Public domain
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We have seen that the atmosphere, the ground-water, and the waters on the surface of the land bring about important changes in its configuration. Ice in its various forms is still another agent of change.
Ice beneath the surface. The wedge- work of ice in the crevices of rock has already been mentioned (p. 103). When the great areas where water freezes during some part of the year are considered, it is clear that the aggregate effect of the freezing of water in the pores and crevices of rock must be great in long periods of time. The water which freezes in the soil also has some effect on the surface. This is shown by the disturbance of the walls of buildings if their foundations do not go below the depth of freezing, and by the working up of stones and bowlders through the soil of the fields, as freezing and thawing succeed each other. The frozen water in the soil makes it solid, and temporarily retards or prevents surface erosion.
Snow is more wide-spread than any other form of ice, but the ice of lakes and rivers is so wide-spread as to be well known. Ice on high mountains and in high latitudes is less familiar.
Ice on lakes and ponds. Since fresh water is densest at 39° Fahr., ice does not commonly form on the surface of a lake until the temperature from top to bottom is reduced to this point. Cooled below 39°, the surface water fails to sink, and reduced to 32°, it freezes. If the lake is small and shallow, it is likely to freeze over completely in any region where the temperature is notably below 32° for any considerable period of time. It is under these circumstances that the lake ice becomes most effective.
Let us suppose a lake in temperate latitudes, where the range of winter temperature is considerable, to be frozen over when the [p. 224] temperature is 20° Fahr. If now the temperature is lowered to -10°, and such a change of temperature is not uncommon in the northern part of the United States, the ice contracts notably. In contracting, it either pulls away from the shores, or cracks. If the former, the water from which the ice is withdrawn quickly freezes; if the latter, water rises in the cracks and freezes there. In either cases, the ice-cover of the lake is again complete. If the temperature now rises to 20° the ice expands, and the solid cover becomes too large for the lake, and must either crowd up on the shores or arch up (wrinkle) elsewhere.
If the water near the shore is very shallow, the ice freezes to the sand, gravel, and bowlders at the bottom. If the land at the shore is very low, the ice in expanding may shove up over it, carrying the debris frozen in its bottom. It may even push up loose gravel and sand in front of its edge if they are present on the shore. Where bowlders are frozen to the bottom of the ice, the shore- ward thrust in expanding has the effect of shifting them in the same direction, and even of lifting them a little above the normal water level. This constant process of concentrating bowlders at the shore-line gives rise to the “walled” lakes (Fig. 182), which are not uncommon in the northern part of the United States. The “wall” does not commonly extend entirely around a lake. In making the walls, the ice driven up by the waves, especially in the spring when the ice is breaking up, co-operates.
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If a lake is bordered by a low marsh, the ice and frozen earth of the latter are really continuous with the ice of the lake, and the push of the latter sometimes arches up the former into distinct anticlines, the frozen part only being involved in the deformation (Fig. 183). A succession of colder and less cold periods may give rise to a succession of such anticlines.1 If the shore is steep, the crowding of the ice against a low cliff of yielding material, such as clay, disturbs all above the shore-line (Fig. 184). Where the cliff is sufficiently resistant, it withstands the push of the ice, and the ice itself is warped and broken.
Ice on rivers. Rivers also freeze over in cold climates, and when the ice breaks up in the spring, the stones and bowlders to which it was frozen in the banks are sometimes floated miles down the river. At Montreal stone buildings 30 to 50 feet square, projecting so as to have river ice form about them, have been moved by the ice of the St. Lawrence.
When the river ice breaks up, masses of it are carried downstream, and may accumulate in vast fields or “jams” behind obstructions in the river. Where a jam forms above a bridge, the bridge may be swept away. The jams also occasion disastrous floods above their sites, and when they break, the waters accumulated above may sweep down the valleys with destructive violence.
Northward-flowing rivers in the northern continents are pecially subject to such floods. The snows of their upper basins often melt while the lower parts of the streams are still frozen over. The free discharge of the upper waters is thus prevented for a time, and freshets follow later.
Ice on the sea. In high latitudes, ice is formed along the shore. Unlike fresh water, sea-water condenses until it free; [p. 227] at a temperature of 26° to 28° Fahr., the variation being due to the degree of salinity of the water.
In polar regions the sea ice attains a depth of several feet, at least as much as eight or ten. Floating ice of much greater thickness is sometimes seen, but it is doubtful if these great thicknesses represent the ice formed by the freezing of undisturbed sea-water. At any rate, the ice formed in winter is often broken up in summer into floating pieces called floe-ice (Fig. 185). Floe-ice is sometimes crowded together in ice-packs, the separate pieces being so jammed together that some of them are ended up and stand high above the water. If the ice-pack of one summer is still far enough north at the end of the warm season, it is frozen together, and its aggregate thickness, made up as it is of blocks of ice some of which are on edge, is far beyond that of normal sea ice.
Snow-fields. Over the larger part of the land, the snow of winter does not endure through the summer, and when it melts it follows the same course as rain; but in cold regions where the fall of snow is heavy, some of it remains unmelted from year to year, and constitutes perennial snow-fields.
High mountains and the lands of high latitudes are the common habitats of snow-fields. In North America there are numerous [p. 228] small snow-fields in the western mountains, from Mexico on the south to Alaska on the north, their number and size increasing in the latter direction. In the United States there are few snowfields south of the parallel of 36° 30’, and most of the many hundreds north of that latitude (excluding Alaska) are small. Farther north, the snow-fields of the western mountains attain greater size. Snow-fields comparable to those of the northwestern part of the United States and British Columbia occur in the higher mountains of Europe and Asia, while in South America there are snow-fields
of small size even in equatorial latitudes, and in the Chilean Andes there are some of considerable size. Small snow-fields occur on the highest peaks of tropical Africa, and in the mountains of New Zealand. For reasons which will appear later, much of every large snow-field is really ice.
Besides these fields of snow in mountain regions, there are fields of much greater extent in polar regions. The greater part of Greenland is covered with a single field of ice and snow, the size of which is estimated at 300,000 to 400,000 square miles (Fig. 186), — an area 400 to 600 times as large as the snow-and-ice-covered area of Switzerland. Numerous islands to the west of North Greenland are also partly covered with snow. In Antarctica there is a still larger field, the largest of the earth. Its area is not even approximately known, but such data as are at hand indicate that it may have an extent ten times as great as that of Greenland.
The only condition necessary for a snow-field is an excess of snow-fall over snow-waste. The lower edge of a snow-field, the [p. 229] snow-line, is dependent chiefly on temperature and snow-fall. In general it does not depart much from the summer isotherm of 32°, though it may be above this isotherm where the snow-fall is light. That the snow-line is not a function of temperature only, is shown by its position in various places. Thus in the equatorial portion of the Andes, the snow-line has an altitude of about 16,000 feet on the east side of the mountains, where the precipitation is heavier, and of about 18,500 feet on the west side, where it is lighter. For the same reason the snow-line in the Himalayas is lower on the south side than on the north. Though temperature and snow-fall are the most important factors controlling the position of the snow-line, both humidity and the movements of the air are of some importance, since both affect the rate of evaporation of snow and ice.
The passage of snow into neve and ice. The snow does not lie on the surface long before it undergoes obvious change. The light flakes are transformed into granules, and the snow becomes “coarsegrained.” The granular character, so pronounced in the last banks of snow in the spring, is even more distinct in perennial snow-fields. This granular snow is called neve. Where the thickness of the snow is great, the neve becomes more compact below, and finally grades into porous ice. Ice is found in some snow-fields at no great depth from the surface.
Structure of the ice. Ice formed beneath a snow-field is in some sense stratified. It is made up of successive falls of snow which tend to retain their individuality to some extent. Thus the snow of one season, or of one period of precipitation, may have been considerably changed before the succeeding fall of snow. Again, the surface of the snow-field at the end of the melting season is often covered with a visible amount of earthy matter, some of which was blown up on the surface during the melting season, and some of which was concentrated at the surface by the melting of the snow in which it was originally imbedded. The amount of earthy matter is often sufficient to define snows of successive years, and makes distinct the stratification which would otherwise remain obscure.
In addition to its stratification, the ice of the deeper portions [p. 230] often acquires a stratiform structure which may perhaps best be called foliation, to distinguish it from the stratification which arises from deposition. The foliation appears to be akin to slaty cleavage or to schistosity, and to result largely from the shearing of one part over another in the course of the movements to which the ice is subjected, as will be illustrated presently.
Texture. The ice derived from the snow is formed of interlocking crystalline grains. The crystalline character is assume! I by the snow-flakes when they form, and the subsequent changes which the snow undergoes seem only to modify the original crystals by building up some and destroying others. By the time the snow is converted into neve, the granules have become coarse, and wherever the ice derived from the neve has been examined, the granular crystalline texture is present. The individual crystals in the ice are usually larger than those of the neve, and more closely grown together. In compact ice the crystals are so intimately interlocked that they are not readily seen by the eye; but when the ice has been honeycombed by partial melting, the granules become partially separated and may be easily seen. While a given mass of snow in a great snow-and-ice-field cannot be followed uninterruptedly through its whole history, yet since the granular texture is pronounced in the neve stage where the granules show eviden of growth, and since the same texture is also pronounced in the last stages of the ice when it is undergoing dissolution, as well as at all observed intermediate stages, it is legitimate to assume that a granular crystalline condition persists throughout all stages of the history of the ice, and that it is a feature of progressive growth.
