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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.

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

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.

EXTERNAL CLIMATIC CONTROLS 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 semipermanent 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.

Latitude 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

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

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.

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

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.

Altitude 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

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

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

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

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.

Continentality 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,

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

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

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

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 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.

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

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

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

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

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

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

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

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.

MAJOR CLIMATIC ELEMENTS The discussion so far has covered the more or less independent climatic controls of latitude, altitude, continentality, and the barrier effect of mountains. These factors, along with ocean currents, pressure conditions, and prevailing winds, control the distribution of sunshine, temperature, humidity, precipitation, and local winds. The climatic elements of sunshine, temperature, and precipitation are essentially dependent variables reflecting the major climatic controls (Thompson 1990). They interact in complex ways to produce the day-to-day weather conditions experienced in different regions. In mountains, these processes frequently occur on small enough scales to be invisible to standard measurement networks used in weather forecasting, while their impact can be serious.

Solar Radiation The effect of the sun becomes more exaggerated and distinct with elevation. The time lag, in terms of energy flow, between stimulus and reaction is greatly compressed in mountains. Looking at the effect of the sun in high mountains is like viewing its effects at lower elevations through a powerful magnifying glass. The alpine environment has perhaps the most extreme and variable radiation climate on earth. The thin clean air allows very high solar intensities, and the topographically complex landscape provides surfaces with a range of different exposures and shadowing from nearby peaks. Although the air next to the ground may heat up very rapidly under the direct rays of the sun, it may cool just as rapidly if the sun's rays are blocked. Thus, in the sun's daily and seasonal march through the sky, mountains experience a continually changing pattern of sunshine and shadow, influencing the energy flux in the ecosystem (Saunders and Bailey 1994; Germino and Smith 2000). The factors to consider are the amount of sunlight received, the quality or kinds of radiation, and the effect of slopes upon this energy.

Amount of Solar Radiation

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

The most striking aspect of the vertical distribution of solar radiation in the atmosphere is the rapid depletion of short-wavelength energy at lower elevations. This attenuation results from the increased density of the atmosphere and the greater abundance of water vapor, carbon dioxide, and particulate matter near the earth‘s surface (Tables 4.1 and 4.2). The atmosphere acts as a filter, reducing the intensity of some wavelengths and screening out others altogether. Consequently, the amount of energy reaching the surface at sea level is only about half that at the top of the atmosphere (Fig. 4.5). High mountains protrude through the lower atmospheric blanket and thus have the potential for receiving much higher levels of solar radiation, as well as cosmic-ray and ultraviolet radiation (Solon et al. 1960). The first, and very vital, screening of solar energy takes place in the stratosphere where most of the ultraviolet radiation from the sun is absorbed by the ozone layer. Greenhouse gases absorb infrared solar radiation, but visible light passes through to the surface except when there is cloud-cover. The visible light is scattering as it strikes molecules of air, water, and dust. Scattering is a selective process, principally affecting the wavelengths of blue light. Have you ever noticed how much bluer or darker the sky looks in high mountains than it appears at lower elevations? That is because there is more water and pollutants at lower altitudes, scattering light of other wavelengths, which dilutes the blue color (Valko 1980). Clouds, of course, are the single most important factor in controlling variable receipt of solar energy at any given latitude and in mountains (Saunders and Bailey 1994). Because of the atmospheric filtering of solar radiation, the more atmosphere the sunlight passes through, the greater the attenuation. Consequently, the sun is most intense when it is directly overhead (90˚) and its rays concentrated in the smallest area. When the sun is only 4˚ above the horizon, solar rays have to penetrate an atmosphere more than twelve times as thick as when the sun is directly overhead. This explains why it is possible to look directly at the orange ball of the sun at sunrise and sunset without being blinded. Since mountains stand above the lower reaches of the atmosphere, the solar radiation is much more intense since it has passed through less atmosphere (Fig. 4.6). Table 4.3 gives values for daily global radiation received at different elevations in the Austrian Alps. Solar intensity increases with altitude under all conditions, but the greatest differential between high and low-level stations occurs when skies are overcast. In summer,

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when skies are clear, there is 21% more radiation at 3,000 m (10,000 ft.) than at 200 m (650 ft.); but when skies are overcast, there is 160% more radiation at the higher elevation. Overcast skies are much more efficient at filtering out shortwave energy, so less reaches the lower elevation (Geiger 1965). The solar constant is defined as the average amount of total radiation energy received from the sun at the top of the atmosphere on a surface perpendicular to the sun's rays (Fig. 4.5). This is approximately 1365 Wm-2 (2 calories per square centimeter per min). At midday under clear skies the total energy flux from the sun in high mountains may approach the solar constant. Angstrom and Drummond (1966) have calculated the theoretical upper limit on high mountains to be 1263 Wm-2 (1.85 cal. cm-2 min-1), but several field investigations have recorded readings even slightly above the solar constant (Turner 1958a; Gates and Janke 1966; Bishop et al. 1966; Terjung et al. 1969a, b; Marcus and Brazel 1974). Turner (1958a) measured instantaneous values as high as 1529 Wm-2 (2.25 cal. cm-2 min-1) in the Alps, 112% of the solar constant! The additional radiation comes from sunlight reflected from cloud bottoms and snow on higher slopes.

Quality of Solar Radiation The alpine environment receives considerably more ultraviolet radiation (UV) than low elevations. If only wavelengths shorter than 320 m. are considered, then alpine areas receive 50% more UV during summer solstice than does sea level (Caldwell 1980). Later in the year, when the sun is lower in the sky (and therefore passes through denser atmosphere), alpine areas receive 120% more UV than areas at sea level (Gates and Janke 1966). The relatively greater quantity of UV received at high elevations has special significance for human comfort and biological processes. A proverb in the Andes says, ―Solo los gringos y los burros caminan en el sol‖ (―Only foreigners and donkeys walk in the sunshine‖). This saying indicates the respect the Andeans give to the efficacy of the sun at high altitudes (Prohaska, 1970). UV has been cited for a number of harmful effects, ranging from the retardation of growth in tundra plants (Lockart and Franzgrote 1961; Caldwell 1968; Runeckles and Krupa 1994) to cancer in humans (Blum 1959). UV is mainly responsible for the deep tans of mountain dwellers and the painful sunburns of neophytes who expose too much of their skin too quickly. The wavelengths responsible for

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

sunburn occur primarily between 280 and 320 m, while those responsible for darkening the skin occur between 300 and 400 m. Wavelengths less than 320 m are known to cause skin cancer and weaken the immune system (Chapman and Werkema 1995). UV has been increasing in alpine areas in recent decades, apparently a response to the depletion of stratospheric ozone (Blumenthaler and Ambach 1990).

Effect of Slopes on Solar Radiation The play of the sun on the mountain landscape is like a symphony. As the hours, days, and seasons follow one another, the sun bursts upon some slopes with all the strength of crescendo while the shadows lengthen and fade into diminuendo on others. The melody is continuous and ever-changing, with as many scores as there are mountain regions, but the theme remains the same. It is a study of slope angle and orientation. The closer to perpendicularly the sun's rays strike a surface, the greater their intensity. The longer the sun shines on a surface, the greater the heating that takes place (Anderson 1998). In mountains, every slope has a different potential for receiving solar radiation. This amount can be measured if the following data are known: latitude, time of year (height of sun), time of day, elevation, slope angle, and slope orientation (Gamier and Ohmura 1968, 1970; Swift 1976; Baily et al. 1989; Bowers and Bailey 1989; Huo and Bailey 1992). The basic characteristics of solar radiation on slopes are illustrated in Figure 4.7. This very useful diagram shows the situation for one latitude at four times of the year, at four slope orientations. They do not include the effects of clouds; diffuse sky radiation, or the receptiveness of different slopes to the sun's rays. The diagram also fails to reveal the shadow effects caused by the presence of ridges or peaks above a location. Most mountain slopes receive fewer hours of sunshine than a level surface, although slopes facing the sun may receive more energy than a level surface (this is particularly true at higher latitudes). In the tropics, level surfaces usually receive a higher solar intensity than slopes because the sun is always high in the sky. Whatever the duration and intensity of sunlight, the effects are generally clearly evident in the local ecology (Fig. 4.8). In the northern hemisphere, south-facing slopes are warmer and drier than north-facing slopes and, under humid conditions, are more favorable for life. Timberlines go higher on south-facing slopes, and the number and

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diversity of plants and animals are greater (Germino and Smith 2000). Humans take advantage of the sunny slopes. In the east-west valleys of the Alps most settlements are located on south-facing slopes. Houses are seldom found within the mid-winter noonday shadow area, although they may go right up to the shadow line (Fig. 4.9; Garnett 1935, 1937). In spring north-facing slopes may still be deep in snow while south-facing slopes are clear. As a result, north-facing slopes have traditionally been left in forest while south-facing slopes are used for high pastures (Fig. 4.10). The environmental differences are so great between the sunny and shady sides of the valley that each mountain speech or dialect in the Alps has a special term for these slopes (Peattie 1936). The most frequently used are the French adret (sunny) and ubac (shade). East- and west-facing slopes are also affected differently by solar radiation. Soil and vegetation surfaces are frequently moist in the morning, owing to higher humidity at night and the formation of dew or frost. On east-facing slopes the sun's energy has to evaporate this moisture before the slope can heat appreciably. By the time the sun reaches the west-facing slope, however, the moisture has already evaporated, so the sun's energy more effectively heats the slope. The driest and warmest slopes are, therefore, those that face toward the southwest rather than strictly south (Blumer 1910). Cloud cover, which varies latitudinally, from season to season, and according to time of day, can make a great deal of difference in the amount of solar energy received on slopes. During storms the entire mountain may be wrapped in clouds; even during relatively clear weather, mountains may still experience local clouds. In winter, stratus clouds and fog are characteristic on intermediate slopes and valleys, but these frequently burn off by midday. In summer, the mornings are typically clear but convection clouds (cumulus) build by mid-afternoon from thermal heating. Consequently, convection clouds result in east-facing slopes receiving greater sunlight while stratus clouds, as described above, allow greater sun on west-facing slopes. As clouds move over mountains, build and dissipate through each day, they have a marked effect upon the amount and character of radiation received. Mountains are composed of a wide range of surface types, snow, ice, water, grassy pastureland, extensive forests, desert shrub, soils, and bare bedrock. This extensive variety of surface characteristics affects the receipt of incoming solar radiation (Miller 1965, 1977; Goodin

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and Isard 1989; Tappenier and Cernusca 1989). The effects of two factors, groundcover and topographic setting, will illustrate this. Dark-colored features, including vegetation, absorb rather than reflect radiation, receiving increased amounts of energy. Snowfields, glaciers, and light-colored rocks have a high reflectivity (albedo), so that much of the incoming shortwave energy is lost. If the snow is in a valley or on a concave slope, reflected energy may bounce from slope to slope, increasing the energy budget of the upper slopes. The opposite occurs on a mountain ridge or convex slope, where the energy is reflected back out into space. Consequently, valleys and depressions are areas of heat build-up and generally experience greater temperature extremes than do ridges and convex slopes. Reflected energy is an important source of heat for trees in the high mountains (Martinec 1987; 1989). Snow typically melts faster around trees because the increased heat is transferred, as longwave thermal energy, to the adjacent surface (Plüss and Ohmura 1997). On a larger scale, the presence of forests adds significantly to the heat budget of snow-covered areas. The shortwave energy from the sun can pass through a coniferous forest canopy, but very little of it escapes again to outer space. The absorbed energy heats the tree foliage and produces higher temperatures than in open areas. This results in rapid melting rates of the regional snowpack (Miller 1959; Martinec 1987). Variation in the components of the surface energy budget provides the main driving force of regional differences in climate. In particular, the relative magnitude of sensible and latent heat fluxes reflects the influence of prevailing weather systems, as well as playing an important role in determining atmospheric temperature and moisture content (McCutchan and Fox 1986; Bailey et al. 1990; Kelliher et al. 1996). These factors in turn have an influence on the development of local wind systems. The surface energy budgets can vary significantly in mountains due to the effects of both complex topography and surface characteristics. When snow or ice are present, energy must first be partitioned to ablation before temperatures rise, and once the snow melts there are large changes in albedo (Cline 1997). These variations affect both the distribution of incoming and outgoing radiation, influencing net radiation, soil heat flux, sensible and latent heat; and producing a range of topo- and microclimates (Barry and Van Wie 1974; Green and Harding 1980; Fitzharris 1989; Germino and Smith 2000).

Temperature

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The decrease of temperature with elevation is one of the most striking and fundamental features of mountain climate. Those of us who are fortunate enough to live near mountains are constantly reminded of this fact, either by spending time in the mountains or by viewing the snowcapped peaks from a distance. Nevertheless, there are many subtle and poorly understood characteristics about the nature of temperature in mountains. Alexander Von Humboldt was so struck by the effect of temperature on the elevational zonation of climate and vegetation in the tropics that he proposed the terms tierra calienfe, tierra templada, and tierra fria for the hot, temperate, and cold zones. These terms, commonplace in the tropics today, are still valid for this region. Their extension to higher latitudes by others, however, under the mistaken assumption that the same basic kinds of temperature conditions occur in belts from the equator to the poles has been unfortunate. This simplistic approach is still used in some textbooks.

Vertical Temperature-Gradient Change of temperature with elevation is called the environmental or normal lapse rate. De Saussure, who climbed Mount Blanc in 1787, was one of the first to measure temperature at different elevations. Since his time many temperature measurements have been made in mountains throughout the world, and almost every one of them has been different (Tabony 1985). The lapse rate varies according to many factors. Nevertheless, by averaging the temperatures at different levels, as well as those measured in the free air by balloon, radiosonde, and aircraft, average lapse rates have been established, ranging from 1˚C to 2˚C (1.8˚F to 3.6˚F) per 300 m (1,000 ft.) (McCutchan 1983). Aside from purposes of gross generalization, however, average lapse rates have little value in mountains. There is no constant relationship between altitude and temperature. Instead, the lapse rate changes continually with changing conditions, particularly the diurnal heating and cooling of the earth‘s surface. For example, the vertical temperature-gradient is normally greater during the day than at night, and greater during the summer than in winter. The gradient is steeper under clear than cloudy conditions, steeper on sun-exposed slopes than shaded ones, and steeper on continental mountains than on maritime mountains (Peattie 1936; Dickson 1959; Tanner 1963; Yoshino 1964a, 1975; Coulter 1967; Marcus 1969). There is also a difference between the characteristics of free-air temperature and that measured on a mountain slope (McCutchan 1983; Richner and Phillips 1984; Pepin and Losleben 2002). Of course, the

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higher and more isolated a mountain peak is, the more closely its temperature will approach that of the free atmosphere (Schell 1934, 1935; Eide 1945; Samson 1965). Table 4.4 shows data for the average decrease of temperature with changing elevation in the Alps, and Figure 4.11 illustrates the temperature changes with elevation in the southern Appalachians of the United States. The temperatures shown are averages, with some interpolation between stations; the actual decrease with elevation is much more variable. A station located on a sunny slope will have a temperature regime different from that of a shaded slope (Fig. 4.23). The disposition of winds and clouds is equally important, as is the nature of the slope surface-whether it is snow-covered, wet or dry, bare or vegetated (Green and Hardy 1979; 1980). A convex slope has qualities of heat retention different from those of a concave slope. A high valley will heat up more during the day (and cool down more at night) than an exposed ridge at the same elevation. Nevertheless, broad averages will smooth out the extremes and individual differences, generally showing a steady and progressive decrease in temperature with increase in elevation.

Mountain Mass (Massenerhebung) Effect Large mountain systems create their own surrounding climate (Ekhart 1948). Similar to the continentality effect, the greater the surface area or land mass at any given elevation, the greater effect the mountain area will have on its own environment. Mountains serve as elevated heat islands where solar radiation is absorbed and transformed into long-wave heat energy, resulting in much higher temperatures than those found at similar altitudes in the free air (Flohn 1968; Chen et al. 1985; Rao and Endogan 1989). Accordingly, the larger the mountain mass, the more its climate will vary from the free atmosphere at any given altitude. This is particularly evident on some of the high plateaus, where treeline and snowline often occur at higher elevations than on isolated peaks at the same elevation. On the broad general level of the Himalayas, at 4,000 m (13,100 ft.) it seldom freezes during summer, while on the isolated peaks at 5,000 m (16,400 ft.) it seldom thaws (Peattie 1936; Tang and Reiter 1989; Brazel and Marcus 1991). An excellent example of the heating effect of large high-altitude land masses is the Mexican Meseta (Fig. 4.12). Radiosonde data indicate higher temperatures in the free atmosphere

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over the plateau than over the Pacific and Gulf coasts up to an elevation of almost 6,000 m (20,000 ft.). The mean annual temperature over the central plateau at 3,000 m (10,000 ft.) is about 3˚C (5.4˚F) warmer than that over coastal stations (Hastenrath 1968). This is largely due to the heating effects of the sun on the larger land mass exposed at higher elevations. In establishing the relationships between mountain mass and the heat balance, continentality, latitude, amount of cloud cover, winds, precipitation, and surface conditions must all be considered. A persistent cloud cover during the summer can prevent a large mountain mass from showing substantial warming. Also, the presence of a heavy snow cover can retard the warming of a mountain area in spring because of surface reflectivity and the amount of initial heat required to melt the snow. The high Sierra Nevada of California are relatively warm compared with other mountain areas, in spite of heavy snowfalls (Miller 1955). This is partially because the extreme clarity of the skies over this region in late summer allows maximum reception of solar energy. In general, the effect of greater mountain mass on climate is somewhat like that of increasing continentality. The ranges of temperature are greater than on small mountains, i.e., the winters are colder and the summers warmer, but the average of these temperatures will generally be higher than the free air at the same altitude. The effective growing climate, especially, is more favorable at the soil surface than in the free air, owing to higher soil temperatures. This is particularly true when there is a high percentage of sunshine (Peattie 1931; Yoshino 1975). Generally, the larger the mountain mass, the higher the elevation at which vegetation grows. The most striking example of this is found in the Himalayas, where plants reach their absolute highest altitude (Zimmermann 1953; Webster 1961; Chen et al. 1985). In the Alps (where the influence of mountain mass, Mussenerhebung, was first observed) the timberline is higher in the more massive central area than on the marginal ranges (Imhof 1900, in Peattie 1936, p. 18). At a more local level, the effects of mountain mass on vegetation development can be observed in the Oregon Cascades. Except for Mount McLoughlin in southern Oregon, timberline is highest and alpine vegetation reaches its best development in the Three Sisters Wilderness area, where three peaks join to form a relatively large land mass above 1,800 m (6,000 ft.) (Price 1978). On the higher but less massive peaks of Mount Hood and Mount Washington a few kilometers to the north, the timberline is 150-300 m (500-1,000 ft.) lower and the alpine

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vegetation is considerably more impoverished. The development of vegetation involves more than climate, of course, since plant adaptations and species diversity are related to the size of the gene pool and other factors (Van Steenis 1961). Nevertheless, vegetation is a useful indicator of environmental conditions and a positive correlation between vegetation development and mountain mass can be observed in most mountain areas (see pp. 266-67). An interesting practical consequence of the mountain mass effect is that rice, basically a tropical plant, can grow at higher altitudes in the subtropics than in the tropics. Rice cultivation goes up to 2,500 m (8,250 ft.) in the high interior valleys of the Himalayas (Fig. 4.13) but only reaches about 1,500 m (5,000 ft.) in the humid tropics. The lower tropical limits are due to the lower cloud level, whereas the higher elevations reached in the Himalayas are due to the greater mountain mass and reduced cloudiness, permitting greater possible sunshine, higher temperatures, and a longer growing season than would otherwise be expected. In general, the upper limit of rice cultivation corresponds closely to the limit of frost during the growing season. At the highest levels in the Himalayas, rice seedlings are germinated and grown inside the houses, since it takes eight months for complete production at this elevation but the growing season is only seven months long (Uhlig 1978).

