Mountain Climate Dynamics: Slope Aspect, Thermal Belts & Orographic Effects

The direction of slope can be critical to the thermal balance, especially in middle to high latitudes. The effect of radiation loading on air temperature is illustrated in Figure 6.14, which shows air temperature and elevation for a forested site in Germany in summer. At this site, elevation ranges from 340 to 540 m with slopes inclined at 27 percent (15°) to 84 percent (40°). On the particular summer day studied, air temperature varied by as much as 3.5°C depending on aspect. The southwest-facing slope was warmest, with temperature up to 22.5°C.

Fig. 6.14. Maximum air temperature in a forest in the harz mountains of Germany on a warm summer day. Thin lines show elevation (m). Thick lines show temperature (°C). Adapted from Geiger (1965, pp. 426)

The northeast-facing slope, with temperature as low as 19.0°C, was coldest. On the north side of the hill, temperature increased from 19.0°C in the east to 21.0°C in the west. This is because the Sun was high above the horizon and illuminated the north-facing slopes late in the afternoon. Direction of slope is especially important at high latitudes. North-facing slopes in interior Alaska receive less solar radiation than south-facing slopes and therefore have cold soil underlain with permafrost (Van Cleve et al. 1983,1986; Viereck etal. 1983).

Another noticeable feature of mountains is the cool air. So long as air is not saturated with water, it cools at a rate of about 1°C per 100 m. It regains heat at the same rate when it descends. This temperature change is called the dry adiabatic lapse rate. If a low-lying site has an air temperature of 30°C, a site located 300 m higher will be exposed to 27°C air. At the top of a 1500 m mountain, the temperature is only 15°C (Figure 6.15a).

Fig. 6.15 air temperature in relation to elevation on a 1500-m mountain. (a) Dry adiabatic cooling. (b) Moist adiabatic cooling with precipitation

The amount of water vapor that air can hold without becoming saturated depends on its temperature (Figure 3.5) . As moist air rises up a mountain and cools, the amount of water vapor it can hold decreases. So long as the air is not saturated, it cools at the dry adiabatic lapse rate. When the air becomes saturated (i.e., relative humidity is 100 percent), some of the water vapor condenses into droplets, forming fog or clouds. This condensation releases heat (the stored latent heat of vaporization), and the cooling of air as it rises decreases to about 0.5°C per 100 m. This is called the moist adiabatic lapse rate.

Figure 6.15b shows changes in air temperature as moist air moves over a mountain. The rising air on the windward slope cools at the dry adiabatic lapse rate until the air becomes saturated, in this example at about 900 m. When saturated, clouds form and the cooling decreases to the moist adiabatic lapse rate. At the summit, the air is 18°C, which is 3°C warmer than if it cooled at the dry adiabatic lapse rate. If, as in this example, precipitation removes water, the air becomes unsaturated and descends on the leeward slope at the dry adiabatic lapse rate.

Because air is warmed by latent heat of condensation as it moves upslope and by adiabatic heating as it moves downslope, it reaches the bottom warmer than it started on the other side - in this case with a temperature of 33°C.

Under the right conditions, temperatures increase with elevation rather than decrease. Mountains and hillslopes develop local wind circulations in response to spatial variation in surface heating. Under calm conditions and clear sky, light winds often blow upslope during the day and downslope at night. During the day, mountain slopes absorb solar radiation. These warm surfaces heat air, which becomes less dense and rises. Air flows upslope from low-lying valleys, ravines, or plains to replace the ascending mountain air. These upslope circulations depend on a temperature contrast and develop most strongly on slopes receiving the greatest amount of solar radiation. For example, an east-facing slope heats up from early morning solar radiation and may develop upslope winds before a west-facing slope. At night, the slopes cool, and cold air near the surface flows downhill and collects in low-lying areas, often forming frost pockets. The effect of cold air drainage on temperature is seen in a study by Hocevar and Martsolf (1971) of early-morning air temperature in relation to elevation in a broad valley in central Pennsylvania. During one particular night, temperature over a 20-km distance varied by as much as 9°C with low elevations colder than high elevations (air temperature increased 3.4°C per 100 m). On average for all nights studied, air temperature increased at a rate of 6.2°C per 100 m.

Warmest nighttime temperatures can be found at mid-slope. Ridgetops are colder due to their high elevation; valleys are cold because of cold air drainage. Figure 6.16 shows the characteristics of this mid-slope thermal belt for Austrian mountains. On the particular mountain studied, daytime temperatures decreased with elevation as expected. Nighttime temperatures increased with elevation up to 800 m; thereafter temperatures decreased with elevation. The warm region at 800 m is the mid-slope thermal belt. Many meteorological factors influence the location of this belt. The thermal belt develops best on clear, calm nights. With high winds or rain, the normal lapse rate occurs. As a result, the elevation with warmest temperature is not constant but varies over time.

As shown in Figure 6.15, the cooling of air with higher elevation often causes condensation and precipitation on the windward side of mountains and dry conditions on the leeward side. This mountain-induced precipitation is called orographic precipitation. The influence of mountains on precipitation is particularly prominent in western United States, along a transect from the Pacific Ocean inland to Colorado. Annual precipitation drops from 1000 mm west of the Pacific Coastal Ranges to about 500 mm on the east side of the mountains. Precipitation on the west side of the Sierra Nevada is 1000 mm, but less than 200 mm on the east side. Similar orographic precipitation occurs in the Rocky Mountains.

Another feature of mountain climates is that mountaintops can often be exceedingly windy. These winds are evident in the twisted, gnarled shapes of exposed trees growing on or near mountaintops. In the United States, the strongest wind recorded (104 m s^-1) occurred on the summit of Mount Washington in New Hampshire (Williams 1994, p. 48). As air flows over the land surface, frictional resistance from vegetation, buildings, and the ground slows the wind. This resistance decreases rapidly with height in the atmosphere so that high altitude winds are stronger than surface winds.