The same principles apply to the moraines. A thin bowlder moraine in high latitudes is sometimes sunk below the surface (Fig. 242). Usually, however, a medial moraine protects the ice beneath from melting, and occasions the development of a ridge of ice beneath itself. As the ice on either side is then lowered by ablation, the moraine matter of the medial belt tends to slide down on either hand. The same is true of the lateral moraines. So far does this spreading go, that in some cases the lower end of a glacier is completely covered with the débris which has spread from the medial and lateral moraines. Examples of this may be seen in almost any region of abundant, long, alpine glaciers.
Dust-wells.—The wind-blown dust sometimes gives rise to peculiar topographic features of small size. The dust is not distributed by the wind with absolute equality, and the surface drainage of the ice tends to aggregate it. Every dust particle acts like a small stone, and where aggregations of dust occur, they melt their way down into the ice, developing holes or “dust-wells” (Fig. 245). These wells rarely reach a depth of more than a few inches, but they may be so numerous that the pedestrian is obliged to watch his steps. This is especially true near the edge of the large ice-caps. It is evident that the depth of these dust wells must be slight, for so soon as they are deep enough to cut off the sun’s rays from the dust at the bottom, the deepening ceases. Other things being equal, they are deeper in low latitudes than in high.
Débris below the surface.—The lower part of a glacier, as well as the upper, carries rock débris. This débris 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 débris, and the stationary débris which is full of ice, there seems to be a nearly complete gradation. The débris 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. These débris sheets are often numerous and usually discontinuous, though groups of such sheets often persist for considerable distances. Débris also occurs to some extent in the ice well above its base. It is sometimes in belts, as seen in section, and sometimes in bunches. These various relations are illustrated by Figs. 227, 229, and 246–249.
Another characteristic of the basal débris-laden part of some glaciers is the foliation of the ice (Figs. 248, 249, etc.). This is especially well shown in the arctic glaciers, the ends and sides of which have steep or vertical faces. The foliation is best developed in the débris zone, though often shown above. The foliation is sometimes minute, consisting of layers of clean ice, an inch or less in thickness, separated by mere films of earthy matter. In extreme cases there are a score or more of laminæ within a foot. Locally, and especially where débris is abundant, the laminæ are much contorted. This is seen both in section (Figs. 248 and 249) and on the surface (Fig. 250).
The temperature of glacier ice may range downward from the freezing point of water much as other solid portions of the earth’s surface, but it has a fixed upper limit at 32° Fahr. (0° C.) because all the heat it receives tending to raise its temperature above that point, is converted into the latent form by the melting of the ice. The range of temperature is greatest at the surface, where it varies from 32° in the summer, to the coldest temperature of the region where the ice occurs. Beneath the surface the range of temperature is more restricted, and increasingly so with increasing depth.
The variation of temperature at the surface is due primarily to the varying temperature of the air. During the cold season, a wave of low temperature (the winter wave), starting at the surface, penetrates the ice, and during the warm season a wave of higher temperature (the summer wave) takes the same course. The day and night waves and other minor variables are, for present purposes, negligible.
The winter wave.—There are but few observations on the internal temperatures of glaciers during the winter season, but it seems certain that the winter wave diminishes rapidly downward and dies out below, much as does the winter wave which affects land surfaces not covered with ice. Conduction alone considered, the temperature of the ice 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.
Assuming that in the high altitudes and high latitudes where glaciers abound, the temperature of the surface may average about −12° Fahr. (about −25° C.) for the winter half of the year, which is about the case for north Greenland, Spitzbergen, and Franz Josef Land, and that the conductivity of the ice in the C. G. S.[130] system is .005, the temperature would be lowered appreciably only about 40 feet below the surface at the close of the winter period, conduction only being considered. How far the internal temperature may be influenced by air forced through the ice by winds and by variations of the barometer is not known and cannot well be estimated. The wave of low temperature descending from the surface in winter would probably become inappreciable before reaching a depth of 60 feet. At this depth the temperature should be about 15° Fahr.—the mean annual temperature of the region.
The summer wave.—The warm wave follows the analogy of the summer wave of ice-free land surfaces much less closely. This is because of the low melting temperature of ice as compared with other forms of solid earth-matter. On this account the summer wave is bi-fold. The one part travels downward by conduction, the other by the descent of water; the one has to do primarily with the temperature before the melting-point of ice is reached; the other, with the temperature after that point is reached; the first conforms measurably to the warm wave affecting other solid earth-matter, while the second is governed by special laws. After the surface portion of the ice is brought to the melting temperature, the additional heat which it receives melts the ice and is transformed from sensible into potential heat. Ice charged 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. In regions where the average temperature is below freezing, the water-wave does not descend so far as the wave of conduction, since the latter descends below the zone where the melting temperature is found.
