Chapter 4: Mountain Climate

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Draft: Chap. 4 Mt. Climate by A. Bach for Mountains and People

Chapter 4: Mountain Climate
Climate is the fundamental factor in establishing a natural environment, it sets the stage upon which all physical, chemical, and biological processes operate. This becomes especially evident at the climatic margins of the earth, i.e., desert and tundra. Under temperate conditions, the effects of climate are often muted and intermingled so that the relationships between stimuli and reaction are difficult to isolate, but under extreme conditions the relationship becomes more evident. Extremes constitute the norm in many areas within high mountains; for this reason, a basic knowledge of climatic processes and characteristics is a prerequisite to an understanding of the mountain milieu.

The climate of mountains is kaleidoscopic, composed of myriad individual segments continually changing through space and time. Great environmental contrasts occur within short distances as a result of the diverse topography and highly variable nature of the energy and moisture fluxes within the system. While in the mountains, have you ever sought refuge from the wind in the lee of a rock? If so, you have experienced the kind of difference that can occur within a small area. Near the margin of a species' distribution, such differences may decide between life and death; thus, plants and animals reach their highest elevations by taking advantage of microhabitats. Great variations also occur within short time spans. When the sun is shining it may be quite warm, even in winter, but if a passing cloud blocks the sun, the temperature drops rapidly. Therefore, areas exposed to the sun undergo much greater and more frequent temperature contrasts than those in shade. This is true for all environments, of course, but the difference is much greater in mountains because the thin alpine air does not hold heat well and allows a larger magnitude of solar radiation to reach the surface.

In more general terms, the climate of a slope may be very different from that of a ridge or valley. When these basic differences are compounded by the infinite variety of combinations created by the orientation, spacing, and steepness of slopes, along with the presence of snow patches, shade, vegetation, and soil, the complexity of climatic patterns in mountains becomes truly overwhelming. Nevertheless, predictable patterns and characteristics are found within this heterogeneous system; for example, temperatures normally decrease with elevation while cloudiness and precipitation increases, it is usually windier in mountains, the air is thinner and clearer, and the sun’s rays are more intense.

The dynamic effects of mountains also have a major impact on regional and local airflow patterns that impact the climates of adjacent regions. Their influence may be felt for hundreds or thousands of kilometers, making surrounding areas warmer or colder, wetter or drier than they would be if the mountains were not there. The exact effect of the mountains depends upon their location, size, and orientation with respect to the moisture source and the direction of the prevailing winds. The 2,400 kilometer long (1,500 mi.) natural barrier of the Himalayas permits tropical climates to extend farther north in India and southeast Asia than they do anywhere else in the world (Tang and Reiter 1984). One of the heaviest rainfall records in the world was measured at Cherrapunji, near the base of the Himalayas in Assam. This famous weather station has an annual rainfall of 10,871 mm (428 in.). Its record for a single day is 1,041 mm (41 in.) as much as Chicago or London receives in an entire year (Kendrew 1961)! On the north side of the Himalayas, however, there are extensive deserts and the temperatures are abnormally low for the latitude. This contrast in environment between north and south is due almost entirely to the presence of the mountains, whose east west orientation and great height prevent the invasion of warm air into central Asia just as surely as they prevent major invasions of cold air into India. It is no wonder that the Hindus pay homage to Siva, the great god of the Himalayas.


Mountain climates occur within the framework of the surrounding regional climate and are controlled by the same factors, including latitude, altitude, continentality, and regional circumstances such as ocean currents, prevailing wind direction, and the location of semi-permanent high and low pressure cells. Mountains themselves, by acting as a barrier, affect regional climate and modifying passing storms. Our primary concern is in the significance of all these more or less independent controls to the weather and climate of mountains.