Inauguration of movement. When the snow and ice in a snowfield become very deep, motion is developed. The ex ad nature of the motion has not been demonstrated to the satisfaction of all who have studied the problem, though much is known about it. Brittle and resistant as ice seems, it may, under proper conditions. be made to exhibit some of the characteristics of a plastic subst a nee. A piece of ice may be made to change its form, and may even be moulded into almost any desired shape if carefully subjected to sufficient pressure, steadily applied through long intervals of [p. 231] time.[1] These changes may be brought about without visible fracture, and have been thought to point to a viscous condition of the ice. There is much reason, however, to question this interpretation. Whatever the real nature of the movement, the aggregate result of the movement in a field of ice is comparable, in a superficial way at least, to that which would be brought about if the ice were capable of moving like a viscous liquid, the motion taking place with extreme slowness. This slow motion of ice in an ice-field is glacier motion, and ice thus moving is glacier ice.
Types. The different shapes of glaciers have given rise to different names. If the surface on which the ice-sheet develops is plane, the ice will move outward in all directions, and ice spreading in all directions from a center is an ice-cap. The glacier covering the larger part of Greenland (Fig. 186) is a good example. The glaciers on some of the flattopped peninsular promontories of the same island are good examples of small ice-caps (Fig. 187) . If ice-caps cover a large part of a continent, as some of those of the past have done, they are called continental glaciers.
Where ice-caps lie on plateaus whose borders are trenched by valleys, ice-tongues from the edge of the ice-cap may extend down the valleys, and constitute one type of valley glacier. A second and more familiar type of valley glacier occupies mountain [p. 232] valleys, and is the offspring of mountain snow-fields. The former type, confined chiefly to high latitudes, are polar or high-latitude glaciers (Fig. 189); the latter are alpine glaciers (Figs. 190-192). The end and side slopes of high-latitude glaciers are, as a rule, much steeper than those of alpine glaciers.
When a valley glacier descends through its valley to the plain beyond, its end spreads. The deploying ends of adjacent glaciers sometimes merge, and the resulting body of ice constitutes a piedmont glacier (Fig. 193). Piedmont glaciers are confined to high latitudes. In some cases the snow-field that gives rise to a glacier is restricted to a relatively small depression in the side of a mountain, or in the escarpment of a plateau. In such cases the snow-field and glacier are hardly distinguishable, and the latter descends but little below the snow-line. Such a glacier, nestled in the face of a cliff, has been called a cliff glacier[2] (Fig. 194). Cliff glaciers often as wide as long, and are always small, and between them and valleys glaciers there are all gradations (Fig. 105). Occasionally the end of a valley glacier, or the edge of an ice-sheet, reaches a precipitous cliff, and the end or edge of the ice breaks off and accumulates like talus below. The fragments of ice may then become a coherent mass by regelation, and the whole may resume motion. Such a glacier is called a reconstructed glacier. The precipitous cliffs of the Greenland coast furnish illustrations.
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Of the foregoing types of glaciers, ice-caps far exceed all others both in size and importance, while valley glaciers outrank the remaining types; but since the valley glaciers are the most familiar, the general phenomena of glaciers will be discussed with primary reference to them.
General Phenomena of Glaciers[3]
Dimensions. Valley glaciers sometimes occupy only the upper parts of mountain valleys, sometimes extend through them, and sometimes push out on the plain beyond. In length they range from a fraction of a mile to many miles. Their thickness is usually measured by scores or hundreds of feet rather than by denominations of a larger order, but the variation is great, and exact measurements are almost wholly wanting. The minimum thickness is that necessary to cause movement, and this varies with the slope, the temperature, and other conditions. There is also much variation in the thickness in different parts of the same glacier. As a rule, it is thinnest in its terminal portion, and thickest at some point between its terminus and its source. Cliff and reconstructed glaciers are comparable in size to the smaller valley glaciers. Piedmont glaciers may attain greater size. An ice-cap is thickest, theoretically, at its center, and thins away to its borders, but its actual dimensions are influenced by the topography of the surface on which it is developed. The Greenland ice-cap rises about 9,000 feet above the sea toward its southern end, and it probably rises higher in the unexplored center of the broader part of the island. The height of the rock surface beneath the ice is unknown, but it is unlikely that it averages half this amount, and hence the ice is probably very thick in the center. The ice cap of Antarctica appears to rest on high land; its thickness is unknown.
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Limits. The ice of a glacier is always moving forward, but the end of a glacier may be retreating, advancing, or remaining stationary, according as the wastage exceeds, falls short of, or equals the forward movement. The position of the lower end of a glacier is therefore determined by the ratio of movement to wastage. Its upper end is generally ill-defined. In a superficial sense, it is the point where the ice emerges from the snow-field; but the lower limit of the snow-field is often ill-defined, and in any case is not the true upper limit of the glacier. The snow-field is really an icefield, covered with snow, and there is movement from it to the tongue of ice in the valley. The ice so moving is, in reality, a part of the glacier.
The lower end of a glacier is usually free from snow and névé in summer, but toward its upper end it is covered with neve, then with snow, and finally merges into the snow-field without having ceased to be a glacier. The term glacier is, however, commonly [p. 238] used to mean merely the more solid portion outside (below) the snow-field.
Movement. The advance of a glacier is too slow, as a rule, to be seen from day to day, and must be detected in other ways. If its end advances, it may override or overturn objects in front of it, or it may move out over ground previously unoccupied. But even when the end of a glacier is not advancing, the movement of the ice may be established by means of stakes or other marks set on its surface. If the positions of these marks, relative to fixed points on the sides of the valley, is determined, they are found after a time to have moved down the valley. Rows of stakes or lines of stones set across a glacier in its upper, middle and lower portions have revealed many facts concerning the movement of the ice.
Generally speaking, the middle of a valley glacier moves more rapidly than its sides, and its top more rapidly than its bottom. In Switzerland, where the glaciers have been studied more carefully than elsewhere, the determined rates of movement range from one or two inches to four feet or more per day. Some of the larger glaciers in other regions move more rapidly, but it does not follow that large glaciers always move faster than small ones. The Muir glacier of Alaska has been found to move some seven feet per day,[4] and some of the glaciers of Greenland have been found to move, in the summer time, 50 or 60 feet per day; but these rates have been observed only where the ice of a large inland area crowds down into a comparatively narrow fiord, and debouches into the sea, and then only in the summer. In the case of the glacier with the highest recorded rate of summer movement, 100 feet per day, the advance was only 34 feet at about the same place in April. The average movement of the border of the inland ice of Greenland is very small, probably less than a foot a week.
Conditions affecting rate of movement. The rate of glaciex movement appears to depend on (1) the depth of the moving ice, (2) the slope of the surf ace over which it moves, (3) the slope of the upper surface of the ice, (4) the topography of its bed, (5) temperature, and (6) the amount of water in the ice. Great thickness, [p. 239] steep slopes, smoothness of bed, a high (for ice) temperature, and abundance of water favor rapid movement. Since some of these conditions, notably temperature and amount of water, vary with the season, the rate of movement of a glacier varies during the year. Other conditions, especially the first of those mentioned above, vary through longer periods of time, and cause corresponding variations in the rate of movement.
A declining upper surface is essential to glacier motion. There are short stretches where this is not the case, and indeed there are occasional places where the upper surface slopes backward, as where the ice is pushed up over a swell in its bed; but such cases are local exceptions, and do not militate against the general truth of the statement that the upper surface of a glacier declines in the direction of motion. A declining lower surface is less necessary. In the case of valley glaciers, the bed does, as a rule, decline in the direction of motion; but that there are exceptions is shown by the deep basins in rock which such glaciers often leave behind them when they retreat. In the great continental glaciers of recent geologic times, the ice frequently moved up slopes for scores, and even hundreds of miles; but in all such cases, the prevailing slope of the upper surface must have been down in the direction of movement.
Fluctuations of glaciers. Observation has shown that the lower ends of glaciers advance and retreat at intervals,[5] and that the periods of advance follow a succession of years when the snowfall has been heavy and the temperature low, while the periods of retreat follow a succession of years when the snowfall has been light and the temperature above the average. The periods of advance and retreat lag behind the periods of heavy and light snowfall respectively, by some years, and a long glacier responds less promptly than a short one.
Likenesses and unlikenesses of glaciers and rivers. Slope, roughness of bed, and volume affect the movement of glaciers somewhat as they affect the movement of rivers. The temperature of [p. 240] water, on the other hand, has little effect on the flow of a river so long as it remains unfrozen; but the effect of temperature on the motion of ice is important. In many cases ,indeed, the temperature, together with the water that is incidental to it, seems to be the chief factor in determining the rate of movement. The way in which its effects are felt will be discussed later.
