Snow, ice, avalanches and glaciers
Factors Influencing Avalanche Occurrence
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- The Avalanche as a Hazard, Avalanche Victims and Avalanche Rescue
- Avalanche Forecasting and Mitigation
- GLACIERS
- Types of Glaciers
- Glacial Climatic Response and Mass Balance
- Glacier Thermodynamics and Hydrology
- Glaciers as Landscape Forming Tools
- Glaciated Mountain Landscapes
Factors Influencing Avalanche OccurrenceSnow is a highly variable material: it occurs under different environmental conditions, it displays vastly different mechanical characteristics as its temperature and microstructure change, and it is susceptible to constant modification once on the ground (Armstrong and Williams 1986, McClung and Schaerer 1993). Modern avalanche forecasters use both their practical experience, and their scientific knowledge of snow and avalanches, to predict the probable occurrence of avalanches in time and space (LaChapelle 1980, McClung 2002). Though many of the tools used for avalanche forecasting have improved, such as remote weather systems and some computer models, most avalanche forecasters still use many techniques pioneered over 50 years ago (Seligman 1936, Atwater 1954, Perla and Martinelli 1976, Armstrong and Williams 1986, McClung and Schaerer 1993). Ultimately, there are three main ingredients needed for avalanching: favored terrain, unstable snow (which is a product of both the weather and the existing snowpack) and a trigger. Terrain Avalanches usually occur in the same places (referred to as avalanche paths) due to the weather and terrain relationships. The area where the avalanche initiates is the starting zone, the area through which the avalanche runs is the track and the area where deceleration and deposition take place is the runout zone (Figure 21). In timbered areas, the path and its component parts are usually easy to identify due to the vegetative damage and clearing (Martinelli 1973). The most important terrain factor for avalanches is the slope angle of the starting zone, with favored angles between 30 and 45o and the majority of avalanches originating on slopes from 36 to 39o (Figure 22). In some unique (e.g. damp) snow climates slope angles can sometimes exceed 45o, and avalanches on slopes steeper than 55º occur in some mountain ranges, but typically continuous sluffing cleans the steeper slopes. Slopes less than about 30o do not generally create enough downslope stress to allow for avalanche initiation, although avalanches in motion can flow down slopes shallower than 10o for quite a distance, and momentum from large avalanches can even carry snow up the opposing canyon wall as with the Battleship avalanche path in the San Juan Mountains of Colorado. Overall, the probability of avalanches increases with steeper slopes to a certain point and then decreases. Many other terrain features favor or inhibit avalanches. A convex slope will be more prone to avalanches than a concave slope: the outward bend of a convex slope puts tensional stress on its snow cover, while a concavity of slope strengthens the snow's cohesion through compression. Slope orientation with respect to both the wind and the sun is critical. Leeward slopes are typically more dangerous since the wind can quickly load large quantities of snow onto the slope, and they are often overhung by cornices, which can fall on the slope and act as triggers (Figure 23). Exposure to the sun greatly influences the behavior of snow. North-facing slopes receive little sun during the winter (in the northern hemisphere). This typically results in cold conditions where instabilities may persist for longer periods. On south-facing slopes, the sun strikes the surfaces at a higher angle, causing kinetic and/or melt-refreeze metamorphism. South-facing slopes can be sites of instability due to near surface faceting (Birkeland, 1998) and wet-snow avalanches (U.S. Dept. Agriculture 1968, pp. 29-31, Armstrong and Williams 1986, McClung and Schaerer 1993). Other important terrain variables include surface roughness and groundcover. A slope strewn with large boulders is not as susceptible to avalanching as a smooth surface, at least until the snow covers the boulders. On the other hand, a smooth grassy slope provides no major surface inequalities to be filled by the snow and offers little resistance to sliding. Densely forested slopes (> 1000 conifer trees per hectare on steeper slopes) generally offer good protection from avalanche initiation (McClung and Schearer 1993). However, avalanches may start above the timbered zone and destroy strips of the forest in their path (Figure 24; cf. Figure 8.25 FIX) and many mountain forests have been cut or destroyed in recent centuries (Aulitzky 1967). Once this happens, these zones become more vulnerable to avalanching and it is difficult for the forest to regenerate since trees in the path of the avalanching are continually damaged or killed (Frutiger 1964, Schaerer 1972, Martinelli 1974). Weather Any weather factors that change the mechanical state of the snowpack may also quickly change the avalanche conditions. The three most important weather factors are new snow (or rain), wind, and changes in temperature. More than 80% of all large slides occur either during or shortly after storms. The more snow, the more weight added to the snowpack and the greater the downhill component of force acting on any weak layers in the existing snowpack or on the interface between the new snow and the previously existing snowpack. The rate of snowfall, or the snowfall intensity, is critically important. If the snow falls slowly enough the snowpack may be able to adjust to the new snow load, but rapid snowfall may quickly overload weak layers before they can adjust, thereby causing avalanches. Rain can also be an important factor since it adds weight to the snowpack without any addition in strength; rain falling on a midwinter snowpack may rapidly initiate a large number of avalanches. Wind is also a critically important weather factor for avalanches. Wind re-distributes snow onto the lee sides of ridges, gullies, and other terrain features where it may pile into thick wind drifts. Even relatively gentle winds (about 15 km/hr (Tremper 2001)) are sufficient to move low-density snow. Wind also breaks snow into smaller particles, which then bond together quickly, forming cohesive wind slabs. Areas of wind deposition are areas where more stress has been added to the snowpack, thereby creating more unstable conditions. Temperature is another important weather factor, affecting both the mechanical strength of falling and accumulated snow. The temperature at the time of snow deposition controls the snow-crystal type, the density, the rate of settlement, and the general cohesion. After deposition, the temperature of the snow layer and of the air above is critical to the rate of settling, compaction, internal creep, and metamorphosis. In general, instabilities in the snowpack tend to persist longer when temperatures are cold, and stabilize more quickly when temperatures are moderate. However, high temperatures permit melting and the loss of cohesion among snow layers, resulting in wet snow avalanches. In addition, changes in temperature may affect how easy it is to trigger avalanches, with warmer temperatures decreasing the stiffness of the snow and allowing stresses to penetrate farther into the snowpack, thereby facilitating triggering by skiers or snowmobilers (McClung and Schweizer 1999).