The distance north or south of the equator governs the angle at which the sun's rays strike the earth, the length of the day, thus the amount of solar radiation arriving at the surface. In the tropics, the sun is always high overhead at midday and the days and nights are of nearly equal length throughout the year. As a result, there is no winter or summer; one day differs from another only in the amount of cloud cover. There is an old adage, "Night is the winter of the tropics." With increasing latitude, however, the height of the sun changes during the course of the year, and days and nights become longer or shorter depending on the season (Fig. 4.1). Thus, during summer solstice in the northern hemisphere (June 21) the day is 12 hours, 7 minutes long at Mount Kenya on the equator; 13 hours, 53 minutes long at Mount Everest in the Himalayas (28˚N lat.); 15 hours, 45 minutes long at the Matterhorn in the Swiss Alps (41˚N lat.); and 20 hours, 19 minutes long at Mount McKinley in Alaska (63˚N lat.) (List 1958). During the winter, of course, the length of day and night at any given location are reversed. Consequently, the distribution of solar energy is greatly variable in space and time. In the polar regions, the extreme situation, up to six months of continuous sunlight follow six months of continuous night.

Although the highest latitudes receive the lowest amounts of heat energy, middle latitudes frequently experience higher temperatures during the summer than do the tropics. This is due to moderate sun heights and longer days. Furthermore, mountains in middle latitudes may experience even greater solar intensity than lowlands, both because the atmosphere is thinner and because the sun's rays strike slopes oriented toward the sun at a higher angle than level surfaces. A surface inclined 20˚ toward the sun in middle latitudes receives about twice as much radiation during the winter as a level surface. It can be seen that slope angle and orientation with respect to the sun are vastly important and may partially compensate for latitude.

The basic pattern of global atmospheric pressure systems reflects the role of latitude in determining climatic patterns (Fig. 4.2). These systems are known as the equatorial low (0˚  20˚ lat.), subtropical high (20˚  40˚ lat.), polar front and subpolar lows (40˚  70˚ lat.), and polar high (70˚  90˚ lat.). The equatorial low and subpolar low are zones of relatively heavy precipitation while the subtropical high and polar high are areas of low precipitation. These pressure zones create the global circulation system (Fig. 4.2). General circulation dictates the prevailing wind direction and types of storms that occur latitudinally. The easterly Trade Winds have warm, very moist convective (tropical) storms, which seasonally follow the direct rays of the sun. The subtropical highs have slack winds and clear skies year round. The subpolar lows and polar front are imbedded in the Westerlies, bringing cool, wet cyclonic storms and large seasonal temperature fluctuations. The cold and dry Polar Easterlies develop seasonally, dissipating in the summer season.

The distribution of mountains in the global circulation system has a major influence on their climate. Mountains near the equator, such as Mount Kilimanjaro in East Africa, Mount Kinabalu in Borneo, or Mount Cotopaxi in Ecuador, are under the influence of the equatorial low and receive precipitation almost daily on their east-facing windward slopes. By contrast, mountains located around 30˚ latitude may experience considerable aridity; as do the northern Himalayas, Tibetan highlands, the Puna de Atacama in the Andes, the Atlas Mountains of North Africa, the mountains of the southwestern United States, and northern Mexico (Troll 1968). Farther poleward, the Alps, the Rockies, Cascades, the southern Andes, and the Southern Alps of New Zealand again receive heavy precipitation on westward slopes facing prevailing Westerlies. Leeward facing slopes and lands down wind are notably arid. Polar mountains are cold and dry year round.

Fundamental to mountain climatology are the changes that occur in the atmosphere with increasing altitude, especially the decrease in temperature, air density, water vapor, carbon dioxide, and impurities. The sun is the ultimate source of energy, but little heating of the atmosphere takes place directly. Rather, solar radiation passes through the atmosphere and is absorbed by the earth’s surface. The earth itself becomes the radiating body, emitting long-wave energy that is readily absorbed by CO2, H2O and other greenhouse gases in the atmosphere. The atmosphere, therefore, is heated directly by the earth, not by the sun. This is why the highest temperatures usually occur near the earth’s surface and decrease outward. Mountains are part of the earth, too, but they present a smaller land area at higher altitudes within the atmosphere, so they are less able to modify the temperature of the surrounding air. A mountain peak is analogous to an oceanic island. The smaller the island and the farther it is from large land masses, the more its climate will be like that of the surrounding sea. By contrast, the larger the island or mountain area, the more it modifies its own climate. This mountain mass effect is a major factor in the local climate (see pp. 77 81).

