ present at concentration N and the effective path length is l, the transmitted radiance L␯ is given by L␯ LT␯

s eyN l ␴ .


ŽFor a definition of radiance, see Chapter 3.C.2.. The absorbed radiance is then given by T yN l ␴ . Ž Labs . ␯ s L␯ 1 y e


Since absorption lines have a finite width and since measurements of absorption in practice are carried out over a small but finite range of frequencies, the absorbed radiance over some interval ⌬¨ is the parameter of interest: Labs LT⌬ ␯ Ž 1 y eyN l ␴ . d␯ . ⌬␯ s


Ž D.

The incident radiation LT␯ usually does not change significantly over one line so that LT⌬ ␯ can be taken as a constant. The remainder of Eq. ŽD. is defined as the equivalent line width, W, Ws

yN l ␴

H Ž1 y e

. d␯ ,


where the integration is carried out over the range of frequencies that encompass the absorption line. The equivalent linewidth can be thought of as the width of a hypothetical absorption line of rectangular shape whose total absorption is equal to that of the real line Žsee Fig. 14.8a.. Equation ŽC. can then be expressed as T Labs ⌬ ␯ s L⌬ ␯ W.

Ž F.

The absorption cross section ␴ depends on the frequency even for a single absorption line due to various line-broadening processes that impart a finite width and particular shape to the absorption line. Figure 14.8b, for example, shows a typical lineshape due to collisional broadening, the Lorentz

FIGURE 14.8 Ža. Meaning of equivalent width, W; Žb. Doppler and Lorentzian line-shapes for equivalent half-widths; Žc. transmission curves for an absorption line for a weak and strong absorber, respectively Žadapted from Lenoble, 1993..

lineshape, and a typical shape due to Doppler broadening. Under atmospheric conditions near the earth’s surface, the linewidth is determined primarily by collisions; i.e., the Lorentz half-width is much larger than the natural linewidth or that due to



Doppler broadening. Collisional broadening becomes less important at the lower pressures found at higher altitudes, so that the Lorentzian half-width and the Doppler half-width become comparable at altitudes of approximately 30᎐40 km. The absorption cross section can be expressed as the product of two factors, an intrinsic line strength S and a shape factor, g x , which depends on the distance x s Ž ¨ y ¨ 0 . from the line center:

␴ s Sg x .

yN lS g x .

Hy⬁ Ž1 y e



a. Weak-Absorber Regime For weak absorptions, i.e., when the combination of concentration and path length Nl is small, Ž1 y eyN lS g . ; NlSg. Equation ŽH. becomes Ws

Hy⬁ Ž NlSg . dx s NlSHy⬁ g x

W s Wcenter q Wwings .


dx s NlS, Ž I .

since the integral is the normalized shape factor, T which is unity. The absorbed radiance Labs ⌬ ␯ s L⌬ ␯ W, so that the net absorption varies linearly with Ž Nl ., i.e., with the column burden of the absorbing gas. This is what is known as the weak-absorber regime and generally applies to such greenhouse gases as the chlorofluorocarbons. However, it should be noted that even for some trace gases such as CF4 , the weak-absorber regime is only obeyed up to ;0.1 ppb Žcurrent atmospheric concentrations of CF4 are ;0.07 ppb; IPCC, 1996.. At 1 ppb, significant deviations are found because CF4 has unusually sharp lines that saturate in the center and that are overlapped by the absorption bands of other atmospheric gases ŽFreckleton et al., 1996.. b. Strong-Absorber Regime However, when the combination of concentration, N, and path length, l, is not small, the weak-absorber approximation is not valid, and the line-shape needs to be taken into account. The reason for this is that in the limit when absorption of light is already saturated at the center of the line, absorption due to added gas occurs only in the wings of the absorption line, which is sensitive to the lineshape. In the case of saturation at the center of the absorption line as shown in Fig. 14.8c, the absorption can be thought of as occurring in two regions, one from


In the center region from x s to x ᎐ s where the absorption is strong, eyN lS g x ª 0 and Wcent s


⬁ g x is the normalized shape factor for which Hy⬁ g x dx s 1. The equivalent linewidth in Eq. ŽE. thus becomes


x s to x ᎐ s around the line center ¨ 0 , where the absorption is saturated, and one in the wings, where the absorption is weaker:


yN lS g x .

Hyx Ž1 y e

dx f 2 x s .



In the wings, ⬁

Wwings s 2


Ž 1 y eyN lS g x . dx,



where the factor of 2 takes into account symmetrical absorption in the wings both on the low- and highfrequency sides of the band center. Thus for strong absorption, W s 2 xs q 2


Ž 1 y eyN lS g x . dx,



and the lineshape in the center of the line is not important, whereas that in the wings at frequencies beyond x s is. Let us return to the definition of equivalent linewidth in Eq. ŽH.. Since lines at most pressures of interest here can generally be described as Lorentzian in shape, the shape factor in Eq. ŽH. is given by ␥L Ž N. g xL s , 2 ␲ Ž x q ␥ L2 . where ␥ L is the Lorentzian line half-width Ži.e., peak width at half-maximum. and x s Ž ¨ y ¨ 0 . as before. This can be substituted into Eq. ŽH. and solved as described in detail elsewhere to get an expression for the equivalent width Že.g., see Liou, 1980; Goody and Yung, 1989; and Lenoble, 1993.. The result for the limit of a strong absorber where absorption in the wings is important can be obtained readily for frequencies in the wings such that x 4 ␥ L . In this case, the lineshape factor reduces to g xL s

␥L ␲ x2



In this region, then, the equivalent linewidth becomes 1r2 WL s 2 Ž 1 y e ŽyN lS␥ L r ␲ x . . dx s 2 Ž NlS␥ L . .




B. CONTRIBUTION OF TRACE GASES TO THE GREENHOUSE EFFECT T Since the absorbed radiance Labs ⌬ ␯ s L⌬ ␯ W, the net absorption in the strong-absorber regime varies as the square root of Ž Nl .. This is the case for O 3 , CH 4 , and N2 O in the atmosphere.

c. Strong Absorptions by CO2 and H2 O The results of a number of laboratory studies wsee Liou Ž1980.; and Goody and Yung Ž1989. for descriptions of thesex have shown that for strong absorptions of CO 2 and H 2 O under conditions similar to those in the atmosphere, the total absorption band absorption area, A, can be described as the sum of three terms: A s C q D ln Ž Nl . q K ln Ž p . .


The three parameters C, D, and K can be obtained by empirical fits to the data and p is the partial pressure of nonabsorbing gases present. Since this total absorption band area is directly related to the equivalent width and hence to the absorbed irradiance, there is a logarithmic dependence of the net absorption on Ž Nl ., which is the case for the strong absorption bands of both water vapor and carbon

For gases that satisfy these conditions, the effects can be proportionately quite large. For example, addition of one molecule of the chlorofluorocarbons ŽCFCs. CFC-11 and CFC-12 is equivalent to the addition of ;10 4 additional molecules of CO 2 due to the stronger absorption cross sections of the CFCs that occur in the atmospheric window and to the dependence of absorption on concentration for the CFCs but on the logarithm of concentration for CO 2 ŽRamanathan et al., 1987.. Figure 14.10 shows the absorption bands and approximate absorption band strengths for a number of molecules found in the troposphere ŽRamanathan et al., 1987; Ramanathan, 1988a, 1988b.. There are many gases that, on the basis of intrinsic absorption strengths in the atmospheric window, can, in principle, contribute to tropospheric heating. However, the third requirement is that they be present in sufficient concentration to lead to significant infrared absorption. Of the molecules shown in Fig. 14.10, the ones that meet all of these requirements are CH 4 , N2 O, the chlorofluorocarbons ŽCFCs. and other halocarbons such as methylchloroform, and some perfluorinated compounds such as SF6 Žsee Chapters 12 and 13.. As we shall see in the next section, the concentra-


dioxide in the atmosphere. As discussed by Goody and Yung Ž1989., the empirically observed logarithmic dependence of absorption on concentration can be shown to be consistent with theoretical expectations based on reasonable assumptions of bandshape and line intensities. It should be noted that the foregoing considerations apply to the major absorption bands. In some cases, weaker absorption bands of the major greenhouse gases can be sufficiently weak to fall in the linear region. This is the case, for example, for light absorption by O 3 in the Chappius band, even if the strong Hartley᎐Huggins band Žsee Chapter 4.B. is saturated Že.g., see Lacis et al., 1990.. These weaker bands can have significant effects on the calculated outgoing infrared radiation. For example, Ho et al. Ž1998. show that much of the reported discrepancy between modeled outgoing long-wavelength radiation and satellite measurements can be attributed to not including weaker absorption bands due to CO 2 at 4.3 ␮ m and O 3 at 14 ␮ m and the weaker O 3 lines located far from the center of the strong 9.6-␮ m band.

tions of all of these ‘‘trace’’ greenhouse gases, as well as CO 2 and O 3 , have been increasing over the past century or more.

2. Trends in Trace Gas Concentrations a. CO2 Carbon is, of course, extensively recycled through the earth system, including both the terrestrial biosphere and the oceans. Figure 14.11 summarizes this cycling and where the reservoirs of carbon are found. Anthropogenic activities contribute to atmospheric carbon mainly in the form of CO 2 emissions from fossil fuel combustion and, to a lesser extent, cement production, which total 5.5 Gt of C per year Žwhere 1 Gt of C s 10 9 metric tons s 10 15 g of carbon.. The amount of carbon in hydrocarbons, including CH 4 , and CO is less than 1% of the total atmospheric carbon ŽIPCC, 1996.. Changes in land use, including biomass burning, also contribute to changing the balance, although the net quantitative contribution is less certain. Land use changes in the tropics during the decade from 1980 to 1990 are estimated to have contributed approximately 1.6 Gt of C per year ŽIPCC, 1996., but this does not



FIGURE 14.9 Warming of troposphere by increased greenhouse gas concentrations Žadapted from Mitchell, 1989..

include a number of potential sinks due to increased growth of forests, which may be associated with increased atmospheric CO 2 and nitrogen deposition Žsee Chapter 7.. For example, C. P. Keeling et al. Ž1996. have measured increases in the annual amplitude of the seasonal CO 2 cycle Žvide infra. in Hawaii and in

FIGURE 14.10 Intrinsic infrared absorption band strengths of some potential greenhouse gases in the atmospheric ‘‘window’’ Žfrom Ramanathan, 1988a, 1988b..

the Arctic, which they attribute to increased uptake of CO 2 by land vegetation during periods of warmer temperatures. wIt is interesting that coverage by snow does not appear to terminate the exchange of CO 2 ŽSommerfeld et al., 1993..x There are different time scales associated with the various emissions and uptake processes. Two terms that are frequently used are turno¨ er time and response Ž or adjustment . time. The turnover time is defined as the ratio of the mass of the gas in the atmosphere to its total rate of removal from the atmosphere. The response or adjustment time, on the other hand, is the decay time for a compound emitted into the atmosphere as an instantaneous pulse. If the removal can be described as a first-order process, i.e., the rate of removal is proportional to the concentration and the constant of proportionality remains the same, the turnover and the response times are approximately equal. However, this is not the case if the parameter relating the removal rate and the concentration is not constant. They are also not equal if the gas exchanges between several different reservoirs, as is the case for CO 2 . For example, the turnover time for CO 2 in the atmosphere is about 4 years because of the rapid uptake by the oceans and terrestrial biosphere, but the response time is about 100 years because of the time it takes for CO 2 in the ocean surface layer to be taken up into the deep ocean. A pulse of CO 2 emitted into the atmosphere is expected to decay more rapidly over the first decade or so and then more gradually over the next century. Figure 14.12 shows what has become classic data illustrating the increase in CO 2 concentrations at Mauna Loa, Hawaii, where continuous measurements have been made since 1958 ŽKeeling et al., 1995.. The concentrations have risen from approximately 315 ppm in the late 1950s to 358 ppm in 1994. The cyclical pattern superimposed on the continuous increase reflects decreased CO 2 concentrations during summer and increased CO 2 during winter in response to seasonal differences in uptake during plant growth. The amplitude of the cyclical pattern in the Northern Hemisphere is largest at the most northerly locations, decreasing from ;15᎐20 to ;3 ppm near the equator, where plant growth is less dependent on season ŽKeeling et al., 1995.. Superimposed on the CO 2 concentration measurements in Fig. 14.12 are the concentrations expected if 55.9% of the cumulative CO 2 emissions from fossil fuel combustion and cement production remained in the atmosphere ŽKeeling et al., 1995.. This percentage was chosen to match the atmospheric observations for the 20-year period between January 1, 1959, and January 1, 1979; the match between the two curves shows that



FIGURE 14.11 Summary of global carbon cycle. Amount Žin gigatons of C s 10 9 metric

tons s 10 15 g of C.. Reservoirs are shown in parentheses, and fluxes Žgigatons of C per year. are indicated by arrows. Note that the time scales associated with the various processes vary Žadapted from IPCC, 1996..

slightly more than half of the CO 2 that has been emitted to date remains in the atmosphere. ŽA detailed analysis of the data in Fig. 14.12 shows a slight anomaly in the 1980s in that the CO 2 concentrations were higher than expected based on industrial emissions; Keeling and co-workers Ž1995. suggest this is due to changes in terrestrial and ocean sinks associated with changes in global temperatures.. Similar increases in CO 2 have been documented at locations around the world, including the South Pole, where measurements have been made since 1957. Most of the remaining emissions of CO 2 that have been removed from the atmosphere have been taken up by land ecosystems, with a small contribution from the oceans ŽTans and White, 1998..

Figure 14.13 shows CO 2 concentrations measured in ice cores at the Byrd Station in Antartica from 5000 years before the present Žbp. to 40,000 years bp ŽAnklin et al., 1997.. The use of ice core data for elucidating atmospheric composition is discussed by Delmas Ž1992. and in more detail in Section E.1. As seen in Fig. 14.13, atmospheric CO 2 concentrations about 5000 years ago were only ;280 ppm. ŽNote that interpretation of such ice core data must be carried out with care since there is evidence that in some cases, CO 2 can be produced in the ice from decomposition of carbonate; e.g., see Smith et al., 1997.. In short, it appears that CO 2 concentrations prior to the industrial age were even smaller than those measured starting in the late 1950s.



FIGURE 14.12 Measured CO 2 concentrations at Mauna Loa, Hawaii, from 1958 to 1994. The line represents the atmospheric CO 2 expected if 55.9% of the cumulative emissions of CO 2 from fossil fuel combustion and cement production remained in the atmosphere Žadapted from Keeling et al., 1995..

Production of atmospheric CO 2 due to the combustion of fossil fuels and land use changes, and uptake by the terrestrial biosphere, all involve stoichiometric changes in O 2 as well, whereas uptake of CO 2 by the oceans does not. Thus, a decrease in O 2 should accompany the increase in atmospheric CO 2 . Detecting small changes in atmospheric O 2 is difficult. However, R. F. Keeling and co-workers Ž1998. have developed the necessary methodology and applied it to the atmosphere ŽKeeling and Shertz, 1992; R. F. Keeling et al., 1996.. For example, Fig. 14.14 shows the anticorrelation observed by Keeling and Shertz Ž1992. at La Jolla, California, between changes in oxygen Žmeasured relative to N2 . and CO 2 in the atmosphere from 1989 to 1992. Similarly, Bender et al. Ž1994. measured changes in O 2 relative to N2 in air bubbles trapped in ice cores at Vlostok, Antarctica, and report an increase in trapped O 2 and a decrease in CO 2 with increasing depth, i.e., as a function of time before the present. As seen in Fig. 14.15, the rate of atmospheric increase of CO 2 has been quite variable over the past

FIGURE 14.13 Concentrations of atmospheric CO 2 measured using gases trapped in ice cores from Byrd Station, Antarctica, from 5000 to 40,000 years before the present Žbp. Žadapted from Anklin et al., 1997..

FIGURE 14.14 Anticorrelation between atmospheric O 2 and CO 2

at La Jolla, California. ␦ O 2 rN2 is defined as 10 6 wŽO 2 rN2 .airr ŽO 2rN2 . ref x y 14, where ŽO 2 rN2 . ref is the ratio for a reference Žadapted from Keeling and Shertz, 1992..

four decades of atmospheric monitoring. The rate of increase at Mauna Loa was ;0.8 ppm per year during the 1960s, 1.3 ppm per year in the 1970s, and 1.5 ppm per year during the 1980s ŽIPCC, 1996.. However, the rate of increase slowed during the 1989᎐1993 period, with the value in 1992 of 0.6 ppm per year being the smallest rate of increase since continuous monitoring began ŽIPCC, 1996; Conway et al., 1994.. The growth rate subsequently increased again to over 2 ppm per year in 1994 ŽIPCC, 1996.. The reasons for the changes in the rate of growth are not clear, but probably not surprising given the complex cycling mechanisms for carbon ŽFig. 14.11.. For example, exchange between the atmosphere and the terrestrial biosphere and the oceans is believed to have substantial year-to-year variability, which can be affected by such events as the El Nino ˜ Southern Oscillation ŽENSO.. Possible reasons for the variability are discussed in detail in the IPCC document Ž1996.. Current data on atmospheric CO 2 concentrations and temporal trends as well as those of other trace gases are available from the U.S. Department of Energy’s Carbon Dioxide Information Analysis Center ŽCDIAC . ŽWorld Wide Web page is http:rrcdiac.esd.ornl.govrcdiac.. CDIAC also has available a number of additional data packages on global change issues such as trends in temperature and precipitation. As discussed earlier, although CO 2 warms the troposphere, it cools the stratosphere since it efficiently radiates infrared out to space. This effect can contribute to changes in the temperature profile in the stratosphere and potentially have a signficant impact



FIGURE 14.15 Rate of growth Žppm per year. of atmospheric CO 2 at Mauna Loa, Hawaii, from 1958 to 1994 Žfrom Keeling and Worf as reported in IPCC, 1996..

on atmospheric circulation processes Že.g., see Rind et al., 1990; and Rind and Lacis, 1993.. b. CH4 Like CO 2 , methane is emitted by both natural and anthropogenic processes. While the major sources are thought to have been identified, there is some uncertainty in the absolute magnitudes of their contributions as well as the factors that affect these ŽCicerone and Oremland, 1988; Fung et al., 1991.. Table 14.1 shows one estimate of methane sources during the mid-1980s, expressed in units of teragrams Ž10 12 g. of carbon per year ŽCrutzen, 1995.. Approximately 60 " 10% of the total emissions Ž370 out of a total of 630 Tg per year. is estimated to be associated with human activities. Of the 370 Tg of C per year, approximately 35% is due to losses during natural gas and oil production and distribution and coal mining. This estimate is reasonably consistent with measurements of the 14 C content of atmospheric methane, since fossil fuel derived methane is depleted in 14 C due to the long time frame for the fuel formation ŽLowe et al., 1988; Wahlen et al., 1989.. The next largest source, approximately 30% of the 370 Tg of C per year, is due to emissions from domesticated ruminant livestock Že.g., see Johnson et al., 1994. and from the decay of animal wastes. Emissions from rice fields Že.g., see Cicerone and Shetter, 1981; Cicerone et al., 1983, 1992; and Tyler et al., 1994. appear to comprise about 20%. The remainder of the 370 Tg of C per year is believed to be due about equally to emissions from sanitary landfills Že.g., see

Bogner and Spokas, 1993. and from biomass burning Že.g., see Hao and Ward, 1993.. It should be noted that the emissions in some cases, for example rice fields and landfills, represent the net flux of emission and microbially mediated oxidation processes so that both need to be understood in assessing the methane budget Že.g., Reeburgh et al., 1993; Bogner and Spokas, 1993.. There were relatively few measurements of atmospheric methane concentrations prior to about 1980, except for a set from 1963 to 1970 by Stephens and co-workers ŽStephens and Burleson, 1969; Stephens, 1985., which were in the 1.37᎐1.57 ppm range. The current global mean concentration of methane is 1.72 ppm, with higher concentrations in the Northern than

TABLE 14.1 Estimated Methane Sources during the Mid-1980s a Source Natural Anthropogenic Gas leakage and oil production Coal mining Rice fields Ruminants Biomass burning Animal wastes Sanitary landfills a b

From Crutzen Ž1995.. Tg of C s 10 12 g of C.

Emissions (Tg of Cr r year) b 260 " 30 370 " 40 85᎐105 25᎐45 20᎐150 65᎐100 20᎐60 20᎐40 20᎐60



FIGURE 14.16 Averaged 3-D methane concentrations in the marine boundary layer. Lines are guides for the eye Žadapted from Dlugokencky et al., 1994a..

in the Southern Hemisphere. Figure 14.16 shows the latitudinal distribution as a function of time from 1983 to 1992 ŽDlugokencky et al., 1994a.. The hemispheric distribution is clearly seen, as is the seasonal cycle due to changes in oxidation by OH and in methane emission sources. Even in this relatively short time span, the increase in the atmospheric concentrations is evident. Unlike CO 2 , there does not appear to be a discernible trend in the amplitude of the seasonal cycles ŽDlugokencky et al., 1997.. Ice core data show that the concentration prior to about the year 1750 was ;700 ppb, less than half of the current global average. The increase appears to have begun in the 1750᎐1800 period Že.g., see Khalil and Rasmussen, 1987, 1994b; Blunier et al., 1993; and Etheridge et al., 1998.. Figure 14.17, for example, shows the concentrations of atmospheric methane for the past approximately 1000 years ŽEtheridge et al., 1998.. The increase in concentration to the present value is well outside the variations of ;70 ppb observed prior to 1750. It is interesting that the increase appears to have begun prior to significant industrial activity, but parallels the increase in population growth in China; Blunier and co-workers Ž1993. suggest that this may be due to associated changes in emissions from rice fields. The rate of increase of atmospheric CH 4 has been variable. From 1978 to 1987, the average growth rate was approximately 16 ppb per year ŽBlake and Rowland, 1988.. However, the global growth rate slowed during the latter part of the 1980s ŽSteele et al., 1992; Dlugokencky et al., 1994a, 1994b, 1998.. Figure 14.18

shows globally averaged methane concentrations ŽFig. 14.18a. as well as the growth rate for CH 4 for latitudes from 82⬚N to 90⬚S from 1984 to 1996 ŽFig. 14.18b. ŽDlugokencky et al., 1998.. Dlugokencky and co-workers suggest that the apparent leveling off of CH 4 may not reflect a change in sources or sinks but rather reflect an approach to steady state. The growth rates in the late 1980s were lower, but there was a sharp rise in 1991 immediately after the Mount Pinatubo volcanic eruption, followed by a sharp decrease to temporarily

FIGURE 14.17 Atmospheric methane concentrations over the past 1000 years. Different symbols represent data from ice cores in Antarctica and Greenland and the Antarctic firm layer. Line from 1978 includes air measurements at Cape Grim, Tasmania Žadapted from Etheridge et al., 1998..



fractionation of both 13 C and D have been used to estimate the fraction of CH 4 oxidized during transport through the coversoils in landfills, for example Že.g., see Bergamaschi et al., 1998a; and Liptay et al., 1998., and to identify sources of CH 4 in the troposphere Že.g., Bergamaschi et al., 1998b; Moriizumi et al., 1998.. Recent studies of the isotopic composition in the upper troposphere show that the methane is enriched in 13 C in a manner that is not consistent with known kinetic isotope effects for CH 4 reactions ŽTyler et al., 1998., again demonstrating the complexity of quantitatively defining the sources and sinks of methane in the atmosphere. c. N2 O

FIGURE 14.18 Globally averaged Ža. CH 4 concentrations and Žb. growth rates from 1984 to 1996 from 82⬚N to 90⬚S latitude Žadapted from Dlugokencky et al., 1998..

negative values. The increase after the eruption is attributed by Dlugokencky and co-workers Ž1996, 1998. to reduced removal rates for CH 4 by its reaction with OH; light scattering by the volcanic aerosol particles and UV absorption by SO 2 decreased UV, and hence decreased OH would be expected in the troposphere. Reasons for the sharp decrease during late 1992 and 1993 are not clear but may involve such factors as smaller natural emissions from wetlands due to lower temperatures following the eruption ŽHogan and Harriss, 1994; Dlugokencky et al., 1998. andror changes in anthropogenically associated source strengths such as decreased emissions from the former U.S.S.R. and decreased biomass burning Že.g., see Dlugokencky et al., 1994b, 1998; Rudolph, 1994; and Crutzen, 1995.. Increased removal by OH due to increased UV associated with stratospheric ozone destruction may also have contributed to these trends in CH 4 ŽBekki et al., 1994.. While quantifying the sources and sinks of CH 4 has been difficult, isotopic measurements of 13 CH 4 and CH x D4 ᎐ x are promising. Various sources have characteristic isotopic signatures; e.g., as mentioned previously, fossil fuel derived CH 4 is depleted in 14 C ŽLowe et al., 1988, 1994; Wahlen et al., 1989.. The sinks of CH 4 , e.g., reaction with OH, reaction with Cl, and uptake by soil bacteria, also exhibit kinetic isotope effects and these have been used to probe the causes of the observed recent changes in CH 4 growth rates Že.g., see Gupta et al., 1996, 1997.. Measurements of isotopic

Nitrous oxide is important not only as a greenhouse gas but, as discussed in Chapter 12, as the major natural source of NO x in the stratosphere, where it is transported due to its long tropospheric lifetime ŽCrutzen, 1970.. The major sources of N2 O are nitrification and denitrification in soils and aquatic systems, with smaller amounts directly from anthropogenic processes such as sewage treatment and fossil fuel combustion Že.g., see Delwiche, 1981; Khalil and Rasmussen, 1992; Williams et al., 1992; Nevison et al., 1995, 1996; Prasad, 1994, 1997; Bouwman and Taylor, 1996; and Prasad et al., 1997.. The use of fertilizers increases N2 O emissions. For pastures at least, soil water content at the time of fertilization appears to be an important factor in determining emissions of N2 O Žand NO. ŽVeldkamp et al., 1998.. Table 14.2 shows one estimate of the contribution of various sources to the N2 O budget ŽBouwman and Taylor, 1996.. While the major source of N2 O is known to be biological, there are several observations that

TABLE 14.2 Estimated Annual N 2 O Budget a Source Soils under natural vegetation and grasslands Arable lands Nitrogen fertilizer use Animal wastes Biomass burning Agricultural waste burning Postclearing enhanced soil flux Fossil fuel combustion and traffic Biofuel combustion Industry Oceans Total sources a

From Bouwman and Taylor Ž1996..