Slabs, weak layers, and bed surfaces often occur in the snowpack, but the snow is not always unstable. For unstable conditions, the stress on the weak layer must exceed its strength. The gravitational force of the pre-existing overlying slab, plus any added weight from new or windblown snow causes the stress on the weak layer. When this exceeds the strength of the weak layer, avalanches result. In reality the mechanisms are more complex since the snowpack is highly spatially variable; cracks initiate in areas of localized weakness within the weak layer before spreading across the slope and triggering the avalanche (Schweizer 1999). Triggers Snow on a slope that is unstable enough to be triggered is called conditionally stable. This condition is particularly dangerous since adding a person on skis, a snowboard, or a snowmobile to the slope may result in an avalanche. The most common triggers for natural avalanches include new or windblown snow, though falling cornices are also important in some areas. Rapidly changing air temperature has occasionally been implicated in avalanche release but its importance has yet to be convincingly demonstrated, though warming temperatures can increase the probability of human-triggered slides (McClung and Schweizer 1999). Explosive blasts can also be used to trigger avalanches as a mitigation measure (see below). Loud sounds have been implicated as avalanche triggers (e.g., the vibration of foot-steps or the sound of voices), and there is said to be an ancient regulation in Switzerland against yodeling during the avalanche season (Allix 1924). However, research suggests this is highly improbable. Even enormous sonic booms are capable of triggering avalanche in only rare cases of extremely unstable snowpacks (Martinelli 1972). The Avalanche as a Hazard, Avalanche Victims and Avalanche RescueOnce humans become involved in the avalanche equation we now have a hazard. It is ironic that despite our greater scientific understanding of avalanches and our considerable investment in their prediction and prevention, the number of accidents continues to increase, primarily because more and more people, especially recreationists, go to the mountains during the winter (Figure 25). This is graphically illustrated by an analysis of avalanche accidents in the United States in a series of volumes called the Snowy Torrents (Williams 1975a, Williams and Armstrong 1984, Logan and Williams 1996). Each situation is objectively evaluated, and the episodes make interesting reading. In the U.S., avalanche fatalities have increased greatly over the last 50 years (1950 to 2001) based on a combination of sources (Williams 1975b as graphed in Price 1981, Armstrong and Williams 1993, CAIC 2002a). Fatalities increased sharply in the late 1970s as ski equipment improved and backcountry skiing became more popular. A second sharp increase in fatalities occurs in the 1990s as ski and snowmobile equipment improved, with the five year average now approaching 30 deaths per year. Since the mid 1990s, snowmobilers as a group have overtaken skiers and climbers in leading avalanche fatalities (Figure 26). Snowmobile technology has improved greatly over the 1990s allowing riders to access more avalanche prone terrain more quickly after storms. Avalanche rescue can be viewed in two distinctly different sets of responses and personnel involved. Response one is the immediate attempt at recovery by the victim’s companions. Due to the likelihood of suffocation, this approach is typically the only successful chance for live recovery. Recent research suggests that over 90% of fully-buried avalanche victims survive the first 15 minutes, but the survival probability drops rapidly to around 30% after 30 minutes (Tremper 2001). The second response category is organized outside rescue such as a ski patrol or search and rescue unit. Even “rapid response” here can mean 24 hours or longer in a backcountry situation so survival in these cases is unlikely unless the victim is in a vehicle or structure. Live avalanche rescue is greatly enhanced by prior training in the use of an electronic avalanche beacon, probe poles and shovel which should be worn or carried by each party member. In addition, each member of the party should realize that they are the victim’s best hope and should not go for help until all other on-site efforts have been exhausted or unless assured help is very close (within minutes). Trained dogs have also become very effective for locating victims. (McClung and Schaerer 1993, Tremper 2001). Avalanche Forecasting and MitigationForecasting Avalanches The safest way to deal with avalanches is to avoid them, and this is possible only by avoiding all snow covered avalanche terrain. This will never occur, however, so long as people wish to live in, play in, and travel through mountains in the winter. Therefore, reducing avalanche accidents relies on avoiding avalanche terrain during times of unstable snowpack conditions, and those times can be best determined through avalanche forecasts. Given the increasing number of avalanche fatalities, there is a need for improved forecast methods and broader forecast area coverage. Scientifically based avalanche forecasting originated in Europe in the early 1900s and has been practiced in the United States since the late 1940s (Atwater 1954, LaChapelle 1980, McClung 2002). Modern avalanche forecasting is a sophisticated yet inexact endeavor. Most avalanche forecasters of today bring an extensive knowledge of mountain meteorology, snow mechanics and terrain analysis to bear on the forecast. Forecasters use meteorological data from remote sites, snowpit and fracture line profile analysis from field sites, the results of explosive tests from nearby ski areas or heli-ski operations, the results of computer simulations for weather and snowpack behavior, sometimes G.I.S. (Geographic Information Systems) analysis of terrain interactions, and, most importantly, a lot of personal experience to derive the forecast for the day. The bottom line is modern avalanche forecasters still use so-called conventional techniques relying on an analysis of terrain with respect to many of the contributory factors defined by Atwater (1954) a half century ago. Currently a number of regional forecast centers exist in the United States (Colorado, Idaho, Montana, Wyoming, Utah, the Pacific Northwest, and Mount Washington in New Hampshire) as well as Canada, New Zealand, mountainous European countries, and China. Most of these centers offer public access to a mountain weather forecast and avalanche hazard assessment through dial-up recordings or Internet sites updated daily. The avalanche hazard assessment system, used internationally, is broken into five classes (CAIC 2002b): LOW (Green) Conditions are generally stable and avalanches are unlikely. Use safe travel techniques anyway, just in case. MODERATE (Yellow) Natural avalanches are unlikely but human-triggered ones are possible. Use caution in steep terrain. CONSIDERABLE (Orange) Natural avalanches are possible and human-triggered ones are probable. Use caution in steep terrain. HIGH (Red) Natural and human-triggered avalanches are likely on a variety of aspects and slope angles. Avoid avalanche terrain including runout zones. EXTREME (Black) Natural and human-triggered avalanches are certain, the hazard is widespread. Avoid avalanche terrain stay on low angle slopes and watch out for small terrain traps.