The density and composition of the air control its ability to absorb and hold heat. The weight or density of the air at sea level (standard atmospheric pressure) is generally expressed as 1013 mb (millibars, or 760 mm [29.92 in.] of mercury). Near the earth, pressure decreases at a rate of approximately 1 mb per 10 m (30 mm/300 m (1 in./1,000 ft.) of increased altitude. Above 5,000 m (20,000 ft.) atmospheric pressure begins to fall off exponentially. Thus, half the weight of the atmosphere occurs below 5,500 m (18,000 ft.) and pressure is halved again in the next 6,000 m (Fig. 4.3).

The ability of air to hold heat is a function of its molecular structure. At higher altitudes, molecules are spaced farther apart, so there are fewer molecules in a given parcel of air to receive and hold heat. Similarly, the composition of the air changes rapidly with altitude, losing water vapor, carbon dioxide, and suspended particulate matter (Tables 4.1 and 4.2). These constituents, important in determining the ability of the air to absorb heat, are all concentrated in the lower reaches of the atmosphere. Water vapor is the chief heat absorbing constituent, and half of the water vapor in the air occurs below an elevation of 1,800 m (6,000 ft.). It diminishes rapidly above this point and is barely detectable at elevations above 12,000 m (40,000 ft.).

The importance of water vapor as a reservoir of heat can be seen by comparing the daily temperature ranges of a desert to that of a humid area. Both areas may heat up equally during the day but, due to the relative absence of water vapor to absorb and hold the heat energy, the desert area cools down much more at night than the humid area. The mountain environment responds in a similar fashion to that of a desert, but is even more accentuated. The thin pure air of high altitudes does not effectively intercept radiation, allowing it to be lost to space. Mountain temperatures respond almost entirely to radiation fluxes, not on the temperature of the surrounding air (although some mountains receive considerable heat from precipitation processes). The sun's rays pass through the high thin air with negligible heating. Consequently, although the temperature at 1,800 m (6,000 ft.) in the free atmosphere changes very little between day and night, next to a mountain peak, the sun's rays are intercepted and absorbed. The soil surface may be quite warm but the envelope of heated air is usually only a few meters thick and displays a steep temperature gradient.

In theory, every point along a given latitude receives the same amount of sunshine; in reality, of course, clouds interfere. The amount of cloudiness is controlled by distance from the ocean, direction of prevailing winds, dominance of pressure systems, and altitude. Precipitation normally increases with elevation, but only up to a certain point. Precipitation is generally heaviest on middle slopes where clouds first form and cloud moisture is greatest, decreasing at higher elevations. Thus, the lower slopes can be wrapped in clouds while the higher slopes are sunny. In the Alps, for example, the outer ranges receive more precipitation and less sunshine than the higher interior ranges. The herders in the Tien Shan and Pamir Mountains of Central Asia traditionally take their flocks higher in the winter than in summer to take advantage of the lower snowfall and sunnier conditions at the higher elevations. High mountains have another advantage with respect to possible sunshine: in effect, they lower the horizon. The sun shines earlier in the morning and later in the evening on mountain peaks than in lowlands. The same peaks, however, can raise the horizon for adjacent land, delaying sunrise or creating early sunsets.


The relationship between land and water has a strong influence on the climate of a region. Generally, the more water dominated an area is, the more moderate its climate. An extreme example is a small oceanic island, on which the climate is essentially that of the surrounding sea. The other extreme is a central location on a large land mass such as Eurasia, far removed from the sea. Water heats and cools more slowly than land, so the temperature ranges between day and night and between winter and summer are smaller in marine areas than in continental areas.