Emissions r year) Tg of Nr 5.7 1.0 1.0 1.0 0.1 0.1 0.4 0.3 0.1 0.5 3.6 13.8



indicate that the sources are not yet entirely characterized. For example, N2 O in the lower stratosphere has been shown to be isotopically enriched in both 18 O and 15 N relative to tropospheric N2 O ŽKim and Craig, 1993.. N2 O emissions from tropical rain forest soils, fertilized soils, and a wastewater treatment facility are lighter in both of these heavy isotopes than tropospheric N2 O ŽYoshinari and Wahlen, 1985; Wahlen and Yoshinari, 1985; Kim and Craig, 1990, 1993.. Either there is a source of N2 O in the stratosphere that selectively produces heavy N2 O or one that in the stratosphere selectively destroys light N2 O ŽJohnson et al., 1995.. Despite a number of studies, the source of this discrepancy is not yet clear ŽMcElroy and Jones, 1996; Wingen and Finlayson-Pitts, 1998; Cliff and Thiemens, 1997; Rahn and Wahlen, 1997., although such processes as enhanced photolysis of the 14 N 14 N 16 O isotopomer ŽYung and Miller, 1997; Rahn et al., 1998. andror formation of N2 O by reaction of highly vibrationally excited O 3 with N2 ŽZipf and Prasad, 1998. have been proposed. One interesting potential source of N2 O is the heterogeneous oxidation of HONO on surfaces ŽWiesen et al., 1995; Pires and Rossi, 1997., which has been observed to form N2 O. This is likely responsible for the observation of significant amounts of N2 O in automobile exhaust, which was shown to be an artifact of sampling ŽMunzio and Kramlich, 1988.. However, it may also occur on aerosol particles in the atmosphere ŽClemens et al., 1997., an area that warrants further investigation. Like CO 2 and CH 4 , the concentrations of N2 O in the atmosphere have also been increasing, from ;275 ppb in the preindustrial era to ;312 ppb in 1994 ŽIPCC, 1996.. Figure 14.19, for example, shows one set of measurements of N2 O over the past 250 years obtained using ice core samples from Antarctica

FIGURE 14.19 Fitted curve for atmospheric N2 O concentrations from 1735 to 1991 obtained from ice cores in Antarctica Žadapted from Machida et al., 1995..

ŽMachida et al., 1995.. The rate of growth has been variable, averaging 0.8 " 0.2 ppb per year from 1977 to 1988 ŽKhalil and Rasmussen, 1992., but was only 0.5 ppb per year in 1993 ŽIPCC, 1996.. d. O3 As discussed in other chapters of this book and summarized in Chapter 16, the formation of tropospheric ozone from photochemical reactions of volatile organic compounds ŽVOC. and oxides of nitrogen ŽNO x . involves many reactions. Concentrations are therefore quite variable geographically, temporally, and altitudinally. Additional complications come from the fact that there are episodic injections of stratospheric O 3 into the troposphere as well as a number of sinks for its removal. Because O 3 decomposes thermally, particularly on surfaces, it is not preserved in ice cores. All of these factors make the development of a global climatology for O 3 in a manner similar to that for N2 O and CH 4 , for example, much more difficult. In addition, the complexity of the chemistry leading to O 3 formation from VOC and NO x is such that modelpredicted ozone concentrations can vary from model to model Že.g., see Olson et al., 1997.. Shortly after the discovery of ozone by Schonbein in ¨ 1839, measurements of this newly discovered atmospheric gas were initiated in a number of locations around the world, including Europe, South America, and North America Že.g., see Bojkov, 1986; Volz and Kley, 1988; McKeen et al., 1989; Anfossi et al., 1991; Sandroni et al., 1992; and Marenco et al., 1994.. Although these very early measurements used wet chemical techniques Žsee Chapter 11., potential interferences can be estimated to make some approximate corrections to the data. Typical annual variations in tropospheric O 3 observed at Moncalieri in Italy from 1868 to 1893, Montsouris, France, from 1876 to 1886, and Zagreb, Croatia, in 1900 are compared in Fig. 14.20 to more recent data Ž1983. at Arkona, on an island believed to be relatively remote, in the Baltic Sea ŽSandroni et al., 1992.. All of the early measurements peak around 10 ppb, whereas 30᎐40 ppb is a typical tropospheric O 3 concentration found essentially everywhere in the world today. Global increases in ozone have also been documented over more recent times, although the geographic distribution and temporal changes are complex Že.g., see Volz et al., 1989; Janach, 1989; Lefohn et al., 1992; Low et al., 1990; Logan, 1994; Jiang and Yung, 1996; and Stockwell et al., 1997.. Berntsen et al. Ž1997. carried out model calculations that showed that significant radiative effects must have resulted from changes in O 3 since preindustrial times in the upper troposphere due to in situ formation from



also solar UV radiation, with associated effects on climate Žsee also Section B.3b.. These phenomena are illustrated in Fig. 14.21a, which shows a model calculation of the change in global surface temperature when 10 Dobson units ŽDU. of O 3 Ž10 DU s total column O 3 equivalent to a layer of thickness 0.1 mm at 273 K and a pressure of 1 atm; see Chapter 12.A. are added one at a time to each of 33 vertical layers of the atmosphere Žassuming no feedbacks.. An increase in the global surface temperature is predicted when the ozone is added in layers up to

FIGURE 14.20 Monthly average O 3 concentrations at Moncalieri, Italy; Montsouris, France; and Zagreb, Croatia in the 1868᎐1900 period and Arkona ŽBaltic Sea. in 1983 Žadapted from Sandroni et al., 1992..

the reactions of precursors transported from the boundary layer. In short, the concentrations of tropospheric ozone, which is also a greenhouse gas, have also increased over the past century, an increase attributed to increased oxides of nitrogen emissions associated with fossil fuel combustion Že.g., see Volz and Kley, 1988; and Janach, 1989.. The potential effects of ozone on climate are particularly complex in that the net effect is very sensitive to the vertical distribution profile, with changes at the tropopause having the largest impact ŽWang et al., 1980; Lacis et al., 1990.. The reason for this is that O 3 near the ground is at temperatures close to those of the earth’s surface. As a result, emission and absorption are occurring at essentially the same temperature, resulting in no contribution to the greenhouse effect. However, because the temperature falls with altitude up to the tropopause, the Boltzmann distribution ŽEq. ŽA.. shifts to smaller relative populations in the excited states. Thus, as discussed earlier, the net emission from O 3 becomes smaller relative to absorption. While the same is true for other greenhouse gases such as CO 2 and CH 4 , their sources and sinks are such they are relatively well mixed in the atmosphere and their vertical distributions are not subject to the variability associated with O 3 . Another important difference is that O 3 absorbs strongly in the UV as well, which leads to heating in the stratosphere, in contrast to CO 2 , which cools it. Thus, changes in the concentrations of ozone and its vertical distribution affect not only infrared but

FIGURE 14.21 Ža. Model-calculated change in global surface temperature when 10 Dobson units ŽDU. of O 3 is added to each of 33 vertical layers of the atmosphere, assuming no feedbacks Žadapted from Lacis et al., 1990.. Žb. Global mean vertical O 3 profile Žadapted from Forster and Shine, 1997.. Žc. Model-calculated change in global surface temperature due to increasing O 3 concentrations sequentially by 10% in each 1-km layer of the atmosphere Žadapted from Forster and Shine, 1997..



;30 km. ŽThe gap around 10 km is an artifact of the calculations. . However, increased O 3 above ;30 km causes cooling due to increased thermal emission to space and to increased absorption of solar radiation before it can reach the earth’s surface Že.g., see Ramanathan et al., 1985.. Addition of a constant, absolute amount of O 3 may be unrealistic, however, in that 10 DU is a large percentage change in the existing O 3 concentration at lower altitudes but a smaller percentage at the altitudes at which the ozone concentration peaks. Figure 14.21b, for example, shows one global mean vertical profile for O 3 Žin Dobson units.. Adding 10 DU at the tropopause corresponds to an increase in O 3 of ;400%, whereas at 24 km, adding 10 DU only increases O 3 by ;60% ŽForster and Shine, 1997.. Figure 14.21c shows a calculated change in surface temperature due to a systematic change in O 3 of 10% in each layer. While O 3 at the tropopause is still important, that at lower and higher altitudes is relatively more important compared to the case in Fig. 14.21a where an absolute increase in O 3 in each layer is assumed. A number of model studies have explored the climate implications of changes in the vertical distribution of ozone Že.g., see Schwarzkopf and Ramaswamy, 1993; Wang et al., 1993; Molnar et al., 1994; and Chalita et al., 1996.. For example, Fig. 14.22a shows a modelcalculated percentage change in tropospheric O 3 in July as a function of altitude and latitude from preindustrial times to the present ŽChalita et al., 1996.. The largest increase in concentration is predicted in the boundary layer, with smaller increases at higher altitudes. However, as seen in Fig. 14.22b, the major contribution to radiative forcing Žsee Section B.3. comes from the relatively small ozone increase predicted for the 6- to 12-km region. Changes in stratospheric ozone also impact atmospheric temperatures through three effects. As discussed in Chapter 3, UV absorption by stratospheric

O 3 and the energy released in the O q O 2 reaction warm the stratosphere. As a result, destruction of stratospheric ozone due to chlorofluorocarbons results in cooling of the stratosphere. From the Boltzmann relationship ŽEq. ŽA.., there is then less downward radiation across the tropopause from ozone at these lower temperatures. In addition, there is the direct effect of a smaller ozone concentration to emit infrared radiation, part of which is in a downward direction and which normally contributes to heating of the troposphere. These two effects of changes in stratospheric ozone lead to cooling of the troposphere. Counterbalancing these effects is the increased solar UV reaching the troposphere due to less absorption by stratospheric O 3 , which is expected to cause surface heating Že.g., see Ramaswamy et al., 1992, 1996; Zhong et al., 1996; and Shine et al., 1998.. It should be noted that while changes in stratospheric ozone can impact tropospheric heating and cooling, greenhouse gases may also impact stratospheric ozone destruction. As described earlier in this chapter, while CO 2 causes warming in the troposphere, it causes cooling in the stratosphere through the efficient emission of infrared to space. Some model calculations suggest that additional stratospheric cooling in the polar regions due to increased greenhouse gases may increase the formation of polar stratospheric clouds. Since these play such a key role in ozone destruction in those regions Žsee Chapter 12.C.5., increased ozone destruction is predicted, particularly in the Arctic, where temperatures are not as routinely cold as in the Antarctic ŽAustin et al., 1992, 1994; Shindell et al., 1998.. ŽIt should be noted, however, that warming of the troposphere by trapping of outgoing terrestrial radiation by PSCs during winter has also been proposed as being important historically at times of high methane concentrations that oxidized to form water; Sloan and Pollard, 1998.. In addition, it has been suggested that the recovery of these regions as the

FIGURE 14.22 Model-calculated percentage Ža. increase in O 3 Žzonal average. in July and Žb. the corresponding contributions to instantaneous radiative forcing calculated as a function of latitude and altitude from preindustrial times to the present Žadapted from Chalita et al., 1996..


emissions of ozone-destroying chlorine and bromine compounds decline may be delayed by about a decade due to this effect of greenhouse gases ŽShindell et al., 1998.. There are a number of factors that affect the ultimate climate response to changes in tropospheric and stratospheric ozone. These include the altitude dependence of the forcing previously discussed, its role in absorbing solar UV in both the stratosphere and troposphere, its depletion through chain reactions of CFCs in the stratosphere, and finally, the large variability in its concentrations geographically, vertically, and temporally. Because of these complexities, the net effect is expected to also vary from one location to another, as well as temporally. The spatial and temporal effects due to ozone formed from emissions from biomass burning over large areas of the tropics are one example ŽPortmann et al., 1997.. A number of studies have addressed the net effect of changes in O 3 on climate, and the reader is referred to them for more detailed information Že.g., see Hauglustaine et al., 1994; Marenco et al., 1994; Lelieveld and van Dorland, 1995; Forster et al., 1996; Chalita et al., 1996; Portmann et al., 1997; Berntsen et al., 1997; van Dorland et al., 1997; Graf et al., 1998; Haywood et al., 1998c; Brasseur et al., 1998; Stevenson et al., 1998; and Wang and Jacob, 1998.. e. CFCs, HCFCs, and HFCs The atmospheric concentrations of chlorofluorocarbons ŽCFCs., hydrochlorofluorocarbons ŽHCFCs., and hydrofluorocarbons ŽHFCs. and the trends in these concentrations are discussed in detail in Chapter 13. In brief, the atmospheric concentrations of CFCs increased as they came into increasing use in the 1950s. As the phase-outs specified in the Montreal Protocol and its amendments have come into play ŽFigs. 12.13 and 13.4., the growth rates have fallen dramatically. For example, that for CFC-12 was ;18 ppt per year prior to the Montreal Protocol but in mid-1993 had fallen to 6᎐7 ppt per year Že.g., see Cunnold et al., 1997. and has since become slightly negative Že.g., Derwent et al., 1998a.. However, the concentrations of their alternatives, the HCFCs and HFCs, are increasing as expected Žsee Fig. 13.6.. Although they do not contribute significantly to radiative forcing at present, they could do so if their emissions approach those of the compounds they are replacing ŽDerwent et al., 1998b.. f. Other Gases Other anthropogenically emitted gases such as CO have also been suggested to contribute to the greenhouse effect Že.g., see Evans and Puckrin, 1995.. CO concentrations also increased during the 1980s but


then decreased from 1988 to 1992 Že.g., see Khalil and Rasmussen, 1984, 1994a; Novelli et al., 1994; and Yurganov et al., 1997.. CO is not believed to directly contribute significantly to the greenhouse effect ŽIPCC, 1996.. However, increasing CO emissions may decrease the OH concentration, which would then increase the concentrations of other greenhouse gases that react with OH, such as CH 4 . For example, Daniel and Solomon Ž1998. estimate that this indirect effect associated with anthropogenic emissions may be as or more significant over the next 15 years than that due to anthropogenic emissions of N2 O.

3. Radiative Forcing by Greenhouse Gases and Global Warming Potentials In the simplest of worlds, the greenhouse gases would exert their influences independent of each other and of other factors such as aerosols and clouds, feedback mechanisms, and ozone depletion. This, of course, is not the case. However, it is useful before examining these ‘‘real-world’’ considerations to consider the direct effects on the radiation balance of the atmosphere of the greenhouse gases. These are commonly expressed in terms of the radiati¨ e forcing. Another tool used for examining the relative effects of various gases on the radiation balance of the atmosphere is the global warming potential. a. Instantaneous and Adjusted Radiati© e Forcing As discussed at the beginning of this chapter, changes in the radiation balance of the atmosphere can occur due to changes either in incoming solar radiation or in the outgoing infrared radiation. Radiati¨ e forcing is defined as a change in the average net radiation at the tropopause due to a particular perturbation of interest. This change Žusually expressed in W my2 . could be in either the incoming or outgoing radiation. Using the flux at the tropopause to define radiative forcing is believed to be appropriate because of the rapid vertical mixing by convection and large-scale processes within the troposphere which closely couples the troposphere and the earth’s surface ŽWang et al., 1995.. As a result, energy absorbed in the troposphere is assumed to be effective in warming the earth’s surface ŽLacis et al., 1990. and the change in the flux at the tropopause can be used to calculate the change in the surface temperature ŽRamanathan, 1976.. Two approaches to calculating radiative forcing due to greenhouse gases have been taken. In the first, the immediate forcing due to increases in the greenhouse gas is calculated without allowing for a change in the stratospheric temperature. This is what is known as the instantaneous radiati¨ e forcing.



The second approach is to calculate the radiative forcing after allowing stratospheric temperatures to readjust to radiative equilibrium, but with the temperatures of the earth’s surface and troposphere, as well as atmospheric moisture, fixed. This is known as the adjusted radiati¨ e forcing. The reason for allowing a stratospheric readjustment is that an increase in a greenhouse gas increases the net radiation absorbed in the troposphere. As a result, there is less upwelling radiation reaching the stratosphere ŽFig. 14.2c.. This causes cooling of the stratosphere, which decreases the net downward radiative flux from the stratosphere at the tropopause, contributing a negative component to the net radiative forcing attributable to an increase in a greenhouse gas ŽWang et al., 1995.. The time for the stratosphere to adjust is of the order of months ŽManabe and Strickler, 1964., so that for the longer term perturbations of interest, this adjustment will occur and decrease the net radiative forcing from the instantaneous value. This is illustrated by the data in Table 14.3, which shows the calculated instantaneous and adjusted radiative forcing attributed to the increase in tropospheric O 3 from preindustrial times to the present ŽBerntsen et al., 1997.. Table 14.3 also illustrates the larger relative importance of absorption of long-wavelength IR by ozone compared to short-wavelength UV. wNote that these calculations represent the contribution due to changes in tropospheric ozone only; as discussed earlier and in the following text, the destruction of stratospheric ozone leads to a negative forcing, i.e., cooling. Indeed, it appears that this cooling effect is likely dominant at the present time ŽHansen et al., 1997b..x Radiative forcing can be calculated for greenhouse gases in a fairly straightforward manner, particularly in the simplest case where there are no feedbacks or indirect effects on the chemistry of the atmosphere. However, translating these radiative forcings into real temperature changes at the earth’s surface is much more uncertain due to the complex feedbacks involving

TABLE 14.3 Radiative Forcing a Due To Changes in Tropospheric Ozone from Preindustrial Times to the Present Time for Clear Skies Calculated Using the Oslo Model b Radiative forcing


Long-wavelength thermal infrared

Short-wavelength solar UV

Instantaneous Adjusted

0.48 0.39

0.42 0.33

0.06 0.06

Global mean and annual mean radiative forcing in W my2 . b Adapted from Bernsten et al. Ž1997.. a

chemistry, physics, and atmospheric dynamics. Empirical relationships of the form of Eq. ŽR., ⌬Ts Ž ⬚C. s ␦ Fa ,

Ž R.

between the change in temperature at the earth’s surface, ⌬Ts , and the adjusted radiative forcing Fa Žin W my2 . are often used ŽHansen et al., 1997b.. The value of ␦ is model dependent, with typical values of 0.3᎐1.1 depending on the model and whether it includes the effects of feedbacks ŽHansen et al., 1997b, 1997d.. This range corresponds to changes in predicted global temperatures from 1.5 to 4.5⬚C for a doubling of CO 2 Žor the equivalent contributions from other greenhouse species.. As discussed in more detail in the following sections, some anthropogenic emissions are expected to cause positive radiative forcings whereas others are negative. Even though the two may be equal in magnitude, giving a net radiative forcing of zero, this does not mean that there will be no effects on climate. For example, negative radiative forcing caused by aerosol particles is expected to occur primarily over continents whereas the positive radiative forcing due to many greenhouse gases occurs globally. In addition, there are temporal differences, with the lifetime of particles typically being of the order of a week while those of many greenhouse gases are of the order of centuries. These differences can lead to impacts on climate, despite a net radiative forcing of zero. As a result of such considerations, radiative forcing is commonly used primarily for comparing the relative potential importance of various gases and particles on climate. b. Absolute and Relati© e Global Warming Potentials As was the case for ozone depletion potentials Žsee Chapter 13.B., the effects of greenhouse gases depend not only on the emissions but also on their lifetimes in the atmosphere ŽKo et al., 1993.. Global warming potentials ŽGWP. express the time-integrated radiative forcing due to the instantaneous emission of a fixed amount Žusually 1 kg. of the gas of interest. Thus, a time-scale horizon ŽTH. that will be considered in assessing the radiative effects of the gas must be specified. Both absolute and relative GWPs have been put forth, where these are defined as follows: Absolute Absolute GWP s



agas w gas x t dt

Ž S.




Caldeira and Kasting Ž1993. show that these two factors largely cancel each other so that using CO 2 as the reference gas is still useful. For the alternatives and proposed replacements for CFCs, CFC-11 has been used in some cases as the reference compound. In interpreting the GWPs, the reader should take note of which compound has been used as the reference. Another index has been proposed as well, a forcing equi¨ alent index Ž FEI . ŽWigley, 1998., defined as


Relative GWP s

agas w gas x t dt


Ž T.



aref w ref x t dt

In Eqs. ŽS. and ŽT., wgasx and wrefx represent the timedependent concentrations of the gas of interest and the reference gas, respectively, which are assumed to decay with characteristic lifetimes or response times after the instantaneous injection of the pulse, and a x Žunits of W my2 per ppb or ppm. is the radiative forcing of the gas or reference per unit increase in their atmospheric concentrations. The value of a x is assumed to be time-independent. The reference gas often used for relative GWPs is CO 2 because it is the major greenhouse gas. As we have seen earlier, the cycling of CO 2 throughout the earth system is complex, occurs with different response times, and is not thoroughly understood in a quantitative manner at present. As a result, uncertainties in how it decays will be translated into uncertainties in the relative GWPs of other greenhouse gases. There may, however, be some ‘‘cancellation of errors.’’ For example, the concentration of atmospheric CO 2 Žwrefx t in Eq. ŽT.. depends in a nonlinear fashion on the amount of total dissolved inorganic carbon in the ocean surface layer because of the equilibria with water Žsee Chapter 8.B. so that relatively less atmospheric CO 2 can be taken up by the oceans as its atmospheric concentrations increase. This would leave relatively more CO 2 in the atmosphere, increasing its greenhouse effect. On the other hand, since the strongest infrared absorption bands of CO 2 are already saturated Žvide supra., the radiative forcing Ž aCO 2 in Eq. ŽT.. decreases as its concentrations increase.