Passive mitigation through terrain modification consists of placing structures such as walls, pylons, dams, and wedges of various designs either in the snow accumulation zone or immediately above the area to be protected (U.S. Dept. Agriculture 1975, McClung and Schaerer 1993). The strategy in the snow accumulation zone is to break up the solid mass of the snow into smaller units, to anchor the snow base, and to create terraces so that there is less effective slope for each snow unit (LaChapelle 1968, p. 1024) (Figure 27). In the runout zone the structures consist of barricades, walls, and wedges to dams or divert the avalanche. Roofs or sheds are frequently constructed over highways and railway lines along avalanche paths (Figure 28). An interesting technique, discovered quite by accident, is the use of alternately spaced earthen mounds. Around the 1960s a construction firm building diversion structures in an avalanche path near Innsbruck, Austria left several mounds of rubble nearby after they finished the job. The following winter, an avalanche came down the slope but broke up in the rubble mounds before reaching the diversion structures. The Austrians quickly seized upon this idea and built several other mound systems (Figure 29). Avalanche mound systems constructed at several places in North America have also been quite successful. The mounds apparently break up and slow the avalanche by dividing it into cross-currents that dissipate its kinetic energy (LaChapelle 1966, p. 96). The other major approach to avalanche mitigation is the active modification of the snow itself. The oldest, and perhaps still the most successful, method of this type is the artificial triggering of avalanches, typically with explosives This is generally done by using charges placed directly on the slope or through artillery fire (Gardner and Judson 1970, Martinelli 1972, Perla 1978, McClung and Schaerer 1993). The use of explosives allows the release of avalanches from a safe distance, and allows avalanche workers to trigger avalanches when no people are present. The traditional artillery weapons are recoilless rifles and mountain howitzers, since they have good range and accuracy. The release of an avalanche requires at least a 75 mm shell, although a 105 mm shell is better. Normal shells penetrate the snow and lose their effectiveness; hence they are generally equipped with fuses to detonate upon impact or slightly above the snow surface (Perla 1978). In some cases the weapons are stationed at critical avalanche areas and lined up with nearby aiming stakes so that they can be blindfired during storms to release the snow from dangerous build-up. An undesirable side effect of using such weaponry is the accumulation of shrapnel, scrap metal, and dud shells on the mountainside (Perla and Martinelli 1976). With these problems and the limited availability of ammunition, traditional military artillery pieces are being replaced by lighter weight and cheaper methods such as the gas-launched Avalauncher that can be either fixed or pick-up truck mounted. A new twist on the artillery approach is the so-called blaster boxes. These devices are essentially gangs of fixed mortars that lob explosive charges into hard-to-reach starting zones. Other innovations include the use of pre-installed vibrators, the inflation of pre-planted airbags, or the use of fixed re-useable explosive devices like the European GazEx units. Another method of snow stabilization is simply to pack the snow down, often referred to as boot packing. This is used at some ski resorts, where skiers and tracked vehicles are constantly packing the newly fallen snow. During the late 1960s and early 1970s the use of organic chemicals (e.g., benzaldehyde) that inhibit the growth of depth-hoar crystals was tried. The chemical was sprayed on the ground before the first snowfall in late autumn; it then moved upward through the snow, coating the crystals and preventing depth hoar from developing. While the method showed reasonable results, it has not been used extensively probably due to its potential ecological effects (LaChapelle 1966). Research is constantly being done to discover new ways of mitigating avalanches. However, the current mitigation and forecasting techniques have clearly reduced the potential hazards many areas, allowing their winter use. Mitigating avalanches is expensive, however, and it will never be possible to protect all people from all avalanches, especially the dispersed backcountry skiers, snowboarders and snowmobilers who play in steep, avalanche-prone terrain. In the long run, the best defense is carefully locating facilities to avoid avalanche terrain, and carefully timing activities to avoid times of high avalanche danger. GLACIERSA glacier is a mass of moving ice created by the accumulation of snow. The transformation of snow into ice is basically a continuation of the processes of metamorphism, densification and expulsion of air. These processes are accomplished by sublimation, melting, refreezing, and compaction of the ice grains as described earlier. Sublimation, melting, and refreezing are most important when the snow is still near the surface in the active layer, and compaction becomes more important after the snow has been buried under successive annual accumulations (Martini, Brookfield and Sadura 2001). As the snow becomes harder and denser the air spaces between the particles are diminished and eventually closed. Once this stage is reached the snow has become glacial ice. The pattern of progression is clear, although the time needed to accomplish it depends upon the temperature, precipitation, and other conditions (the time is shorter where there is greater warmth and moisture). First, newly fallen snow turns into pea-sized melt-freeze polycrystals of ice at the end of the season (corn snow). The corn snow then becomes firn (also called neve) as it survives from one year to the next. By this time the ice crystals are somewhat denser, with smaller air spaces between them. Firn (or neve) represents an intermediate stage in the progress toward glacial ice, but several more years are required to complete the process. The difference between firn and glacial ice is not always clearly marked, but they can usually be distinguished by the color and