Energy Balance Of The Earth And Its Atmosphere

As the Earth sweeps through space on its annual revolution around the Sun, it intercepts a minute fraction of the energy emitted by this all-important star. Because the Earth is (approximately) spherical and rotates on an axis inclined (at present) 23.4° to the plane across which it moves around the Sun, (the ecliptic) energy receipts vary greatly from one part of the globe to another. Furthermore, the pattern of energy receipts is constantly changing. These differential energy receipts are the fundamental driving force of the atmospheric circulation. If solar output is assumed to be invariant, the spatial and temporal patterns of energy receipts impinging on the outer atmosphere can be calculated (Fig. 2.7; Newell and Chiu, 1981). However, for conditions near the surface of the Earth, the role of the atmosphere must be considered because it greatly diminishes potential solar radiation receipts. A consideration of energy exchanges in the Earth-atmosphere system also provides some insight into the potentially important factors involved in climatic variations and variability. For the system as a whole, energy receipts at the outer limits of the atmosphere during the course of a year are 342 W m"2 (Fig. 2.8).

As radiation penetrates the atmosphere, as a global average 77 W m"2 (23%) is either reflected from cloud tops or scattered upward by molecules and particulate matter in the air. Because the Earth's surface is also reflective, another ~9% of incoming solar radiation is returned to space without heating the atmosphere or the Earth's surface. A further 67 W m"2 (20%) is absorbed by ozone, by water vapor and water droplets in clouds, and by particulates, thereby raising the temperature of the atmosphere.2 Thus, only approximately half of the energy impinging on the outer atmosphere reaches the surface, where it is absorbed, increasing the surface temperature. Energy is re-radiated from the Earth's surface at longer wavelengths (terrestrial radiation), much of which is absorbed by water vapor and carbon dioxide in the atmosphere (the greenhouse effect). This is eventually re-radiated by the atmosphere and ultimately lost to space. Only -39% of the energy absorbed by the Earth's surface (66 W m"2) is lost by radiative emissions in this way. The balance, or net

2 Because the atmosphere absorbs short-wave solar radiation as well as long-wave radiation from the Earth, it also emits long-wave radiation both upward and downward (counter-radiation). Overall, however, there is a net loss of long-wave radiation (66 W nr2) from the Earth to space via the atmosphere.

FIGURE 2.7 Distribution of solar radiation at top of atmosphere (in Watt hours per square meter). Apparent position of Sun overhead at noon (declination) is shown by dotted line.

Redacted Solar Radiation 107 Win'

Incoming

Solar Radiation 342 W m1

■ Reflected try Clouds. \ Aerosol and . . A.Aln . ispr -

Emitted by/ i Atmosphere 165

I \ Absorbed by 67 Atmosphere

324 Sack Radiation

Oulooing

Lonuwave

Radiation

235 W m1

Atmospheric

^siGreen hoose j Gases

FIGURE 2.8 Mean annual radiation and energy balance of the Earth. Of the 168 W m"2 absorbed by the Earth's surface, there is a net radiative loss of 66 W m"2 (the net loss from long-wave emissions, less counter-radiation from the atmosphere to the surface).The balance of energy at the surface (the net radiation) is transferred by latent heat transfer (evaporation) and sensible heat transfer ("thermals" in this diagram) (from Kiehl andTrenberth, 1997).

radiation, is transferred to the atmosphere via sensible and latent heat transfers. Sensible heat flux (H) involves the transfer of heat directly from the surface to layers of air immediately adjacent to it by the processes of conduction and convection. Latent heat flux (LE) involves the transfer of heat from the surface via the evaporation of water; as water evaporates from the surface, latent heat is extracted, only to be released to the atmosphere later when the water condenses. This is the most important mechanism by which energy is transferred from the Earth to the atmosphere, accounting for ~46% (78 W m 2) of the incoming energy absorbed by the Earth's surface (Fig. 2.8). The relative importance of sensible and latent heat mechanisms in the transfer of heat from the Earth's surface is sometimes characterized by the Bowen ratio (H/LE); high values (>10) are typical of desert areas where values of latent heat flux are very low, whereas low Bowen ratios (<1) are typical of oceanic areas where most energy is transferred through the evaporation of water.

The global mean values for the energy balance provide a basis for appreciating the importance of a number of parameters in the climate system. Consider, for example, the role of cloudiness in global energy receipts. On a global scale approximately one-fifth of all energy entering the atmosphere is reflected by cloud tops as a result of their extremely high albedo. Small variations in global cloud cover, or even of cloud type, may thus have very large consequences for global energy balance but we have no clues from the paleoclimatic record as to how cloudiness may have varied through time on a global scale (Bradley et al., 1993). Albedo is of particular significance at the Earth's surface, and this is particularly apparent when zonal (latitude band) averages are considered (Fig. 2.9). The distribution of snow and ice

FIGURE 2.9 Latitudinal distribution of seasonal average surface albedo (averaging around latitude bands, i.e., zonally).-, estimates by Kukla and Robinson (1980); °-----estimates by Hummel and Reck (1979)

(Kukla and Robinson, 1980).