From Fig. 196 it will be seen that a valley glacier is an elongate body of ice, following the curves of the valley in streamlike fashion. It has its origin in the snows collected on the mountain heights, and it works its way down the valley in a manner which, in the aggregate, is similar to the movement of a stiff liquid. The likeness to a river extends to many details. Not only does the center move faster than the sides, and the upper part faster than the bottom, as in the case of streams, but the movement is more rapid in narrow parts of the valley and slower in the broader parts. These and other likenesses, some of which are apparent rather than real, have given rise to the view that glacier ice moves like a stiff viscous liquid.
But while the points of likeness between glaciers and rivers arc several, their differences are at least equally numerous and significant. Though the trains of debris on the surface (the dark bands in Fig. 196) pass nearer the projecting points of the valley walls and farther from the receding bends like the central currents of streams, they do not conform in detail to the course of the winding [p. 241] valley, nor is there evidence of rotatory motion, as in the water of a river. Furthermore, the glacier is readily fractured, as the numerous gaping crevices on many glaciers show. The crevasses are sometimes longitudinal, sometimes transverse, and sometimes oblique. In the case of arctic glaciers, longitudinal crevassing is especially conspicuous. Crevasses appear to be developed wherever there is appreciable tension, and the causes of tension are many. An obvious cause is an abrupt increase of gradient in the bed (Fig. 198). If the change of gradient is considerable, an ice-fall or cascade results, and the ice may be greatly riven. Transverse crevasses at the margin sometimes appear to be the result of the tension developed on a curve. Oblique crevasses on the surface near the sides are commonly ascribed to the tension between the fastermoving center and the slower-moving margins, and in like manner cracks that rise obliquely from the bottoms are attributed to the tension between the faster-moving parts above and the slowermoving parts below. All these crevasses indicate strains to which a liquid, whose pressures are equal in all directions, does not offer a close analogy. Longitudinal crevasses may affect both the nai row part of a glacier and its deploying end, and are the result of tension developed by movement within the ice itself, to which, again, rivers offer no analogy. All cracks show that the glacier is a very brittle body, incapable of resisting even very moderate strains brought to bear upon it very slowly. In its behavior under tension, therefore, a glacier is notably unlike a river.
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Surface moraines. The surface of a glacier is often affected by rock debris, which is sometimes disposed in the form of belts or moraines (Fig. 199). The surface moraines may be lateral, medial, or terminal. A lateral moraine is any considerable accumulation of debris in a belt on the side of a glacier. A medial moraine is a similar accumulation at some distance from the margins, but not necessarily in or even near the middle. There may be several medial [p. 243] moraines on one glacier, in which case some of them may be far from the center. In valley glaciers, the surface terminal moraine often connects two lateral moraines, making a loop roughly concentric with the end of the glacier.
Besides the surface moraines, there may be scattered bowlders and bits of rock of various sizes on the ice, and, in addition to the coarse material, there is often some dust which has been blown upon the ice.
Relief due to surface debris. The debris on the ice affects its topography by influencing the melting of the ice beneath and about it. The rock debris absorbs heat more readily than the ice. A thin piece of stone lying on the ice is warmed through by the sun’s rays, and, melting the ice beneath, sinks, just as a piece of black [p. 244] cloth on snow will sink because of the increased melting beneath it. Though a good absorber of heat, rock is a poor conductor, and so the lower surface of a thick mass of stone is not warmed notably, and the ice beneath, being protected from the sun, is melted more slowly than that around it. The result is that the bowlder presently stands on a protuberance of ice (Fig. 200). When its pedestal Incomes high, the oblique rays of the sun and the warm air surround* ing it cause it to waste away, and the capping bowlder falls.
The same principles apply to the moraines. A surface moraine usually protects the ice beneath from melting, and causes the development of a ridge of ice beneath itself. As the ice on either side is lowered by ablation, the moraine matter tends to slide down [p. 245] on either hand. So far does this spreading go, that in some cases the lower end of a glacier is completely covered with the debris whici has’ spread from the medial and lateral moraines.
Debris below the surface. The lower part of a glacier, as well as the upper, carries rock debris. This debris is sometimes so abundant, especially near the ends and edges of the ice, that it is difficult to locate the bottom of the glacier; for between the moving ice which is full of debris, and the stationary debris which is full of ice, there seems to be a nearly complete gradation. The debris in the lower part of arctic glaciers, and to some extent of others, is often disposed in thin sheets sandwiched in between layers of clean ice. Debris also occurs to some extent in the ice far above its base, sometimes in sheets and sometimes in bunches. These various relations are illustrated by Figs. 201 and 202.
Temperature
In winter, the surface of the ice becomes as cold as the air above it, and the cold of the surface penetrates downward. This [p. 246] descent of the surface cold may be called the winter wave of low temperature. The degree of cold becomes rapidly less below the surface, and at no great depth the temperature is not much affected by the seasons. In middle latitudes, the zone of no seasonal variation of temperature on land is only 50 or 60 feet below the soil. Its depth in ice is not known, but it is not probable that the cold of winter has much effect below a similar depth. Conduction alone considered, the temperature of the ice at the depth where the cold wave dies out, should correspond, approximately, to the mean annual temperature of the region, provided that temperature is below the melting-point of ice. At this depth, the temperature of the ice in middle Greenland should be about 20° Fahr. — the mean annual temperature of the region.
With the coming of summer, the temperature of the surface of the ice may be raised to 32°, and this relatively high temperature makes itself felt below the surface as the warm wave, if the temperature of ice can ever be said to be warm. The relatively high temperature of the surface makes itself felt below both by conduction and by the descent of water. The former affects the ice at all temperatures, the latter only after the melting-point of the ice at the surface is reached. The first conforms measurably to the warm wave affecting other solid earth-matter, while the second is governed by other laws. After the surface portion of the ice is brought to t he melting temperature, the additional heat which it receives melts ice. and is transformed from sensible into potential heat. Ice charged [p. 247] with water is potentially, but not sensibly, warmer than ice which has just reached the melting temperature. The warm wave of conduction dies out below, like the cold wave. The warm wave descending by the flow of water stops where the freezing temperature of water is reached, except where crevasses extend to greater depths.
The foregoing considerations warrant the generalization that glaciers normally consist of two zones (1) an outer or upper zone of varying temperature, and (2) an under zone of nearly constant temperature. The under zone obviously does not exist where the thickness of the ice is less than the thickness of the zone of fluctuating temperature, as may be the case in very thin glaciers and in the thin ends and edges of all glaciers.
The temperature of the bottom. The internal heat of the earth is slowly conducted to the base of a glacier where it melts the ice at the estimated average rate of about one-fourth of an inch per year. It is probable that in all thick glaciers the temperature of the bottom is constantly at the melting-point. In glaciers or in parts of glaciers so thin as to lie wholly within the zone of fluctuating temperature, the temperature of the bottom is obviously not constant.
Temperature of the interior of the ice. The variation of temperature of the surface of a glacier has already been stated to lie between a maximum of 32° Fahr. and the minimum temperature of the region where the glacier occurs. In the zone of varying temperature, the variation is less and less with increasing depth. In the zone of constant temperature, the range is from the mean annual temperature of the region at the top of the zone (provided jui’s is not above the melting-point of ice at this depth) to the melting temperature of the ice at the bottom. Within these limits the range may be great or slight.
Considering only the effects of the external seasonal temperatures and the internal heat of the earth, it appears that all the ice in the zone of constant temperature in the lower end of a typical alpine glacier should have a melting temperature constantly, for the average annual temperature of regions where the ends of such glaciers occur is usually above 32° Fahr., and this determines a temperature of 32° Fahr., approximately, at the top of the zone, [p. 248] while a melting temperature is maintained at the bottom by the earth’s interior heat. In thin glaciers of very cold regions, the lowtemperature descending from the surface may so far overcome the effect of internal heat as to keep the bottom of the ice at a freezing temperature during the winter season; but in all other cases the ice at the bottom of the under zone probably has a melting temperature, while that above is probably colder, except where the ice is warmed by the descent of surface water.
In the higher altitudes and in the polar latitudes, where glaciers are chiefly generated, the mean annual temperature of the surface is usually below the melting-point of ice. Here the temperature of the ice between the top and bottom of the zone of constant temperature must be below the melting-point, on the average, unless heat enough is generated in the interior of the ice to offset the effect of the temperature above. For example, where the mean annual temperature is 20° Fahr. or lower, as in middle Greenland, the mean temperature in the zone of constant temperature should range from 20° Fahr. at the top, to 32° Fahr. (or a little less) below; i. e., it should average some 6° below the melting-point, disregarding the effect of water descending from the surface. Under these conditions, all the ice in the zone of constant temperature, except that at its bottom, must be permanently below the melting-point; but it is worthy of especial note that much of it is but little below. In alpine glaciers, the part of the ice affected by this constant lowtemperature (below freezing) is presumed to be chiefly that which lies beneath the snow-fields. In polar glaciers, the low temperature probably prevails beneath the surface not only throughout the great ice-caps, but also in the marginal glaciers which descend from them.