density of the material. If there are air spaces between the ice crystals and the ice has a whitish color when viewed in mass, it is firn. On the other hand, if the material has a massive structure with no air spaces between the ice crystals, and a vitreous appearance reflecting and transmitting a blue or greenish color (due to the absorption of red wavelengths) , it is glacial ice and has attained densities between 0.700 and 0.914 (Seligman 1936, p. 118). Types of GlaciersThe full classification of glaciers, which includes both Alpine (mountain) and continental types, is a bit involved and we will consider only the most basic forms of purely mountain glaciers here (Benn and Evans 1998, Martini, Brookfield and Sadura 2001). Mountain or Alpine glaciers range from small cirque glaciers occupying isolated depressions on mountain slopes to major icefields covering all but the highest peaks (Figures 30 and 31). Cirque glaciers are typical of tropical and middle-latitude mountains, while icefields are restricted to subpolar and polar areas. Intermediate between these is the valley glacier, which heads in an accumulation basin and extends down-valley for some distance (Figure 32; cf. Figure 34). Where the ice is sufficient to flow through the valley and accumulate at the base of the mountains, it may spread out upon reaching the flats to form a spatulate tongue; this is a piedmont glacier. The forms of alpine glaciers result from both topography and climate. A glacier cannot develop if the slopes are too steep, since the snow cannot accumulate, even if climatic conditions are favorable. At the opposite extreme, it is unlikely that a glacier would develop on an exposed level upland of limited size, because of wind and sun exposure. Topography can be viewed as the initial mold into which the snow and ice must fit, while climate determines at what level and to what extent glaciers develop in any given topographic situation. In the simplest terms, all that is required for a glacier to form is for more snow to fall than melts. This may be accomplished by combinations of various environmental factors. Consider the differences in energy flux and temperature vs. precipitation regimes in mountains at various latitudes. Mid-latitude mountains receive heavy amounts of snow but summers are very warm, resulting in quick melting and relatively rapid turnover within the system. By contrast, polar mountains receive so little precipitation that the contribution of rime and hoarfrost is often greater than that of snow. At the same time, there is little or no melting. Calving-off of icebergs when glaciers move into the sea is the principal method of depletion. At the other extreme, tropical mountains often display a curious situation: the lower part of the glacier receives more precipitation than the higher part (owing to the zone of maximum precipitation), and melting may take place every day of the year rather than just during summer. Consequently, tropical glaciers are quite short. There are also major differences in environmental conditions through time. Glacial Climatic Response and Mass BalanceToday's glaciers are only a vestige of what existed during the height of the Pleistocene epoch, nevertheless, active mountain glaciers still occur in all latitudes. The Pleistocene is the most recent of about seven “ice ages” the earth has seen and it represents 1.6 to 2 million years of major fluctuation in environmental conditions which “ended” about 10,000 years ago (Boellstorff 1978). During this time at least four major advances of the ice are thought to have taken place in the northern hemisphere (Martini, Brookfield and Sadura 2001). Continental glaciers developed and moved into the middle latitudes of the continent and mountain glaciers grew, advanced downslope and spread into the surrounding lowlands. It is generally felt that each of the major glacial advances coincided with a period of lower temperatures, but the exact requirements for glacial growth are complex and may vary for different regions. Vastly different conditions can develop, depending on the deployment of moisture and temperature regimes between winter and summer. For example, lowering the winter temperature in continental areas may lead to less, not more, snow. This is confirmed by the very small amount of precipitation received in the polar regions. Consequently, if the winter temperatures are lowered in a continental area while the summer temperatures remain the same, the result may be glacial retreat. On the other hand, a small amount of summer cooling, from increased cloud cover may allow the normal snowfall that would otherwise have melted to persist through the season. In some areas this change alone causes increased snow accumulation and glacial growth. Similarly, increased snowfall (with no marked change in temperature) permits some snow survival even under summer temperatures, when the regular amount would normally melt. It can be seen that glaciers are created not simply by the lowering of temperature but by the interplay of different climatic factors. Nevertheless, temperature remains the crucial factor, and there is evidence that temperatures were several degrees lower during the height of the ice ages (see p. 130FIX). Once formed, a glacier responds to and reflects changing climatic conditions. A glacier’s “state of health” from a climatic point of view can be determined by analysis of its mass balance. Whether a glacier grows, retreats, or maintains itself depends upon its mass balance, or budget. This is determined by total snow accumulation as opposed to what is lost through ablation (i.e. melting, calving, evaporation and sublimation). A simple indicator of glacial status on a year-to-year basis is the location of the annual snowline, or firn limit, which represents the maximum extent of summer melting (Figure 33). Since fim is snow at least one year old, the firn limit is the zone dividing this year's snow from last (or, in some cases, fresh snow from glacial ice). The firn limit on a glacier is generally quite distinct near the end of the summer and can be identified by field examination or from aerial photographs (Figure 34). A similar but more sophisticated approach to the study of glacial mass-balance is use of the equilibrium line (or