FIGURE 2.9 Latitudinal distribution of seasonal average surface albedo (averaging around latitude bands, i.e., zonally).-, estimates by Kukla and Robinson (1980); °-----estimates by Hummel and Reck (1979)

(Kukla and Robinson, 1980).

dominates this pattern (see Fig. 2.4) and is largely responsible for the large energy deficits at high latitudes (i.e., higher radiative losses than gains, accommodated by energy transfers from low latitudes). Only during the last 25 years have satellites provided a global perspective on snow- and ice-cover variations, both seasonally and interannually. Although the records are quite short, it is clear that variations in snow and ice extent from year to year can alter area-weighted hemispheric surface albedo by 3-4% (compare the interannual troughs, or peaks, in Fig. 2.10), which may influence atmospheric circulation in subsequent seasons, providing a positive feedback to the system (Groisman et al., 1994a,b). Over longer time periods, changes in surface albedo have been very large, and their effects on albedo must

30°N) and Southern Hemisphere ocean (south of 50°S) expressed as departures from the 1974-78 monthly means. Units in percentage albedo. Year ending in December is marked. Note opposite trends in the two hemispheres (from Kukla, 1979).

have been profound. Not only did continental ice sheets and more extensive sea ice (Table 2.5) increase global albedo but the more extensive deserts and savanna grasslands at the time of glacial maxima would have accentuated this effect.

The significance of atmospheric C02 and water vapor is also apparent from Fig. 2.8; these gases play a vital role in global energy balance because of their relative opacity to terrestrial radiation. An increase in C02 would reinforce this energy exchange, increasing atmospheric temperatures. However, many other interactions and consequences would also ensue and it is this complexity that makes forecasts of the climatic impact of C02 increases so difficult (Dickinson et al., 1996).

This thumbnail sketch of the radiation balance of the Earth-atmosphere system is very much a simplification of reality. Most importantly, there are large regional differences in values of net radiation and of latent and sensible heat flux due to the geography of the earth (distribution of continents and oceans, surface relief, vegetation, and snow cover) and the basic climatic differences from one region to another (principally variations in cloud cover and type) (Budyko, 1978). This is readily apparent from a consideration of annual energy balance components for the Earth's surface, shown as zonal averages in Table 2.6, and mapped in Figs. 2.11-2.13. Net radiation varies from near zero at high latitudes to >140 kcal cm'V1 (186 W m 2) over parts of the tropical and equatorial oceans (Fig. 2.11). On the continents, net radiation is lower than the zonal average due to higher albedo of thé surface (e.g., in desert regions) or because of higher cloud amounts, which reduce surface radiation receipts (Table 2.6). For the Earth as a whole (Table 2.6, bottom line) 84% of net radiation is accounted for by latent heat expenditures (66 of 79 kcal cm 2 a"1, or 88 of 105 W rrr2). If we just consider the oceans, however, 90% of net radiation is utilized in evaporation compared to only 54% (27 of 50 kcal cm"2 a"1, or 36 of 66 W m"2) on

TABLE 2.5 Maximum Extent of Land-Based Ice Sheets During the Pleistocene

TABLE 2.5 Maximum Extent of Land-Based Ice Sheets During the Pleistocene

North America

16.22

Greenland

2.30

Europe

7.21

Asia

3.95

South America

0.87

Australasia

0.03

Antarctica

13.81

From Flint (1971) and Hollin and Schilling (1981).

Note that not all areas experienced maximum ice cover at the same time during the Pleistocene. It is therefore not appropriate to total these values. Also, seasonal snow cover and sea-ice extent are not included, so these figures represent minimum changes in the area of the overall cryo-sphere (see Tables 2.3 and 2.4).

From Flint (1971) and Hollin and Schilling (1981).

Note that not all areas experienced maximum ice cover at the same time during the Pleistocene. It is therefore not appropriate to total these values. Also, seasonal snow cover and sea-ice extent are not included, so these figures represent minimum changes in the area of the overall cryo-sphere (see Tables 2.3 and 2.4).

