Cqi Q Q

To determine the rate of change in internal energy and its conversion into potential energy in the atmosphere, appearing on the left-hand side of this equality, as well as the rate of sensible heat change in the ocean and latent heat in the atmosphere and sea ice, we use monthly averaged values of air temperature and air specific humidity systematized by Oort (1983), estimates of the rate of change in heat content {c0m0T0} in a 275-metre layer of ocean, from Levitus (1982), and data on the sea ice area and volume represented in Table 2.1. The calculation results, together with estimates, from Figure 2.2, of the global average net radiation flux at the upper atmospheric boundary are shown in Figure 2.5. They have a lot in common with the results of similar calculations published by Ellis et al. (1978). In both cases the amplitude and phase of seasonal variations of the net radiation flux at the upper atmospheric boundary and the rate of sensible heat change in the ocean are in close proximity, which points to the key role of the ocean in the formation of the seasonal variability of the global heat budget. It turns out that the rate of change in internal energy and its conversion into the potential energy of the atmosphere, and the rate of latent heat change in the atmosphere (their combination will be referred to, for the sake of abbreviation, as rate of energy change) are much less than the net radiation flux at the upper atmospheric boundary. Finally, the rate of energy change and the net radiation flux are out of phase with each other. This last fact, as well as the fact that the maximum and minimum of the rate of energy change fall on cold and warm semiannual periods of the Northern Hemisphere, result from the different ocean/land area ratio in the two hemispheres.

It should be emphasized that the rate of energy change in the atmosphere, even amounting to the latent heat change in the sea ice, cannot balance the discrepancy between the net radiation flux at the upper atmospheric boundary and the rate of sensible heat change in the ocean. This is caused by incompleteness and inaccuracy of initial information and by errors in determining the global heat budget components as small differences of large values. But in attempting to define the cause of the discrepancy we

Figure 2.5 Seasonal variability of the global heat budget in the ocean-atmosphere system: (1) the net radiation flux at the upper atmospheric boundary; (2) the rate of change in sensible heat in the ocean; (3) the rate of change in latent heat in sea ice; (4) the rate of change in latent and sensible heat in the atmosphere.

Figure 2.5 Seasonal variability of the global heat budget in the ocean-atmosphere system: (1) the net radiation flux at the upper atmospheric boundary; (2) the rate of change in sensible heat in the ocean; (3) the rate of change in latent heat in sea ice; (4) the rate of change in latent and sensible heat in the atmosphere.

have identified ourselves with the well-known opinion that the active land layer has no marked effect on the seasonal variability of the global heat budget. Let us check this. Let the heat capacity of a unit column of the active land layer be equal to 1.2 x 107 J/m2 K, let the continent fraction be equal to 0.3 and let the mean temperature change in the active layer be equal to 2 x 10"7 K/s. Then the rate of heat content change 3{cLmLrL}/L/dr will be approximately 0.6 W/m2, that is, it cannot eliminate the imbalance in the global heat budget.

The relation between the components of the heat budget presented above is fulfilled for the planet as a whole but not for its separate parts. Examine, for example, the northern and southern polar regions, bounded by the 70° S and 70°N parallels respectively. In these regions the net radiation flux at the upper atmospheric boundary during most of the year is much larger than the rate of energy change in the atmosphere, and this rate is comparable to the rate of latent heat change in snow and sea ice. The question is: what balances heat losses? There is an unambiguous answer - they are balanced only by the meridional energy transport (the transport of sensible and latent heat, and potential energy) in the ocean-atmosphere system. This can be seen from Table 2.2. The data in Table 2.2 are interesting in more than one respect: they allow us to establish the distinctions in the formation of the heat budget for two polar regions. We discuss the most remarkable distinctions below.

In summer when the net radiation flux at the upper atmospheric boundary in the northern polar region is small, the meridional energy transport is basically balanced by the latent heat release due to ice and snow melting, and by the increase in ocean heat content, whereas in the southern polar region the meridional energy transport is distributed almost equally between the latent heat release due to snow and ice melting and the emission into space. On the other hand, in winter, heat radiation losses at the upper atmospheric boundary in the northern polar region are large, and 2/3 of these losses are compensated by the meridional energy transport and 1/3 of them are compensated by a reduction in the ocean heat content and by the heat release due to ice formation. In the southern polar region the radiation losses at the upper atmospheric boundary are balanced solely by the meridional energy transport. And again the main cause of distinctions is the different ratio between ocean-land areas.

Let us discuss the time-space distribution of the rate of heat content change in the ocean, bearing in mind that this rate (owing to the smallness of its analogue in the atmosphere) might characterize not only the ocean, but also the entire ocean-atmosphere system. The most interesting features of this distribution (Figure 2.6) are the increase in seasonal variations in middle latitudes in the Northern Hemisphere compared with the same latitudes in the Southern Hemisphere, the absence of marked seasonal variations in the equatorial region and in high latitudes of the Northern and Southern Hemispheres, and, finally, the domination of extreme values in the Northern Hemisphere compared with analogous values in the Southern Hemisphere.