Temperature Inversion Temperature inversions are ubiquitous in landscapes with marked relief, and anyone who has spent time in or around mountains is certain to have experienced their effects. Inversions are the exception to the general rule of decrease in temperature with elevation. During a temperature inversion the lowest temperatures occur in the valley and increase upward along the mountain slope. Eventually, however, the temperatures will begin to decrease again, so that an intermediate zone, the thermal belt, will experience higher night temperatures than either the valley bottom or the upper slopes (Yoshino 1984). Cold air is denser and therefore heavier than warm air. As slopes cool at night, the colder air begins to slide down slope, flowing underneath and displacing the warm air in the valley. Temperature inversions are best developed under calm, clear skies, where there is no wind to mix and equalize the temperatures and the transparent sky allows the surface heat to be rapidly radiated and lost to space (Blackadar 1957). Consequently, the surface becomes colder than the

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air above it, and the air next to the ground flows downslope. These slope winds are further explored on pp. 34. The cold air will continue to collect in the valley until an equilibrium between the temperatures of the slopes and the valleys has been established. If the valley is enclosed, a pool of relatively stagnant colder air may collect, but if the valley is open there may be a continuous movement of air to the lower levels, leading to the development of pollution problems (Whiteman and KcKee 1978; Nappo et al 1989). The depth of the inversion depends on the characteristics of the local topography and the general weather conditions, but it is generally not more than 300-600 m (1,000-2,000 ft.) in depth. Figure 4.14 demonstrates a temperature inversion in Gstettneralm, a small enclosed basin at an elevation of 1,270 m (4,165 ft.) in the Austrian Alps, about 100 km (62 mi.) southwest of Vienna. Because of the local topographic situation and the "pooling" of cold air, this valley experiences some of the lowest temperatures in Europe, even lower than the high peaks (Schmidt 1934). The lowest temperature recorded at Gstettneralm is -51˚C (-59.8˚F) while the lowest temperature recorded at Sormblick at 3,100 m (10,170 ft.) is -32.6˚C (-26.7˚F). As might be expected, distinct vegetation patterns are associated with these extreme temperatures. Normally, valley bottoms are forested and trees become stunted on the higher slopes, eventually being replaced by shrubs and grasses still higher up, but the exact opposite occurs here. The valley floor is covered with grass, shrubs, and stunted trees, while the larger trees occur higher up. An inversion of vegetation matches that of temperature (Schmidt 1934). A similar vegetative pattern has been found in the arid mountains of Nevada, where valley bottoms support sagebrush, while higher up is a zone of pinyon and juniper woodland. Higher still the trees again disappear (Billings 1954). The pinyon/juniper zone, the thermal belt, is sandwiched between the lower night temperatures of the valley bottom and those which occur higher up on the slopes. Human populations have taken advantage of thermal belts for centuries, particularly in the cultivation of frost-susceptible crops such as vineyards and orchards. In the southern Appalachians of North Carolina, the effect of temperature inversions is clearly displayed by the distribution of the fruit orchards (Cox 1920,1923; Dickson 1959; Dunbar 1966). During the winter, the valleys are often brown with dormant vegetation, while the mountain tops at 1,350 m (4,430 ft.) may be white with snow. In between is a strip of green that marks the thermal belt.

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Frost is common in the valley, but in the thermal belt they cultivate a sensitive Isabella grape which has apparently grown for years without danger from frost (Peattie 1936). A similar situation exists in the Hood River Valley of Oregon, on the north side of Mount Hood. Cherries are grown on the slopes of this valley in a sharply delimited thermal belt between the river and the upper slopes. With increased demand, more fruit trees are being planted in marginal areas, but their success is questionable, since the risk of frost is much greater.

Temperature Range The temperature difference between day and night and between winter and summer generally decreases with elevation (Fig. 4.15; Linacre 1982). This is because of the relatively greater distance from the heat source, the broad level of the earth's surface. Like the analogy of a marine island and the dominating influence of the ocean, the higher and more isolated a mountain, the more its temperature will reflect that of the surrounding free air. Temperature in mountains is largely a response to solar radiation. The free air, however, is essentially non-responsive to the heating effects of the sun, particularly at higher altitudes. A mountain becomes heated at the surface but there is a rapid temperature-gradient in the surrounding air. As a result, only a thin boundary layer or thermal shell surrounds the mountain, its exact thickness depending on a variety of factors (e.g., solar intensity, mountain mass, humidity, wind velocity, surface conditions, and topographic setting). Ambient temperatures are normally measured at a standard instrument-shelter height of 1.5 m (5 ft.). Such measurements generally show a progressive decline in temperature and a lower temperature range with elevation (Table 4.4; Figs. 4.11, 4.15). There is a vast difference between the temperature conditions at a height of 1.5 m (5 ft.) and immediately next to the soil surface, however. Paradoxically, the soil surface in alpine areas may experience higher temperatures (and therefore a greater temperature range) than the soil surface of low elevations, due to the greater intensity of the sun at high elevations (Anderson 1998). At an elevation of 2,070 m (6,800 ft.) in the Alps, temperatures up to 80˚C (176˚F) were measured on a dark humus surface near timberline on a southwest-facing slope with a gradient of 35˚ (Turner 1958b). This is comparable to the maximum temperatures recorded in hot deserts! At the same time, the air temperature at a height of 2 m (6.5 ft.) was only 30˚C (86˚F), a difference of 50˚C (90˚F). Such

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high surface temperatures may occur infrequently and only under ideal conditions, but temperatures somewhat less extreme are characteristic, and demonstrate the vast differences that may exist between the surface and the overlying air (Fig. 4.16). The soil surface in the alpine tundra will almost always be warmer during the day than the air above it. It may also become colder at night, although the differences are far less at night than during the day. The low growth of most alpine vegetation may be viewed as an adaptation to take advantage of these warmer surface conditions. In fact, several studies have shown that tundra plants may suffer more from high temperatures than from low temperatures (Dahl 1951; Mooney and Billings 1961). Temperature ranges vary not only with elevation, but on a latitudinal basis as well. The contrast in daily and annual temperature ranges is one of the most important distinguishing characteristics between tropical and mid-latitude or polar climates. The average annual temperatures of high tropical mountains and polar climates are similar. The average annual temperature of El Misti in Peru at 5,850 m (19,193 ft.) is -8˚C (18˚F), which is comparable to many polar stations. The use of this value alone is grossly misleading, however, since there are vast differences in the temperature regimes. Tropical mountains experience a temperature range between day and night that is relatively greater than any other mountain area, due to the strongly positive heating effect of the sun in the tropics. On the other hand, changes in temperature from month to month or between winter and summer are minimal. This is in great contrast to middle-latitude and polar mountains, which experience lower daily temperature ranges with latitude, but are increasingly dominated by strong seasonal gradients. Knowledge of the differences between these temperature regimes is essential to an understanding of the nature and significance of the physical and biological processes at work in each latitude. Figure 4.17a depicts the temperature characteristics of Irkutsk, Siberia, a subpolar station with strong continentality. The most striking feature of this temperature regime is its marked seasonality. The daily range is only 5˚C (9˚F), while the annual range is over 60˚C (108˚F). This means that during winter, which lasts from October to May, the temperatures are always below freezing, while in summer they are consistently above freezing. The period of stress for organisms, then, is concentrated into winter. An alpine station at this latitude would have essentially the same temperature regime except for a relatively longer period with negative

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temperatures and a shorter period with positive temperatures. More poleward stations would show an even smaller daily temperature range (Troll 1968). Such a temperature regime stands in great contrast to that of tropical mountains. Figure 4.17b shows the temperature characteristics of Quito, Ecuador, located on the equator at an elevation of 2,850 m (9,350 ft.). The isotherms on the graph are oriented vertically, indicating very little change between winter and summer, but with a marked contrast between day and night. The average annual range is less than 1˚C (1.8˚F), while the average daily range is approximately 11˚C (19.8˚F). This beautifully demonstrates the saying, "Night is the winter of the tropics"; night is indeed the only winter the humid tropics experience. This is particularly true if the station is high enough for freezing to occur. The lower limit of frost is determined principally by latitude, mountain mass, continentality, and the local topographic situation. In the equatorial Andes it exists at about 3,000 m (10,000 ft.). This elevation decreases with latitude; the point where frost begins to occur in the lowlands is normally taken as being the outer limits of the tropics. In North America the frost line runs through the middle of Baja California and eastward to the mouth of the Rio Grande, although it is highly variable from year to year. The frost line in tropical mountains is much more sharply delineated. In Quito, Ecuador, at 2,850 m (9,350 ft.), frost is practically unknown. The vegetation consists of tropical evergreen plants which blossom continuously; farmers plant and harvest crops throughout the year. By an elevation of 3,500 m (11,500 ft.), however, frost becomes a limiting factor (Troll 1968). At an elevation of 4,700 m (15,400 ft.) on El Misti in southern Peru, it freezes and thaws almost every day of the year. The fundamental relationships between these disparate freeze-thaw regimes are demonstrated in Figure 4.18. Each of the sites selected has a similar average annual temperature of -8˚C to -2˚C (18˚F to 28˚F) but the daily and annual ranges are markedly different. Yakutsk, Siberia, experiences strong seasonality, with a frost-free summer period of 126 days, but in winter the temperatures remain below freezing for 197 days. Alternating freezing and thawing take place during 42 days in the spring and fall. At Sormblick in the Alps, the winter season is much longer (276 days), with a very short summer during which freezing and thawing can occur at any time. El Misti, however, is dominated by a freeze-thaw regime that operates almost every day throughout the year. This type of weather has been characterized as "perpetual spring": the

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sun melts the night frost every morning and the days are quite pleasant. The twelve-hour day adds to the impression of spring (McVean 1968). It can be seen that these different systems provide greatly contrasting frameworks for the survival of plants and animals, as well as for the development of landscapes.

Humidity and Evaporation Water vapor constitutes less than 5% of the atmosphere but it is by far the single most important component with regard to weather and climate. It is highly variable in space and time. Water vapor provides energy for storms and its abundance is an index of the potential of the air for yielding precipitation; it absorbs infrared energy from the sun and reduces the amount of shortwave energy reaching the earth; it serves as a buffer from temperature extremes; and it is important biologically, since it controls the rate of chemical reactions and the drying power of the air. The moisture content of the atmosphere decreases rapidly with increasing altitude. At 2,000 m (6,600 ft.) it is only about 50 percent of that at sea level; at 5,000 m (16,400 ft.) it is less than 25%; and at 8,000 m (26,200 ft.) the water-vapor content of the air is less than 1% of that at sea level (Table 4.2). Within this framework, however, the presence of moisture is highly variable. This is true on a temporal basis, between winter and summer, day and night, or within a matter of minutes when the saturated air of a passing cloud shrouds a mountain peak (McCutchan and Fox 1986; Huntington et al. 1998). It is also true on a spatial basis, between high and low latitudes, a marine and a continental location, the windward and leeward sides of a mountain range, or north and south-facing slopes. The general upward decrease in water-vapor content, and the variations that occur, are illustrated by the east and west sides of the tropical Andes (Fig. 4.19). The contrast in absolute humidity between these two environments is immediately apparent, although the difference decreases with elevation and probably disappears altogether above the mountains. Imata, Salcedo, and Arequipa on the west have only about half the water-vapor content of stations on the east (Cerro do Pasco, Pachachaca, Huancayo, Bambamarca). Values similar to those at Arequipa occur at elevations 2,000 m (6,600 ft.) higher on the east side (e.g., at Pachachaca). During the wet season, however, the absolute humidity at Arequipa may be two to three times higher than during the dry season. This corresponds to an elevational difference of up to 3,000 m (10,000 ft.).

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The decrease in water vapor with altitude may seem somewhat difficult to explain, since it is well-known that precipitation increases with elevation. The two phenomena are not directly related, however. Precipitation results from the lifting of moist air from lower elevations upward into an area of lower temperature. Increasing precipitation does create a more humid environment in mountains, at least for part of the year and up to certain elevations, but eventually signs of aridity increase. Aridity at high elevations is due, in part, to lower barometric pressure, stronger winds, porous well-drained soils, and the intense sunlight. The greater aridity of high elevation is evident from the plants and animals, many of which have adapted to a dry environment. Thick, corky bark and waxy leaves are common in alpine plants (Isard and Belding 1986). Mountain sheep and goats and their cousins, the llama, guanaco, alpaca, chamois, and ibex, are all able to live for prolonged periods on little moisture. Geomorphologically, aeolian processes become increasingly important in higher landscapes, and the low availability of moisture is reflected in soil development (Litaor 1987). One of the physiological stresses reported by climbers on Mount Everest is a dryness of the throat and a general desiccation. The establishment of sanatoriums in alpine areas to utilize the intense sunlight and clean, dry air was mentioned earlier (Hill 1924). Air-dried meat is a provincial dish in the high Engadine, and pemmican and jerky were both important in the mountains of western North America. In the Andes, an ancient method exists for the production of dried potatoes (chuho) in the high dry air above 3,000 m (10,000 ft.). Permanent settlement of the higher elevations apparently depended upon the development of this technique of food preservation (Troll 1968). Mummification of the dead was practiced in the Andes and in the Caucasus. The lower absolute humidities and the tendency toward aridity at higher altitudes suggest greater evaporation rates with elevation. However, this may not be true, since the few studies of alpine evaporation have conflicting results (reviewed in Barry, 1992). Several studies do indicate an increase of evaporation with elevation (Hann 1903; Church 1934; Matthes 1934; Peattie 1936; Henning and Henning 1981; Sturman and Tapper 1996). For example, Matthes (1934), in discussing the development of the dimpled surfaces (sun cups) of snowfields above 3,600 m (12,000 ft.) in the Sierra Nevada of California, states that ablation (the combined processes of wasting away of snow and ice) is caused entirely by evaporation, since melting does not occur at this elevation. Two years of water balance data from a high elevation (2,800-3,400 m) lake in the

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Sierra Nevada show that evaporation accounts for 19-32% of the ablation (Kattelman and Elder 1991). Snowfall contributed 95% of the precipitation and 80% of the evaporative (sublimation) losses came from snowcover. Similar results were reported from the alpine zone of the White Mountains of California (Beatty 1975). However, other studies in the have shown that evaporation does not exceed 10% of the total ablation (Kehrlein et al. 1953). Whichever of these observations is accepted as being the more general, it should be noted that these particular alpine areas are exceptionally, if not uniquely, dry environments, with high solar intensities, strong winds, and persistent subfreezing temperatures (Terjung et al. 1969a; LeDrew 1975). The bulk of investigations on snowfields and glaciers in other regions have tended to show that evaporation is relatively unimportant in total ablation. In some cases, evaporation may actually inhibit ablation, owing to the heat it extracts (Howell 1953; Martinelli 1960; Hoinkes and Rudolph 1962; Platt 1966). In addition, long-term studies of evaporation in measurement pans and lakes at different elevations in the western United States have shown that evaporation decreases with elevation (Fig. 4.20; Shreve 1915; Blaney 1958; Longacre and Blaney 1962; Peck and Pfankuch 1963). Evaporation and the factors that control it in a natural environment are exceedingly complex (Horton 1934; Penman1963; Gale 1972; Calder 1990). The rate depends upon temperature, solar intensity, atmospheric pressure, the available quantity of water (soil moisture), the degree of saturation of the air, and wind. One of the problems in measuring the rate of evaporation is the availability of moisture. In a lake or evaporating pan, the available moisture is for all practical purposes unlimited, but this is not true for most surfaces in high mountains. Rainfall is generally lost to the surface by drainage through porous soil or by runoff on steep slopes. As a result, there is frequently little surface moisture available for evaporation, no matter how great the measured rates are from an evaporation pan. For this reason, the determination of evapotranspiration, the loss of water to the air from both plant and soil surfaces, has become an increasingly attractive approach (Thomthwaite and Mather 1951; Penman 1963; Rao et al. 1975; Henning and Henning 1981). The single most important factor in controlling the decrease of evaporation with elevation is temperature, both of the evaporation surface and of the air directly above it (Konzelman et al. 1997; Huntinton et al. 1998). While it is true that soil surfaces exposed to the sun at high elevations may reach exceptionally high temperatures, this is a highly variable condition

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(Anderson 1998; Germino and Smith 2000). During periods of high sun intensity and high soil temperatures, the potential for evaporation may be considerable, especially when the wind is blowing (Isard and Belding 1986). Generally, however, the lower temperatures of higher altitudes are more than sufficient to compensate for the decreasing water-vapor content and lower barometric pressure, so that the vapor pressure gradient is likewise decreased (Bailey et al. 1990). In other words, the relative humidity (ratio of water vapor in the air to the maximum amount it could hold at that temperature) increases with decreasing temperature, and it is the relative humidity that really determines the rate of evaporation. This is illustrated by the surprising fact that the water-vapor content of the air in the Sahara Desert is two to three times greater than that over the Rocky Mountains during clear summer weather. Owing to the higher temperatures in the Sahara, however, the relative humidity is usually not more than 20-30%, compared to 40-60% for the Rockies. Consequently, the evaporation rate in the Sahara far exceeds that of the Rockies, even though there is more actual moisture in the desert. An inverse relationship exists between air temperature and relative humidity. This can be seen by comparing measurements taken in mountains during day and night, and at various slope exposures (Fig. 4.21). The greatest contrasts occur on south-facing slopes in the northern hemisphere. Under the higher temperatures that prevail during the day, relative humidity varies very little with elevation, although it is lowest in the valley bottom. At night there is considerable contrast because of the temperature inversion that develops in the valley, resulting in lower temperatures and high relative humidities. The lowest relative humidity occurs immediately above the temperature inversion, in the thermal belt, where temperatures are higher (Hayes 1941). The difference in relative humidity between the two slopes gradually decreases with elevation. Local wind circulation can also greatly affect water-vapor content: descending air brings dry air from aloft, while ascending air carries moist air upward from below. At night colder air tends to descend through air drainage, but during the day the slopes are warmed and the air rises. Under these conditions, the normal inverse temperature/relative-humidity relationship may be overridden. Even though the summit air is cool at night, the motion of the descending air lowers the relative humidity. During the day, however, when temperatures are higher and

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relative humidity would normally decrease, it may actually increase, because the valley breezes carry moist air up the mountain slopes. This frequently results in afternoon clouds and precipitation (Schell 1934).

Precipitation The increase of precipitation with elevation is well-known. It is demonstrated in every country of the world, even if the landforms involved are only small hills. In many regions an isohyetal map with its lines of equal precipitation will look similar to a topographic map composed of lines of equal elevation (Fig. 4.22). Of course, the data on which most precipitation maps are based are scanty, so that considerable interpolation may be necessary, particularly in the areas of higher relief (Peck and Brown 1962; Kyriakidis et al. 2001). Precipitation does not always correspond to landforms. In some cases, maximum precipitation may occur at the foot or in advance of the mountain slopes (Reinelt 1968; Barry 1992). In some regions and under certain conditions, valleys may receive more rainfall than the nearby mountains (Sinclair et al. 1997). In many higher alpine areas, precipitation decreases above a certain elevation, with the peaks receiving less than the lower slopes. Wind direction, temperature, moisture content, storm and cloud type, depth of the air mass and its relative stability, orientation and aspect, and configuration of the landforms are all contributing factors in determining location and amount of precipitation (Sinclair 1994; Ferretti et al. 2000; McGinnis 2000; Drogue et al. 2002). The complex topographic arrangement and often high relief of mountains creates complex meso- and micro-scale three-dimensional circulation and cloud formations, leading to complex spatial patterns of precipitation within mountainous regions (Bossert and Cotton 1994; Cline et al. 1998; Garreaud 1999; Germann and Joss 2001; 2002). Great variations in precipitation occur within short distances; one slope may be excessively wet while another is relatively dry. The terms "wet hole" and "dry hole" may be used in this regard. Jackson Hole, Wyoming, is located in a protected site at the base of the Grand Tetons. The mountains receive 1,400 mm. (55 in.) but Jackson Hole, only 16 km (10 mi.) away, receives 380 mm (15 in.). The most fundamental reason for increased precipitation with elevation is that landforms obstruct the movement of air and force it to rise. This is part of a complex of processes known as the orographic effect (from the Greek oros, meaning "mountain," and graphein, "to describe").

