glacial landform


glacial landform

Introduction
 any product of flowing ice and meltwater. Such landforms are being produced today in glaciated areas, such as Greenland, Antarctica, and many of the world's higher mountain ranges. In addition, large expansions of present-day glaciers (glacier) have recurred during the course of Earth history. At the maximum of the last ice age, which ended about 20,000 to 15,000 years ago, more than 30 percent of the Earth's land surface was covered by ice. Consequently, if they have not been obliterated by other landscape-modifying processes since that time, glacial landforms may still exist in regions that were once glaciated but are now devoid of glaciers.

      Periglacial features, which form independently of glaciers, are nonetheless a product of the same cold climate that favours the development of glaciers, and so are treated here as well.

General considerations
      Before describing the different landforms produced by glaciers and their meltwater, the glacial environment and the processes responsible for the formation of such landforms is briefly discussed.

Types of glaciers
      There are numerous types of glaciers, but it is sufficient here to focus on two broad classes: mountain, or valley, glaciers and continental glaciers, or ice sheets, (including ice caps). For information about other types, see the articles ice and glacier (Glacier National Park).

      Generally, ice sheets are larger than valley glaciers. The main difference between the two classes, however, is their relationship to the underlying topography. Valley glaciers are rivers of ice usually found in mountainous regions, and their flow patterns are controlled by the high relief in those areas. In map view, many large valley glacier systems, which have numerous tributary glaciers that join to form a large “trunk glacier,” resemble the roots of a plant. Pancakelike ice sheets, on the other hand, are continuous over extensive areas and completely bury the underlying landscape beneath hundreds or thousands of metres of ice. Within continental ice sheets, the flow is directed more or less from the centre outward. At the periphery, however, where ice sheets are much thinner, they may be controlled by any substantial relief existing in the area. In this case, their borders may be lobate on a scale of a few kilometres, with tonguelike protrusions called outlet glaciers. Viewed by themselves, these are nearly indistinguishable from the lower reaches of a large valley glacier system. Consequently, many of the landforms produced by valley glaciers and continental ice sheets are similar or virtually identical, though they often differ in magnitude. Nonetheless, each type of glacier produces characteristic features and thus warrants separate discussion.

Glacial erosion
      Two processes, internal deformation and basal sliding, are responsible for the movement of glaciers under the influence of gravity (see glacier). The temperature of glacier ice is a critical condition that affects these processes. For this reason, glaciers are classified into two main types, temperate and polar, according to their temperature regime. Temperate glaciers are also called isothermal glaciers, because they exist at the pressure-melting point (the melting temperature of ice at a given pressure) throughout their mass. The ice in polar, or cold glaciers, in contrast, is below the pressure-melting point. Some glaciers have an intermediate thermal character. For example, subpolar glaciers are temperate in their interior parts, but their margins are cold-based. This classification is a broad generalization, however, because the thermal condition of a glacier may show wide variations in both space and time.

      Internal deformation, or strain, in glacier ice is a response to shear stresses arising from the weight of the ice (ice thickness) and the degree of slope of the glacier surface. Internal deformation occurs by movement within and between individual ice crystals (slow creep) and by brittle failure (fracture), which arises when the mass of ice cannot adjust its shape rapidly enough by the creep process to take up the stresses affecting it. The relative importance of these two processes is greatly influenced by the temperature of the ice. Thus, fractures due to brittle failure under tension, known as crevasses, are usually much deeper in polar ice than they are in temperate ice.

      The temperature of the basal ice is an important influence upon a glacier's ability to erode its bed. When basal temperatures are below the pressure-melting point, the ability of the ice mass to slide on the bed (basal sliding) is inhibited by the adhesion of the basal ice to the frozen bed beneath. Basal sliding is also diminished by the greater rigidity of polar ice: this reduces the rate of creep, which, in turn, reduces the ability of the more rigid ice to deform around obstacles on the glacier bed. Thus, the flow of cold-based glaciers is predominantly controlled by internal deformation, with proportionately low rates of basal sliding. For this reason, rates of abrasion are commonly low beneath polar glaciers, and slow rates of erosion commonly result. Equally, the volume of meltwater is frequently very low, so that the extent of sediments and landforms derived from polar glaciers is limited.

      Temperate glaciers, being at the pressure-meeting point, move by both mechanisms, with basal sliding being the more important. It is this sliding that enables temperate glaciers to erode their beds and carve landforms so effectively. Ice is, however, much softer and has a much lower shear strength than most rocks, and pure ice alone is not capable of substantially eroding anything other than unconsolidated sediments. Most temperate glaciers have a basal debris zone from several centimetres to a few metres thick that contains varying amounts of rock debris in transit. In this respect, glaciers act rather like sheets of sandpaper; while the paper itself is too soft to sand wood, the adherent hard grains make it a powerful abrasive system. The analogy ends here, however, for the rock debris found in glaciers is of widely varying sizes—from the finest rock particles to large boulders—and also generally of varied types as it includes the different rocks that a glacier is overriding. For this reason, a glacially abraded surface usually bears many different “tool-marks,” from microscopic scratches to gouges centimetres deep and tens of metres long. Over thousands of years glaciers may erode their substrate to a depth of several tens of metres by this mechanism, producing a variety of streamlined landforms typical of glaciated landscapes.

