AMERICAN SCIENCE SERIES—ADVANCED COURSE

GEOLOGY

BY
THOMAS C. CHAMBERLIN AND ROLLIN D. SALISBURY

Heads of the Departments of Geology and Geography, University of Chicago
Members of the United States Geological Survey
Editors of the Journal of Geology

IN THREE VOLUMES

VOL. I.—GEOLOGIC PROCESSES AND THEIR RESULTS

SECOND EDITION, REVISED

NEW YORK
HENRY HOLT AND COMPANY
1909

In the preparation of this work it has been the purpose of the authors to present an outline of the salient features of geology, as now developed, encumbered as little as possible by technicalities and details whose bearings on the general theme are unimportant. In common with most writers of text-books on geology, the authors believe that the subject is best approached by a study of the forces and processes now in operation, and of the results which these forces and processes are now bringing about. Such study necessarily involves a consideration of the principles which govern the activities of geologic agencies. These topics are presented in Volume I, and prepare the way for the study of the history of past ages, which is outlined in Volume II.

The general plan of the work has been determined by the experience of the authors as instructors. Little emphasis is laid on the commonly recognized subdivisions of the science, such as dynamic geology, stratigraphic geology, physiographic geology, etc. The treatment proceeds rather from the point of view that the science is a unit, that its one theme is the history of the earth, and that the discussions of dynamic geology, physiographic geology, etc., apart from their historical bearing, lose much of their significance and interest. The effort has, therefore, been to emphasize the historical element, even in the discussion of special themes, such as the work of rivers, the work of snow and ice, and the origin and descent of rocks. This does not mean that phases of geology other than historical have been neglected, but it means that an effort has been made to give a historical cast to all phases of the subject, so far as the topics permit.

Throughout the work the central purpose has been not merely to set forth the present status of knowledge, but to present it in such a way that the student will be introduced to the methods and spirit of the science, led to a sympathetic interest in its progress, and prepared to receive intelligently, and to welcome cordially, its future advances. Where practicable, the text has been so shaped that the student mayiv follow the steps which have led to present conclusions. To this end the working methods of the practical geologist have been implied as frequently as practicable. To this end also there has been frankness of statement relative to the limitations of knowledge and the uncertainty of many tentative conclusions. In these and in other respects, the purpose has been to take the student into the fraternity of geologists, and to reveal to him the true state of the development of the science, giving an accurate and proportionate view of the positive knowledge attained, of the problems yet unsolved, or but partially solved, and of solutions still to be attained.

The theoretical and interpretative elements which enter into the general conceptions of geology have been freely used, because they are regarded as an essential part of the evolution of the science, because they often help to clear and complete conceptions, and because they stimulate thought. The aim has been, however, to characterize hypothetical elements as such, and to avoid confusing the interpretations based on hypothesis, with the statements of fact and established doctrines. Especial care has been taken to recognize the uncertain nature of prevalent interpretations when they are dependent on unverified hypotheses, especially if this dependence is likely to be overlooked. If this shall seem to give prominence to the hypothetical element, it should also be regarded as giving so much the more emphasis to that which is really trustworthy, in that it sets forth more frankly that which is doubtful. Hypothetical and unsolved problems have been treated, so far as practicable, on the multiple basis; that is, alternative hypotheses and alternative interpretations are frequently presented where knowledge does not warrant positive conclusions.

In many cases the topics discussed will be found to be presented in ways differing widely from those which have become familiar. In some cases, fundamentally new conceptions of familiar subjects are involved; in others, topics not usually discussed in text-books are stated with some fullness; and in still others, the emphasis is laid on points which have not commonly been brought into prominence. Whether the authors have been wise in departing to this extent from beaten paths, the users of the volumes must decide.

The work is intended primarily for mature students, and is designed to furnish the basis for a year’s work in the later part of the collegev course. By judicious selection of material to be presented and omitted, the volumes will be found useful for briefer courses, and by the use of the numerous references to the fuller discussions of special treatises, they may be made the basis for more extended courses than are commonly given in undergraduate work. The attempt has also been made to make the volumes readable, in the belief that many persons not in colleges or universities will be interested in following a connected account of the earth’s history, and of the means by which that history is recorded and read. Antecedent elementary courses in geology will not be necessary to the use of these volumes, though such courses may be helpful.

The arrangement of themes adopted is such as to bring to the fore processes with which all students are immediately in contact, and which are available for study at all seats of learning. The commoner geologic agents, such as the atmosphere and running water, have been elaborated somewhat more fully than is customary, and the common rather than the exceptional phases of the work of these agents have been emphasized, both because of their greater importance and their universal availability. The text has been so shaped as to suggest field work in connection with these topics especially, since work of this sort is everywhere possible.

After the preliminary outline, which is intended to give some idea of the scope of the science, and of its salient features, and to show the relations of the special subjects which follow, the order of treatment is such as to pass from the commoner and more readily apprehended portions of the subject to those which are less readily accessible and more obscure. Following the same general conception, the treatment of the topics is somewhat graded, the earlier chapters being developed with greater simplicity and fullness, while the later are somewhat more condensed.

Many acknowledgments are made in the text and foot-notes, but it is impossible to adequately acknowledge all the sources which have been drawn upon, since the whole body of literature has been laid under contribution. The authors especially acknowledge the generous assistance of Professor J. P. Iddings in connection with the chapter on The Origin and Descent of Rocks; of Dr. F. R. Moulton, Professor C. S. Slichter, Professor L. M. Hoskins, Mr. A. C. Lunn, and Mr. W. H. Emmons in connection with mathematical problems; of Professor C. R.vi Barnes in connection with the geologic functions of life; and of Professor Julius Stieglitz in connection with chemical subjects.

The illustrations have been selected from numerous sources, which are usually acknowledged in the text. Especial acknowledgment is due to the U. S. Geological Survey for the use of numerous photographs and maps, and to Mr. G. A. Johnson, who has made many of the drawings reproduced in Volume I. The authors are under even larger obligations for assistance in the preparation of Volume II, for which acknowledgment will be made in the proper place.

University of Chicago, January, 1904.

Geology treats of the structure of the earth, of the various stages through which it has passed, and of the living beings that have dwelt upon it, together with the agencies and processes involved in the changes it has undergone. Geology is essentially a history of the earth and its inhabitants. It is one of the broadest of the sciences, and brings under consideration certain phases of nearly all the other sciences, particularly those of astronomy, physics, chemistry, zoology, and botany. It also embraces the earlier expressions of mental development and of life-relationships, chiefly as found in the lower animals.

Subdivisions.—Naturally so broad a science has many special aspects which constitute subdivisions, in a sense, though they are rather dominant phases than independent sections. That phase which treats of the outer relations of the earth is Cosmic or Astronomic geology; that which treats of the constituent parts of the earth and its material is Geognosy, of which the most important branch is Petrology, the science of rocks. That branch which investigates the structural arrangement of the material, or “the architecture of the earth,” is Geotectonic, or Structural geology; while that which deals with the surface changes and topographic forms, that is, with the face of the earth, is Physiographic geology. The study of the fossils that have been preserved in the rocks, and of the faunas and floras that these imply, constitutes Paleontologic geology, or Paleontology. The treatment of the succession of events forms Historical geology. This is chiefly worked out by the succession of beds laid down in the progress of the ages, which constitutes Stratigraphic geology. The treatment of causes, agencies, and processes is the function of Dynamic or Philosophic geology.

Besides these there are special applications which give occasion for other terms, as Economic geology, which is concerned with the industrial applications of geologic knowledge; Mining geology, which2 is a sub-section of economic geology, relating to the application of geologic facts and principles to mining operations; Atmospheric geology, Glacial geology, and others that define themselves, and are for the greater part but limited aspects of the broad science.

Dominant processes.—Three sets of processes, now in operation on the surface of the lithosphere, have given rise to most of the details of its configuration, and even many of its larger features. These processes have been designated diastrophism, vulcanism, and gradation. Diastrophism includes all crustal movements, whether slow or rapid, gentle or violent, slight or extensive. Many parts of the land, especially along coasts, are known to be slowly sinking relative to the sea-level, while other parts are known to be rising. The fact that rocks originally formed beneath the sea now exist at great elevations, and the further fact that areas which were once land are now beneath the sea, are sufficient evidence that similar changes have taken place in the past. Vulcanism includes all processes connected with the extrusion of lava and other volcanic products, and with the rise of lava from lower to higher levels, even if not extruded. Vulcanism and diastrophism may be closely associated, for local movements at least are often associated with volcanic eruptions, and more considerable movements may be connected with the movements of subsurface lavas, even when the connection is not demonstrable. Gradation includes all those processes which tend to bring the surface of the lithosphere to a common level. Gradational processes belong to two categories—those which level down, degradation, and those which level up, aggradation. The transportation of material from the land, whether by rain, rivers, glaciers, waves, or winds, is degradation and the deposition of material, whether on the land or in the sea, is aggradation. Degradation affects primarily the protuberances of the lithosphere, while aggradation affects primarily its depressions.

Astronomic Geology.

The earth as a planet.—Though supremely important to us, the earth is but one of the minor planets attendant upon the sun, and is in no very special way distinguished as a planetary body. Of the eight planets, four, Jupiter, Saturn, Uranus, and Neptune, are much larger than the earth, while three, Mars, Venus, and Mercury, are smaller. There are a host of asteroids, but all together they do not equal the mass of the smallest planet. The average mass of the eight planets3 is more than fifty times that of the earth, while the largest, Jupiter, is more than three hundred times as massive as the earth. The earth’s position in the group is in no sense distinguished. It is neither the outer nor the inner, nor even the middle planet. Even in the minor group to which it belongs, it is neither the outermost nor the innermost member, though in this group it is the largest. Its average distance from the sun is about 92.9 million miles, and this fixes its revolution at 365¼ days, for its period of revolution is directly dependent on its distance from the sun, and is necessarily longer than the revolutions of the inner planets and shorter than those of the outer planets. Its rotation in twenty-four hours is not far different from that of its neighbor Mars, but is much slower than the more distant and larger planets, Jupiter and Saturn, which rotate in about ten hours. Comparison cannot be made with the innermost and outermost planets, because their rotations are not yet satisfactorily determined. The plane of the earth’s revolution lies near the common plane of the whole system, but this is not peculiar, as all of the planets revolve in nearly the same plane. Only a few of the small asteroids depart notably from this common plane. This has an important bearing on theories of the origin of the system, since this close coincidence of the planes of the orbits is not consistent with any haphazard aggregation of the material. Of similar importance is the fact that all of the planets revolve in the same direction and in ellipses that do not depart widely from circles. The eccentricity of the earth’s orbit is only about ¹⁄₆₀. This eccentricity varies somewhat, due to the disturbing influences of the other planets, and this variation has been regarded by some geologists as an influential cause of climatic changes, but its adequacy to produce great effects has been doubted by others. The inclination of the earth’s axis, now about 23½°, holds an intermediate position, some of the planets having axes more inclined, as Saturn, 26⅚°, and others less inclined, as Jupiter, 3°. The inclination of the axis is subject to trivial variations at present, and in the long periods of the past has possibly changed more notably. This possible change has also been thought to be a cause of climatic variation, but its efficiency has not been demonstrated.

