The transformation of snow into ice glacier is along process that involves a number of stages. The entire series of transformations is referred to as a metamorphosis or diagenesis (Chorley, Schumm & Sugden, 1985). The initial stages involving accumulation of snow include snowfall, which is characterized by snow of density between 0. 05 and 0. 07 g/cm3 (Chorley, Schumm & Sugden, 1985). This is followed by settling of the snow, which involves loose pieces of tiny ice particles that break up and collapse to the ground (Chorley, Schumm & Sugden, 1985).
The fall of snow to the ground is facilitated by the weight of the overlying mass of snow and the process of partial melting, which makes the snow damp and causes it to settle (Menzies, 2002; Chorley, Schumm & Sugden, 1985). The preceding process in diagenesis is consolidation or densification of the large mass of snow that falls on various rock surfaces over a certain period of time (Menzies, 2002). The process of densification is facilitated by dry conditions (Chorley, Schumm & Sugden, 1985). Under the dry conditions, there is increased packing of the snow and eventual rounding or sintering of tiny ice particles (Menzies, 2002).
Sintering involves a continuous process of fusion and pressing out of air from the snow particles due to the compression caused by further accumulation of snow (Chorley, Schumm & Sugden, 1985). The rounding process reduces the permeability of the tiny ice particles, and this eventually leads to recrystallization (Chorley, Schumm & Sugden, 1985; Menzies, 2002). The ice droplets that form from snow are affected by diurnal and annual temperature changes that lead to alternating freeze-thaw actions.
The freeze-thaw processes are referred to as nivation and they lead to the conversion of snow into ice crystals (Menzies, 2002). The crystallization process can also be facilitated by effects of winds that drift the snow particles towards one side (Menzies, 2002). For instance, snow has a higher propensity to accumulate on the northeastern slopes of highland areas because winds usually sweep the particles of snow near the top of the peaks and causes them to settle on the leeward side of the highland (Chorley, Schumm & Sugden, 1985; Menzies, 2002).
The process of transformation of snow into firn can take a long period of time, even over one hundred years (Menzies, 2002; Chorley, Schumm & Sugden, 1985). But in cases where surface melting and percolation occur rapidly, the process may be fast because the melt water that accumulates from the snow on the surface further percolates into the snow and solidifies to form ice at the bottom (Menzies, 2002). This reduces the porosity of the snow and consequently leads to reduction ion size of the ice grains and accelerates the process of transformation (Chorley, Schumm & Sugden, 1985).
The melt water at the bottom of snow or ice undergoes a series of regelation processes, which are much similar to nivation (Menzies, 2002). Nivation involves a series of freezing and thawing processes that are induced by changes in temperature and weight of the accumulated mass of snow (Menzies, 2002). Thus if there is more freezing than thawing, more ice is likely to be formed in one location. On the other hand, regions that are characterized by frequent thawing are likely to have thin glacier because the melt water spreads over a relatively large area and freezes thereafter (Menzies, 2002).
Glacier ice is generally a polycrystalline substance that is impermeable to air at it microscopic level but contains air bubbles and other tiny inclusions (Chorley, Schumm & Sugden, 1985; Menzies, 2002). It is characterized by an arrangement of layers, each of which represents a layer formed in a given period (Menzies, 2002). Factors that contribute to the temperature of glacier ice; difference between warm (temperate) and cold (polar/sub polar) glaciers The temperature of glacier ice is affected by heat emanating from three major sources (Magnu?
sson et al, 2007; Menzies, 2002). One is surface heat that is generated from sources such as the sun as well as any other form of heat emanating from the atmosphere, which has a direct influence on the temperature of the glacier ice (Menzies, 2002). The second source of heat is the base of the glacier ice, which include the rock surface on which the glacier rests (Magnu? sson et al, 2007; Chorley, Schumm, & Sugden, 1985). It is noteworthy that the rock base is affected by factors such as geothermal heat generated from the layers below (Chorley, Schumm, & Sugden, 1985).