From these theoretical considerations we may deduce the generalization that in the zone of constant temperature within the area of glacier growth the temperature of the ice is generally below the melting-point, while within the area of wastage the temperature of the corresponding zone is generally at the melting-point.
Compression and friction as causes of heat. The foregoing conclusions are somewhat modified by dynamic sources of heat. Both the compression arising from gravity, and the friction developed [p. 249] where there is motion, are causes of heat. Since friction occurs only when motion takes place, the heat which it generates is secondary and may, for present purposes, be neglected. But compression produces heat at the point of compression, and also lowers the meltingpoint slightly. If the compression is considerable in ice which is already near the melting temperature, the granules may be warmed to the melting-point wrhere they press each other. In this case melting will take place at the points of compression, and the water so produced will flow to points of less pressure, and be re-frozen immediately. Melting at the points of compression would result in some yielding of the mass, and in some shifting of the pressure to new points, where compression and melting wrould again take place. From considerations already adduced, it appears that the temperature in some parts of every considerable body of ice must be such as to permit these changes. The heat due to compression and friction may modify the theoretical conclusions deduced above.
Summary. If the foregoing generalizations are correct, (1) the surface of a glacier is likely to be melted during the summer; (2) its immediate bottom is slowly melting all the time (unless the thickness of the ice is less than the thickness of the zone of annual variation) ; (3) its sub-surface portion in the zone of waste is generally melting, owing to descending water, compression, and friction; while (4) its sub-surface portion in the zone of growth is probably below the melting-point except as locally brought to that temperature by compression, friction, descending water, and, at the bottom, by conduction from the rock beneath.
Glacier motion is not under discussion at this point, but it may be pointed out that since there must be motion in the area of growth in order to supply the loss in the area of waste, the fundamental cause of motion must be operative in bodies of ice whose mean temperature is below the melting-point, unless the dynamic sources of heat are considerable.
Drainage. Some of the water produced by surface melting forms little streams on the ice. Sooner or later they plunge into crevasses or over the sides and ends of the glacier. In the former case, they may melt or wear out well-like passages (moulins) in the ice, and even in the rock beneath. Much of the surface water sinks [p. 250] into the ice. The depth to which it penetrates is undetermined by observation, but it doubtless goes down to the zone of constant temperature in all cases, and still lower in some cases, as where there are crevasses, and where the temperature is as high as 32° Fahr. all the way to the bottom. At considerable distances above the line of perennial snow there is little water either from melting or from rain, and hence relatively slight penetration. Below the line of perennial snow, and for a short distance above it, there is more melting and rain, and here it is probable that the water often penetrates to the bottom of the ice during the melting season, even independently of crevasses.
Once within the glacier, the course of the water is variable. Exceptionally it follows definite englacial channels, as shown by springs or streams issuing from the ice at some point above its bottom. Oftener it descends or moves forward through the irregular openings which the accidents of motion have developed. If it reaches a level where the temperature is below its freezing-point, it congeals. Otherwise it remains in cavities or descends to the bottom. The water produced by melting within the glacier probably follows a similar course. So far as these waters descend to the bottom, they join those produced by basal melting, and issue from the glacier with them. In alpine glaciers, the waters beneath the ice often unite in a common stream in the axis of the valley, and hollow out a tunnel. Thus the Rhone River is already a considerable stream where it issues from beneath the Rhone Glacier. In the glaciers of high latitudes, subglacial tunnels are less common, and the drainage is in streams along the sides of the glaciers or through the debris beneath and about them.
At the end of the glacier, all waters, whether they have been superglacial, englacial, or subglacial, unite to bear away the silt, sand, gravel, and even small bowlders set free from the ice, and to spread them in belts along the border of the ice, or in trains stretching down the valleys below, forming glacio-fluvial deposits.
[p. 251]
Erosion and Transportation
Glaciers abrade the valleys through which they pass, carry forward the material which they remove from the surface, and wear, grind, and ultimately deposit it. Like other agents of gradation, therefore, their work includes erosion, transportation, and deposition.
Getting load. If the snow-field which is to become a glacier accumulates on a rough surface covered with abundant rock debris, the glacier already has a basal load when it begins to move, for the snow covers, surrounds, and includes such loose blocks of rock as project above the general surface, and envelops all projecting points of rock within its field. Wrhen the ice begins to move, it carries forward this loose rock in its bottom, and tears off the weak points of the enveloped rock-projections. It may also drag along some of the earthy matter of the surface on wrhich the ice forms, and to which the ice is frozen. In addition to the basal and subglacial load which the glacier has at the outset, there may be surface debris which has fallen on the snow or ice from cliffs above. This is likely to be the case where steep cliffs tower above the snow-fields. If debris descending to the glacier in this way is unburied, it is superglacial, but if it has been buried by subsequent falls of snow, it is englacial.
Once in movement, the ice not only moves the debris to which it was originally attached, but it gathers new load. This it acquires partly by the rasping effect of its rock-shod bottom, and partly by its power of plucking off or quarrying out considerable blocks of rock from its sides and bottom. This plucking process is at its best where the ice passes over the cliffs of jointed rock, but is not confined to such situations. The steep bed of a valley glacier is often worn more by plucking than by rasping. The advancing ice gets some material, too, especially loose debris, by freezing to it, for the frozen water in the soil, etc., becomes continuous with the ice above, and the two move together. Superglacial material may be acquired during movement, as well as before it, by the fall of debris from cliffs, or by the descent of avalanches.
[p. 252]
Conditions influencing rate of erosion. (1) A somewhat obat motive configuration of the surface over which the ice moves is necessary for effective erosion. Ice wears a flat surface relatively little since there is little for it to get hold of. Ice sometimes overrides such a surface, burying its soil and even more or less of its herbaceous vegetation. Erosion is probably at its maximum, so far as influenced by topography, when the surface is rough enough to offer notable catchment for the base of the ice, but not so rough as to impede its motion seriously. The amount of roughness favorable for the greatest erosion increases with increasing thickness of the ice.
Other conditions which influence erosion by ice nvv (2) the amount of loose or slightly attached debris on the surface; (3) the resistance of the rock; (4) the slope of the surface; (5) the thickness of the ice; (6) the rate of its movement; and (7) the abundance and character of the debris which it has to work with. The effect of most of these conditions is evident, but the effect of the last La less simple. Clean ice passing over a smooth surface of solid rock would have little effect upon it; but a rock-shod glacier [p. 253] will abrade the same surface notably. The effect of this abrasion is shown in the grooves and scratches (striae) which the stones in the bottom of the ice inflict on the surface of the rock over which they pass (Figs. 203 and 204). At the same time, the stones in the ice are themselves worn both bv abrasion with the bottom, and with one another (Fig. 205). It does not follow, however, that erosion is greatest when there is most material in the bottom of the ice; for with increase of debris there may be decrease of motion,[6] and decrease of motion interferes seriously with the efficiency of erosion. When any considerable thickness of ice at the bottom of a glacier is full of debris, the loaded portion may approach stagnancy, while the cleaner ice above shears over it. A moderate but not excessive load of debris is, therefore, favorable for great erosion. Something depends, too, on the character of the load. Coarse, hard, and angular debris is a more effective instrument of abrasion than fine, soft, or rounded material. So far as plucking is concerned, rate of motion is probably more important than load. Other things being equal, ice which is unyielding, as when full of debris, would be more effective in plucking than ice which is more yielding.
[p. 254]
In connection with the resistance of the rock, it should be noted that resistance is not a matter of hardness and softness simply. Rock which is affected by cleavage planes, whether joints or bedding planes or both, is subject to erosion by plucking, as well as by the abrasion effected by the load in the bottom of the ice. On steep slopes, plucking is probably more important, on the whole, than wear by the debris carried.
So far as concerns the ice itself, erosion is not most effective at the end of a valley glacier, or at the edge of an ice sheet, for here the strength of movement is too slight and the load too great; nor is the most effective erosion at the source or near it, for though the ice here may be thick, the movement is slow and the load likely to be slight. Ice alone considered, erosion is most effective somewhere between the source and the terminus of a glacier, and probably much nearer the latter than the former. The conditions of the surface over which the ice passes may be such as to make the place of greatest erosion vary widely.
Summary. In summary it may be said that rapidly moving ice of sufficient thickness to be working under goodly pressure, shod with a sufficient but not excessive quantity of hard-rock material, passing over non-resistant formations possessing a topography [p. 255] of sufficient relief to offer some resistance, and yet too little to retard the progress of the ice seriously, will erode most effectively.
Varied nature of glacial debris. From its mode of erosion it will be seen that a glacier may be charged with various sorts of material.
At its bottom there may be (1) bowlders which the ice has picked up from the surface, or which it has broken off from projecting points of rock over which it has passed; (2) smaller pieces of rock of the size of cobbles, pebbles, etc., either picked up by the ice from its bed or broken off from larger masses; (3) the fine products (rock-flour) produced by the grinding of the debris in the ice on the rock-bed over which it passes, and similar products resulting from the rubbing of stones in the ice against one another; and (4) sand, clay, soil, vegetation, etc., derived from the surface overridden. Thus the materials which the ice carries (called drift) are of all grades of coarseness and fineness, from huge bowlders to fine clay. The coarser materials may be angular or round at the outset, and their forms may be changed and their surfaces striated as they are moved forward. Whether one sort of material or another predominates depends primarily on the nature of the surface overridden.