Turning to an analysis of the annual mean meridional sensible heat transfer in the ocean-atmosphere system (Figures 2.7(a) and 2.8(a)), we should note first of all that in the tropics the meridional sensible (as well as latent) heat transport in the atmosphere is directed to the equator, and because of this the tropics are a powerful source of heat for the atmosphere such that compensation for the meridional sensible and latent heat transport is realized here mainly by the meridional potential energy transport in the atmosphere,

Table 2.2. Components of the heat budget in the northern and southern polar regions (according to Nakamura and

Oort, 1988)

Table 2.2. Components of the heat budget in the northern and southern polar regions (according to Nakamura and

Oort, 1988)

Components

I

II

III

IV

V

Month VI VII

VIII

IX

X

XI

XII

Annual mean value

Net radiation flux at the upper atmospheric boundary (W/m2)

-164

-146

-122

-78

-37

1

5

-50

-126

-166

-164

-162

-100.7

-10

-72

-106

-148

-135

-139

-130

-123

-99

-63

-38

-14

-89.6

Rate of energy change in the atmosphere

-2

4

16

25

25

21

3

-16

-29

-26

-14

-7

0.0

(W/m2)

-1

-10

-13

-15

-14

-9

-5

1

8

21

24

11

-0.1

Rate of heat content change in the ocean

-29

-27

-11

11

29

32

23

10

0

-6

-12

-21

0.0

(W/m2)

6

2

-2

-4

-5

-5

-4

-4

-1

3

6

8

0.0

Meridional energy

113

94

94

96

82

81

85

88

102

116

114

116

98.4

transport (W/m2)

~33

55

78

105

126

139

143

128

12Ï

113

~64

953

Resulting heat flux at the underlying surface

50

56

44

7

-20

-61

-86

-55

-5

25

36

39

2.4

(W/m2)

-42

-13

11

17

9

7

7

6

5

2

-8

-33

Rate of latent heat change in snow and sea ice (W/m2)

29 63 45

Noie: values in numerators correspond to the northern polar region, and those in denominators correspond to the southern polar region.

Figure 2.6 Time-space distribution of the rate of change in heat content (W/m2) in the World Ocean, according to Kagan and Tsankova (1986).

that the meridional sensible heat transport in high latitudes of the atmosphere and ocean is directed to the poles, that such orientation is an effect of the domination of the absorbed short-wave radiation over upward long-wave radiation in low latitudes and of the inverse ratio in high latitudes, and, finally, that on the equator the meridional sensible heat transport in the ocean takes on near-zero and positive values, indicating the presence of a small heat transport from the Southern to the Northern Hemisphere.

The seasonal variability of the meridional sensible heat transport in the ocean-atmosphere system can be judged by parts (b) in Figures 2.7 and 2.8. The first thing to command our attention is the striking distinctions between time-space distributions of the meridional transport in the atmosphere and the ocean. We mean, firstly, localization of maxima at different latitudes in the atmosphere and ocean (in the atmosphere the maximal variability is coordinated with the equator, and in the ocean it is coordinated with the middle latitudes of both hemispheres), and, secondly, the constancy, during

(a) and its seasonal variability (b). Positive values indicate northward transport; negative values indicate southward transport. (After Carissimo et al, 1985, and Oort and Rasmusson, 1971.)

the annual cycle, of the direction of the meridional transport in middle and high latitudes of the atmosphere and its changing at all latitudes in the ocean.

It is an open secret that the meridional heat transport in the ocean is still the least examined element of the climate, so it is worth dwelling at some length on the analysis of its time-space distribution. Figure 2.8 (part (h)) shows the well-known similarity of seasonal variations of MHT0 in the Northern and Southern Hemispheres. In both hemispheres the meridional heat transport reaches a maximum in winter and is directed to the North Pole at this time of the year. In summer, there is a marked reduction and change in direction of MHT0 everywhere with the exception of the tropical region of the ocean. The extreme values of MHT0 occur in the tropics and fall during February and March in the Northern Hemisphere and during October and November in the Southern Hemisphere. The secondary maxima displaced from the primary maxima further from the equator fall during August and September in the Northern Hemisphere and during May and June in the Southern Hemisphere. This last fact points to the presence of the strongly pronounced semi-annual harmonic.

Let us mention two more distinguishing features of the seasonal variability

'Figure 2.8 Annual mean meridional sensible heat transport (W/m2) in the World Ocean (a) and its seasonal variability (b), according to Kagan and Tsankova (1987). Positive values indicate northward transport; negative values indicate southward transport.

of MHT0. The question is, firstly, about the intensification of the poleward heat transport in spring as compared with autumn in the Northern Hemisphere, and in summer as compared with winter in the Southern Hemisphere, and, secondly, about the excess (about twice) of the period with heat transport directed from the Northern Hemisphere to the Southern Hemisphere as compared with the period with reverse direction of meridional heat transport.

And, finally, the comparison of the annual mean distributions in Figures 2.7 and 2.8 is indicative of the commensurability of the meridional sensible heat transports in the atmosphere and the ocean and, mainly, of their absolutely different natures. This is indicated by the fact that the maximum poleward heat transport in the ocean is in low latitudes where meridional gradients of water temperatures are relatively small, and in the atmosphere in middle latitudes, where meridional gradients of air temperature are large by way of contrast.

2.4 Moisture budget

Let us start, as we did in the preceding section, with the derivation of equations for the moisture budget in separate subsystems of the climatic system. For the atmosphere and sea ice these equations are derived by division of (2.3.4) and (2.3.5) by L(pA/pQ) and L^pjpo), respectively, where p, and pQ are densities of sea ice and fresh water. Continuing, we obtain the following expressions:

0 0

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