The geothermal heat per se can melt approximately a 6mm layer of ice in a year, thus causing a corresponding change in temperature (Menzies, 2002). The third source of heat for the glacier ice is internal friction (Menzies, 2002). This occurs as a result of differential movement of ice masses within a glacier and at the base of the glacier (Menzies, 2002). It is estimated that if the movement occurs at a rate of 20m of ice per year, it can generate an amount of heat equal that is produced by geothermal heating from rocks (Chorley, Schumm, & Sugden, 1985).
In addition to the above differences, the temperature of ice glacier ice is affected by the thickness of the glacial mass (Menzies, 2002). Hence, smaller amounts of glaciers are more likely to be warmed more than thick masses of glaciers should any variations in amount of heat occur (Menzies, 2002). On the other hand, thicker masses of glacier are likely to be colder than thinner ones due to controlled regelation processes (Menzies, 2002).
As a result of the difference in sources of heat for glacier ice, there is a significant difference in the amount of heat retained by glaciers ice, hence, glaciers can be classified as warm glaciers(mostly found in temperate areas ) cold glaciers (found in polar and sub polar areas) (Chorley, Schumm & Sugden, 1985; Menzies, 2002; Magnu? sson et al, 2007). The behavior of glacier ice is usually connected with the glacier ice’s temperature (Ramsay & Huber, 1983). Thus, a glacier whose temperature is below the conventional pressure melting point is referred to as cold or polar glacier ice (Ramsay & Huber, 1983).
In contradistinction, glacier ice that has a temperature sufficiently close to the melting point which leads to the formation of melt water is referred to as warm or temperate glacier ice (Chorley, Schumm & Sugden, 1985; Menzies, 2002; Magnu? sson et al, 2007). Warm glaciers are formed whenever there is an amount of heat that can suffice raising of ice temperature to the conventional pressure melting point (Menzies, 2002). This is commonly experienced when firn on rock surfaces melt during summer (Menzies, 2002). In some instances, the melting may be so pronounced that the whole glacier ice occurs as melt water (Menzies, 2002).
Warm ice may also occur at the base of glacier whose other surfaces comprise cold ice (Magnu? sson et al, 2007). The formation of cold glaciers occurs in two main instances. The first one involves the accumulation of firn in environments that are so cold that there is very little melting during summer (Menzies, 2002). Such conditions are prevalent in the Antarctic region, the Greenland and several mountains in Europe, mainly the Alps (Magnu? sson et al, 2007). When no melting occurs, the firn usually accumulates at temperature that relates to the mean annual temperature (Menzies, 2002; Magnu?
sson et al, 2007). Hence, Polar Regions such as the Antarctica are characterized by very low temperature in the range of -30 to 60? C (Chorley; Schumm & Sugden, 1985). The second situation that results in the formation of cold glacier ice involves cooling of the surfaces of glaciers by the winter cold (Chorley; Schumm & Sugden, 1985). This characterizes all glaciers during the winter season (Menzies, 2002). Processes that contribute to the motion of ice Movement of ice is caused by a number of factors and involves different processes as described below. Internal deformation
Under this phenomenon, ice usually collapses under the influence of its own weight due to gravity (Chorley; Schumm & Sugden, 1985). As this occurs, the mass of ice crystals contained in the glacier gradually changes shape but does not completely break or melt, which causes it to flow along any slope (Chorley; Schumm & Sugden, 1985). According to Menzies (2002), thicker and warmer glaciers are likely to flow faster than colder and thinner glaciers. Although internal deformation is a significant cause of glacial movement, the process generally very slow, in the range of only few meters a year (Chorley; Schumm & Sugden, 1985; Magnu?
sson et al, 2007; Menzies, 2002). Basal sliding Basal sliding is a phenomenon that occurs when water accumulates at the base of a glacier (Menzies, 2002). The water is usually formed at the bottom or close to the base a glacier because of the pressure exerted by the overlying mass of ice (Chorley; Schumm & Sugden, 1985). The melt water at the base of the glacier causes a reduction in friction between ice particles and between ice and the rock surface (Menzies, 2002). This ultimately causes the ice mass to move at a faster rate (Chorley; Schumm & Sugden, 1985). Soft sediment deformation
While water is the ubiquitous agent involved in facilitating sliding of glacier ice, it is not the only cause of sliding (Chorley; Schumm & Sugden, 1985). The other materials associated with the rocks and glaciers too play significant roles in the movement of ice (Menzies, 2002; Magnu? sson et al, 2007). For instance, rock debris held under ice can increase the movement of the glacier (Menzies, 2002) unlike when the rock particles are tightly packed. In addition, if a glacier rests on a loose surface, the surface is likely to collapse and cause rapid movement of the ice (Chorley; Schumm & Sugden, 1985; Magnu?
