A part of the atmospheric precipitation falls as snow, and this, like the rain, does its appropriate work in degrading the land. Over the larger part of the land surface the snow of the winter does not endure through the succeeding summer, and when it melts it follows the same course as the precipitation which falls as rain; but in cold regions where the fall of snow is heavy some of it remains unmelted and constitutes perennial snow-fields.
Snow-fields.—Mountain heights and polar lands are the most common habitats of snow-fields, though they are not confined to these situations. In North America there are numerous 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 snow-fields south of the parallel of 36° 30′, and most of the many hundreds north of that latitude (excluding Alaska) are small (Pl. XVIII, Fig. 1, Washington, Lat. 48° 5′, Long. 121° 5′; Fig. 2, Lat. 41° 25′, Long. 122° 12′. From Glacier Peak and Shasta Special Quadrangles, U. S. Geol. Surv.). Farther north, especially in Alaska, the snow-fields of the western mountains attain much greater size. In Europe snow-fields comparable to those of the northwestern part of the United States and British Columbia occur in the Alps (Fig. 221), the Pyrenees, the Caucasus, and the Scandinavian mountains. In Asia snow-fields occur in the Himalayas and in many of the high mountains farther north, from Turkestan on the southwest nearly to the coast on the northeast. In South America there are snow-fields of small size even in equatorial latitudes, and farther south 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 considerable snow-field is really ice.
In addition to these limited fields of snow in mountain regions, there are fields of much greater extent covering wide expanses of plain and plateau in the polar regions. The greater part of the island of Greenland is covered with a single field of ice and snow, the size of which is variously estimated at 300,000 to 400,000 square miles (Fig. 222)—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, the areas of the snow-fields far exceeding those of most mountain regions. In Antarctica there is believed to be 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 of 3,000,000 or 4,000,000 square miles.
The only condition necessary for a snow-field is an excess of snowfall over snow waste. The lower edge of a snow-field, the snow-line, is dependent chiefly on temperature and snowfall. In general it does not depart much from the summer isotherm of 32°, though it may be well above this isotherm where the snowfall is light. That the snow-line is not a function of temperature only is shown by its position in various places. In the equatorial portion of the Andes, for example, 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 3000 or 4000 feet lower on the south side than on the north.
While in equatorial regions the snow-line has an altitude of 15,000 to 18,000 feet, it approaches or even reaches sea-level in the latitude of Antarctica and North Greenland. In intermediate latitudes it has an intermediate position.
While temperature and snowfall 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 névé and ice.—The snow does not lie on the surface long before it undergoes obvious change. The light flakes soon begin to be transformed into granules, and the snow becomes “coarse-grained.” The granular character, so pronounced in the snow of the last banks which remain in the spring in temperate latitudes, is even more distinct in perennial snow-fields, either at the surface or just beneath it. This granular snow is called névé, or firn. Still deeper beneath the surface, where the thickness of the snow is great, the névé becomes more compact and finally coherent, and grades into porous ice. This gradation is accomplished at no great depth, though the thicknesses of snow and névé are by no means constant.
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 the form of layers. This follows from two or three conditions. The snow of one season, or of one period of precipitation, may have been considerably changed before the succeeding fall of snow. So also 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 and dropped on the surface during the melting season, and some of which was concentrated in that position 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, or perhaps of minor periods of precipitation, and makes distinct the stratification which would otherwise remain obscure. The snowfall of successive years has been estimated by this means[122] where the snow is exposed in crevices in the snow-field.
In addition to its rude stratification, the ice of the deeper portions 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 result mainly 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 present from the beginning, for it is assumed by the snowflakes when they form, and the subsequent changes seem only to modify the original crystals by building up some and destroying others. By the time the snow is converted into névé, the granules have become coarse, and wherever the ice derived from the névé has been examined, the granular crystalline texture is present. The individual crystals in the ice are usually larger than those of the névé, and more closely grown together. In the fresh unexposed ice the crystals are so intimately interlocked that they are not readily seen except under a polarizing microscope, but when the ice has been honeycombed by partial melting, the granules become partially separated and may be easily seen. Fig. 223 shows quantities of them which have been washed down from the surface, and disposed as cones at its base. While a given mass of snow in a great snow- and ice-field cannot be followed consecutively through its whole history, yet since (1) the granular texture is pronounced in the névé stage where the granules show evidences of growth, and since (2) 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, and since (3) the crystals are, on the average, larger in proportion to the lateness of the stage of their history, while (4) experiment has shown that granules grow under the conditions which exist in snow-fields, and (5) that they persist under very considerable pressure, it is legitimate to assume that a granular crystalline condition persists throughout all stages, and is a feature of progressive growth.