The foregoing considerations warrant the generalization that glaciers normally consist of two zones (1) an outer or upper zone of fluctuating 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. This may be the case in very thin glaciers in very cold regions, 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. The temperature of melting is a little below 32° Fahr. since pressure lowers the melting-point at the rate of .0133° Fahr. (.0075° C.) for one atmosphere of pressure. At the bottom of a mile of ice therefore the melting temperature is about 30.2° Fahr. (−1° C.) It is probable that in all thick glaciers the temperature of the bottom is constantly maintained at the melting-point. This may be indicated by the streams which issue from beneath glaciers during the winter, though this criterion is hardly decisive since the issuing waters may be derived partly or wholly from the rock beneath. 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 shown to lie between a maximum of 32° Fahr. and the minimum temperature of the region where the glacier occurs. Lower, in the zone of fluctuating temperature, the variation is less, and where the zone of fluctuating temperature passes into the zone of constant temperature, variation ceases. The thickness of the zone of fluctuating temperature varies with the temperature of the region where the glacier occurs, being greatest where the winters are coldest. In the case of all glaciers except thin ones in very cold regions, the temperatures within the zone of constant temperature range from the mean annual temperature of the region at the top of the zone (provided this 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.
If we consider 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 constant melting temperature, for the average temperature of regions where the ends of such glaciers occur is usually above 32° Fahr., and this determines a temperature of 32° Fahr. (or a little less) at the top of the zone, while a melting temperature is maintained at the bottom by the earth’s interior heat. In thin glaciers of very cold regions, where the zone of constant temperature has relatively slight thickness, the low temperature 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. In all other cases, the ice at the bottom of the under zone has a melting temperature, while that above is probably colder.
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, on the average, be below the melting-point, 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, Spitzbergen, and Franz Josef Land, and at certain high altitudes in more southerly latitudes, 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 about 6° below the melting-point. 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 perhaps worthy of especial note that much of it is but little below. In alpine glaciers the part of the ice affected by this constant low temperature (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 glacial 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. The compression arising from gravity, and the friction developed 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. Compression not only lowers the melting-point slightly, but it produces heat at the point of compression. Where the ice is granular, the compression, due to weight, takes place at the contacts of the grains. At intermediate points the pressure tends to cause them to bulge, and this has the effect of lowering the temperature of the bulging points. If therefore the compression be considerable, the granules may be warmed to the melting-point where they press each other, while at other points their temperature may be lower. In this case melting will take place at the points of compression, and the moisture so produced will be transferred to the adjacent parts of the granule and immediately refrozen. 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 would again take place. Thus the melting, the refreezing, and the attendant movement might go on until the limits of the power of gravity in this direction were reached. 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 depression and friction may modify the theoretical deductions drawn above from atmospheric and internal influences.
Summary.—If the foregoing generalizations be 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 be less than the thickness of the zone of annual variation or of permanent freezing temperature); (3) its subsurface portion in the zone of waste is generally melting, owing to descending water, compression, and friction; while (4) its subsurface portion in the zone of growth is probably below the melting-point except as locally brought to that temperature by compression, friction, and descending water, and at the bottom by conduction from the rock beneath.
Movement under low temperature.—Glacier motion will not be discussed at this point, but one of the bearings of the preceding conclusions on glacier motion may be pointed out. Since there must be motion in the area of growth to supply the loss in the area of waste, the fundamental cause of motion must be operative in bodies of ice the mean temperature of which is below the melting-point, unless the dynamic sources of heat are considerable. This fundamental cause does not exclude the coöperation of causes that work only (1) at the melting temperature, or (2) where the ice is bathed with water, or (3) in the plane of contact between wet ice above and dry ice below. These may be auxiliary causes which abet the fundamental one in producing the more rapid movement of warm seasons, or in bringing about the especially rapid movement in situations where there is abundant water, or in inducing the shearing which is such a remarkable feature of arctic glaciers.
Evaporation.—The ice wastes by evaporation as well as by melting, and while the former process is far less important than the latter, its results are probably larger than is commonly apprehended. One of the most remarkable features of some of the deposits of ancient glaciers is the slight evidence they afford of escaping waters. The most plausible explanation seems to lie in the supposition that the ice was largely wasted by evaporation. This conclusion finds support in many places in the presence of a mantle of fine silt over the drift, the silt being apparently composed of dust blown upon the ice. It is supposed to imply aridity in the region about the ice. If a sufficient mantle of dust were spread over the border zone of the ice, and if the air were very dry, nearly all the water melted on the surface of the ice might be held back by the dust-wells until the water was evaporated or absorbed.