FEI gas Ž t . s

⌬ ECO 2Ž t . ⌬ Egas Ž t .


where ⌬ EgasŽ t . is the emissions reduction in the gas of interest calculated year-by-year that is needed to give the same change in radiative forcing as changes in CO 2 emissions, ⌬ ECO 2Ž t .. Table 14.4 summarizes the estimated total direct radiative forcing calculated for the period from preindustrial times to 1992 for CO 2 , CH 4 , N2 O, and O 3 ŽIPCC, 1996.. The estimate for CH 4 includes the effects due to its impacts on tropospheric ozone levels or on stratospheric water vapor, both of which are generated during the oxidation of methane. That shown for O 3 is based on the assumption that its concentration increased from 25 to 50 ppb over the Northern Hemisphere. The total radiative forcing due to the increase in these four gases from preindustrial times to the present is estimated to be 2.57 W my2 . Also shown are the relative global warming potentials, using CO 2 as the reference and for the two time horizons of 20 and 100 years, respectively ŽIPCC, 1996.. The apparently disproportionate effects of CH 4 , N2 O, and O 3 relative to CO 2 are due to the fact that CO 2 was present from natural processes in large concentrations even in preindustrial times and is such a strong

TABLE 14.4 Direct Radiative Forcings and Global Warming Potential for the Major Greenhouse Gases a Relative to CO 2 Concentration Gas



CO 2 CH 4 N2 O O3 Total

278 ppm 0.7 ppm 275 ppb 10᎐20 ppb

356 ppm 1.71 ppm 311 ppb 30᎐50 ppb


Total radiative forcing c (W m I2 ) 1.56 0.47 f 0.14 0.4 b 2.57

Relative global warming potential time horizon d 20 years

100 years

1 56 e 280 ᎏ

1 21e 310 ᎏ

Adapted from IPCC Ž1996.. Assuming an increase from 25 to 50 ppb throughout the Northern Hemisphere; see discussion for range and regional effects. c Change from preindustrial times to 1992. d Referenced to CO 2 and assuming CO 2 concentration is constant at mid-1990s levels. e Includes estimated effects of CH 4 on production of tropospheric O 3 and stratospheric H 2 O. f Does not include indirect effects on tropospheric O 3 and stratospheric H 2 O. b



FIGURE 14.23 Global warming potentials ŽGWP. relative to CO 2 for N2 O, C 2 F6 , HFC-134a, and HCFC-225ca as a function of time Žadapted from IPCC, 1996..

absorber of infrared radiation that its absorption is already saturated at the peak of the absorption bands. As discussed earlier, the absorption due to CO 2 thus depends on the logarithm of its concentration, whereas it depends on the square root of the concentrations of the other three gases. The importance of the time horizon over which the radiative forcing is considered is further illustrated by the data in Fig. 14.23, which shows the GWPs relative to CO 2 for N2 O, C 2 F6 , HFC-134a ŽCH 2 FCF3 ., and HCFC-225ca ŽCF3 CF2 CHCl 2 .. The lifetime of HCFC225ca at ;2.5 years is much shorter than for the other gases, and as a result, the GWP decreases relatively rapidly while that of N2 O, with a lifetime of ;120

years, increases over the same time frame. Thus, while the GWP of N2 O is initially an order of magnitude less than that for HCFC-225ca, it exceeds it after about four decades. Table 14.5 shows the direct radiative forcing calculated for some CFCs, HCFC-22, and some other chlorine-containing gases due to the increase in their atmospheric concentrations from preindustrial times Žwhen the concentrations of most of them were zero. to 1992. The direct radiative forcing due to the two most commonly used CFCs in the past, CFC-11 and CFC-12 Žsee Chapter 12.C.1., is approximately 8% of the total radiative forcing due to CO 2 , CH 4 , N2 O, and O 3 ŽTable 14.4.. However, the chlorinated and brominated gases exert not only direct radiative effects through their absorption of infrared radiation but also, as discussed in Section B.2, indirect effects through their destruction of stratospheric ozone Žsee Chapter 12.. Decreases in stratospheric ozone have three effects: Ž1. less downward radiation across the tropopause from the lower O 3 concentrations; Ž2. lowered stratospheric temperatures, which also decrease the downward radiation through a decrease in the population of the emitting excited states Žsee Eq. ŽA..; and Ž3. an increase in the solar UV radiation reaching the earth’s surface Že.g., see Ramaswamy et al., 1996.. The net effect is a decrease in the downward radiative flux at the tropopause, i.e., a negative radiative forcing Že.g., see Ramaswamy et al., 1992.. For example, Fig. 14.24 shows one estimate of the changes in the adjusted radiative forcing due to combined changes in total stratospheric and tropospheric O 3 compared to those to the other major greenhouse gases for the period from 1979 to 1994 ŽHansen et al., 1997b.. The negative radiative forcing associated with the destruction of stratospheric ozone

TABLE 14.5 Direct Radiative Forcings for CFCs, HCFC-22, and Other Chlorine-Containing Gases a



Direct radiative forcing (W m I2 )

0 0 0 0 0 0 ;600 0 0

268 503 82 20 -10 100 ;600 132 135

0.06 b 0.14 0.02 0.007 -0.003 0.02 0 0.01 0.007

Concentration (ppt) Gas CFC-11 ŽCCl 3 F. CFC-12 ŽCCl 2 F2 . CFC-113 ŽCCl 2 FCClF2 . CFC-114 ŽCClF2 CClF2 . CFC-115 ŽCF3 CClF2 . HCFC-22 ŽCHClF2 . CH 3 Cl CCl 4 CH 3 CCl 3 a

Adapted from IPCC Ž1996.. Christidis et al. Ž1997. suggest this is too small by ;30% based on their new infrared absorption cross section measurements. b



FIGURE 14.24 Calculated changes in adjusted radiative forcings from 1979 to 1994 due to changes in O 3 Žboth stratospheric and tropospheric based on SAGE data. and other greenhouse gases Žadapted from Hansen et al., 1997b..

counterbalances about half of the total positive forcing due to the other greenhouse gases. It should be noted that this estimated change in forcing due to the change in O 3 is quite sensitive both to the change in the vertical profile and to the change in the temperature

profile, which, of course, are related Že.g., see Zhong et al., 1996; and Shine et al., 1998.. For example, using the changes in O 3 derived using a different data set ŽSAGErTOMS data. gives a predicted net change in radiative forcing due to O 3 of y0.2 W my2 compared to y0.28 W my2 shown in Fig. 14.24. In addition, as seen later when all known anthropogenic perturbations are considered together, other estimates show the positive forcing by tropospheric ozone being larger than the negative forcing due to stratospheric ozone. The chlorinated and brominated compounds, therefore, are expected to exert not only a direct positive radiative forcing Žleading to warming. by their interaction with terrestrial infrared radiation in the usual manner of a greenhouse gas but also an indirect negative radiative forcing Žleading to cooling. due to changes in stratospheric ozone which changes the radiative flux at the tropopause. The net effect is thus expected to be less than that predicted from the direct radiative forcings in Table 14.5, which only take into account the direct greenhouse gas effect. Daniel et al. Ž1995., for example, have estimated the global warming potentials for a variety of chlorine- and bromine-containing gases involved in stratospheric ozone destruction, taking into account the indirect negative forcing as well as the direct positive forcing. Table 14.6 summarizes both the direct GWPs and the net values where the indirect, negative forcing due to ozone destruction has been taken into account. There are a number of simplifying assumptions that have been made in such calculations, but they are illustrative of the importance of considering the indirect as well as

TABLE 14.6 Direct and Net Global Warming Potentials for Some Chlorine- and Bromine-Containing Gases and Halons Relative to CO 2a Time horizon 2010







CFC-11 ŽCCl 3 F. CFC-12 ŽCCl 2 F2 . CFC-113 ŽCCl 2 FCClF2 . HCFC-22 ŽCHClF2 . CH 3 Cl CCl 4 CH 3 CCl 3 Halon-1301 ŽCF3 Br. Halon-1211 ŽCF2 ClBr. CH 3 Br

4360 6930 4390 3740 26.6 b 1730 q320 q5420 ᎏ 12.8 b

1900 5720 2940 3430 ᎏ y1290 y670 y19,740 y18,960 y4280

3170 6750 4020 1330 8.1b q1150 q90 q4460 ᎏ 3.9 b

1120 5540 2630 1200 ᎏ y1210 y200 y18,140 y11,050 y1190


Except where noted, adapted from Daniel et al. Ž1995.. Calculations based on a bromine enhancement factor of ␣ s 40 globally Žsee Chapter 12.D for discussion of ␣ . and assuming phase-out of emissions as scheduled in the Copenhagen amendments to the Montreal Protocol Žsee Chapter 13.A for a description of these.; note that WMO Ž1999. recommends ␣ s 60. b From Grossman et al. Ž1997..



the direct effects of anthropogenic emissions. For example, as discussed in Chapter 12.D, atom for atom, bromine is much more effective in destroying stratospheric ozone than is chlorine. This leads to large changes in the predicted GWPs when the indirect effects of ozone destruction are taken into account. For example, the relative GWP for Halon-1301 ŽCF3 Br. changes from net heating for the direct effect to net cooling when the indirect effect of ozone destruction is taken into account. Figure 14.25 illustrates the estimated relative contributions to the direct heating effect and to the indirect cooling effect in 1990 and 2040, respectively ŽDaniel et al., 1995.. The enhanced impact of the halons on cooling is particularly apparent. The net contribution of halocarbons through about 2080 is still estimated by Daniel et al. Ž1995. to be significant, however, in the range of 0.15᎐0.25 W my2 . Such calculations also illustrate that the net effects of ozone-depleting gases on climate should change as a function of time due to changes in the chemistry. For example, Daniel et al. Ž1995. assume in their calculations that the stratospheric ozone loss prior to 1980 was negligible. As a result, the contribution of CFCs to radiative forcing until 1980 was estimated to be positive, in the range q0.05 to q0.10 W my2 per decade. This rate of increase in radiative forcing should then have fallen as the indirect negative radiative forcing from ozone destruction came into play. Figure 14.26

FIGURE 14.26 Estimated direct tropospheric heating Žq. effects of chlorine- and bromine-containing gases and indirect cooling Žy. effects due to their destruction of stratospheric ozone, assuming a bromine enhancement factor of ␣ s 40 Žfrom Daniel et al., 1995..

illustrates their calculated temporal changes due to these two effects. As stratospheric ozone depletion becomes less severe in the future, the indirect, negative radiative forcing that has been in part counterbalancing the positive forcing will decrease; as a result, the net change in radiative forcing in the coming decades due to halocarbons is expected to be more steep than if their only effect was through direct positive radiative forcing, leading to heating ŽSolomon and Daniel, 1996; Hansen et al., 1997b.. As discussed in Chapter 13, the short- and long-term replacements for CFCs and halons in use or contemplated for use at the present time are typically compounds containing hydrogen Žto shorten their tropospheric lifetimes, hence decreasing the amounts that reach the stratosphere. andror fluorine Žwhich does not participate in stratospheric ozone destruction to a significant extent.. One concern with respect to these substitutes is their potential effect on climate since they are also greenhouse gases. Table 14.7 summarizes direct radiative forcings calculated for some of these compounds, as well as relative GWPs using either CFC-11 or CO 2 as the reference compound. Clearly, their contributions to radiative forcing may be significant and hence are taken into account in assessing the overall impacts of CFC replacements.

C. AEROSOL PARTICLES, ATMOSPHERIC RADIATION, AND CLIMATE CHANGE FIGURE 14.25 Partitioning of estimated direct tropospheric heating effect of stratospheric ozone-depleting gases in the years 1990 and 2040 and of the indirect cooling effect due to stratospheric ozone depletion Žfrom Daniel et al., 1995..

Aerosol particles have been thought for many years to play a role in global climate. For example, as discussed in Section E.1, while the mean global surface



TABLE 14.7 Direct Radiative Forcings per Unit Concentration for Some Fluorinated Compounds a and Relative Global Warming Potentials Taking GWP(CFC-11) = 1.0 b, c or GWP (CO 2 ) = 1.0

Gas CFC-11 ŽCFCl 3 . HFC-134a ŽCH 2 FCF3 . HFC-134 ŽCF2 HCF2 H. HFC-125 ŽCHF2 CF3 . HFC-152a ŽCH 3 CHF2 . HFC-23 ŽCHF3 . HFC-32 ŽCH 2 F2 . HFC-41 ŽCH 3 F. HFC-227ea ŽC 3 HF7 . HFC-245ca ŽC 3 H 3 F5 . HFC-236fa ŽCF3 CH 2 CF3 . HFC-236ea ŽCF3 CHFCHF2 . CF4 C 2 F6 C 3 F8 SF6

Radiative forcing per ppb (W m I2 ppb I1 )

Time horizon GWP relative to CFC-11

GWP relative to CO 2

20 years

100 years

20 years

100 years

d 0.17 0.18 0.20 0.11 0.18 0.11 0.02 0.26 0.20 0.24

Ž1.0. 0.58 0.50 c 0.99 0.093 2.0 0.47 0.10 0.81

Ž1.0. 0.29 0.22 c 0.83 0.036 3.2 0.18 0.038 0.73

0.10 f 0.22 f 0.23 f 0.64

0.88 f 1.26 f 0.97 f

1.70 f 2.37 f 1.88 f

Ž1.0. 3400 2900 4600 460 9100 2100 490 4300 1800 5610 e 2200 e 3660 f 5210 f 4040 g 16300

Ž1.0. 1300 1000 2800 140 11700 650 150 2900 560 6940 e 710 e 4690 f 6650 f 5150 f 23900

a Adapted from IPCC Ž1996.. For recent computations that are in reasonable agreement with these values, see Papasavva et al. Ž1997.. b GWPs relative to CFC-11 s 1.0 from Pinock et al. Ž1995. except where indicated. GWPs relative to CO 2 from IPCC Ž1996. assuming current CO 2 concentrations. c From Christidis et al. Ž1997. relative to CFC-11 s 1.0. d Reported as 0.22 W my2 in IPCC Ž1996.. Christidis et al. Ž1997. report a revised value of 0.285 W my2 ppby1. e From Gierczak et al. Ž1996.. f From Roehl et al. Ž1995..

temperature has increased by ;0.3᎐0.6⬚C over the past century, the period from 1940 to the mid-1970s showed no such trend ŽIPCC, 1996.. Indeed, this period was characterized by cooler than normal temperatures, which has often been qualitatively attributed to the scattering of incoming solar radiation by pollutionderived aerosol particles. In addition, as discussed in Chapter 9.A.4, elemental carbon in the particles, and perhaps some organics as well, also absorb solar radiation. A third effect discussed herein is the absorption and emission of long-wavelength infrared radiation by the inorganics found in mineral dust. These are the so-called direct effects of aerosol particles on global climate. In addition, aerosol particles have indirect effects. The most important of these is their effect on cloud properties, since clouds obviously also have major effects on climate. In addition, since heterogeneous chemistry can occur on aerosol particles Žsee Chapter 5., it is possible that such chemistry can alter the concentrations of other contributors to the climate system, such as the greenhouse gases. One example is the formation of N2 O from reactions of HONO on the surface of aerosol particles Žsee Chapter 7.C..

As will be evident from the following, this is an area in which there are a number of uncertainties and which at the time of writing is a very active area of research. For reviews, see Penner et al. Ž1994., Schwartz Ž1994., Andreae Ž1995., Charlson and Heintzenberg Ž1995., IPCC Ž1996., National Research Council Ž1996., and Andreae and Crutzen Ž1997..

1. Direct Effects a. Scattering of Solar Radiation One of the important properties of aerosol particles is their ability to scatter light. The diameter of many airborne particles is approximately that of the wavelength of visible light, so that Mie scattering occurs. ŽAs the particles become smaller, light scattering approaches the Rayleigh limit.. As discussed in Chapter 9.A.4, the intensity of Mie light scattering is a complex function of the wavelength of the incident light, the size and composition of the particle, and the scattering angle. However, for spherical particles of known composition and size, Mie theory can be used to predict the fraction of incident light that is scattered in various directions. One can therefore estimate how much light



is scattered back in the upward direction into space, which is the important parameter for aerosol particle direct scattering effects on global climate. This cooling effect due to direct scattering of light back out to space has been dubbed the ‘‘whitehouse effect’’ ŽSchwartz, 1996.. The net effect of such scattering depends on the underlying surface ŽAndreae et al., 1995; Haywood and Shine, 1997.. If the underlying surface is dark Že.g., oceans., backscattering by aerosol particles increases the planetary albedo, leading to cooling. However, if the surface is already highly reflecting, with an albedo greater than about 0.5, e.g., snow, scattering by particles can lead to a decrease in net reflection, especially at small solar zenith angles. The reason for this is that part of the light reflected from the surface is backscattered by the aerosol particles. This light then travels back through part of the atmosphere, undergoing enhanced absorption by particles Žas well as gases., before being reflected as well as partly absorbed again at the surface. This increased opportunity for absorption due to multiple scattering results in a net decrease in the reflectance and therefore a positive value for the radiative forcing ŽHaywood and Shine, 1997.. However, given what is currently known about aerosols and typical surface albedos, backscattering is believed to predominate on a global basis ŽCharlson et al., 1992b; Haywood and Shine, 1997.. As seen in Chapter 9, aerosol particles can, indeed most often do, have many chemical components and a wide range of particle sizes. However, sulfate is one of the most common species found in particles worldwide, being generated not only from the oxidation of anthropogenically derived SO 2 but also from the oxidation of sulfur compounds produced in biological processes Žsee Chapter 8.. However, in terms of global climate change caused by human influences, estimation of the effects of sulfate particles generated from the SO 2 emissions, 90% of which occur in the industrialized Northern Hemisphere, has been of greatest interest. As we shall see, however, recent studies suggest that carbonaceous aerosol particles and soil dust may also be quite important Že.g., Tegen et al., 1997.. Over the past several decades, there have been a number of estimates of scattering of light by tropospheric sulfate particles, with the aim of quantifying the magnitude of the expected cooling effect due to scattering of solar radiation compared to warming due to the greenhouse gases discussed earlier Že.g., see Rasool and Schneider, 1971.. One of the simplest approaches first taken was a box model in which the change in the short-wavelength radiative forcing averaged over the globe, ⌬ FR , is treated in terms of a

combination of effectsrterms ŽCharlson et al., 1991, 1992a; Penner et al., 1994.: 1 ⌬ FR s y F Ž1 y A c . 4 T

ž /

2 = 2T 2 Ž 1 y R s . ␤ ␣ RH f RH BSO 42y . Ž U.

The negative sign means that the direct effect of aerosols is expected to result in cooling of the troposphere, rather than heating as do the greenhouse gases discussed earlier. The terms in Eq. ŽU. are summarized in the following; also shown are the percentages of the total radiative forcing attributed to each factor by Charlson et al. Ž1992a. and Penner et al. Ž1994. and the uncertainty factor they assigned to each. ŽSubsequent estimates using global climate models will be discussed shortly.. 䢇

The incoming solar intensity: 1r4FT , where FT is the average total incoming solar light intensity per unit area normal to the direction of propagation outside the earth’s atmosphere, 1368 W my2 , and the 1r4 factor takes into account that the solar energy is spread over the entire surface of the earth Žsee Fig. 14.2a.. The fraction of the earth that is not covered by clouds, which is where increased scattering by aerosol particles is expected to be most important Žsee later, however, for a brief discussion of this assumption.: Ž1 y A c . s 0.39 with an uncertainty of a factor of 1.1, where A c is the fraction covered by clouds. The fraction of incident radiation that is scattered upward Ž ␤ ., the fraction ŽT . of light transmitted through the atmosphere above the aerosol layer, and the surface albedo Ž R s , which is the fraction of incident light that is reflected at the earth’s surface ., including multiple reflections between the surface and the aerosol layer: 2T 2 Ž1 y R s . 2 ␤ . Charlson et al. Ž1992a. and Penner et al. Ž1994. estimated T 2 s 0.58 with an uncertainty factor of 1.4, Ž1 y R s . 2 s 0.72 with an uncertainty factor of 1.2, and ␤ s 0.3 with an uncertainty factor of 1.3. A model intercomparison study by Boucher et al. Ž1998. shows that the linearity in Ž1 y R s . 2 holds up to R s ; 0.2᎐0.4; above this, model predictions for the dependence on this term diverge significantly. The light scattering efficiency of sulfate at a particular relative humidity ŽRH. chosen as a reference: ␣ RH ; this was estimated as 4.8 = 10 2 m2 moly1 with an uncertainty of 1.4. The increase in scattering due to an increase in the RH above the reference value: f RH ; this was estimated as 1.7 with an uncertainty factor of 1.2. The amount of atmospheric aerosol sulfate, expressed as the column burden: BSO 42y. This term can


be expressed as a combination of the source strength of SO 2 emitted and confined to an area A, QSO 2 in moles of S per year, the fraction that is oxidized to sulfate particles, Ysulfate , and the lifetime for sulfate particles in the atmosphere, ␶sulfate : BSO 42ys QSO 2 Ysulfate␶sulfaterA.

Ž V.

In the Charlson et al. Ž1992a. and Penner et al. Ž1994. estimates, BSO 2ys 3.3 = 10 5 mol my2 , with 4 uncertainties in QSO 2 , Ysulfate , and ␶sulfate of 1.15, 1.5, and 1.5, respectively. ŽNote, however, that some subsequent studies have suggested that direct radiative forcing may vary nonlinearly with the sulfate concentration due to chemical interactions with other particle constituents; e.g., see West et al., 1998.. The net results of such calculations vary depending on the input assumptions. For example, use of the inputs cited above for each of the terms gives ⌬ FR s y0.9 W my2 , with a possible range from y0.4 to y2 W my2 . However, application of different models gives different absolute values for the radiative forcing. For example, Penner et al. Ž1998. calculate global average values from y0.55 to y0.81 W my2 , depending on the assumptions regarding the effects of relative humidity. Boucher et al. Ž1998. compared the results for direct radiative forcing by sulfate particles predicted by 12 different groups using 15 different models. Overall, the predicted forcings for various input assumptions of particle radii, surface albedo, and solar zenith angle were similar. The largest differences were for high surface albedos; the range of predicted globally averaged direct forcing by particles with 0.17-␮ m radius over a surface with albedo 0.05᎐0.15 was 27% Žexpressed as a percentage of the median value.. The largest uncertainties are associated with the aerosol radiative properties ␤ , ␣ RH , and f RH , which is particularly sensitive to the treament of the dependence on relative humidity Žvide infra., and with the amount of sulfate Ži.e., in BSO 42y . available to scatter light. Other sensitivity analyses also suggest that uncertainties in Ysulfate and ␶sulfate , and hence in BSO 42y, as well as in the relative humidity factor are large sources of uncertainty Že.g., see Pan et al., 1997.. Not included in Eq. ŽU. is the important fact that sulfate is not distributed evenly over the globe but is concentrated in urbanrindustrialized areas of the Northern Hemisphere and in downwind regions. As a result, there is a great deal of spatial as well as temporal variability in the direct forcing expected from sulfate aerosol Žsee later discussion.. In addition, it should be noted that there are correlations between some of the terms in Eq. ŽU., so they do not all vary independently.


Recall from Chapter 9.A.4 that Mie scattering of light favors the forward direction and depends on particle size and composition and the wavelength of light. The fraction of light that is scattered in an upward direction relative to the earth’s surface is therefore expected to be a strong function of the direction of the incident light, i.e., the solar zenith angle. This is illustrated in Fig. 14.27, which shows calculated values of ␤ as a function of the solar zenith angle and particle radius r for light at 550 nm. As expected, the fraction of light scattered upward is smallest at a small zenith angle Ži.e., overhead sun. and highest Ž;50%. for the sun on the horizon. It is also high for the smallest particles. It should also be noted that most calculations to date of the backscattered fraction Ž ␤ . assume spherical aerosol particles, which may not be the case, particularly at low RH, where dry particles have shapes determined by their crystal structures. For example, Pilinis and Li Ž1998. have shown that this assumption can cause significant underestimation Žas much as a factor of three. of the calculated direct aerosol forcing, particularly at small solar zenith angles. Because the index of refraction enters into Mie scattering calculations, one would expect that the light scattering efficiency at a particular RH, ␣ RH , would depend on the particle composition as well as radius. As discussed in Chapter 9.A.4, this has indeed been shown to be the case. However, given that tropospheric particles are a complex mixture of inorganic and organic compounds, it is not clear that the linear and independent relationship between light scattering and the mass scattering coefficient for each species described by Eq. ŽEE. in Chapter 9 holds ŽSloane, 1986; White, 1986; Hegg et al., 1993, 1994.. Water uptake affects light scattering by particles due to changing the particle size and the index of refraction. This has been discussed in detail in Chapter 9.B.4 and the reader is referred to that section for more

FIGURE 14.27 Calculated fraction of 550-nm light scattered upward Ž ␤ . as a function of solar zenith angle Ž ␪ . and particle radius Ž ␮ m.. The refractive index of the particles is 1.4 Žadapted from Nemesure et al., 1995..



details. The important point is that the combination Ž ␣ RH f RH . is a very sensitive function of RH, and indeed, this appears to be the most important aerosol parameter in direct forcing ŽNemesure et al., 1995; Pilinis et al., 1995.. Figure 14.28 shows one calculation of ␣ RH f RH as a function of RH for ŽNH 4 . 2 SO4 , which is a common form of sulfate in tropospheric aerosols ŽNemesure et al., 1995.. The effect of relative humidity is large, increasing from ;1.5 m2 per g of sulfate for dry particles to ;50 m2 per g at 97% RH at a wavelength of 600 nm. Finally, as expected from Mie theory and demonstrated in Fig. 14.28, the light scattering efficiency varies widely over the wavelength region of interest in the troposphere. The total emissions of SO 2 due to anthropogenic processes are believed to be relatively well known Žsee Chapter 2.A.4.. A great deal is also known about the processes oxidizing tropospheric SO 2 to sulfate Žsee Chapter 8.. In brief, while OH oxidizes SO 2 in the gas phase, oxidation in the liquid phase of fogs and clouds is generally more important. In the latter case, O 3 is a major oxidant at high pH values, but as the droplet becomes acidified, this slows down. Over most of the pH range characteristic of particles in the troposphere, H 2 O 2 is an effective oxidant and hence is believed to be the major contributor to sulfate formation. SŽIV. may also be oxidized in the marine boundary layer by HOCl and HOBr ŽVogt et al., 1996.. This means that the fraction of SO 2 forming sulfate will depend on such factors as the availability of liquid phase in the form of fogs and clouds, as well as the availability of oxidants such as H 2 O 2 , which is itself a complex function of the VOC᎐NO x chemistry discussed throughout this book. In addition, a large contributor to the removal of sulfate is wet deposition, which, due to its spatial and temporal variability, results in a corresponding variability in the tropospheric lifetime of sulfate. Because of all of these factors, there is significant uncertainty

FIGURE 14.28 Calculated mass scattering efficiency term Ž ␣ RH f RH . as a function of wavelength for 0, 40, and 97% RH. Particle dry radius taken to be 0.096-␮ m ŽNH 4 . 2 SO4 Žadapted from Nemesure et al., 1995..

associated with both Ysulfate and ␶sulfate , and hence in the column burden of sulfate used in such calculations. Finally, the term Ž1 y A c . in Eq. ŽU., where A c is the fraction of the earth’s surface covered by clouds, contains the implicit assumption that the direct scattering by aerosol particles is only significant in cloud-free regions. It is not clear, however, that this is the case. For example, while Haywood et al. Ž1997a. and Haywood and Shine Ž1997. report only a small contribution Ž5%. to aerosol particle direct forcing in cloudy regions, Boucher and Anderson Ž1995. report a much larger effect, with 22% of the total arising from cloudy areas. Other studies fall in between these values Že.g., see Haywood and Ramaswamy, 1998.. For example, one calculation using a 1-D model suggests that cirrus clouds above a layer of particles of ŽNH.4 SO4 can enhance the direct forcing at a solar zenith angle of 0⬚ from y2.0 to y3.3 W my2 for their assumed set of conditions ŽLiao and Seinfeld, 1998.. This is due to the fact that the overlying cloud scatters the incoming beam so that the light incident on the aerosol layer is effectively at larger solar zenith angles than 0⬚, leading to increased upward scattering Žsee the dependence of the upscattered fraction, ␤ , on solar zenith angle due to enhanced forward Mie scattering by particles, Fig. 14.27.. A similar enhancement is predicted for a cloud layer of 100-m thickness assumed to be at an altitude of 900 m and embedded in a layer of aerosol taken to be uniformly distributed from 0 to 5 km. However, for much thicker clouds Že.g., 1000 m., the direct aerosol forcing due to sulfate may be decreased due to partial blocking of the incoming solar light intensity by the cloud ŽLiao and Seinfeld, 1998.. In short, the approach summarized in Eq. ŽU. provided a useful first approach to establishing that direct scattering of light by sulfate particles could be important in counterbalancing the expected warming due to the increased greenhouse gases. However, given the large number of variables that enter into each term in this equation and their spatial and temporal variability, the development and application of more sophisticated models are clearly needed. For some examples of such models, see Kiehl and Briegleb Ž1993., Taylor and Penner Ž1994., Cox et al. Ž1995., Mitchell et al. Ž1995., Boucher and Anderson Ž1995., Meehl et al. Ž1996., C. C. Chuang et al. Ž1997., Haywood and Shine Ž1997., Haywood et al. Ž1997a, 1997b., van Dorland et al. Ž1997., Tegen et al. Ž1997., Schult et al. Ž1997., Haywood and Ramaswamy Ž1998., and Penner et al. Ž1998.. For example, Fig. 14.29 shows one calculation of direct radiative forcing using a global climate model, GCM ŽPenner et al., 1998.. Due to the preponderance of anthropogenic SO 2 emissions in the Northern Hemisphere, the direct radiative forcing due to sulfate