TABLE 2.6 Mean Latitudinal Values of the Heat Balance Components of the Earth's Surface (W m"2)0

Latitude

R

Land LE

P

R

Ocean LE

P

R

Earth LE

P

70-60°N

29

21

8

30

41

29

-40

29

27

15

-12

60-50

42

30

12

57

62

25

-31

49

44

17

-12

50-40

60

33

27

85

89

21

-25

72

60

24

-12

40-30

77

31

46

119

127

19

-27

101

86

31

-16

30-20

85

25

60

147

145

9

-7

125

100

28

-3

20-10

98

42

56

161

155

97

-4

145

126

21

-3

10-0

105

76

29

165

138

9

17

151

123

13

15

0-10°S

105

81

24

169

131

8

29

154

119

12

23

10-20

100

60

40

162

150

12

0

149

130

19

0

20-30

94

37

57

145

141

15

-11

133

117

24

-8

30-40

82

38

44

122

109

15

-1

117

101

19

-3

40-50

58

29

29

96

68

8

20

94

66

9

19

50-60

46

29

17

61

46

12

3

61

46

12

3

Earth as a whole6

66

36

30

121

109

12

0

105*

88*

17*

From Budyko (1978).

"R is the radiative flux of heat (radiation balance of the Earth's surface) equal to the difference of absorbed shortwave radiation and the net long-wave radiation outgoing from the Earth's surface; LE is the heat expenditure for evaporation (L is the latent heat of vaporization, E is the rate of evaporation); P is the turbulent flux of heat between the Earth's surface and the atmosphere; Fg is the heat income resulting from heat exchange through the sides of the vertical column of a unit section going through the Earth's surface with the ambient layers.

b Values for the earth as a whole are slightly different from those given in Fig. 2.8, which is based on more recent satellite-derived data (Kiehl and Trenberth, 1997). These recent estimates give R = 102, LE = 78, and P = 24, so the zonal average values given here will no doubt require some revision. Nevertheless, the broad patterns depicted in this table will not change substantially, and the values are likely to be correct to within ± 10%.

the continents. In fact, in extremely arid areas, latent heat transfer may account for only 15-20% of the net radiation (see Figs. 2.11 and 2.12). In those areas, sensible heat flux is of primary importance (Fig. 2.13). For the continents as a whole, 46% of net radiation is utilized in sensible heat transfers. Over the oceans, sensible heat flux is only important at high northern latitudes where northward-flowing currents bring warm water into contact with cold polar air masses (see Fig. 2.13). Ocean currents themselves play a very important role in energy transport, as is clear from column 8 in Table 2.6. "Excess" heat is transferred from equatorial and tropical regions to higher latitudes where the energy thereby made available may even exceed net radiation at the surface (e.g., 60-70° N; see Figs. 2.11 and 2.13).

From this overview of the energy balance of different regions it is only a short step to consider how components of the energy balance of some areas may have varied in the past, and how human activities may affect the energy balance of some

FIGURE 2.1 I Radiative balance (net radiation, RJ of the Earth's surface (in kcal cm"2 yr"'). Note discontinuities at ocean/land boundaries (from Budyko, 1978).
FIGURE 2.12 Expenditure of latent heat for evaporation (latent heat flux, I, in kcal cm"2yr') (from Budyko, 1978).
FIGURE 2.13 Sensible heat flux between Earth's surface and the atmosphere (in kcal cm"2yr') (from Budyko, 1978).

areas in the future. Of course, it will only be possible to do this in a crude way because the energy balance of any one site is a function of a great many variables, including parts of the climate system far from the site in question. Nevertheless, some general points can be made. Consider, for example, the vast Saharan Desert region. At the present time, net radiation in this area averages -53 kcal cm 2 a"1 (70 W m~2) with a Bowen ratio of -8 (Table 2.7; Baumgartner, 1979). During the early to mid-Holocene, the area was wetter and supported a sparse grassland vegetation cover, increasing to savanna along the Sahelian margin to the south (Fabre and Petit-Maire, 1988; Lezine, 1989); if modern analogies are any guide, the area would have had a lower albedo, higher net radiation, and much lower Bowen ratio. Other desert regions also experienced similar changes in vegetation and hence in energy balance (though changes elsewhere were commonly greatest at the last glacial maximum). As

TABLE 2.7 Energy Balance for Different Surfaces (W nr2)

R

L

H

a

HIL

Tropical rainforest

110

85

25

13

0.3

Savanna

65

40

25

33

0.6

Desert

70

8

62

46

8.0

From Baumgartner (1979).

From Baumgartner (1979).

deserts and semideserts today occupy more than 10% of the continental area, such changes had major consequences for the energy balance of the world as a whole. It also seems likely that overgrazing and desertification of marginal environments today, as well as the destruction of tropical forest ecosystems, will bring about marked changes in the energy balance of low latitudes, with possible global consequences.

The energy balance changes associated with alterations in natural vegetation do not, of course, provide information on why the environmental changes occurred in the first place. However, they do provide important baseline data for computer models of the general circulation at particular time periods in the past (e.g., Kutzbach et al., 1993a) and point to potentially important feedbacks between the atmosphere and underlying surface once the vegetation changes have occurred. The development of a particular vegetation type may, in fact, bring about changes in the energy balance that would favor persistence of the "new" vegetation type (Charney et al, 1975; Foley et al., 1994).

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