      Several other processes of glacial erosion are generally included under the terms glacial plucking or quarrying. This process involves the removal of larger pieces of rock from the glacier bed. Various explanations for this phenomenon have been proposed. Some of the mechanisms suggested are based on differential stresses in the rock caused by ice being forced to flow around bedrock obstacles. High stress gradients are particularly important, and the resultant tensile stresses can pull the rock apart along pre-existing joints or crack systems. These pressures have been shown to be sufficient to fracture solid rock, thus making it available for removal by the ice flowing above it. Other possibilities include the forcing apart of rock by the pressure of crystallization produced beneath the glacier as water derived from the ice refreezes (regelation) or because of temperature fluctuations in cavities under the glacier. Still another possible mechanism involves hydraulic pressures of flowing water known to be present, at least temporarily, under nearly all temperate glaciers. It is hard to determine which process is dominant because access to the base of active glaciers is rarely possible. Nonetheless, investigators know that larger pieces of rock are plucked from the glacier bed and contribute to the number of abrasive “tools” available to the glacier at its base. Other sources for the rock debris in glacier ice may include rockfalls from steep slopes bordering a glacier or unconsolidated sediments overridden as a glacier advances.

Glacial deposition
      Debris in the glacial environment may be deposited directly by the ice ( till) or, after reworking, by meltwater streams (outwash). The resulting deposits are termed glacial drift.

      As the ice in a valley glacier moves from the area of accumulation to that of ablation, it acts like a conveyor belt, transporting debris located beneath, within, and above the glacier toward its terminus or, in the case of an ice sheet, toward the outer margin. Near the glacier margin where the ice velocity decreases greatly is the zone of deposition. As the ice melts away, the debris that was originally frozen into the ice commonly forms a rocky and/or muddy blanket over the glacier margin. This layer often slides off the ice in the form of mudflows. The resulting deposit is called a flow-till by some authors. On the other hand, the debris may be laid down more or less in place as the ice melts away around and beneath it. Such deposits are referred to as melt-out till, and sometimes as ablation till. In many cases, the material located between a moving glacier and its bedrock bed is severely sheared, compressed, and “over-compacted.” This type of deposit is called lodgment till. By definition, till is any material laid down directly or reworked by a glacier. Typically, it is a mixture of rock fragments and boulders in a fine-grained sandy or muddy matrix (non-stratified drift). The exact composition of any particular till, however, depends on the materials available to the glacier at the time of deposition. Thus, some tills are made entirely of lake clays deformed by an overriding glacier. Other tills are composed of river gravels and sands that have been “bulldozed” and striated during a glacial advance. Tills often contain some of the tools that glaciers use to abrade their bed. These rocks and boulders bear striations, grooves, and facets, and characteristic till-stones are commonly shaped like bullets or flat-irons. Till-boulders of a rock type different from the bedrock on which they are deposited are dubbed “erratics.” In some cases, erratics (erratic) with distinctive lithologies can be traced back to their source, enabling investigators to ascertain the direction of ice movement of ice sheets in areas where striations either are absent or are covered by till or vegetation.

      Meltwater deposits, also called glacial outwash, are formed in channels directly beneath the glacier or in lakes and streams in front of its margin. In contrast to till, outwash is generally bedded or laminated (stratified drift), and the individual layers are relatively well sorted according to grain size. In most cases, gravels and boulders in outwash are rounded and do not bear striations or grooves on their surfaces, since these tend to wear off rapidly during stream transport. The grain size of individual deposits depends not only on the availability of different sizes of debris but also on the velocity of the depositing current and the distance from the head of the stream. Larger boulders are deposited by rapidly flowing creeks and rivers close to the glacier margin. Grain size of deposited material decreases with increasing distance from the glacier. The finest fractions, such as clay and silt, may be deposited in glacial lakes or ponds or transported all the way to the ocean.

      Finally, it must be stressed that most glacier margins are constantly changing chaotic masses of ice, water, mud, and rocks. Ice-marginal deposits thus are of a highly variable nature over short distances, as is much the case with till and outwash as well.

Erosional landforms

Small-scale features of glacial erosion
      Glacial erosion is caused by two different processes: abrasion and plucking (see above). Nearly all glacially scoured erosional landforms bear the tool-marks of glacial abrasion provided that they have not been removed by subsequent weathering. Even though these marks are not large enough to be called landforms, they constitute an integral part of any glacial landscape and thus warrant description here. The type of mark produced on a surface during glacial erosion depends on the size and shape of the tool, the pressure being applied to it, and the relative hardnesses of the tool and the substrate.

Rock polish
      The finest abrasive available to a glacier is the so-called rock flour produced by the constant grinding at the base of the ice. Rock flour acts like jewelers' rouge and produces microscopic scratches, which with time smooth and polish rock surfaces, often to a high lustre.