[p. 256]
The topographic effects of glacial erosion. In passing through its valley, an alpine glacier deepens it, widens its lower part, and smooths its slopes up to the limit of the ice. It tends to make a V-shaped valley (Fig. 206) U-shaped (Fig. 207). The change in topography at the upper limit of glaciation is often marked (Fig. 208).
The deepening of a valley by glacial erosion may throw its [p. 257] tributaries out of topographic adjustment. Thus if a main valley is lowered 100 feet by glacial erosion while its tributary is not deepened, the lower end of the latter will be 100 feet above the former when the ice disappears. Such a valley is called a hanging valley (Fig. 209). Such valleys are common in regions which were recently glaciated, like the western mountains of North America.
Ice-caps which overspread the surface irrespective of valleys and hills tend to reduce the angularities of the surface. Hills and ridges are cut down and smoothed (Figs. 210 and 211); but since [p. 258] valleys parallel to the direction of movement are deepened at the same time, it is doubtful if the relief of the surface is commonly reduced by the erosion of an ice-cap.
Fiords. A glacier descending into the head of a bay may gouge it out to a very considerable depth, and cause its head to advance into the land. When the ice finally melts, the bay, if narrow, deep, and long, with high slopes, is called a fiord. Many of the fiords of coasts in high latitudes have arisen in this way, and some of the glaciers of these coasts are now making fiords. Fiords arise in other ways also.
The positions in which debris is carried. As a result of the methods by which a glacier gets its load, debris is carried in three positions: (1) the basal or subglacial, (2) the englacial, and (3) the superglacial. The material picked up or rubbed off from the surface over which the ice moves is normally carried forward in the bottom of the ice, and is therefore basal; that which falls on the surface is usually carried on the surface, and is therefore superglacial. Either basal or superglacial drift may become englacial, as wTe shall see.
The basal load of a glacier is constantly being mixed with new accessions derived from ground over which the ice is passing, and this admixture tells the story of the work done by the bottom of the ice. The superglacial material, on the other hand, is normally borne from the place of origin to the place of deposition without such intermixture. It is doubtful if much debris is moved along beneath (that is, strictly below the bottom of) the ice, though the movement of the ice would tend to drag along the loose material of its bed. If drift were carried forward in such position, it would be strictly subglacial.
Transfers of load. While the origin of the load usually determines its position at the outset, exceptions and complications arise from the transfer of load from one position to another, and from the gradation of one horizon to another.
Most of the debris gathered by ice is acquired at its bottom. While such material is basal at the outset, some of it may find itself above the bottom a little later. Thus when ice passes over a hill, the bottom of the ice rends debris from its top. To the lee of [p. 259] the hill the ice from either side may close in under that which came over the top. The debris derived from the top of the hill by the bottom of the overriding ice will then be well up in the ice. It has passed from an initial basal to a subsequent englacial position; but the change does not usually involve an actual rise of the material. If carried upward at all, the upward movement is temporary only, and incident to the passage of the ice up over the hill, or to other local causes.
Superglacial debris may obviously become englacial by falling into crevasses or by being carried down by descending waters, and either superglacial or englacial debris may become basal by the same means. There is less ice-free land in immediate association with ice-caps than with valley glaciers, and the ice-free land about the borders of an ice-cap is less likely to be in the form of cliffs above it. Except at their edges where the ice is thin, the surfaces of ice-caps are comparatively clean, for there is little land which rises above them in such position as to furnish surface debris.
Englacial material may become superglacial by surface ablation. In this case the drift does not rise, but melting brings the surface of the ice down to it. This occurs chiefly at the end or edge of the ice, where the surface melting is greatest. Englacial material plucked or rasped from an elevation over which the ice has passed is liable to be disposed in a longitudinal belt in the ice in the lee of the elevation itself. By surface ablation this material may reach the surface at some point below its source (Fig. 212), and be disposed as a medial moraine. Such a moraine has an origin very different from that of a medial moraine formed by the junction of two lateral moraines of superglacial origin. Englacial debris, especially that near the bottom, may also become basal by the melting of the bottom of the ice.
[p. 260]
Drift is sometimes transferred from a basal to an englacial and then to a superglacial position, by upward movement. Such transfer is the more remarkable because the specific gravity of rock is two and a half to three times that of ice, so that the normal tendency of rock is to sink in ice.
In arctic glaciers, and probably in others, some material which has been basal becomes englacial by being sheared forward over ice in front of it. So far as observed, this takes place chiefly where the ice in front of the plane of shearing lies at a lower level than that behind, as where the surface of an upland falls off into a valley, or where a boss of rock shelters the ice in its lee from the thrust of the overriding ice (Fig. 213).
At the borders of arctic glaciers the lower layers are not infrequently turned up, as shown in Fig. 214. Where the layers turn up at the end of a glacier, basal and englacial debris is carried to the surface by actual upward movement, and a terminal moraine or a series of terminal moraines is sometimes aggregated where the [p. 261] upturned layers of ice outcrop at the surface (Fig. 215). That the material of these moraines was originally basal is abundantly demonstrated in many cases by the bruised and scratched condition of the bowlders and pebbles, and sometimes by the nature of the material itself. The upturning sometimes affects the edges of glaciers (Fig. 216) as well as their ends, and the material thus brought to the surface gives rise to lateral moraines. Sometimes also there is an upturning of the ice along a longitudinal zone well back from the lateral margins (Fig. 216), and the material so borne to the surface in such a zone gives rise to a medial moraine.
[p. 262]
The upturning of ice here referred to has been observed only at or near the terminus of the ice. It is perhaps due in part to the resistance of frozen morainic or other material beneath and in front of the edge. To this should probably be added the effect of the great rigidity of the ice due to the low external temperature during the larger part of the year, while the interior, with its higher temperature, remains more fluent. But even this probably leaves the explanation inadequate.
Wear of drift in transit. Drift carried at the bottom of the ice is much worn, for the materials in transit abrade one another and are abraded by the bed over which they pass. Englacial drift is subject to less wear, because it is commonly more scattered. Superglacial drift is worn little or none while it lies on the surface of the ice; but in so far as superglacial or englacial drift is derived from the basal load, it may show the same evidences of wear as the basal drift itself. Superglacial drift often reveals its history in this way.
During the advance of a glacier, deposition may take place both (1) beneath the body of the ice, and (2) beneath its end and edges.
1. Beneath the body of the ice deposition takes place where the topography favors lodgment, or where the ice is overloaded.
[p. 263]
The topography favoring deposition is much the same as that favoring erosion, but the two processes are not favored at the same points.’ Erosion is greatest on the “stoss” side (the side against which the ice advances) of an obstruction, and deposition on the lee side (Fig. 217). The ice is likely to be overloaded (1) just beyond a place where conditions have favored the gathering of a heavy load, and (2) where the ice is rapidly thinning. On the whole, the deposition of material beneath the main body of a glacier is much more than balanced by erosion in the same position.
2. At and near the end of a glacier, deposition goes on faster than elsewhere, chiefly because of the rapid melting, and therefore the thinning and weakening of the ice. If the end of the glacier is stationary in position, drift is being continually brought to it and left there, for it is to be remembered that the ice continues to move though the end is stationary. If the glacier moves forward 500 feet per year, and if its end is melted at the same rate, all the debris in the 500 feet of ice which is melted is deposited, and all except that which has been washed away has been deposited at and beneath the end of the glacier (Figs. 218-219). If the end of the glacier is retreating, the retreat means that the waste at the end exceeds the forward movement. If the ice advances 500 feet per year and is melted back 600 feet in the same time, all the debris carried by the 600 feet which has been melted has been deposited, and largely in the narrow zone (100 feet) from which the ice has receded. If the end of the glacier is advancing 500 feet per year while it is being melted but 400 feet, all the drift in the 400 feet melted has been deposited, and chiefly at or beneath the immediate margin of the ice. To the marginal and sub-marginal accumulations made in this way, the material carried on the ice is added whenever the ice is melted from beneath it. Deposition beneath the lateral margins of a glacier is much the same as beneath its terminus (Fig. 221).
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[p. 266]
Types of Moraines
The terminal moraine. The thick accumulation of drift made at the end of a glacier or at the edge of an ice sheet, especially where its end or edge is stationary or nearly so for a considerable time, is the terminal moraine. Terminal moraines are of more importance, relatively, in connection with ice-caps than in connection with valley glaciers, for streams are more effective in destroying the moraines of the latter than those of the former.
The ground moraine. When a glacier disappears, all its debris is deposited. All the drift deposited beneath the advancing ice, and all deposited from the base of the ice during its dissolution, constitutes the ground moraine. The thickness of the ground moraine is notably unequal. In general, it is thicker toward the terminus oi the glacier and thinner toward its source, but considerable portions of a glacier’s bed are often left without debris when the ice melts. In general, the ground moraine is thinner than the terminal moraine, and less irregularly disposed. The ground moraine is of relatively slight importance in valley glaciers as compared with ice-caps, since conditions for erosion under the body of a valley glacier are, on the average, better than under an ice-sheet, and those for deposition less favorable.