FIGURE 14.29 Calculated direct radiative forcing due to sulfate aerosol particles Žadapted from Penner et al., 1998..

aerosol is also predicted to occur in the Northern Hemisphere, particularly over the eastern United States, central and eastern Europe, and southeastern Asia. This model predicts a global average direct sulfate aerosol forcing of y1.2 W my2 in the Northern Hemisphere compared to only y0.26 W my2 in the Southern Hemisphere, where less than 10% of the anthropogenic sulfur emissions occur. Haywood and Ramaswamy Ž1998. predict very similar average values for the direct radiative forcing by sulfate, y1.4 W my2 in the Northern Hemisphere and y0.24 W my2 in the Southern Hemisphere, although there are some quantitative differences in the specific geographical dependencies. It should be noted that the magnitude of the predicted forcing is quite sensitive to treatment of relative humidity ŽRH. in the model because of the effects on particle size and optical properties Že.g., Haywood and Shine, 1995; Haywood and Ramaswamy, 1998; Ghan and Easter, 1998; Haywood et al., 1998a; Penner et al., 1998.. For example, in the calculations by Penner et al. Ž1998., when the particle properties were held fixed at the values for 90% RH for 90᎐99% RH, the predicted direct radiative forcing for sulfate particles decreased from y1.18 W my2 to y0.88 W my2 for the Northern Hemisphere and from y0.81 to y0.55 W my2 globally. Volcanic eruptions provide one test of the relationship between light scattering by sulfate particles and the resulting change in temperature, since they generate large concentrations of sulfate aerosol in the lower stratosphere and upper troposphere. These aerosol

particles cause tropospheric cooling by backscattering solar radiation out to space. In principle, they can also cause tropospheric and stratospheric warming by absorbing and reemitting the long-wavength terrestrial infrared ŽMcCormick et al., 1995; Robock and Mao, 1995.; calculations show that the warming increases substantially with particle size and should be sufficiently large to counterbalance the cooling from light scattering at area-weighted mean radii larger than ;2 ␮ m ŽLacis et al., 1992.. However, since the sulfate particles formed by gas-to-particle conversion of SO 2 are sub-␮ m Že.g., see Fig. 12.28., the backscattering effect leading to cooling generally predominates in the troposphere. Model calculations predicted that surface cooling of ;0.5⬚C should have occurred after the Mount Pinatubo eruption due to the scattering of solar radiation back out to space, and indeed, this is what was observed Že.g., Hansen et al., 1992; Dutton and Christy, 1992; Minnis et al., 1993; Lacis and Mishchenko, 1995; Robock and Mao, 1995; McCormick et al., 1995; Parker et al., 1996; Saxena and Yu, 1998.. Figure 14.30, for example, shows the predicted and measured monthly surface temperatures obtained prior to and following the eruption, demonstrating excellent agreement between observations and the model; the observed stratospheric warming is also in excellent agreement with model calculations Že.g., Lacis and Mishchenko, 1995; Angell, 1997.. Temperature trends on a smaller, regional scale have also been attributed to the effects of such volcanic eruptions, although obtaining statistically robust results is difficult in these cases due to natural variability Že.g., Saxena et al.,



FIGURE 14.30 Measured Ž ᎏ . and model-predicted Ž ⭈⭈⭈ . change in monthly mean temperatures Ža. at the earth’s surface and Žb. in the stratosphere Žobservations at 30 mbar and 10⬚S, model results for the 10- to 70-mbar layer from 8⬚S to 16⬚S. Žadapted from Lacis and Mishchenko, 1995..

1997.. Additional discussion of the effects of volcanic eruptions on surface temperatures is found in Section D.3. Finally, while sulfate is a major component of aerosol particles, particularly in the Northern Hemisphere where fossil fuel burning is extensive, it is certainly not the sole particle component. As discussed in detail in Chapter 9, nitrate, organics and elemental carbon, inorganic soil elements, and sea salt components are all found in tropospheric particles to varying extents, depending on the region. Some regional sources such as biomass burning may also have quite widespread and global impacts Že.g., Kaufman et al., 1991; Penner et al., 1992.. These other sources will also scatter light and contribute to direct aerosol forcing Že.g., see Andreae et al., 1995.. While less attention has been paid to assessing the contribution of these other components ŽPenner et al., 1994., there are increasing indications that these may prove to be quite important in direct forcing by aerosol particles Že.g., see Andreae et al., 1995; Penner et al., 1998.. For example, Fig. 14.31 shows one set of calculations of direct radiative forcing by sulfate particles as well as those from biomass burning and from fossil fuel combustion in the Northern Hemisphere, Southern Hemisphere, and the global average, respectively ŽPenner et al., 1998.. In this case, the fossil fuel particles were assumed to contain both black carbon, which absorbs radiation and hence has a positive radiative forcing Žsee following section., and organic carbon, which scatters light Žnegative radiative forcing.. The biomass particles were assumed to take up water as if they contained 30% ŽNH 4 . 2 SO4 by mass and to scatter light. Figure 14.31 illustrates that the direct radiative forcing

FIGURE 14.31 Calculated direct radiative forcing by sulfate, biomass, and fossil fuel black carbon ŽBC. q organic carbon ŽOC. particles in the Ža. Northern Hemisphere, Žb. Southern Hemisphere, and Žc. global average Žadapted from Penner et al., 1998..

by biomass particles is predicted to be large in the Southern Hemisphere and negative because of their contribution to light scattering. However, over some regions, e.g., the Sahara desert, such particles are predicted to lead to positive radiative forcing because they are over an already highly reflecting surface Žsee the earlier discussion.. ŽNote that the size distribution assumed for the particles affects the absolute value pre-


dicted for radiative forcing; for example, input of a different size distribution than was used for Fig. 14.31 for the biomass particles gave a minimum direct radiative forcing in the Southern Hemisphere of y0.52 W my2 , significantly more negative than shown in Fig. 14.31b.. The contribution of fossil fuel black and organic carbon leads to a positive radiative forcing because of the absorption of solar radiation by black carbon Žvide infra.. Organic constituents of particles are also likely to make a contribution to light scattering. For example, Li et al. Ž1998. studied aerosol particles over the east coast of Canada and found that unidentified species, likely organics, account for a large fraction Ž;2r3. of the mass in the 0.1- to 1-␮ m range. In air masses with origins over the continental United States, these unidentified compounds were calculated to be responsible for 45᎐80% of the direct backscatter coefficient. There is also evidence of new particle formation likely involving organics from biological processes in coastal regions. For example, nucleation of new particles in the ultrafine size range Ž1.5᎐ 5 nm. has been observed in coastal regions at Mace Head, Ireland, and the Outer Hebrides and was correlated to solar radiation and low tide ŽO’Dowd et al., 1998.. This suggests that photochemical processes may lead to new organic particle formation and that the precursors may be biogenics from the shore regions. The contribution of carbonaceous components to tropospheric aerosols off the east coast of the United States has been reported to vary from ;10% at low altitudes to )90% of the total aerosol mass at altitudes of ;3 km ŽNovakov et al., 1997.. These carbonaceous components include both organics and elemental carbon, the latter estimated to be ;10% of the total carbonaceous mass. The larger fraction at higher altitudes may reflect more rapid removal of the inorganic components such as sulfate and nitrate through wet deposition at the surface. These carbonaceous materials contributed 66 " 16% of the light scattering coefficient ŽHegg et al., 1997., i.e., were major contributors to the direct effect of aerosol particles. Figure 14.32 shows some typical contributions to the light scattering measured during these studies ŽHegg et al., 1997.. While the total scattering, shown here in terms of the aerosol optical depth, varied from one set of measurements to another, it can be seen that liquid water in the particles, carbonaceous compounds, and sulfate were consistently the largest contributors, and in that order. Light absorption was a minor contributor to the aerosol optical depth in these studies. Since these measurements were made close to the earth’s surface, the relative importance of organics when averaged over the troposphere may, however, be somewhat less.


FIGURE 14.32 Typical contributions of aerosol liquid water, carbonaceous compounds, and sulfate in the lower troposphere of the east coast of the United States to the total aerosol optical depth. The contribution of light absorption is also shown. The different bars represent different sets of measurements during different flights Žadapted from Hegg et al., 1997..

It should be noted that while the focus of most of these studies has been on the contributions of species that are believed to have changed over time, e.g., those associated with human activities, there are also contributions due to natural processes. The latter are presumably not changing with time, or at least not at a significant rate compared to those due to anthropogenic activities Žsee discussion in Section E, however.. For example, the carbonaceous aerosols measured by Novakov et al. Ž1997b. and Hegg et al. Ž1997. over the east coast of the United States may have been largely natural in origin Žsee Chapter 9.C.2. and hence not expected to show secular changes associated with anthropogenic activities. Similarly, Murphy et al. Ž1998. have measured the size-dependent composition of aerosol particles in the marine boundary layer in the Southern Ocean and shown that the contribution of sea salt particles to backscattering dwarfs that due to nonsea salt sulfate Žnss.. In addition, these particles comprised more than 50% of the cloud condensation nuclei, CCN Žvide infra.. ŽNote, however, that since the concentration of sea salt particles is highest in the marine boundary layer, this contribution to backscattering will not be characteristic of the global troposphere.. Sea salt particles are generated by natural processes, so that this contribution to backscattering has undoubtedly always been present and will not contribute to a secular trend. However, understanding its contribution, as well as those from other natural processes, in scattering of solar radiation is important to place the contribution of anthropogenically derived aerosol particles in perspective. In addition, as discussed later with respect to indirect effects of sea salt



particles, the effects of such natural aerosols often act in a synergistic manner with anthropogenic aerosol particles. In short, the direct effects of aerosol particles in terms of backscattering solar radiation out to space and hence leading to cooling are reasonably well understood qualitatively and provided the aerosol composition, concentrations, and size distribution are known, their contribution can be treated quantitatively as well. However, major uncertainties exist in our knowledge of the physical and chemical properties, as well as the geographical and temporal variations, of aerosol particles and it is these uncertainties that primarily limit the ability to accurately quantify the direct effects at present. b. Absorption of Solar Radiation Depending on their chemical composition, aerosol particles can not only scatter incoming solar radiation but, in some cases, absorb it as well. This absorbed energy is converted to heat, which can contribute to warming of the troposphere. In addition, since energy absorbed by such particles does not reach the surface but heats the atmosphere directly, changes in the lapse rate may result as well, and this can contribute to global change by altering atmospheric circulation patterns Že.g., see Penner et al., 1994; and Tegen et al., 1997.. Although sulfate aerosols do not appreciably absorb incoming solar radiation, elemental or black Žgraphitic. carbon particles do. In addition, mineral dust particles absorb in the visible, primarily due to the presence of iron compounds such as hematite ŽFe 2 O 3 . ŽPatterson, 1981; Sokolik et al., 1993.. As discussed in Chapter 9.A.4, some recent studies suggest there may be a previously unrecognized, but substantial, contribution of complex organics to light absorption ŽMalm et al., 1996.. However, since relatively little is known about the nature and concentrations of these complex organics, and the dust contribution to solar radiation absorption is generally assumed to be small compared to that of elemental carbon Žwhich may not be justified in some regions; e.g., see Patterson, 1981., we focus here on the contribution of elemental carbon. As described in Chapter 9 Žsee also Penner and Novakov, 1996., tropospheric particles containing carbon are often referred to as ‘‘carbonaceous aerosols.’’ The form of the carbon may be organic or elemental, the latter often being referred to as graphitic or black carbon due to its strong absorption of visible light. The expression given in Eq. ŽU. for direct radiative forcing by aerosol particles can be modified to include contri-

butions due to absorption ŽChylek ´ and Wong, 1995.: ⌬ FR s y


FT Ž 1 y A c . T 2

ž / 4

2 = Ž 1 y R s . 2 ␤␶scat y 4 R s␶abs ,


where ␶scat s ␣ RH f RH BSO 42y is the effective aerosol optical scattering depth and ␶abs is the optical depth due to absorption. For no absorption, ␶abs s 0 and Eq. ŽW. reduces to Eq. ŽU. as expected. Note that for absorption when scattering is small, ⌬ FR becomes positive, i.e., warming results. In short, carbonaceous aerosols can both scatter solar radiation, causing negative radiative forcing, and absorb light, leading to positive forcing. The net effect depends on the composition of the aerosols as well as their particle size and vertical distribution. Several groups have carried out detailed global 3-D model studies of carbonaceous aerosol particles and their effects on radiation balance Že.g., Liousse et al., 1996; Haywood and Ramaswamy, 1998; Penner et al., 1998.. For example, Figure 14.33 shows the model estimates by Penner et al. Ž1998. for direct radiative forcing due to aerosol particles from fossil fuel combustion. These were assumed to include both black carbon, which absorbs solar radiation, and organic carbon, which scatters it. The black carbon over Europe, Asia, and, to a lesser extent, the eastern United States leads to positive radiative forcing in these areas due to direct absorption of solar radiation. Similar results are predicted by other models, although the absolute values of the radiative forcing may differ somewhat, depending on the details of the particular model used and the emission inventory chosen for the black carbon Že.g., see Schult et al., 1997; Haywood and Ramaswamy, 1998; and Penner et al., 1998.. Figure 14.34 shows one model prediction for the net radiative forcing due to a combination of sulfate, biomass, and fossil fuel particles containing both black and organic carbonaceous compounds ŽPenner et al., 1998.. The net result is predicted negative radiative forcing over the industrialized areas of the Northern Hemisphere due to aerosol particles Žremember this does not include the positive forcing due to greenhouse gases.. Positive radiative forcing is predicted over the highly reflecting ice-covered surfaces at high latitudes and over a small portion of Asia. The predicted average values of radiative forcing were y0.65 to y1.07 W my2 for the Northern Hemisphere, y0.33 to y0.51 W my2 for the Southern Hemisphere, and y0.51 to y0.88 W my2 for the global average. Jayaraman et al. Ž1998. measured the aerosol optical depth, aerosol size distribution, and the solar flux close to the coast of India, over the Arabian Sea, and then



FIGURE 14.33 Calculated direct radiative forcing for the combination of black carbon and organic carbon from fossil fuel combustion. The numbers shown are the lower limits of the ranges included in each area. The boundaries of the regions are "10, "5, "2, "1, "0.5, "0.2, "0.1, and 0. Thus y5 represents the y5 to y10 W my2 region, y2 represents the y2 to y5 W my2 region, etc. Žadapted from Penner et al., 1998..

over the more remote Indian Ocean. The aerosol optical depth increased from 0.1 or less in the remote region to 0.2᎐0.4 over the Arabian Sea to values up to 0.5 close to the Indian coast. This paralleled the trend in aerosol mass concentrations, which varied from a few ␮ g my3 over the Indian Ocean to ;80 ␮ g my3 near the coast. The data indicated that the contribution of light absorption by aerosols near the coast was

larger than that over the remote ocean, suggesting a contribution from black carbon and perhaps other organics associated with anthropogenic activities. The presence of clouds can also affect the net light absorption by black carbon, indeed even more than for sulfate ŽHaywood and Shine, 1997; Haywood and Ramaswamy, 1998; Liao and Seinfeld, 1998.. For example, Liao and Seinfeld Ž1998. calculate that the net heating

FIGURE 14.34 Model-predicted net direct radiative forcing due to sulfate and carbonaceous particles. The numbers shown are the lower limits of the ranges included in each area. The boundaries of the regions are "10, "5, "2, "1, "0.5, "0.2, "0.1, and 0. Thus, y5 represents the y5 to y10 W my2 region, y2 represents the y2 to y5 W my2 region, etc. Žadapted from Penner et al., 1998..



effect can be enhanced by as much as a factor of three in the presence of a low, thick stratus cloud below the particles, due to the enhanced scattering of solar radiation back to the absorbing particles Žsee Chapter 3.C.2g.. On the other hand, a thick cirrus cloud above the black carbon reduces its heating effect for an overhead sun because of reduced transmission of the direct solar radiation. As a result of this sensitivity to the location of clouds relative to the carbon particles, knowing the vertical distribution of black carbon is important Že.g., Haywood and Shine, 1997.. There is another interaction of absorbing aerosol particles with clouds that has been proposed. Hansen et al. Ž1997b, 1997d. suggest there is a ‘‘semi-direct’’ effect in that warming caused by light absorption may reduce cloud cover and result in net warming. Their calculations suggest that at values of the single-scatter albedo Ždefined as the fraction of extinction that appears as scattered radiation. below about 0.85, this semi-direct effect predominates and heating, rather than cooling, results. In summary, aerosol particles containing compounds that can absorb solar radiation, such as elemental Žor black. carbon and possibly some organic compounds as well, can also contribute to direct radiative forcing. This absorption of solar radiation generally results in positive radiative forcing. This effect occurs simultaneously with scattering, which results in negative radiative forcing. It should be noted that such particles also affect the amount of solar radiation reaching the low troposphere that is available for photochemistry Že.g., see Haywood and Shine, 1997.. See Chapter 3.C.2f for a more detailed discussion of this issue.

However, this is not the case for airborne particles composed of crustal materials formed by erosion processes. As discussed in Chapter 9.C, mineral dust consists primarily of such crustal materials. Despite the fact that soil dust particles tend to be quite large, of the order of a micron and larger, they can be carried large distances. These particles not only scatter and absorb solar radiation but also absorb long-wavelength infrared emitted by the earth’s surface. Figure 14.35, for example, shows the real Ž n. and imaginary Ž k . parts of the index of refraction Ž␩ s n y ki . of three samples of dust collected in the Barbados, but thought to be transported from the Sahara Desert ŽVolz, 1973., from Afghanistan ŽSokolik et al., 1993., and from Whitehill, Texas, in the southwestern United States ŽPatterson, 1981., respectively ŽSokolik et al., 1998.. Regions of absorption in the infrared due to some common dust components are also shown: the asymmetric C᎐O stretch of carbonate in calcite near 7 ␮ m seen in the Texas dust, the asymmetric Si᎐O᎐Si

c. Absorption of Long-Wa© elength Infrared Species such as sulfate and black carbon can absorb the long-wavelength thermal infrared emitted by the earth’s surface, leading to positive radiative forcing. In principle, the same is true of other infrared-absorbing particle components such as nitrate, ammonium, formate, acetate, and oxalate for the bands that are not in the region of the spectrum that is already saturated ŽMarley et al., 1993.. It has been proposed that if these species andror sulfate are present at sufficiently high concentrations in particles, for example in or downwind of urban areas, they can contribute to radiative forcing by direct absorption of infrared on local to regional scales ŽMarley et al., 1993; Gaffney and Marley, 1998.. On a global basis, Haywood and Shine Ž1997. and Haywood et al. Ž1997a. estimate that the contribution of sulfate and black carbon to long-wavelength direct forcing is at least an order of magnitude less than that due to the scattering, and in the case of black carbon, absorption of solar radiation.

FIGURE 14.35 Ža. Real Ž n. and Žb. imaginary Ž k . parts of the index of refraction Ž␩ s n y ki . of some atmospheric dust samples from the Sahara collected in Barbados, Afghanistan, and Whitehill, Texas. Regions of strong absorption of some known common dust components are also shown Žadapted from Sokolik et al., 1998..


stretch in quartz near 9 ␮ m, and absorptions due to kaolinite wAl 2 Si 2 O5 ŽOH.4 x in the 8.5- to 12-␮ m region ŽSalisbury et al., 1992.. Many of these crustal materials also absorb in the 15- to 25-␮ m region; the major infrared absorption of hematite Ž ␣-Fe 2 O 3 . is also in this region. This absorption also leads to positive radiative forcing. Figure 14.36, for example, shows a model estimate of the contribution to radiative forcing by dust particles due to scattering of solar radiation and absorption of infrared radiation ŽTegen et al., 1996.. As expected, the effects are calculated to be the largest in areas having the highest dust, around the Arabian Sea and over the Atlantic Ocean off the coast of Africa. Figure 14.36 shows that the calculated effect of infrared absorption by mineral dust particles can be equal to or greater than that of scattering of solar radiation. Alpert et al. Ž1998. have proposed that the previously unrecognized contribution of dust particles to heating of the atmosphere was responsible for inaccuracies in weather prediction models over the eastern tropical North Atlantic Ocean. While dust particles are a natural component of the atmosphere, the amount of airborne dust is believed to have increased due to anthropogenic surface land modifications such as deforestation, cultivation, and shifts in vegetation Že.g., see Tegen and Fung, 1994, 1995; Tegen and Lacis, 1996; Li et al., 1996; Tegen et al., 1996; and Tegen et al., 1997.. These activities may be responsible for as much as half of the total airborne dust ŽTegen and Fung, 1995.. Based on their measurements of North African dust transported to the Barbados, Li et al. Ž1996. estimate that over a 10-year period, dust contributed about 56% of the total light scattering. Similarly, Tegen et al. Ž1997. estimate using a global transport model that scattering and absorption of light by submicron soil


dust were about as important as that due to carbonaceous aerosols and scattering by sulfate aerosols. Sokolik and Toon Ž1996. point out that on a regional basis, direct radiative forcing due to dust aerosols can significantly exceed that of sulfate. Light scattering by mineral dust has been shown to be relatively insensitive to the relative humidity ŽLi-Jones et al., 1998.. Because scattering and absorption contribute to radiative forcing in opposite directions ŽFig. 14.36., the positive radiative forcing and negative radiative forcing largely cancel at the top of the atmosphere. For example, for the model study shown in Fig. 14.36, the global mean net radiative forcing at the top of the atmosphere due to dust associated with disturbances from anthropogenic processes was only q0.09 W my2 . However, the net radiative effect of this dust at the surface is calculated to be negative, y1 W my2 , since both scattering and absorption reduce the sun intensity reaching the ground ŽTegen et al., 1996.. Finally, the absorption of infrared radiation by mineral dusts leads to direct heating of the atmosphere, which may alter atmospheric circulation processes ŽTegen et al., 1996.. In short, the combination of absorption and scattering of light by mineral dusts, combined with an increase in these due to anthropogenic activities, has the potential to contribute to climate change. However, many uncertainties need to be removed before these effects can be confidently quantified. For example, the infrared absorption depends on the composition of the dust and as seen in Fig. 14.35, this can be quite variable from location to location and even as a function of time from one source. This one effect alone can lead to a large variability in the predicted effects on radiative forcing ŽSokolik et al., 1998..

2. Indirect Effects of Aerosol Particles a. Clouds

FIGURE 14.36 Predicted mean radiative forcing at the top of the atmosphere due to mineral dust from scattering, absorption, and their total Žadapted from Tegen et al., 1996..

In addition to the direct effects on radiative forcing due to scattering and absorption of light, aerosol particles also have indirect effects, which may, in many instances, be more important than the direct radiative forcing. These indirect effects are based on the ability of some Žbut as we shall see, not all. aerosol particles to act as cloud condensation nuclei, CCN. This changes the number concentration of droplets in clouds and their size distribution, which can alter the precipitation rate. In addition, such changes in the cloud characteristics are believed to alter the lifetime and extent of the cloud Že.g., see Cess et al., 1997; and Lohmann and Feichter, 1997.. As discussed in more detail shortly, clouds decrease the incoming solar radiation by reflecting a significant amount back out to space Žthe



predominant effect., but high clouds can also lead to tropospheric warming through interaction with the longwave terrestrial thermal radiation. In addition, there are some data, which are presently controversial, suggesting that clouds absorb solar radiation directly to a larger extent than expected. If proven true, this can lead to thermal heating and effects on atmospheric circulation processes that are greater than have been understood to the present. Twomey suggested in 1974 that anthropogenic emissions could affect cloud properties and albedo, i.e., have an indirect effect on global climate. Attention was further drawn to such indirect effects in 1987 when Charlson, Lovelock, Andreae, and Warren proposed a series of feedbacks involving dimethyl sulfide emitted by phytoplankton in seawater, CCN, and clouds. Dimethyl sulfide ŽDMS. is known to be oxidized in part to sulfate Žsee Chapter 8.E.1., which acts as a source of CCN and hence affects cloud properties, including albedo. Thus, DMS and its oxidation products such as methanesulfonate have been shown in a number of studies to correlate with CCN ŽDurkee et al., 1991; Hegg et al., 1991a,b; Ayers and Gras, 1991; Berresheim et al., 1993; Quinn et al., 1993; Putaud et al., 1993; Lawrence, 1993; Andreae et al., 1995; Ayers et al., 1995.. The relationship between DMS and CCN may not be linear, however. For example, the model of Pandis et al. Ž1994. predicts that small DMS emission fluxes Ž-1.3 ␮ mol my2 dayy1 . do not lead to new CCN since much of the oxidation of SO 2 from DMS occurs in existing sea salt particles Že.g., see Sievering et al., 1992; Chameides and Stelso, 1992.. However, at fluxes ) 2.3 ␮ mol my2 dayy1 , new CCN are predicted by the model to be formed in an approximately linear relationship. Average DMS fluxes measured over the North Atlantic have been observed that span this range, from 1.2 to 12 ␮ mol my2 dayy1 , with peaks up to 39 ␮ mol my2 dayy1 ŽTarrason ´ et al., 1995.. Over the Southern Ocean, a range from 0.2 to 5 ␮ mol my2 dayy1 has been measured ŽAyers et al., 1995.. This suggests that the effect of DMS on CCN may vary geographically and seasonally. The seasonal cycle of CCN has also been shown to be correlated with that of cloud optical depth in one remote marine area ŽBoers et al., 1994., and the isotope composition of non-sea salt sulfate over remote regions of the southern Pacific Ocean has been shown to be consistent with a DMS source ŽCalhoun et al., 1991.. Based on such correlations, it is reasonable to assume that the Twomey proposal is applicable, i.e., that anthropogenic emissions of SO 2 and other species that form particles in the atmosphere may contribute to CCN and hence have indirect effects on climate.