Striations
      These are scratches visible to the naked eye, ranging in size from fractions of a millimetre to a few millimetres deep and a few millimetres to centimetres long. Large striations produced by a single tool may be several centimetres deep and wide and tens of metres long.

      Because the striation-cutting tool was dragged across the rock surface by the ice, the long axis of a striation indicates the direction of ice movement in the immediate vicinity of that striation. Determination of the regional direction of movement of former ice sheets, however, requires measuring hundreds of striation directions over an extended area because ice moving close to the base of a glacier is often locally deflected by bedrock obstacles. Even when such a regional study is conducted, additional information is frequently needed in low-relief areas to determine which end of the striations points down-ice toward the former outer margin of the glacier. On an outcrop scale, such information can be gathered by studying “chatter marks (chatter mark).” These crescentic gouges and lunate fractures are caused by the glacier dragging a rock or boulder over a hard and brittle rock surface and forming a series of sickle-shaped gouges. Such depressions in the bedrock are steep-sided on their “up-glacier” face and have a lower slope on their down-ice side. Depending on whether the horns of the sickles point up the glacier or down it, the chatter marks are designated crescentic gouges or lunate fractures. Another small-scale feature that allows absolute determination of the direction in which the ice moved is what is termed knob-and-tail. A knob-and-tail is formed during glacial abrasion of rocks that locally contain spots more resistant than the surrounding rock, as is the case, for example, with silicified fossils in limestone. After abrasion has been active for some time, the harder parts of the rock form protruding knobs as the softer rock is preferentially eroded away around them. During further erosion, these protrusions protect the softer rock on their lee side and a tail forms there, pointing from the knob to the margin of the glacier. The scale of these features depends primarily on the size of the inhomogeneities in the rock and ranges from fractions of millimetres to metres.

P-forms and glacial grooves
      These features, which extend several to tens of metres in length, are of uncertain origin. P-forms (P for plastically molded) are smooth-walled, linear depressions which may be straight, curved, or sometimes hairpin-shaped and measure tens of centimetres to metres in width and depth. Their cross sections are often semicircular to parabolic, and their walls are commonly striated parallel to their long axis, indicating that ice once flowed in them. Straight P-forms are frequently called glacial grooves, even though the term is also applied to large striations, which, unlike the P-forms, were cut by a single tool. Some researchers believe that P-forms were not carved directly by the ice but rather were eroded by pressurized mud slurries flowing beneath the glacier.

Erosional landforms of valley glaciers (glacial valley)
      Many of the world's higher mountain ranges—e.g., the Alps, the North and South American Cordilleras, the Himalayas, and the Southern Alps in New Zealand, as well as the mountains of Norway, including those of Spitsbergen—are partly glaciated today. During periods of the Pleistocene, such glaciers were greatly enlarged and filled most of the valleys with ice, even reaching far beyond the mountain front in certain places. Most scenic alpine landscapes featuring sharp mountain peaks, steep-sided valleys, and innumerable lakes and waterfalls are a product of several periods of glaciation.

      Erosion is generally greater than deposition in the upper reaches of a valley glacier, whereas deposition exceeds erosion closer to the terminus. Accordingly, erosional landforms dominate the landscape in the high areas of glaciated mountain ranges.

Cirques (cirque), tarns, U-shaped valleys, arêtes, and horns
      The heads of most glacial valleys are occupied by one or several cirques (or corries). A cirque is an amphitheatre-shaped hollow with the open end facing down-valley. The back is formed by an arcuate cliff called the headwall. In an ideal cirque, the headwall is semicircular in plan view. This situation, however, is generally found only in cirques cut into flat plateaus. More common are headwalls angular in map view due to irregularities in height along their perimeter. The bottom of many cirques is a shallow basin, which may contain a lake. This basin and the base of the adjoining headwall usually show signs of extensive glacial abrasion and plucking. Even though the exact process of cirque formation is not entirely understood, it seems that the part of the headwall above the glacier retreats by frost shattering and ice wedging (see below Periglacial landforms (glacial landform)). The rock debris then falls either onto the surface of the glacier or into the randkluft or bergschrund. Both names describe the crevasse between the ice at the head of the glacier and the cirque headwall. The rocks on the surface of the glacier are successively buried by snow and incorporated into the ice of the glacier. Because of a downward velocity component in the ice in the accumulation zone, the rocks are eventually moved to the base of the glacier. At that point, these rocks, in addition to the rock debris from the bergschrund, become the tools with which the glacier erodes, striates, and polishes the base of the headwall and the bottom of the cirque.

      During the initial growth and final retreat of a valley glacier, the ice often does not extend beyond the cirque. Such a cirque glacier is probably the main cause for the formation of the basin scoured into the bedrock bottom of many cirques. Sometimes these basins are “over-deepened” several tens of metres and contain lakes called tarns (tarn).