Inauguration of movement.—Eventually the increase in depth of snow and ice in a snow-field gives rise to motion. The exact nature of the motion has not yet been demonstrated to the satisfaction of all investigators. Brittle and resistant as ice seems, it exhibits, under proper conditions, some of the outward characteristics of a plastic substance. Thus it 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 time.[123] 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, as will be seen later, to question this interpretation of the ultimate nature of the movement. Whatever this may be, the mass 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.
If both the surface on which the ice-sheet develops and its surroundings be essentially plane, as may happen in high latitudes, and if the snow- and ice-field be symmetrical in shape, the outward movement will be approximately equal in all directions, and the area covered by the spreading ice-field will remain more or less circular. If the ice-field rests on a steeply inclined surface, like a mountain slope, the movement becomes one-sided in conformity to the slope. If the surface, otherwise plane, be affected by valleys parallel to the direction of movement, the ice in the valleys will be deeper than that on the divides between them, and its movement stronger. In the valleys, therefore, the ice will advance farther than elsewhere before being melted, and the outline of the ice will become lobate, the lobes occupying the depressions. These general relations are shown in Figs. 224 and 225. If the depressions be wide and shallow, the lobes will be broad and short (Fig. 226); if the depressions be narrow and deep, the lobes will be relatively narrow and Long. If the snow and ice rest on a surface consisting chiefly of steep valleys and sharp ridges, as is common on mountain slopes, the snow and ice are chiefly gathered in the valleys, and take a linear form.
These different forms give rise to different terms. The ice which spreads with some approach to equality in all directions from a center is a glacier, is indeed the type of the greatest glaciers, but is commonly called an ice-cap. The same name is applied to any glacier in which there is movement in all directions from the center, even though its shape departs widely from a circle. The glacier covering the larger part of Greenland (Fig. 222) is a good example of a large ice-cap, and the glaciers on some of the flat-topped peninsular promontories of the same island are good examples of small ones (Fig. 224). If ice-caps cover a large part of a continent, as some of those of the past have done, they are often called continental glaciers.
Where ice-caps are developed on plateaus whose borders are trenched by valleys, ice-tongues from the edge of the ice-cap often extend down into the valleys and give rise to one type of valley glacier (Figs. 224 and 227). A second and more familiar type of valley glacier occupies mountain valleys, and is the offspring of mountain snow-fields (Fig. 228). The former are confined chiefly to high latitudes, and are distinguished as polar or high-latitude glaciers (Figs. 227 and 229); the latter are known as alpine glaciers (Figs. 228 and 230). The end and side slopes of polar 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 deploys, forming a fan (Fig. 231). The deploying ends of adjacent glaciers sometimes merge, and the resulting body of ice constitutes a piedmont glacier (Fig. 232). At the present time, 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. In many cases it does not even enter the narrow valley which leads out from the depression occupied by the snow-field. Such a glacier is nestled in the face of a cliff, and may therefore be called a cliff glacier[124] (Figs. 233 and 234). The snow-field of a cliff glacier is sometimes no more than a great snowdrift, accumulated through successive years. Cliff glaciers are often as wide as long, and are always small, and between them and valley glaciers there are all gradations (Fig. 235). 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 ice fragments may then again 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.
Of the foregoing types of glaciers, the ice-caps far exceed all others both in size and importance, while the valley glaciers out-rank, in the same respects, the other types; but since the valley glaciers are the most familiar type, the general phenomena of glaciers will be discussed with primary reference to them.
Dimensions.—Glaciers which occupy valleys leading down from snow-fields sometimes reach the upper parts of the valleys only, 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, and though their width is usually much less than their length, the reverse is sometimes the case (Figs. 233, 234, and 235). Their thickness is usually measured by hundreds of feet rather than by denominations of other orders, 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 a glacier. As a rule, it is thinnest in its terminal portion, and thickest at some point intermediate between this and its source, but nearer the latter than the former. Cliff and reconstructed glaciers are comparable in size to the smaller valley glaciers. Piedmont glaciers may attain greater size.
An ice-cap is theoretically thickest at its center and thins away to its borders, but its actual dimensions are influenced by the topography on which it is developed. The Greenland ice-cap is known to rise about 9000 feet above the sea, and it probably reaches considerably higher than this in the unexplored center of its broad dome. The height of the land surface beneath is unknown, but it is unlikely that it averages half this amount, and hence the ice is probably 5000 feet or more thick in the center. There is reason to think that it is much thicker in Antarctica.