Its satellite.—The earth is peculiar in having one unusually large satellite, which has a mass ¹⁄₈₁ of its own. The great planets have several satellites whose combined mass exceeds that of the moon, and perhaps in some few cases the individual satellites may be larger than the moon, but they do not sustain so large a ratio to their planets,4 for Titan, probably the largest, is only ¹⁄₄₆₀₀ of the mass of Saturn. There is little doubt that the moon has played an important part in the history of the earth. It is the chief agency in developing oceanic tides, and it possibly also develops a body tide in the earth itself. These tides act as a brake on the rotation of the earth and tend to reduce its rate, and thereby to lengthen the day. While this may have been counteracted in some measure by the shrinkage of the earth, which tends to increase its rate of rotation, it has been held by eminent physicists and geologists that the rotation of the earth has been greatly lessened during its history, and that a long train of important consequences has resulted. If the contraction of the earth has been sufficient to offset this lessening, the tidal brake must be credited with the prevention of the excessive speed of rotation which would otherwise have been developed. The tides are efficient agencies in the shore wear of the oceans, and in the distribution of marine sediments, and these, it will be seen later, are important elements in the formation of strata.

Dependence on the sun.—By far the most important external relation of the earth, however, is its dependence on the sun. The earth is a mere satellite of the sun, less than ¹⁄₃₀₀₀₀₀ of its mass, and hence under its full gravitative control. The earth is dependent on the sun for nearly all its heat and light, and, through these, for nearly all of the activities that have given character to its history. It is too much to say that all activities on the surface of the earth are solely dependent on those of the sun, for a certain measure of heat and light and other energy is derived from other bodies, and a certain not inconsiderable source of energy is found in the interior of the earth itself; yet all of these are so far subordinate to that great flood of energy which comes from the sun that they are quite inconsequential. The history of the earth in the past has been intimately dependent upon that of the sun, and its future is locked up with the destiny of that great luminary. Geology in its broadest phases can therefore scarcely be separated from the study of the sun, but this falls within the function of the astronomer rather than the geologist.

Meteorites.—There are a multitude of small bodies passing through space in varying directions and with varying velocities and occasionally encountering the earth, to which they add their substance. Some of these meteorites revolve about the sun much as if they were minute planets, but some of them come from such directions and with such velocities as to show that they do not belong to the solar family. Some5 consist almost wholly of metal, chiefly iron alloyed with a small percent. of nickel (holosiderites); some consist of metal and rock intimately mixed (syssiderites and sporadosiderites); and some consist wholly of rock (asiderites). The rock is usually composed of the heavier basic minerals, though some meteorites consist largely of carbonaceous material. Besides meteorites, there is little doubt that wandering gaseous particles strike the earth, but this is beyond the reach of present demonstration. The amount of substance added to the earth by these meteorites and gases in recent times is relatively slight compared with the whole body of the earth. What contribution may have come to the earth in earlier times from such sources is a matter of hypothesis which will be discussed later.

Geognosy.

The constitution of the earth.—Turning from its external relations to the earth itself, a natural threefold division is presented: (1) the atmosphere, (2) the hydrosphere, and (3) the lithosphere.

I. The Atmosphere.

The atmosphere is an intimate mixture of (1) all those substances that cannot take a liquid or solid state at the temperatures and pressures which prevail at the earth’s surface, together with (2) such transient vapors as the various substances of the earth throw off. The first class form the permanent gases of the atmosphere, and consist of nitrogen about 79 parts, oxygen about 21 parts, carbon dioxide about .03 part, together with small quantities of argon, neon, xenon, krypton, helium, and other rare constituents. The second class are the transient and fluctuating constituents of the atmosphere, chief among which is aqueous vapor, which varies greatly in amount according to temperature, pressure, and other conditions. To this are to be added volcanic emanations and a great variety of volatile organic substances. Theoretically, every substance, however solid, discharges particles which may transiently become constituents of the atmosphere. Practically, only a few of these exist in such quantity as to be appreciable. Dust and other suspended matter are usually regarded as impurities rather than constituents of the atmosphere, but they play a not unimportant part by affecting its temperature and luminosity, and by facilitating the condensation of moisture.

6

Mass and extent.—The total mass of the atmosphere is estimated at five quadrillion tons, or ¹⁄₁₂₀₀₀₀₀ of the mass of the earth. It is relatively dense at the surface of the earth and decreases in density outwards in a manner difficult of absolute determination, so that the actual height of the appreciable atmosphere is not positively known. The true conception of the atmosphere is perhaps that of a tenuous envelope exerting a pressure of about fifteen pounds per square inch at the sea-level, and thinning gradually upwards until it reaches a tenuity which is inappreciable, but perhaps not ceasing absolutely until the sphere of gravitative control of the earth is passed, about 620,000 miles from the lithosphere. In the lower portion, according to the kinetic theory of gases, the molecules fly to and fro, colliding with each other with almost inconceivable frequency, and with very short paths between successive collisions, but in the upper rare portion some of the molecules bound outwards, and do not strike other molecules, and hence pursue long elliptical paths until the gravity of the earth overcomes their momentum, when they return, perhaps to bound off again or to force other molecules to do so. This fountain-like nature of the outer part of the atmosphere makes any sharp definition of its limit impracticable. Some molecules are believed to be shot away at such speed that they do not return. Beyond about 620,000 miles from the surface of the lithosphere, the differential attraction of the sun is greater than that of the earth, and if the attraction of the earth does not turn the molecules back before reaching this distance, they are almost certain to be lost to the earth.

The measurement of heights by the aneroid barometer, which is much used in practical geology, is dependent on the lessening of pressure as the instrument is carried upward.

Geologic activity.—The atmosphere is the most mobile and active of the three great subdivisions of the earth, and when its indirect effects through the agency of water, as well as its direct effects, are considered, it is to be regarded as one of the most effective agencies of change. It acts chemically upon the rock substance of the earth, causing induration in some instances, but more often inducing disintegration and change of composition by means of which rock is reduced to soil, or soil-like material, and rendered susceptible of easy removal by winds and waters. When in motion the atmosphere acts mechanically on the surface of the earth, transporting dust and sand, and by the friction of these it abrades the surface. It is chiefly effective, however, in furnishing the7 conditions for water action. Partly by its mechanical aid, but chiefly by securing the right temperature, it is a necessary factor in the action of rains, streams, glaciers, and the various forms of moving water upon land. So also, on the ocean, wave action is essentially dependent on the winds. In the absence of atmospheric propulsion, wave action would be chiefly confined to the tides and to occasional earthquake impulses, and would lose nearly all its efficiency. Stream action and wave action, which are the most declared of the geological agencies, are therefore to be credited as much to the atmosphere as to the hydrosphere, since the action is a joint one to which both envelopes are essential.

A thermal blanket.—A function of the atmosphere of supreme importance is the thermal blanketing of the earth. In its absence the heat of the sun would reach the surface with full intensity, and would be radiated back from the surface almost as rapidly as received, and only a transient heating would result. During the night an intensity of cold would intervene scarcely less severe than the temperature of space. In penetrating the atmosphere certain portions of the radiant energy of the sun are absorbed. Of the remainder which reaches the surface of the earth, a part is transformed into vibrations of lower intensity, which are then more effectively retained by the atmosphere. The air thus distributes and equalizes the temperature. The two constituents of the atmosphere which are most efficient in this work are aqueous vapor and carbon dioxide, and the climate of the earth is believed to have been very greatly affected by the varying amounts of these constituents in the atmosphere, as well as by the total mass of the atmosphere.

The function of the atmosphere in sustaining life and promoting all that depends on life is too obvious to need comment.

The special geological action of the atmosphere will be discussed in the next chapter.

II. The Hydrosphere.

About 1300 quadrillion tons of water lie upon the surface of the solid earth. This equals about ¹⁄₄₅₄₀ part of the earth’s mass. Were the surface of the solid earth perfectly spheroidal, this would constitute a universal ocean somewhat less than two miles deep. Owing to the inequalities of the rock surface, the water is chiefly gathered into a series of great basins or troughs occupying about three-fourths (72%)8 of the earth’s surface. These basins are all connected with each other and act as a unit, so that anything which changes the level of the water in one changes the level of all. This helps to make a common record of all great movements of the earth’s body, for the level of the ocean determines where the detritus from the land shall lodge, and hence where the edge of the marine beds shall be formed. This will appear more clearly when the formation of marine strata is discussed.

Oceanic dimensions.—The surface area of the ocean is estimated by Murray at 143,259,300 square miles. Of this, somewhat more than 10,000,000 square miles lie on the continental shelf, i.e., lap up on the borders of the continental platforms. This shows that the great basins are somewhat more than full. If about 600 feet of the upper part of the ocean were removed, the true ocean basins would be just full, and the surfaces of the true continental platforms would be dry land. The area of the true oceanic basins is about 133,000,000 square miles, and that of the true continental platforms about 64,000,000 square miles. Under about 20% of the ocean area, the bottom sinks to depths between 6000 and 12,000 feet; under about 53% it sinks to depths between 12,000 and 18,000 feet; and under the remaining 4% it ranges from 18,000 feet down to about 30,000. The last includes those singular sunken areas known as “deeps,” and sometimes called anti-plateaus, as they extend downward from the general ocean bottom much as the plateaus protrude upwards from the general land surface.

Besides the ocean, the hydrosphere includes all the water which constitutes the surface streams and lakes, together with that which permeates the pores and fissures of the outer part of the solid earth; but altogether these are small in amount compared with the great ocean mass.