Charlson et al. Ž1987. drew attention to the potential importance of feedback loops in which increased DMS emissions lead to increased sulfate CCN, increased clouds, and cloud albedo, followed by changes in the temperature and solar radiation at the surface. Changes in temperature and solar radiation might then alter DMS production, although whether in terms of increased or decreased emissions is uncertain. Despite many studies, the details and importance of such feedbacks remain to be elucidated Že.g., see Schwartz, 1988; Baker and Charlson, 1990; Lin et al., 1992; Hegg, 1990, 1993a, 1993b; Chameides and Stelson, 1992; Raes and Van Dingenen, 1992, 1995; Sievering et al., 1992; Lin et al., 1993a, 1993b; Russell et al., 1994; Pandis et al., 1994, 1995; Raes, 1995; Capaldo and Pandis, 1997; and Andreae and Crutzen, 1997.. For example, Bates and Quinn Ž1997. report that the DMS concentration in the surface seawater of the equatorial Pacific was relatively insensitive to changes in the properties of the atmosphere Že.g., cloud cover and precipitation. and oceans Že.g., sea surface temperature, mixed layer depths, and upwelling rates. associated with El Nino᎐Southern Os˜ cillation events. The following sections focus on the potential indirect effects of aerosol particles due to anthropogenic contributions, which, unlike the natural emissions, are expected to provide a contribution that changes with time. Effect of aerosol particles on cloud drop number concentrations and size distributions Clouds and fogs are characterized by their droplet size distribution as well as their liquid water content. Fog droplets typically have radii in the range from a few ␮ m to ;30᎐40 ␮ m and liquid water contents in the range of 0.05᎐0.1 g my3 . Clouds generally have droplet radii from 5 ␮ m up to ;100 ␮ m, with typical liquid water contents of ;0.05᎐2.5 g my3 Že.g., see Stephens, 1978, 1979.. For a description of cloud types, mechanisms of formation, and characteristics, see Wallace and Hobbs Ž1977., Pruppacher Ž1986., Cotton and Anthes Ž1989., Heymsfield Ž1993., and Pruppacher and Klett Ž1997.. There are several basic physical᎐chemical principles involved in the ability of aerosol particles to act as CCN and hence lead to cloud formation. These are the Kelvin effect Žincreased vapor pressure over a curved surface . and the lowering of vapor pressure of a solvent by a nonvolatile solute Žone of the colligative properties.. In Box 14.2, we briefly review these and then apply them to the development of the well-known Kohler curves that determine which particles will grow ¨ into cloud droplets by condensation of water vapor and which will not.



BOX 14.2

KELVIN EFFECT, VAPOR PRESSURE LOWERING, ¨ HLER CURVES AND THE KO 1. Kel¨ in effect. Recall that the change in free energy of a gas due to a change in pressure Žat constant T . from P1 to P2 is given by ⌬G s nRT ln Ž P2rP1 . .

Ž X.

Consider the situation in Fig. 14.37 in which a number of moles dn is transferred from a bulk liquid with vapor pressure P0 to a droplet of the same liquid having radius r and over which the vapor pressure is P. Assuming each system is at equilibrium, the free energies of the liquid and gas in each case must be equal. The change in free energy, dG, for transferring dn moles from the bulk liquid to the droplet is therefore the same as the accompanying free energy change for the gas. From Eq. ŽX., this is given by dG s dn RT ln Ž PrP0 . .

Ž Y.

However, there is also a free energy change due to the increase in surface area, dA, of the droplet caused by the transfer of the dn moles. The surface tension of the liquid, ␥ , is the work required per unit change in surface area to expand a surface against the intermolecular forces that tend to minimize the surface. The area of the initial droplet is 4␲ r 2 and hence the change due to a small change dr caused by transfer of dn moles is 8␲ r dr. The free energy change in the droplet due to an increase in its surface area is therefore dG s ␥ dA s 8␲ r␥ dr.

Ž Z.

This transfer of dn moles causes a volume change dV s 4␲ r 2 dr. If the molecular weight of the compound is MW and liquid density is ␳ , then ␳ s ŽMW. dnrŽ4␲ r 2 dr .. This expression can be used to

FIGURE 14.37 Basis of Kelvin effect for increased vapor pressure over small liquid droplets.

replace dr in Eq. ŽZ.: dG s 8␲ r␥ Ž MW. dnr4␲ r 2␳ s dnw 2␥ Ž MW. rr ␳ x . Ž AA. Combining Eqs. ŽY. and ŽAA. gives the Kelvin equation: ln Ž PrP0 . s w 2␥ Ž MW. rr ␳ RT x

Ž BB .

or PrP0 s exp w 2␥ Ž MW. rr ␳ RT x . This Kelvin equation says that the vapor pressure over a droplet depends exponentially on the inverse of the droplet radius. Thus, as the radius decreases, the vapor pressure over the droplet increases compared to that over the bulk liquid. This equation also holds for water coating an insoluble sphere ŽTwomey, 1977.. This has important implications for nucleation in the atmosphere. Condensation of a vapor such as water to form a liquid starts when a small number of water molecules form a cluster upon which other gaseous molecules can condense. However, the size of this initial cluster is very small, and from the Kelvin equation, the vapor pressure over the cluster would be so large that it would essentially immediately evaporate at the relatively small supersaturations found in the atmosphere, up to ;2% ŽPruppacher and Klett, 1997.. As a result, clouds and fogs would not form unless there was a preexisting particle upon which the water could initially condense. Such particles are known as cloud condensation nuclei, or CCN. While water is a major component of tropospheric particles, and hence largely determines the surface tension Ž␥ ., organics found in particles may act as surfactants Žsee Chapter 9.C.2.. In this case, their segregation at the air᎐water interface could potentially lead to a substantial surface tension lowering of such particles, which would lead to a lower equilibrium water vapor pressure over the droplet ŽEq. ŽBB.. and hence activation at smaller supersaturations. This possibility is discussed in more detail in the next section. 2. Vapor pressure lowering. Raoult’s law says that the vapor pressure of a solution component, A, whose pure vapor pressure is PAT is proportional to



its mole fraction in solution, xA , i.e., in a two-component solution of A and B, to xA s Ž1 y x B .: PA s xA PAT s Ž 1 y x B . PAT

i.e., PA rPAT s 1 y x B

or alternatively PAT

y PA s ⌬ P s

x B PAT .


Thus, if a nonvolatile solute is dissolved in water, the vapor pressure of water is lowered by an amount proportional to the mole fraction of dissolved solute, taking into account any dissociation that occurs Žvide infra.. It should be noted that this assumes ideal solution behavior. As we have seen in Chapter 9, there are a variety of dissolved solutes in atmospheric particles, which will lower the vapor pressure of droplets compared to that of pure water. As a result, there is great interest in the nature and fraction of water-soluble material in atmospheric particles and their size distribution Že.g., Eichel et al., 1996; Novakov and Corrigan, 1996; Hoffmann et al., 1997.. This vapor pressure lowering effect, then, works in the opposite direction to the Kelvin effect, which increases the vapor pressure over the droplet. The two effects are combined in what are known as the Kohler curves, ¨ which describe whether an aerosol particle in the atmosphere will grow into a cloud droplet or not under various conditions. 3. Kohler cur¨ es. Calculation of the mole fraction ¨ of dissolved solute, x B , in a water droplet requires knowing the number of moles of water and of dissolved solute. Take a two-component solution such as NaCl in water, where the solute dissociates into i ions Ž i s 2 for NaCl.. Assume m B grams of salt of molecular weight MWB are dissolved in water to form a solution of density ␳s . The number of moles of dissolved ions is im B rMWB . The number of moles of water for a drop of volume V s Ž4r3.␲ r 3 is w ␳s V y m B xrMWA , where MWA is the molecular weight of water. The mole fraction of dissolved

The term supersaturation, S, defined as Ž PArPAT y 1. is often expressed in the form of percent supersaturation, i.e., as 100Ž PA rPAT y 1., where PA and PAT are defined in Box 14.2. The relationship between the equilibrium vapor pressure over the droplet and that over the bulk liquid wEq. ŽHH.x is often expressed in a simplified form using the supersaturation: Ss

a r


b r3


Ž II .

solute ions ŽB. is then given by Eq. ŽDD.: xB s

im B rMWB


4 3

␲ r ␳s y m B . rMWA q im B rMWB 3

. Ž DD.

Using Eq. ŽCC., the vapor pressure lowering due to m B grams of dissolved salt that forms i ions per dissolved molecule is therefore given by PA PAT

im B rMWB


4 3


. ␲ r ␳s y m B . rMWA q im B rMWB Ž EE . 3

For dilute solutions, where m B is small, this reduces to PA PAT

s1y s1y

im B rMWB 4 3


␲ r ␳srMWA 3

b r3


im B MWA 4 3

␲␳s MWB



1 r3


where b s w im B MWArŽŽ4r3.␲␳s MWB .x. This vapor pressure lowering by the solute acts simultaneously with, and counteracts, the vapor pressure increase due to the Kelvin effect wEq. ŽBB.x. Multiplying the two, the net result for the vapor pressure above a solution containing a dissolved solute is given by PA PAT

s 1y

b r



2␥ MWA r ␳s RT


Ž GG .

Applying the approximation e x s 1 q x q x 2r2!q ⭈⭈⭈ and using only the first two terms, Eq. ŽGG. becomes PA PAT

s 1y

b r



a r


a r


b r3

, Ž HH.

where a s 2␥ MWA r␳s RT and the ry4 term has been omitted since it is small compared to the other three terms for radii of atmospheric interest.

Plots of S against radius are known as Kohler cur¨ es. ¨ Figure 14.38a shows a schematic diagram of such a curve. A more detailed thermodynamic treatment of Kohler curves is given by Reiss and Koper Ž1995.. ¨ Typical values of supersaturation found in clouds are between about 0.2 and up to 2%. For fogs, the values are lower by about an order of magnitude, typically between about 0.02 and 0.2% ŽPruppacher and Klett, 1997..


FIGURE 14.38 Ža. Schematic diagram of traditional Kohler curve, ¨

T where S s PA rPA y 1 is the supersaturation and r is the radius of the droplet. Žb. Kohler curves for 30-nm dry particle of ŽNH 4 . 2 SO4 : ¨ Ž1. traditional curve; Ž2. for a 500-nm CaSO4 particle Žslightly soluble. and ŽNH 4 . 2 SO4 as for curve 1; Ž3. as for curve 2 but in the presence of 1 ppb HNO3 which is taken up by the particle Žadapted from Kulmala et al., 1997..

The dependence of S on the radius r is such that at small radii, the second term due to the vapor pressure lowering dominates and S is negative. In this region, air with RH below 100% is in equilibrium with the particle; such diagrams are often plotted as a function of RH in this region, instead of S. At large radii, the first term due to the Kelvin effect dominates, and S becomes positive and ultimately reaches a maximum before decreasing again. The region to the left of the peak is known as the haze region for reasons that will


become apparent shortly, whereas that to the right is known as the cloud droplet region. Take as an example, a small dry particle of NaCl of a given mass Ž m B . that is introduced into air at a water vapor pressure corresponding to SA in Fig. 14.38a. Assuming that the RH is above the deliquescence point of NaCl, ;75% at 25⬚C, the particle will take up water, dissolve, and form a stable droplet of radius rA . Similarly, if the air saturation ratio increases to S B , the particle will, under equilibrium conditions, take up water and grow to radius r B . These particles are then in stable equilibrium with water in the air. Say the particle at point B loses some water molecules and starts to shrink. The equilibrium supersaturation for the smaller particle is lower than for the original particle. However, the supersaturation of the surrounding air remains higher, so that water will condense back out on the particle to bring it back to its original size. Similarly, if the particle at B gains some water molecules and the radius starts to increase, the value of S required to maintain this new size would be larger than that of the surrounding air and water would evaporate to restore the equilibrium size. In short, particles to the left of the peak in Fig. 14.38a do not tend to shrink or grow. Because they are generally in the 0.1- to 1-␮ m size range which scatters light efficiently Žsee Chapter 9.A.4., these particles are known as haze particles or droplets. These often occur at relative humidities below 100%. Consider, however, a particle at point D in Fig. 14.38a. If it gains some water molecules and the radius starts to increase, the surrounding air will have a larger supersaturation than the required equilibrium value of S for this larger particle. As a result, water will condense out on the droplet, causing it to grow further. Particles that lie to the right of the peak are thus in an unstable equilibrium Že.g., see Reiss and Koper, 1995. and can therefore activate into cloud droplets from the condensation of water. There are two questions with respect to potential indirect effects of aerosol particles on properties of clouds: Ž1. What are the sources of new particles? Ž2. How do these new particles grow to sufficient size Ž)50 nm. to act as CCN? The first issue is that of formation of new particles. As discussed in Chapter 9.B, nucleation of gases to form new particles in the atmosphere is not well understood. The observed rates of nucleation of H 2 SO4 , for example, greatly exceed the calculated rates. An important contributor to the formation of new particles in the boundary layer ŽBL. under some conditions appears to be exchange between the BL and the free troposphere Že.g., Davison et al., 1996; Raes et al., 1997; Clarke et al., 1997.. For example, some of the DMS from the oceans can be carried to the free



troposphere, where it is oxidized to sulfate, generating new CCN. Mixing of the air mass back into the BL may then quench the formation of new CCN in that region by scavenging the low-volatility species such as H 2 SO4 before they can nucleate to form new particles Že.g., see Slinn, 1992; and Raes, 1995.. The second issue is how these new particles grow into a sufficient size that they can act as CCN. The size to which the particles must grow to act as CCN is determined under equilibrium conditions by the Kohler ¨ curves. The peak values of S and r on the Kohler ¨ curves ŽFig. 14.38. are known as the critical values, Sc and rc Žsee Problem 6.. Whenever the supersaturation of the air mass is greater than Sc , condensation occurs to form cloud and fog droplets. However, if it is less than Sc , particles with radii less than the critical radius rc will maintain an equilibrium size Že.g., point B in Fig. 14.38a. and not form clouds or fogs. Of course, in the real atmosphere, there are a variety of initial particle sizes containing varying amounts of dissolved solutes, and the supersaturation of the air mass also changes with time. However, the Kohler curves provide a basis ¨ for understanding which particles can grow into clouds and fogs and which will not. It is important to note that the Kohler relationship ¨ assumes equilibrium. However, under real atmospheric conditions, the system may not be at equilibrium. P. Y. Chuang et al. Ž1997. and Hallberg et al. Ž1998. point out that if the time scale for growth of cloud droplets is larger than that for the particle to reach equilibrium, the growth of CCN into cloud droplets may be controlled by kinetics, rather than equilibrium. They suggest that ignoring this potential kinetic limitation may lead to overestimating the number of cloud droplets that will form from a given number of CCN. In addition, measurement techniques for CCN using cloud chambers may not give an accurate assessment of CCN under ambient conditions since the range of time scales encounted in air is much larger than that used in the measurements. At present, it is not clear how much of a problem such kinetic limitations present in the atmosphere. wSome parameterizations in use, however, do take into account the kinetic limitations Že.g., C. C. Chuang et al., 1997..x In addition, these traditional Kohler curves do not ¨ take into account the effects of slightly soluble solutes or gases that can dissolve in the particles. Figure 14.38b compares the traditional Kohler curves for a 30-nm dry ¨ particle of ŽNH 4 . 2 SO4 at 298 K with that for the same particle but containing a 500-nm core of slightly soluble CaSO4 ŽKulmala et al., 1997.. The increased particle size reduces the Kelvin effect contribution and a minimum in the curve reflects the point at which all of the CaSO4 dissolves. Also shown is a Kohler curve for the ¨

case where 1 ppb of gaseous HNO3 is present and is taken up by the particles. It can be seen that the equilibrium supersaturation is less than 1 for sizes up to ;10 ␮ m; i.e., for this more complex Žbut realistic . case, cloud droplets can form at RH below 100%. Possible mechanisms of growth for small particles into a sufficient size that they are on the right side of the Kohler curve include uptake of small particles into ¨ existing cloud droplets where in-cloud oxidation of gaseous species such as SO 2 to sulfate occurs. Evaporation of the cloud then leaves a larger particle containing the additional oxidation products Že.g., see Hoppel and Frick, 1990; Hegg, 1990; Van Dingenen et al., 1995; and Hoppel et al., 1996.. Other possible mechanisms of particle growth include coagulation of fine particles or the growth of existing particles by condensation of low-volatility products Že.g., see Lin et al., 1992, 1993a, 1993b; Hegg, 1990, 1993.. However, coagulation is not expected to be important in remote areas since the number concentration of particles is much smaller than in polluted areas where coagulation can be important ŽLin et al., 1992.. As we have already seen, the critical supersaturation Sc corresponding to the peak of the Kohler curve ¨ depends on a number of parameters unique to the aerosol particle. Thus, at a given supersaturation some particles will form cloud droplets and some will not. As a result, the total number of CCN will vary with the supersaturation used in the CCN measurement. This is illustrated in Fig. 14.39, which shows the concentration of CCN measured in Antarctica as a function of the percentage supersaturation for CCN that grow into droplets larger than 0.3 and 0.5 ␮ m, respectively ŽSaxena, 1996.. This particular set of measurements

FIGURE 14.39 CCN concentrations measured during a cloud event at Palmer Station, Antarctica, as a function of percent supersaturation Ž%S .. The two lines represent particles that grew to sizes greater than 0.3 and 0.5 ␮ m, respectively Žfrom Saxena, 1996..


was carried out during a burst of CCN production due to cloud processing Žvide infra.. The relationship between the number concentration Ž N . of CCN and the supersaturation Ž S . is often expressed in the form N s CS k , where C and k are empirical coefficients characteristic of the particular air mass ŽPruppacher and Klett, 1997.. However, alternate forms such as N s N0 w 1 y eyŽ BS . x , k

where B and k are empirical coefficients and N0 is the number concentration of CCN at infinite supersaturation, have been suggested to match data from laboratory studies where the aerosol composition is relatively simple ŽJi and Shaw, 1998.. As discussed earlier, if organics congregate at the air᎐water interface of particles and act as surfactants, they can lead to a reduced Kelvin effect and hence activation at lower supersaturations. In addition, if they dissolve, they will also contribute to Raoult’s law effects of the dissolved species. As discussed by Shulman et al. Ž1996., the dissolution of organics in particles may lead to modifed Kohler curves having two maxima ¨ instead of one. If the second maximum is at a higher supersaturation than the first, droplets could become partially activated. They also suggest that with the distribution of chemical compositions and particle sizes in the atmosphere, the region beyond the maxima could be relatively flat in shape, rather than having a steep negative slope. This would result in the formation of cloud droplets with radii corresponding to some characteristic metastable size. The effects of surfactants on particles is discussed in detail in Chapter 9.C.2b. In some cases, the interaction of organic surfactants with particles has been observed to enhance water uptake and hence the cloud nucleating properties of the particle, whereas in others, it inhibits water uptake. An example of the former case is a study by Cruz and Pandis Ž1998., who coated particles of ŽNH 4 . 2 SO4 with the C 5 dicarboxylic acid, glutaric acid. A coating of glutaric acid on ammonium sulfate increased the size of the particle and increased its cloud nucleating properties, but in a manner consistent with that expected from Kohler curves for a two-com¨ ponent solution assuming the organic does not alter the surface tension. On the other hand, coating with the water-insoluble dioctyl phthalate did not alter the activation of the inorganic salt. Similarly, Kotzick et al. Ž1997. and Weingartner et al. Ž1997. reported that oxidation of carbon or diesel soot particles by O 3 increased their CCN activation properties due to the formation of polar surface groups. As discussed in detail in Chapter 9, Saxena et al. Ž1995. have measured the hygroscopic behavior of par-


ticles in the Los Angeles area Ži.e., urban. and in the Grand Canyon, Arizona Ži.e., nonurban.. They found that organics in the urban particles inhibit water uptake, i.e., are hydrophobic in nature, whereas those in the nonurban particles are hydrophilic, i.e., increase water uptake. It may be that these differences are due to the formation of smaller, more oxidized, water-soluble organics during long-range transport to the nonurban site. Saxena et al. Ž1995. suggest that hydrophobic organics in the urban aerosol may form a surfactant film on the particles that inhibits water uptake. For example, difunctional acids such as oxalic and adipic acids have been shown to slow the rate of evaporation of water from droplets once they are sufficiently concentrated by the initial evaporation of water ŽShulman et al., 1997.. At any rate, as might be expected, the effects of organics on water uptake for real atmospheric particles clearly can be negative or positive, depending on their particular composition. As expected from the earlier discussion of the Kohler ¨ curves, not all particles act as CCN. For example, only about 15᎐20% of the Aiken nuclei Žsee Chapter 9.A.2. in a marine air mass off the coast of Washington state acted as CCN at 1% supersaturation ŽHegg et al., 1991b.. Similarly, in a marine air mass in Puerto Rico, between 24 and 70% of the particles measured at 0.5% supersaturation before cloud formation led to cloud droplet formation ŽNovakov et al., 1994.. Gillani, Leaitch, and co-workers Ž1995. carried out a detailed study of the fraction of accumulation mode particles Ždiameters from 0.17 to 2.07 ␮ m. that led to cloud droplet formation in continental stratiform clouds near Syracuse, New York. When the air mass was relatively clean, essentially all of the particles were activated to form cloud droplets in the cloud interior and the number concentration of cloud droplets increased linearly with the particle concentration. However, when the air mass was more polluted, the fraction of particles that were activated in the cloud interior was significantly smaller than one. This is illustrated by Fig. 14.40, which shows the variation of this fraction Ž F . as a function of the total particle concentration, Ntot . In the most polluted air masses Žas measured by large values of Ntot ., the fraction of particles activated was 0.28 " 0.08, whereas in the least polluted, it was as high as 0.96 " 0.05. The reason for this is likely that in the more polluted air masses, the higher number of particles provided a larger sink for water vapor, decreasing the extent of supersaturation. In short, while anthropogenically produced particles can act as cloud condensation nuclei, only a fraction of them actually do so. This fraction can be close to one



FIGURE 14.40 Fraction Ž F . of aerosol particles that are activated to form cloud droplets as a function of the total number of particles Ž Ntot .. The horizontal line represents the 50th percentile for 10 sets of aircraft measurements. The 5th, 25th, 75th, and 95th percentiles are also shown Žadapted from Gillani et al., 1995..

or as much as an order of magnitude smaller, depending on a number of factors, the most important of which are the particle sizes, the total particle concentration, and the local cooling rate in the cloud ŽGillani et al., 1995.. An increase in aerosol particles that can act as CCN can increase the number of cloud droplets and their size distribution, both of which can affect the light scattering properties of clouds and hence climate. We first briefly discuss the effects of clouds on climate and then the potential impacts of anthropogenic aerosols on the formation and properties of clouds. Clouds and global climate. Clouds in the troposphere interact with both solar and terrestrial radiation in complex ways, and either warming or cooling can result Že.g., see Ramanathan et al., 1989; Ramaswamy and Ramanathan, 1989; Fouquart et al., 1990; Harrison et al., 1990; Liou, 1992; King, 1993; Hartmann, 1993; Rossow and Zhang, 1995; Ramanathan, 1995; Crutzen and Ramanathan, 1996; and Baker, 1997.. Thus, marine stratiform clouds found in the boundary layer backscatter solar radiation, leading to negative radiative forcing and a net cooling. Although such clouds also absorb terrestrial infrared radiation, they emit at about the same temperature as the earth’s surface. As a result, as discussed for the greenhouse gases, there is little net effect Žsee Section A.2.. However, cirrus clouds and deep convective cloud anvils found near the tropopause emit long-wavelength infrared out to space ŽFig. 14.2c. at the colder temperatures characteristic of this region of the atmosphere. Because this energy emission is at lower temperatures, the net infrared emission out to space when they are present is smaller, leading to a positive radiative forcing, i.e., to warming ŽTwomey, 1991.. As a result, the net radiative forcing

due to low clouds over the oceans is generally negative, whereas it is positive over some continental regions with high clouds. On a global basis, the mean net effect is negative, ;y20 W my2 ŽBaker, 1997.. For reviews of the relationship between clouds and climate, and anthropogenic effects on them, see Hobbs Ž1993a, 1993b., King Ž1993., Hartmann Ž1993., Andreae Ž1995., Schwartz and Slingo Ž1996., Schwartz Ž1996., and Baker Ž1997.. Since some aerosol particles, which may be solids or liquids, e.g., H 2 SO4 , can serve as cloud condensation nuclei ŽCCN., increased particle emissions from anthropogenic processes have the potential for increasing the number of CCN. The concentration of droplets in a cloud is therefore expected to increase, although not necessarily in a linear fashion, with the increased concentrations of aerosol particles. The formation of a larger number of droplets for a given liquid water content will lead to each droplet being smaller, i.e., shift the size distribution to smaller droplets. This effect increases the cloud albedo and hence can contribute to global climate change ŽTwomey, 1974, 1977a,b, 1991; Twomey et al., 1984.. Evidence for these effects is discussed shortly. Finally, the shift to smaller drop sizes may decrease the precipitation rate from clouds, increasing their lifetimes and hence the average amount of cloud cover ŽAlbrecht, 1989; Lohmann and Feichter, 1997.. This is also expected to have a significant effect on global climate. It should be noted that while aerosol particles affect clouds by serving as CCN for cloud formation, the reverse is also true, i.e., clouds also affect the formation and size distribution of aerosol particles. For example, the oxidation of SO 2 to sulfate in clouds generates larger particles whose light scattering cross sections are larger than for smaller particles formed by gas-phase oxidation processes Že.g., see Lelieveld and Heintzenberg, 1992.. Aerosol particles incorporated into a cloud droplet will reappear as particles when the cloud evaporates. However, if new aerosol constituents are formed by in-cloud oxidation Že.g., of SO 2 to sulfate., the size of the resulting particle will be larger than the original particle. As discussed earlier, this is a potentially important process for growing aerosol particles, which are too small to serve as CCN at low values of Smax Že.g., for marine stratus ., into sufficiently large sizes Ži.e., to the right of the peak in Fig. 14.38. that they can act as CCN under the appropriate conditions of supersaturation Že.g., see Hegg, 1990; and Kaufman and Tanre, ´ 1994.. For example, as seen in Fig. 14.41, aerosol particle number size distributions in the clean marine boundary layer outside of clouds are often observed to have a bimodal distribution. The larger mode above 0.1 ␮ m