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Forced ascent of air is most effective when mountains are oriented perpendicular to the prevailing winds; the steeper and more exposed the slope, the more rapidly air will be forced to rise. As air is lifted over the mountains it is cooled by expansion and mixing with cooler air at higher elevations. The ability of air to hold moisture depends primarily upon its temperature, warm air can hold much more moisture than cold air. The temperature, the pressure, and the presence of hygroscopic nuclei in the atmosphere tend to concentrate the water vapor in its lower reaches. This is why most clouds occur below 9,000 m (30,000 ft.), and why those that do develop higher than this are usually thin and composed of ice particles and yield little or no precipitation. When the air holds as much moisture as it can (relative humidity is 100%), it is said to be saturated. Condensation is a common process in saturated air, and the temperature at which condensation takes place is called the dew point. Ground forms of condensation, i.e., fog, frost, and dew, are caused by cooling of the air in contact with the ground surface, but condensation in the free atmosphere, i.e., clouds, can only result from rising air. The key to forming clouds and creating precipitation, therefore, is rising air. This may be brought about by one of several ways. The driving force may be convection (thermal heating), where the sun warms the earth's surface and warm air rises until clouds begin to form. Such clouds may grow to great size since they are fed from below by relatively warm, moist rising air, until the moisture content within the clouds becomes too great and it is released as precipitation. Convectional rainfall is best displayed in the humid tropics where water vapor is abundant, but it occurs in all climates. The air may also be forced to rise by the passage of cyclonic storms, where warm and cold fronts lift moist, warm air over cooler, denser air. This takes place primarily in the middle latitudes in association with the polar front (Fig. 4.2). Although both of these processes can operate without the presence of mountains, their effectiveness is greatly increased on windward sides of mountains and decrease on leeward sides. For example, a passing storm may drop a certain amount of precipitation on a plains area, but when the storm reaches the mountains, a several-fold increase in precipitation typically occurs on the windward side, while a marked decrease generally takes place on the leeward side. One has only to compare the distribution of world precipitation with the location of mountains to see their profound influence (Fig. 4.22). Almost every area of heavy rainfall is

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associated with mountains. In general, any area outside the tropics receiving more than 2,500 mm (100 in.) and any area within the tropics receiving more than 5,000 mm (200 in.) is experiencing a climate affected by mountains. The examples of Cherrapunji, Assam; Mount Waialeale, Hawai‘i; and the Olympic Mountains were given earlier. Many others could be added: Mount Cameroon, West Africa, the Ghats along the west coast of India, the Scottish Highlands, the Blue Mountains of Jamaica, Montenegro in Yugoslavia, and the Southern Alps of New Zealand. The list could go on and on. The reverse is also true, for to the lee of each of these ranges is a rain-shadow in which precipitation decreases drastically (Manabe and Broccoli 1990; Broccoli and Manabe 1992). The western Ghats receive over 5,000 mm (200 in.) but immediately to their lee on the Deccan Plateau the average amount of precipitation is only 380 mm (15 in.). The windward slopes of the Scottish Highlands receive over 4,300 mm (170 in.) but the amount decreases to 600 mm (24 in.) on the lowlands around the Moray Firth. The Blue Mountains on the northeast side of Jamaica face the Trade Winds and receive over 5,600 mm (220 in.), while Kingston, 56 km (35 mi.) to the leeward, receives only 780 mm (31 in.) (Kendrew 1961). Mountains, therefore, not only cause increased precipitation, but also have the reciprocal effect of decreasing precipitation. Despite these useful generalities, many local and regional variations occur within mountains. The complex local topography creates funneling effects that can increase atmospheric moisture content and precipitation, even downwind from the funnel (Sinclair et al. 1997). High peaks or ridges within a range can create ‗mini-rain-shadow‘ zones even in the center of a range (Garreaud 1999). The central portion of the north Cascades Range in Washington receives ~100 cm less precipitation than the surrounding ridges (Kresch 1994). Significant quantities of precipitation can fall on the leeside of mountains due to spillover effects (Sinclair et al. 1997; Thompson et al. 1997; Chater and Sturman 1998). Many precipitation maps of mountainous regions do not indicate this internal variability due to the lack of data points and interpolation between existing station data using generalized elevation-precipitation relationships (Kyriakidis et al. 2001). Additionally, seasonal and interannual variability of storm tracks and storm intensities can create non-elevational precipitation patterns (Lins 1999). The movement of air up a mountain slope, creating clouds and precipitation, may be due simply to the wind, but it is usually associated with convection and frontal activity. Rising air

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cools at a rate of 3.05˚C (5.5˚F) per 300 m (1,000 ft.) (dry adiabatic rate) until the dew point is reached and condensation occurs (Fig. 4.39). Thereafter, the air will cool at a slightly lower rate (wet adiabatic rate) because of the release of the latent heat of condensation. If, upon being lifted, the air has a high relative humidity, it may take only slight cooling to reach saturation, but if it has a low relative humidity it may be lifted considerable distances without reaching the dew point. Conversely, if the air is warm, it often takes considerable cooling to reach dew point but then may yield copious amounts of rainfall, whereas cool air usually needs only slight cooling to reach dew point but also yields far less precipitation. After the air has passed over the mountains precipitation decreases or may cease as the air descends. As air descends it gains heat at the same rate at which it was cooled initially 3.05˚C per 300 m (5.5˚F per 1,000 ft.), since it is being compressed and moving into warmer air (Fig. 4.39). Such conditions are not conducive to precipitation. The orographic effect involves several distinct processes: (1) forced ascent, (2) blocking (or retardation) of storms, (3) the triggering effect, (4) local convection, (5) condensation and precipitation processes, and (6) runoff.

Forced Ascent Forced ascent is the most important precipitation process in mountains; after all, rainfall increases with elevation and is greater on windward than on leeward slopes. The process may be most clearly seen in coastal mountains, like the Olympics, that lie athwart moisture-laden winds. Other processes contribute to the total precipitation, of course, and differentiation among them is difficult. In order to explain the amount and distribution of rainfall caused strictly by forced ascent it is necessary to consider the atmospheric conditions from three different perspectives (Sawyer 1956; Sarker 1966; Browning and Hill 1981). First is the large-scale synoptic pattern that determines the characteristics of the air mass crossing the mountains, its depth, stability, moisture content, wind speed, and direction (Sinclair 1994; McGinnis 2000). Second is the microphysics of the clouds, the presence of hydroscopic nuclei, the size of the water droplets, and their temperature, which will determine whether the precipitation will fall as rain or as snow or will evaporate before reaching the ground (Andersson 1980; Meyers et al 1995; Uddstrom et al. 2001). Third, and most important, is the air motion with respect to the mountain (Bates, 1990;

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Bossert and Cotton 1994; Tucker and Crook 1999). Will it blow over, or around, the mountain? This will determine to what depth and extent the air mass at each level is lifted. It is not realistic, for example, to assume that the air is lifted the same amount at all levels. The solution to these problems involves atmospheric physics and the construction of dynamic models (Myers 1962; Sarker 1966, 1967; Sinclair 1994; Thompson et al. 1997; Susong et al. 1999; Drogue et al. 2002). The simplest system is that of coastal mountains with moisture-laden winds approaching from the ocean. As the air is lifted from sea-level, the resulting precipitation is clearly due to the landforms (Colle and Mass, 1996). Exceptions may occur in areas where the mountains are oriented parallel to the prevailing winds and/or where the frontal systems resist lifting. In southern California, for example, precipitation is often heavier in the Los Angeles coastal lowlands than in the Santa Inez and San Gabriel mountains due to the blocking of storms. The orographic component of precipitation increases only when the approaching air mass is unstable; under stable conditions, the wind will flow around the mountains (which are oriented east-west), so there is no significant orographic lifting and the precipitation is due entirely to frontal lifting. The mountains apparently receive less rainfall than the lowlands under these conditions because the shallow cloud-development does not allow as much depth for falling precipitation particles to grow by collision and coalescence with cloud droplets before reaching the elevated land. The situation becomes even more complex in interior high elevation areas where there is more than one source region and storms enter the area at various levels in the atmosphere. Such a situation exists in the Wasatch Mountains of Utah (Williams and Peck 1962; Peck 1972a; Sassen and Zhao 1993). It has long been known that precipitation in this region is highly variable; the valleys may receive greater amounts than the mountains during any given storm or season (Clyde 1931). The average over a period of years, however, does show an increase with elevation (Price and Evans 1937; Lull and Ellison 1950). The greater precipitation in valleys is apparently associated with certain synoptic situations, particularly when a "cold low" is observed on the upper-air charts. These occur as closed lows on the 500-millibar pressure chart, i.e., at a height of about 5,500 m (18,000 ft.), and are associated with large-scale upward (vertical) movement of air which is not displayed in normal cold or warm-front precipitation (Schultz et al. 2002). Under these conditions, precipitation may occur with relatively little dependence on orographic lifting, compared to other storm types (Williams and Peck 1962).

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Blocking of storms By retarding or hindering the free movement of storm systems, mountains can cause increased precipitation (Kimura and Manins 1988). Storms often linger for several days or weeks as they slowly move up and over the mountains, producing a steady downpour (Kimura and Manins 1988; Gan and Rao 1994). This is best displayed in the middle latitudes with high-barrier mountains. Winter storms linger with amazing persistence in the Cascades and in the Gulf of Alaska before they pass across the mountains or are replaced by another storm. Storms of similar character in the Great Plains travel much more rapidly, since there are no restrictions to their movement. The countries surrounding the Alps are ideally located with respect to storm blocking. Switzerland frequently experiences lingering torrential rains during the summer (Bonacina 1945; Chen and Smith 1987). In northern Italy, between the Alps and the Apennines, heavy and persistent rains are associated with the "lee depressions" caused by the interception of polar air by the Alps (Grard and Mathevet 1972; Pichler and Steinacker 1987).

The Triggering Effect Although little mention has so far been made of it, one important variable influencing the amount of precipitation is the stability of the air, that is, its resistance to vertical displacement. This is controlled primarily by temperature. When there is a low environmental lapse rate, i.e., less than 1.4˚C per 300 m (2.5˚F per 1,000 ft.), as there frequently is at night in mountains, the air is stable. Stable air resist lifting in mountains will often move down slope. During the day, when the sun warms the slopes and the surface air is heated, the environmental lapse rate increases and the air will have a tendency to rise, frequently producing afternoon clouds. When the lapse rate exceeds the dry adiabatic rate of 3.05˚C per 300 m (5.5˚F per 1,000 ft.), a condition of absolute instability prevails. Under these conditions even slight lifting of the air by a landform is enough to "trigger" it into continued lifting on its own accord. If it then begins to feed upon itself through the release of latent heat of condensation, it can yield considerable precipitation (Bergeron 1965; Thornthwaite 1961; Revell 1984). As a result of this effect, thunderstorms can develop, even on small hills in the path of moist unstable air (Schaaf et al. 1988).

Local Convection

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

Clouds commonly form over mountains during the day, especially in the summer, when nights and early mornings are clear but by mid-morning clouds begin to build, often culminating in thunderstorms with hail and heavy rain (Fuquay 1962; Baughman and Fuquay 1970; Flohn 1974). This has been well-documented for the base of the Colorado Rockies, where the higher peaks of the Front Range provide a "heated chimney effect" in the initiation of thunder and hailstorms (Harrison and Beckwith 1951; Beckwith 1957; Banta and Schaaf 1987). Mountains serve as elevated heat-islands during the day, since their surfaces can be warmed to a similar temperature as surrounding lowlands (Raymond and Wilkening 1980). As a consequence, the air at a given altitude is much warmer over the mountains than over the valley (MacCready 1955). The lapse rate above the peaks, therefore, is considerably greater than in the surrounding free air, resulting in actively rising air. Glider pilots have long taken advantage of this fact (Scorer 1952, 1955; Ludlam and Scorer 1953). Airline pilots, on the other hand, make every effort to avoid the turbulence associated with unstable air over mountains (Reiter and Foltz 1967, Colson 1963,1969). Rarely, given weak synoptic conditions, local mountain convection can become organized into a meso-scale convective complex (Tucker and Crook 1999). These strong storms can reinforce themselves, spawning severe thunderstorms and even tornadoes (mountainadoes). Clouds and thunderstorms initiated in the Front Range frequently drift eastward, continuing to develop as they move onto the plains, and producing locally heavy precipitation (Chung et al. 1976; McGinley 1982). A study in the San Francisco Mountains north of Flagstaff, Arizona, suggests that clouds may increase in volume by as much as ten times after drifting away from a mountain source (Glass and Carlson 1963; Banta and Schaaf 1987). Most of the clouds observed in this area were small cumuli that eventually dissipated once removed from their supply of moist, rising air, but a large cumulonimbus could maintain itself independently of the mountains and result in storms at some distance away. Fujita (1967) found that there was a ring of low precipitation about 24 km (15 mi.) in diameter encircling these mountains, with an outer ring of heavier precipitation. During the day the rainfall is over the mountains but at night it falls over the lowlands because the mountains are relatively cold. A "wake effect" due to wave action created by airflow over the mountains may be partly responsible for the inner ring of light precipitation (Fujita 1967). A similar phenomenon occurs adjacent to the Rockies, on the Great

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

Plains, where there is a second peaking of thunderstorm activity in the early evening (Bleeker and Andre 1951).
Well-studied mountain convection phenomena are found in the San Francisco and Santa Catalina Mountains of Arizona. A number of studies have traced the initiation and development of convection and cumulus clouds over the range (Braham and Draginis 1960; Orville 1965; Fujita 1967). Figure 4.23 shows the change in temperature and moisture content over the Santa Catalina Mountains from early morning to midmorning. Note that the south-facing slopes show considerably more thermal convection than the north-facing slopes. On this particular day the base of the clouds was about 4,500 m (15,000 ft.), so the sun was not blocked and could continue to shine on the slopes to feed the thermal convection (Braharn and Draginis 1960).

The height of the cloud base is very important to the development of convection in mountains, since once the sun is blocked the positive effect of solar heating is eliminated. The height of the cloud base is also critical to the distribution of precipitation, as is demonstrated in the San Gabriel Mountains, California (see p. 95). If the cloud base is below the level of the peaks, as it usually is in the winter, when forced ascent occurs, cloud growth and precipitation will take place mainly on the windward side. In summer, however, the base of convection clouds is generally much higher. Mountains, as sites of natural atmospheric instability, are ideal areas for artificial stimulation of precipitation. The considerable efforts that have been made in this regard have met with varied success, depending upon technique and local atmospheric conditions (Mielke et al. 1970; Chappell et al. 1971; Hobbs and Radke 1973; Grant and Kahan 1974; Deshler et al. 1990; Meyers et al. 1995; Long and Carter 1996). Most of the projects have been aimed at increasing the snowpack for runoff during the summer. This appears to be a desirable objective, but the ecological implications of such undertakings are far-reaching (Weisbecker 1974; Steinhoff and Ives 1976). For example, the Portland General Electric Company of Portland, Oregon, hired a commercial firm during the winter of 1974/75 to engage in cloud-seeding on the eastern side of the Cascades. The objective was to increase the snowpack in the Deschutes River watershed, where they have two dams and power-generating plants. Considerable success was apparently achieved, but problems arose when residents of small towns at the base of the mountains were suddenly faced with a marked increase in snow. There were new problems of transportation and of snow removal, as well as other hardships for the local people. Greater snowfall meant greater profits for the power company but it also meant greater expenses for the local people. Objections were raised in the courts, and the project was eventually halted. The positive effects of such programs must always be balanced against the negative. In our efforts to manipulate nature we

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

are made increasingly aware of how little we understand the effects of our actions on natural systems. This is especially true of the mountain environment (Steinhoff and Ives 1976).

Condensation Processes The presence of fog or clouds near the ground may result in increased moisture. Water droplets in fog and clouds are usually so small that they remain suspended, and even a slight wind will carry them through the air until they strike a solid object and condense upon it. You have experienced this, if water droplets have ever formed on your hair and eyebrows as you passed through a cloud or fog. Fog drip and rime deposits, which form at subfreezing temperatures, are responsible for an appreciable amount of the moisture in mountains, since elevated slopes are often in contact with clouds.

Clouds. Cloud cover is generally more frequent and thicker over mountains than over the surrounding lowlands (Uddstrom et al. 2001). Forced lifting of moist air mass over the topographic barrier is the primary cause, although it may be augmented by convective processes. A slowing of storm movement by the blocking effect also leads to an increase in cloud watercontent (Pedgley 1971). Cloud type in mountain areas is primarily determined by synoptic characteristics. In middle and high latitudes, stratiform clouds are common, especially during winter in the absence of convection. These clouds often envelope the ground as hill fog. Middle latitude summers, continental, subtropical, and tropical areas typically have cumulus clouds associated with convection. A problem relating to cloud data from mountain stations is the clouds often engulf the observer, obstructing the view of the cloud forms. Likewise, cloud tops can be below the station. A number of cloud forms are unique to mountain environments (Ludlam 1980). All of them are stationary clouds, which continually dissipate on the lee edge of the cloud and reform on the upwind edge, thus appear to remain in the same location for long periods. A cap or crest cloud forms over the top of an isolated peak or ridge. They resemble a cumulous cloud, although are often streamlined, or have streamers of cirrus forms. They sit near or just below the summit level, appearing like a hat atop the peak. Banner clouds are cap cloud which extent downwind from the peak like a flag waving in the wind. This form is sometimes difficult to distinguish from

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streamers of snow blowing from summits. Lenticular clouds are lens-shaped clouds formed in regular spaced bands parallel to the mountain barrier on the lee side (Figs. 4.39-4.42). These streamlined cloud features form by the interaction of high velocity winds with the mountain barriers (see p. 119). Stratification of humidity in the atmosphere can result in multi-storied lenticular clouds, forming a ‗pile of plates‘ or ‗pile of pancakes‘ (Fig. 4.42). These sometimes eerie looking clouds might be responsible for the ‗flying saucer‘ scare of the 1950s, which originated from a sighting of ―a disc-shaped craft skimming along the crest of the Cascades Range in Washington‖ (Arnold and Palmer, 1952).