equilibrium line altitude – ELA). Since the ELA is calculated from measurements of snow density, water equivalent, ablation loss, and other internal qualities, it does not always coincide with the firn limit. The equilibrium line altitude marks the zone on the glacier where the mass of the glacier stays approximately the same during the year. The area above the equilibrium line receives an excess of winter snow, resulting in increased mass. The area below the equilibrium line loses more to ablation than is gained by accumulation resulting in decreased mass. If the mass gain above the line equals the mass loss below the line, the glacier is in a (rare) steady-state condition; if the mass gain above the equilibrium line exceeds the mass loss below the line, the result is mass transfer to lower levels and glacial growth, whereas if mass loss exceeds gain, the glacier is shrinking. Over a period of several years, the obvious standard for judgment is the advance or retreat of the glacial tongue itself (Posamentier 1977). The problems involved with conducting formal glacial mass balance studies are discussed by Meier (1962). In spite of these problems, numerous mass balance observations have been conducted in many mountain and subpolar locations. The National Snow and Ice Data Center (NSIDC) reports ” …about 70 percent of the observations come from the mountains of Europe, North America and the former Soviet Union. Mass balance on more than 280 glaciers has been measured at one time or another since 1946” (NSIDC 2002). The most striking fact of glacial mass balance behavior in the last century has been widespread glacial retreat (Figures 35, 36 and 37). There have been short cooling periods with glacial advance as in the 1920s (Hoinkes 1968), and from the 1940s, through the 1960s (Meier 1965, p. 803) but as of the late 1900s and early 2000s the global temperatures continue to rise and most glaciers continue to shrink (Charlesworth 1957, Flint 1971, Leggett 1990, Dyurgerov and Meier 1997a, Dyurgerov and Meier 1997b, Haeberli et. al. 1998). The NSIDC report continues “….although we only have a continuous record from about 40 glaciers since the early 1960s. These results indicate that in most regions of the world, glaciers are shrinking in mass. For the period 1961-1998 "small" glaciers lost approximately 7 meters in thickness, or the equivalent of more than 4,000 cubic kilometers of water. The Global Glacier Mass Balance graph (Figure 37) contains data for average global mass balance for each year from 1961 to 1998 as well as the plot of the cumulative change in mass balance, expressed in cubic kilometers of water, for this period” (NSIDC 2002). Especially alarming is the rate of loss of tropical glaciers (Thompson 2001). Climatic variations and glacial fluctuations such as this are the norm rather than the exception over long periods of time. It was formerly thought that the major ice age advances (stadials) had each lasted about 100,000 years and had been separated by somewhat longer interglacials (interstadials) with warmer and drier climates. However, it is now believed that there were more frequent stadial and interstadial periods each lasted only 10,000 to 30,000 years (Emiliani 1972, Woillard 1978). The last major ice advance melted about 15,000 years ago and we are now in an interglacial period. Shorter-term climatic fluctuations have continued to occur within this framework, however. For example, the final melting of the continental ice was followed by a distinctly warm and dry period, known as the hypsithermal, which lasted from 4,000 to 10,000 years ago (Deevey and Flint 1957). The next major change was a widespread advance of mountain glaciers 2,000-4,000 years ago (Denton and Karlen 1973). Subsequent climatic fluctuations, most notably a warming trend about 1,000 years ago, were followed by a period of glacial advance during the Little Ice Age in the seventeenth and eighteenth centuries (Lamb 1965). Glacial advance during this period did considerable damage to farmland and villages in the Alps and in the mountains of Norway and we have Dutch Masters paintings of people skating on the frozen canals of Holland (Grove 1972, Messerli et al. 1978). The period of modern glacial retreat witnessed within the twentieth century apparently reflects warming and amelioration of conditions following the Little Ice Age. While these temporal generalizations apply to most high mountains on the broad scale, recent evidence reinforces the idea that mountain glaciation is often asynchronous in different mountain ranges even though they may be relatively close neighbors (Gillespie and Molnar 1995, Benn and Owen 1997). It is, of course, too soon to know where this will end. Is it just another small deviation from the norm, or are we in fact near the end of the interglacial period and on the verge of another ice age? If the Milankovitch cycles (long term variations in the Earth’s orbit) are responsible for ice age / interglacial variations over the last 2 million years we should expect to return to glacial conditions over tens of thousands of years (Imbrie et. al. 1992, 1993). See the full volume of Quaternary Research (Vol. 2, No.3, 1972) for some further discussion on this topic. Glacier Thermodynamics and HydrologyGlaciers can be classified based on thermal conditions at the surface and at the base of the ice. Polar glaciers remain well frozen throughout, are frozen to the rock of their beds and move very slowly. Temperate glaciers are at the pressure melting point, are warm based and hence thawed from their bed and thus are free to slide and flow more rapidly. The intermediate condition is Subpolar which are temperate in their inner regions but have cold based margins. While this classification makes important implications in the behavior of glaciers in different climates, it is simplistic and often a single glacier may display multiple thermal behaviors and it is important to avoid lumping entire glaciers into a single thermal classification (Boulton 1972, Sugden 1977, Denton and Hughes 1981, Robin 1976, Paterson 1994). The portions of glaciers at the pressure melting point can develop very complex internal melt water routing systems that take surface melt water and feed it through a network of englacial and basal tunnels