Geologic activity.—Of all geological agencies water is the most obvious and apparently the greatest, though its efficiency is conditioned upon the presence of the atmosphere, upon the relief of the land, and upon the radiant energy of the sun. Through the agency of rainfall, of surface streams, of underground waters, and of wave action, the hydrosphere is constantly modifying the surface of the lithosphere, while at the same time it is bearing into the various basins the wash of the land and depositing it in stratified beds. It thereby becomes the great agency for the degradation of the land and the building up of the basin bottoms. It works upon the land partly by dissolving soluble portions of the rock substance, and partly by mechanical action. The9 solution of the soluble part usually loosens the insoluble, and renders it an easy prey of the surface waters. These transport the loosened material to the valleys and at length to the great basins, meanwhile rolling and grinding it and thus reducing it to rounder forms and a finer state, until at length it reaches the still waters or the low gradients of the basins and comes to rest. The hydrosphere is therefore both destructive and constructive in its action. As the beds of sediment which it lays down follow one another in orderly succession, each later one lying above each earlier one, they form a time record. And as relics of the life of each age become more or less imbedded in these sediments, they furnish the means of following the history of life from age to age. The historical record of geology is therefore very largely dependent upon the fact that the waters have thus buried in systematic order the successive life of the ages. Aside from this, the means of determining the order of events of the earth’s history are limited and more or less uncertain.

The special processes of the hydrosphere in its various phases will be the subject of discussion hereafter (Chaps. III, IV, VI). Suffice it here to recognize its great function in the constant degradation of the land, and in the deposition of the derived material in orderly succession in the basins.

Chief horizons of activity.—The great horizons of geological activity are (1) the contact zone between the atmosphere and the hydrosphere, chiefly the surface of the ocean, (2) the contact zone between the hydrosphere and the lithosphere, chiefly the shore belts, and (3) the contact zone of the atmosphere and surface waters, with the face of the continents. It is in these three zones that the greatest external work is being done and has been done in all the known ages.

III. The Lithosphere.

The atmosphere and hydrosphere are rather envelopes or shells than true spheres, though in some degree both penetrate the lithosphere. The lithosphere, on the other hand, is a nearly perfect oblate spheroid with a polar diameter of 7899.7 miles, and an equatorial diameter of about 26.8 miles more. Its equatorial circumference is 24,902 miles, its meridional circumference 24,860 miles, its surface area 196,940,700 square miles, its volume 260,000,000,000 cubic miles, and its average specific gravity about 5.57. The oblateness of the spheroid is an accommodation to the rotation of the earth, the centrifugal force at the10 equator being sufficient to cause the specified amount of bulging there. Computations seem to indicate that the accommodation is very nearly what would take place if the earth were in a liquid condition, from which the inference has been drawn that it must have been in that condition when it assumed this form, and must have continued essentially liquid until it attained its present rate of rotation, since, if the earth once rotated at a much higher speed, the flattening at the poles and the bulging at the equator must have been correspondingly greater. It is thought by others, however, that the plasticity of the earth is such that it would at all times assume a close degree of approximation to the demands of rotation, even if the interior were in a solid condition. By still others it is thought that the contraction of the earth has tended to accelerate the rotation about as much as the tides have tended to retard it, and that it has undergone little change of form.

Irregularities.—It is only in a general view, however, that there is a close approximation to a perfect spheroidal surface. In detail there are very notable variations from it. Geodetic surveys seem to have shown that the equatorial diameters are not all equal, even when the measurements are reduced to sea-level, but research along this line has not reached a sufficient stage of completeness to permit satisfactory discussion. It is, however, highly probable that the ocean surface as well as the average land surface is warped out of the perfect spheroidal form to some notable degree. This is very likely due to inequalities in the density of the earth’s interior. The fact that the larger portion of the water is gathered on one side of the globe, while the land chiefly protrudes on the opposite side, is very possibly due to unequal specific gravity in the interior of the earth.

The most obvious departure from a spheroidal form is found in the protrusion of the continents and in the sinking away of the earth surface under the oceans. As these inequalities present themselves to-day, they are known as continental platforms and ocean basins. These do not correspond accurately with the present land and water surfaces. About the continental lands there is a submerged border extending some distance out from the shore, and constituting a sea-shelf beyond which the surface descends rapidly to the great depths of the ocean. This slightly submerged portion, known as the continental shelf, belongs as properly to the continent as the adjacent low lands which are not submerged. The submergence of the edge of this shelf at present is usually about 100 fathoms, so that if the upper 600 feet 11of the ocean were removed, the outlines of the land would correspond quite closely with the border of the true continental platform.

It is customary to look upon the protrusions of the continents as the great features of the earth’s surface, but in reality the oceanic depressions are the master phenomena. In breadth, depth, and capacity they much exceed the continental protrusions, and if the earth be regarded as a shrunken body, the settling of the ocean bottoms has doubtless constituted its greatest surface movement. From the estimates of Murray, Gilbert has derived the following tables, showing the relative areas of the lithosphere above, below, and between certain levels.[1]

From these estimates it appears that if the surface were graded to a common level by cutting away the continental platforms and dumping the matter in the abysmal basins, the average plane would lie somewhere near 9000 feet below the sea-level. The continental platform may be conceived as rising from this common plane rather than from the sea-level.

Epicontinental seas.—Those shallow portions of the sea which lie upon the continental shelf, and those portions which extend into the interior of the continent with like shallow depths, such as the Baltic 12Sea and the Hudson Bay, may be called epicontinental seas, for they really lie upon the continent, or at least upon the continental platform; while those other detached bodies of water which occupy deep depressions in the surface are to be regarded as true abysmal seas, as, for example, the Mediterranean and Caribbean seas and the Gulf of Mexico, whose bottoms are as profound as many parts of the true ocean basin itself.

Diversities of surface.—The bottoms of the oceanic basins are diversified by broad undulations which range through many thousands of feet, but they are not carved into the diversified forms that give variety to land surfaces. The ocean bottoms are also diversified by volcanic peaks, many of which rise to the surface and constitute isolated islands. Some of them have notable platforms at or near the surface, cut by the waves or built up by the accumulation of sediment and of coralline and other growths about them. Aside from these encircling platforms, the solid surface usually shelves rapidly down to abysmal depths, so that the islands constitute peaks whose heights and slopes would seem extraordinary if the ocean were removed.

The surface of the land is diversified in a similar way by broad undulations and volcanic peaks, and also by narrower wrinklings and foldings of the crust; but all of these irregularities have been carved into diversified and picturesque forms by subaërial erosion. In this respect the surface of the land differs radically from the bed of the sea. The agencies which have produced the continental platforms and abysmal basins, and the great undulations and foldings, as well as the volcanic extrusions that mark them, are yet subjects of debate. Here lie some of the most difficult problems of geology, but these cannot be stated with sufficient brevity to find a place here.

The surface mantle of the lithosphere.—The surface of the lithosphere is very generally mantled by a layer of loose material composed of soil, clay, sand, gravel, and broken rock. This loose material is sometimes known as mantle rock, and sometimes as rock waste. On the land, mantle rock is often composed of the disintegrated products of underlying rock formations. It represents the results of the recent action of the atmosphere, of water, of changes of temperature, and of other physical agencies acting on the outer part of the rock sphere. The surface of this mantle is being constantly removed by wind and water, but as constantly renewed by continued decomposition of the rock below. In some areas, especially in the northern part13 of North America and the northwestern part of Europe, the soil graduates down into an irregular sheet of mixed clay, sand, gravel, and bowlders, known as drift. From this and other evidence it is inferred that at a time not greatly antedating our own, ice, chiefly in the form of glaciers, spread extensively over the high latitudes of the northern hemisphere. In some parts of the earth the surface is still covered by fields of snow and ice, comparable to those which formed the drift. In still other places, especially along the flood plains of streams, the mantle rock consists of deposits made by streams which were unable to carry their loads of sediment to the sea.

The crust of the lithosphere.—Much of the detritus washed down from the land finds its way to bodies of standing water, and beneath lakes and seas the mantle of loose material is made up largely of the gravel, sand, and mud derived from the land. Before deposition these materials are more or less assorted and arranged in layers by waves and currents. When consolidated they constitute rock. The weathering of the rocks of the land, the wearing away of the resulting detritus, and its deposition beneath standing water, are among the most important processes of geologic change.

On the land, the mantle of loose material is sometimes absent, and in such places the surface of solid rock of the crust appears. Bare surfaces of rock are most commonly seen where the topography is rough, especially on the slopes of steep-sided valleys and mountains, and on the slopes of cliffs which face seas or lakes. Solid rock, without covering of soil or loose material of any sort, is also frequently seen in the channels of streams, especially where there are falls or rapids.

We have but to note the effects of a vigorous shower on a steep slope, or of a swift stream on its channel, or of waves on the cliffs which face lakes and seas, to understand at least one of the reasons why loose materials are frequently absent from steep slopes. The very general exposure of solid rock where conditions favor surface erosion suggests that rock is everywhere present beneath the soil or subsoil. Fortunately there is an easy way of testing the universality of the crust beneath the mantle. In all lands inhabited by civilized peoples there are numerous wells and other excavations ranging from a few feet to several hundred feet in depth, and occasional wells and mine-shafts reach depths of several thousand feet. Even in shallow excavations rock is often encountered, and in most regions excavations as much as two or three hundred feet deep usually reach rock, and no really deep boring14 has ever failed to find it. It may, therefore, be accepted as a fact that the upper surface of the solid rock is nowhere far below the surface.

Concerning the thickness of the crust, if there be any true crust at all, little is known by direct observation. The deepest valleys, such as the canyon of the Colorado, and the shafts and borings of the deepest mines and wells, give knowledge of nothing but rock. The deepest excavations extend rarely more than a mile below the surface. It is certain that rock of known kinds extends to far greater depths.

The interior.—Concerning the great interior of the earth, little is known except by inference. From the weight of the earth,[2] it is inferred that its interior is much more dense than its surface. From its behavior under the attraction of other bodies, it is believed to be at least as rigid as steel, and its interior cannot, therefore, be liquid, in the usual sense of that term. From the phenomena of volcanoes, and from observations on temperature in deep borings, it is inferred that its interior is very hot. Further inferences concerning its character are less simply stated, and will be referred to later.

The solid part of the earth is therefore composed of (1) a thin layer of unconsolidated or earthy material, a few feet to a few hundred feet in thickness, covering (2) a layer or zone many thousands of feet, and probably many miles, thick, composed of solid rock comparable to that exposed at numerous points on the surface, and (3) a central mass, to which the preceding layers are but a shell, composed of hot, dense, and rigid rock, the real nature of which is not known by observation.