      In contrast to the situation in a stream valley, all debris falling or sliding off the sides and the headwalls of a glaciated valley is immediately removed by the flowing ice. Moreover, glaciers are generally in contact with a much larger percentage of a valley's cross section than equivalent rivers or creeks. Thus glaciers tend to erode the bases of the valley walls to a much greater extent than do streams, whereas a stream erodes an extremely narrow line along the lowest part of a valley. The slope of the adjacent valley walls depends on the stability of the bedrock and the angle of repose of the weathered rock debris accumulating at the base of and on the valley walls. For this reason, rivers tend to form V-shaped valleys. Glaciers, which inherit V-shaped stream valleys, reshape them drastically by first removing all loose debris along the base of the valley walls and then preferentially eroding the bedrock along the base and lower sidewalls of the valley. In this way, glaciated valleys assume a characteristic parabolic or U-shaped cross profile, with relatively wide and flat bottoms and steep, even vertical sidewalls. By the same process, glaciers tend to narrow the bedrock divides between the upper reaches of neighbouring parallel valleys to jagged, knife-edge ridges known as arêtes (arête). Arêtes also form between two cirques facing in opposite directions. The low spot, or saddle, in the arête between two cirques is called a col. A higher mountain often has three or more cirques arranged in a radial pattern on its flanks. Headward erosion of these cirques finally leaves only a sharp peak flanked by nearly vertical headwall cliffs, which are separated by arêtes. Such glacially eroded mountains are termed horns, the most widely known of which is the Matterhorn in the Swiss Alps.

Hanging valleys
      Large valley glacier systems consist of numerous cirques and smaller valley glaciers that feed ice into a large trunk glacier. Because of its greater ice discharge, the trunk glacier has greater erosive capability in its middle and lower reaches than smaller tributary glaciers that join it there. The main valley is therefore eroded more rapidly than the side valleys. With time, the bottom of the main valley becomes lower than the elevation of the tributary valleys. When the ice has retreated, the tributary valleys are left joining the main valley at elevations substantially higher than its bottom. Tributary valleys with such unequal or discordant junctions are called hanging valleys. In extreme cases where a tributary joins the main valley high up in the steep part of the U-shaped trough wall, waterfalls may form after deglaciation, as in Yosemite and Yellowstone national parks in the western United States.

Paternoster lakes
      Some glacial valleys have an irregular, longitudinal bedrock profile, with alternating short, steep steps and longer, relatively flat portions. Even though attempts have been made to explain this feature in terms of some inherent characteristic of glacial flow, it seems more likely that differential erodibility of the underlying bedrock is the real cause of the phenomenon. Thus the steps are probably formed by harder or less fractured bedrock, whereas the flatter portions between the steps are underlain by softer or more fractured rocks. In some cases, these softer areas have been excavated by a glacier to form shallow bedrock basins. If several of these basins are occupied by lakes along one glacial trough in a pattern similar to beads on a string, they are called paternoster (Latin: “our father”) lakes by analogy with a string of rosary beads.

Roches moutonnées (roche moutonnée)
      These structures are bedrock knobs or hills that have a gently inclined, glacially abraded, and streamlined stoss side (i.e., one that faces the direction from which the overriding glacier impinged) and a steep, glacially plucked lee side. They are generally found where jointing or fracturing in the bedrock allows the glacier to pluck the lee side of the obstacle. In plan view, their long axes are often, but not always, aligned with the general direction of ice movement.

Rock drumlins
      A feature similar to roches moutonnées, rock drumlins are bedrock knobs or hills completely streamlined, usually with steep stoss sides and gently sloping lee sides. Both roches moutonnées and rock drumlins range in length from several metres to several kilometres and in height from tens of centimetres to hundreds of metres. They are typical of both valley and continental glaciers. The larger ones, however, are restricted to areas of continental glaciation.

Erosional landforms of continental glaciers
      In contrast to valley glaciers, which form exclusively in areas of high altitude and relief, continental glaciers, including the great ice sheets of the past, occur in high and middle latitudes in both hemispheres, covering landscapes that range from high alpine mountains to low-lying areas with negligible relief. Therefore, the landforms produced by continental glaciers are more diverse and widespread. Yet, just like valley glaciers, they have an area where erosion is the dominant process and an area close to their margins where net deposition generally occurs. The capacity of a continental glacier to erode its substrate has been a subject of intense debate. All of the areas formerly covered by ice sheets show evidence of areally extensive glacial scouring. The average depth of glacial erosion during the Pleistocene probably did not exceed a few tens of metres, however. This is much less than the deepening of glacial valleys during mountain glaciation. One of the reasons for the apparent limited erosional capacity of continental ice sheets in areas of low relief may be the scarcity of tools available to them in these regions. Rocks cannot fall onto a continental ice sheet in the accumulation zone, because the entire landscape is buried. Thus, all tools must be quarried by the glacier from the underlying bedrock. With time, this task becomes increasingly difficult as bedrock obstacles are abraded and streamlined. Nonetheless, the figure for depth of glacial erosion during the Pleistocene cited above is an average value, and locally several hundreds of metres of bedrock were apparently removed by the great ice sheets. Such enhanced erosion seems concentrated at points where the glaciers flowed from hard, resistant bedrock onto softer rocks or where glacial flow was channelized into outlet glaciers.