The lateral moraines. Lateral moraines are the product of \ alley glaciers. The lateral moraines on such glaciers are let down on the surface beneath when the ice melts; but the lateral moraines in a valley from which the ice has melted are not merely the lateral moraines which were on the glacier at a given time. They are often far more massive than any which ever existed on the ice (Figs. 216, 221, and 222). Furthermore, much of the material of lateral moraines left after the ice is gone is as well worn as that of the ground moraine. The massive lateral moraines left after the ice is gone are made up chiefly of the drift accumulated beneath the lateral margins of the valley glaciers. This accumulation ifl the result of the lateral motion of the ice from the center to the side of the valley. Such sub-lateral accumulations are akin to terminal moraines. Some of the lateral moraines of ancient valley glaciers. such as those of the Uinta, Wasatch, and Bighorn mountains are [p. 267] several hundred feet high, or even as much as a thousand. In northern Italy a lateral moraine is said to be about 2,000 feet high.[7] Most of the material which was englacial during its transportation becomes either subglacial or superglacial before deposition, for it reaches the bottom or the top of the ice before being deposited. Only where the ends or edges of a glacier are vertical or nearly so, as in the arctic regions, does deposition take place from the englacial position directly.
Distinctive nature of glacial deposits. The deposits made by glaciers are distinctive. In the first place, the ice does not assort its material, and bowlders, cobbles, pebbles, sand, and clay are confusedly commingled (Fig. 223). In this respect, the deposits of ice differ notably from the deposits of water. Furthermore, many stones of the drift show the peculiar type of wear which glaciers inflict. Though notably worn, they are not rounded like the stones carried by rivers. Many of them have sub-angular forms with planed and beveled faces, the planes being striated and bruised (Fig. 205). Absence of stratification, physical heterogeneity, and the striation of at least a part of the stones are among the most distinctive characteristics of glacial drift. A not less real though [p. 268] less obvious characteristic is the constitution of the fine material, for it is, as a rule, the product of rock grinding, not of rock decay.
Glaciated rock surfaces. Another distinctive mark which a glacier leaves behind it is the character of the surface of the rock on which the drift rests. This is generally smoothed by the severe abrasion to which it has been subjected, and the smoothed surfaces (Figs. 203 and 224) are marked by grooves ami si rise, similar to those on the stones of the drift (Fig. 205). Other distinctive features of a glaciated area are the rounded bosses of rock (roches moutonnées, Fig. 225), the rock basins, the lakes (Fig. 226), ponds, and marshes, and the peculiar topographies resulting from the unequal erosion, and the still more unequal deposition of the drift (Figs. 227 and 228, and Pis. XVI and XVII). Surface bowlders, often unlike the underlying formations of rock, and sometimes in peculiar and apparently unstable positions, are still another mark of a glaciated area (Fig. 229).
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[p. 272]
The constant but unequal waste of glaciers has already been referred to. The streams to which the melting of the ice gives rise are unusually laden with gravel, sand, and silt derived from the ice. Since the mud is often light-colored, the streams are some times described as “milky.” Where the amount of material carried is great, much of it is dropped at a slight distance from the ice, the coarsest being dropped first. Glacial streams are, as a rule, aggrading streams, and therefore develop alluvial plains, called valley trains (Fig. 230), or, where they enter lakes, bays, or other streams, deltas. In its transportation, the river-borne drift is assorted, and after "its deposition it is stratified. True glacial deposits in the upper part of a mountain valley are, therefore, often connected with glacio-fluvial deposits farther down the valley. The silt, sand, and gravel of valley trains can often be distinguished from valley deposits of non-glacial origin by the fact that they are more largely of undecayed rock material, especially if deposited recently.
[p. 273]
From the ice-sheet, numerous streams flow, spreading their debris in front of the terminal moraine, forming a broad fringing sheet or “apron” (outwash plain) along it. Outwash plains have much in common with piedmont alluvial plains. They differ from valley trains chiefly in being shorter, wider, and not confined to a valley. Where streams of considerable size form tunnels under the ice, the tunnels may become more or less filled with water-worn debris, and when the ice melts, the aggraded channels appear as ridges of gravel and sand, known as eskers (Fig. 231). It has been thought that eskers represent deposits formed in superglacial channels; but this is probably rarely if ever the case, for the surface [p. 274] streams usually have high gradients, swift currents, and smooth bottoms, and hence give little opportunity for lodgment. Furthermore, ice-sheets, in connection with which eskers are chiefly developed, usually have no surface material except at the immediate edge where the ice is thin and its layers upturned.
At the mouths of ice-tunnels or ice-channels, especially where they end against terminal moraines, and in the re-entrant angles of the edge of the ice, sands and gravels arc liable to be bunched in quantity, giving rise, after the adjacent ice has melted, to peculiar hills and hollows of the knob-and-baao type. The hills and short ridges of stratified drift formed in this way are known as fames.
[p. 275]
Much stratified drift (gravel, sand, and silt) deposited by glacial streams’ has no distinctive topographic form, and therefore no special name.
When glaciers advance into water the depth of which approaches their thickness, their ends are broken off (Fig. 232), and the detached masses float away as icebergs (Fig. 233) . Many of the bergs are overturned, or at least tilted, as they set sail. If this does not happen at the outset, it is likely to occur later as the result of the melting, wave-cutting, etc., which shift their centers of gravity. The great majority of bergs do not float far before losing all trace of stony and earthy debris; but the finding of glaciated pebbles in dredgings far south of all glaciers shows that bergs occasionally carry stones far from land.
The key to the study of the structure and motion of glacier ice is based on the view that a glacier is a mass of crystalline rock — the purest and simplest type of crystalline rock known. It is made of a single simple mineral, ice, which always has a crystalline structure.
The development of ice from snow. The fundamental conception of a glacier is best developed by tracing the growth of its constituent crystals. When water solidifies from the vapor of the atmosphere, it takes the form of separate crystals (Fig. 234). The flakes are rarely perfect, but they are always crystals. Snow crystals may continue to grow so long as they are in the atmosphere ; but if they pass through a dry stratum of air, or a stratum whose temperature is above 32° Fahr., they suffer from melting and evaporation. When they reach the ground, the processes of growth and decadence continue, and the crystals grow or shrink according to circumstances.
A glacier is a colossal aggregation of crystals grown from snowflakes to granules of greater size and more compact form. The microscopic study of snow reveals the mode of change from flakes to granules. The slender points and angles of new-fallen flakes melt and evaporate more than the central portions. The water [p. 276] melted from the periphery of a flake may gather about its center, and, if the temperature is right, freeze there. Similarly, evaporation from the periphery is probably followed by condensation about the center. These are first steps toward the pronounced granulation so conspicuous in snow which has lain long on the ground. Measured from day to day, the larger granules beneath the surface of coarse-grained snow are found to be growing. When the temperature of the atmosphere is above the melting-point, the growth is faster than when the air is colder, but there is an increase in the average size of the granules, and a decrease in their number, under all conditions of temperature. A part of the increase of the larger granules appears to come from the diminution and destruction of the smaller ones; another part doubtless comes from the moisture of the atmosphere which penetrates the snow and is condensed there, and still another part from the descent of moisture derived from surface melting.
[p. 277]
Deep beneath the surface of a large body of snow, the larger part of the growth of the large granules is probably at the expense of the small ones. To understand how this takes place, it should be noted that the free surface of every granule is constantly throwing off particles of water-vapor (i. e., evaporating); that the rate of evaporation increases with the sharpness of the curve of the surface, and that the smaller the particles, the sharper the curve; that the surface of a granule is liable to receive and retain molecules evaporated from other granules, and that, other things being equal, the retention of particles is more common on the surfaces which are least curved. It follows that the larger granules of less curvature will lose less and gain more, on the average, than the smaller granules of greater curvature. The larger granules therefore grow at the expense of the smaller. Furthermore, small granules melt more readily than large ones, and where the temperature is nicely adjusted between melting and freezing, the smaller may lose while the larger gain.
Another factor that affects the growth of the granules is pressure and tension. The granules are compressed at the points of contact and put under tension elsewhere, and pressure and tension are, on the average, likely to be relatively greatest for the smallest granules. Tension increases the tendency to evaporation, and the capillary spaces adjacent to the points of contact probably favor condensation. Pressure reduces the melting-point, while tension raises it. [p. 278] Though the effect of this is slight, it is to be correlated with the much more important fact that compression produces heat which mayraise the temperature of the ice to the melting-point at some points, while tension may reduce the temperature to or below freezing, at other adjacent points. There is therefore a tendency for the ice to melt at the points of contact and compression, and for the water so produced to refreeze at adjacent points where the surface is under tension. This process becomes effective beneath a considerable body of snow, and here the granules gradually lose the spheroidal form assumed in the early stages of granulation, and become irregular polyhedrons, interlocked into a mass of more or less solid ice.