The same principle applies to alpine landscapes, but is intensified by the barrier effect of mountains. We have already noted this effect in the Himalayas between India and China. The Cascades in the Pacific Northwest of the United States provide another good example. This range extends north south at right angles to the prevailing westerly wind off the Pacific Ocean. As a result, western Oregon and Washington have a marine dominated climate characterized by moderate temperatures, cloudiness, and persistent winter precipitation (Schermerhorn 1967). The eastern side of the Cascades, however, experiences a continental climate characterized by hot summers and cold winters with low precipitation. In less than 85 km (50 mi.) across the Cascades the vegetation changes from lush green forests to dryland shrubs and grasses (Price 1971a). This spectacular transect provides eloquent testimony to the vast differences in climate that may occur within a short horizontal distance. The presence of the mountains increases the precipitation in western Oregon and Washington at the expense of that received on the east side. Additionally, the Cascades inhibit the invasion of cold continental air to the Pacific side. At the same time, their obstruction of mild Pacific air allows the continental climate to extend much closer to the ocean than it otherwise would (Church and Stephens 1941). It must be stressed that the significance of mountains in accentuating continentality depends upon their orientation with respect to the ocean and prevailing winds. Western Europe has a climate similar to the Pacific Northwest, but the east west orientation of the European mountains allows the marine climate to extend far inland, resulting in a milder climate throughout Europe.

The effect of continentality on mountain climate is much like that on climate generally. Mountains in the interior of continents experience more sunshine, less cloudiness, greater extremes in temperatures, and less precipitation than mountains along the coasts. This would seem to add up to a more rigorous environment, but there may be extenuating circumstances. The extra sunshine in continental regions tends to compensate for the lower ambient temperatures, while the greater cloudiness and snowfall in coastal mountains tend to make the environment more rigorous for certain organisms than is suggested by the moderate temperatures of these regions. The fact that trees generally grow to higher altitudes on continental mountains than coastal mountains is a good, if rough, indication of the importance of these compensating circumstances to regional mountain climate and ecology (see pp. 277 82). People, too, find that the bright sunshine typical of high mountain slopes can make the low air temperatures of the alpine environment tolerable. During the winter in the Alps, for instance, when it is cloudy and rainy in the surrounding lowlands and foggy in the lower valleys, the mountain slopes and higher valleys may bask in brilliant sunshine. It is for this reason that lodges and tourist facilities in the Alps are generally located higher up on the slopes and in high valleys. Health resorts and sanatoriums also take advantage of the intense sunlight and clean dry air of the high mountains (Hill 1924).

Barrier Effects

Several examples of how mountains serve as barriers have already been given. The Himalayas and Cascades are both outstanding climatic divides that create unlike conditions on their windward and leeward sides. All mountains serve as barriers to a greater or lesser extent, depending on their size, shape, orientation, and relative location. Specifically, the barrier effect of mountains can be grouped under the following subheadings: (1) damming, (2) deflection and funneling, (3) blocking and disturbance of the upper air, (4) forced ascent, and (5) forced descent.


Damming of stable air occurs when the mountains are high enough to prevent the passage of an air mass across them. When this happens, a steep pressure gradient may develop between the windward and leeward sides of the range (Stull 1988). The effectiveness of the damming depends upon the depth of the air mass and the elevation of the lowest valleys or passes (Smith 1979). A shallow, ground hugging air mass may be effectively dammed, but a deep one is likely to flow through higher gaps and transverse valleys to the other side. In the Los Angeles Basin of southern California, for example, the San Gabriel, San Bernardino, and San Jacinto Mountains act as dams for marine air blowing from the Pacific Ocean. As the automobile-based culture of southern California pollutes the air, the pollution can only be vented as far east as the towns of San Bernardino and Riverside at the base of the mountains. In the absence of a strong wind system, the pollution can build up to dangerous levels as the air stagnates behind the mountain barrier.