      As a continental glacier expands, it strips the underlying landscape of the soil and debris accumulated at the preglacial surface as a result of weathering. The freshly exposed harder bedrock is then eroded by abrasion and plucking. During this process, bedrock obstacles are shaped into streamlined “whaleback” forms, such as roches moutonnées and rock drumlins (see above). The adjoining valleys are scoured into rock-floored basins with the tools plucked from the lee sides of roches moutonnées. The long axes of the hills and valleys are often preferentially oriented in the direction of ice flow. An area totally composed of smooth whaleback forms and basins is called a streamlined landscape.

      Streams cannot erode deep basins because water cannot flow uphill. Glaciers, on the other hand, can flow uphill over obstacles at their base as long as there is a sufficient slope on the upper ice surface pointing in that particular direction. Therefore the great majority of the innumerable lake basins and small depressions in formerly glaciated areas can only be a result of glacial erosion. Many of these lakes, such as the Finger Lakes in the U.S. state of New York, are aligned parallel to the direction of regional ice flow. Other basins seem to be controlled by preglacial drainage systems. Yet, other depressions follow the structure of the bedrock, having been preferentially scoured out of areas underlain by softer or more fractured rock.

      A number of the largest freshwater lake basins in the world (e.g., the Great Lakes or the Great Slave Lake and Great Bear Lake in Canada) are situated along the margins of the Precambrian shield of North America. Many researchers believe that glacial erosion was especially effective at these locations because the glaciers could easily abrade the relatively soft sedimentary rocks to the south with hard, resistant crystalline rocks brought from the shield areas that lie to the north. Nonetheless, further research is necessary to determine how much of the deepening of these features can be ascribed to glacial erosion, as opposed to other processes such as tectonic activity or preglacial stream erosion.

      Fjords (fjord) are found along some steep, high-relief coast-lines where continental glaciers formerly flowed into the sea. They are deep, narrow valleys with U-shaped cross sections that often extend inland for tens or hundreds of kilometres and are now partially drowned by the ocean. These troughs are typical of the Norwegian coast, but they also are found in Canada, Alaska, Iceland, Greenland, Antarctica, New Zealand, and southernmost Chile. The floor and steep walls of fjords show ample evidence of glacial erosion. The long profile of many fjords, including alternating basins and steps, is very similar to that of glaciated valleys. Toward the mouth, fjords may reach great depths, as in the case of Sogn Fjord in southern Norway where the maximum water depth exceeds 1,300 metres. At the mouth of a fjord, however, the floor rises steeply to create a rock threshold, and water depths decrease markedly. At Sogn Fjord the water at this “threshold” is only 150 metres deep, and in many fjords the rock platform is covered by only a few metres of water. The exact origin of fjords is still a matter of debate. While some scientists favour a glacial origin, others believe that much of the relief of fjords is a result of tectonic activity and that glaciers only slightly modified preexisting large valleys. In order to erode Sogn Fjord to its present depth, the glacier occupying it during the maximum of the Pleistocene must have been 1,800 to 1,900 metres thick. Such an ice thickness may seem extreme, but even now, during an interglacial period, the Skelton Glacier in Antarctica has a maximum thickness of about 1,450 metres. This outlet glacier of the Antarctic ice sheet occupies a trough, which in places is more than one kilometre below sea level and would become a fjord in the event of a large glacial retreat.

Depositional landforms

Depositional landforms of valley glaciers
Moraines (moraine)
      As a glacier moves along a valley, it picks up rock debris from the valley walls and floor, transporting it in, on, or under the ice. As this material reaches the lower parts of the glacier where ablation is dominant, it is concentrated along the glacier margins as more and more debris melts out of the ice. If the position of the glacier margin is constant for an extended amount of time, larger accumulations of glacial debris (till; see above) will form at the glacier margin. In addition, a great deal of material is rapidly flushed through and out of the glacier by meltwater streams flowing under, within, on, and next to the glacier. Part of this streamload is deposited in front of the glacier close to its snout. There, it may mix with material brought by, and melting out from, the glacier as well as with material washed in from other, nonglaciated tributary valleys. If the glacier then advances or readvances after a time of retreat, it will “bulldoze” all the loose material in front of it into a ridge of chaotic debris that closely hugs the shape of the glacier snout. Any such accumulation of till melted out directly from the glacier or piled into a ridge by the glacier is a moraine. Large valley glaciers are capable of forming moraines a few hundred metres high and many hundreds of metres wide. Linear accumulations of till formed immediately in front of or on the lower end of the glacier are end moraines. The moraines formed along the valley slopes next to the side margins of the glacier are termed lateral moraines. During a single glaciation, a glacier may form many such moraine arcs, but all the smaller moraines, which may have been produced during standstills or short advances while the glacier moved forward to its outermost ice position, are generally destroyed as the glacier resumes its advance. The end moraine of largest extent formed by the glacier (which may not be as extensive as the largest ice advance) during a given glaciation is called the terminal moraine of that glaciation. Successively smaller moraines formed during standstills or small readvances as the glacier retreats from the terminal moraine position are recessional moraines.