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 into the ice. Its ready penetration is aided by the “dust-wells” which mark the surface of many glaciers. In north Greenland wells which contain six or eight inches of water at the end of a warm day are often dry in the morning. The water has leaked out and passed to lower levels. From these and other harmonious observations it is inferred that the superficial part of a glacier at least is readily penetrated by water. The depth to which surface water penetrates is undetermined, but it doubtless varies greatly, not only in different glaciers, but in different parts of the same glacier, and in the same part at different times. 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 there is much melting and much rain, and here it is probable that the water sometimes, perhaps usually, 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 (Fig. 251). 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 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 débris 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. These are the most common of the glacio-fluvial deposits.
The work accomplished by glaciers is distinctive, for while like rivers, they abrade the valleys through which they pass, carry forward the material which they remove from the surface, and wear, grind, and ultimately deposit it, and while their work therefore includes erosion, transportation, and deposition, their method is peculiar.
Getting load.—If the surface on which the snow-field which is to become a glacier accumulates be rough and covered with abundant rock débris, as such surfaces usually are, the glacier already has a basal load when its movement begins, 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. When the snow becomes ice and the ice begins to move, it carries forward the loose rock already imbedded in it, and tears off the weak points of the enveloped rock-projections. It may perhaps also move some of the soil and mantle rock of the original surface to which it is frozen. In addition to the subglacial load which the glacier thus has at the outset, there may be a surface load which has fallen on the snow or ice from cliffs above. This is especially true of mountain-valley glaciers. If this has been buried by snow and ice it is englacial; if it lies on the surface it is superglacial.
Once in movement, the ice carries away the débris to which it was originally attached, and at the same time gathers new load from the same area. The new load is acquired partly by the rasping effect of the rock-shod ice on its bed, and partly by its rending power which, under favorable circumstances, may quarry out considerable blocks of rock. This “plucking” process is at its best where the ice passes over cliffs of jointed rock or steep-sloped hills.
As the ice advances into new territory it acquires additional basal load, partly by rasping, partly by plucking, partly by freezing to it, and partly in other ways. One of these ways may be illustrated by the sequence of events when the end of a glacier advances on a very large bowlder. As the ice approaches it, the reflection of heat from it melts the adjacent edge of the ice, making a slight reëntrant. With farther advance, the ice closes in against and around the bowlder, and finally carries it along in the bottom of the moving mass. In some cases, especially when its advance is rapid, the ice may push débris in front of itself. Even where this is the case, the amount of material pushed forward is generally slight, partly because the extreme edge of the ice often fails to rest on the land in summer, when the movement is greatest, being melted from below by the heat of the surface over which it is spreading (see Fig. 252), and partly because the earth in front of the glacier is frozen during a large part of the year. In this condition, the earthy matter has greater resistance than the ice, and the latter rides over it. Superglacial material may be acquired during movement by the fall of débris from cliffs, or by the descent of avalanches.
Conditions influencing rate of erosion.—An obstructive attitude of the surface toward the movement of the ice is as necessary for effective erosion as the movement of the ice itself. Advancing over a flat surface, ice ordinarily inflicts but little wear, since there is little for it to get hold of. So slight is the abrasive power of ice under these conditions that it frequently overrides and buries the soil with more or less of its herbaceous vegetation. But while a certain measure of roughness of surface is favorable for glacial erosion, the topography may be so uneven as to seriously impede the ice. Erosion is probably at its maximum, so far as influenced by topography, when the roughness of the surface is such as to offer notable catchment for the basal ice, but not such as to impede its motion very seriously. The amount of relief favorable for the greatest erosion increases with increasing thickness of the ice.