ultimately disgorging a large volume of water in an outlet stream near the snout of the glacier. Since the flow of surface water into these glacial plumbing systems is strongly temperature and radiation dependent, the discharge from the outlet stream varies considerably over a diurnal period. Maximum melt occurs shortly after the daily maximum temperature and minimum melt occurs shortly after sunrise following the typical diurnal temperature swing (Gerrard 1990). Given the lag time in response through the internal channels of the ice, similar discharge swings propagate downstream.Glacial Movement Glacial movement is determined by the thickness of the ice, the ice temperature, the steepness of the glacial surface, the condition at the bed (frozen or thawed) and the configuration of the underlying and confining topography. When the thickness of the ice exceeds about 20 m (60 ft.) internal deformation can occur and the glacier can move. In general, greatest movement of the ice takes place in the center of the glacier and decreases toward the edges. Movement is also greatest at the surface and decreases with depth. On a longitudinal basis, movement is greatest near the center at the equilibrium line and least at the head and terminus. The area above the equilibrium line is the zone of accumulation; the area below is the zone of ablation (Figure 33). Therefore, if a glacier is to maintain its form and profile, transfer of mass must be greatest in this zone (Paterson 1969, p. 64). Movement in the accumulation area is generally greatest in winter because of increased snow load, while movement in the ablation zone is greatest in the summer because temperatures are higher and there is more meltwater to serve as lubrication. Velocities are highly variable, ranging from a few centimeters to several meters per day. In steep reaches of the glacier, particularly in ice falls where the ice cascades over cliffs, velocities may be much higher. Greatest velocities occur in the so-called surges of glaciers, where speeds exceeding 100 m (330 ft.) per day may occur for short periods of time. This still little-understood phenomenon has received increasing attention within the last several decades (Meier 1969). Glaciers are thought to move through one or two basic mechanisms (internal deformation and basal sliding) depending on whether the portion of the glacier being considered is frozen to its bed (cold based) or thawed at the base (warm based). Several competing theories have surfaced to describe the internal deformation of ice including the kinematic theory (ice as a viscous fluid with laminar flow), the hydrodynamic theory (ice as a Newtonian viscous fluid) and the plastic flow theory (ice as a plastic material) and the processes involved in glacial flow are still not completely agreed upon (Benn and Evans 1998, Martini, Hooke 1998, Brookfield and Sadura 2001). Historically, it was believed that glaciers deformed much like a viscous liquid with laminar flow. In the latter half of the twentieth century it was realized that ice behaves more like a quasi-plastic polycrystalline solid where there is deformation due to flow or creep, as in the creep of metals. This idea has been mathematically formalized in what is known as Glens Flow Law (Glen 1955, Paterson 1994, Hooke 1998). Ice, of course, is much weaker than most crystalline solids and deforms easily through the action of gravity producing shear stress on its mass causing intragranular yielding. Here the ice crystals yield to shear stress by gliding over one another along basal planes within the lattices of the ice crystals. The individual ice crystals should become internally elongated, but since no such deformations of crystals is found in glaciers, a progressive recrystallization apparently accompanies the deformation (Sharp 1960, p. 46, Hooke 1998). The flow is largely a result of directional shear stress, so the ice is equally plastic throughout (rather than being more plastic at the base owing to greater confining pressures, as is sometimes thought). The primary factors controlling the rate of internal deformation are the depth of the ice and the surface slope of the glacier, the temperature of the ice. The steepness of the bedrock slope beneath the ice is less important, since plastic flow may continue even where there are bedrock depressions and obstacles (Paterson 1969, p. 78, 1994). The other major mechanism involved in glacier movement is that of basal sliding, which involves the slippage of ice en-masse over the rock surface at its base. The abrasions and striations left on bedrock across which glaciers have moved are evidence for this kind of movement (Figure 38). The processes involved are even less well understood than those of plastic flow, since the base of a glacier is inaccessible to direct observation except in rare cases. The important control on basal sliding is the temperature of the ice at the base and the presence of water to serve as a lubricant. Basal sliding does not generally occur in polar glaciers, since the ice freezes to the underlying rock surface. In other regions the temperature of the glacial ice is higher and water may be present along the base. Water may also be released when ice reaches the pressure melting-point. This happens when an obstacle is encountered during glacial movement, the ice is compressed on the upstream side of the obstruction, and the increased pressure causes melting. The meltwater then flows around the obstacle and refreezes to the downstream side where the pressure is less (referred to as regelation). The process is maintained by the latent heat of fusion (given off upon refreezing), that is transmitted by conduction from the freezing area to the melting area, where it helps maintain the melting. Regelation operates only on small obstacles 1-2 m (3-7 ft.) in length, however, because the heat cannot be effectively transmitted through larger features. On larger obstacles the ice undergoes greater deformation and movement is probably due mainly to plastic flow, since the ice immediately next to the obstacle must travel farther and faster in order to keep up with the surrounding glacier mass. The larger the obstacle, the more rapid the ice deformation and movement near the bedrock interface (Weertman 1957, 1964).