Flutes
      The depositional equivalent of erosional knob-and-tail structures (see above) are known as flutes. Close to the lower margin, some glaciers accumulate so much debris beneath them that they actually glide on a bed of pressurized muddy till. As basal ice flows around a pronounced bedrock knob or a boulder lodged in the substrate, a cavity often forms in the ice on the lee side of the obstacle because of the high viscosity of the ice. Any pressurized muddy paste present under the glacier may then be injected into this cavity and deposited as an elongate tail of till, or flute. The size depends mainly on the size of the obstacle and on the availability of subglacial debris. Flutes vary in height from a few centimetres to tens of metres and in length from tens of centimetres to kilometres, even though very large flutes are generally limited to continental ice sheets.

Depositional landforms of continental glaciers
      Many of the deposits of continental ice sheets are very similar to those of valley glaciers. Terminal, end, and recessional moraines are formed by the same process as with valley glaciers (see above), but they can be much larger. Morainic ridges may be laterally continuous for hundreds of kilometres, hundreds of metres high, and several kilometres wide. Since each moraine forms at a discreet position of the ice margin, plots of end moraines on a map of suitable scale allow the reconstruction of ice sheets at varying stages during their retreat.

      In addition to linear accumulations of glacial debris, continental glaciers often deposit a more or less continuous, thin (less than 10 metres) sheet of till over large areas, which is called ground moraine. This type of moraine generally has a “hummocky” topography of low relief, with alternating small till mounds and depressions. Swamps or lakes typically occupy the low-lying areas. Flutes (see above) are a common feature found in areas covered by ground moraine.

      Another depositional landform associated with continental glaciation is the drumlin, a streamlined, elongate mound of sediment. Such structures often occur in groups of tens or hundreds, which are called drumlin fields. The long axis of individual drumlins is usually aligned parallel to the direction of regional ice flow. In long profile, the stoss side of a drumlin is steeper than the lee side. Some drumlins consist entirely of till, while others have bedrock cores draped with till. The till in many drumlins has been shown to have a “fabric” in which the long axes of the individual rocks and sand grains are aligned parallel to the ice flow over the drumlin. Even though the details of the process are not fully understood, drumlins seem to form subglacially close to the edge of an ice sheet, often directly down-ice from large lake basins overridden by the ice during an advance. The difference between a rock drumlin and a drumlin is that the former is an erosional bedrock knob (see above), whereas the latter is a depositional till feature.

Meltwater deposits (outwash)
      Much of the debris in the glacial environment of both valley and continental glaciers is transported, reworked, and laid down by water. Whereas glaciofluvial deposits are formed by meltwater streams (river), glaciolacustrine sediments accumulate at the margins and bottoms of glacial lakes and ponds.

Glaciofluvial deposits
      The discharge of glacial streams is highly variable, depending on the season, time of day, and cloud cover. Maximum discharges occur during the afternoon on warm, sunny summer days, and minima on cold winter mornings. Beneath or within a glacier, the water flows in tunnels and is generally pressurized during periods of high discharge. In addition to debris washed in from unglaciated highlands adjacent to the glacier, a glacial stream can pick up large amounts of debris along its path at the base of the glacier. For this reason, meltwater streams issuing forth at the snout of a valley glacier or along the margin of an ice sheet are generally laden to transporting capacity with debris. Beyond the glacier margin, the water, which is no longer confined by the walls of the ice tunnel, spreads out and loses some of its velocity. Because of the decreased velocity, the stream must deposit some of its load. As a result, the original stream channel is choked with sediments, and the stream is forced to change its course around the obstacles, often breaking up into many winding and shifting channels separated by sand and gravel bars. The highly variable nature of the sediments laid down by such a braided stream reflects the unstable environment in which they form. Lenses of fine-grained, cross-bedded sands are often interbedded laterally and vertically with stringers of coarse, bouldery gravel. Since the amount of sediment laid down generally decreases with distance from the ice margin, the deposit is often wedge-shaped in cross section, ideally gently sloping off the end moraine formed at that ice position and thinning downstream. The outwash is then said to be “graded to” that particular moraine. In map view, the shape of the deposit depends on the surrounding topography. Where the valleys are deep enough not to be buried by the glaciofluvial sediments, as in most mountainous regions, the resulting elongate, planar deposits are termed valley trains. On the other hand, in low-relief areas the deposits of several ice-marginal streams may merge to form a wide outwash plain, or sandur.