The moisture of the air may be of importance in the process, even far below the surface. Under severe wind pressure, air penetrates porous bodies appreciably, as the “breathing” of soils, “blowing-wells,” and “blowing-caves” teach us, for all these things are expressions of the effective penetration and extrusion of the air, under variations of barometric pressure. In the snow-fields, and in the more granular portions of glaciers near their heads, the porosity is doubtless sufficient to make this process effective. The probable effect is two-fold: (1) The condensation within the ice of moisture from the air at some times, and (2) escape of moisture from within the ice at others. These alternating processes are attended by oscillations of temperature which may involve melting and freezing, and these processes produce granular change.
Whether these processes furnish an adequate explanation of the changes or not, all gradations may be observed from snowfiakes into granular neve, and thence into the granules of glacier Lee, ranging in size up to that of walnuts, and even beyond. In coherence, these aggregations vary from the neve (coarse-grained snow) stage where the grains are small and spheroidal, to the ice stage where the cohesion is strong through the interlocking growths of the large granules.
Structure and arrangement of the crystals in glacier ice. A crystal of ice is made up of a series of plates arranged ;if eight angles to the principal axis of the crystal. These plates may be likened to a pile of cards, the principal axis being represented by a line verticil to them. If a cube be cut from a large crystal of ice, it will behave [p. 279] much like a cube cut from the pile of cards. If it is so placed that its plate’s are horizontal (Fig. 235), and if it is rested on supports at two edges and heavily weighted in the middle, it will sag, the plates sliding slightly over one another. In this case the cube offers considerable resistance to deformation. If the cube is so placed that the plates are on edge, each reaching from support to support (Fig. 236), it will offer very great resistance to deformation. But if the plates are vertical and transverse to the line joining the supports, as in Fig. 237, the middle portion will sag under moderate weighting by the sliding of the plates on one another, and in a comparatively short time the middle portion may be pushed entirely out, dividing the cube.[8] These properties have been made the basis for the hypothesis that glacier motion is primarily the result of the slipping of the plates of the crystal on one another. This hypothesis might have much in its favor if the crystals of ice in the glacier were all oriented so that the plates were parallel with the bed of the glacier; but this is not the case. The crystals, starting from snowflakes, have their axes turned in various directions according to the accidents of their fall; and as the snow develops into ice, the principal axes of the granules continue to lie in all directions.
There does appear to be a tendency, however, for the crystals of ice to approach parallelism in the basal, terminal part of a glacier. From such observations on this point as are available, it is probable that the parallel orientation is due (1) partly to the vertical pressure [p. 280] of ice, and (2) partly to the shearing and foliation of the ice, or to the causes which produce these structures. In this position, then, the slipping of crystal plates on one another may be a factor in glacial motion.
The Probable Fundamental Element in Glacier Motion
Melting and freezing. There seems to be no escape from the conclusion that the primal cause of glacier motion is one which may operate even under the relatively low temperatures, the relatively dry conditions, and the relatively granular textures which affect the heads of glaciers. These considerations lead to the view that movement there takes place by the minute individual movements of the grains upon one another. While they are in the spheroidal form, as in the neve, this would not seem to be difficult. They may rotate and slide over each other as the weight of the neve increases. After they become interlocked by further growth, rotation and sliding must be more difficult. The amount of motion required of an individual granule is surprisingly small. In order to account for a movement of three feet per day in a glacier six miles long, the mean motion of the average granule relative to its neighbor would be roundly, 1/10000 of its own diameter per day; in other words, it should change its relations to its neighbors to the extent of its diameter once in about thirty years. A change of such slowness under the conditions of granular alteration can scarcely be thought incredible, or even improbable, in spite of the interlocking which the granules may develop. It is conceivable that there may be motion between the granules comparable to that between shot in great quantities in similar positions; but the movement is supposed to be effected chiefly by the temporary passage of minute portions of the granules into the fluid form at the points of greatest compression, the transfer of the water thus produced to adjoining points, and its reeolidification. The points of greatest compression are obviously those whose yielding most promotes motion, and the successive yielding of points which come in succession to oppose motion most (and thus to receive the greatest stresses), permits continuous motion. It ionly necessary to assume that the gravity of the accumulated mass is sufficient to produce a little temporary liquefaction at the points [p. 281] of greatest stress, the result being accomplished not so much by the lowering of the melting-point as by the development of heat by pressure.
This conception of glacial movement involves the momentary liquefaction of minute portions of the ice, while the mass as a whole remains rigid, as its crystalline nature requires. Instead of assigning a slow viscous fluidity like that of asphalt to the whole mass, which seems inconsistent with its crystalline character, it assigns a free fluidity, momentarily, to a succession of particles that form only a minute fraction of the whole at any instant.
This conception is consistent with the retention of the granular condition of the ice, with the heterogeneous (in the main) orientation of the crystals, with the rigidity and brittleness of the ice, and with its strictly crystalline character, a character which a viscous liquid does not possess, however much its high viscosity may make it resemble a rigid body.
Accumulated motion in the terminal part of a glacier. However slight the relative motion of one granule on its neighbor, the granules in any part of a glacier partake in the accumulated motion of all parts nearer the source, and hence all except those at the head are thrust forward. Herein appears to lie the distinctive nature of glacial movement. Each part of a stream of water feels (1) the hydrostatic pressure of neighboring parts (theoretically equal in all directions), and (2) the momentum of motion, but not the thrust of the water up stream. This is probably one of the fundamental differences between water flow and glacier motion.
Lava streams are good examples of viscous fluids flowing in masses comparable to those of glaciers, on similar slopes, and, in the later stages of motion, at similar rates; but their special modes of flow and their effects on the sides and bottoms of their paths are radically different from those of glaciers. Forceful abrasion, and particularly the rigid holding of imbedded stones which score and groove the rock beneath, is unknown in lava streams, and is scarcely conceivable. There is, so far as we know, no experimental or natural evidence that any viscous fluid, in the ordinary sense of that term, detaches and picks up fragments and holds them firmly as graving tools in its base so as to cut deep, long, straight grooves [p. 282] in the hard bottom over which it flows. It would seem that competency on the part of a viscous body to do this peculiar class of work so distinctive of glaciers, should be demonstrated before the viscous theory of glacial movement is accepted as even a good working hypothesis. In contrast with viscous movement, it is conceived that a glacier is thrust forward rigidly by internal elongation, and that it is sheared forcibly over its sides and bottoms f leaving its distinctive marks upon them.
Auxiliary Elements
Shearing. In the lower part of a glacier, where the thrusts are greatest, normally, where the granules arc fewest and their interlocking most intimate, shearing takes place within the ice. This is illustrated by Figs. 238 and 239. The shearing results in the foliation of the ice, and in the dragging of debris along the planes of shear. Thus the ice becomes loaded in a special englacial, or basoenglacial fashion, as previously mentioned and illustrated.
[p. 283]
Within the zone of shearing, the gliding planes of the crystals may come into effective function. It is thought that the combined effect of the vertical pressure, the forward thrust, and the basal drag of the ice, may be to increase the number of granules whose gliding planes are parallel to the glacier’s bottom. At any rate, such an arrangement in the basal portion of the Greenland glaciers at their borders is said to exist. It is conceived that where strong thrusts are brought to bear upon such a mass of granules, thosewhose gliding planes are parallel to the direction of thrust are strained with sufficient intensity to cause the plates to slide over each other, while those which are not parallel to the direction of thrust are either rotated into parallelism — when they also yield — or are pressed aside out of the plane of shear. As previously noted, shearing is observed to occur chiefly where the ice below the plane of shearing is protected more or less from the force of the thrust, as in the lee of a hill or mass of debris. It perhaps also occurs at the top of the basal zone of ice so loaded with debris that it is incapable of ready movement.
[p. 284]
It is probable, also, that sharp differential strains and shearing are developed at the level where the surface water of the warm season, descending into the ice, reaches the zone of freezing; for the expanding of the freezing water at the upper limit of the cold zone may cause the layer expanded by it to shear over that below. As the level of freezing descends with the advance of the warm season, the zone of shearing sinks. This is a phase of shearing developed in a special horizon relatively near the surface. It may be noted that expansion at the zone where descending water freezes, not only leads to shear, but to the development of surface cracks, for the surface is stretched as the zone below expands. In the course of years, the cracks developed in this way may become wide crevasses, limited below by the depth of the zone of freezing in summer.
High temperature and water. In the zone of waste, the higher temperature and the greater abundance of water lend their aid to the fundamental agencies of movement; and there is need for these aids to promote a proportionate movement, for here the granules are more intimately interlocked, and the ice more compact and inherently more solid and rigid. The average temperature is near the melting-point (p. 247), and during the warm season the ice is bathed in water all the time, so that the necessary changes in the crystals are facilitated. Under these conditions, movement takes place more readily than in the more open granular ice of lower temperature and drier state.