Deflection and Funneling

When an air mass is dammed by a mountain range, the winds can be deflected around the mountains if topographic gaps exist. Deflected winds can have higher velocities as their streamlines are compressed, the so-called ‘Bernoulli-effect’ (Davidson et al. 1964; Chen and Smith 1987). In winter, polar continental air coming down from Canada across the central United States is channeled to the south and east by the Rocky Mountains. Consequently, the Great Plains experience more severe winter weather than does the Great Basin (Church and Stephens 1941; Baker 1944). Similarly, as the cold air progresses southward, the Sierra Madre Oriental prevents it from crossing into the interior of Mexico. The east coast of Mexico also provides an excellent example of deflection in the summer: the northeast trade winds blowing across the Gulf of Mexico cannot cross the mountains and are deflected southward through the Isthmus of Tehuantepec, where they become northerly winds of unusual violence (Hurd 1929). Maritime air from the northeastern Pacific is deflected north and south around the Olympic Mountains (Fig. 4.4). To the north of the Olympics where wind is also deflected south from the Vancouver Island Ranges, these winds converge into a topographic funnel of the Strait of Juan de Fuca, resulting in much higher wind speeds (Ramachandran et al. 1980). A similar phenomenon occurs around the Southern Alps of New Zealand, with winds funneled through Cook Strait between the islands (Reid 1996; 1997; Sturman and Tapper 1996). These perturbations to the local airflow influence transit storms, making local forecasts difficult. The same funneling effect occurs over mountain passes as winds are deflected around peaks or ridges on either side of the pass. In the Los Angeles Basin example given above, the San Gorgonio Pass (750 m) is the lowest divide through the damming mountains. Wind speeds average 7.2 m/s and are very consistent, resulting in very active aeolian processes and a booming wind power generating industry (Williams and Lee 1995).

Blocking and Disturbance of the Upper Air

High pressure areas prevent the passage of storms. Large mountain ranges such as the Rockies, Southern Alps and Himalayas are very efficient at blocking storms, since they are often the foci of anti-cyclonic systems (because the mountains are a center of cold air), the storms must detour around the mountains (Kimurak and Manins 1988; McCauley and Sturman 1999). In addition to the effect of blocking, mountains cause other perturbations to upper air circulation and subsequent effects on clouds and precipitation (Chater and Sturman 1998). This occurs on a variety of scales: locally, with the wind immediately adjacent to the mountains; on an intermediate scale, creating large waves in the air; and on a global basis, with the larger mountain ranges actually influencing the motion of planetary waves (Bolin 1950; Gambo 1956; Kasahara 1967; Carruthers and Hunt 1990; Walsh 1994) and the transport momentum of the total circulation (White 1949; Wratt et al. 1996). Disturbance of the air by mountains generally creates a wave pattern much like that found in the wake of a ship. This may result in the kind of clear air turbulence feared by airline pilots (Alaka 1958; Colson 1963) or it may simply produce lee waves with their beautiful lenticular (standing wave) clouds, associated with mountains the world over (Fig. 4.41; Scorer 1961). An area of low precipitation occurs immediately lee of the Rocky Mountains: the area immediately to the lee is frequently cloud-free and receives low precipitation, while regions farther east are cloudy and wetter. This pattern corresponds to an intermediate-scale wave whose trough is located close to the lee of the mountains and whose ridge is located over the eastern United States (Reiter et al. 1965; Dirks et al. 1967 Durran 1990; Czarnetzki and Johnson 1996).

Mountains have additional influence on the location and intensity of the jet streams, which have vastly important effects on the kind of weather experienced at any particular place and time. The jet streams may also split to flow around the mountains; they rejoin to the lee of the range, where they often intensify and produce storms (Reiter 1963; Buzzi et al. 1987). In North America these storms, known as "Colorado Lows" or "Alberta Lows," reach their greatest frequency and intensity in the spring season, sometimes causing heavy blizzards on the Great Plains and Prairie provinces. The tornadoes and violent squall lines that form in the American Midwest also result from the great contrasts in air masses which develop in the confluence zone to the lee of the Rockies (McClain 1958; Henz 1972; Chung et al 1976).