Glaciers contain a number of interesting features resulting from the transformation of snow to ice and from downslope movement. Most of these are beyond our present concern, but three of them, crevasses , ogives and moraines, require mention. A crevasse is a crack in the ice that may range up to 15 m (50 ft.) in width, 35 m (115 ft.) in depth, and several tens to hundreds of meters in length. Most are smaller than this, especially in temperate mountain glaciers, where the average crevasse is only 1-2 m (3-7 ft.) wide and 5-10 m (16-33 ft.) deep (Figure 39). Crevasses are among the first structural features to appear on a glacier and may develop anywhere from the head to the terminus. Crevasse formation is primarily a response to tensional stress, so their distribution, size, and arrangement provide useful information on the flow behavior of the ice (Sharp 1960, p. 48). Crevasses occur most often where the middle and the sides of the glacier move at different rates, or where the ice curves around a bend, or where the slope steepens and the rate of movement increases (Figures 32 and 33). Crevasses are most often transverse to the direction of flow, but they can be oriented in any direction. They are also largely restricted to the surface, where the ice is more brittle and fractures easily; the greater pressure at depth results in closure by plastic flow. A special type of crevasse develops at the upslope end of the glacier where the ice pulls away from the rocky headwall. This is known as the bergschrund (Figure 40). Rock debris from the headwall and valley sides falls into the bergschrund and other crevasses and becomes incorporated into the glacier, often not to be seen again until it is released by glacial melting at the terminus. The presence of crevasses, therefore, increases the efficiency of rock transport. Crevasses also hasten ablation by increasing the glacier's surface area, by the pooling of meltwater, and by disaggregating the ice near the terminus. Crevasses pose great danger to travel across glaciers. This is particularly true after a fresh snow has bridged the surface, hiding the underlying chasms from view. For this reason glacial travel is usually attempted only by experienced teams using ropes (Graydon 1996). Another type of ice structure that is interesting is the ogive. Ogives are arcuate bands in the glacier containing a light and a dark ice pair per band. Ogives usually form down slope from an ice fall (Figure 41) and are thought to represent the annual flow of ice through the fall, the dark portion of the band from the summer and the light portion of the band from the winter (Sharp 1981). One of the most conspicuous surface features of mountain glaciers is the linear accumulation of rocky debris oriented in the direction of flow. Known as lateral and medial moraines, these accumulations result from rocks that have fallen onto the ice, ablated from the edges of the ice, and from the debris input of tributary glaciers. When a smaller ice stream joins a larger glacier, it usually carries with it a load of rocky debris along its edges (lateral moraine) that becomes incorporated into the ice as a vertical partition between the two ice masses. The material then becomes a medial moraine on the main glacier (Figure 42). What we see is only the surface expression of the rock debris, which extends into the ice, frequently all the way to the bottom (except for material contributed by smaller ice streams that join at shallower levels) (Figure 43). The presence of moraines on the ice alters the mass balance, since the rock material is dark in color and can absorb more of the sun's energy. On the other hand, if the rocky burden is thick enough it may serve as an insulative cover and inhibit local melting of the underlying ice. This results in the ice on either side melting more quickly, leaving the moraine exposed as a higher ridge. On very large glaciers, moraines can reach heights of up to 40 m (130 ft.) (Flint 1971, p. 108). As the ridge builds through differential melting, some of the rocky material may slide or tumble onto the ice; in this way the moraine is widened and the underlying ice is again exposed to melting. The moraines gradually widen toward the terminus, eventually ending up as a jumble of rock debris covering the terminus of the glacier (called an ablation moraine). If the glacier is retreating, the underlying ice may melt, leaving the rocks lying about in heaps. On the other hand, isolated masses of ice may be preserved indefinitely under the debris as ice-cored moraine. This is essentially the end of the journey for the larger rock material. The finer debris, however, can be transported farther through the action of glacial melt-streams and wind. Glaciers as Landscape Forming ToolsMechanisms of Glacial Erosion When a glacier moves over an area, the ice undergoes plastic deformation to fill every nook and cranny. Movement at the ice-rock interface results in modification of the underlying surface through glacial erosion and transport. The primary processes are abrasion, crushing and plucking or quarrying. Abrasion is the scratching, gouging, and grooving of the surface as the ice, carrying rock particles as tools, moves across it. Obviously, this is most effective when the rocks in the ice are harder than the surface over which they are passing. Pure ice or ice containing softer rocks is relatively ineffective at abrasion, although it may produce smoothed and shiny surfaces (glacial polish). A bedload of fine material will result in tiny scratches and smoothing of surfaces, while large embedded rocks can produce scratches several centimeters deep. Striations are found in greatest abundance on gently inclined terrain where the ice was forced to ascend, since in that way greater pressure is placed on the glacial base. Striations provide good evidence for the direction of glacial movement; some caution should be used in their interpretation, however, since they can be caused by other processes, e.g., avalanches and mass movement. Crushing is the pulverization of rock due to the glacial mass above. Plucking or quarrying is generally considered to be the most potent erosional tool of glaciers. Plucking involves regelation and the lifting and incorporation of surface rubble and bedrock segments into the moving ice. Plucking is aided in its work by crushing and frost-weathering, which operates in front of the glacier, producing frost-shattered rock with many cracks and crevices. As the ice moves, it easily incorporates the loose material and the ice undergoes plastic deformation around the larger rocks until they too are swept along with the mass. This debris becomes part of the glacier's bedload and serves as a tool for abrasion. Plucking also operates when ice reaches the pressure melting point on the upstream side of obstacles and the water moves downslope and refreezes in cracks in the bedrock, creating a bond between the glacial ice and the rock; the continued movement of the ice plucks the individual segments from the bedrock. This process gives an asymmetric profile to the