sson et al, 2007) The flow field of glaciers is estimated by the three dimensional flow field that combines the interferometric radar line-of-sight (LOS) velocities in the ascending format (va ) with the descending format (vd ), as well as the equation of mass continuity (Magnusson et al, 2007). Thus, the velocities of the vectors vn, ve and vu (where n, e and u represent north, east and up) can be derived from the following equations (Magnu’sson et al, 2007): vecos Фa sin ? a + vn sin Фa sin ? a- vucos ? a= va……………………(1) vecos Фd sin ? d + vn sin Фdsin ? d- vucos ? d= vd……………………………
(2) ……………………… (3), where S represents the surface elevation of the hill surface in which the glacier is located, H refers to the thickness of the glacier, and F represents the ratio of average velocity in any vertical column of ice to the velocity of the ice glacier at the surface (Magnusson et al, 2007). In areas that have isothermal glaciers such as the Vatnajo? kull in Iceland, the value of F is always expected to fall in the range between 0. 8 and 1, where the lower value represents deformation alone whereas the upper value corresponds to the effect of sliding alone (Magnusson et al, 2007).
Glacier mass balance Glacier balance is the difference between the amount of accumulation of glacier per year and the glacier’s level of ablation (Addison, 2002). Thus, it is a measure of the difference between the amount of glacier that accumulates during winter and the amount that is removed by melting and ablation during summer (Menzies, 2002). If the rate of accumulation is higher than the rate of removal, then there a positive mass balance because the volume is increased; but if removal exceeds accumulation then the mass balance is negative because of the decreased volume (Addison, 2002).
Mass gain or mass loss is affected by accumulation and ablation zones. On each glacier, a zone of accumulation occurs on the area just above the firn (Menzies, 2002). This is the zone where snow accumulates progressively until it exceeds any losses that arise due to ablation and other phenomena such as evaporation, melting, and sublimation (Menzies, 2002). The accumulation zone can also be identified as a surface of a glacier generally locates on higher slopes on which the net buildup of snow occurs and consequently changes into firn and later into glacier ice (Addison, 2002).
On the other hand, an ablation zone is an area of wastage on a glacier where the yearly loss of ice or snow through various mechanisms exceeds yearly gain of the same materials and accumulation on the surface (Addison, 2002). The ablation zone is usually characterized by melt water that progressively carries away materials contained in the glacier (Menzies, 2002). Processes that influence glacier mass balance Mass balance is affected by a number of processes such as evaporation, melting, iceberg calving and sublimation (Chorley; Schumm & Sugden, 1985).
Evaporation leads to a reduction in the amount of water available for processes such as ice flow (Chorley; Schumm & Sugden, 1985). Melting is a fundamental process that is responsible for most of the ablation that occurs within glaciers (Chorley; Schumm & Sugden, 1985). To begin with, melting is responsible for the melt water that usually reduces friction between ice particles or between ice particles and the rock surfaces bearing ice (Addison, 2002). This leads to mass flows from ablation zones to lower accumulation zones.
Thus, as melt water moves from one ablation zone, it carries with it suspended material and refreezes in other zones lower than the initial point (Chorley; Schumm & Sugden, 1985). On the other hand, sublimation also influences mass balance albeit to a small extent due to rapid changes of substances in glaciers from one physical state to another (Menzies, 2002). Relationship between glacier mass balance and rate of ice motion and glacial erosion There is a critical relationship between mass balance and ice motion as well as glacial erosion (Nagle & Spencer, 2002).
