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Manaus

Valentia

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Figure 3.17 Mean annual temperature regimes in various climates. Manaus, Brazil (equatorial), Valentia, Ireland (temperate maritime) and Toronto, Canada (temperate continental).

chapter). Plate B shows a false-colour satellite thermal image of the western North Atlantic showing the relatively warm, meandering Gulf Stream. Maps of sea-surface temperatures are now routinely constructed from such images.

6 Effect of elevation and aspect

When we come down to the local scale, differences in the elevation of the land and its aspect (that is, the direction that the surface faces) strongly control the amount of solar radiation received.

High elevations that have a much smaller mass of air above them (see Figure 2.13) receive considerably more direct solar radiation under clear skies than do locations near sea-level due to the concentration of water vapour in the lower troposphere (Figure 3.19). On average in middle latitudes the intensity of incident solar radiation increases by 5 to 15 per cent for each 1000 m increase in elevation in the lower troposphere. The difference between sites at 200 and 3000 m in the Alps, for instance, can amount to 70 W m-2 on cloudless summer days. However, there is also a correspondingly greater net loss of terrestrial radiation at higher elevations because the low density of the overlying air results in a smaller fraction of the outgoing radiation being absorbed. The overall effect is invariably complicated by the greater cloudiness associated with most mountain ranges, and it is therefore impossible to generalize from the limited data available.

Figure 3.20 illustrates the effect of aspect and slope angle on theoretical maximum solar radiation receipts at two locations in the northern hemisphere. The general effect of latitude on insolation amounts is clearly shown, but it is also apparent that increasing latitude causes a relatively greater radiation loss for north-facing slopes, as distinct from south-facing ones. The radiation intensity on a sloping surface (I) is

I =I cos i s o where i = the angle between the solar beam and a beam normal to the sloping surface. Relief may also affect the quantity of insolation and the duration of direct sunlight when a mountain barrier screens the sun from valley floors and sides at certain times of day. In many Alpine valleys, settlement and cultivation are noticeably concentrated on southward-facing slopes (the adret or sunny side), whereas northward slopes (ubac or shaded side) remain forested.

7 Variation of free-air temperature with height

Chapter 2C described the gross characteristics of the vertical temperature profile in the atmosphere. We will now examine the vertical temperature gradient in the lower troposphere.

Vertical temperature gradients are determined in part by energy transfers and in part by vertical motion of the

160° 140° 120" 100° 80° 60° 40° 20° 0° 20° 40° 60" 80° 100° 120° 140° 160° 180°

100° ¡40° 120" 100° 80° 60° 40° 20° 0° 20" 40° 60° 80° 100° 120° 140° 160° 180°

Figure 3.18 World temperature anomalies (i.e. the difference between recorded temperatures °C and the mean for that latitude) for January and July. Solid lines indicate positive, and dashed lines negative, anomalies.

100° ¡40° 120" 100° 80° 60° 40° 20° 0° 20" 40° 60° 80° 100° 120° 140° 160° 180°

Figure 3.18 World temperature anomalies (i.e. the difference between recorded temperatures °C and the mean for that latitude) for January and July. Solid lines indicate positive, and dashed lines negative, anomalies.

1000 1 100 1200 1300 1400

ENERGY FLUX DENSITY (Wm~2)

Figure 3.19 Direct solar radiation as a function of altitude observed in the European Alps. The absorbing effects of water vapour and dust, particularly below about 3000 m, are shown by comparison with a theoretical curve for an ideal atmosphere.

Source: After Albetti, Kastrov, Kimball and Pope; from Barry (1992).

1000 1 100 1200 1300 1400

ENERGY FLUX DENSITY (Wm~2)

Figure 3.19 Direct solar radiation as a function of altitude observed in the European Alps. The absorbing effects of water vapour and dust, particularly below about 3000 m, are shown by comparison with a theoretical curve for an ideal atmosphere.

Source: After Albetti, Kastrov, Kimball and Pope; from Barry (1992).

air. The various factors interact in a highly complex manner. The energy terms are the release of latent heat by condensation, radiative cooling of the air and sensible heat transfer from the ground. Horizontal temperature advection, by the motion of cold and warm airmasses, may also be important. Vertical motion is dependent on the type of pressure system. High-pressure areas are generally associated with descent and warming of deep layers of air, hence decreasing the temperature gradient and frequently causing temperature inversions in the lower troposphere. In contrast, low-pressure systems are associated with rising air, which cools upon expansion and increases the vertical temperature gradient. Moisture is an additional complicating factor (see Chapter 3E), but it remains true that the middle and upper troposphere is relatively cold above a surface low-pressure area, leading to a steeper temperature gradient.

The overall vertical decrease of temperature, or lapse rate, in the troposphere is about 6.5°C/km. However, this is by no means constant with height, season or location. Average global values calculated by C. E. P. Brooks for July show increasing lapse rate with height: about 5°C/km in the lowest 2 km, 6°C/km between 4 and 5km, and 7°C/km between 6 and 8 km. The seasonal regime is very pronounced in continental regions with cold winters. Winter lapse rates are generally small and, in areas such as central Canada or eastern Siberia, may even be negative (i.e. temperatures increase with height in the lowest layer) as a result of excessive radiational

Figure 3.20 Average direct beam solar radiation (W m-2) incident at the surface under cloudless skies at Trier, West Germany, and Tucson, Arizona, as a function of slope, aspect, time of day and season of year.

Source: After Geiger (1965) and Sellers (1965).

Figure 3.20 Average direct beam solar radiation (W m-2) incident at the surface under cloudless skies at Trier, West Germany, and Tucson, Arizona, as a function of slope, aspect, time of day and season of year.

Source: After Geiger (1965) and Sellers (1965).

cooling over a snow surface. A similar situation occurs when dense, cold air accumulates in mountain basins on calm, clear nights. On such occasions, mountain summits may be many degrees warmer than the valley floor below (see Chapter 6C.2). For this reason, the adjustment of average temperature of upland stations to mean sea-level may produce misleading results. Observations in Colorado at Pike's Peak (4301 m) and Colorado Springs (1859 m) show the mean lapse rate to be 4.1°C/km in winter and 6.2°C/km in summer. It should be noted that such topographic lapse rates may bear little relation to free air lapse rates in nocturnal radiation conditions, and the two must be carefully distinguished.

In the Arctic and over Antarctica, surface temperature inversions persist for much of the year. In winter the Arctic inversion is due to intense radiational cooling, but in summer it is the result of the surface cooling of advected warmer air. The tropical and subtropical deserts have very steep lapse rates in summer causing considerable heat transfer from the surface and generally ascending motion; subsidence associated with high-pressure cells is predominant in the desert zones in winter. Over the subtropical oceans, sinking air leads to warming and a subsidence inversion near the surface (see Chapter 13).

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