Varieties of rock in crust.—If the mantle of soil, subsoil, and glacial rubbish were stripped from the land, the surface beneath would be found to be made up of a great variety of rocks, all of which may be grouped into two great classes. About four-fifths of the land surface would be of rock arranged in layers, and the other fifth would be of crystalline rock, generally without distinct stratification, and often bearing evidence of the effects of high temperature.

Stratified rocks.—The composition of most stratified rocks corresponds somewhat closely with the composition of sediments now being carried from the land and being deposited in the sea. Their arrangement in layers is the same, and the markings on the surfaces of the layers, such as ripple-marks, rill-marks, wave-marks, etc., are identical. Furthermore, the stratified rocks of the land, like the recent sediments 15of the sea, frequently contain the shells and skeletons of animals, and sometimes the impressions of plants. Most of the relics of life found in the stratified rocks belonged to animals or plants which lived in salt water. Because of their structure, their composition, their distinctive markings, and the remains of life which they contain, it is confidently inferred that most of the stratified rocks which lie beneath the mantle rock of the land were originally laid down in beds beneath the sea, and that the familiar processes of the present time furnish the key to their history.

Conformability.—When the stratified rocks exposed by the removal of the mantle rock are examined, the successive beds are sometimes found to lie on one another in regular succession, showing that they were laid down one after another, without change in the attitude of the surface on which they were deposited. Such rocks are conformable (the beds of series a, Fig. 1). In other cases it would be seen that certain beds overlie the worn surfaces of lower beds, the layers of which may have a different angle of inclination (series a, Fig. 1, is unconformable on series b). Such relations show that the lower series of beds was disturbed and eroded before the overlying beds were deposited on them. Such series of rocks are unconformable.

Relative ages.—The structure and relations of rocks lead to inferences as to their relative ages. In the case of stratified rocks it is obvious that overlying beds were deposited later than those below, and where there is unconformity it is evident that an interval of time elapsed between the deposition of the unconformable series. Another and in some respects more important means of telling their order of16 formation is found in the remains of life entrapped in the water-laid sediments. Whatever life existed in the waters in which the sediments were deposited was liable to burial, and if it was possessed of hard parts, such as bones, teeth, shells, hard integuments, etc., these parts, or at least their impressions, were likely to be preserved in the sediments. Even tracks and imprints of perishable parts are sometimes preserved. All these relics, which we call fossils, give indications of the kinds of life which existed when the beds were formed. The fossils of the youngest beds show that the life which existed when they were deposited was quite like that of the present time. The fossils of the next older and lower beds show greater departure from present types. This series of changes continues downward as lower and lower beds are studied, until beds at considerable depths contain no relics of existing species but, in lieu thereof, forms of more primitive types. Some of these earlier types are clearly the ancestors of more modern forms, while others seem to have no living descendants. Going still deeper, the fossils indicate life of more and more primitive types, until they depart very widely from the living forms, and seem to be but remotely ancestral. So the beds may be followed downward until the lowest, which contain distinct evidences of life, are reached.

It should be understood that it is not possible to proceed directly downward through the whole succession of bedded rocks, but that the edges of the various beds may be found here and there where they have been brought to the surface by warpings or tiltings, or exposed by the wearing away of the beds which once overlay them. The full series of strata is made out only by putting together the data gathered throughout all lands, and even when this is done an absolutely complete series cannot yet be made out or, at least, has not been.

The crystalline rocks.—The crystalline rocks which would appear if the mantle rock were removed are of two types, igneous and metamorphic. Igneous rocks may be loosely defined as hardened lavas. Metamorphic rocks are those which are greatly changed from their original condition. Either stratified or igneous rocks may become metamorphic.

Igneous rocks sustain various relations to the stratified rocks, as illustrated by Fig. 2. From these relations it is possible to tell something of the order of their formation. Where the stratified rocks are broken through by lavas, it is obvious that the stratified rocks were formed first, and the lavas intruded later. Lava sheets intruded17 between beds of stratified rock can be told from those which flowed out on the surface and were subsequently buried, for in the former case the sedimentary rocks, both above and below the igneous rock, were affected by the heat, while in the latter case only those below were so affected.

More commonly than otherwise the metamorphic rocks (Fig. 3) lie beneath the sedimentary beds and are often broken through by the igneous rocks. From their position in many places their great age may be inferred, but locally, especially where dynamic action has been severe, relatively young rocks are metamorphic.

Four great sedimentary eras.—The water-laid series represents four great eras in the history of the earth, as shown by the relics of life imbedded in them. Beginning with the latest, these are the Cenozoic (recent life), during which the life took on its modern aspect; the Mesozoic (middle life), during which the life bore a mediæval aspect; the Paleozoic (ancient life), during which the life belonged to older types; and the Proterozoic (earlier life), during which it is inferred that much life prevailed, though its record is very imperfect. It may safely be assumed to have been more primitive than that of the Paleozoic, as it was earlier. Each of these great divisions embraced several lesser periods or epochs, and these again are subdivided more and more closely according to the degrees of refinement to which studies are carried. The chief of these subdivisions are given in the table on page 19, and others will come under consideration in the historical chapters.

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In these four great series of sedimentary rocks there are, here and there, intrusions of igneous rocks, and in some places the sedimentary beds have been metamorphosed into crystalline rocks by heat and pressure. This is particularly true in the lowest of these series, the Proterozoic, where a large part of the sediment is metamorphosed, and where there is much igneous rock, but it is still clear that the main portion of this series was originally water-laid sediment, and so it belongs to the sedimentary series rather than the Archean, in which the sediments are the minor rather than the main factor. It has, however, usually been classed with the Archean, and it is certainly not always easy to draw the dividing line. In a sense it may be regarded as a transition series.

The Archean complex.—Beneath the dominantly sedimentary but partly metamorphic and igneous series there is a very complex group of rocks largely of metamorphosed igneous origin, though containing some metamorphosed sediments. These extend downwards to unknown depths. While all the great formations are occasionally bent and broken, these lowest ones are almost everywhere warped, folded, and contorted, often in the most intricate way. They have been very generally mashed and sheared by enormous pressure, so that they have become foliated, and their original character is much masked. They therefore form a series of great obscurity and complexity. As they are at the bottom of the known series, they have been called the “Fundamental gneiss” and the “Basement complex,” but as the part which we see is not the true base nor the true foundation, it is safer to call them simply the Archean (very ancient) complex. As life appears to have been present during a part at least of the period of its formation is referred to the Archeozoic era.

Beyond and below this series, the structure of the earth is a matter of inference. Vast as are the preceding series, they together form relatively but a thin shell on the outer surface of the globe.

The foregoing series are diagrammatically expressed in Fig. 4, and systematically presented to the eye in the following table.

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GENERAL TABLE OF GEOLOGIC DIVISIONS.

Cenozoic		Present.
		Pleistocene.
		Pliocene.
		Miocene.
		Oligocene.
		Eocene.
		Transition (Arapahoe and Denver).
Mesozoic		Upper Cretaceous.
		Lower Cretaceous (Comanche or Shastan).
		Jurassic.
		Triassic.
Paleozoic		Permian.
		Coal Measures or Pennsylvanian.
		Subcarboniferous, or Mississippian.
		Silurian.
		Devonian.
		Ordovician.
		Cambrian.
		Great interval.
Proterozoic		Keweenawan.
		Interval.
		Animikean or Penokean. (Upper Huronian of some authors).
		Interval.
		Huronian.
		Great interval.
Archeozoic		Archean Complex.		Great Granitoid Series. (Intrusive in the main, Laurentian.)
				Great Schist Series. (Mona, Kitchi, Lower Keewatin, Coutchiching, Lower Huronian of some authors.)

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The purpose of this general survey is to bring the salient features of the earth’s structure into view preparatory to entering in more detail into the study of particular processes and special formations and to lay a foundation for the fuller apprehension of the successive stages of the history of the earth, which constitutes the chief purpose of geological study. It is now advisable to turn to the detailed consideration of individual processes and specific structures. The complexity of the actions involved in the history of the earth is so great that such separate consideration at the outset is helpful.

While it is convenient to regard the lithosphere as the earth proper, and the atmosphere as its envelope, the latter is as truly a part of the planet as the former, and its activities and its history are as truly subjects of geological study as the formation of the rocks. This view is in no way vitiated by the fact that the special study of the atmosphere is set apart under the name Meteorology, for in the same way the special study of rocks is set apart under the name Petrology, that of ancient life under the name Paleontology, and that of other phases of the subject under other names. The atmosphere is one of the three great formations of the earth, and as a geological factor takes its place beside the hydrosphere and the lithosphere. It has played a part in the history of the earth comparable to that of the water, though its mass is less and its record more elusive. Unsubstantial as the atmosphere seems when contrasted with the liquid and solid portions of the earth, its extreme mobility and its chemical activity compensate for its lightness and tenuity, and give it a function of the first order of importance.

The atmosphere plays a direct part as (1) a mechanical and (2) a chemical agent, and at the same time serves an indirect function in furnishing favorable conditions under which (3) solar radiation produces temperature effects, and (4) evaporation gives origin to precipitation and stream effects, and furnishes the necessary conditions for land plants and animals, and the important influences that spring from them.

This chapter is devoted to the work of the atmosphere in these and some less notable phases. The consideration of the origin and history of the atmosphere will receive attention later.

A. THE ATMOSPHERE AS A DIRECT AGENCY.

I. Mechanical Work.

The mechanical work of the atmosphere is accomplished chiefly through its movement. A feeble breeze is competent to move particles22 of dust, and winds of moderate velocity to shift sand. Exceptionally strong winds sometimes move small pebbles, but winds of sufficient force to move larger pieces of rock are rare. It follows, therefore, that the impact of the wind has little direct effect except on surfaces covered with dust and dry sand.

The transportation of material by the wind is limited by the size of the particles to which it has access. Dust particles expose more surface to the wind relative to their mass than sand grains. Winds which are unable to carry sand may still carry dust, and winds which are able to shift sand no more than a trivial distance may blow dust great distances.

The common conception of wind as a horizontal movement of some part of the atmosphere is not altogether accurate. Every obstacle against which wind blows causes deflections of its currents, and some of these deflections are upward. Furthermore, there are exceptional winds, in which the vertical element predominates. Particles of dust are often involved in these upward currents, and by them carried to great heights, and in the upper air are transported great distances.

Transportation and deposition of dust.[3]—The universality of the transportation of dust by the wind is well known. No house, no room, and scarcely a drawer can be so tightly closed but that dust enters it, and the movements of dust in the open must be much more considerable. The visible dustiness of the atmosphere in dry regions during wind-storms is adequate and familiar proof of the efficiency of the wind as a transporter of dust.