      If the ice margin stabilizes at a recessional position during glacial retreat, another valley train or sandur may be formed inside of the original one. Because of the downstream thinning of the outwash at any one point in the valley, the recessional deposit will be lower than and inset into the outer, slightly older outwash plain. Flat-topped remnants of the older plain may be left along the valley sides; these are called terraces (river terrace). Ideally each recessional ice margin has a terrace graded to it, and these structures can be used in addition to moraines to reconstruct the positions of ice margins through time. In some cases where the glacier either never formed moraines or where the moraines were obliterated by the outwash or postglacial erosion, terraces are the only means of ice margin reconstruction.

      Streams that flow over the terminus of a glacier often deposit stratified drift in their channels and in depressions on the ice surface. As the ice melts away, this ice-contact stratified drift slumps and partially collapses to form stagnant ice deposits. Isolated mounds of bedded sands and gravels deposited in this manner are called kames (kame). Kame terraces form in a similar manner but between the lateral margin of a glacier and the valley wall. Glacial geologists sometimes employ the term kame moraine to describe deposits of stratified drift laid down at an ice margin in the arcuate shape of a moraine. Some researchers, however, object to the use of the term moraine in this context because the deposit is not composed of till.

      In some cases, streams deposit stratified drift in subglacial or englacial tunnels. As the ice melts away, these sinuous channel deposits may be left as long linear gravel ridges called eskers (esker). Some eskers deposited by the great ice sheets of the Pleistocene can be traced for hundreds of kilometres, even though most esker segments are only a few hundred metres to kilometres long and a few to tens of metres high.

      Kettles (kettle), potholes, or ice pits are steep-sided depressions typical of many glacial and glaciofluvial deposits. Kettles form when till or outwash is deposited around ice blocks that have become separated from the active glacier by ablation. Such “stagnant” ice blocks may persist insulated under a mantle of debris for hundreds of years. When they finally melt, depressions remain in their place, bordered by slumped masses of the surrounding glacial deposits. Many of the lakes in areas of glacial deposition are water-filled kettles and so are called kettle lakes. If a sandur or valley train contains many kettles, it is referred to as a pitted outwash plain.

Glaciolacustrine deposits
      Glacial and proglacial lakes (lake) are found in a variety of environments and in considerable numbers. Erosional lake basins have already been mentioned, but many lakes are formed as streams are dammed by the ice itself, by glacial deposits, or by a combination of these factors. Any lake that remains at a stable level for an extended period of time (e.g., hundreds or thousands of years) tends to form a perfectly horizontal, flat, terracelike feature along its beach. Such a bench may be formed by wave erosion of the bedrock or glacial sediments that form the margin of the lake, and it is called a wave-cut bench. On the other hand, it may be formed by deposition of sand and gravel from long-shore currents along the margin of the lake, in which case it is referred to as a beach ridge. The width of these shorelines varies from a few metres to several hundred metres. As the lake level is lowered due to the opening of another outlet or downcutting of the spillway, new, lower shorelines may be formed. Most former or existing glacial lakes (e.g., the Great Salt Lake and the Great Lakes in North America) have several such shorelines that can be used both to determine the former size and depth of now-extinct or shrunken lakes and to determine the amount of differential postglacial uplift because they are now tilted slightly from their original horizontal position.

      Where a stream enters a standing body of water, it is forced to deposit its bedload. The coarser gravel and sand are laid down directly at the mouth of the stream as successive, steeply inclined foreset beds. The finer, suspended silt and clay can drift a bit farther into the lake, where they are deposited as almost flat-lying bottomset beds. As the sediment builds out farther into the lake (or ocean), the river deposits a thin veneer of subhorizontal gravelly topset beds over the foreset units. Because the foreset–topset complex often has the shape of a triangle with the mouth of the stream at one apex, such a body of sediment is called a delta. Many gravel and sand pits are located in deltas of former glacial lakes.

      The flat-lying, fine-grained bottomset beds of many large former glacial lakes filled in and buried all of the pre-existing relief and are now exposed, forming perfectly flat lake plains. Cuts into these sediments often reveal rhythmically interbedded silts and clays. Some of these so-called rhythmites have been shown to be the result of seasonal changes in the proglacial environment. During the warmer summer months, the meltwater streams carry silt and clay into the lakes, and the silt settles out of suspension more rapidly than the clay. A thicker, silty summer layer is thus deposited. During the winter, as the surface of the lake freezes and the meltwater discharge into it ceases, the clays contained in the lake water slowly settle out of suspension to form a thin winter clay layer. Such lacustrine deposits with annual silt and clay “couplets” are known as varves.

Periglacial landforms
      In the cold, or periglacial (near-glacial), areas adjacent to and beyond the limit of glaciers, a zone of intense freeze-thaw activity produces periglacial features and landforms. This happens because of the unique behaviour of water as it changes from the liquid to the solid state. As water freezes, its volume increases about 9 percent. This is often combined with the process of differential ice growth, which traps air, resulting in an even greater increase in volume. If confined in a crack or pore space, such ice and air mixtures can exert pressures of about 200,000 kilopascals (29,000 pounds per square inch). This is enough to break the enclosing rock. Thus freezing water can be a powerful agent of physical weathering. If multiple freeze-and-thaw cycles occur, the growth of ice crystals fractures and moves material by means of frost shattering and frost heaving, respectively. In addition, in permafrost regions (see below) where the ground remains frozen all year, characteristic landforms are formed by perennial ice.