Fog Drip. Fog drip is most significant in areas adjacent to oceans with relatively warm, moist air moving across the windward slopes. In some cases, the moisture yield from fog drip may exceed that of mean rainfall (Nagel 1956). The potential of clouds for yielding fog drip depends primarily upon their liquid content, the size of the cloud-droplet spectrum, and the wind velocity (Grunow 1960; Vermeulen et al. 1997). The amount that occurs at any particular place depends upon the nature of the obstacles encountered and their exposure to the clouds and wind. For example, a tree will yield more moisture than a rock, and a needle-leaf tree is more efficient at "combing" the moisture from the clouds than a broadleaf tree (Cavelier and Goldstein 1989; Vermeulen et al. 1997). A tall tree will yield more moisture than a short one, and a tree with front-line exposure will yield more than one surrounded by other trees. The tiny fog droplets are intercepted by the leaves and branches and grow by coalescence until they become heavy enough to fall to the ground, thereby increasing soil moisture and feeding the ground-water table. If the trees are removed, of course, this source of moisture is also eliminated. Many tropical and subtropical mountains sustain so-called "cloud forests," which are largely controlled by the abundance of fog drip (Cavelier and Goldstein 1989). Along the east coast of Mexico in the Sierra Madre Oriental, luxuriant cloud forests occur between 1,300-2,400 m (4,300-7,900 ft.). The coastal lowlands are arid by comparison, as is the high interior plateau beyond the mountains. Measurements in both of these drier areas show little increase in available moisture due to fog drip, whereas on the middle and upper slopes the process boosts moisture by more than 50% at one site located at 1,900 m (6,200 ft.), the increase over rainfall was 103%

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

(Vogelman 1973). These cloud forests were at one time much more extensive, but they have been severely disturbed by humans and are now in danger of being eliminated. On the northeast slopes of Mauna Loa, Hawai‘i, at 1,500-2,500 m (5,000-8,200 ft.), above the zone of maximum precipitation, fog drip is likewise a major ecological factor in the floristic richness of the forests. During a twenty-eight-week study, fog drip was found to provide 638 mm (25.3 in.) of moisture at an elevation of 1,500 m (5,000 ft.); and at 2,500 m (8,200 ft.) it provided 293 mm (11.5 in.), which was 65% of the direct rainfall (Fig. 4.24; Juvik and Perreira 1974; Juvik and Ekern 1978). The contribution of fog drip on middle and upper mountain slopes in the lower latitudes is clearly a major factor in the moisture regime. The relationship between the cloud forest and fog drip is essentially reciprocal. The trees cause additional moisture in the area. At the same time, the trees apparently need the fog drip in order to survive. This is particularly true in areas with a pronounced dry season, at which time fog drip provides the sole source of moisture for the plants. In the middle latitudes, fog drip is less critical to the growth of trees, but it can still be important (Grunow 1955; Costin and Wimbush 1961; Vogelmann et al. 1968). This can be seen in the mountains of Japan, where there is heavy fog at intermediate altitudes (Fig. 4.25).

Rime Deposits. Rime is formed at subfreezing temperatures when supercooled cloud droplets are blown against solid obstacles, freezing on them (Hindman 1986; Berg 1988). Rime deposits tend to accumulate on the windward side of objects (Fig. 4.26). The growth rate is directly related to wind velocity. In extreme cases the rate of growth may exceed 2.5 cm (1 in.) per hour, although a typical rate is usually less than 1 cm (0.4 in.) per hour (Berg 1988). Rime deposits can reach spectacular dimensions and, by their weight, cause considerable damage to tree branches, especially if followed by snow or freezing rain. Trees at the forest edge and at timberline frequently have their limbs bent and broken by this process; power lines and ski lifts are also greatly affected (Fig. 4.26). One study in Germany measured a maximum hourly growth of 230 g per m (8.1 oz. per 3.3 ft.) on a power-line cable (Waibel 1955, in Geiger 1965). The stress caused by this added weight may cause a power failure if the supporting structures are not properly engineered.

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Rime accumulation is a severe obstacle to the maintenance of mountain weather stations because instruments become coated, making accurate measurements extremely difficult. Some instruments can be heated or enclosed in protected housing, but the logistical problems of accurately monitoring the alpine environment are very great. The U.S. Weather Bureau Station on Mount Washington, New Hampshire, where rime-forming fogs are frequent and the wind is indefatigable, exemplifies the problems encountered (Smith 1982). This mountain has been nominated as having the worst weather in the world (Brooks 1940). It is foggy over 300 days a year, or about 87% of the time; wind velocities there average 18 m/sec. (40 mph) with frequent prolonged spells of 45 m/sec. (100 mph) and occasional extremes of over 90 m/sec. (200 mph) (Pagliuca 1937; Smith 1982). Few investigations have been made concerning the moisture contribution of rime. It is known to be generally somewhat less than fog drip, but it may nevertheless be significant. A study in the eastern Cascades of Washington indicates that timbered areas above 1,500 m (5,000 ft.) receive an added 50-125 mm (2-5 in.) of moisture per year from this source (Berndt and Fowler 1969). Considerably greater amounts have been measured in Norway (Table 4.5). Rime is found primarily in middle-latitude and polar mountains, although it also occurs at the highest elevations in the tropics. Like fog drip, it is most effective on forest-covered slopes that provide a large surface area for its accumulation. At very high altitudes and at latitudes where total precipitation is low, rime deposits on glaciers and snowfields may constitute the primary source of the water taken from the air.

Zone of Maximum Precipitation Precipitation is generally thought to increase only up to a certain elevation, beyond which it decreases (Lauer 1975; Miller 1982). The argument is that the greatest amount of precipitation will usually occur immediately above the cloud level because most of the moisture is concentrated here. As the air lifts and cools further, the amount of precipitation will eventually decrease, because a substantial percentage of the moisture has already been released on the lower slopes (Miniscloux et al. 2001). In addition, the decreased temperature and pressure at higher elevations reduce the capacity of the air to hold moisture. The water-vapor content at 3,000 m (10,000 ft.) is only about one-third that at sea level. Forced ascent also plays a part, since the air,

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seeking the path of least resistance, will generally move around the higher peaks rather than over them. The concept of a zone of maximum precipitation was developed over a century ago from studies in tropical mountains and in the Alps (Hann 1903). Other studies seemed to confirm the concept and its application to other areas (Lee 1911; Henry 1919; Peattie 1936; Lauer 1975). The elevation of maximum precipitation varies geographically, depending upon the synoptic setting (Barry 1992; McGinnis 2000). Tropical mountains tend to have precipitation maxima at lower elevations, with the maximum zone rising with decreasing annual totals (Fig. 4.28). In middle latitudes, the general trend is for precipitation to increase with elevation, often to the highest observation station (Schermerhorn 1967; Hanson, 1982; Alpert, 1986; Marwitz 1987). Using precipitation and accumulation data for western Greenland, a zone of maximum precipitation was found at ~2,400 m at 69˚N latitude and lowering northward to ~1,500 m at 76˚N (Ohmura 1991). The existence of such a zone has been challenged, as calculations of the amount of precipitation necessary to maintain active glaciers in high mountains and observations of relatively heavy runoff from small alpine watersheds seem to call for more precipitation in certain mountain areas than climatic station data would indicate (Court 1960, Anderson 1972; Slaymaker 1974). Currently, the situation is moot, the problem being one of measurement. There are very few weather stations in high mountains, and even where measurements are available, their reliability is questionable (Sevruk 1986; 1989). As one author says, "Precipitation in mountain areas is as nearly unmeasureable as any physical phenomenon" (Anderson 1972, p. 347). This is particularly true at high altitudes with strong winds. Not surprisingly, many studies have shown that wind greatly affects the amount of water collected in a rain gauge (Fig. 4.27; Court 1960; Brown and Peck 1962; Hovind 1965; Rodda 1971). Considerable effort has been made to alleviate this problem by the use of shields on gauges, by location in protected sites, by use of horizontal or inclined gauges, and by the use of radar techniques (Storey and Wilm 1944; Harrold et al. 1972; Peck 1972b; Sevruk 1972; Rango et al. 1989). To measure snow is even more difficult, since the wind not only drives falling snow but redistributes it after it is on the ground (Goodinson et al. 1989). Correction factors have been developed for certain types of gauges (Goodinson et al. 1989; Sevruk 1989; Kyriakidis et al.

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2001). There are also problems in storage and melting of snow for water equivalency, as well as the losses due to evaporation. The major problem, however, is accurate monitoring of snowfall. Small clearings are used in conifer forests, and above timberline snow fences are increasingly being used to enclose and shield the gauges. This still does not guarantee accurate measurements, but shielded gauges (whether for rain or snow) do record greater amounts of precipitation than unshielded gauges in the same location (Goodinson et al. 1989). For example, the University of Colorado has since 1952 operated a series of weather stations in the Front Range of the Rocky Mountains (Marr 1967; Marr et al. 1968a, b). The measured precipitation amounts from the two highest sites above treeline increased abruptly in 1964 when snow fences were erected around the recording gauges. Before the gauges were shielded, the average annual amount was 655 mm (25.8 in.); it jumped to 1,021 mm (40.2 in.) and 771 mm (30.3 in.), respectively, after the snow fence was installed (Barry 1973). The data now show an absolute increase in precipitation with increasing elevation (Table 4.6). More reliable instrumentation in the Alps has led to similar results, at least up to an elevation of 3,000 m (10,000 ft.) (Flohn 1974; Schmidli et al. 2002). Studies of snow accumulation at still higher elevations, in the Saint Elias Mountains, Yukon Territory, indicate decreasing amounts beyond 3,000 m (10,000 ft.), although there is a steady increase at lower elevations at least up to 2,000 in (6,000 ft.) (Murphy and Schamach 1966; Keeler 1969; Marcus and Ragle 1970, Marcus 1974b). Snow accumulation in alpine watersheds can be investigated more thoroughly by collecting depth and density data from snow pits, which can be converted to water equivalent (Østrem and Brugman 1991). In North America, an extensive network of over 1,200 snow courses are surveyed on a monthly by the Natural Resources Conservation Service (NRCS, formally the Soil Conservation Service) of the U.S. Department of Agriculture (NRCS 1997). Snow courses are where depth and density is measured manually to estimate annual water availability, spring runoff, and summer streamflows. In more remote locations of the western United States, a network of over 650 automated snow reporting stations have been installed. The SNOTEL (Snow Telemetry) network uses an air filled pillow attached to a pressure gauge to measure snowpack weight, which is transmitted via VHF signals to a data collection station (NRCS 1997). Combining field measurements of snow-water-equivalency with topographic data

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(slope and aspect) and net radiation, estimates of watershed snowpack water content can be modeled (Elder et al. 1989; Susong et al. 1999). Another problem with precipitation analysis in mountains is that many weather stations are located in valleys. Uncritical use of these data may lead to erroneous results (Benizou 1989). Valleys oriented parallel to the prevailing winds may receive as much or more precipitation than the mountains on either side, while valleys oriented perpendicular to the prevailing winds may be "dry holes" (Collie and Mass 1996; Neiman et al. 2002). In addition, local circulation systems between valleys and upper slopes may result in valleys being considerably drier than the ridges (see p. 111). For example, in parts of the Hindu Kush, Karakoram, and Himalayas, many valleys are distinctly arid (Schweinfurth 1972; Troll 1972b). These contrast sharply with the adjacent mountains, where large glaciers exist. Some glaciologists have estimated an average annual precipitation of over 3,000 mm (120 in.) for the glacial area, compared to 100 mm (4 in.) in the valleys, data that seem to support the idea of a steady increase of precipitation with elevation (Flohn 1968, 1969a, 1970). On the other hand, it is argued that little precipitation is required to maintain a glacier under such low temperatures, owing to the relatively small losses to be expected through ablation (Hock et al. 2002). Several studies have provided evidence for a zone of maximum precipitation at about 2,000 m (6,600 ft.) along the southern slope of the Himalayas (Dhar and Narayanan 1965; Dalrymple et al. 1970; Khurshid Alam 1972). The high, sheltered inner core of the Himalayas is arid (Troll 1972c). In the tropics, decrease of precipitation above a certain elevation is much better established (Fig. 4.28; Lauer 1975). The precipitation falls principally as rain, with snow or rime on the highest peaks, and tropical mountains experience considerably less wind than in middle latitudes. As a result, simple rainfall measurements are more dependable. The zone of maximum precipitation varies according to location. In the tropical Andes and in Central America it lies between 900-1,600 m (3,000-5,300 ft.) (Hastenrath 1967; Weischet 1969; Herrmann 1970). Mount Cameroon in West Africa near the Gulf of Guinea receives an annual rainfall of 8,950 mm (355 in.) on the lower slopes but less than 2,000 mm (80 in.) at the summit. The zone of maximum precipitation occurs at 1,800 m (6,000 ft.) (Lefevre 1972). In East Africa, measurements on Mount Kenya and Mount Kilimanjaro show an increase up to the montane forest belt at 1,500 m (5,000 ft.) and then a sharp decrease (Fig. 4.28). The maximum zone receives about 2,500 mm (100 in.) but less than 500 mm (20 in.) falls on the summit areas. The effects of low rainfall, high sun-intensity, and porous soils give the alpine belt a desert-like appearance, although both summit areas support small glaciers (Hedberg 1964; Thompson 1966; Coe 1967). Desert-like conditions exist at the summits of many tropical mountains (Fig. 4.29). On the islands of Indonesia and on Ceylon the zone of maximum precipitation varies between 900-1,400 m (3,000-4,600 ft.) (Domrös 1968; Weischet 1969), while it lies between 600-900 m

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(2,000-3,000 ft.) in Hawai‘i (Blumenstock and Price 1967; Juvik and Peffeira 1974; Nullet and McGranaghan 1988). The decrease immediately above the zone of maximum precipitation is counteracted somewhat by the presence of fog drip, however, since this is a zone of frequent cloudiness (Fig. 4.24). The vertical distribution of precipitation illustrates yet another environmental distinction between tropical and extratropical mountains. The presence of a zone of maximum precipitation is well established for the tropics, but is less defined in the middle latitudes. Although there are insufficient measurements to settle the question categorically, evidence from mass-balance studies on glaciers, runoff from mountain watersheds, and improved methods of instrumentation seem to indicate that precipitation continues to increase with altitude in middle latitudes at least up to 3,000-3,500 m (10,000-11,000 ft.). The decrease beyond moderate elevations in the tropics is explained by the dominance there of convection rainfall, which means that the greatest precipitation occurs near the base of the clouds. Where forced ascent is important the level may be somewhat higher, but it does not vary over a few hundred meters. In many tropical areas an upper air inversion composed of dry, stable air tends to restrict the deep development of clouds. This is the case on Mount Kenya and on Kilimanjaro, as well as on Mauna Loa and Mauna Kea in Hawai‘i (Juvik and Perreira 1974; Ramage and Schroeder 1999). The continued increase of precipitation with elevation in the middle latitudes is somewhat more difficult to explain. The water-vapor content of the air decreases at the higher levels just as it does in the tropics. Precipitation in middle-latitude mountains is caused primarily by forced ascent rather than convection, however. Orographic lifting becomes stronger as the wind grows stronger, and wind velocity increases markedly in middle latitudes. Apparently, this factor is more than enough to compensate for the absolute decrease in water content. The water-vapor transport is at its maximum in the westerlies at about the 700-millibar level, i.e., at 3,000 m (10,000 ft.). Consequently, it is postulated that precipitation continues to increase at least up to this level (Havlik 1969). In tropical mountains, however, the wind tends to decrease with elevation above 11000 m (3,300 ft.), so the decrease in water content of the air becomes more effective and precipitation decreases beyond this point (Weischet 1969; Rohn 1974).

Runoff

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Mountain surface runoff is related to topographic, biotic, pedologic, and particularly the climatic characteristics of a watershed (Miller 1982; Alford 1985). In particular, seasonality of precipitation inputs, temperatures, snowpack characteristics, and non-precipitation water sources (i.e. groundwater and glaciers) are important variables in determining the amount of water flowing down a mountain stream (Martinec 1989; Lins 1999; Peterson et al. 2000). Globally, mountain runoff displays significant temporal and spatial variation. The temporal heterogeneity arises from the intraannual, interannual, and secular changes in temperature, precipitation and other climatic factors (Rebetez 1995; Lins 1999). Spatial heterogeneity is due to climatic, topographic, biotic, land-use, and pedologic variability within and between mountains, making generalities about mountain hydrology difficult. The two generalities of mountain hydrology are the generation of flood events and the influences of snowpack meltwater. Steep slopes, generally thin soils and generally high rain intensities in mountains often result in elevated rates of overland flow compared lowlands (Miller 1982; Dingman 1993). During lingering or moisture-ladened storms the delivery rate of runoff to main rivers can exceed the channel capacity and a flood results. Flooding typically occurs with regular intervals for a particular watershed, determined by its hydrologic characteristics (Castro and Jackson 2001). While floods can cause damage to mountain valleys, often their impacts are felt more strongly in the lowlands adjacent to the mountains, where flood magnitude is often larger due to the cumulative flow of many tributaries. Snow meltwater has four principal impacts on watershed hydrology: lowering stream temperature, sudden contributions to discharge resulting from rapid melting (rain on snow events), an increase in melt-season discharge and decrease in snow-accumulation season discharge, and a decrease in annual and especially seasonal variations in runoff (Male and Gray 1981). The changes in average seasonal discharge due to snowmelt are illustrated in Figure 30. Differences in discharge in these side-by-side mountainous watersheds of the same size are due to elevational effects on snowfall (Bach in review). During the month of October, the beginning of the wet season, both basins have similar discharges. Between November and about April two differences become apparent in the flow characteristics. The higher elevation basin discharge becomes smaller in volume and less variable than the lower basin (Fig. 4.30). The decrease in volume is due to more precipitation falling in the form of snow and accumulating in the upper

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basin, while rain falls throughout the winter in the lower basin and runs off. The greater temporal variability (as indicated by the wiggles in the line) in the lower basin is due to the chaotic timing of storms throughout the period of record. In March, temperatures begin to warm and the snowmelt season begins, a few weeks earlier in the lower basin (Fig. 4.30). The variability of discharge decreases in both basins (lines become smoother), indicating a change from storm event dominated runoff to temperature driven snowmelt runoff (Peterson et al. 2000). The peak in the snowmelt flood of the lower basin occurs about one month earlier and is only 70% the size of the upper basin, reflecting the difference in snowpack volume. These runoff characteristics are further exasperated by the presence of glaciers in the watershed (Fountain and Tangborn 1985). Besides the daily to weekly variations caused by storm events, and the seasonal variations through out a year, streamflow regimes in mountains are prone to interannual and secular variations related to large-scale patterns of climate variations such as the El Niño/Southern Oscillation (Trenberth 1999). The type of climatic variations vary around the globe, but generally result in extreme weather and climate conditions such as flooding, droughts, different storm frequency and precipitation, or changes in snowmelt season (Karl et al. 1999). Global warming is likely to increase the frequency and magnitude of these climatic variations and their impacts on the hydrological system (Barry 1990; Trenberth 1999). High-elevation snowsheds are important to the regional water supply, as they provide water for domestic, industrial and agricultural users; recreation; hydroelectric power; habitat; and produce flood hazards. Rapid population growth, increasing environmental concerns, and resulting changes in the character of water demands have led to increased competition for water even under normal flow conditions. Water management practices, storage infrastructure, and patterns of use are tuned to the expected range of variation in surface runoff and groundwater availability (Robinson 1977). The abundant surface water supply from mountainous regions has promoted a historic reliance on this resource in adjacent lowlands. Effective water development planning and policy making must recognize how changes in upper watershed conditions will impact lowland water resources (Hulme et al. 1999). New reservoirs and water transfer systems require considerable lead time to plan and construct. Such structures will be necessary to deal with changing water supplies conditions that will exist in the future.