To begin with, the net amount of glacier retained in a particular zone determines the amount of ice that is available for melting (Menzies, 2002). Secondly, the weight of the glacier also determines the possibility of occurrence of processes such as internal deformation, basal sliding and soft sediment deformation (Menzies, 2002). All these factors have a direct impact on the rate at which glacier ice moves. A crucial factor of mass balance is that the size of glaciers usually determines the rate of erosion through processes such as plucking (Menzies, 2002).
The larger the size of the glacier higher the intensity of erosion it can facilitate and vice versa (Menzies, 2002). In addition, the amount of glacier determines the melt water that can be produced to enhance ablation and deformation (Menzies, 2002). Factors that influence processes of glacial abrasion and plucking Glacial abrasion occurs when a bedrock beneath a glacier or mass of ice is eroded by the materials that are embedded in the glacier as well as on the sides of the glacier (Nagle & Spencer, 2002). Several factors affect the level of bedrock erosion that can be done through abrasion.
One is the instance where basal sliding occurs (Nagle & Spencer, 2002). Along this line, it is noteworthy that cold-based glaciers are not able to abrade because they completely frozen and have no possibility of moving (Nagle & Spencer, 2002). Nevertheless, glaciers that have a lot of melt water have a high possibility of moving, hence abrading the surface of the bedrock with their embedded materials (Menzies, 2002). The amount of debris available for the abrading process is also significant in determining the level of abrasion (Menzies, 2002).
This also implies that if there is a lot of debris at the base of a glacier, more abrasion is likely to occur. Of equal importance is the fact that the velocity at which a glacier moves across a rock surface influences the level of erosion (Nagle & Spencer, 2002). Thus, steep zones are likely to experience more abrasion than flat zones (Menzies, 2002). Moreover, the amount of glacier plays an important role in determining the rate of abrasion since it determines the level of pressure exerted to produce melt water as well as the amount pressure with which the debris can abrade a rock surface (Menzies, 2002).
Plucking or quarrying involves removal of much larger debris than is done in abrasion (Nagle & Spencer, 2002). The process is most affected by processes of freezing and thawing since the melt water produced plays a pivotal role in the transportation of debris (Nagle & Spencer, 2002). Thus, is addition to the factors involved in abrasion, the nature of joints in the bedrock determines the level of plucking (Menzies, 2002). Contributions of abrasion and plucking in the creation of erosional landforms
Abrasion and plucking processes are significant in the creation of a number of erosional landforms such as cirques, crag and tail, roche moutonnees, pyramidal peaks and aretes among other features (Nagle & Spencer, 2002). Cirques are formed on the north- or east- sides of mountains where ablation is most experienced as a result of abrasion and plucking (Nagle & Spencer, 2002). The crag and tail is formed when a large resistant object obscures the flow of ice due to abrasion or plucking (Menzies, 2002). Material in the lee of the obstruction is usually protected from abrasion or plucking by a crag, thus forming a tail (Nagle & Spencer, 2002).
Another important feature formed by abrasion and plucking is the roche moutonnee (Nagle & Spencer, 2002). This landform is characterized by the notable effects of plucking and abrasion in that it has a smooth side and a rugged side, both which are shaped by the two processes. The smooth side is facilitated by abrasion while the rugged side down a valley of glacier is created by plucking as the ice moves (Nagle & Spencer, 2002). References Addison, K. (2002). Fundamentals of the Physical Environment: Third Edition. London: Routledge Chorley, R. J. ; Schumm, S A. & Sugden, D. E. (1985). Geomorphology.
London: Routledge Magnu? sson, E. ; Rott, H. ; Bjo? rnsson, H. ; & Pa? lsson, F. (2007). The impact of jo? kulhlaups on basal sliding observed by SAR interferometry on Vatnajo? kull, Iceland. Journal of Glaciology, 53 (181): 232-240 Menzies, J. (2002). Modern and past glacial environments: a revised student edition 2ndedition. New York: Butterworth-Heinemann Nagle, G & Spencer, K. (2002). Advanced Geography through diagrams: AS & A Level Geography through diagrams. Oxford: Oxford University Press Ramsay, J. G. & Huber, M I. (1983). The techniques of modern structural geology. New York: Academic PressSample Essay of EduBirdie.com