Rption

10 20 40 60 100 400 1000 4 000 10.000

THICKNESS OF CLOUD (METRES)

10 20 40 60 100 400 1000 4 000 10.000

THICKNESS OF CLOUD (METRES)

Figure 3.6 Percentage of reflection, absorption and transmission of solar radiation by cloud layers of different thickness.

Source: From Hewson and Longley (1944). Reprinted with permission. Copyright © CRC Press, Boca Raton, Florida.

^ LATITUDE

Figure 3.7 The average receipt of solar radiation with latitude at the top of the atmosphere and at the earth's surface during the June solstice.

The effect of cloud cover also operates in reverse, since it serves to retain much of the heat that would otherwise be lost from the earth by long-wave radiation throughout the day and night. In this way, cloud cover lessens appreciably the daily temperature range by preventing high maxima by day and low minima by night. As well as interfering with the transmission of radiation, clouds act as temporary thermal reservoirs because they absorb a certain proportion of the energy they intercept. The modest effects of cloud reflection and absorption of solar radiation are illustrated in Figures 3.5 to 3.7.

Global cloudiness is not yet known accurately. Ground-based observations are mostly at land stations and refer to a small (~ 250 km2) area. Satellite estimates are derived from the reflected short-wave radiation and infra-red irradiance measurements, with various threshold assumptions for cloud presence/absence;

typically they refer to a grid area of 2500 km2 to 37,500 km2. Surface-based observations tend to be about 10 per cent greater than satellite estimates due to the observer's perspective. Average winter and summer distributions of total cloud amount from surface observations are shown in Figure 3.8. The cloudiest areas are the Southern Ocean and the mid- to high-latitude North Pacific and North Atlantic storm tracks. Lowest amounts are over the Saharan-Arabian desert area (see Plate 1). Total global cloud cover is just over 60 per cent in January and July.

4 Effect of latitude

As Figure 3.4 has already shown, different parts of the earth's surface receive different amounts of solar radiation. The time of year is one factor controlling this, more radiation being received in summer than in winter because of the higher altitude of the sun and the longer days. Latitude is a very important control because this determines the duration of daylight and the distance travelled through the atmosphere by the oblique rays of the sun. However, actual calculations show the effect of the latter to be negligible near the poles, due apparently to the low vapour content of the air limiting tropospheric absorption. Figure 3.7 shows that in the upper atmosphere over the North Pole there is a marked maximum of solar radiation at the June solstice, yet only about 30 per cent is absorbed at the surface. This may be compared with the global average of 48 per cent of solar radiation being absorbed at the surface. The explanation lies in the high average cloudiness over the Arctic in summer and also in the high reflectivity of the snow and ice surfaces. This example illustrates the complexity of the radiation budget and the need to take into account the interaction of several factors.

A special feature of the latitudinal receipt of radiation is that the maximum temperatures experienced at the earth's surface do not occur at the equator, as one might expect, but at the tropics. A number of factors need to be taken into account. The apparent migration of the vertical sun is relatively rapid during its passage over the equator, but its rate slows down as it reaches the tropics. Between 6°N and 6°S the sun's rays remain almost vertically overhead for only thirty days during each of the spring and autumn equinoxes, allowing little time for any large buildup of surface heat and high temperatures. On the other hand, between 17.5° and 23.5° latitude the sun's rays shine down almost

Figure 3.8 The global distribution of total cloud amount (per cent) derived from surface-based observations during the period 1971 to 1981, averaged for the months June to August (above) and December to February (below). High percentages are shaded and low percentages stippled.

Source: From London et al. (1989).

Figure 3.8 The global distribution of total cloud amount (per cent) derived from surface-based observations during the period 1971 to 1981, averaged for the months June to August (above) and December to February (below). High percentages are shaded and low percentages stippled.