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Winds Mountains are among the windiest places on earth. They protrude into the high atmosphere, where there is less friction to retard air movement. There is no constant increase in wind speed with altitude, but measurements from weather balloons and aircraft show a persistent increase at least up to the tropopause where, in middle latitudes, the wind culminates in the jet streams. Similar increases occur in mountains, although the conditions at any particular site are highly variable. Wind speeds are greater in middle latitudes than in tropical or polar areas, in marine than in continental locations, in winter than in summer, during the day than at night, and, of course, the velocity of the wind is dependent on the local topographic setting and the overall synoptic conditions (Smith 1979; Gallus et al. 2000). The wind is usually greatest in mountains oriented perpendicular to the prevailing wind, on the windward rather than the leeward side, and on isolated, unobstructed peaks rather than those surrounded by other peaks. The reverse situation may exist in valleys, since those oriented perpendicular to the prevailing winds are protected while those oriented parallel to the wind may experience even greater velocities than the peaks, owing to funneling and intensification (Ramachandan et al. 1980). Table 4.7 lists the mean monthly wind speeds during the winter for several representative mountain stations in the northern hemisphere. Mountains greatly modify the normal wind patterns of the atmosphere (Smith 1979; Bossert and Cotton 1994). Their effect may be felt for many times their height in both horizontal and vertical distance. The question of whether the wind speed is greater close to mountains or in the free air has long been problematic. The two basic factors that affect wind speeds over mountains operate in opposition to on another. The vertical compression of airflow over a mountain causes acceleration of the air, while frictional effects cause a slowing. Frictional drag in the lowest layers of the atmosphere is caused by the interaction of air with individual smallscale roughness elements (i.e. vegetation, rocks, buildings or landforms < 10 m dimensions) and by influence of larger topographic features and vegetation canopies (Richard et al. 1989; Taylor et al. 1987; 1989; Walmsley et al. 1989). It is generally believed that the wind near mountains is greater, because of compression and forcing of the air around the peak like water around a rock in a stream (Schell 1935, 1936; Conrad 1939; Ryan 1977; Woodbridge et al. 1987; Taylor et al. 1989). However, studies in the Alps indicate that the wind speed on these mountain summits

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averages only about one-half that of the free air (Wahl 1966; Davies and Phillips 1985). Both of these situations may in fact occur; much depends upon the stability of the air mass and the size and configuration of the mountain. Generally, the more stable the air, the greater the compression, because the air will resist lifting in its passage, and this will result in increased wind speeds near the surface (Lee et al. 1987). On the other hand, if the air is unstable, it will tend to rise on its own accord as it is forced up over the mountain, and this will result in greater wind speeds aloft. The vertical velocity-gradient of the wind is largely a function of the interplay between compressional and frictional effects (Gallus et al. 2000). Compression tends to create greater velocities near the surface, decreasing upward, whereas friction tends to cause lower velocities near the surface, increasing upward (Carruthers and Hunt 1990). Consequently, the wind speeds in any given mountain area may have very different distributions in time and space (Schell 1936). The sharpest gradient in wind speed usually occurs immediately above the surface. Wind speed doubles or triples within the first few meters, but the vegetation and surface roughness make a great difference in the absolute velocity (Fig. 4.31). At a height of 1 m (3.3 ft.) the wind speed in a closed-forest immediately below timberline is less than half that in the open tundra just above timberline (Richard et al. 1989). The low-lying foliage of alpine vegetation does not produce much frictional drag on the wind, so the wind can reach quite high velocities close to the ground. There is nevertheless a sharp gradient within the first few centimeters of the surface, and most alpine plants escape much of the wind (Warren-Wilson 1959). A reciprocal and reinforcing effect is operative here: taller vegetation tends to reduce the wind speed and provide a less windy environment for plants, while low-lying alpine vegetation provides little braking effect, so the wind blows freely and becomes a major factor of stress in the environment. Under these conditions the presence of microhabitats becomes increasingly important. Surface roughness caused by clumps of vegetation and rocks creates turbulence and hence great variability in wind speed near the surface (Fig. 4.32). In the illustration, wind speed at a height of 1 m (3.3 ft.) above the grass tussock is 390 cm/sec., while closer to the ground it is 50 cm/sec. on the exposed side of the tussock and 10 cm/sec. on the lee side (Fig. 4.32a). Similar conditions exist with the eroded soil bank, except that wind speeds are higher on the exposed side and there is more eddy action and reverse flow to the lee. The restriction of the vegetation to

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the lee of the soil bank is largely due to the reduced wind speed there (Fig. 4.32b). The wind follows a similar pattern across the rock, with small eddies developing in depressions and to the lee (Whalsley et al. 1989). A mat of vegetation occupies the center depression where wind speeds are less (Fig. 4.32c). Wind is clearly an extreme environmental stress; in many cases it serves as the limiting factor to life. What may be the two most extreme environments in mountains are caused by the wind: late-lying snowbanks, where the growing season is extremely short, and windswept, dry ridges. Both of these environments become more common and more extreme with elevation, until eventually the only plants are mosses and lichens-or perhaps nothing at all (but see p. 292). Trees on a windswept ridge may be ―flagged‖ with the majority of branch growth on the protected lee-side (Yoshino 1973; Fig. 8.17 and 8.18). In the extreme conditions within the krummholz (crooked wood) zone, trees take on a prostate cushion form (Fig. 8.19) The redistribution of snow by the wind is a major feature of the alpine environment. The wind speed necessary to pick snow up from the surface and transport it, depends upon the state of the snow cover, including temperature, size, shape and density of the snow particles and the degree of intergranular bonding (Tabler 1975). For loose, unbound snow the typical velocity is about 5 m/s, while a dense bonded snow cover requires velocities in excess of 25 m/s. Blowing snow can abrade surfaces, causing erosion to snow cover and flagging trees. Once the wind velocity lowers, the snow is deposited into dune-like features called drifts. Drifts are found in the lee-eddy of obstacles of all sizes (e.g trees, ridges, fences; Fig. 4.27). To control blowing snow, snow fences and other barriers are specially engineered to maximize deposition, and carefully placed to reduce the hazard of blowing snow or drifts (Ring 1991). In mountains, snow redistribution by wind is strongly affected by meso- and micro-scale topography and vegetation (Föhn 1980; Meister 1987; Tesche 1988). Topographic traps fill in with snow, where it may survive late into spring or summer due to its depth and temperature inversions. The deposition of drifts is an important component to alpine water storage and spring runoff (Elder et al. 1989). Many glaciers have a significant component of their accumulation from snow blown over crests (Föhn 1980; Pelto, 1996). There are two overall groups or types of winds associated with mountains. One type originates within the mountains themselves. These are local, thermally induced winds given

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distinct expression by the topography. The other type is caused by obstruction and modification of winds originating from outside the mountain area. The first type is a relatively predictable, daily phenomenon, while the second is more variable, depending as it does on the vagaries of changing regional wind and pressure patterns.

Local Wind Systems in Mountains Winds that blow upslope and upvalley during the day and downslope and downvalley at night are common. Albrecht von Haller, author of Die Alpen, observed and described these during his stay in the Rhône Valley of Switzerland from 1758 to 1764. Since then many studies (summarized by Defant 1951; Geiger 1965; and Rohn 1969b; Barry 1992) have been made on thermally induced winds. The driving force for these winds is differential heating and cooling which produces air density differences between slopes and valleys and between mountains and adjacent lowlands (McGowan and Sturman 1996a). During the day, slopes are warmed more than the air at the same elevation in the center of the valley; the warm air, being less dense, moves upward along the slopes. Similarly, mountain valleys are warmed more than the air at the same elevation over adjacent lowlands, so the air begins to move up the valley. These are the same processes that give rise to convection clouds over mountains during the day and provide good soaring for glider pilots (and birds). At night, when the air cools and becomes dense, it moves downslope and downvalley under the influence of gravity. This is the flow responsible for the development of temperature inversions. Although they are interconnected and part of the same system, a distinction is generally made between slope winds, and larger mountain and valley wind systems (Fig. 4.33).

Slope Winds. Slope winds consist of thin layers of air, usually less than 100 m (330 ft.) thick. In general, the upslope movement of warm air during the day is termed anabatic flow, and the downslope movement of cold air during the night is referred to as katabatic flow, or a gravity or drainage wind. The upslope flow of air during the day is associated with surface heating and the resulting buoyancy of the warm air (Vergeiner and Dreiseitl 1987). The wind typically begins to blow uphill about one-half hour after sunrise and reaches its greatest intensity shortly after noon (Fig. 4.33a). By late afternoon the wind abates and within a half hour after sunset reverses to

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blow downslope (Fig. 4.33c). Katabatic winds in the strict sense are local downslope gravity flows caused by nocturnal radiative cooling near the surface under calm, clear-sky conditions, or by the cooling of air over a cold surface such as a lake or glacier. The extra weight of the stable layer, relative to the ambient air at the same altitude, provides the mechanism for the flow. Since slope winds are entirely thermally induced, they are better developed in clear weather than in clouds, on sun-exposed rather than on shaded slopes, and in the absence of overwhelming synoptic winds. Local topography is important in directing these winds; greater wind speeds will generally be experienced in ravines and gullies than on broad slope (Defant 1951; Banta and Cotton 1981; McKee and O‘Neal 1989). Downslope winds form better at night and during the winter, when radiative cooling dominates the surface energy system (Horst and Doran, Barr and Orgill 1989). The down slope flow of cold air is analogous to that of water, since it follows the path of least resistance and always gravitates toward equilibrium, but water has a density 800 times greater than air (Bergen 1969). Even with a temperature difference of 10˚C (18˚F), the density of cold air is only 4% greater than warm air, unlike the rapid flow of water due to gravity, the displacement of warm air by cold air is a relatively slow process (Geiger 1969). Katabatic winds begin periodically as the layer of air just above the surface cools, then slides downslope (Papadopoulos and Helmis 1999). The cycle is repeated when the radiative cooling rebuilds the downslope pressure gradient. This pulsating downslope flow depends on the temperature difference between the katabatic layer and the valley temperature (McNider 1982). Surges of cold air are commonly observed on slopes greater than 10°, and are referred to as ―air avalanches‖ (Scaetta 1935; Geiger 1969). A final steady velocity will be achieved once a certain temperature has been reached (Papadopoulos and Helmis 1999). Further down a drainage basin a steady velocity will be reached, maintained, and increased through out the night as individual slope winds accumulate down basin in a similar fashion to tributaries in a stream (Neff and King 1989; Porch et al. 1989). Closed basins, even created by dense forest cover, can trap the cold air creating cold pockets or ―frost hollows‖ (Thompson 1986; Neff and King 1989). These temperature inversions can reach 30°C below the ambient atmosphere, persisting for weeks to months, effectively trapping atmospheric contaminants until sufficient winds can clear the air (Geiger 1965; McGowan and Sturman 1993; Iijima and Shinoda 2000). These have obvious significance for a number of human activities,

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such as agriculture, forestry, tourism and air pollution. In the wine-producing regions of Germany, hedges are frequently planted above the vineyards to deflect cold air from upslope (Geiger 1969). Upslope winds form best during the day and during the summer when surfaces are radiatively warmed (Banta 1984; 1986). The upslope wind does not rise far above the ridge tops since it is absorbed and overruled by the regional prevailing wind. The upward movement of two slope winds establishes a small convection system in which a return flow from aloft descends in the center of the valley (Figs. 4.32, 4.33; cf. Fig. 4.33a). This descending flow brings from aloft drier air that has been heated slightly by compression and thus is strongly opposed to cloud formation. For this reason the dissipation of low-lying fog and clouds generally takes place first in the center of the valley (Fig. 4.34). If the valley is deep enough, the dry descending air can produce markedly arid zones. In the dry gorges and deep valleys of the Andes of Bolivia and in the Himalayas, the vegetation ranges from semi-desert shrubs in valley bottoms to lush forests on the upper slopes where clouds form (Troll 1952, 1968; Schweinfurth 1972).

Mountain and Valley Winds. The integrated effects of slope generated flows produce mountain and valley winds, blowing longitudinally up and down the main valleys, essentially at right angles to the slope winds (Whiteman 1990; Clements 1999). They are all part of the same system, however, and are controlled by similar thermal responses. The valley wind (blowing from the valley toward the mountain) is interlocked with the upslope winds, and both begin after sunrise (Buettner and Thyer 1962, 1965; Banta 1984; 1986; Fig. 4.33b). Valley winds involve greater thermal contrast and a larger air mass than slope winds, however, so they attain higher wind speeds. In the wide and deep valleys of the Alps, the smooth surfaces left by glaciation allow maximum development of the wind. The Rhône Valley has many areas where the trees are wind-shaped and flagged in the upvalley direction (Yoshino 1964b). Mountain winds (blowing from the mountains down valley) are associated with the nocturnal downslope winds and can be very strong and quite cold in the winter (Porch et al. 1989; Whiteman 1990; Fig, 4.31d). As with slope winds, a circulation system is established in mountain and valley winds. The return flow from aloft (called an anti-wind) can frequently be found immediately above the valley wind (Bleeker and Andre 1951; Defant 1951). This concept was formerly only theoretical,

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but study of valley winds near Mount Rainier, Washington, using weather balloons, clearly identified the presence of anti-winds (Fig. 4.35; Buettner and Thyer 1965). This wind system beautifully demonstrates the three-dimensional aspects of mountain climatology: next to the surface are the slope and mountain-valley winds; above them is the return flow or anti-wind; and above this is the prevailing regional gradient wind (McGowan and Sturman 1996a; Clements 1999). During clear weather all of these may be in operation at the same time, each moving in a different direction.

Other Local Mountain Winds. An important variant of the thermal slope wind is the glacier wind, which arises as the air adjacent to the icy surface is cooled and moves downslope due to gravity. The glacier wind has no diurnal period but blows continuously, since the refrigeration source is always present. It reaches its greatest depth and intensity at mid-afternoon, however, when the thermal contrast is greatest. At these times the cold air may rush downslope like a torrent. During the day the glacier wind frequently collides with the valley wind and slides under it (Fig. 4.36). At night it merges with the mountain wind that blows in the same direction (Defant 1951). In mountains like the Rockies or Alps, with small valley glaciers, the glacier winds are fairly shallow, but when glaciers are as extensive as they are in the St. Elias Mountains or the Alaska Range, the wind may be several hundred meters in depth (Marcus 1974a). Glacier winds have a strong ecological effect, since the frigid temperatures are transported downslope with authority and the combined effect of wind and low temperatures can make the area they dominate quite inhospitable. In a valley with a receding glacier, these winds can entrain the unconsolidated till, sand-blasting vegetation and rocks (into ventifacts) down valley (Bach 1995). Another famous local wind in mountains is the Maloja wind, named after the Maloja Pass in Switzerland between the Engadine and Bergell valleys (Hann 1903; Defant 1951; Whipperman 1984). This wind blows downvalley both day and night and results from the mountain wind of one valley reaching over a low pass into another valley, where it overcomes and reverses the normal upvalley windflow. This anomaly occurs in the valley with the greater temperature-gradient and the ability to extend its circulation into the neighboring valley across the pass. Thus, the wind ascends from the steep Bergell Valley and extends across the Maloja

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Pass downward into the Engadine Valley to St. Moritz and beyond. A similar situation exists in the Davos Valley, Switzerland (Flohn 1969b). A related phenomenon occurs in coastal areas where a strong sea-breeze moves inland and over low passes in such a way that the wind blows down the lee mountain slope during the day. This is well-developed on the asymmetric escarpments of the Western Ghats in India. In the equatorial Andes, cool air from the Pacific moves inland in a shallow surface layer overflowing the lower passes into the valleys beyond, producing relatively cool flows down the east side of the range (Lopez and Howell 1967). In some cases these winds are forced up the opposite slopes in a hydraulic-jump phenomenon (Gaylord and Dawson 1987), producing afternoon rainfall (Fig. 4.37). Other examples of local winds could be given, since every mountainous country has its own peculiarities, but those mentioned suffice to illustrate their general nature.

Mountain Winds Caused by Barrier Effects As mentioned earlier in the chapter, mountains can act as barriers to the prevailing general circulation of the atmosphere. The barrier effect introduces turbulence to the winds, increasing and decreasing speeds, changes directions, and modifies storms (reviewed earlier). Once the wind passes over the mountain crest however, it will do one of two things: flow down the lee-side or stay lifted in the atmosphere (Durran 1990). Most commonly the wind will fall down the lee-side of mountains under the influence of gravity. These surface winds are sometimes collectively termed fall winds, but are known by a variety of local terms because they have long been observed in many regions downwind from mountains, and have associated with them distinct weather phenomenon. When the winds leave the surface in a hydraulic jump, they often travel through the atmosphere in a wave motion, producing unique cloud-forms.

Foehn Wind. Of all the transitory climatic phenomena of mountains the foehn wind (pronounced "fern" and sometimes spelled föhn) is the most intriguing. Many legends, folklore, and misconceptions have arisen about this warm, dry wind that descends with great suddenness from mountains. The foehn, known in the Alps for centuries, is a feature common to all major mountain regions formed as synoptic winds blow over a mountain crest and down the lee-side (Barry 1992). In North America it is called the "Chinook;" in the Argentine Andes the "zonda;"

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in New Zealand the "Canterbury north-wester;" in New Guinea the ―warm braw;‖ in Japan the ―yamo oroshi;‖ in the Barison Mountains of Sumatra the ―bohorok;‖ the ―halny wigtr‖ in Poland; the ―autru‖ in Romania; other mountain regions have their own local names for it (Brinkman 1971; Forrester 1982). The ―Santa Ana‖ of southern California forms in a similar fashion (Kasper 1981). The foehn produces distinctive weather: gusts of wind, high temperatures, low humidity, and very transparent and limpid air (Brinkman 1971; Pettre 1982). When viewed through the foehn, mountains frequently take on a deep blue or violet tinge and seem unnaturally close and high, because light rays are refracted upward through layers of cold and warm air. The bank of clouds that typically forms along the crest line is associated with the precipitation failing on the windward side. This bank of clouds remains stationary in spite of strong winds and is known as the foehn wall (when viewed from the lee side). The following is an early naturalist's description of the foehn in Switzerland: In the distance is heard the rustling of the forests on the mountains. The roar of the mountain torrents, which are filled with an unusual amount of water from the melting snow, is heard afar through the peaceful night. A restless activity seems to be developing everywhere, and to be coming nearer and nearer. A few brief gusts announce the arrival of the foehn. These gusts are cold and raw at first, especially in winter, when the wind has crossed vast fields of snow. Then there is a sudden calm, and all at once the hot blast of the foehn bursts into the valley with tremendous violence, often attaining the velocity of a gale which lasts two or three days with more or less intensity, bringing confusion everywhere; snapping off trees, loosening masses of rock; filling up the mountain torrents; unroofing houses and barns- a terror in the land. (Quoted in Hann 1903, p. 346)

The primary characteristics of the foehn are a rapid rise in temperature, gustiness, and an extreme dryness that puts stress on plants and animals and creates a fire hazard (Ives 1950; Brinkman 1971; Marcus et al. 1985; Spronken-Smith 1998). Forests, houses, and entire towns have been destroyed during foehn winds clock up to 195 km/h (Reid and Turner 1997). For this reason, smoking and fires (even for cooking) have traditionally been forbidden in many villages in the