Application. The co-operation of the auxiliary agencies of motion with the fundamental ones appears to explain the peculiarities of glacial movement. In regions of intense cold, where a dry state and low temperature prevail, as in the heart of Greenland, the snow-ice mass may attain an extraordinary thickness. Here the burden of movement seems to be thrown almost wholly upon compression, with the slight aid of molecular changes due to internal evaporation and allied inefficient processes. Since the temperature in the upper part of the ice is very adverse (p. 245), the compression must be great before it becomes effective in melt ing the ice, and hence the great thickness of the snow before motion is considerable. -Similar conditions more or less affect the heads of Alpine glaciers, though here the high gradients favor motion with [p. 285] lesser thicknesses of ice. In the lower reaches of Alpine glaciers, where the temperatures are near the melting-point, and where the ice is bathed in water much of the time, movement may take place in the ice which is thin and compact.
If the views here presented are correct, there is also, at all points below the source, the cooperation of rigid thrust from behind, with the tendency of the mass to move on its own account. The latter is controlled by gravity, and conforms in its results to laws of liquid flow. The former is a derived factor, and is a mechanical thrust. This thrust is different from the pressure of the upper part of a liquid stream on the lower part, because it is transmitted through a body whose rigidity is effective, while the latter is transmitted on the hydrostatic principle of equal pressure in all directions. Thrust would be most effective toward the end or edge of a glacier.
Corroborative Phenomena
The conception of the glacier and its movement here presented explains some of the anomalies that otherwise seem paradoxical. If the ice is always a rigid body which yields only as its interlocking granules change their form by loss and gain, a rigid hold on the imbedded rock at some times, and a yielding hold at others, is intelligible. Stones in the base of a glacier may be held with great rigidity when the ice is dry, scoring the bottom with much force, while they may be rotated with relative ease when the ice is wet. In short, the relation of the ice to the bowlders in its bottom varies radically according to its dryness and temperature. A dry glacier is a rigid glacier. A dry glacier is necessarily cold, and a cold glacier is necessarily dry.
It is difficult to explain the furrows and grooves cut by glaciers in firm rock if the ice is so yielding as to flow under its own weight on a surface which is almost flat. If the mass is really viscous, its hold on its imbedded debris should also be viscous, and a bowlder in the bottom should be rotated in the yielding mass when its lower point catches on the rock beneath, instead of being held firmly while a groove is cut. This is especially to the point since viscous fluids flow by a partially rotary movement.
On the view here presented, a glacier should be more rigid in [p. 286] winter than in summer. The total thickness of a glacier should experience this rigidity of winter 'at its ends and edges, where the ice is thin enough to permit the low temperature to affect its bottom. The motion in these parts during the winter is, therefore, very small.
In this view, also, may be found an explanation of the movement of glaciers for considerable distances up-slope, even when the surface of the ice, as well as its bed, is inclined backwards. So far does this go, that superglacial streams sometimes run for some distance backwards, i. e., toward the heads of the glaciers, while in other places surface waters are collected into ponds and lakelets. Such a slope of the surface of ice is not difficult to understand if the movement is due to thrust from behind, or if it is occasioned by internal crystalline changes acting on a rigid body; but it must be regarded as very remarkable if the movement of the ice is that of a fluid body, no matter how viscous, for the length of the acclivity is sometimes several times the thickness of the ice. Crevassing and other evidences of brittleness and rigidity find a ready elucidation under the view that the ice is really a solid body at all times, and that its apparent fluency is due to the momentary fluidity of small portions of its mass assumed in succession as compression demands.
In addition to the considerations already adduced, it may be urged that a glacier does not flow as a stiff liquid because its granules are not habitually drawn out into elongated forms, as are cavities in lavas, and plastic lumps in viscous bodies. Flowage lines comparable to those in lavas are unknown in glaciers.
All this is strictly consistent with our primary thesis, that a glacier is crystalline rock of the purest and simplest type, and that it never has other than the crystalline state. This strictly crystaline character is incompatible with viscous liquidity.
Other Views of Glacier Motion
While these views of glacial motion seem to us to best accord with the known facts, they are not to be regarded as established in scientific opinion, or as the views most commonly held. The mode of glacial motion has long been a mooted question, and is still so [p. 287] regarded. The main alternative interpretations that have been entertained are the following:
(1) In the early days of glacial studies De Saussure thought that glaciers slid bodily on their beds.
(2) Charpentier and Agassi z referred the movement to the expansion of descending water freezing within the glacier.
(3) Rendu and Forbes, followed by many modern writers, believed ice to be viscous, and that in sufficiently large masses it flowrs under the influence of its own weight, like pitch or asphalt.
(4) .Others, realizing the fundamental difference between crystalline ice and a true viscous body, have fallen back on a vague notion of plasticity, which scarcely amounts to a definite hypothesis at all.
(5) Tyndall urged that the movement was accomplished by minute repeated fracturing and regelation, appealing to the fact that broken pieces of ice slightly pressed together at melting temperatures freeze together, but neglecting the fact that this would destroy the integrity of the crystals.
(6) Moseley assigned the movement to a bodily expansion and contraction of the glacier, analogous to the creeping of a mass of lead on a roof.
(7) James Thompson demonstrated that pressure lowers the melting-point, and while this effect is so small as probably to be ineffectual, it is correlated with the very important fact that compression may cause melting, which is not the case in most other rocks. He recognized that under pressure partial liquefaction took place, that the water so liberated might be re- frozen as it escaped from pressure, and appears to have regarded this as a vital factor.
(8) Croll held that the movement was due to a consecutive series of molecular changes somewhat like the chain of chemical combinations in electrolysis.
(9) Hugi, Eli de Beaumont, Bertin, Forel, and others thought that the growth of the granules was the leading factor in the ice movement.
(10) McConnel and Mugge have made the gliding planes of the ice crystals serve an important function in glacial movement.
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It will be seen that the principle of partial liquefaction for which Thompson laid the basis, the crystallization of descending water, urged by Charpentier and Agassiz, and the granular growth on which Hugi, Beaumont, Forel, and others founded their hypotheses, are incorporated in the view already presented. Probably the agencies on which some of the other views are based may also be participants in producing glacial motion, sometimes as incidental factors, and sometimes perhaps as important ones, for under different conditions, different agencies may play roles of varying importance. For example, in going over the brinks of precipices of sufficient height, glaciers break into fragments which are recemented below, and the “reconstructed” glacier moves on as before. Here fracture and regelation are evident, though they are hardly causes of motion. The movement of the gliding planes of the ice crystals over each other, which has been looked upon as a special kind of viscoid movement, probably plays a part in shearing. But neither of these is probably a large factor in ordinary glacial movement, and it seems highly improbable that any of them are essential factors in the movements in the snowfields where glacial motion begins.
Map work. The effects of glaciation are well shown on many of the topographic maps of the U. S. Geological Survey. Lists of such maps are found on pages 230 and 289 of the junior author’s Physiography, Advanced Course. Plates XCV to CXXIX, found in Professional Paper 60, U. S. Geological Survey, present a number of maps illustrating this topic.
For an account of experiments illustrating the mobility of ice see Aitkin, Am. Jour. Sci. Vols. V, 1873, p. 305, and XXXIV, 1887, p. 149, and Nature, Vol. XXXIX, p. 203. ↩︎
Jour, of Geol., Vol. Ill, p. 888. ↩︎
The following list includes many of the more available articles and treatises on existing glaciers; others are referred to in the following P&ge4
Alaskan glaciers: Reid, (1) Nat. Geog. Mag., Vol. IV, pp. 19 55 Sixteenth Ann. Rept., U. S. Geol. Surv., Part I, pp. 421-461. Russell. (1) Nat. Geog. Mag., Vol. Ill, pp. 176-188; (2) Jour, of Geol., Vol. I, pp. 219-245.
Glaciers in the United States: (1) Russell, Eighteenth Ann. Kept., I 8. Geol. Surv., Part II, pp. 379-409; (2) Glaciers of North America.
Greenland glaciers: Chamberlin, Jour, of Geol., Vol. II. pp 768-788; Vol III, pp. 61-69, 198-218, 469-480, 565-582, 668-681, and 833-843; Vol. IV. pp. 582-592. Salisbury, Jour, of Geol., Vol. IV, pp. 769-810.
General: Shaler and Davis, Illustrations of the Earth’s Surface. ↩︎
Reid. Natl. Geog. Mag., Vol. IV, p. 44. ↩︎
Reid. Variations of Glaciers. Jour, of Geol., Vols. Ill, p. 278; V, p. 378; VI, p. 473; VII, p. 217; VIII, p. 154; IX, p. 250; X, p. 313; XI, p. 285; XIII, p. 313; XIV, p. 402; and XV, pp. 46 and 664. ↩︎
Russell. Jour, of Geol., Vol. Ill, p. 823. ↩︎
Geikie. The Great Ice Age, 3d ed., p. 529. ↩︎
McConnel and Mi’igge. On the Plasticity of Glaciers and Other Ice. Proc. Roy. Soc, Vol. XLIV, 1888, pp. 331-67 (with D. A. Kidd); Vol. XLVIII, 1890, pp. 259, 260; Vol. XLIX, 1891, pp. 323-43. ↩︎