underlying obstacles: the stoss (upstream) side is smoothed and gentle, while the lee (downstream) side becomes steep and irregular, owing to the quarrying which has taken place. Such features provide excellent evidence for the direction of glacial movement. The landscape that extends above the glacial ice is a product of both frost-shattering and glacial erosion (Russell 1933). Frost shattered rocks eventually tumble onto the ice surface for further transport. Glacial erosion constantly takes place as well. A glacier can be thought of as a huge malleable mass completely smothering the surface and picking up loose rock and soil as it moves along. In this way new surface is continually being exposed to the erosive power of the ice. The load a glacier can carry is almost unlimited; a large glacier can easily transport rocks as big as a house. Mechanisms of Glacial Transport Rock material incorporated into the flow of glacial ice can be transported in one of three modes: supraglacial (on top of the ice surface), englacial (within the glacial ice) or basal (at the bottom of the glacier). The most important notion to keep in mind about glacially transported and deposited sediment is that ice can carry any size particle anywhere in its flow including huge boulders right on the surface! In addition to material carried directly by the ice, sediment can also be transported by melted water through the complex plumbing system within the glacier. Sediment carried by melted water is subject to the same hydraulics as sediment in rivers and hence displays different characteristics from sediments laid down directly from the melting ice. The most pronounced difference is the sorted nature of the glacio-fluvial sediments compared with the unsorted nature of the ice-laid deposits. Glaciers will generally continuously bury surface rock material under deepening layers of snow and ice in the accumulation zone and will generally continuously expose melted out rock material in the ablation zone. All of this takes place while the glacier is also moving from the accumulation zone to the ablation zone. The net result is a set of curved flow streamlines that are nested from the surface at the ELA to the sole of the glacier at the head and toe of the glacier (Figure 33). This has the practical effect of taking rocks falling onto the glacier surface near the head on a long trip deep into the glacier at or near the sole of the ice before releasing them from their icy surroundings near the snout. On the other hand, rocks falling onto the surface of the glacier just above the ELA are taken on a short shallow ride into the glacier before re-emerging just down slope of the ELA. A dramatic demonstration of this effect can be found in the disappearance and subsequent discovery of the infamous missing airliner named “Stardust”. Stardust vanished without a trace in 1947 during a trans-Andean flight from Buenos Aries, Argentina to Santiago, Chile. The disappearance was so immediate and complete that the loss of Stardust became attributed to a UFO abduction. In actuality the plane had crashed in poor weather at the head of the Tupangato Glacier and became quickly entombed by an avalanche triggered by the impact. Thus Stardust vanished and began a long slow ride through the bowels of the glacier following the flow streamlines only to re-emerge in 2000 to be discovered by climbers ascending the nearby peak (NOVA 2001). Mechanisms of Glacial Deposition Using the transport mechanisms described above, glaciers may liberate their load (collectively called drift) in one of three ways: first by deposition directly from the melting ice (the deposit material is then referred to generically as glacial till), second by intermediate deposition from meltwater within the ice then by melting of the ice (referred to as glacio-fluvial deposits), or third, by direct deposition from meltwater below the terminus of the ice (referred to as outwash). The importance of knowing these depositional mechanisms is evident when trying to interpret various landforms found in previously glaciated terrain. Only by understanding these processes can anomalous features such as huge boulders of alien rock type littering a landscape (glacial erratics) or hundred foot high sinuous mounds of sorted stream deposits running miles across a landscape (eskers) can be explained. Glaciated Mountain LandscapesThe landscapes of glaciated mountains are among the most distinctive and striking on earth. The features and forms created by ice sculpture are very different from those caused by running water, and glaciated mountains possess a ruggedness and grandeur seldom achieved in un-glaciated mountains. For most of us, the visual image of high mountains is typified by glaciated landscapes with their pyramidal peaks, jagged sawtooth ridges, amphitheater-like basins, and deep elongated valleys where occasional jewel-blue lakes sparkle amid surrounding meadows. It is a landscape largely inherited from the past, when the ice was much more extensive than now. In the western United States alone there were over 75 separate high-altitude glacial areas (Figure 44). Cirque or valley glaciers occupied most of these areas, but in some areas there were mountain ice-caps. The largest areas of former glaciation are in the Yellowstone-Grand Teton-Wind River ranges, the Sierra Nevada, the Colorado Rockies, and the Cascades (Flint 1971, pp. 471-74). Mountains farther north (i.e. the Canadian Rockies, the coast ranges, and the Alaska and Brooks Range) were almost totally inundated while the Yukon River Valley remained ice-free. The most characteristic and dominant feature of mountain glaciation is erosion. Glacial erosion in mountains is facilitated by the channeling of ice into pre-existing valleys which accentuate its depth and velocity. For this reason glaciers erode deeper in mountain areas than the former ice sheets did in continental areas often exceeding erosive depths of 600 m (2,000 ft.) (Flint 1971, p. 114). There is a sharp contrast between the appearance of glaciated uplands and valleys. The ice is thinner on upper surfaces and prone to earlier melting than that in the valleys, where the ice is deeper and more sheltered. The higher surfaces are thus exposed to prolonged weathering. Typical features include sharp, angular ridges and peaks, and accumulations of frost-loosened rock. By contrast, the valleys (glacial troughs) are so smoothed and shaped by the ice that very few sharp or rugged features remain. An exception occurs where the entire upland surface has been overrun by ice so that both upland and valley are smoothed by the ice. The Scottish Highlands and the Presidential Range in New Hampshire are examples (Goldthwait 1970). Features of deposition, moraines and glacio-fluvial debris, are largely restricted to the lower elevations and generally mark the point of maximum extent of the ice or places where the glacier remained for the longest periods or where it re-advanced slightly as it receded. Download 197.56 Kb. Share with your friends:
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