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Alps during the foehn. In some cases special guards (Föhnwächter) were appointed to enforce the regulations. In New Zealand foehn winds commonly entrain glacial sediments causing dust storms and degrading grasslands (McGowan and Sturman 1996b). The foehn is purported to cause various psychological and physiological reactions, including a feeling of depression, tenseness, and irritability, and muscular convulsions, heart palpitations, and headaches. The suicide rate is said to rise during the foehn (Berg 1950). These symptoms have rarely been observed in North America, and medical explanations remain elusive. In spite of its disadvantages, the foehn is generally viewed with favor, since it provides respite from the winter's cold and is very effective at melting snow (Ashwell and Marsh 1967), a fact reflected in many local sayings from the Alps: "if the foehn did not interfere, neither God nor his sunshine would ever be able to melt the winter snows"; "The foehn can achieve more in two days than the sun in ten"; "The wolf is going to eat the snow tonight" (De La Rue 1955, pp. 36-44). In North America, the value of the chinook for the Great Plains was poignantly illustrated by the painting "Waiting for a Chinook," by Charles M. Russell. During the winter of 1886, cattle and sheep died by the thousands in Montana in one of the worst snowstorms on record. Russell, a cowboy on a large ranch there, received a letter from his alarmed employers in the East, asking about the condition of their stock. Instead of writing a reply, he made a watercolor sketch of a nearly starved steer standing in deep snow unable to find food, with coyotes waiting nearby (Fig. 4.38). The picture soon became famous and so did Russell; "Waiting for a Chinook" remains his best-known painting (Weatherwise 1961). The causes of the foehn are complex. One of the early explanations in the Alps was that the warm dry wind came from the Sahara Desert. The wind was usually from the south, so this seemed a perfectly logical solution, until one day somebody climbed to the side of the mountain from which the foehn was coming and found that it was raining there, a very unlikely effect for a Saharan wind to produce! The Austrian climatologist Julius Hann (1866, 1903) is given credit for the true explanation. When air is forced up a mountain slope, it is cooled at the dry adiabatic rate, 3.05˚C per 300 m (5.5˚F per 1,000 ft.) until the dew point is reached and condensation begins. From this point on, the air is cooled at a lower rate (wet adiabatic rate) of approximately 1.7˚C (3˚F) (Fig. 4.39). On the lee side of the summit, precipitation ceases and the air begins to descend. Under these conditions, the air is warmed at the dry adiabatic rate, 3.05˚C per 300 m

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(5.5˚F per 1,000 ft.) the entire length of its descent. Consequently, the air has the potential for arriving at the valley floor on the leeward side warmer than its original temperature at the same elevation on the windward side (Fig. 4.39). While this general model has been widely demonstrated, many foehn winds involve site specific processes (Barry 1992). The foehn develops only under specific pressure conditions (Hoinka 1985; McGowan and Sturman 1996b). The typical situation is a ridge of high pressure on the windward side and a trough of low pressure on the leeward, creating a steep pressure-gradient across the mountain range. Under these conditions the air may undergo the thermodynamic process just described in a relatively short time. In order for there to be a true foehn, however, the wind must be absolutely warmer than the air it replaces (Brinkman 1971). Either side of the mountains may experience a foehn, depending upon the orientation of the range and the development of the pressure systems. In the Alps there is a south foehn that affects the north side of the mountains and comes from the Mediterranean, and a north foehn that comes from northern Europe and affects the south side of the Alps. Because of its original warmth, the south foehn is much more striking and more frequent than the north foehn, which has to undergo much greater warming to make itself felt (Defant 1951). Similarly, in western North America most chinook winds occur on the east side of the mountains, because of the prevailing westerly wind and its movement over the Pacific Ocean, which is considerably warmer in winter than the continental polar air characteristic of the Great Basin and High Plains. Chinooks do occur, although less frequently, on the western side of the mountains (Ives 1950; Cook and Topil 1952; McClain 1952; Glenn 1961; Longley 1966,1967; Ashwell 1971; Riehl 1974; Bower and Durran 1986).

Bora, Mistral, and Similar Winds. Like the foehn, these winds descend from mountains onto adjacent valleys and plains but, unlike the foehn, they are cold. Compressional heating occurs, but it is insufficient to appreciably warm the cold air that blows from an interior region in winter across the mountains to an area that is normally warmer. These winds and others like them are basically caused by the exchange of unlike air across a mountain barrier. The ―bora‖ is a cold, dry north wind on the Adriatic coast of Dalmatia. It reaches its most intense development in winter and originates from high-pressure, cold continental air in southwestern Russia that results in air movement southward across Hungary and the Dinaric Alps

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(Smith 1987; Durran 1990). Ideal conditions for the bora exist when a southerly wind has brought exceptionally warm conditions to the Adriatic coast, and relatively large temperatureand pressure-differentials exist between the coast and the interior. Under these conditions, the cold continental air may move down the pressure gradient, steepened by the presence of the mountains, with extraordinary violence (Yoshino 1975; Pettre 1982). It frequently reaches gale force, especially when channeled through narrow valleys and passes. The bora has been known to overturn haywagons, tear off roofs, and destroy orchards. It is even claimed that it once overturned a train near the town of Klis (De La Rue 1955). The ―mistral‖ occurs in Provence and on the French Mediterranean coast (Jansa 1987). It is caused by the movement of cold air from high-pressure areas in the north and west of France toward low-pressure areas over the Mediterranean between Spain and Italy in the Gulf of Lyon. It is as violent as the bora, or more so, since it must pass through the natural constriction between the Pyrenees and the western Alps (Defant 1951). The mistral was known in ancient times. The Greek geographer Strabo called it "an impetuous and terrible wind which displaces rocks, and hurls men from their chariots" (De La Rue 1955, p. 32). Its effects extend throughout Provence and may be felt as far south as Nice. Like the bora, the mistral poses a major problem for fruit production, and great expenditures of human labor have gone into constructing stone walls and other windbreaks to protect the orchards (Gade 1978). The greatest wind velocities occur in the Rhône Valley, where wind speeds of over 145 km (90 mi.) per hour have been recorded. Although the bora and mistral are the most famous, similar cold dry winds occur in many mountain areas (Forrester 1982). The ―bise‖ (breeze) at Lake Geneva between the Alps and the French ―Jura‖ is of the same type, and numerous examples could be cited from the large mountain gaps and passes of Asia (Flohn 1969b). ―Helm‖ winds blow down from the Pennine Chain in north-central England, often creating rolls of clouds. ―Sno‖ winds fill the fjords of Scandinavia during winter and the ―oroshi‖ blows near Tokoyo (Yoshino 1975). In North America, the "northers" of the Gulf of Tehuantepec, Mexico, are a similar phenomenon (Hurd 1929). Another example is the exchange of the cold air during winter between the east and west sides of the Cascade Mountains along the Columbia River Gorge (recently term the ―Coho‖ wind) and the Fraser River Valley. This is most pronounced when an outbreak of cold Arctic air moves southward and banks up against the east side of the Cascades, causing great temperature-

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and pressure-contrasts between the cold continental air and the relatively warm Pacific air. At these times cold air is forced through these sea-level valleys at high velocities and brings some of the clearest and coldest weather of the winter to the cities of Vancouver, British Columbia, and Portland, Oregon (both located at valley mouths). At other times the cold air may force the warm coastal air aloft and produce locally heavy snowfall. The most dominant feature, however, is the cold and ferocious wind that leaves its mark on the landscape in the brown and deadened foliage of needle-leaf trees and in the strongly flagged and wind-shaped trees on exposed sites (Lawrence 1938).

Lee Waves. The behavior of airflow over an obstacle depends largely on the vertical wind profile, the stability structure, the shape of the obstacle and the surface roughness (Stull 1988; Barry 1992; Romero et al 1995). When wind passes over an obstacle, its normal flow is disrupted and a train of waves may be created that extends downwind for considerable distances (Figs. 4.26 and 4.39). The major mountain ranges produce large-amplitude waves that extend around the globe (Hess and Wagner 1948; Gambo 1956; Nicholls 1973; Vosper and Parker 2002). On a smaller scale, these waves take on a regional significance reflected in their relationship to the foehn, in distinctive cloud forms, in upper-air turbulence and downwind climate (Scorer 1961, 1967; Reiter and Foltz 1967; Wooldridge and Ellis 1975; Smith 1976; Durran 1990; Reynolds 1996; Reinking et al 2000). The amplitude and spacing of lee waves depends on the wind speed and the shape and height of the mountains, among other factors. An average wavelength is between 2-40 km (1-25 mi.), the vertical amplitude is usually between 1-5 km (0.6-3 mi.), and occurs at altitudes of 300-7,600 m (1,000-25,000 ft.) (Hess and Wagner 1948; Durran 1990). Wind speeds within lee waves are quite strong, frequently exceeding 160 km (100 mi.) per hour (Scorer 1961). The most distinctive visible features of lee waves are the lenticular (lens-shaped) or lee-wave clouds that form at the crests of waves (Fig. 4.41). These are created when the air reaches dew point and condensation occurs as the air moves upward in the wave (Ludlam 1980). The clouds do not form in the troughs of the waves, since the air is descending and warming slightly (Fig. 4.40). The relatively flat cloud-bottoms represent the level of condensation, and the smoothly curved top follows the outline of the wave crest. The clouds are restricted in vertical extent by overlying stable air (Vosper and Parker 2002). Lenticular clouds are relatively

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stationary (hence the name "standing-wave clouds"), although the wind may be passing through them at high speeds. Lee-wave clouds frequently develop above one another, as well as in horizontal rows (Fig. 4.41; cf. Fig. 4.40). They typically consist of 1-5 clouds and extend only a few kilometers downwind, but satellite photography has revealed series of 30-40 clouds extending for several hundred kilometers (Fig. 4.43; Fritz 1965; Bader et al. 1995). Much of the early knowledge about lee waves was acquired by glider pilots who found to their surprise that there was often greater lift to the lee of a hill than on the windward side. The pilots had long made use of upslope and valley winds, but by this method could never achieve a height of more than a couple of hundred meters above the ridges. In southern England, the members of the London Gliding Club had soared for years in the lift of a modest 70 m (230 ft.) hill, never achieving more than 240 m (800 ft.). After discovering the up-currents in the lee wave, however, one member soared to a height of 900 m (3,000 ft.), thirteen times higher than the hill producing the wave (Scorer 1961). German pilots were the first to explore and exploit lee waves fully. In 1940, one pilot soared to 11,300 m (37,400 ft.) in the lee of the Alps. The world's altitude record of 13,410 m (44,255 ft.) was set in 1952 in the lee of the Sierra Nevada of California. This range has one of the most powerful lee waves in the world, owing to its great altitudinal rise and the clean shape of its east front (Scorer 1961). Another aspect of lee waves is the development of rotors. These are awesome roll-like circulations that develop to the immediate lee of mountains, usually forming beneath the wave crests (Fig. 4.40). The rotor flow moves toward the mountain at the base and away from it at the top (Tampieri 1987). It is marked by a row of cumulus clouds but, unlike ordinary cumulus, they may contain updrafts of 95 km (60 mi.) per hour (Fig. 4.44). The potential of such a wind for damage to an airplane can well be imagined. The height of the rotor clouds is about the same as that of the crest cloud or foehn wall. The rotating motion is thought to be created when the lee waves reach certain amplitude and frictional drag causes a roll-like motion in the underlying air (Fig. 4.40; Scorer 1961, 1967). Several other kinds of turbulence may be associated with lee waves, particularly when the wave train produced by one mountain is augmented by that of another situated in the right phase relationship (Fig. 4.45). In some cases they cancel each other; in others they reinforce each other. Wind strength and direction are also important, since a small change in either one can alter the

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wave length of two superposed wave trains so that they become additive and create violent turbulence (Scorer 1967; Lilly 1971; Lester and Fingerhut 1974; Neiman et al 2001). One example of this is where the energy in standing waves is caused to "cascade" down from a wavelength of 10 km (6 mi.) to only a few hundred meters (Reiter and Foltz 1967).

Microclimates In addition to the climatic characteristics reviewed above, it should be emphasized there are substantial variations in climates over very short distances within mountains. Mountain environments are exceedingly spatially complex in terms of vegetation types and structures, geology, soils, and topography. All vary in composition (i.e. species, canopy characteristics or rocktypes), and variations occur across a range of slopes and aspects. The climate over each of these surfaces, or microclimate, can differ significantly due to the variations net radiation, soil and air temperature; humidity, precipitation accumulation (amount and form) and soil moisture; and winds (Barry and Van Wie 1974; Green and Harding 1980; Fitzharris 1989). Large differences in temperature, moisture and wind can be found within a few meters, or even centimeters (Turner 1980; McCutchan and Fox 1986). The thin atmosphere at high elevation means surfaces facing the sun on a clear day can warm dramatically, but shaded surfaces remain cold (Fig. 4.23; Germino and Smith 2000). Other effects may arise according to valley orientation with respect to the mountain range, valley cross-profile, and the affect of winds and cold air drainage. The effect of aspect in generating slope winds can exceed the influence of elevation on wind velocity and temperature (McCutchan and Fox 1986). The mosaic of microclimates determines the local variability in ecosystem processes. The distribution of vegetation zones, and even individual species may follow the distribution of microclimates (Fig. 4.8; Canters, et al 1991; Roberts and Gilliam 1995; Parmesan 1996). A simple classification using solar receipt, wind exposure, depth of winter snow cover and density and height of vegetation cover, can help to characterize alpine microclimates (Turner 1980). On this basis the following general microclimates can be differentiated: Sunny, windward slope – solar radiation and windspeeds high Sunny, lee slope – solar radiation high, windspeeds low Shaded, windward slope – solar radiation low, windspeeds high

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Shaded, lee slope – solar radiation and windspeeds low. Certainly there are gradients between these categories. Precipitation and runoff inputs with alter the soil moisture regime of each site, but generally the list goes from dry to moist. Vegetation creates its own microenvironment by creating shade and windbreaks (Fig. 4.32). In association with wind regime is a recurring pattern of snow accumulation in the lee of obstacles. These snowdrifts add to soil moisture during the melt season, and protect trees from freezing in winter (Wardle 1974). The resolution of most weather station networks in mountains is far too coarse to capture the spatial variability of climates in mountains. Maps of climatic variables are often interpolated from existing meager data sets, using assumed or empirical relationships with elevation (Peck and Brown 1962; Kyriakidis et al. 2001). These models are unable to demonstrate local deviations in trends, and when combined with map scale, microclimates are typically eliminated from most maps of mountains. Likewise, vegetation maps of mountainous areas rarely show the small patches of vegetation that occurs in microclimatic habitats. While these features may be un-mappable, they are certainly observable in mountains, adding to the spender of multifaceted mountain environment.

Climate Change and Variability Variability of climatic phenomenon is an important natural component of earth's climate system. Climatic variability (occurrence of certain climatic events) is different than climatic change, which is a permanent change in climatic conditions. However, changes in variability are a likely result of climatic change. The middle and high latitudes inherently have very variable climates since they are influenced by large seasonal changes in energy. The equatorial region experiences little variability, as it has nearly the same energy fluxes year round. Reflecting the complexity of the climate system, most regions of the world show different patterns and magnitudes of variability and trends through time (Karl et al. 1999). All temporal climate records demonstrate some degree of interannual variability (e.g. Karl et al. 1999; Liu and Chen 2000; Kane 2000; Peterson and Peterson 2001). Every mountain location has its record high and low temperature, snowfall, rain event, drought, and wind speed. While these extreme events are rare, they often occur with a greater frequency, and with more

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extreme magnitudes in mountainous regions, than in lowlands (Frei and Schar 2001). Extreme storm events are exasperated by the topographic setting of mountains, producing even higher precipitation totals, lower temperatures, and higher wind velocities. Extreme precipitation events in mountains are of significance because they lead to hazards, such as downstream flooding, soil erosion and mass-movements on slopes (Ives and Messerli 1989; Rebetez et al. 1997). Temperature, precipitation and the resulting runoff variations are often related to distant forcing mechanisms such as the El Niño/Southern Oscillation (Dettinger and Cayan 1995; Cayan et al. 1998; 1999; Diaz et al. 2001; Clare et al. 2002; Rowe et al. 2002). Several other periodic, yet chaotic perturbations to the climate system have been linked to increased climatic variability (McCabe and Fountain 1995; Mantua et al. 1997; Fowler and Kilsby 2002). Among the regional differences in variability, the following consistent temporal trends emerge in data sets over the last century: the number of extremely warm summer temperatures has increased a small amount, the number of extremely cold winter temperatures has clearly decreased (with fewer frost days), and mean summer season precipitation has increased, especially an increase in heavy precipitation events (Karl et al. 1999). All of these general trends have temporally reversed during the period of record. So while variability is expected, changes in frequency of occurrence of extreme events is recognized as a signal of ongoing climate change (NAST 2001). Climate changes are well-documented to have occurred in the geologic past, as illustrated by the glacial and inter-glacial climates of the Pleistocene (COHMAP 1988; Petit et al. 1999). General scientific consensus states that the climate is currently changes, namely warming due to anthropogenic inputs of greenhouse gases to the atmosphere (IPCC 2001; NAST 2001). Different magnitudes of warming, and even cooling, are predicted for different mountainous regions of the world (IPCC 2001). Precipitation, in particular, is predicted to both increase and decrease in different regions due to changes in general circulation (Schroeder and McGuirk 1998). Climate models in mountainous regions, however, tend to be rather poor, due to coarse resolution, topographic smoothing and local effects not captured by the models (Brazil and Marcus 1991; Sinclair 1993). In mountains, higher temperatures would cause both a higher percentage of annual precipitation to fall as rain (i.e. higher snowlines), as well as accelerate summer ablation (Barry 1990; Groisman etal. 1999). Characterizing the exact climatic impacts to any mountain

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site is difficult. We can however, demonstrate that past, and likely future climatic changes and variations are likely to have major impacts in mountain environments. Mountain and glacier environments are especially sensitive to climate changes and variability (Barry 1990; Willis and Bonvin 1995). Many climate changes have been detected in mountain records (e.g. Shrestha et al. 1999; Cayan et al. 2001; Pepin and Loaleben 2002). Changes in winter precipitation and summer temperatures will alter the rate and extent to which snowlines migrate up and down slope and contribute to glacier mass-balance and runoff (Rebetez 1995; Clare et al. 2002). Seasonal snow packs in the Northern Hemisphere have significantly declined over recent years (Cayan 1996; Robinson and Frei 2000). Glaciers are likely to experience negative mass, which will contribute more water to melt-season runoff and cause the glacier to retreat. Glacier recession will have an impact on local climatic conditions, such and energy and moisture exchanges and the generation of local winds. Measurements of alpine glacier mass-balances globally have documented retreats in recent decades (Haeberli et al. 1989; Marsten et al. 1989; Harper 1993; Bedford and Barry, 1995; Chambers 1997; Cogley and Adams 1998; McGabe and Fountain, 1995; Pelto 1996; Rabus and Echelmeyer 1998; McCabe et al. 2000). As a result, downstream runoff characteristics (i.e. seasonality and magnitude) may change appreciably over the next several decades. If glaciers entirely disappear from mountains, than melt-season, especially late melt-season discharge will decrease substantially (Fig. 4.30). Glaciers are estimated to provide 6-20% of annual runoff in some rivers (Aizen et al. 1995; Bach in review). Even climate changes in non-glacierized, low mountains can have a significant impact on municipal water supplies (Frie et al. 2002). Land-use changes in mountains, especially urbanization, logging and hydrolake development can have significant impacts upon the regional and microclimates in mountains (Goulter 1990; Roberts and Gilliam 1995; McGowan and Sturman 1996a). These environmental disturbances can have long-term influences on climates since they change the surface characteristics and energy and moisture fluxes. Hydrolakes have been found to moderate temperatures, increase atmospheric water vapor content and precipitation, and increase windiness by decreasing surface roughness and developing their own wind systems (Goulter 1990; McGowan and Sturman 1996a).