FIGURE 14.41 Typical particle size number distributions for marine aerosols outside of clouds where the total aerosol number concentration was - 500 cmy3 Žadapted from Anderson et al., 1994..

has been attributed to aerosol particles that have been ‘‘cloud processed’’ ŽHoppel et al., 1986; Hoppel and Frick, 1990; Anderson et al., 1994.. That is, these particles served as CCN upon which clouds formed, followed by aqueous-phase reactions and evaporation of the cloud droplets to leave larger particles. Indeed, in some studies Že.g., Van Dingenen et al., 1995., it is assumed that particles in this mode can be taken as a measure of the CCN that were available for cloud formation in the prior cloud event. As discussed in Chapter 9.A.2, such aqueous-phase processes in the atmosphere are also believed to lead to two peaks in accumulation mode particles in urban areas. The albedo Ž R . of a thick, boundary layer cloud that does not absorb solar radiation over a surface with zero albedo can be approximated ŽTwomey, 1991; Schwartz and Slingo, 1996; Baker, 1997. by ␶ Ž JJ . Albedo s R f a q␶ 1yg



The value of the factor a is usually taken as 1 or 2 and ␶ is the optical depth of the cloud, defined by IrI0 s ey␶ , where I and I0 refer to the transmittance of direct solar radiation in the presence and absence of the cloud, respectively. The optical depth can be approxi2 mated by ␶ ( 2␲ reff Nh, where reff is an effective average droplet radius for scattering of solar radiation, h is the thickness of the cloud, N is the number of cloud droplets per unit volume, and g is the asymmetry parameter for single scattering, defined as the average of the cosine of the scattering angle. A typical value of g is ;0.85 for cloud droplets. ŽNote that the cloud height h and N are not independent, since the number of cloud droplets affects the precipitation rate, and this alters the cloud height; e.g., see Pincus and Baker,

1994.. Typical cloud albedos for thick clouds in the boundary layer are ;0.5 over the ocean in midlatitudes; i.e., half of the incoming solar radiation is scattered back out to space ŽBaker, 1997.. This approximation, Eq. ŽJJ., illustrates why a change in the number of cloud droplets and their size affects the cloud albedo and hence the radiative forcing Žsee Problem 9.. An important aspect of quantifying the indirect effects of anthropogenic emissions is the recognition that the changes generated in the cloud albedo are not constant for all clouds but rather depend on the particular cloud properties. For example, the effects of an absolute increase in CCN on a cloud with a low droplet number concentration will be larger than for one with a larger droplet number concentration, even if the two have the same albedo Že.g., see Platnick and Twomey, 1994; and Taylor and McHaffie, 1994.. This ‘‘susceptibility’’ has been expressed in different ways, for example as dRrdŽln N .. Platnick and Twomey Ž1994. have derived an expression applicable to nonabsorbing clouds for the sensitivity Ž dRrdN . of cloud albedo, R, to changes in the number of cloud droplets, N, which they term cloud susceptibility: dR dN


␦R ␦␶

r v3

4␲␳ w 9W


Ž KK .

2 In Eq. ŽKK., ␶ ( 2␲ reff Nh is the optical thickness defined above, r v is the volume-weighted moment of the cloud droplet size distribution, which can be approximated by reff , ␳ w is the density of liquid water, and W is the liquid water content of the cloud. Using Eq. ŽJJ., the term ␶ Ž ␦ Rr␦␶ . can be shown to be equal to RŽ1 y R . Žsee Problem 10.. Platnick and Twomey Ž1994. have applied Eq. ŽKK. to marine clouds off the coast of California and southern Africa, to fogs in central California, and to ship tracks. Figure 14.42 shows a typical range of susceptibilities as a function of cloud droplet size. The measured susceptibilities in these studies covered three orders of magnitude, from 5 = 10y5 cm3 for fogs to 0.8 = 10y3 cm3 for marine clouds off south Africa and 2 = 10y2 cm3 for thin stratus clouds off the California coast. Similarly, Taylor and McHaffie Ž1994. report cloud susceptibilities in the range from 10y4 to )8 = 10y3 at various locations around the world. The highest susceptibilities were those with the smallest aerosol particle concentrations below the cloud base. As the particle concentration increased beyond ;500 cm3, the susceptibility was relatively constant at ;5 = 10y4 cm3. This means that the addition of new particles to a relatively clean air mass is far more effective than for a polluted one in terms of the effect on clouds. In short,



FIGURE 14.42 Cloud susceptibilities Žlogarithmic scale. and cloud droplet radii for stratus clouds off Africa and California and from fogs in California Žadapted from Platnick and Twomey, 1994..

it is important in quantifying the indirect effects of anthropogenic emissions that the individual cloud properties be taken into account. It should be noted that while most studies have focused on the effects of sulfate on stratocumulus clouds, model studies suggest that sulfate may also exert significant effects on convective clouds in tropical systems as well ŽAndronache et al., 1998.. The first major link between the indirect effects of aerosol particles and climate is whether there has been an increase in particles and in CCN due to anthropogenic activities. As discussed in Chapter 2, anthropogenic emissions of particles and of gas-phase precursors to particles such as SO 2 have clearly increased since preindustrial times, and it is reasonable that CCN have also increased. Ice core data provide a record of some of the species that can act as CCN. Not surprisingly, sulfate and nitrate in the ice cores have increased substantially over the past century ŽMayewski et al., 1986, 1990; Laj et al., 1992; Fischer et al., 1998.. For example, Figure 14.43 shows the increases in sulfate and nitrate since preindustrial times in an ice core in central Greenland ŽLaj et al., 1992.. Sulfate has increased by ;300% and nitrate by ;200%. This suggests that sulfate and nitrate CCN also increased, although not necessarily in direct proportion to the concentrations in the ice core measurements. Modeling studies by Langner et al. Ž1992. suggest that at most 6% of the anthropogenic SO 2 emissions can form new particles, since removal of SO 2 by direct deposition is large Ž;50%. and the portion that is oxidized in clouds does not lead to new particles. Taking these factors into account, Langner et al. Ž1992. estimate that new sulfate particles may have doubled since preindustrial times. There is other evidence that supports increased CCN due to increased anthropogenic emissions. Thus, typical

FIGURE 14.43 Ža. Sulfate and Žb. nitrate Žin ␮ equiv Ly1 . in central Greenland ice cores from ;1750 to 1985 Žadapted from Laj et al., 1992..

concentrations of CCN over industrialized continental areas are much larger than those in more remote regions. Over remote oceans, CCN concentrations measured using typical supersaturations of 0.7᎐1.25% are typically -100 cmy3 Že.g., Hegg et al., 1991a,b; Hudson and Li, 1995; Saxena, 1996. whereas over industrialized continents, number concentrations as high as 5000 cmy3 have been observed Že.g., Hudson, 1991; Pruppacher and Klett, 1997.. Another piece of evidence for anthropogenic emissions leading to increased CCN and hence effects on cloud properties such as albedo and extent is found in ‘‘ship tracks.’’ These are lines of clouds that trace ship movements, either in initially cloud-free regions ŽConover, 1966; Platnick and Twomey et al., 1994. or superimposed on preexisting clouds ŽCoakley et al., 1987.. Emissions associated with the ship exhausts serve as CCN. This allows clouds to form where the background CCN concentration is too small for cloud formation. Alternatively, the CCN can modify existing cloud properties in the exhaust plume by changing the number and size distribution of the cloud droplets as well as the liquid water content Že.g., Ferek et al., 1998.. For example, Fig. 14.44 shows the cloud number concentration Ž N ., effective radius Ž reff ., and liquid water content measured simultaneously during an aircraft flight through two ship tracks ŽKing et al., 1993.. In addition, the upwelling and downwelling radiation at 744 nm and at 2.2 ␮ m, respectively, are shown. Consistent with the foregoing discussion, the upwelling radiation at 744 nm increased in the ship tracks while the



downwelling radiation decreased. However, both the upward and downward infrared radiation at 2.2 ␮ m decreased due to increased absorption of this wavelength in the cloud. It is noteworthy that the cloud susceptibility, i.e., change in reflectance per change in cloud droplet number, in ship tracks has been measured to be smaller than outside of the tracks, as expected from the earlier discussion ŽPlatnick and Twomey, 1994.. Another piece of evidence supporting the relationship between anthropogenic emissions and CCN is the observation by Hudson and Li Ž1995. of higher CCN concentrations associated with higher O 3 levels and with higher particle concentrations in air masses near the Azores. Falkowski et al. Ž1992. and Kim and Cess Ž1993. also report enhanced cloud albedos near continental coastal regions having higher sulfate concentrations. A number of field studies have quantitatively examined the relationship between the CCN number concentration andror cloud droplet concentration and the mass concentration of non-sea salt sulfate Žnss.. Figure 14.45 shows one summary of some of these studies in the form of a log᎐log plot ŽVan Dingenen et al., 1995.. Given the variety of measured parameters and wide range of conditions encompassed by these data ŽCCN

FIGURE 14.45 Log᎐log plot of CCN or cloud droplet concentra-

FIGURE 14.44 Effect of ship emissions on Ža. cloud number concentration, N, Žb. effective cloud droplet radius, reff , Žc. cloud liquid water content, LWC, and Žd, e. down- and upwelling radiation at Žd. 744 nm and Že. 2.2 ␮ m Žadapted from King et al., 1993..

tion measured in a number of field studies as a function of non-sea salt sulfate Žnss.. Data are from ŽI. Cape Grim, Tasmania, CCN measured at 0.46% supersaturation ŽGras, 1989; Ayers and Gillett, 1989.; Ž^. Puerto Rico, CCN measured at 0.5% supersaturation ŽNovakov and Penner, 1993.; Ž䉫. western Washington, CNN measured at 0.3% supersaturation ŽBerresheim et al., 1993.; Ž`. the Azores ŽHegg et al., 1993.; Ž=. Whiteface Mountain, New York, cloud droplets measured ŽPueschel et al., 1986.; Ž). central Ontario, Canada, cloud droplets measured ŽLeaitch et al., 1992.; and Ž䢇. North Atlantic, cloud-processed accumulation mode particles measured ŽVan Dingenen et al., 1995. Žadapted from Van Dingenen et al., 1995..



at different supersaturations versus cloud droplet concentrations, the large range in sulfate concentrations, etc.., such correlations Ž r 2 s 0.42. suggest that it is possible to relate particle mass to CCN and cloud droplet number concentration, and ultimately to changes in cloud albedo, albeit within a large uncertainty. However, some caution is needed in applying such correlations, as might be expected, particularly given our current relatively rudimentary understanding of the indirect effects of aerosol particles on clouds. For example, the data in Fig. 14.45 include both CCN and cloud droplet number concentrations, which both appear to be correlated to nss. However, this is not always the case. While CCN and non-sea salt mass concentrations were observed to be highly correlated at a marine site in Puerto Rico, cumulus cloud droplet number concentrations were not, and stratocumulus cloud droplet number concentrations showed a very low sensitivity to nss ŽNovakov et al., 1994.. Similarly, Anderson and co-workers Ž1994. did not observe a convincing relationship between the cloud droplet number concentration and the aerosol number or volume at a coastal mountain site in the state of Washington. Entrainment and mixing processes in the clouds may have played major roles in these Žand, of course, many other. studies. Finally, some studies suggest that particularly at very small sulfate concentrations, a wide range of CCN can be observed Že.g., Hegg, 1994., which may be related to non-sulfate species acting as CCN. Indeed, while the initial focus on the indirect effects of anthropogenic aerosols has been on sulfate, there is increasing evidence that other species may also not only contribute significantly to CCN but actually dominate it under many circumstances. For example, although CCN at 1% supersaturation were correlated with sulfate in air masses over the northeast Pacific and the northeast Atlantic, the slope of the curve relating the two was much higher for the relatively clean northeast Pacific ŽHegg et al., 1993.. The authors suggest that this is consistent with DMS as the major source of sulfate over the Pacific. Over the Atlantic, however, the slope of CCN versus sulfate was smaller and there was a significant intercept, suggesting that much of the CCN was not formed from sulfate. Observations at a coastal site in the state of Washington also suggested that components other than sulfate may be important in CCN formation at 0.9% supersaturation ŽBerresheim et al., 1993.. Novakov and Penner Ž1993. measured the mass size distributions of sulfur, organic carbon, and chlorine Žcharacteristic of sea salt. as well as the CCN concentration Žat 0.5% supersaturation ., nss, and Aitken nuclei concentrations at a mountain peak in Puerto Rico.

They concluded that about 63% of the CCN at this site was due to organic aerosol particles, possibly due to some unspecified anthropogenic sources. Similar measurements at Point Reyes, California, gave variable contributions of organics and sulfate to CCN, ranging from 4 to 78% for organic particles, from 19 to 64% for sulfate, and from 9 to 31% for NaCl in sea salt particles ŽRivera-Carpio et al., 1996.. While Andrews et al. Ž1997. suggest that the organic aerosol particles in the Puerto Rican studies may have originated from the rain forest below the mountain peak sampling site, subsequent studies at the mountain site and at a site in the Atlantic Ocean suggest that a large fraction of this organic aerosol may originate from the ocean ŽNovakov et al., 1997a.. A similar oceanic source of CCN measured in Antarctica was suggested by Saxena Ž1996.. It is particularly interesting that most of the organic aerosol particles in the Puerto Rican study were water soluble; in addition, their average mass concentration Ž390 ng my3 . was larger than that of sulfate Ž270 ng my3 .. This combination of water solubility and relatively high mass fraction suggests that the organic particles may be particularly effective as CCN ŽNovakov et al., 1997a.. Recent studies have provided additional evidence for the contribution of organics to CCN. For example, Matsumoto et al. Ž1998. measured CCN at 0.5 and 1% supersaturations, along with the aerosol particle composition and size distribution at the Ogasawara Islands in the northwest Pacific Ocean. In agreement with earlier studies, air masses affected by continental emissions had CCN concentrations of ;150᎐1000 cmy3 Žat 1% supersaturation . compared to 30᎐150 cmy3 for clean air masses. Sulfate, nitrate, and ammonium in the particles were correlated with 222 Rn, which is a tracer of continental air masses. Oxalate was also found in the particles, primarily in the accumulation mode Ž-1.1 ␮ m., and was highly correlated with 222 Rn, indicating an anthropogenic source. However, formate and acetate were not well correlated with 222 Rn, suggesting marine biogenic sources for these species. A major contributor to the aerosol mass Ž;80% of the total mass. was unspecified water-soluble organics. As discussed in Section C.1a, sea salt particles in the marine boundary layer have been shown to likely play a major role in backscattering of solar radiation ŽMurphy 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 ŽGong 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 Kohler curves ŽFig. 14.38., this means that ¨ the size of other particles such as non-sea salt sulfate Žnss. 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 ŽO’Dowd et al., 1997a, 1997b.. Sea salt particles also provide an aqueous medium for the oxidation of SO 2 to sulfate Že.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 SO 2 to sulfate, are pH dependent. This is in part due to the decreasing concentrations of dissolved SŽIV. in the aqueous phase as the pH falls and in part due to the dependence of the reaction kinetics on pH. Seawater is basic ŽpH ;8. so that newly formed sea salt particles are likely basic when initially formed. Under these conditions, oxidation by O 3 is important Žsee 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 H 2 O 2 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 ŽKeene and Savoie, 1998, 1999.. Oxidation by HOCl, believed to be an important intermediate in halogen chemistry in the marine boundary layer, has been proposed to be important as well Že.g., see Vogt et al., 1996; Keene et al., 1998; and Chapter 8.C.3.. The significance of this oxidation of SŽIV. 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 Že.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 HNO3 and HCl onto particles may alter their ability to act as CCN Že.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 Ž N .. As seen in Eqs. ŽJJ. and ŽKK., 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 Ž1967. measured cloud droplet number concentrations and condensation nuclei at 0.5% supersaturation below the base of clouds upwind Žover the ocean. and downwind of a region in which sugar cane was burning. The average concentration of CCN was 280 cmy3 over the ocean but 750 cmy3 downwind of the fire; the cloud droplet number concentration similarly increased from 300 to 920 cmy3 . 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 Že.g., Twomey and Warner, 1967.. Similarly, Martin and co-workers Ž1994. measured aerosol particles in the size range from 0.05 to 1.5 ␮ 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 ŽMartin 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. Ž1995. discussed earlier. In addition, the relationship is much more scattered, indicating that the chemical



FIGURE 14.46 Average cloud droplet number concentration as a function of subcloud aerosol particle concentration Ž0.05᎐1.5 ␮ m. in marine Ž䢇. and continental ŽI. air masses Žadapted 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 Žwith diameters )0.018 ␮ m. and, in particular, with the concentration of light-absorbing materials in the ice crystals ŽStrom ¨ and Ohlsson, 1998.. 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 Žsee Eqs. ŽJJ. and ŽKK... Figure 14.47, for example, shows the size distribution for cloud droplets measured in urban and nonurban air around Denver, Colorado ŽAlkezweeny et al., 1993.. The median volume diameter was 14 ␮ m 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 cmy3 compared to 22 cmy3 for the nonurban cloud. Analogous results have been reported in other studies Že.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 ␮ m larger than over continental regions; in addition, marine cloud droplet radii in the Southern Hemisphere are ;1 ␮ m larger than in the Northern Hemisphere ŽHan et al., 1994.. Hudson and Li Ž1995. 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 Ž1989. 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 ŽLohmann and Feichter, 1997.. Clearly, this too could play a role in global climate change. The studies by Hudson and Li Ž1995. also reported evidence for this effect in that the number of ‘‘drizzle drops’’ with diameters )50 ␮ m was smaller Žby an order of magnitude. in the cloud in the polluted air mass ŽTable 14.8.. Related to this is the observation by Parungo et al. Ž1994. that there has been an increase in total oceanic clouds from 1930 to 1981, with the change in the Northern Hemisphere Ž2.3%. being about double that for the Southern Hemisphere Ž1.2%..

TABLE 14.8 Some Cloud and Subcloud Aerosol Properties Measured in Polluted and Clean Air Masses near the Azores a Property Aerosol particle concentration CCN concentration at S s 0.7% at S s 0.04% Cloud drop concentration b Mean cloud drop diameter b Drizzle drop concentration c a

FIGURE 14.47 Cloud droplet size distributions for stratiform clouds in the Denver area for urban and nonurban air masses Žadapted from Alkezweeny et al., 1993..

Clean air

Polluted air

176 cmy3

806 cmy3

116 cmy3 2 cmy3 10᎐100 cmy3 18 ␮ m 800 Ly1

668 cmy3 119 cmy3 220᎐370 cmy3 7᎐9 ␮ m 80 Ly1

Adapted from Hudson and Li Ž1995.. Diameters 3᎐66 ␮ m; upper end overlaps lower end of drizzle drop range. c Diameters 50᎐600 ␮ m defined as drizzle drops. b


Parungo and co-workers suggest this may be due to the indirect effects of increasing SO 2 emissions. The indirect effect of aerosols on climate, which at present contributes a major uncertainty in understanding anthropogenic perturbations on climate, is a very active area of research. For some typical model treatments of this indirect effect and how it interacts with those due to other, simultaneous, perturbations, see, for example, Jones et al. Ž1994., Hansen et al. Ž1997a᎐d., C. C. Chuang et al. Ž1997., Lohmann and Feichter Ž1997., and Pan et al. Ž1998.. Figure 14.48 shows one assessment ŽHansen et al., 1997d. of the contributions of anthropogenic emissions to the average global radiative forcing from preindustrial times to the present as well as that due to changes in solar intensity over the past 200 years Žsee Section D.3.. The contributions due to an increase in tropospheric O 3 from preindustrial times to 1980 and that due to stratospheric ozone destruction from 1979 to 1995 are predicted to essentially cancel out. Three contributions due to changes in tropospheric aerosol particles are included. Desert aerosols give a positive radiative forcing because of their absorption of light discussed earlier, whereas sulfate and biomass particles scatter light, leading to a negative radiative forcing. The indirect effect of particles on clouds has very large uncertainties associated with it and is shown as y1 W my2 in Fig. 14.48. Finally, changes in vegetation are estimated to have contributed y0.2 W my2 , due the reduction in the area of forests, which are dark.


As discussed in IPCC Ž1996., the confidence level associated with these values ranges from high for the greenhouse gases to very low for tropospheric aerosols, and in particular for the indirect effects. For example, the calculations of Penner et al. Ž1998. suggest a larger direct radiative forcing due to sulfate aerosol particles Žy0.81 W my2 . than that shown in Fig. 14.48 and a global average contribution of q0.16 W my2 for fossil fuel black and organic carbon particles. The contribution of changes in the solar flux and uncertainties in this are discussed in Section D.3. The uncertainties in the indirect effects on clouds are very large. As discussed earlier, increased CCN can alter the properties of clouds in several ways that can impact climate. Thus, they can lead to changes in cloud albedo and, in addition, alter the size distribution of cloud droplets, changing the precipitation rate and hence cloud lifetime. For example, Lohmann and Feichter Ž1997. carried out model studies of the indirect effects of sulfate on clouds and predicted increases in shortwave cloud forcing ranging from y1.4 to y4.8 W my2 . A significant portion of the effects was due to changes in cloud lifetime; for example, for the y1.4 W my2 case, about 40% was attributed to changes in the cloud lifetime and 60% to changes in cloud albedo. It is important to note that such globally and annually averaged estimates of contributions to radiative forcing are not expected to be the sole measures of effects on climate. The inference may be mistakenly drawn that negative radiative forcing, e.g., through

FIGURE 14.48 Calculated radiative forcings due to changes in greenhouse gases, particles, clouds, solar radiation, and vegetation from preindustrial times to 1995. That due to changes in stratospheric ozone is for the 1979᎐1995 period Žadapted from Hansen et al., 1997d..



tropospheric aerosol particles, may largely counterbalance the positive radiative forcing due to the greenhouse gases and hence there will be no net change in climate. This is not expected to be the case since the effects operate on different geographical and temporal scales. Thus, many of the greenhouse gases Že.g., CO 2 , CH 4 , N2 O, and the CFCs. are sufficiently long-lived to be globally distributed. Their contributions to radiative forcing vary geographically, from about 3 W my2 over hot regions such as the Sahara to ;0.6 W my2 over the South Pole ŽNational Research Council, 1996.. Shorter-lived greenhouse gases such as O 3 have much more spatial and temporal variability, with associated differences in their contributions to radiative forcing. However, the contribution of all greenhouse gases to radiative forcing operates both day and night since it involves their interaction with terrestrial radiation. On the other hand, aerosol particles from anthropogenic activities tend to be concentrated over or near industrial regions in the continents. Because both the direct and indirect effects of particles are predominantly in terms of scattering solar radiation, their effects are expected primarily during the day. For example, model studies by Sinha and Harries Ž1997. have explored a hypothetical case in which CO 2 is doubled to 710 ppm and the amount of tropospheric aerosol is increased about a factor of four, giving no net change in the predicted equilibrium surface temperature. However, even with a predicted net surface temperature change of zero, significant effects on climate are still predicted. The solar radiation at the surface at mid and low latitudes is predicted for this hypothetical case to decrease by as much as y6 W my2 in January. Similarly, the vertical distribution of the rate of total radiative heating is predicted to change by more than 4% at some altitudes, which would be expected to lead to changes in the lapse rate, potentially affecting atmospheric circulation processes. In addition to the differences in geographical distribution of the greenhouse gases compared to the aerosol particles and the day᎐night differences, there are also differences in their temporal behavior. As discussed earlier, typical residence times for sulfate particles are about a week, whereas that of CO 2 is about 100 years. As a result, the impacts of sulfate aerosols are almost immediately manifested, whereas those due to CO 2 occur over decades to centuries ŽSchwartz, 1993.. Hansen and co-workers have carried out modeling studies that examine the effects of various perturbations, both anthropogenic and natural, on climate ŽHansen et al., 1997a᎐c.. The altitude and geographical location of the forcings are shown to be important determinants of the effects on climate, rather than simply the magnitude of the forcing. For example, the

addition or removal of heat in the upper troposphere is partially compensated by changes in radiation to space, which does not occur close to the earth’s surface ŽHansen et al., 1997b.. In short, while net radiative forcing is a convenient means for examining the potential importance of various anthropogenic perturbations for climate, it cannot be used in an additive manner for gases and aerosol particles to predict the ultimate impacts. b. Heterogeneous Chemistry In© ol© ing Climate Species Another potential contribution of aerosol particles to global climate is that of heterogeneous chemistry. For example, particle surfaces could in principle destroy greenhouse gases such as ozone that are surface sensitive. Another example is the formation of greenhouse gases such as N2 O on surfaces. Thus, nitrous acid ŽHONO., which is itself formed by heterogeneous reactions on surfaces, has been shown to react on acid surfaces to generate N2 O by a mechanism that is not well understood Že.g., see Wiesen et al., 1995; Pires et al., 1996; and Pires and Rossi, 1997.. Given the present lack of understanding of the reaction mechanism, it is not possible to assess the importance of such heterogeneous chemistry for N2 O formation in the atmosphere. However, it does illustrate the potential for heterogeneous chemistry on aerosol particles to impact global climate through the effect on gas-phase species.