The splitting of the jet streams by the Himalayas has the effect of intensifying the barrier effect in this region and produces a stronger climatic divide. In addition, the presence of the Himalayas reverses the direction of the jet streams in early summer. The Tibetan Highlands act as a "heat engine" in the warm season, with a giant chimney in their southeastern comer through which heat is carried upward into the atmosphere. This causes a gradual warming of the upper air above the Himalayas during the spring, which weakens and finally eliminates the subtropical westerly jet. The easterly tropical jet then replaces the subtropical jet during the summer. Thus, the Himalayas are intimately connected with the complex interaction of the upper air and the development of the Indian monsoon (Flohn 1968; Hahn and Manabe 1975; Reiter and Tang 1984; Tang and Reiter 1984; Kurtzbach et al. 1989).

Forced Ascent

When moist air blows perpendicular to a mountain range, the air is forced to rise; as it does, it is cooled. Eventually the dew point is reached, condensation occurs, clouds form, and precipitation results (see p. 94). This increased cloudiness and precipitation on the windward slope is known as the orographic effect (Browning and Hill1981). Some of the rainiest places in the world are mountain slopes in the path of winds blowing off relatively warm oceans. There are many examples and could be given from every continent, but the mountainous Hawai’ian Islands will serve as an illustration. The precipitation over the water around Hawai’i averages about 650 mm (25 in.) per year, while the islands average 1,800 mm (70 in.) per year. This is largely due to the presence of mountains, many of which receive over 6,000 mm (240 in.) per year (Nullet and MacGranaghan 1988). At Mount Waialeale on Kauai, the average annual rainfall reaches the extraordinary total of 12,344 mm (486 in.), i.e., 12.3 m (40.5 ft.)! This is the highest recorded annual average in the world (Blumenstock and Price 1967). In the continental United States, the heaviest precipitation occurs at the Hoh Rain Forest on the western side of the Olympic Mountains in Washington, where an average of 3,800 mm (150 in.) or more is received annually as storms are funneled up valleys oriented towards winter storm tracks (Fig. 4.4; Phillips 1972; Collie and Mass 1996).

Forced Descent

Atmospheric pressure conditions determine whether the air, after passing over a mountain barrier, will maintain its altitude or whether it will be forced to descend. If the air is forced to descend, it will be heated by compression (adiabatic heating) and will result in clear, dry conditions. This is a characteristic phenomenon in the lee of mountains and is responsible for the famous foehn or chinook winds (see pp. 114 19). The important point here is that the descent of the air is induced by the barrier effect and results in clear dry conditions that allow the sunshine to reach the ground with much greater intensity and frequency than it otherwise would. This can produce "climatic oases" in the lee of mountain ranges, e.g., in the Po Valley of Italy (Thams 1961).

Although heavy precipitation may occur on the windward side of mountains where the air is forced to rise, the leeward side may receive considerably less precipitation because the air is no longer being lifted (it is descending) and much of the moisture has already been removed. The so-called rainshadow effect is an arid area on the leeward or down-wind side of mountains. To the lee of Mount Waialeale, Kauai, precipitation decreases at the rate of 3,000 mm (118 in.) per 1.6 km (1 mi.) along a 4 km (2.5 mi.) transect to Hanalei Tunnel (Blumenstock and Price 1967). In the Olympic Mountains, precipitation decreases from the windward side to less than 430 mm (17 in.) at the town of Sequim on the leeward, a distance of only 48 km (30 mi.) (Fig. 4.4; Phillips 1972). Since both of these leeward areas are maritime, they are still quite cloudy; under more continental conditions, there would be a corresponding increase in sunshine as precipitation decreases, especially where the air is forced to descend on the leeward side.

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