Under special circumstances, opportunity is afforded for rough determinations of the distance and height to which wind-blown dust is transported. Snow taken from snow-fields in high mountain regions is found to contain a small amount of earthy matter. Its particles are often found to be in part volcanic, even when the place whence the snow was taken is scores or even hundreds of miles from the nearest volcano. There is probably no snow-field so high, or so far from volcanoes, but that volcanic dust reaches it. If this be true of all snow-fields, it is probably true of all land surfaces. In the great Krakatoa[4] eruptions of 1883 large quantities of volcanic ash (pulverized lava) were projected to great heights into the atmosphere. The coarser particles 23soon settled; but, caught by the currents of the upper atmosphere, many of the finer particles were transported incredible distances. Through all their long journey, the particles of dust were gradually settling from the atmosphere, but not until the dust had traveled repeatedly round the earth did its amount become so small as to cease to make its influence felt in the historic red sunsets which it occasioned.[5] Some of this dust completed the circuit of the earth in 15 days.

In various parts of Kansas and Nebraska[6] there are very considerable beds of volcanic dust, locally as much as 30 feet thick, which must have been transported from volcanic vents by the wind, though there are no known centers of volcanic action, past or present, within some hundreds of miles of some of the localities where the dust occurs. These beds of volcanic dust, so far from its source, may serve as an illustration of the importance of atmospheric movements as a geological force.

Volcanic dust is shot into the atmosphere rather than picked up by it. Dust picked up by the wind is perhaps transported not less widely than volcanic dust, but, after settling, its point of origin is less readily determined. It would perhaps be an exaggeration to say that every square mile of land surface contains particles of dust brought to it by the wind from every other square mile, but such a statement would probably involve much less exaggeration than might at first be supposed.

Examples of extensive deposits of dust other than volcanic are also known. In China there is an extensive earthy formation, the loess, sometimes reaching 1,000 feet in thickness, which von Richtofen believes to have been deposited by the wind.[7] This conclusion has, however, not passed unchallenged.[8] The loess of some other regions has been referred to the same origin, and some of it is quite certainly eolian.[9]

The transportation of dust is important wherever strong winds blow over dry surfaces, free or nearly free of vegetation, and composed of earthy matter. Its effects may be seen in such regions as the sage-brush plains of western North America. The roots of the 24sage-brush hold the soil immediately about them, but between the clumps of brush, where there is little other vegetation, the wind has often blown away the soil to such an extent that each clump of brush stands up several inches, or even a foot or two, above its surroundings (Fig. 5). Such mounds are often partly due to the lodgment of dust about the bushes.

Where the earthy matter is moist, the cohesion of the particles is great, and the wind cannot pick them up. Furthermore, if the surface is generally moist, it is likely to be covered with vegetation which protects it against the wind. But even where vegetation is prevalent the wind finds many a vulnerable point. Thus on the edges of plains or plateaus facing abrupt valleys, the wind attacks the soil from the side, and in such situations all earthy matter may be stripped25 from the underlying rock for considerable distances from the edge of the cliff (Fig. 6). This may be seen at numerous points on the lava plateaus of Washington.

The presence of dust in the upper atmosphere during a rain-storm is sometimes the occasion of phenomena which are often misinterpreted. If there be abundant dust in the atmosphere through which rain-drops or snowflakes fall, much of it is gathered up by them, and the water is thereby rendered turbid and the snow discolored. Here is to be found the explanation of “mud-rain,” “blood-rain” (red dust), etc.

Since dust is carried to a considerable extent in the upper atmosphere, its movements and its deposition are little affected by obstacles on the surface of the land. A building or a hedge can only affect the lodgment of that part of the atmospheric dust which comes in contact with it or is swept into its lee. Since most obstacles on the surface of the solid part of the earth reach up but slight distances into the atmosphere, the dust of the greater part of the air settles without especial reference to them, and is spread more or less uniformly over the surface on which it falls.

Much of the dust transported by the wind is carried out over seas or lakes and falls into them. By this means, sedimentation is doubtless going on at the bottom of the whole ocean, and at the bottoms of all lakes. While means of determining the amount of dust blown into the sea are not at hand, it is safe to say that, were such determinations possible, the result, if stated in terms of weight, would be surprising.

Transportation and deposition of sand.—In its transportation by the wind, sand is not commonly lifted far above the surface of the26 land, and its movement is therefore more generally interfered with by surface obstacles than is the movement of dust. A shrub, a tree, a fence, a building, or even a stone may occasion the lodgment of sand in considerable quantity, though it has little effect on the lodgment of dust. The effect of obstacles is illustrated by Fig. 7 (see also Fig. 5). If the obstacle which occasions the lodgment of sand presents a surface which the wind cannot penetrate, such as a wall, sand is dropped abundantly on its windward side as well as on the leeward; but if it be penetrable, like an open fence, the lodgment takes place chiefly on its leeward side. In cultivated regions cases are known where, in a few weeks of dry weather, sand has been drifted into lanes in the lee of hedges to the depth of two or three feet, making them nearly impassable to vehicles.

Formation of dunes.—In contrast with dust deposited from the atmosphere, wind-blown sand is commonly aggregated into mounds and ridges in the process of lodgment. These mounds and ridges are dunes. Once a dune is started, it occasions the further lodgment of sand, and is a cause of its own growth. Dunes sometimes reach heights of 200 or 300 feet, but they are much more commonly no more than 10 or 20 feet in height. On plane surfaces, there is a limit in height above which they do not rise, though the limit is different under different conditions. The velocity of wind at the bottom of the air is not so great as that higher up, and as a dune is built up, a level is presently reached where the stronger upper winds sweep away as much sand as is brought to the top. The very even crests of many dune ridges are probably to be accounted for in this way. Wind-blown or eolian sand, not piled up in heaps or ridges, is somewhat widespread, but does not constitute dunes.

Shapes of dunes.[10]—Dunes may assume the form of ridges or of hillocks. The ridges may be transverse to the direction of the prevailing wind or parallel with it. Where dunes assume the form of hillocks rather than ridges, a group of them may be elongate in a direction parallel to the dominant wind, or at right angles thereto. The shape assumed by a dune or a group of dunes depends on the abundance of the sand, the strength and direction of the wind, and the shape of the obstacle which occasions the lodgment.

27

The incipient stages of dune formation are readily seen in many dry, sandy regions. The dune is likely to start in the lee of some obstacle, and to be elongate in the direction of the wind, especially if the wind be strong relative to the supply of sand. This shape is permanently28 preserved if the proper relations between the supply of sand and strength and direction of wind are preserved. In the dune region of the Indian desert[11] the prevailing winds are alternately the southwest and northeast monsoons, the former being the stronger. The supply of sand comes from the southwest. Near the southwest coast the dune ridges are parallel to the direction of the wind (Fig. 8); in the interior, where the winds are less strong, the dunes are transverse to it (Fig. 9); while between the districts where these two types prevail intermediate forms occur. The transverse dune ridges (Fig. 9) are said to be the result of the lateral growth and erosion of longitudinal dunes.[12] In regions of changeable winds the shape of the dunes is subject to great variation. Dunes are sometimes crescentic, the convexity facing the wind (Fig. 10).

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Along coasts, dune ridges are often transverse to the wind, and groups of dune hillocks are frequently elongate in the same direction. Here the source of supply of the sand is itself an elongate belt, often transverse to the dominant wind, and the resulting dunes often have great length transverse to the wind. Where the wind has strong mastery over the sand, the longitudinal tendency is seen, even along coasts.[13]

The shapes of dunes in section, like the shapes in ground-plan, depend on the relative strength and constancy of the winds and the supply of sand. With constant winds and abundant drifting sand, dunes are steep on the lee side (bc, Fig. 11), where the angle of slope is the angle of rest for the sand. It rarely exceeds 23° or 24°.[14] Under the same conditions the windward slope is relatively gentle (ab, Fig. 11). If the winds be variable so that the windward slope of one period becomes the leeward slope of another, and vice versa, this form is not preserved. Thus, by reversal of the wind, the section abc, Fig. 11, may be changed to adc. If the winds and the supply of sand be equal, on the average, from opposite directions, the slopes should, on the average, be equal, though perhaps unequal after any particular storm. The steep slopes of new-made dunes are lost after the sand has ceased to be blown. At some points where the winds erode (scour) more than they deposit, new profiles are developed (Figs. 12 and 13). The erosion profiles may be very irregular if the dunes are partially covered with vegetation. The effect of vegetation in restraining wind erosion is shown in Fig. 14, where plants have preserved a remnant of a dune.

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The topographic map.—Since dunes as well as other topographic features are conveniently represented on contour maps, and since such maps will be used frequently in the following pages, a general explanation of them is here introduced.

“The features represented on the topographic map are of three distinct kinds: (1) inequalities of surface, called relief, as plains, plateaus, valleys, hills, and mountains; (2) distribution of water, called drainage, as streams, lakes, and swamps; (3) the works of man, called culture, as roads, railroads, boundaries, villages, and cities.

Relief.—All elevations are measured from mean sea-level. The heights of many points are accurately determined, and those which are most important are given on the map in figures. It is desirable, however, to give the elevation31 of all parts of the area mapped, to delineate the horizontal outline, or contour, of all slopes, and to indicate their grade or degree of steepness. This is done by lines connecting points of equal elevation above mean sea-level, the lines being drawn at regular vertical intervals. These lines are called contours, and the uniform vertical space between each two contours is called the contour interval. On the maps of the United States Geological Survey the contours and elevations are printed in brown (see Plate II).

The manner in which contours express elevation, form, and grade is shown in the following sketch and corresponding contour map, Fig. 15.

The sketch represents a river valley between two hills. In the foreground is the sea, with a bay which is partly closed by a hooked sand bar. On each side of the valley is a terrace. From the terrace on the right a hill rises gradually, while from that on the left the ground ascends steeply in a precipice. Contrasted with this precipice is the gentle descent of the slope at the left. In the map each of these features is indicated, directly beneath its position in the sketch, by contours. The following explanation may make clearer the manner in which contours delineate elevation, form, and grade:

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1. A contour indicates approximately a certain height above sea-level. In this illustration the contour interval is 50 feet; therefore the contours are drawn at 50, 100, 150, 200 feet, and so on, above sea-level. Along the contour at 250 feet lie all points of the surface 250 feet above sea; and similarly with any other contour. In the space between any two contours are found all elevations above the lower and below the higher contour. Thus the contour at 150 feet falls just below the edge of the terrace, while that at 200 feet lies above the terrace; therefore all points on the terrace are shown to be more than 150 but less than 200 feet above sea. The summit of the higher hill is stated to be 670 feet above sea; accordingly the contour at 650 feet surrounds it. In this illustration nearly all the contours are numbered. Where this is not possible, certain contours—say every fifth one—are accentuated and numbered; the heights of others may then be ascertained by counting up or down from a numbered contour.