D. SOME OTHER FACTORS AFFECTING GLOBAL CLIMATE As discussed at the beginning of this chapter, the focus here is on the relationship between atmospheric chemistry and global climate change, rather than on the magnitude of this change and its causes. However, to place the role of atmospheric chemistry in context, we briefly treat in this section some other important factors known to be involved, or thought to be potentially involved, in climate change. For a more detailed discussion of these and other related issues, see IPCC Ž1996..

1. Absorption of Solar Radiation by Clouds Although clouds form on existing particles, the ‘‘solutions’’ formed are quite dilute and hence are not expected to absorb solar radiation significantly. As a result, it has been commonly accepted that clouds will predominantly reflect solar radiation and that absorption will not be significant. However, as early as 1951, it was suggested that clouds appeared to be absorbing


FIGURE 14.49 Absorption of light from an overhead sun by water associated with a 1-km stratus cloud with its top at an altitude of 2 km. The solid line is the absorption due to liquid water, the dashed line water vapor inside the cloud, and the dotted line water vapor in a column in the atmosphere Žadapted from Davies et al., 1984..

significantly as well ŽFritz, 1951.. This has been dubbed the ‘‘cloud absorption anomaly’’ wsee reviews by Stephens and Tsay Ž 1990. , Liou Ž1992 ., and Ramanathan and Vogelmann Ž1997.x. Although this area might not be considered to fall in the realm of ‘‘atmosphere chemistry’’ per se, it is clearly potentially very important in the relationship between clouds and global climate. Figure 14.49 shows the absorption of light from an overhead sun by liquid cloud droplets, water vapor inside the cloud, and water vapor in a column in the atmosphere for a 1-km stratus cloud whose top is 2 km above the ground ŽDavies et al., 1984; see also Goldstein and Penner, 1964.. There are small amounts of absorption in the tail end of the red region of the visible attributed to water vapor in and outside the cloud. The absorption increases into the near-IR Žthe region from ;780 to 2500 nm or 12,800᎐4000 cmy1 . and mid-IR Ž2.5᎐50 ␮ m or 4000᎐200 cmy1 . where liquid water in the cloud absorbs Že.g., see Evans and Puckrin, 1996.. Several different approaches have been taken to investigate whether there is more absorption of visible light by clouds than expected based on current models of radiative transfer in the atmosphere. Some of these approaches and results are discussed in Box 14.3. While there thus appears to be evidence for apparent excess absorption of solar radiation by clouds, there is substantial controversy over whether this is indeed true absorption or whether there is some other explanation for the discrepancies Že.g., see Stephens and


Tsay, 1990; Imre et al., 1996; Stephens, 1996; Cess and Zhang, 1996; Pilewskie and Valero, 1996; and Ramanathan and Vogelmann, 1997.. For example, Li et al. Ž1995. also analyzed solar flux surface and satellite data over a 4-year period to obtain values of the ratio of shortwave cloud forcing at the surface to that at the top of the atmosphere. These varied from about 1.4 in the tropics, in agreement with Cess et al. Ž1995., to values less than 1 in polar regions. They concluded that, within the uncertainties, their analysis does not provide support for excess cloud absorption of solar radiation ŽLi et al., 1995; Li and Moreau, 1996., although Zhang et al. Ž1997. suggest there may have been some unrecognized complexities in the analysis of the satellite data. Similarly, Chou et al. Ž1998. used measurements of surface radiative fluxes and satellite radiance data in the Pacific warm pool region to conclude that the effect of clouds was similar to that expected, i.e., that the excess absorption, if it exists, is small. Based on aircraft measurements, Francis et al. Ž1997. suggested that the excess cloud absorption was insignificant. In comparing satellite and ground-based observations of solar flux to model predictions, Arking Ž1996. also found no evidence for significant cloud excess absorption, although he reported a discrepancy between models and measurements that is not found over the central equatorial Pacific Ocean under clear skies ŽConant et al., 1997.. Imre et al. Ž1996. used collocated satellite and surface observations of shortwavelength fluxes at a site in Oklahoma to probe for the contribution of enhanced absorption by clouds, but found none. They suggested that uncertainties and biases in the analyses, particularly in the clear-sky references used, can give rise to apparent excess cloud absorption that is an artifact of the analysis. In short, whether excess absorption of solar radiation by clouds even occurs and why there are discrepancies between models and measurements remain controversial. For example, absorption into the visible region can be enhanced by the presence of strongly absorbing species such as soot, either in the cloud droplets themselves or as aerosol particles suspended between the cloud droplets, i.e., as interstitial aerosol particles Že.g., see Stephens and Tsay, 1990; Chylek ´ and Hallett, 1992; and Mel’nikova and Mikhaylov, 1994.. However, as discussed shortly, the enhanced absorption of solar radiation by clouds has been reported in many locations globally, including in remote regions, and hence appears less likely to be so directly associated with anthropogenic emissions. It has also been suggested that measurements of cloud drop size distributions have missed the presence of larger ‘‘drizzle drops’’



BOX 14.3

SOME INVESTIGATIONS OF THE CLOUD ABSORPTION ANOMALY One approach to the cloud absorption anomaly has been to examine the energy budget of the socalled ‘‘warm pool’’ area in the western Pacific Ocean, which is the region from approximately 140⬚E to 170⬚E and 10⬚N to 10⬚S ŽRamanathan et al., 1995.. This area has relatively high annual mean sea-surface temperatures ŽSST., up to 302.5 K, resulting in a humid and cloudy atmosphere with which is associated frequent deep convection. The annual mean heat transport Ž D . both horizontally and vertically out of the ocean surface mixed layer is known to be small, D F 20 W my2 . To give a constant temperature of the mixed layer, this small heat transport, D, out of the layer must be balanced by the net energy flux Ž H . at the ocean surface. As shown in Fig. 14.50, H is a balance between positive contributions from incoming solar radiation, S, and negative contributions from net outgoing thermal radiation, F Ži.e., up minus down., and evaporative Ž E . and turbulent sensible Ž h. heat fluxes, i.e., H s S y F y Ž E q h.. Using available data to estimate values for all of the energy terms except S leads to a

calculated value for the solar input of 175 W my2 . However, the solar radiation input under clear skies is well known to be 275 W my2 . If the total solar input is treated as the sum of incoming solar radiation and a contribution due to clouds, then the latter must be y100 W my2 . The outgoing solar radiation at the top of the atmosphere in this location based on 5 years of data from the Earth Radiation Budget Experiment ŽERBE. on the Earth Radiation Budget Satellite ŽERBS. is estimated to be y66 W my2 Žthe negative sign indicating cooling by reflection of solar radiation by clouds.. The ratio of the shortwave cloud energy flux at the surface to that at the top of the atmosphere must be 100r66 s 1.5, representing a long-term average for this effect. A similar conclusion is reached using direct measurements of solar fluxes at the top of the atmosphere ŽTOA. and at surface sites under clear compared to cloudy conditions Že.g., Cess et al., 1995, 1996b; Evans et al., 1995.. Figure 14.51a shows the absorptance, defined as the fraction of the down-

FIGURE 14.50 Schematic of energy balance in ‘‘warm pool’’ in western Pacific Ocean used to deduce the net effect of clouds on solar radiation. All numbers are given in W my2 Žadapted from Ramanathan et al., 1995..



FIGURE 14.51 Ža. Fraction of downward solar radiation at the top of the atmosphere that is absorbed, measured at Boulder, Colorado, under clear or cloudy skies compared to CCM2 model predictions; Žb. same as Ža. but in warm pool area over the central Pacific compared to the CCM3 model; Žc. measured maximum albedo at the top of the atmosphere ŽTOA. at three locations compared to CCM2 model predictions Žadapted from Cess et al., 1995, 1996b; and Valero et al., 1997a..

ward solar radiation at the top of the atmosphere that is absorbed, at Boulder, Colorado, under clear and cloudy skies ŽCess et al., 1995, 1996b.. Also shown are the values predicted using one climate model ŽCommunity Climate Model version 2, CCM2., which does not include enhanced cloud absorption of solar radiation. The agreement between observations and model predictions is excellent for clear skies. However, for cloudy skies the

observed absorption is, on average, ;30% larger than the model predictions. Similar conclusions are reached using data over the warm pool ŽFig. 14.51b.. Such data give values for the ratio of the shortwave cloud forcing at the surface to that at the top of the atmosphere of 1.5, the same as that obtained using the energy balance approach over the warm pool. Related to this are the maximum albedos measured at the top of the atmosphere at three loca-



tions, which are compared to CCM2 model predictions in Fig. 14.51c. These values were obtained from plots of broadband Ž0.2᎐5 ␮ m. albedo Ž ␣ . measured at the top of the atmosphere against the transmittance of the atmosphere ŽT ., determined from the ratio of downward solar flux at the surface to that at the top of the atmosphere. Extrapolation to zero transmittance gives the maximum mean values of albedo shown in Fig. 14.51c. The maximum measured value is ;0.7, whereas the maximum for model predictions for this limit of thick clouds is ;0.8, perhaps indicative of some unaccounted absorption by clouds. Another approach to this problem has been to make simultaneous flux measurements above and below clouds using aircraft Že.g., see Pilewskie and Valero, 1995, 1996; Valero et al., 1997a; Zender et al., 1997.. The use of radiometers covering different spectral regions in such studies has proven useful in deconvoluting the contributions of absorption in the visible and near-IR, respectively. Figure 14.52a shows absorptances measured in one such study in Okla-

homa, and Fig. 14.52b shows the corresponding measured values of transmittance ŽValero et al., 1997a.. Three different values of absorptance are shown in Fig. 14.52a: Ž1. broadband absorptance from 224 nm to 3.91 ␮ m; Ž2. the contribution of near-infrared absorptances, in this study defined as the region from 680 nm to 3.3 ␮ m; and Ž3. spectral band measurements at 500 nm Ž10-nm width.. The difference between the broadband and near-infrared measurements is a measure of the contribution of absorption in the visible region, and it is seen that it increases with increasing cloudiness. Consistent with the evidence for increased absorption with increased cloudiness is the simultaneous decrease in transmittance ŽFig. 14.52b.. Valero et al. Ž1997a. suggest that if the absorption were primarily in the near-IR, one would expect the contribution of the near-IR to transmittance seen for the cloudiest conditions in Fig. 14.51b to be relatively small; that it is still substantial supports the contention that absorption in the visible is important. However, it is interesting that the absorptance at 500 nm ŽFig. 14.51a. shows

FIGURE 14.52 Ža. Absorptance and Žb. transmittance measured on days with varying degrees of cloudiness using aircraft colocated above and below the clouds where broadband Ž224 nm to 3.91 ␮ m., near-IR Ž680 nm to 3.30 ␮ m., and spectral band Ž10-nm width centered at 500 nm. measurements were made Žadapted from Valero et al., 1997a..


no evidence of increased absorption with increasing cloudiness; whatever the absorbing species is Žif indeed there is one; see below. it does not appear to absorb light at 500 nm. The difference between the broadband measured absorption Ž224 nm to 3.91 ␮ m. and that from the near-IR radiometer Ž680 nm to 3.3 ␮ m. was used as a measure of the visible light absorption ŽZender et al., 1997.. On the cloudiest day during these studies, about 25% of the shortwave absorption was attributable to visible light, compared to 10% predicted by a model ŽZender et al., 1997.. These studies also report values for the ratio of the shortwave cloud forcing at the surface to that at the top of the

ŽWiscombe and Welch, 1986.. Cloud absorption is related to droplet size in a complex way ŽStephens and Tsay, 1990. so that errors in droplet size measurements can alter the model-predicted absorption by a cloud. The treatment of absorption due to water vapor is another possibility. As discussed by Crisp Ž1997., the treatment of water vapor in models is simplified and may not properly reflect, for example, continuum absorptions between major bands in the near-IR. Model calculations suggest that the presence of a thin, saturated layer of water vapor above the clouds, for example, leads to increased absorption by 2᎐6% ŽDavies et al., 1984; Podgorny et al., 1998.. However, the Crisp calculations indicate that this cannot account for all of the observed excess absorption. In addition, it has been suggested that the excess absorption is really due to radiation escaping from the sides of clouds, which would not be observed in measurements carried out above and below an isolated cloud ŽAckerman and Cox, 1981.. However, Valero et al. Ž1997a. argue that this should not give significant errors for continuous measurements made over clouds and analyzed with sufficiently long averaging times. Finally, whether the models properly capture the radiative effects, e.g., due to the heterogeneity of drop sizes in clouds and the cloud shape, is not clear Že.g., see Chou et al., 1995; Lubin et al., 1996; Byrne et al., 1996; and Loeb and Davies, 1996.. For example, while many models assume plane parallel cloud geometry, the absorption for other shapes such as wavy, broken clouds can be different by as much as 10᎐15%, depending on the solar zenith angle ŽPodgorny et al., 1998.. One important aspect of radiation and clouds that may ultimately prove to be important in this issue of excess cloud absorption is the very long effective path lengths for light inside clouds due to multiple scatter-


atmosphere in the range of 1.36᎐1.65, in agreement with the warm pool energy balance studies, but again significantly larger than model-predicted values of 1.12᎐1.14. Zender et al. Ž1997. also show that several other measures of cloud absorption, such as the slope of plots of albedo versus transmittance, are consistent with excess absorption of solar radiation that is not included in the models. Measurements of the single-scattering albedo Žthe fraction of incident energy that appears as scattered radiation. inside a marine stratocumulus cloud implied a small contribution from enhanced absorption as well ŽKing et al., 1990..

ing processes. For example, Pfeilsticker et al. Ž1997. measured the absorption of solar radiation by the oxygen collision complex ŽO 2 . 2 under clear and cloudy sky conditions. The absorption is sufficiently weak that it is in the linear, ‘‘weak-absorber’’ regime Žsee Section A.3.. Based on the Beer᎐Lambert law, the effective path length can be calculated from the known absorption cross section and atmospheric concentrations, combined with measurements of its atmospheric absorbance. For one cloud, for example, this approach gave an extra effective path length of 135 km, about an order of magnitude larger than the clear-sky geometrical path length! The presence of even a relatively weak absorption that is not accounted for in the models could have a disproportionate effect with such large effective path lengths Že.g., see Kondrat’ev et al., 1996a,b.. If this ‘‘excess absorption’’ by clouds is ultimately shown to be a real phenomenon, then an increased cloud formation and extent due to anthropogenic emissions may alter the radiative balance of the atmosphere not only through increased reflectance but also through increased absorption of solar radiation. Such an effect could impact atmospheric temperatures, their vertical distribution, and circulation, as well as surface wind speeds and the surface latent heat flux ŽKiehl et al., 1995.. Hence establishing if this is truly excess absorption, and if so, its origins, is a critical issue that remains to be resolved.

2. Feedbacks: Water Vapor, Clouds, and the ‘‘Supergreenhouse Effect’’ Given the complexity of the ocean᎐atmosphere᎐ biosphere system, it is not surprising that there are a



number of feedbacks that greatly complicate the accurate prediction of the effects of anthropogenic or natural emissions. For example, warming is expected to lead to decreased amounts of clouds and hence to changes in cloud contributions to radiative forcing wsee a comparison of model predictions for cloud feedbacks in Cess et al. Ž1996b.x. The DMS᎐CCN᎐cloud formationr reflectance discussed earlier is another such example. Another important example involves water vapor in the atmosphere. Water vapor is the most important greenhouse gas, and its concentration in the atmosphere is a function of temperature as given by the Clausius᎐Clapeyron equation: ln

P2 P1


⌬ H vap





1 T1

rather than changes in SST. In addition, Chou et al. Ž1998. found that the regions of maximum sea surface temperature did not necessarily coincide with those of maximum cloudiness, as expected if such feedbacks were operative. Some of these uncertainties are summarized by Stephens and Slingo Ž1992. and further discussion is found in Ramanathan and Collins Ž1992, 1993., Lau et al. Ž1994., Ramanathan et al. Ž1994., and Fu et al. Ž1993.. However, there are additional data supporting the relationship between the greenhouse effect and SST. As discussed by Valero et al. Ž1997b., the greenhouse effect Ž G . can be expressed as 4


Ž LL .

In Eq. ŽLL., ⌬ H vap is the heat of vaporization of water and R is the gas constant. Thus the vapor pressure of water has an exponential dependence on temperature. This suggests that there may be a water vapor feedback associated with global climate change. If the atmosphere warms, for example due to increased greenhouse gases such as CO 2 , increased concentrations of gaseous water are expected in accordance with Eq. ŽLL.. The increased water vapor traps more thermal infrared radiation, warming the atmosphere further Že.g., Raval and Ramanathan, 1989; Stenchikov and Robock, 1995.. However, there may be additional feedbacks that limit what would otherwise be a runaway system. For example, Ramanathan and Collins Ž1991. proposed that there is a natural ‘‘thermostat’’ mechanism over the warm pool in the Pacific Ocean that limits the sea surface temperature ŽSST. from rising above 305 K. This mechanism consists of triggering deep convection when the SST exceeds ;300 K, resulting in the formation of thick anvil clouds. Reflection of solar radiation back to space then acts to cool this region, providing a negative feedback and acting like a thermostat. Consistent with this thermostat mechanism, Waliser et al. Ž1993. report that for SST between 299 and 303 K in the western Pacific, the frequency of highly reflective clouds and decreases in outgoing longwave radiation are correlated with the SST. Interestingly, between 303 and 305 K, these relationships reversed; i.e., the frequency of highly reflective clouds decreased and outgoing longwave radiation increased with SST. This hypothesis has been somewhat controversial. For example, evaporative cooling of the ocean surface ŽFu et al., 1992. and large-scale dynamical feedbacks ŽWallace, 1992. have also been suggested as being important in the feedback, suggesting that changes in cloud cover reflect changes in atmospheric circulation

G s ␴ Ž SST. y Fq,


where ␴ is the Stefan᎐Boltzmann constant Ž5.67 = 10y8 W my2 Ky4 ., ␴ ŽSST. 4 is the thermal emission from the surface Žsee Section A.1., and Fq is the outgoing radiation flux at the top of the atmosphere. The rate of change of thermal emission by the ocean surface with SST is given by 4 d w ␴ Ž SST. x

d Ž SST.


s 4␴ Ž SST. ,


which for an ocean temperature of 300 K, typical of the tropics, is 6.1 W my2 Ky1 . This is what one would expect for the increase in thermal emission from the ocean surface as the SST increases. If the outgoing flux at the top of the atmosphere remains constant, the rate of change of the greenhouse effect with SST should also be about 6.1 W my2 Ky1 . However, measured values in the tropics under clear skies exceed this value by a factor of about two, which has been dubbed the ‘‘supergreenhouse effect.’’ For example, Fig. 14.53 shows G as a function of SST for SST ) 300 K measured using airborne infrared

FIGURE 14.53 Measured values of the clear-sky greenhouse effect G wsee Eq. ŽMM.x using measured upwelling infrared irradiance at an altitude corresponding to 191 mbar as a function of sea surface temperature ŽSST. over the central equatorial Pacific for SST ) 300 K Žadapted from Valero et al., 1997b..


radiometers over the central equatorial Pacific to obtain Fq and a combination of satellite and in situ data to obtain SST ŽValero et al., 1997b.. It is clear that the greenhouse effect increases approximately linearly with SST. The slope of such plots gave values of dGrdŽSST. from 13.5 to 15.3 W my2 Ky1 , more than double the 6.1 W my2 Ky1 expected for the increase in the ocean surface thermal emission wEq. ŽNN.x. Furthermore, the extent of the area over which this occurred was quite large, about half of the tropical ocean between 20⬚N and 20⬚S. The absorbed energy is radiated back to the surface to further contribute to surface warming ŽValero et al., 1997b.. In addition to these feedbacks involving water vapor and clouds, others are expected involving ice and snow. These surfaces are highly reflecting so that if warming leads to increased exposure of the underlying, darker surfaces, further warming will occur to give a positive feedback.

3. Solar Variability Because it is the sun that drives the earth’s energy balance ŽFig. 14.2., even small variations in its output can significantly alter the earth’s climate. Orbital variations of the earth relative to the sun which resulted in changes in the geographical distribution of solar radiation and, to a lesser extent, small changes Ž-1%. in the annual and global average solar intensity are believed to have affected global climate over the approximately past million years. This is often referred to as the Milankovitch mechanism Že.g., see Imbrie and Imbrie, 1979; Crowley and North, 1991; Lindzen, 1994; and Bryant, 1997.. These solar variations have characteristic periodicities of ;20,000, 40,000, 100,000, and 400,000 years, respectively. These changes in solar insolation, and particularly in its geographical distribution, are expected to have affected climate by altering circulation patterns and heat transport in the atmosphere ŽLindzen, 1994.. On a much shorter time scale, the radiant energy from the sun, the ‘‘solar constant,’’ currently averages 1368 W my2 . However, there is natural variability around this mean due to bright solar faculae and dark sunspots. In particular, there is a solar cycle approximately 11 years in length that occurs with an amplitude for total irradiance changes of about 0.1%; it is sometimes treated in terms of a ‘‘Hale cycle’’ of ;22 years in length Že.g., see Wilson, 1998.. The variation in intensity during the solar or Hale cycles is not constant across the solar spectrum, but is larger in the UV Že.g., see Lean et al., 1995b, 1997.. In addition, changes in solar output are modified by the atmosphere before reaching the earth’s surface, so that the magnitude of


FIGURE 14.54 Annual average number of sunspots from 1880 to 2000, showing the 11-year cycle Žadapted from Cliver et al., 1998..

changes due to the solar cycle depends on wavelength, latitude, and altitude ŽHaigh, 1994.. This may have indirect effects on the troposphere by altering stratospheric chemistry Žsee discussion by Robock Ž1996. and references therein .. Figure 14.54, for example, shows the annual average number of sunspots from 1880 to the present, which clearly shows this cycle ŽCliver et al., 1998.. Both the sunspot number and the aa geomagnetic index have been used as proxies for the solar cycle. For the relatively short time period covered by available instrumental temperature records, both the sunspot number and the aa geomagnetic index are correlated to surface temperature Že.g., see Cliver et al., 1998; and Wilson, 1998.. However, there is increasing evidence that longer term solar variations are measurable over the past few centuries as well, and understanding these is very important for discerning anthropogenic effects on global climate Že.g., Lean et al., 1995a; Willson, 1997.. Figure 14.55, for example, shows a reconstruction of total solar irradiance from 1610 to the present, in which the 11-year cycle and a component having much longer term variability are both included ŽLean et al., 1995a.. This two-component model is consistent with the Maunder Minimum that occurred during the years from 1645 to 1715 ŽEddy, 1976.. During this period, the 11-year cycle did not occur for a number of decades, which was also the coldest period in the ‘‘Little Ice

FIGURE 14.55 Reconstructed total solar irradiance from 1610 to 1995 using an 11-year solar cycle plus a longer term component of variability Žadapted from Lean et al., 1995a..



Age’’ from 1450 to 1850. The long-term component has been scaled to agree with the estimate of an overall increase in total irradiance from the Maunder Minimum to the present of 0.24% ŽLean et al., 1992, 1995b.. Lean et al. Ž1995a. have used this reconstruction to estimate how much of the increase in surface temperatures in the Northern Hemisphere can be explained due to solar variability. They concluded that about half of the observed Northern Hemisphere surface temperature increase of 0.55⬚C since 1860 is due to solar variability but that it only accounts for about a third of the 0.36⬚C increase since 1970. Similar conclusions have been reached by a number of researchers Že.g., see Kelly and Wigley, 1992; Schlesinger and Ramankutty, 1992; Scuderi, 1993; Crowley and Kim, 1996; Solanki and Fligge, 1998; Cliver et al., 1998; and Wilson, 1998.. However, Frohlich and Lean Ž1998. have ¨ reexamined the solar irradiance record since 1978; they concluded that the irradiances in 1986 and 1996 were similar and that changes in the solar flux during this period could not have contributed significantly to the observed changes in global mean surface temperature. In short, it is clear that variations in solar output have played a major role in determining the earth’s climate in the past, and understanding and quantifying this variability are critical for understanding anthropogenic influences on global climate. The observed temperature increases over approximately the past three decades are larger than expected from solar variability and have been interpreted by many researchers in this field to be the first signs of anthropogenic perturbations on climate.