Other conditions which influence erosion by ice are (1) the amount of loose or slightly attached débris on the surface, (2) the resistance of the rock, (3) the slope of the surface, (4) the thickness of the ice, (5) the rate of movement, and (6) the abundance and character of the débris which the ice has to work with. The effect of the first five of these conditions is evident. The effect of the last is less simple. Clean ice passing over a smooth surface of solid rock has little effect upon it; but a rock-shod glacier will abrade the same surface notably. The effect of this abrasion is shown in the grooves and scratches (striæ) which the stones in the bottom of the ice inflict on the surface of the rock over which they pass (Figs. 253, 255, and 256). At the same time the stones in the ice are themselves worn both by abrasion with the bottom, and with one another (Fig. 254). It does not follow, however, that the more material in the bottom of the ice the greater the erosion it effects; for with increase of débris there may be decrease of motion[131] and, beyond a certain point, the decrease of motion seriously interferes with the efficiency of erosion. When any considerable thickness of ice at the bottom of the glacier is full of débris, this loaded basal portion may approach stagnancy, and the lower limit of considerable movement may lie between the loaded ice below and the relatively clean ice above. A moderate, but not an excessive load of débris is, therefore, favorable for great erosion. Something depends, too, on the character of the load. Coarse, hard, and angular débris is a more effective instrument of erosion than fine, soft, or rounded material. The adverse influence of the overloading of the ice on its motion has been likened to the stiffening of a viscous liquid by the addition of foreign matter, but it may better perhaps be referred to the destruction of the granular-crystalline continuity on which glacier motion probably depends.
From the preceding statement, it is evident that erosion is not equally effective at all points beneath a glacier. 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 may here be thick, the movement is slow and the load likely to be slight. Ice conditions only being considered, erosion is most effective somewhere between the source and the terminus, and probably much nearer the latter than the former. The conditions of the surface over which the ice passes may be such as to vary the place of greatest erosion widely. Thus in an Alpine glacier, erosion may be most effective at the Bergschrund because the slope here favors “plucking.” Here, notable amphitheatres (cirques) are sometimes excavated. After the glacier disappears, the bottom of the cirque is often seen to contain rock basins (Fig. 257). Glacial cirques abound in mountains where glaciers once existed, but from which they have now disappeared. The cirques of the Bighorn mountains of Wyoming (Pl. XIX) are examples.
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 incoherent or soft formations possessing a topography of sufficient relief to offer some resistance, and yet too little to retard seriously the progress of the ice, will erode most effectively.
Varied nature of glacial débris.—From its mode of erosion it will readily be seen that the bottom of a glacier may be charged with various sorts of material. 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 débris 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 (drift) are of all grades of coarseness and fineness, from large bowlders to fine clay. The coarser material may be angular or round at the outset, and its form may be changed and its surface striated as it is moved forward. Whether one sort of material or another predominates, depends primarily on the nature of the surface overridden.
The topographic effects of glacial erosion.—In passing through its valley, an alpine glacier deepens and widens its bottom and smooths its slopes up to the upper limit of the ice. It tends to change a V-shaped valley (Fig. 258) into a U-shaped one (Fig. 259). The change in topography at the upper limit of glaciation is often marked (Figs. 260 and 261).
The deepening of a valley by glacial erosion may throw its 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 (Figs. 262 and 263). Such valleys are of common occurrence in regions recently glaciated, but now ice-free. Examples are common in the western mountains of North America and elsewhere.
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. 264 and 265); but since 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 narrow bay may gouge out the bay to a very considerable depth, causing its head to recede. When the ice finally melts, the bay may be a fiord. Thus have arisen the glacial features of many of the fiords of high-latitude coasts, and many of the glaciers of those coasts are now making fiords (Fig. 266). Fiords also arise in other ways. Coasts indented by fiords are likely to be bordered by islands.
The positions in which débris is carried.—As a result of the methods by which a glacier becomes loaded, there are three positions in which the débris is carried: (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 base of the ice; while that which falls on the surface is usually carried in the form of surface moraines. In the former position the drift is basal; in the latter, superglacial. It is doubtful if much débris is moved along beneath (that is, strictly below the bottom of) the ice, though the movement of the latter would have a tendency to drag or urge along with it the loose material of its bed. If drift were carried forward in such positions, it would be strictly subglacial.
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 englacial and superglacial material, on the other hand, is normally borne from the place of origin to the place of deposition without such intermixture. It is a case of “local” versus “through” transportation.
Transfers of load.—While the origin of the load usually determines its position, exceptions and complications arise from the transfer of load from one position to another, and from the gradation of one horizon into another.
Most of the débris 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 (Fig. 267) the bottom of the ice rends débris from the top of the hill. When it descends from one level to another there is a similar result (Fig. 268). To the lee of the hill the ice from either side may close in under that which came over the top, in which case the débris derived from the top of the hill by the bottom of the overriding ice will be well up in the ice. It has passed from an initial basal to a subsequent englacial position. The change does not usually involve an actual rise of the material, but rather a decline. If carried upward at all, the upward movement is temporary only, and incident to the passage of the ice over the hill, or to other local causes. The englacial débris may be little or much above the basal zone according to the height of the elevation overridden.