2. Contours define the forms of slopes. Since contours are continuous horizontal lines conforming to the surface of the ground, they wind smoothly about smooth surfaces, recede into all reëntrant angles of ravines, and project in passing about prominences. The relations of contour curves and angles to forms of the landscape can be traced in the map and sketch.

3. Contours show the approximate grade of any slope. The vertical space between two contours is the same, whether they lie along a cliff or on a gentle slope; but to rise a given height on a gentle slope one must go farther than on a steep slope, and therefore contours are far apart on gentle slopes and near together on steep ones.

For a flat or gently undulating country a small contour interval is used; for a steep or mountainous country a large interval is necessary. The smallest interval used on the atlas sheets of the Geological Survey is 5 feet. This is used for regions like the Mississippi delta and the Dismal Swamp. In mapping great mountain masses, like those in Colorado, the interval may be 250 feet. For intermediate relief contour intervals of 10, 20, 25, 50, and 100 feet are used.

Drainage.—Watercourses are indicated by blue lines. If the streams flow the year round the line is drawn unbroken, but if the channel is dry a part of the year the line is broken or dotted. Where a stream sinks and reappears at the surface, the supposed underground course is shown by a broken blue line. Lakes, marshes, and other bodies of water are also shown in blue, by appropriate conventional signs.

Culture.—The works of man, such as road, railroads, and towns, together with boundaries of townships, counties and states, and artificial details, are printed in black.”[15]

Topography of dune areas.—From what has been said, it is clear that the topography of dune regions may vary widely, but it is always distinctive. Where the dunes take the form of ridges (Fig. 1, Pl. II), the ridges are often of essentially uniform height and width for considerable distances. If there are parallel ridges, they are often separated 33by trough-like depressions. Where dunes assume the form of hillocks (Figs. 2 and 3, Pl. II), rather than ridges, the topography is even more distinctive. In some regions, depressions (basins) are associated with the dune hillocks. Occasionally they are hardly less notable than the dunes themselves. A somewhat similar association of hillocks and basins is locally developed by other means, but dunes are made up of sand and usually of sand only, while the composition of similarly shaped hillocks and depressions shaped by other agencies is notably different.

In Fig. 1, Plate II (Five Mile Beach, 8 miles northeast of Cape May, N. J.), the contour interval is 10 feet. There is here but one contour line (the 10-foot contour), though this appears in several places. Since this line connects places 10 feet above sea-level, all places between it and the sea (or marsh) are less than 10 feet above the water, while all places within the lines have an elevation of more than 10 feet. None of them reaches an elevation of 20 feet, since a 20-foot contour does not appear. It will be seen that some of the elevations in Fig. 1 are elongate, while others have the form of mounds.

Fig. 2 (Pl. II) shows dune topography along the Arkansas River in Kansas (near Larned); Fig. 4, dune topography in Nebraska (Lat. 42°, Long. 103°), not in immediate association with a valley or shore; and Fig. 3 shows irregular ridge-like dunes at the head of Lake Michigan. In Fig. 2 the contour interval is 20 feet. All the small hillocks southeast of the river are dunes. Some of them are represented by one contour and some by two. The altitude of the region is considerable, the heavy contour representing an elevation of 2100 feet; but the dunes themselves are rarely more than 20 feet above their surroundings. In Fig. 4, where the contour interval is also 20 feet, there are, besides the numerous hillocks, several depressions (basins). These are represented by hachures inside the contour lines. In some cases there are intermittent lakes (blue) in the depressions. The heavy contour at Spring Lakes in this figure is the contour of 4300 feet. There are two depression contours (4280 and 4260) below it. The bottom of the depression is therefore lower than 4260, but not so low as 4240. In Fig. 3 the contour interval is 10 feet, and the dune ridges north of Miller are more than 50 feet high. The dune ridges here have helped to determine the position of this branch of the Calumet River, and have blocked its former outlet. The present drainage is to the westward.

Migration of dunes.—By the continual transfer of sand from its windward to its leeward side, a dune may be moved from one place to another, though continuing to be made up, in large part, of the same sand. In their migration dunes sometimes invade fertile lands, causing so great loss that means are devised for stopping them. The simplest method (employed in France and Holland) is to help vegetation to get a foothold in the sand. The effect of the vegetation is to pin the34 sand down. As a dune ridge along a coast travels inland, another may be formed behind it. Successions of dune ridges are thus sometimes formed.

A remarkable instance of the migration of a sand dune is recorded on the Kurische Nehrung on the north coast of Germany. The Nehrung consists of a long narrow neck of land composed of sand, lying off the main coast. At the beginning of this century there was a notable dune ridge on one side. Since that time it has migrated a considerable distance, and in its migration it has been brought into the relationships illustrated in the accompanying diagrams (Fig. 16). In 180035 the dune ridge was on one side of a church, which was then in use. In 1839 the ridge had been so far shifted to the leeward as to completely bury the church, and in 1869, its migration had progressed so far as to again discover the building.[16]

When dunes migrate into a timbered region they bury and kill the trees (Fig. 17). In one instance on the coast of Prussia a tall pine forest, covering hundreds of acres, was destroyed during the brief period between 1804 and 1827.[17] At some points in New Jersey orchards have been so far buried within the lifetime of their owners that only the tops of the highest trees are exposed. Trees and other objects once buried may be again discovered by farther migration of the sand (Figs. 18 and 19).[18]

Eolian sand, not aggregated into distinct dunes, is often destructive. Even valleys and cities are sometimes buried by it. Drifting sands had so completely buried Nineveh two centuries after its destruction that its site was unknown.

Distribution of dunes.—Dunes are likely to be developed wherever dry sand is exposed to the wind. Their favorite situations are the 36dry and sandy shores of lakes and seas, sandy valleys, and arid sandy plains.

Along coasts, dunes are likely to be extensively developed only where the prevailing winds are on shore. Thus about Lake Michigan, where the prevailing winds are from the west, dunes are abundant and large on the east shore, and but few and small on the west. In shallow water, shore currents and storm waves often build up a reef of sand a little above the normal level of the water. When the waves subside, the sand dries and the wind heaps it up into dunes. This sequence of events is in progress at many points on the Atlantic Coast. Sandy Hook, New Jersey, and the “beaches” farther south started as barrier ridges. When the waves had built them above normal water-level, the wind re-worked the sand, piling it up into mounds and hillocks (Fig. 1, Pl. II). Such dune belts a little off shore are sometimes turned to good account. They are usually separated from the mainland by a shallow lagoon. Where land is valuable, the lagoon is sometimes filled in, making new land, thus anticipating the result which nature would achieve more slowly. This has been done at some points on the western coast of Europe.

Dunes are likely to occur along stream valleys (Fig. 2, Pl. II), if their bottoms or slopes are of sand, and not covered by vegetation. Dunes along valleys are usually on the side toward which the prevailing winds blow. Thus they are more common on the east side of the Mississippi than on the west. Dunes may be formed in the valley37 bottoms, but the sand is often blown up out of the valley and lodged on the bluffs above.

Apart from these special classes of situations, any sandy region the surface of which is dry is likely to have its surface material shifted by the wind and piled up into dune ridges or hillocks (Fig. 4, Pl. II). Dunes probably reach their greatest development in the Sahara, where some of them take the form of hillocks, and some the form of ridges. Travelers in that region report that dune ridges are sometimes encountered the faces of which are so high and steep as to be difficult of ascent, and that parties have been obliged to travel miles along their bases before finding a break where crossing was practicable.

Wind-ripples.—The surface of the dry sand over which the wind has blown for a few hours is likely to be marked with ripples (Fig. 20) similar to those made on a sandy bottom beneath shallow water, under the influence of waves. Like ripple marks made by the water, wind-ripples have one side (the lee) steeper than the other. While the ripples are, as a rule, but a fraction of an inch high, they throw much38 light on the origin of the great dune ridges. If the ripples be watched closely during the progress of a wind-storm, they are found gradually to shift their position. Sand is blown up the gentler windward slope to the crest of the ridge and falls down on the other side. The moment it falls below the crest of the ridge to leeward, it is protected against the wind, and is likely to lodge. Wear on the windward side is about equal to deposition on the leeward, and the result is the orderly progression of the ripples in the direction in which the wind is blowing, just as in the case of dune ridges.

Abrasion by the wind.—While the effect of the wind on sandy and dusty surfaces may be considerable, its effect on solid rock is relatively slight and accomplished, not by its own impact, but by that of the material it carries. The effect of blown sand on rock surfaces over and against which it is driven is perhaps best understood by recalling the effects of artificial sand-blasts, by means of which glass is etched. In a region where sand is blowing, exposed surfaces of rock suffer from a multitude of blows struck by the sand grains in transit. The result is that such rock surfaces are worn, and worn in a way peculiar to the agency accomplishing the work. If the rock be made up of laminæ which are of unequal hardness, the blown sand digs out the softer ones, leaving the harder projecting as ridges between them. Adjacent masses of harder and softer rock of whatever thickness are similarly affected. The sculpturing thus effected on projecting masses of rock is often picturesque and striking (Figs. 21 and 22), and is most common in arid regions. Details of wind-carving are shown in Fig. 23.

U. S. Geol. Surv.

Scale, 1+ mile per inch.

U. S. Geol. Surv.

Scale, 2+ mile per inch.

U. S. Geol. Surv.

Scale, 1+ mile per inch.

U. S. Geol. Surv.

Scale, 2+ mile per inch.

39

Sand drifted over loose stones lying on the surface often develops flat or flattish faces or facets on them. These facets are likely to be three in number, and the exposed portion of the stone is likely to develop a sort of pyramidal shape, the three flattish surfaces being mutually limited by tolerably well-defined lines (Fig. 24). Thus arise the three-faceted stones (Dreikanter of the Germans) commonly seen where sands have been long in movement.

Not only does the drifting sand wear the surface over which it passes and against which it strikes, but the grains themselves are worn in the process. They are liable to be broken as they strike rock surfaces, and they are likely to strike one another in the atmosphere. In both cases they are subject to wear, and so to reduction to a finer and finer state.