4. Volcanic Eruptions As discussed in Section C.1a, major volcanic eruptions have been observed to alter the earth’s climate through injection of large amounts of SO 2 into the stratosphere. There it is oxidized to sulfate particles that scatter incoming solar radiation, leading to cooling at the earth’s surface. These particles also absorb long-wavelength terrestrial infrared radiation, warming the stratosphere ŽFig. 14.30.. While this absorption of infrared increases the downward emission of infrared from the stratosphere into the troposphere, i.e., causes a positive radiative forcing, the effect is much smaller than the direct scattering of solar radiation. As a result, the major overall net effect of volcanic eruptions is cooling ŽRobock and Mao, 1995.. However, it should be noted that the effect is somewhat geographically and temporally variable. For example, Robock and Mao Ž1995. have examined climate records since about 1850 and correlated them to volcanic eruptions both before and after removal of the

effects of the El Nino᎐Southern Oscillation ŽENSO. ˜ signal. While the effect of volcanic eruptions on the global mean surface temperature is cooling, there are circumstances where the effect is not only smaller than the mean, but warming was observed. For example, in the first winter following a number of different volcanic eruptions, the Northern Hemisphere and Eurasia on average warmed, in contrast to northern Africa and southwestern Asia, which cooled. Robock and Mao propose that the warming is due to changes in the winter circulation pattern, associated with an enhanced polar vortex, which lowers the extent of normal winter cooling. Figure 14.56 demonstrates the overall cooling effect of volcanic eruptions in the Northern Hemisphere over the past six centuries, reconstructed using the effects of temperature on tree ring densities ŽBriffa et al., 1998.. The relationship between the average summer monthly mean land and marine temperatures in the Northern Hemisphere and tree ring density was determined for the period from 1881 to 1960. This was then applied to measured tree ring densities to obtain the temperature anomalies for the entire 600 years compared to the 1881᎐1960 period. Some of the major volcanic eruptions are also marked on the diagram, clearly demonstrating their association with significant cooling. Similar conclusions have been reached using ice core data Že.g., White et al., 1997; Clausen et al., 1997; Taylor et al., 1997; Zielinski et al., 1997.. Volcanic eruptions provide an opportunity for testing not only our current understanding of the direct effects of aerosol particles due to backscattering but also the sensitivity of the climate system to such perturbations. Thus, after the initial short-term effects on temperature, the coupled atmosphere᎐ocean᎐land system responds on a longer time scale through a complex set of feedback mechanisms. As discussed by Lindzen and Giannitsis Ž1998., the effects of multiple volcanic eruptions should provide a better test of our understanding of such feedbacks than is provided by a single eruption.

5. Oceans Oceans have an enormous effect on climate through many different mechanisms that are beyond the scope of this book. Globally, oceans absorb heat and greenhouse gases such as CO 2 from the atmosphere ŽFig. 14.11., both moderating such changes Že.g., Schneider et al., 1997; Bush and Philander, 1998. and providing a time lag in the response to atmospheric perturbations Že.g., Wigley, 1995.. Other phenomena such as the El Nino᎐Southern Oscillation ŽENSO. and the North At˜ lantic Oscillation ŽNAO. clearly also have substantial



FIGURE 14.56 Temperature anomalies in the Northern Hemisphere compared to the 1881᎐1960 mean, calculated based on measured changes in tree ring densities. The 95% confidence limit is ; "0.3⬚C. The line shows bidecadal smoothed levels. Arrows on the lower axis mark some of the major volcanic eruptions Žadapted from Briffa et al., 1998..

large-scale impacts on climate Že.g., see Enfield and Mayer, 1997; and Hurrell and Van Loon, 1997.. While the atmosphere and oceans are closely linked, how changes in one impact changes in the other is not as clear Že.g., Broecker and Denton, 1990.. The oceans have well-documented circulation systems that play a major role in determining climate. For example, the Atlantic thermohaline circulation, often referred to as the ‘‘conveyor,’’ consists of a complex combination of ocean currents that result in the transport of warmer surface waters from the North Pacific into the Indian Ocean, around the African continent, and into the northern Atlantic ŽBroecker, 1997.. This provides a source of heat to air masses moving east in the winter, resulting in much warmer winters in Europe than would otherwise be the case. As discussed by Broecker Ž1997., this conveyor appears in the past to have jumped from one mode of operation to another, initiating substantial and rapid global climate changes. Furthermore, he suggests that it is possible that substantial increases in greenhouse gases such as CO 2 may also initiate such changes in the thermohaline circulation, with associated, and perhaps surprising, effects on global climate. For example, increased water vapor at high latitudes and the associated increased precipitation, combined with melting glaciers, due to global warming could provide a layer of less dense surface water in the northern Atlantic. Since the conveyor is driven by high-density salt water, this could shut down this global

ocean circulation system. Such a shutdown is expected to lead to cooling in the Northern Hemisphere but warming in the Southern Hemisphere since the heat transport associated with the conveyor no longer occurs ŽKerr, 1998.. That is, anthropogenic emissions that one normally associates with greenhouse warming may trip the ocean᎐atmosphere system in such a way that cooling could result in the Northern Hemisphere and warming in the Southern Hemisphere. Such feedbacks, hypothesized to have been triggered by closing of the Panamanian Isthmus, have been postulated to explain the Northern Hemisphere glaciation that occurred about 3 million years ago ŽDriscoll and Haug, 1998.. Clearly, this is an area that needs to be explored further.

E. OBSERVATIONS OF CLIMATE CHANGES 1. Observed Temperature Trends a. Trends o© er the Past Century One of the obvious manifestations of anthropogenic emissions is expected to be an increase in the temperature of the air and sea surface ŽSST.. As a result, there have been many analyses of such temperatures, for which there are substantial records based on instrumental measurements made in a number of locations



back to approximately 1860 and in at least one location, Armagh Observatory, North Ireland, to 1795 ŽWilson, 1998.. A review of these data, as well as more limited temperature data at higher altitudes in the troposphere and stratosphere, is found in Bradley and Jones Ž1993. and IPCC Ž1996.. Figure 14.57 shows the globally averaged temperature anomalies for land and sea surface measurements from 1861 to 1994, relative to the 1961᎐1990 period ŽIPCC, 1996; Jones et al., 1994.. Such data indicate there has been an increase in near-surface temperatures of ;0.3᎐0.6⬚C over this period, with an uncertainty of about 0.15⬚C. Measurements of underground temperatures from 358 boreholes in central Europe, southern Africa, Australia, and eastern North America show a similar temperature trend ŽPollack et al., 1998.. However, the increase has not been continuous, with substantial increases in temperature occurring between about 1920 and 1940, followed by a decrease and then an increase to the present time. The increase in the global average temperature over the past 40 years has been about 0.2᎐0.3⬚C. Similarly, the changes in temperature vary geographically and seasonally. For example, warming has occurred in the Northern Hemisphere over the continents, while cooling has occurred over the midlatitude North Pacific and over the northwestern Atlantic. The geographical and seasonal dependencies are summarized in IPCC Ž1996.. While the surface temperatures have clearly been increasing, some satellite measurements have suggested that the air temperatures in the troposphere have been cooling at altitudes where this was not expected. However, this is controversial Že.g., Pielke et al., 1998a, 1998b.. For example, Wentz and Schabel Ž1998. have shown that the loss of satellite altitude with

time can introduce an artifact into the data, which, if not corrected, leads to artifact cooling trends. The five warmest years for which there are surface temperature records have all been since 1990 ŽJones et al., 1998., with the most recent year for which there are data Žat the time of writing., 1997, being the warmest in the past century Žsee Kerr, 1998, and references therein .. Mann et al. Ž1998. have used a variety of indirect indicators for temperature Že.g., ice core data; see later. over the past 600 years in the Northern Hemisphere and report that mean annual temperatures for three of the eight years up to and including 1995 are higher than any since 1400 A.D. An interesting aspect of the surface temperature changes is that in many locations, particularly continental regions, the minimum daily temperature has increased more than the maximum daily temperature Že.g., Hansen et al., 1997d.. As a result, the daily temperature range has decreased in these regions. Figure 14.58, for example, shows globally averaged maximum and minimum temperatures, as well as the diurnal temperature range, from 1950 to 1993 based on approximately 4100 nonurban stations ŽEasterling et al., 1997.. These are expressed as deviations from the mean for all stations in 5⬚ = 5⬚ latitude᎐longitude grid boxes during the period from 1961 to 1985. The trend in the maximum temperature is 0.82⬚C per century, but that in the minimum is larger, 1.79⬚C per century. As a result, the diurnal temperature range decreases, with a slope of y0.79⬚C per century.

FIGURE 14.58 Global nonurban annual average temperature FIGURE 14.57 Global average temperature anomaly Ž ⌬T . for land and sea surface measurements relative to the period from 1961 to 1990 Žadapted from IPCC, 1996..

anomalies for the Ža. maximum temperature, Žb. minimum temperature, and Žc. diurnal range of temperatures from 1950 to 1993 for ;4100 stations in both the Northern and Southern Hemispheres Žadapted from Easterling et al., 1997..


These trends vary, depending on location. For example, the diurnal temperature range did not decrease over mid-Canada or parts of southwest Asia, southern Africa, the interior of Australia, the western tropical Pacific Islands, and Europe ŽEasterling et al., 1997.. Similarly, in some European mountain locations both the minimum and maximum of daily temperatures have been observed to increase ŽWeber et al., 1994. while in India, the maximum increased but there was no trend in the minimum ŽKumar et al., 1994.. There are many possible reasons for the decrease in the diurnal temperature range, where it occurs ŽKukla and Karl, 1993.. These include the urban heat island effect, which is strongest at night. However, the data in Fig. 14.58 excluded population centers of 50,000 or more; including data from urban areas leads to a slight increase in the slopes of maximum, minimum, and diurnal range of temperatures. Another potential contributor is an increase in cloudiness, which during the day scatters incoming solar radiation and leads to cooling. At night, the ground cools by thermal emission of infrared, which is counterbalanced in part by downward thermal infrared emission from atmospheric constituents. The downward emission is greatest when clouds and higher water vapor concentrations are present, leading to less net cooling of the surface at night and an increase in the minimum temperature. Soil moisture, which is affected by irrigation, drying of wetlands, deforestation, etc., is another factor. Evaporative cooling in moist soils occurs in the afternoon, but moist soils are warmer at night, leading to a decreased diurnal temperature range. Finally, anthropogenic aerosol particles and greenhouse gases may also affect the diurnal temperature range. As discussed in detail earlier, aerosol particles cool during the day by scattering incoming solar radiation. At night, they can contribute to warming by absorbing terrestrial infrared radiation, but this effect is small relative to the daytime cooling effect. Greenhouse gases such as CO 2 , of course, absorb the terrestrial infrared, leading to warming, an effect that does not vary as strongly with time of day as do those involving solar radiation. As a result, increased CO 2 leads to increased heating without a strong diurnal variation, whereas aerosol particles lead primarily to cooling during the day. These two effects can lead to a reduced diurnal temperature range, although if these were the only two effects operating, one would expect the reduced temperature range to be due more to changes in daytime temperatures. This is the opposite of what has been observed in many locations ŽFig. 14.58.. Stenchikov and Robock Ž1995. suggest that feedbacks may actually be more important than these direct


effects in reducing the diurnal temperature range. For example, in a warmer climate, more evaporation of water occurs, leading to increased gas-phase water vapor concentrations and possibly increased clouds. Stenchikov and Robock Ž1995. suggest that the major effect of such feedbacks in reducing the diurnal temperature range is not through the usual greenhouse effect but rather through increased absorption of solar radiation in the near-IR by the increased atmospheric water. Hansen et al. Ž1995, 1997d. have modeled various contributions to changes in the diurnal temperature range and the increase in global temperatures and concluded that the observed changes are only consistent with a combination of factors. These include a contribution from direct forcing by anthropogenic aerosols and an increase in cloud cover Žthe indirect effect of aerosols discussed earlier. primarily over continental regions, which are about 50% of the magnitude of anthropogenic greenhouse gas global forcings, but in the opposite direction. As discussed throughout this chapter, there are a variety of anthropogenic emissions that are expected to lead to warming Že.g., the greenhouse gases. or to cooling Že.g., increased aerosol particle concentrations .. To quantify such contributions to changes in surface temperatures, it is necessary first to understand and account for changes due to natural processes such as the solar variability discussed earlier. How to do so in an accurate and reliable manner remains a very complex and controversial area Že.g., Lindzen, 1994; Jones, 1995; Mahasenan et al., 1997; Legates and Davis, 1997; Wigley et al., 1997.. However, the weight of evidence at the present time suggests that the observed recent increase in global mean surface temperature is in part due to anthropogenic influences Že.g., see IPCC, 1996; Santer et al., 1996; Overpeck et al., 1997; Kaufmann and Stern, 1997; and Mann et al., 1998.. b. Temperatures and Other Proxies for Climate Change o© er the Past ;10 5 Years One approach to elucidating the contribution of natural variability to recent temperature trends is to examine markers for temperature over much longer time scales, prior to the industrial revolution. A major source of such data is ice cores Žsee also Section B.2a.. These ice cores provide a record of climate and atmospheric composition for at least 110,000 years, for which there is agreement among various studies. Data are available for 250,000 years before the present Žbp., but there is some uncertainty in the dating of the layers corresponding to these older ice core depths ŽChappellaz et al., 1997..



FIGURE 14.59 Schematic diagram of uptake and incorporation

FIGURE 14.60 Relationship of ice core depth to years before the

of gases and aerosol particles into ice cores Žadapted from Delmas, 1992..

present Žbp. and to the type of ice for the Central Greenland Ice Sheet. The Holocene and Wisconsin periods are also marked Žadapted from Gow et al., 1997..

Figure 14.59 shows schematically the processes by which gases and particles are trapped from the atmosphere into snow and ice at high latitudes ŽDelmas, 1992.. As snow is deposited, the surface is initially quite porous. As more snow accumulates, it compacts the underlying layers, forming a porous structure known as firn. Atmospheric gases continue to penetrate the porous firn. At some depth Žtypically ;100 m., there is a close-off zone in which recrystallization starts to seal off the pores, trapping the atmospheric constituents in bubbles in the ice. At larger depths, the bubbles are completely sealed off and the trapped gas is preserved. As snow accumulation continues, the ice below is further compacted and the ice sheet spreads down and out. At depths larger than ;1200 m, the hydrostatic pressure is sufficiently large that the air is forced into the ice to form clathrates so that distinct bubbles are no longer evident ŽMiller, 1969.. Clathrates are solid ice lattices that incorporate another molecule such as CO 2 or CH 4 into their crystal lattice Že.g., see Kvenvolden, 1993.. Once an ice core is drilled, various depths can be dated and the air trapped in the ice as either bubbles or clathrates is recovered for analysis, for example, by crushing the ice sample ŽWilson and Long, 1997.. The bubbles found at larger depths therefore correspond to older atmospheres. Figure 14.60 shows the relationship between the ice core depth and age Žin

number of years before the present time, bp. as well as the characteristics of the ice for samples from the Greenland Ice Sheet Project 2, GISP2 ŽGow et al., 1997.. Also shown are the periods corresponding to the Holocene Žthe past 10,000 years. and the Wisconsin ice age, for which some data are shown below. Ice at depths of ;260, 1000, 2430, and 2759 m corresponds to ages of 1000, 5100, 50,000, and 103,000 years bp, respectively ŽGrootes and Stuiver, 1997.. Because the firn is ventilated by atmospheric air while the bubbles are forming over a period of time and ice depths, the air eventually trapped in the bubbles is a time-integrated sample that is younger than the snow deposit itself. For example, in one recent study ŽSmith et al., 1997., the air bubbles were, on average, 220᎐700 years younger than the ice in which they were embedded, but the difference can be as much as several thousand years Že.g., see Rommelaere et al., 1997.. These exchange processes with the atmosphere, gas diffusion, and the porosity and tortuosity of the ice pores have to be taken into account in relating the depth of the core to the age of the trapped air. As seen earlier ŽSection B.2., air trapped in these ice cores can be recovered and analyzed to provide a snapshot of the composition of the atmosphere tens or even hundreds of thousands of years ago Žbut note cautions with respect to potential artifacts, e.g., in situ formation of CO 2 from carbonate in the bubbles.. In



addition, ice core composition can be used to infer the local temperature of the atmosphere when the snowrice was deposited, so-called paleothermometry ŽDelmas, 1992.. The isotopic composition, particularly the 18 Or16 O and DrH ratios, of the ice is related to the temperature at the level of the precipitating cloud that generated the snowrice. Isotopic fractionation occurs during the natural water cycle and this leads to a relationship between the isotopic composition and the precipitation temperature ŽDansgaard, 1964.. Once this relationship is established, the isotopic composition of water in the ice core can be used to estimate the corresponding atmospheric temperature. Such relationships, using 18 O as an example, are usually expressed in the form T s a ␦ 18 O q b,

Ž OO .

where ␦ 18 O in per mil Ž%. is defined as 1000Ž R y R 0 .rR 0 , R is the isotope ratio of the sample, and R 0 is the ratio of a standard sample. In the case of oxygen, the standard is usually standard mean ocean water ŽSMOW.. The values of the slope and intercept Ž a and b, respectively. vary from location to location and likely with time as well ŽCuffey and Clow, 1997; Jouzel et al., 1997.. However, it appears that such relationships are still useful for inferring historic temperatures ŽJouzel et al., 1997; Salamatin et al., 1998.. Ice core data have provided evidence that quite rapid and large oscillations in climate have occurred over the period of record. Figure 14.61, for example, shows the temperature changes in central Greenland for the past ;110,000 years. Recent millenia are characterized by relatively small rates of temperature change. Indeed, in summarizing the results of the Greenland Summit Ice Core Projects ŽGISP2 and GRIP. published in a special issue of the Journal of Geophysical Research ŽVol. 102 ŽC12., pp. 26315᎐26886, November 30, 1997., Hammer, Mayewski, Peel, and Stuiver state The ice-core records tell a clear story: humans have come of age agriculturally and industrially in the most stable climate regime of the last 110,000 years. However, even this relatively stable period is marked by change . . . wwhich isx more characteristic of the Earth’s climate than is stasis.

Figure 14.61 illustrates the much larger changes, more than 20⬚C from one extreme to the other, that have occurred historically when viewed over these long time periods. Associated with decreases in temperature are decreased methane concentrations, increased dust loadings, and decreased snow accumulations Že.g., see Fig. 14.62.. ŽIt should be noted that the amplitude of the temperature changes can vary from site to site; for example, Dahl-Jensen et al. Ž1998. showed that the

FIGURE 14.61 Calculated temperature changes over the past 100,000 years from the European Greenland Ice Core Program ŽGRIP. Žadapted from Jouzel et al., 1997..

amplitude of the temperature record from the Dye 3 site in Greenland 865 km south of the GRIP site and 730 m lower in elevation was 50% larger than that at the GRIP site.. Some of the indicators of climate change have shown very rapid changes, over decades or less, during particular time periods in the past. For example, Fig. 14.62 shows ice core measurements of Ža. electrical conductivity, Žb. snow accumulation rate, and Žc. concentration of calcium at the start of the Holocene period some 11,600 years ago ŽTaylor et al., 1997.. Electrical conductivity is a measure of the acidity of the ice, since Hq is the major charge carrier; the direct current is measured between two electrodes that have a voltage difference of several thousand volts. As seen in Fig. 14.62a, such measurements provide excellent spatial and hence time resolution. The transition from the Wisconsin to the Holocene period is seen to be characterized by an increase in electrical conductivity of the ice core and a decrease in calcium. The two are inversely related since high concentrations of CaCO 3 neutralize strong acids, decreasing the conductivity. The rate of accumulation of snow ŽFig. 14.62b. also increases. These changes occurred in less than about two decades. Taylor and co-workers also point out that the data in Fig. 14.62 suggest there is a ‘‘flicker’’ just prior to the rapid transition to the alternate climate state.



FIGURE 14.62 Evidence of rapid climate changes at the start of the Holocene period ;11,600 years before the present Žbp.: Ža. electrical conductivity; Žb. rate of accumulation of snow; Žc. calcium concentration. The depth of the ice cores is shown on the bottom axis and the corresponding age in years before the present on the top axis Žadapted from Taylor et al., 1997..

In short, ice core and other long-term records show that there have been dramatic climate changes in the past, some of them within or shorter than a typical human life span. Separating out such natural variability from anthropogenic perturbations remains a major challenge, particularly when it is possible that the anthropogenic emissions may act to hasten or ‘‘jolt’’ the climate system into a relatively rapid transition from one state to another.

2. Other Climate Changes

with multiple feedbacks between them. Some possible future scenarios based on the state of the science as of about 1995 are described in the IPCC Ž1996. document. Figure 14.63 shows one model estimate for temperature changes due only to direct radiative forcing by CO 2 from 1990 to 2100 based on three scenarios ŽWigley, 1998.. While as we have seen, many gases contribute to radiative forcing, the calculations in Fig. 14.63 are based on expressing these changes in terms of equivalent CO 2 reductions. Note that this does not take into account other contributing factors such as aerosol particles which may contribute in the opposite

There are a variety of other climate changes that might be expected to occur simultaneously with changes in temperature. These include changes in precipitation, an increase in the mean sea level, and more variability in the climate. As discussed in detail in IPCC Ž1996., changes in precipitation patterns and cloudiness have been noted over the past approximately four decades and there is evidence that the sea level has risen by ;10᎐25 cm. The IPCC document should be consulted for detailed evidence for these effects and their possible relationship to anthropogenic perturbations.

F. THE FUTURE As seen throughout this chapter, the parameters controlling climate are extremely varied and complex,

FIGURE 14.63 Calculated temperature changes relative to 1990 for existing policies Ž ᎏ ., for a 5% decrease in equivalent CO 2 as required by Kyoto protocol from 1990 to 2010 followed by no further emissions reductions Ž ⭈⭈⭈ ., and for further reductions of 1% per year Žcompounded. from 2010 to 2100 Ž ᎐ ᎐ ᎐ . assuming a climate sensitivity of 2.5 K for a doubling of CO 2 Žadapted from Wigley, 1998..


direction. Hence, Fig. 14.63 should be taken as illustrative of the effects on direct radiative forcing by gases and not the net result of all contributing factors and feedbacks. Some of complexities of climate are reflected in the natural variability, observed through such ‘‘proxies’’ as the ice core and tree-ring records. However, there are reasons, based on our current state of scientific understanding, to believe that anthropogenic activities that have resulted in changes in atmospheric composition may affect climate locally, regionally, and globally. Despite the uncertainties in the deconvolution of anthropogenic and natural contributions to the relatively recent observed global climate changes, the weight of scientific opinion at the end of the twentieth century is reflected in the IPCC Ž1996. summary: Nevertheless, the balance of evidence suggests that there is a discernible human influence on global climate.

Elucidating this influence remains a challenge for the twenty-first century.

G. PROBLEMS 1. Using gas kinetic molecular theory, show that under typical atmospheric conditions of pressure and temperature corresponding to an altitude of 5 km Žsee Appendix V. collisional deactivation of a CO 2 molecule will be much faster than reemission of the absorbed radiation. Take the collision diameter to be 0.456 nm and the radiative lifetime of the 15-␮ m band of CO 2 to be 0.74 s ŽGoody and Yung, 1989.. 2. Assume that a typical greenhouse gas absorbs infrared radiation at 10 ␮ m resulting in a vibrational energy change in the molecule. If there were a temperature change in the upper troposphere Žwhere the temperature is typically 220 K. of q5⬚C, estimate using the Boltzmann equation the factor Ži.e., ratio. by which the emission of energy out to space would change. 3. Calculate for liquid water the factor by which the vapor pressure increases over droplets of the following sizes compared to that over the bulk liquid at 298 K: Ža. 1, Žb. 0.1, Žc. 0.01, and Žd. 0.001 ␮ m. If typical supersaturations of water vapor in the atmosphere are of the order of 0.1%, which of these could be stable in the atmosphere? The surface tension of water at room temperature is 72 dyn cmy1 . 4. Calculate for sulfuric acid Ž␥ s 55.1 dyn cmy1 , ␳ s 1.84. the percentage increase in vapor pressure compared to the bulk over droplets of the following sizes at 298 K: Ža. 1, Žb. 0.1, Žc. 0.01, and Žd. 0.001 ␮ m. The vapor pressure of sulfuric acid is sufficiently low


that it exists in the atmosphere primarily as a liquid, although there will be very small concentrations in the gas phase. By what percentage will the gas-phase concentration increase over that of the bulk liquid if H 2 SO4 exists primarily in 10-nm particles? 5. Derive an expression for r 0 in Fig. 14. 38, corresponding to an RH of 100% or a supersaturation of zero, in terms of a and b in Eq. ŽII.. 6. Derive expressions for the critical radius rc and the critical supersaturation Sc at the peak of the Kohler ¨ curve in Fig. 14.38 in terms of the a and b parameters in Eq. ŽII.. What is the relationship between rc and r 0? 7. Calculate the critical radius rc and critical supersaturation Sc for activation into a cloud droplet of a 10y1 5-g NaCl particle. Assume the surface tension is 72 dyn cmy1 and the liquid density is that of water. 8. Repeat Problem 7 for a 10y1 6-g ŽNH 4 .SO4 particle. 9. Ža. Using Eq. ŽJJ., calculate the albedo of a 250-m-thick cloud with 5 = 10 7 droplets per m3 and an effective mean radius for light scattering of 10 ␮ m. Take g s 0.8. Žb. Now assume that the number of CCN have decreased sufficiently that the number of cloud droplets is 5 = 10 4 droplets my3 but the liquid water content is the same. Calculate the new effective radius and the cloud albedo. 10. As seen in Eq. ŽKK., cloud susceptibility depends on ␶ Ž ␦ Rr␦␶ .. Using Eq. ŽJJ., show that ␶ Ž ␦ Rr␦␶ . s RŽ1 y R ..

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