Source: From London et al. (1989).

vertically for eighty-six consecutive days during the period of the solstice. This longer interval, combined with the fact that the tropics experience longer days than at the equator, makes the maximum zones of heating occur nearer the tropics than the equator. In the northern hemisphere, this poleward displacement of the zone of maximum heating is enhanced by the effect of con-tinentality (see B.5, this chapter), while low cloudiness associated with the subtropical high-pressure belts is an additional factor. The clear skies allow large annual receipts of solar radiation in these areas. The net result of these influences is shown in Figure 3.9 in terms of the average annual solar radiation on a horizontal surface at ground level, and by Figure 3.10 in terms of the average daily maximum shade temperatures. Over land, the highest values occur at about 23°N and 10-15°S. Hence the mean annual thermal equator (i.e. the zone of maximum temperature) is located at about 5°N. Nevertheless, the mean air temperatures, reduced to mean sea-level, are related very broadly to latitude (see Figures 3.11A and B).

5 Effect of land and sea

Another important control on the effect of incoming solar radiation stems from the different ways in which land and sea are able to profit from it. Whereas water has a tendency to store the heat it receives, land, in contrast, quickly returns it to the atmosphere. There are several reasons for this.

Figure 3.9 The mean annual global solar radiation (Q + q) (W m-2) (i.e. on a horizontal surface at ground level). Maxima are found in the world's hot deserts, where as much as 80 per cent of the solar radiation annually incident on the top of the unusually cloud-free atmosphere reaches the ground.

Figure 3.9 The mean annual global solar radiation (Q + q) (W m-2) (i.e. on a horizontal surface at ground level). Maxima are found in the world's hot deserts, where as much as 80 per cent of the solar radiation annually incident on the top of the unusually cloud-free atmosphere reaches the ground.

Source: After Budyko et al. (1962).

Figure 3.10 Mean daily maximum shade air temperature (C). Source: After Ransom (1963).
160° 140° 120° 100° 80° 60° 40° 20° 0° 20° 40° 60° 60° 100° 120° 140° 160° 180°

160° 140° 120° 100° 80° 60° 40° 20° 0° 20° 40° 60° 60° 100° 120° 140° 160° 160°

Figure 3.11 (A) Mean sea-level temperatures (°C) in January. The position of the thermal equator is shown approximately by the line dashes. (B) Mean sea-level temperatures (°C) in July. The position of the thermal equator is shown approximately by the line dashes.

160° 140° 120° 100° 80° 60° 40° 20° 0° 20° 40° 60° 60° 100° 120° 140° 160° 160°

Figure 3.11 (A) Mean sea-level temperatures (°C) in January. The position of the thermal equator is shown approximately by the line dashes. (B) Mean sea-level temperatures (°C) in July. The position of the thermal equator is shown approximately by the line dashes.

Figure 3.12 Average annual snow-cover duration (months). Source: Henderson-Sellers and Wilson (1983).

A large proportion of the incoming solar radiation is reflected back into the atmosphere without heating the earth's surface. The proportion depends upon the type of surface (see Table 3.2). A sea surface reflects very little unless the angle of incidence of the sun's rays is large. The albedo for a calm water surface is only 2 to 3 per cent for a solar elevation angle exceeding 60°, but is more than 50 per cent when the angle is 15°. For land surfaces, the albedo is generally between 8 and 40 per cent of the incoming radiation. The figure for forests is about 9 to 18 per cent according to the type of tree and density of foliage (see Chapter 12C), for grass approximately 25 per cent, for cities 14 to 18 per cent, and for desert sand 30 per cent. Fresh snow may reflect as much as 90 per cent of solar radiation, but snow cover on vegetated, especially forested, surfaces is much less reflective (30 to 50 per cent). The long duration of snow cover on the northern continents (see Figure 3.12 and Plate A) causes much of the incoming radiation in winter to spring to be reflected. However, the global distribution of annual average surface albedo (Figure 3.13A) shows mainly the influence of the snow-covered Arctic sea ice and Antarctic ice sheet (compare Figure 3.13B for planetary albedo).

The global solar radiation absorbed at the surface is determined from measurements of radiation incident on the surface and its albedo (a). It may be expressed as

where the albedo is a percentage. A snow cover will absorb only about 15 per cent of the incident radiation, whereas for the sea the figure generally exceeds 90 per cent. The ability of the sea to absorb the heat received also depends upon its transparency. As much as 20 per cent of the radiation penetrates as far down as 9 m (30 ft). Figure 3.14 illustrates how much energy is absorbed by the sea at different depths. However, the heat absorbed by the sea is carried down to considerable depths by the turbulent mixing of water masses by the action of waves and currents. Figure 3.15, for example, illustrates the mean monthly variations with depth in the upper 100 metres of the waters of the eastern North Pacific (around 50°N, 145°W); it shows the development of the seasonal thermocline under the influences of surface heating, vertical mixing and surface conduction.