Limits.—The ice of a glacier is always moving forward (neglecting temporary halts), but the end of a glacier may be retreating, advancing or remaining stationary, according as the rate of wastage is greater, less, or just equal to the forward movement of the ice. The position of the lower end of the 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, since there must be movement from the granular mass of ice beneath the snow to make up for the waste below, and the moving ice beneath the snow-field which feeds the tongue of ice in the valley is just as really a part of the glacier as the more consolidated portion in the valley below. If a definite upper limit for an alpine glacier is to be named, it should probably be the Bergschrund, a gaping crevasse, or series of crevasses which sometimes open near the precipitous slope of the peak or cliff where the snow-field lies. The Bergschrund is formed by the moving of the lower part of the snow-field away from the portion above.
The lower end of a glacier is usually free from snow and névé in summer, but, traced toward its source, it first becomes covered with névé, then with snow, and finally merges into the snow-field without having ceased to be a glacier. The term glacier is, however, commonly used to mean merely the more solid portion outside (below) the névé.
Movement.—The fact of glacier movement is established in various ways, the most obvious being by the advance of its lower end. Such advance is too slow to be seen from day to day, and is only detected when the lower end of the glacier overrides or overturns objects in front of it, or moves 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 the surface. If the positions of these marks relative to fixed points on the sides of the valley be 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 the 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 (Fig. 236), but in some cases, especially in large glaciers, there are found to be two or more main lines of movement, with belts of lesser movement between. The top of a glacier moves, on the whole, more rapidly than the bottom, though the observations made do not show that the rate of movement diminishes regularly downward, and it probably does not so diminish in many cases. 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 seven feet or more per day,[126] and some of the glaciers of Greenland have been found to move, in the summer time, 50, 60, or even more feet per day. A single estimate as high as 100 feet per day has been made; but these high rates have been observed only where the ice of a large inland area crowds down into a comparatively narrow fjord, 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. Rink says that “between 62° and 68° 30′, the edge of the inland ice is almost stationary for a remarkably long distance.”[127] The observations of the authors between 77° and 78° were of like import. Probably the average movement of the border of the Greenland ice-cap is less than one foot a week.
Conditions affecting rate of movement.—The rate of glacier movement appears to depend on (1) the depth of the moving ice; (2) the slope of the surface over which it moves; (3) the slope of the upper surface of the ice; (4) the topography of the bed over which it passes; (5) the temperature; and (6) the amount of water which falls upon it or is carried to it by the drainage of its surroundings, in addition to that produced by the melting of the glacier itself. Great thickness, a steep slope, much water, smoothness of bed, and a high (for ice) temperature favor rapid movement. Since some of these conditions, notably temperature and amount of water, vary with the season, the rate of movement for any given glacier is not constant throughout the year. Other conditions, especially the first of those mentioned above, vary through longer periods of time, and occasion periodic variations in the rate of movement.
Since the volume of ice concerned influences the rate of movement, anything which changes the volume affects the rate. An excess of snowfall with favorable conditions for its preservation for a period of years, increases the volume of ice, and tends to accelerate its movement. A deficiency in snowfall, or in its preservation, as from high average temperature or from aridity, diminishes the quantity of ice, and so retards the movement. An acceleration of velocity causes the ice to move down the valley farther before being melted, that is, causes the end of the glacier to advance, while a decrease of velocity produces the opposite effect. As a matter of fact, the lower ends of glaciers advance for a period of years and then retreat, to advance again at a later time.[128] Observation has shown 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. Present knowledge seems to point to a period of 35 to 40 years as the time within which a cycle of fluctuation, that is, an advance and a retreat, takes place.
A declining upper surface is essential to glacier motion. There are short stretches where this is not the case, and indeed there are particular places where the upper surface slopes backward.[129] This may occur where the ice is pushed up over a swell in its bed, or is crowded up against any considerable obstacle; but such cases are no more than local exceptions, and do not militate against the truth of the general 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 local 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 upper surface must have declined in the direction of movement. With a given thickness of ice, the greater the decline of its lower surface in the direction of motion, the more rapid its progress. A rough bed, or a crooked course retards the motion of a glacier, while a smooth bottom and a straight course facilitate it.
Slope, roughness of bed, and volume affect the movement of glaciers somewhat as they affect the movement of rivers. The temperature of the 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 most 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.
Likenesses and unlikenesses of glaciers and rivers.—Many of the characteristics of a valley glacier may be understood from the study of the accompanying figure (Fig. 237) of the White (Alaska) glacier. From this figure it will be seen that the glacier is an elongate river-like body, following the curves of the valley in stream-like fashion. It has its origin in the snows collected on the mountain heights seen in the distance, 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 constricted portions of the valley and slower in the broader parts. These and other likenesses, some of which are apparent rather than real, have given origin to the view that glacier ice moves like a stiff viscous liquid.