The erosion accomplished by the wind is therefore of various sorts. The impact of the wind itself picks up the fine materials which are already loosened, thus wearing down the surface from which they are removed; the materials picked up wear the rock surfaces against which they are blown, and the transported materials themselves suffer reduction in transit.

Effects of wind on plants.—Another effect of wind work is seen in the uprooting of trees (Fig. 25). The uprooting disturbs the surface41 in such a way as to make loose earth more readily accessible to wind and water. The uprooting of trees on steep slopes often causes the descent of considerable quantities of loose rock and soil. Again, organisms of various sorts (certain types of seeds, germs, etc.), as well as dust and sand, are extensively transported by the wind. While this is important biologically its geological effects are remote.

Indirect effects of the wind.—Other dynamic processes are called into being or stimulated by the atmosphere. Winds generate both waves and currents, and both are effective agents in geological work. The results of their activities are discussed elsewhere.

II. The Chemical Work of the Atmosphere.

The chemical work of the atmosphere (including solution and precipitation from solution) is principally accomplished in connection with water, a dry atmosphere having relatively little direct chemical effect on rock or soils.

Precipitation from solution.—The water in the soil is constantly evaporating. Such substances as it contains in solution are deposited where the water evaporates, and where evaporation is long continued42 without re-solution of the substances deposited, the surface becomes coated with an efflorescence of mineral matter. Conspicuous examples are found in the alkali plains of certain areas in the western part of the United States. Since the alkaline efflorescence is the result of evaporation it is connected with the atmosphere, but the material of the efflorescence was brought to its present position by water. The principle involved is illustrated by the white efflorescence which frequently appears on brick walls during the dry days which follow a drenching rain. The water penetrates the brick and mortar and dissolves something of their substance, and when it is evaporated from the surface the material in solution is left behind.

In arid regions the deposition of substances other than alkali is common. The percolating waters dissolve whatever is soluble, and when they evaporate their mineral content is left. The pebbles and stones of the arid plains have in many places become heavily coated with mineral matter deposited in this way, and not infrequently cemented into conglomerate. One of the commonest mineral substances found in such situations is lime carbonate. In some cases it was doubtless derived by solution from limestone beds beneath the surface, but this is not always the case. It often encrusts the bits of lava on lava plains where it can hardly have been derived from limestone. The faces of cliffs of granite or gneiss, hundreds and even thousands of feet above all other sorts of rock,[19] are sometimes spotted with patches of lime carbonate. In the first case the lime carbonate was derived by chemical change from the lava, and in the second, from the granite or gneiss (see Carbonation below), but its present position is the result of evaporation.

Oxidation.—In the presence of moisture the oxygen of the air enters into combination with various elements of the soil and rocks. This is oxidation. No other common mineral substance shows the results of oxidation so quickly and so distinctly as iron. The oxidized portion is loose and friable, and a mass of iron exposed to a moist atmosphere will ultimately crumble away. This change is comparable to other less obvious changes taking place in many minerals at and below the surface. Oxidation generally involves the disintegration of the rock concerned. Its effects in this direction will be referred to in other connections.

43

Carbonation.—The production of lime carbonate from rock containing calcium compounds, but not in the form of carbonates, is known as carbonation, and is one of the important chemical changes effected by the carbon dioxide of the atmosphere in coöperation with water. In the process of carbonation the original minerals of complex composition are decomposed and simpler ones usually formed. Volumetric changes are involved, which often lead to the disruption of the rock (see Ground water). Furthermore, carbonates are among the more soluble minerals, and their production therefore brings some of the rock materials into a soluble condition, and their extraction through solution tends still further to disintegrate the rock. The carbonation of crystalline rocks is therefore a disintegrating process, and will be considered further in its many concrete applications.

Other chemical changes.—A third chemical process which often accompanies oxidation and carbonation is hydration. This is effected by water rather than by air, and will be considered in that connection. In general it leads to the disintegration of the minerals and rocks affected. The chemical effects of nitric acid, etc., developed through the agency of atmospheric electricity, and the corresponding effects of the gases and vapors which issue from volcanoes, many of them chemically active, are to be mentioned in this connection.

Conditions favorable for chemical changes.—Conditions are not everywhere equally favorable for the chemical work of the atmosphere. In general, high temperatures facilitate chemical action, and, other things being equal, rocks are more readily decomposed by atmospheric action in warm than in cold regions. Chemical activity is probably greater where the climate is continuously warm than where there are great changes of temperature. Changes of temperature, on the other hand, tend to disrupt rock, and thus increase the amount of surface exposed to chemical change. Since nearly all the chemical changes worked by the atmosphere on the rocks are increased by the presence of moisture, the chemical activity of the atmosphere is greater in moist than in dry regions.

B. THE ATMOSPHERE AS A CONDITIONING AGENCY.

The most obvious mechanical work of the atmosphere is effected by the wind, but mechanical results of great importance, conditioned by the atmosphere, are also effected when the air is still.

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I. Temperature Effects.

When the sun shines on bare rock its surface is heated and expanded, and the expanded particles crowd one another with great force. Since rock is a poor conductor of heat its surface is heated and expanded notably more than parts beneath the surface. It follows that strains are set up between the expanded outer portion and the cooler and less expanded parts within. In the cooling of the same rock mass it is the outermost portion which cools first and fastest, and, contracting as it cools, strains are again set up between the outer part, which is cooled more, and the inner part, which is cooled less. The result may be illustrated by the effect of cold water on hot glass, or of hot water on cold glass. In either case the fracture is the result of the sudden and considerable differential expansion or contraction. Since the heating and cooling of rock are much slower than the heating and cooling of glass under the conditions mentioned, the rupturing effects are less conspicuous, but none the less real. The actual effects of temperature changes are illustrated by familiar phenomena. The surface portions of bowlders exposed to the sun are frequently seen to be shelling off (Fig. 26). The loosened concentric shells may be a fraction of an inch, or sometimes even several inches in thickness. This process of exfoliation affects not only bowlders, but bare rock surfaces wherever exposed to the sun (Figs. 27, 28). It is often conspicuous on the faces of cliffs.

45

Several conditions, some of which are connected with the atmosphere and some with the rock, determine the efficiency of this process. Since the breaking of the rock results from the expansion and contraction due to its changes of temperature, it follows that, other things being equal, the greater the change, the greater the breaking; but the suddenness of the temperature change is even more important than its amount. It follows that great daily, rather than great annual, changes of temperature[20] favor rock-breaking, though with changes of a given frequency their effectiveness is greater the greater their range. A partial exception to this generalization should be noted. If abundant moisture is present in the pores and cracks of the rock a change of temperature from 45° to 35° (Fahr.) might be far less effective in breaking the rock than a change from 35° to 25° in the same time, for in the latter case the sudden and very considerable expansion (about one-tenth) which water undergoes on freezing is brought into play. This may be called the wedge-work of ice. The daily range of temperature is influenced especially by latitude, 46altitude, and humidity. Other things being equal, the greatest daily ranges of temperature occur in high-temperate latitudes, though to this general statement there are local exceptions, depending on other conditions. High altitudes favor great daily ranges of temperature, so far as the rock surface is concerned (see Figs. 29, 30), for though the rock becomes heated during the sunny day, the thinness and dryness of the atmosphere allow its heat to radiate rapidly at night. Here, too, the daily range of temperature is likely to bring the wedge-work of ice into play. Since the south side of a mountain (in the northern hemisphere) is heated more than the north, it is subject to the greater daily range of temperature, and the rock on this side suffers the greater disruption. Similarly, rock surfaces on which the sun shines daily are subject to greater disruption than those much shielded by clouds. Isolated peaks, because of their greater exposure, are subject to rather greater daily ranges of temperature than plateaus of the same elevation.

The daily range of temperature is also influenced by humidity. Because of the effect of water vapor in the atmosphere on insolation47 and radiation, a rock surface becomes hotter in the day and cooler at night beneath a dry atmosphere than beneath a moist one. Aridity therefore favors the disruption of rock by changing temperatures.

Turning from the conditions of the atmosphere which affect the disruption of rock to the conditions of the rock which influence the same process, several points are to be noted. In the first place, the disrupting effects of changes of temperature are slight or nil where the solid rock is protected by soil, clay, sand, gravel, snow, or other incoherent material. If the constituent parts of the loose material are coarse, like bowlders, their surfaces are affected like those of larger bodies of rock. The color of rock, its texture and its composition, also influence its range of daily temperature by influencing absorption and conduction. Dark-colored rocks absorb more heat than light-colored ones, and compact rocks are better conductors than porous ones. Great absorption48 and slow conduction favor disruption. A given range of temperature is unequally effective on rocks of different mineral composition. In general crystalline rocks (igneous and metamorphic) are more subject to disruption by this means than sedimentary rocks, partly because they are more compact, but especially because they are made up of aggregates of crystals of different minerals which, under changes of temperature, expand and contract at different rates, while the common sedimentary rocks are made up largely of numerous particles of one mineral.

The freezing of water in the pores of rock is effective in disrupting them only when the pores are essentially full at the time of freezing. Otherwise there is room for the expansion attending the freezing. If the pores of the rock are large, the expansion on freezing may force out sufficient water to balance the increase of volume, even though the rock was completely saturated. If the pores be very small the water passes out less readily, and if the rock is saturated, freezing is more likely to be attended with disruption.[21]

49

In view of these considerations the breaking of rock by changes of temperature should be greatest on the bare slopes of isolated elevations of crystalline rock, where the temperature conditions of temperate latitudes prevail, and where the atmosphere is relatively free from moisture. All these conditions are not often found in one place, but the disrupting effects of changing temperatures are best seen where several of them are associated (Figs. 29, 30, and 31).

The importance of this method of rock-breaking has rarely been appreciated except by those who have worked in high and dry regions. Climbers of high mountains know that almost every high peak is covered with broken rock to such an extent as to make its ascent dangerous to the uninitiated. High serrate peaks, especially of crystalline rock, are, as a rule, literally crumbling to pieces (Fig. 31). The piles of talus which lie at the bases of steep mountain slopes are often hundreds of feet in height, and their materials are often in large part the result of the process here under discussion. In mountain regions where atmospheric conditions favor sudden changes of temperature, the sharp reports of the disruption of rock masses are often heard. Masses of rock, scores and even hundreds of pounds in weight, are frequently thus detached and started on their downward course.[22] Small pieces of rock are of course much more commonly broken off than large ones. The disruption of rock by changes of temperature is not usually the result of a single change of temperature, but rather of many successive expansions and contractions.