A measure of the difference between the subsurfaces of land and sea is given in Figure 3.16, which shows ground temperatures at Kaliningrad (Königsberg) and sea temperature deviations from the annual mean at various depths in the Bay of Biscay. Heat transmission in the soil is carried out almost wholly by conduction, and the degree of conductivity varies with the moisture content and porosity of each particular soil.

165°W 120°W 60° W 0° 60" E 120°E 180°E

Figure 3.13 Mean annual albedos (per cent): (A) At the earth's surface. (B) On a horizontal surface at the top of the atmosphere.

Source: After Hummel and Reck; from Henderson-Sellers and Wilson (1983), and Stephens et al. (1981), by permission of the American Geophysical Union.

Figure 3.13 Mean annual albedos (per cent): (A) At the earth's surface. (B) On a horizontal surface at the top of the atmosphere.

Source: After Hummel and Reck; from Henderson-Sellers and Wilson (1983), and Stephens et al. (1981), by permission of the American Geophysical Union.

Air is an extremely poor conductor, and for this reason a loose, sandy soil surface heats up rapidly by day, as the heat is not conducted away. Increased soil moisture tends to raise the conductivity by filling the soil pores, but too much moisture increases the soil's heat capacity, thereby reducing the temperature response. The relative depths over which the annual and diurnal temperature variations are effective in wet and dry soils are approximately as follows:

Diurnal variation

Annual variation

Wet soil

G.5 m

9 m

Dry sand

G.2 m

3 m

Figure 3.14 Schematic representation of the energy spectrum of the sun's radiation (in arbitrary units) that penetrates the sea surface to depths of 0.1, 1, 10 and 100 m. This illustrates the absorption of infra-red radiation by water, and also shows the depths to which visible (light) radiation penetrates.

Source: From Sverdrup (1945).

Figure 3.14 Schematic representation of the energy spectrum of the sun's radiation (in arbitrary units) that penetrates the sea surface to depths of 0.1, 1, 10 and 100 m. This illustrates the absorption of infra-red radiation by water, and also shows the depths to which visible (light) radiation penetrates.

Source: From Sverdrup (1945).

Figure 3.15 Mean monthly variations of temperature with depth in the surface waters of the eastern North Pacific. The layer of rapid temperature change is termed the thermocline.

Source: After Tully and Giovando (1963). Reproduced by permission of the Royal Society of Canada.

Figure 3.15 Mean monthly variations of temperature with depth in the surface waters of the eastern North Pacific. The layer of rapid temperature change is termed the thermocline.

Source: After Tully and Giovando (1963). Reproduced by permission of the Royal Society of Canada.

However, the actual temperature change is greater in dry soils. For example, the following values of diurnal temperature range have been observed during clear summer days at Sapporo, Japan:

Sand

Loam

Peat

Clay

Surface

40°C

33°C

23°C

2I°C

5 cm

20

19

14

14

15 cm

7

6

2

4

The different heating qualities of land and water are also accounted for partly by their different specific heats. The specific heat (c) of a substance can be represented by the number of thermal units required to raise a unit mass of it through 1°C (4184 J kg-1 K-1). The specific heat of water is much greater than for most other common substances, and water must absorb five times as much heat energy to raise its temperature by the same amount as a comparable mass of dry soil. Thus for dry sand, c = 840 J kg-1 K-1.

If unit volumes of water and soil are considered, the heat capacity, pc, of the water, where p = density (pc = 4.18 X 106 J m-3 K-1), exceeds that of the sand approximately threefold (pc = 1.3 X 1.6 J m-3 K-1) if the sand is dry and twofold if it is wet. When this water is cooled the situation is reversed, for then a large quantity of heat is released. A metre-thick layer of sea water being cooled by as little as 0.1°C will release enough heat to

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