But while the points of likeness between glaciers and rivers are several, their differences are at least equally numerous and significant. The trains of débris on the surface (the dark bands in the illustration), like the central currents of streams, pass nearer the projecting points of the valley walls and farther from the receding bends; but beyond this point the analogy fails, for the trains of débris on the ice do not conform in detail to the courses of the currents of a winding stream, nor is there evidence of the rotatory motion that characterizes river water. 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 this tension are many. An obvious cause is an abrupt increase of gradient in the bed (Fig. 238). If the change of gradient be considerable, an ice-fall or cascade results, and the ice may be greatly riven (Fig. 228). Below the cascade, the surface may bristle with wedges and pinnacles of ice (séracs, Fig. 239). Transverse crevices at the margin sometimes appear to be the result of the tension developed on a curve. Oblique crevices on the surface near the sides are commonly ascribed to the tension between the faster-moving center and the slower-moving margins, and in like manner crevasses that rise obliquely from the bottoms are attributed to the tension between the faster-moving portions above and the slower-moving portions 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 river-like 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. Somewhat similar cracks develop in the outer crust of asphalt, when a mass of it is allowed to stand and spread; but in this case there is evaporation of the volatile ingredients, giving to the outer part relative rigidity and brittleness, while the inner part remains more fluent. The analogy is therefore not perfect and probably not really illustrative. The crevices may be narrow or wide, and both narrow and wide may be found in the same glacier. The narrow crevices that never open much are the most significant, as they show that very little stretching is needed to satisfy the tension. The opening of a gaping crevice is sometimes the work of weeks, and, in the slow-moving glaciers of high latitudes, sometimes the work of successive seasons. All this shows that the glacier is a very brittle body, incapable of resisting even very moderate strains brought to bear upon it very slowly. Had the ice even moderate ductility, it would adapt itself to tension brought to bear upon it so slowly as are many of the tensions which produce crevassing. In its behavior under tension therefore a glacier is notably unlike a river.
Topography.—Many of the minor irregularities of the surface of a glacier are the result of crevassing. After the ice is crevassed, the sun’s rays and the air which has been warmed by them penetrate the openings and melt the ice. The melting is most rapid at the top, and decreases downward. The result is that the sections of ice between adjacent crevasses are narrowed into wedges. If there be cross-crevassing, as is common, points instead of wedges result. As the sort of surface shown in Fig. 239 develops, any débris which was on the ice slides into the crevices, and the upper surface becomes clean.
Where ice is crevassed transversely, and where melting is not rapid, the crevasses may close as the ice moves forward, and the regelation of adjoining faces heals the rents in the surface. Even in this case, however, the surface is likely to be more or less undulating because of the waste on the sides of the crevices before they are closed. After regelation, surface ablation tends to obliterate the protuberances.
The topography of the surface of the ice is affected by other conditions. All parts of the surface of the ice are not equally compact, and the least compact portions melt most rapidly, giving rise to depressions, while the more solid parts occasion protuberances. Both depressions and protuberances may be regular or irregular in form (Figs. 240 and 241). Undulations of the bed often show themselves in the surface of the ice as suggested by Fig. 242. In such cases, ponds or lakelets sometimes accumulate on the surface of the ice. The topography of the ice in such cases seems to show that the ice is forced up slope.
Surface moraines.—The surface of a glacier is often affected by débris of one sort or another, and this also influences its topography. The débris is sometimes disposed in the form of belts or moraines (Figs. 237, 243). The surface moraines may be lateral, medial, or terminal. A lateral moraine is any considerable accumulation of débris 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 near the middle. There may be several medial moraines on one glacier, in which case some of them may be far from the center. In alpine glaciers, the surface terminal moraine is less well-defined; in polar glaciers it often connects two lateral moraines, making a loop roughly concentric with the terminus of the glacier.
Besides the surface moraines, which represent belted aggregations of débris, 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 débris.—The débris on the ice affects its topography by influencing the melting of the subjacent and adjacent ice. The rock débris absorbs heat more readily than the ice. A small and 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 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 large mass of stone is not notably warmed. The ice beneath it is protected from the direct rays of the sun, and is therefore melted more slowly than that around it. The result is that the bowlder presently stands on a protuberance of ice (Fig. 244). When its pedestal becomes high, the oblique rays of the sun and the warm air surrounding it cause it to waste away, and the capping bowlder falls. In high latitudes, the great obliquity of the rays sometimes allows them to strike under isolated bowlders. In this case, they are warmed from below, and thus aid rather than hinder the melting of the ice.