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Since many organisms living in mountains survive near their tolerance range for climatic conditions, even minor climatic changes could have a significant impact on alpine ecosystems (Grabher et al. 1994; Graumlich 1994; Parmesan 1996; Peterson 1998; Gottfried et al. 1998; Mizuno 1998; NAST 2001). Vegetation zones will migrate altitudinally in response to warming temperatures, possibly eliminating some biomes, although the adaptations will likely be more complex (Rochefort et al. 1994; Neilson and Drapek 1998). Treelines in many mountainous regions have been responding to recent temperature changes (MacDonald et al. 1998; Kullman and Kjallgren 2000; Marlow et al. 2000; Pallatt et al. 2000; Peterson and Peterson 2001; Klasner and Fagre 2002). Trees are invading into meadows (Rochefort et al. 1994; Gavin and Brubaker, 1999; Wearne and Morgan 2001). Complex topography will result in habitat fragmentation and the creation of barriers to migration, making it difficult for some species to adapt and allowing others, often evasive species, to expand their range. There is a chance that Quaking Aspen and Engleman Spruce of the North American western mountains might not survive under projected climate changes (Hansen et al. 2001). In response to the habitat changes, wildlife also migrates to find appropriate climatic niches (Happold 1998; Hansen et al. 2001; Wang et al. 2002). Because of microclimatic complexity, populations or individuals could readily be insolated on individual slopes or peaks, as the mountain environment increases in fragmentation (Fig. 4.8; Neilson and Drapek 1998). Since this climate shift is occurring rapidly, some species may not be able to adapt or migrate quickly enough (Grabher et al. 1994). It is probable that some alpine and cold-water fish species will not survive climatic changes, and new water temperatures will allow for the invasion of nonnative fish species (Grimm et al. 1997). Pacific salmon, which migrate to and spawn in some mountains, have experienced population fluctuations related to climate (Mantua et al. 1997; Downton and Miller 1998). In the Columbia River systems, the projected impacts of global warming are warmer water temperatures and earlier snowmelt peak flows, which are likely to further impact the beleaguered salmon population and related ecosystems (Miller 2000)

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Cut from bibliography Asp, 1956 Bolin 1950 Hales 1933 Hough 1945 Lawrence 1939 Lynott 1966 Gallimore and Lettau, 1970 Gutman and Schwerdtfeger 1965 Miller 1977 Montieth 1965 Suzuki 1965 Tanner and Fuchs 1968 A. H. Thompson 1967 Add to Biblio Aizen, V.B. E.M. Aizen, and J.M. Melack 1995. Climate, snow cover, glaciers, and runoff in the Tien Shan, Central Asia, Water Resources Bulletin, 31: 1113-1129. Alford, D. (1985) Mountain hydrologic systems, Mountain Research and Development, 5: 349363. Alpert, P. 1986. Mesoscale indexing of the distribution of orographic precipitation over high mountains, Journal of Applied Meteorology, 25: 532-545. Anderson, R.S. 1998. Near-surface thermal profiles in alpine bedrock: implications for frost weathering of rock, Arctic and Alpine Research, 30: 362-372. Andersson, T. 1980. Bergeron and the oreigenic (orogrpahic) maxima of precipitation, Pure and Applied Geophysics, 119: 558-576. Arnold, K. and R. Palmer 1952. The Coming of the Saucers, Amherst, WI: Palmer Publications. Bach, A.J. (1995) Aeolian Modifications of glacial moraines at Bishop Creek, eastern California. In Desert Aeolian Geomorphology (V.P. Tchakerian, Ed.), Chapman and Hall: London, pp. 179197. Bach, A.J. (in review) Estimating high elevation snowshed contributions to the Nooksack River watershed, North Cascades, Washington. The Geographical Review. Bader, M.J., Forbes, J.R., Grant, J.R., Lilley, R.B. and A.J. Waters 1995. Images in weather forecasting: practical guide for interpreting satellite and radar data. Cambridge : University Press.

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Bailey, W.G., IR. Saunders, and J.D. Bowers 1990. Atmosphere and surface control on evaporation from alpine tundra in the Canadian cordillera, In: H. Lang and A. Musy (eds.), Hydrology of Mountainous Regions, International Association of Hydrologic Sciences Publication no. 193, pp. 59-64. Bailey, W.G., E.J. Weick, and J.D. Bowers 1989. The radiation balance of alpine tundra, Plateau Mountains, Alberta, Canada, Arctic and Alpine Research, 21: 126-134. Banta, R.M. 1984. Daytime boundary-layer evolution over mountainous terrain. Part 1. Observations of the dry conditions, Monthly Weather Review, 112: 340-356. Banta, R.M. 1986. Daytime boundary-layer evolution over mountainous terrain. Part 2. Numerical studies of upslope flow duration, Monthly Weather Review, 114: 112-130. Banta, R. and W.R. Cotton 1981. An analysis of the structure of local wind systems in a broad mountain basin, Journal of Applied Meteorology, 20: 1255-1266 Banta, R.M. and C.L.B. Schaaf 1987. Thunderstorm genesis zones in the Colorado Rocky Mountains as determined by traceback of geosynchronous satellite images, Monthly Weather Review, 115: 463-476. Barr, S. and M.M. Ogill 1989. Influence of external meteorology on nocturnal valley drainage winds, Journal of Applied Meteorology, 28: 497-517. Barry, R.G. 1990. Changes in Mountain Climate and Glacio-Hydrologic Responses. Mountain Research and Development 10: 161-170. Barry, R.G. 1992. Mountain Weather and Climate, 2nd edition, London: Routledge. Barry, R.G. and C.C. Van Wie 1974, Topo- and microclimatology in alpine areas, In: J.D. Ives and R.G. Barry (eds.), Arctic and Alpine Environments, London: Methuen, pp. 73-83. Bates, G.T. 1990. A case study of the effects of topography on cyclone development in the western United States. Monthly Weather Review, 118: 1808–1825. Bedford, D.P. and R.G. Barry 1995. Glacier trends in the Caucasus, 1960s to 1980s, Physical Geography, 15: 414-424. Benizou, P. 1989. Taking topography into account for network optimization in mountainous areas, In: B. Suvruk (ed.) Precipitation Measurements, WMO/IAHS/ETH Workshop on Precipitation Measurement, Zurich, pp. 307-312, Swiss Federal Institute of Technology. Berg, N.H. 1988 Mountain-top riming at sites in California and Nevada, U.S.A., Arctic and Alpine Research, 30: 429-447.

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Rango, A., J. Martinec, A.T.C. Chang, J.L. Foster, and V. van Katwijk 1989. Average water equivalent of snow in a mountain basin using microwave and visible satellite data, IEEE Transactions of Geoscience and Remote Sensing, 27: 740-745. Rao, G.V., and S. Erdogan 1989. The atmospheric heat source over the Bolivian Plateau for a mean January, Boundary-Layer Meteorology, 46: 13-33. Raymond, D.J. and M.H. Wilkening 1980. Mountain induced convection under fair weather conditions, Journal of Atmospheric Sciences, 37: 2693-2706. Rebetez, M. 1995. Seasonal relationships between temperature, precipitation and snow cover in a mountainous region, Theoretical and Applied Climatology, 54: 99-106. Rebetez, M., R. Lugon, and P.A. Baeriswyl 1997. Climate change and debris flows in high mountain regions: the case study of the Ritigraben torrent (Swiss Alps), Climatic Change, 36: 371-389. Reid, S.J. 1996. Pressure gradients and winds in Cook Strait, Weather and Forecasting, 11: 476488. Reid, S.J. 1997. Modelling of channelled winds in the high wind areas of New Zealand, Weather and Climate, 17: 3-22. Reid S.J. and R. Turner 1997. Wind storms, Tephra, 16: 24-32. Reinking, R.F., J.B. Snider, J.L. Coen 2000. Influences of storm-embedded orographic gravity waves on cloud liquid water and precipitation. Journal of Applied Meteorology, 39: 733–759. Reiter, E.R. and M.Tang 1984. Plateau effects on diurnal circulation patterns, Monthly Weather Review, 112: 638-651. Revell, C.G. 1984. Annual and Diurnal Variation of Thunderstorms in New Zealand and Outlying Islands, Miscellaneous Publication 170, Wellington: New Zealand Meteorological Service. Reynolds, D.W. 1996. The effects of mountain lee waves on the transport of liquid propanegenerated ice crystals. Journal of Applied Meteorology,. 35: 1435–1456. Richard, E., P. Mascart, and E.C. Nickerson 1989. On the role of surface friction in downslope windstorms, Journal of Applied Meteorology, 28: 241-251. Richner, H. and P.D. Phillips 1984. A comparison of temperature from mountaintops and the free atmosphere- their diurnal variation and mean difference, Monthly Weather Review, 112: 13281340.

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Ring, S.L. 1991. Snow-drift modeling and control, In: A.H. Perry and L.J. Symons (eds.) Highway Meteorology, London: E & FN Spon., pp. 77-90. Roberts, M.R. and F.S. Gilliam 1995. Patterns and mechanisms of plant diversity in forested ecosystems: Implications for forest management, Ecological Applications, 5: 969-977. Robinson, P.J. 1997. Climate change and hydropower generation, International Journal of Climatology, 17: 983-996. Robinson, D.A. and A. Frei 2000. Seasonal variability of Northern Hemisphere snow extent using visible satellite data, The Professional Geographer, 52: 307-315. Rochefort, R.M., R.L. Little, A. Woodward, and D.L. Peterson 1994. Changes in subalpine tree distribution in western North America: A review of climate and other factors, The Holocene, 4: 89-100. Romero, R., S. Alonso, E.C. Nickerson, C. Ramis 1995. The influence of vegetation on the development and structure of mountain waves. Journal of Applied Meteorology,. 34: 2230– 2242. Rowe, H.D., Dunbar, R.B., Mucciarone, D.A., Seltzer, G.O., Baker, P.A., and S. Fritz 2002. Insolation, moisture balance and climate change on the South American Altiplano since the last glacial maximum, Climatic Change, 52: 175-199. Runeckles, V.C. and S.V. Krupa 1994. The impact of UV-B Radiation and ozone on terrestrial vegetation, Environmental. Pollution, 83: 191-213. Ryan, B.C. 1977. A mathematical model for diagnosis and prediction of surface winds in mountainous terrain, Journal of Applied Meteorology, 16: 571-584 Sassen, K. and H. Zhao 1993. Supercooled liquid water clouds in Utah winter mountain storms: cloud-seeding implications of a remote-sensing dataset, Journal of Applied Meteorology, 32: 1548–1558. Saunders, I.R. and W.G. Bailey 1994. Radiation and energy budgets of alpine tundra environments of North America, Progress in Physical Geography, 18: 517-538. Schaaf, C.L.B., J. Wurman, and R.M. Banta 1988. Thunderstorm-producing terrain features, Bulletin of the American Meteorological Society, 69: 272-277. Schermerhorn, V.P. 1967. Relations between topography and annual precipitation in Western Oregon and Washington. Water Resources Research 3: 707-711.

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Schmidli, J., C, Schmutz, C, Frei, H, Wanner, C, Schär 2002. Mesoscale precipitation variability in the region of the European Alps during the 20th century, International Journal of Climatology, 22: 1049-1074. Schroeder, S.R. and J.P. McGuirk 1998. Widespread tropical atmospehric drying from 1979 ti 1995m Geophysical Research Letters, 25: 1301-1304. Schultz, D. M., W. J. Steenburgh, R J. Trapp, J. Horel, D.E. Kingsmill, L.B. Dunn, W.D. Rust, L. Cheng, A. Bansemer, J. Cox, J. Daugherty, D.P. Jorgensen, J. Meitín, L. Showell, B.F. Smull, K. Tarp, M. Trainor 2002. Understanding Utah Winter storms: The Intermountain Precipitation Experiment. Bulletin of the American Meteorological Society, 83: 189–210. Sevruk, B. 1986. Correction of precipitation measurements, In: B. Sevruk (ed.) Proceedings, International Workshop on the Correction of Precipitation Measurements, Instruments and Observing Methods Report no. 24 (WMO/TD no. 104), pp. 13-23, Geneva: World Meteorological Organization. Sevruk, B. 1989. Precipitation Measurement, WMO/IAHS/ETH Workshop on Precipitation Measurement, Zurich, Swiss Federal Institute of Technology. Shrestha, A.B., Ca.P. Wake, P.A. Mayewski, and J.E. Dibb 1999. Maximum temperature trends in the Himalaya and its vicinity: An analysis based on temperature records from Nepal for the period 1971–94. Journal of Climate, 12: 2775–2786. Sinclair, M. R. 1993. A diagnostic study of the extratropical precipitation resulting from Tropical Cyclone Bola, Monthly Weather Review, 121: 2690-2707. Sinclair, M. R. 1994. A diagnostic model for estimating orographic precipitation. Journal of Applied Meteorology, 33: 1163–1175. Sinclair, M.R., D.S. Wratt, R.D. Henderson, and W.R. Gray 1997. Factors affecting the distribution and spillover of precipitation in the Southern Alps of New Zealand—A Case Study. Journal of Applied Meteorology, 36: 428–442. Smith, A.A. 1982. The Mount Washington Observatory- 50 years old, Bulletin of the American Meteorological Society, 63: 986-994. Smith, R.B. 1979. The influence of mountains on the atmosphere, Advances in Geophysics, 21: 87-230. Smith, R.B. 1987. Aerial observations of the Yugoslavian bora, Journal of Atmospheric Sciences, 44: 269-297. Stull, R.B. 1988. An Introduction to Boundary Layer Meteorology, Dordrecht: Kluwer Academic Publishers.

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Sturman, A.P. 1987. Thermal influences on airflow in mountainous terrain, Progress in Physical Geography, 11: 183-206. Sturman., A.P. and N.J. Tapper 1996. The Weather and Climate of Australia and New Zealand, Melbourne: Oxford University Press. Susong, D., D. Marks, and D. Garen 1999. Methods for developing time-series climate surfaces to drive topographically distributed energy- and water-balance models, Hydrological Processes, 13: 2003-2021. Tabler, R.D. 1975. Predicting profiles of snowdrifts in topographic catchmets, Proceedings of the 43rd Western Snow Conference, pp. 87-97. Tabony, R.C. 1985. The variation of surface temperature with altitude, Meteorological Magazine, 114: 37-48. Tampieri, F. 1987. Separation features of boundary-layer flow over valleys, Boundary-Layer Meteorology, 40: 295-307. Tang, M. and E.R. Reiter 1984. Plateau monsoons of the northern hemisphere: a comparison between North America and Tibet, Monthly Weather Review, 112: 617-637. Tappenier, U. and A. Cernusca 1989. Canopy structure and light climate of different alpine plant communities, Theoretical and Applied Climatology, 40: 81-92. Taylor, P.A., P.J. Mason, and E.F. Bradley 1987. Boundary-layer flow over low hills, BoundaryLayer Meteorology, 37: 107-132. Taylor, P.A., R.I. Sykes and P.J. Mason 1989. On the parameterization of drag over small-scale topography in neutrally-stratified boundary layer flow, Boundary-Layer Meteorology, 48: 409422. Tesche, T.W. 1988. Numerical simulation of snow transport, deposition and redistribution, Proceedings of the Western Snow Conference 56th Annual Meeting, pp. 93-103. Thompson, B.W. 1986. Small-scale katabatics and cold hollows, Weather,41: 146-153. Thompson, C.S., M.R. Sinclair and W.R. Gray 1997. Estimating long-term annual precipitation in a mountainous region from a diagnostic model, International Journal of Climatology, 17: 997-1007. Thompson, W.F. 1990. Climate related landscapes in world mountains: Criteria and map, Zeitschrit fur Geomorphologie, Supplement, 78: 92pp. Trenberth, K.E. 1999. Conceptual framework for changes of extremes of the hydrological cycle with climate change. Climatic Change, 42: 327-339.

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Tucker, D.F. and N.A. Crook 1999. The generation of a mesoscale convective system from mountain convection. Monthly Weather Review, 127: 1259–1273. Turner, H. 1980. Types of microclimate at high elevations, In: U. Benecke and M.R. Davis (eds.), Mountain Environments and Subalpine Tree Growth, Wellington Forest Research Institute, New Zealand Forest Service, pp. 21-26. Uddstrom, M.J., J.A. McGregor, W.R. Gray, and J.W. Kidson 2001. A high-resolution analysis of cloud amount and type over complex topography. Journal of Applied Meteorology, 40: 16– 33. Vermeulen A.T., G.D. Wyers, F.G. Romer, N.F.M. van Leewen, N.F.M. Draaijers, and J.W. Erimim 1997. Fog deposition on a coniferous forest in the Netherlands, Atmospheric Environment, 31: 388-396. Valko, P. 1980. Some empirical properties of solar radiation and related parameters, In: An Introduction to Meteorological Measurements and Data Handling for Solar Energy Applications, Chapter 8, DOE/ER-0084, U.S. Department of Energy. Vergeiner, I. and E. Dreiseitl 1987. Valley winds and slope winds- observations and elementary thoughts, Meteorology and Atmospheric Physics, 36: 264-286. Vosper, S.B. and D.J. Parker 2002. Some perspectives on wave clouds, Weather, 57: 3-7. Walsh, K. 1994. On the influence of the Andes on the general circulation of the Southern Hemisphere. Journal of Climate, 7: 1019–1025. Wang, G., Hobbs, N.T., Singer, F.J., Ojima, D.S., and B.C. Lubow 2002. Impacts of climate changes on elk population dynamics in Rocky Mountain National Park, Colorado, U.S.A., Climatic Change, 54: 205-223. Wearne, L.J. and J. W. Morgan 2001. Recent forest encroachment into subalpine grasslands near Mount Hotham, Victoria, Australia Arctic, Antarctic, and Alpine Research, 33: 369-377. Whalsley, J.L. P.A. Taylor, and J.R. Salmon 1989. Simple guidelines for estimatingwind speed variations due to small-scale topographic features- an update, Climatological Bulletin, 23: 3-14. Whiteman, C.D. and T.B. KcKee 1978. Air pollution implications of inversion descent in mountain valleys, Atmospheric Environment, 2: 2151-2158. Williams, S.H. and J.A. Lee 1995. Aeolian saltation transport rate: An example of the effect of sediment supply, Journal of Arid Environments, 30: 153-160 Willis, I. and J.M. Bonvin 1995. Climate change in mountain envirionments: Hydrological and resource implicaitons, Geography, 247-261.

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Whipperman, F. 1984. Airflow over and in broad valleys: channeling and counter-current, Contributions to Atmospheric Physics, 57: 184-195. Whiteman, C.D. 1990. Observations of thermally developed wind systems in mountainous terrain, In: W. Blumen (ed.) Atmospheric Processes over Complex Terrain, Meteorological Monograph 23(45), pp. 5-42, Boston: American Meteorological Society. Woodbridge, G.L. D.G. Fox and R.W. Furman 1987. Airflow patterns over and around a large three-dimensional hill, Meteorology and Atmospheric Physics, 37: 259-270. Wratt, D.S., R.N. Ridley, M.R. Sinclair, H. Larsen, S.M. Thompson, R. Henderson, G.L. Austin, S.G. Bradley, A. Auer, A.P. Sturman, I.F. Owens, B.B. Fitzharris, B.F. Ryan, and J.F. Gayet 1996. The New Zealand Southern Alps experiment, Bulletin of the American Meteorological Society, 77: 683-692. Yoshino, M.M. 1973. Studies on wind-shaped trees: their classification, distribution nad significance as a climatic indicator, Climatological Notes, 12: 1-52. Yoshino, M.M. 1984. Thermal belt and cold air drainage on the mountain slope and cold air lake in the basin at quiet, clear night, Geojournal, 8: 235-250.