Felsenmeers, talus, and rock glaciers
      In nature, the tensional strength of most rocks is exceeded by the pressure of water crystallizing in cracks. Thus, repeated freezing and thawing not only forms potholes in poorly constructed roads but also is capable of reducing exposed bedrock outcrops to rubble. Many high peaks are covered with frost-shattered angular rock fragments. A larger area blanketed with such debris is called a felsenmeer, from the German for “sea of rocks.” The rock fragments can be transported downslope by flowing water or frost-induced surface creep, or they may fall off the cliff from which they were wedged by the ice. Accumulations of this angular debris at the base of steep slopes are known as talus. Owing to the steepness of the valley sides of many glacial troughs, talus is commonly found in formerly glaciated mountain regions. Talus cones are formed when the debris coming from above is channelized on its way to the base of the cliff in rock chutes. As the talus cones of neighbouring chutes grow over time, they may coalesce to form a composite talus apron.

      In higher mountain regions, the interior of thick accumulations of talus may remain at temperatures below freezing all year. Rain or meltwater percolating into the interstices between the rocks freezes over time, filling the entire pore space. In some cases, enough ice forms to enable the entire mass of rock and ice to move downhill like a glacier. The resulting massive, lobate, mobile feature is called a rock glacier. Some rock glaciers have been shown to contain pure ice under a thick layer of talus with some interstitial ice. These features may be the final retreat stages of valley glaciers buried under talus.

permafrost, patterned ground, solifluction deposits, and pingos
      Permafrost is ground that remains perennially frozen (see permafrost). It covers about 20–25 percent of the Earth's land surface today. The “active layer” of soil close to the surface of permafrost regions undergoes many seasonal and daily freeze-thaw cycles. The constant change in the volume of water tends to move the coarser particles in the soil to the surface. Further frost heaving arranges the stones and rocks according to their sizes to produce patterned ground. Circular arrangements of the larger rocks are termed stone rings. When neighbouring stone rings coalesce, they form polygonal stone nets. On steeper slopes, stone rings and stone nets are often stretched into stone stripes by slow downhill motion of the soggy active layer of the permafrost. In other areas, patterned ground is formed by vertical or subvertical polygonal cracks, which are initiated in the soil by contraction during extremely cold winters. During the spring thaw of the active layer, water flows into these cracks, freezes, and expands. This process is repeated year after year, and the ice-filled cracks increase in size. The resulting ice wedges are often several metres deep and a few tens of centimetres wide at the top. Along the sides of ice wedges, the soil is deformed and compressed. Because of this disturbance and sediment that may be washed into the crack as the ice melts, relict patterned ground may be preserved during a period of warmer climate long after the permafrost has thawed. Today, relict patterned ground that formed during the last ice age exists more than 1,000 kilometres to the south of the present limit of permafrost.

      When the active layer of permafrost moves under the influence of gravity, the process is termed gelifluction. The soft flowing layer is often folded and draped on hillsides and at the base of slopes as solifluction, or gelifluction, lobes.

      In some permafrost areas, a locally abundant groundwater supply present at a relatively shallow depth may cause the exceptional growth of ice within a confined area. The sustained supply of liquid water results in the expansion of an increasingly large, lens-shaped ice body. These conical mounds, or pingos (pingo), may be several tens of metres high and hundreds of metres in diameter.

Edward B. Evenson Gunnar Schlieder

Additional Reading
The classic text on glacial geology is Richard Foster Flint, Glacial and Quaternary Geology (1971), encyclopaedic coverage including an extensive bibliography. Recent hypotheses and observations on glacial erosion and deposition are included in David Drewry, Glacial Geologic Processes (1986), even though the coverage of glacial landforms is not complete. David E. Sugden and Brian S. John, Glaciers and Landscape: A Geomorphological Approach (1976, reprinted 1984), is an excellent detailed introduction to glacial landforms and the processes that shaped them. More theoretical emphasis can be found in Clifford Embleton and Cuchlaine A.M. King, Glacial Geomorphology, 2nd ed. (1975), and Periglacial Geomorphology, 2nd ed. (1975). A.L. Washburn, Geocryology: A Survey of Periglacial Processes and Environments (1979), contains numerous explanatory photographs and diagrams. A collection of articles is found in Cuchlaine A.M. King (ed.), Periglacial Processes (1976). The most comprehensive and up-to-date account of glacial geomorphology and sedimentology is Douglas I. Benn and David J.A. Evans, Glaciers and Glaciation (1998). A detailed discussion of the formation of permafrost can be found in Stuart A. Harris, The Permafrost Environment (1986), and in Peter J. Williams and Michael W. Smith, The Frozen Earth: Fundamentals of Geocryology (1989).Edward B. Evenson Gunnar Schlieder

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Universalium. 2010.

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