Figure list Fig. 4.1. Length of day light received at each latitude during summer (left) and winter (right) solstice in the northern hemisphere. (After Rumney 1968; p. 90) Fig. 4.2. The general distribution of global atmospheric pressure systems and general circulation of the atmosphere. These winds dictate global climatic patterns associated with latitude. The general latitudinal climatic zones are shown along the right side of the diagram. Fig. 4.3. Generalized profile showing the decrease of atmospheric pressure with altitude. (Adapted from several sources) Fig. 4.4 The influence of the Olympic Mountains on the wind field and precipitation. The arrows are flow lines indicating wind direction. Distance between the flow lines indicates relative speed, the closer they are to one another the faster the wind in that region. Notice that the flow lines are evenly spaced over the Pacific Ocean. As they are deflected through the Strait of Juan de Fuca, the wind speed increases. Also notice that winds are funneled up the western valleys of

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the Olympics, concentrating moist air and increasing precipitation at the Hoh Rain Forest (3800 mm), while Sequim only receives 430 mm in the rain shadow. (Author) Fig. 4.5. Spectral distribution of direct solar radiation at the top of the atmosphere and at sea level. Calculations are for clear skies with the sun directly overhead. Also shown is the spectral distribution of cloud light and sky light. The graph is plotted on a wave number scale in cm-1 that is the reciprocal of the wavelength and is directly proportional to the frequency of light, to allow display of the full spectrum (a wavelength plot has difficulty including the visible and infrared together). The total area under the upper curve is the solar constant, 2.0 cal. cm-2 min-1 (1365 W/m2). (After Gates and Janke 1966, p. 42) Fig. 4.6. Spectral transmissivity of the atmosphere at 4,200 m (14,000 ft.) and at sea level for latitude 40˚N at summer and winter solstice. The attenuation shown here is for clear skies and is due entirely to ozone absorption. When the effects of dust, water vapor, and other impurities are included, the difference in transmissions between high and low elevations becomes considerably greater. (After Gates and Janke 1966, p. 45) Fig. 4.7. Direct solar radiation (Cal. cm-2 hr-1) received on different slopes during clear weather at 50˚ N. lat. Three slopes are shown: north, south, and east-facing (west would be a mirror image of east), for summer and winter solstice and equinox (vernal is a mirror image of autumnal). The lefthand side of each diagram shows the distribution of solar energy on a horizontal surface (0˚ gradient) and is therefore identical for each set of 3 in the same column. The righthand side of each diagram represents a vertical wall (90˚ gradient). The top of each diagram shows sunrise and the bottom shows sunset. As can be seen, the north- and south-facing slopes experience a symmetrical distribution of energy, while the east and west reveal an asymmetrical distribution. Thus, on the east-facing slope during summer solstice the sun begins shining on a vertical cliff at about 4:00 a.m and highest intensity occurs At 8:00 a.m. By noon the cliff passes into shadow. The opposite would hold true for a west-facing wall: it would begin receiving the direct rays of the sun immediately past noon. The bottom row of diagrams illustrates a south-facing slope. During equinox days and nights are equal, so the distribution of energy is equal. During winter solstice the sun strikes south-facing slopes of all gradients at the same time (sunrise), but during summer the sun rises farther to the northeast, so some time elapses before it can shine on a south-facing slope. This difference in time increases with steeper slopes: for example, a 30˚ south-facing slope would receive the sun at about 5:00 a.m. and would pass into shadow at about 6:30 p.m., while a 60˚ south-facing slope would receive the sun at 6:30 a.m. (11/2 hrs. later) and the sun would set at 5:30 p.m. (1 hr. earlier). On a north-facing slope (top row of diagrams) during summer, slopes up to 60˚ receive the sun at the same time, but if the slope is greater than 60% the sun cannot shine on it at noon; hence the "neck" cut out of the righthand margin. Steep north-facing slopes at this latitude would only receive the sun early in the morning and late in the evening. During the winter solstice only north-facing slopes with gradients of less than 15˚ would receive any sun at all. (After Geiger 1965, p.374) Fig. 4.8. Topo- and micro-climatic influences of slope and aspect on vegetation types. The northern hemisphere example is given where more solar receipt on south-facing slopes warms

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temperatures to where forest is replaced by grass. North-facing slopes are shaded and cooler with more soil moisture retention and thicker forests. On a larger scale, forests move down valleys following moisture and cooler temperatures created by cold air drainage. (After Kruckeberg 1991) Fig. 4.9. Settlement in relation to noonday shadow areas during winter in the upper Rhône Valley, Switzerland. (From Garnett 1935, p. 602) Fig. 4.10. View of an east-west valley near Davos, Switzerland, showing settlement and clearing on the sunny side (south-facing), while the shady side (north-facing) is left in forest. (Larry Price) Fig. 4.11. Mean annual temperature with altitude in the southern Appalachian Mountains. Dots represent U.S. Weather Bureau First Order Stations in Tennessee and North Carolina. Temperatures were calculated for period 1921-1950. (Adapted from Dickson 1959, p. 353) Fig. 4.12. Distribution of mean annual temperature (˚C) in a transect across the Mexican Meseta from Mazatlan to Veracruz. The temperature over the plateau at 3,000 m (10,000 ft.) is about 3˚C (5.4˚F) higher than over the coastal stations, owing to greater heating of the elevated land mass. (Adapted from Hastenrath 1968, p.123) Fig. 4.13. Rice terraces on steep slopes in the Himalayas, near the upper limit for rice cultivation. Most are dry terraces; those in lower left are fed by a spring in the slope and are used for growing wet rice. A village is situated among the dry terraces in the upper part of the slope. The somewhat muted terraces to the right are apparently former terraces that have been abandoned. (Harold Uhlig, University of Giessen) Fig. 4.14. Cross-section of an enclosed basin, Gstettneralm, in the Austrian Alps, showing a temperature inversion in early spring. Elevation of valley bottom is 1,270 m (4,165 ft.). Note increase in temperature (˚C) with elevation above valley floor, especially the rapid rise directly above the pass. This results from the colder air flowing into a lower valley at this point. (After Schmidt 1934, p. 347) Fig. 4.15. Diurnal temperature range at different elevations on Mount Fuji, Japan. The difference between high and low altitudes is much more exaggerated in winter (left) than in summer (right). (After Yoshino 1975, p. 193) Fig. 4.16. Vertical profile of soil and air temperatures (˚C) under clear skies on a well-drained alpine tundra surface at 3,580 m (11,740 ft.) in the White Mountains of California. Note the tremendous gradient occurring immediately above and below the soil surface. The slightly higher temperatures at a depth of 25-30 cm (10-12 in.) are a result of the previous day's heating and are out of phase with present surface conditions. (After Terjung et al. 1969a, p. 256) Fig. 4.17. Daily and seasonal temperature distribution in a subarctic continental (a) and alpine tropical (b) climate. The opposite orientation of the isotherms reflects the fundamental differences in daily and seasonal temperature ranges in the two contrasting environments. The

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subarctic continental station (a) experiences a small daily temperature range (read vertically) but a large annual range (read horizontally). Conversely, the high altitude tropical station (b) experiences a much greater daily temperature range than the annual range. (Adapted from Troll 1958a, p. 11) Fig. 4.18. Freeze-thaw regimes at different latitudes and altitudes. Frost-free days indicate the number of days when freezing did not occur, ice days are those when the temperature was continually below freezing, and frost alternation days are the days when both freezing and thawing occurred. Note that the greatest number of these occur in tropical mountains. (Adapted from Troll 1958a, pp. 12-13) Fig. 4.19. Average annual absolute humidity (mass of water vapor per unit volume, g/m3) with elevation on the humid eastern and arid western side of the tropical Andes. Horizontal lines provide a measure of the annual range of the monthly means of absolute humidity. The extremes are largely a reflection of the wet and dry seasons. Profiles are calculated as a function of height, according to starting values at Lima and Amazonas, based on empirical formulas obtained from observations in the Alps. The tropical-station data indicate that the decrease in vapor density with height is less pronounced than in middle latitudes. (Adapted from Prohaska 1970, p. 3) Fig. 4.20. Mean annual evaporation from reservoirs at different elevations in the Sierra Nevada of central California. (After Longacre and Blaney 1962, p. 42) Fig. 4.21. Diagrammatic representation of daily changes in relative humidity with altitude on northand south-facing forested slopes during August in the mountains of northern Idaho. Dotted line represents the altitude where minimum relative humidities occur at different times during the twenty-four-hour cycle. Note that both the highest and lowest relative humidities occur in the valley bottoms, where the greatest temperature extremes are also found. (Adapted from Hayes 1941, p. 17) Fig. 4.22. Annual average precipitation. Fig. 4.23. Cross-section of atmosphere above the Santa Catalina Mountains near Tuscon, Arizona, on a summer day in 1965. Measurements were made by flying transects across the range in an instrument-equipped airplane. Profiles show changes in mixing ratio (a measure of humidity) and temperature (˚K) at the different altitudes before sunrise (6:15 AM) and after sunrise (10:41 AM). Note that considerable warming, increased humidity, and increased instability of the air, all develop after sunrise, especially on the south side of the range. This leads to convectional lifting, cloud formation, and localized precipitation over the mountains. (After Braham and Draginis 1960, pp. 2-3) Fig. 4.24. Contribution of fog drip to precipitation during twenty-eight-week study period (October 1972 to April 1973) on the forested northeast slopes of Mauna Loa, Hawai‘i. Numbers show precipitation totals in millimeters. Those in parentheses indicate fog drip. Percentages are the relative amounts contributed to the total by fog drip at each station. (After Juvik and Perreira 1974, p. 24)

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Fig. 4.25. Relationship between the number of foggy or cloudy days and elevation in the mountains of Japan. (Dots represent data from weather stations at various altitudes.) The elevation of greatest cloudiness is 1,500-2000 m (5,000-6,000 ft.), where clouds develop almost daily, especially in August. This is caused by the inflow of cool marine air at these levels. The actual height of maximum cloudiness varies from one season to another and from one mountain range to another. (After Yoshino 1975, p. 205) Fig. 4.26. Rime accumulation on newly constructed Palmer ski lift at 2,380 m (7,800 ft.) on the south side of Mount Hood, Oregon. The heavy rime resulted in discontinuation of lift construction until the following summer. (Bob McGown, December 1977) Fig. 4.27. The effects of a precipitation gauge on surface wind-flow. In the first case (a) the wind may tend to speed up next to the gauge since it must travel farther to get around the obstacle. The lower illustrations (b and c) show that turbulence caused by surface roughness may result in upflow or downflow at the gauge orifice, depending on its location with respect to surrounding topography and wind direction. The lee-eddy created in each situation is a location of snow and dust deposition due to slow (reversed) wind speeds. (Adapted from Peck 1972b, p. 8) Fig. 4.28. Generalized profiles of mean annual precipitation (cm) vs. elevation (m) in the tropics. The shaded area shows the zone of maximum precipitation. (Adapted from Lauer 1975) Fig. 4.29. The "alpine desert" at 4,400 m (14,500 ft.) on Mount Kilimanjaro. View is toward the east from the saddle between Kibo and Mawenzi (pictured). (O. Hedberg, 1948, University of Uppsala) Fig. 4.30. Influence of snowpack and glacier cover on runoff is illustrated by data from side-byside basins of the same size, but different elevations. Water year (Oct. - Sept.) hydrographs showing mean (1938-1999) daily discharge for high elevation and low elevation subbasins of the Nooksack River, Washington. The high elevation, North Fork has a mean elevation of 1311 m, 6% glacier cover, and a mean annual discharge of 22.0 m3/s. The lower elevation, Soutth Fork has a mean elevation of 914 m, no glacier cover, and a mean annual discharge of 20.8 m3/s. (Daily discharge data from U.S.G.S, figure by author) Fig. 4.31. Wind velocity with height above a tundra surface. Note how wind speed increases with distance above the ground, one reason why alpine plants grow so close to the ground. (From Warren-Wilson 1959, p. 416) Fig. 4.32. Wind behavior in relation to microtopography in the Cairngorm Mountains, Scotland. The stippled area represents vegetation. Vertical scale is roughly equivalent to the horizontal. (a) Air movement across a grassy tussock. (b) The movement of air over a rock with a depression occupied by vegetation. (c) A wind-eroded bank. Note the eddies that develop to the lee of small obstacles: wind speed is greatly reduced in these areas and vegetation is better developed. (Adapted from Warren-Wilson 1959, pp. 417-18)

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Fig. 4.33. Schematic representation of slope winds (open arrows) and mountain and valley winds (black arrows). (a) and (b) Day conditions. (c) and (d) Night conditions. (After Defant 1951, p. 665, and Hindman 1973, p. 199) Fig. 4.34. Valley fog in the Coast Range of northern California beginning to dissipate as slope winds strengthen and the return flow develops in the center of the valley. Top photo taken at 9:58 A.M.; bottom photo taken at 10:07 A.m. (Edward E. Hindman, U. S. Navy) Fig. 4.35. Graphic representation of slope and valley winds. The view on the left is looking upvalley at midday. Slope winds are rising along the slopes, while the valley wind and anti-wind are moving opposite each other, up and down the valley. The illustration on right provides a vertical cross-section of the same situation, viewed from the side. The valley wind and anti-wind essentially establish a small convection system. The regional gradient wind is shown blowing above the mountains. If the regional wind is very strong, of course, it may override and prevent development of the slope and valley winds. (Adapted from Buettner and Thyer 1965, p. 144) Fig. 4.36. Idealized cross-section of wind movement in a valley with a glacier near its head. Glacier wind is shown moving downslope in a thin zone immediately next to the ice. Valley wind blows upslope and rides over the glacier wind. At elevations above the mountains the regional gradient-wind may be blowing in still another direction. (After Geiger 1965, p. 414) Fig. 4.37. Diagrammatic representation of typical late-afternoon weather conditions along the western slopes of the Colombian Andes, 5˚N Lat. The Andes provide a barrier to the prevailing easterly wind flow, allowing a thin layer of cool, moist Pacific air to move inland. This causes much cooler conditions and also transports moisture for the formation of clouds as it moves over the ridges. In the Cauca Valley, thunderstorms often result from air flowing down the slopes of the western Andes with enough velocity so that it is forced up the adjoining slopes of the central Andes. This produces a "hydraulic jump" that provides the impetus for cloud formation and thunderstorm activity. (After Lopez and Howell 1967, p. 31) Fig. 4.38. "Waiting for a Chinook," by Charles M. Russell. This small watercolor was sent in a letter to Russell's employers in 1886 to announce the emaciated condition of their cattle. (Courtesy of Montana Stockgrowers Association) Fig. 4.39. Diagrammatic representation of classical development of a foehn (chinook) wind. Temperatures at different locations are based on the assumption that air at the base of mountain on windward side is 10˚C (50˚F). By the time the air has undergone the various thermodynamic processes indicated in its journey across the mountains it reaches the base on the leeward side at 18.1˚C (64.6˚F). (Author) Fig. 4.40. Lee waves resulting from air passing across a mountain barrier. Lee-wave clouds often form at the ridge of the waves. Rotors may develop nearer the ground in the immediate lee of the mountain. (Adapted from Scorer 1967, p. 93)

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Fig. 4.41. Lee-wave clouds forming over the Front Range of the Colorado Rockies. View is toward the west, so wind is southwesterly (from left to right). (Robert Bumpas, National Center for Atmospheric Research) Fig. 4.41. Multi-storied lee-wave clouds forming to the lee of the Front Range of the Colorado Rockies. Formation of lenticular clouds above one another in this fashion indicates different wave amplitudes and increasing instability of the air. (Robert Bumpas, National Center for Atmospheric Research) Fig. 4.43. Satellite photo of the northwestern United States, showing extensive lee-wave cloud development from the lee of the Cascades in Washington and Oregon through the intermountain west of Idaho and Utah. Photo was taken 8 December 1977 from a weather satellite at an attitude of 4,320 km (2,700 mi.) at 40˚N. lat. and 140˚W long. Resolution, or size of features which may be identified, is 1.6 km (1 mi.). (National Oceanic and Atmospheric Administration) Fig. 4.44. Photograph of a rotor along the east face of the Sierra Nevada, California. This powerful roll-like circulation of the air is operating beneath the flat, thin clouds. Dust is being lifted from the floor of Owens Valley to a height of 4,800 m (16,000 ft.). (Robert Symons, courtesy of R. S. Scorer) Fig. 4.45. Air current over mountains, showing superposition of lee-wave trains. The mountain ridge (indicated by dashed line of mountain form) produces a certain wave pattern (dashed streamline) and the other mountain (solid line) produces a different wave pattern (continuous streamline). Together the mountains have the effect of creating an obstacle (indicated by the continuous line). In the upper diagram the wavelength is such that the wave trains cancel out; in the lower diagram the amplitude is doubled. Since the wavelength is determined by the flow of air across the ridge, the same air-stream could produce either large-amplitude lee waves or none at all, depending on its direction. (After Scorer 1967, p. 76) Table 4.1. Average density of suspended particulate matter in the atmosphere with changing elevation (Landsberg 1962, p. 114). Table 4.2. Average water-vapor content of air with elevation in the middle latitudes (Landsberg 1962, p. 110). Table 4.3. Average daily global radiation totals (cal. cm-2 d-1) received on a horizontal surface at different elevations in the Austrian Alps. Data include diffuse and reflected energy as well as direct solar radiation (Geiger 1965, p. 444). Table 4.4.Temperature conditions with elevation in the eastern Alps (after Geiger 1965, p. 444). Table 4.5. Accumulation of rime deposits near Haldde Observatory, Norway. The larger amounts at Talviktoppen and Store Haldde are due to higher wind velocity and cloud frequency at these elevations (Kikler 1937, in Landsberg 1962, p. 186).

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Table 4.6. Average annual precipitation at four ridge sites in a transect up the Front Range of the Colorado Rockies during 1965-1970 (Barry 1973, p. 96). Table 4.7. Mean monthly wind speeds during winter at selected mountain weather-stations, in order of decreasing velocity. Readings were taken above tree-line or in treeless areas but anemometers were located at various heights above the ground (after Judson 1965, p. 13).

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Table 4.4. Mean Air Temperature Annual Number of (˚C) Frost Continuous Elevation January - 1.4 19.5 - 2.5 18.3 - 3.5 17.1 - 3.9 16.0 - 3.9 14.8 - 3.9 13.6 - 4.1 12.4 - 4.9 11.2 - 6.1 9.9 - 7.1 8.7 - 8.2 7.2 - 9.2 5.9 -10.3 4.6 -11.3 3.2 -12.4 1.8 Annual Frost-Free AlternationFrost July 9.0 8.0 7.1 6.4 5.7 4.9 4.0 2.8 1.6 0.4 -0.8 -2.0 -3.3 -4.5 -5.7 Year 20.9 20.8 20.6 19.9 18.7 17.5 16.5 16.1 16.0 15.8 15.4 15.1 14.9 14.5 14.2 Range 272 267 250 234 226 218 211 203 190 178 163 146 125 101 71 Days 67 97 78 91 86 84 81 78 76 73 71 68 66 64 62 Days Days 26 1 37 40 53 63 73 84 99 114 131 151 174 200 232

200 400 600 Soo 1,000 1,200 1,400 1,600 1,800 21000 2,200 2,400 2,600 2,8W 3,000

Table 4.7 Elevation Location

Monthly Wind Speed (mph) Nov. Dec. Jan. Feb. Mar.Apr. 42 25 27 26 15 is is is 42 36 29 24 19 is 15 12 47 39 25 ZZ 20 is 17 19 37 49 24 21 17 15 17 15 43 41 26 20 16 13 34 36 25 17 17 10

Mount Fuiiyama, Japan 3,776 Mount Washington, N.H. 1,909 Jungfraujoch, Switzerland 3,575 Niwot Ridge, Colo. 3,749 21 25 Pic du Midi, France 2,860 Sonnblkk, Austria 3,106 22 16 Berthoud Pass, Colo. 3,621 Mauna Loa, Hawai‘i 3,399


				
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