The sharp needle-like peaks which mark the summits of most high mountain ranges (Fig. 32) are largely developed by the process here outlined. The altitude at which the serrate topography appears varies with the latitude, being, as a rule, higher in low latitudes and lower in high. But even in the same latitude it varies notably with the isolation of the mountains and with the aridity of the climate. Thus within the United States the sharply serrate summits appear in some places, as in Washington and Oregon, at altitudes of 6000 to 10,000 feet, while in the isolated Wichita range of Oklahoma, much farther south, but in a much drier climate, the same sort of topography is developed at altitudes of 2500 to 3000 feet.

50

Even in low latitudes and moist climates the effects of temperature changes are often seen. Thin beds of limestone at the bottom of quarries are sometimes so expanded by the heat of the sun as to arch up and break.[23] In desert and arid regions,[24] whatever the altitude, the effects of temperature changes are often striking.

The disruption of rock by changes of temperature is one phase of weathering. It tends to the formation of a mantle of rock waste, which, were it not removed, would soon completely cover the solid rock beneath and protect it from further disruption by heating and cooling; but the loose material thus produced becomes an easy prey to running water, so that the work of the atmosphere prepares the way for that of other eroding agencies.

II. Evaporation and Precipitation.

Perhaps the most important work of the atmosphere as a dynamic agent lies in its function as the medium for the circulation and distribution51 of water. Atmospheric temperature is the primary factor governing evaporation, an important factor in the circulation of the vapor after it is formed, and controls its condensation and precipitation.

The average amount of annual precipitation on the land is variously estimated at from forty to sixty inches, the lesser figure being probably more nearly correct. Since much of this water falls at high altitudes, the work which it accomplishes in getting back to the sea is great. The water which falls on the land, if withdrawn wholly from the ocean, would exhaust that body of water in 10,000 to 15,000 years if none of it returned. The work of evaporation is of course not done by the atmosphere, though the atmosphere determines the effect of the solar energy which does the work.[25]

The precipitation is distributed with great inequality, and this inequality affects both the rain and the snow. Some regions have heavy precipitation and some light; some regions have much rain and little snow; others have much snow and little rain; others have rain and no snow, and still others have snow and little or no rain. The amount and distribution of rain and snow determine the size and distribution of streams and glaciers, and streams and glaciers are the most important agencies modifying the surface of the land.

It is impossible to separate sharply the geologic work of the water of the atmosphere from that of other waters; but so long as moisture is in the atmosphere (including the time of its precipitation) its effects are best considered in connection with the atmosphere.

The mechanical work of the rain.—In falling the rain washes the atmosphere, taking from it much of the dust, spores, etc., which the winds have lifted from the surface of the dry land. Not only this, but in passing through the atmosphere the water dissolves some of its gases, and perhaps particles of soluble solid matter. When therefore the falling water reaches the surface of the land it is no longer pure, and some of the gases it has taken up in its descent enable it to dissolve various mineral matters on which pure water has little effect.

As it falls on the surface of the land the rain produces various effects of a mechanical nature. In the first place, it leaves on the 52surface the solid matter taken from the air. The amount of material, thus added to any given region in any particular shower is trivial, but in the course of long periods of time the total amount of material washed out of the air must be very great.

Every rain-drop strikes a blow. If the drops fall on vegetation, they have little effect, but if they fall on sand or unprotected earthy matter they cause movements of the particles on one another, and this movement involves friction and wear. While the results thus effected are inconsiderable in any brief period of time, they are not so insignificant when the long periods of the earth’s history are considered.

Clayey soils contract and often crack on drying. Falling on such a soil when it is dry the rain causes it to expand, and the cracks are healed by lateral swelling. The same soils are baked under the influence of the sun, and when in this condition are softened and made more mobile by the falling of rain. Under the influence of the expansion and contraction occasioned by wetting and drying, the soils and earths on slopes creep slowly downward. When rain falls on dry sand or dust the cohesion is at once increased, and shifting by the wind is temporarily stopped.

After the water has fallen on the land its further work cannot be looked upon as a part of the work of the atmosphere; but any conception of the geological work of the atmosphere which did not recognize the fact that the waters of the land have come through the atmosphere would be inadequate. The work of the water after it has been precipitated from the atmosphere must be considered in another chapter.

III. Effects of Electricity.

Another dynamic effect conditioned by the atmosphere is that produced by lightning. In the aggregate this result is inconsequential; yet instances are known where large bodies of rock have been fractured by a stroke of lightning, and masses many tons in weight have sometimes been moved appreciable distances. Incipient fusion in very limited spots is also known to have been induced by lightning. Where it strikes sand it often fuses the sand for a short distance, and, on cooling, the partially fused material is consolidated, forming a little tube or irregular rod (a fulgurite) of partially glassy matter. Fulgurites are usually only a few inches in length, and more commonly 54than otherwise a fraction of an inch in diameter. Strictly speaking these results are the effect of the electricity of the atmosphere rather than of the atmosphere itself, but they are best mentioned in this connection.

Allusion has already been made to the chemical changes in the atmosphere occasioned by electric discharges.

SUMMARY.

Weathering.—The result of all atmospheric processes, whether physical or chemical, by which surface rock is disrupted, decomposed or in any way loosened, is weathering. This convenient term also includes similar results effected by ground water, plants, etc. The tendency of weathering is to produce a mantle of residuary earth over55 solid rock. Weathering by mechanical means tends to produce material which, though in a finer state of division, is still like the original rock in chemical composition. Weathering by chemical means tends to produce a mantle made up chiefly of the less soluble parts of the rock from which it was derived. All processes of weathering prepare material for transportation by wind and water.

Many considerations determine the thickness which the mantle of weathered rock (mantle rock) attains. Some of these considerations have to do with the atmosphere, and some with drainage. Since the latter are, on the whole, more important, this matter will be discussed in connection with the work of water (Chapters III and IV).

Familiar phenomena, both of land and sea, reveal the constant activity and importance of water as a geologic agent. Even when there is no precipitation the moisture in the air influences its activity in certain ways. Just as iron “rusts” more readily in moist air than in dry, so changes in other mineral substances are influenced by atmospheric humidity. Where precipitation takes place the results are more obvious. The passing shower works changes in the surface of the land, striking in proportion to the rate and amount of precipitation. The rains feed the streams, and every stream is modifying its bed, and with increasing rapidity as its current is swollen. Even the moisture which is precipitated as snow works its appropriate results. Before it melts it protects the surface against other agents of change; but if it accumulates in sufficient quantity in appropriate situations, it may give rise to avalanches and glaciers, which, like running water, degrade the surface over which they pass.

A part of the water which falls as rain, and a part of that which results from the melting of snow and ice, sinks into the soil and into the rock below, becoming ground water. It is this ground water which especially justifies the name hydrosphere, often applied to the waters of the earth, for it literally forms a spherical layer in the outer portion of the solid part of the earth. During the stay of the water beneath the surface it effects changes in the rocks through which it passes, dissolving mineral matter here and depositing it there, substituting one substance for another in this place, and effecting new chemical combinations in that. Slow as these processes are, they have worked wondrous changes in the course of the earth’s history.

When the waters are gathered together in ponds, lakes, and oceans, they are still active, and the results of their activity are seen along the shores, where winds and waves produce their chiefest effects. Even the ocean currents, far from land, and the processes of the deep sea, are not without their effect on the course of geological history.

The work of the surface waters, ground (underground) waters, standing waters, and ice will be considered in order.

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RAIN AND RIVER EROSION.

Rain and river erosion began when the first rains fell on land surfaces. Neither the location nor the nature of the first land surface is known. There is little reason to believe that the ocean was ever universal, but there is reason to believe that most land areas have at some time or other been covered by the sea. The prevalent conception that land areas which were once submerged came into existence by being elevated above sea-level, should be supplemented by the alternative conception that submerged areas may have become land by the depression of the ocean basins, thus drawing off the water from the areas where it was shallow. Thus in Fig. 36 the sinking of the sea-bottom from a to b would lower the surface of the water from cc′, to dd′, and draw off the water from the surfaces cd and c′d′.

Without attempting to picture the character of the original land our study of subaërial erosion may begin with an area which has just been changed from sea bottom to land. What is the nature of such a land surface? Of what material is it composed, and what is the character of its topography? Concerning its constitution something may be inferred from the nature of the deposits now found at the bottom of the sea. Near the shore and in shallow water they often consist of gravel and sand, though other materials are not wanting. Far from shore and in deep water they consist for the most part of fine sediments, some of which were washed or blown from the land, some of which came from the shells and other secretions of marine animals, some from volcanoes, and some from various other sources. The topography of the newly emerged land may have had some likeness to the topography of the sea bottom. The numerous soundings which have been made over large areas of the sea have shown that its bottom is, as a rule, free from the numerous small irregularities which affect58 the surface of the land. They seem to show that a large part of the ocean bottom is so nearly flat that, if the water were removed, the eye would hardly detect irregularities in the surface. This statement does not lose sight of the fact that the ocean bottom is, in certain places, markedly irregular. Volcanic peaks and striking irregularities of other sorts abound in some places. Nevertheless if the bottom of the sea could be seen as the land is, its most striking feature, taken as a whole, would be its apparent flatness.

With the topography of the sea bottom the topography of the land is, in its details, in sharp contrast. In order to get at the history of the latter, we may study the sequence of events which would follow the emergence of a portion of the former.

Subaërial Erosion without Valleys.

For the sake of emphasizing the fundamental principles involved in the work of running water, a hypothetical case will first be studied in some detail, even at the risk of elaborating processes already understood. The principles themselves will find application later in relations which are much less simple.

Let it be assumed that the area of newly emerged land is a circular dome-shaped island. The simplest possible condition is represented by assuming its slope to be the same in all directions from the center, and its materials to be absolutely homogeneous. Such an island would be subject to all the forces ordinarily operating on land surfaces. The chief agency tending to modify land surfaces is atmospheric precipitation. It will be assumed that the rain falls on the surface of the island with absolute equality at all points, and that all other forces which affect it operate equally everywhere.

The rain falling on a land area disappears in various ways; part of it evaporates, part of it sinks, and part of it runs off over the surface. If the island be composed of fine and unconsolidated materials, such as clay, the water which runs off over the surface will carry sediment down to the sea. If the island be composed of solid rock instead, exposure to the air will cause it to decay, and the products of decay, such as sand and mud, will suffer a like fate.

For the sake of a clear understanding of the processes involved, two cases may be postulated; one in which the waters of the sea remove the sediment washed down from the hypothetical island as fast as it reaches the shore, and one in which they allow it to accumulate59 without let or hindrance. In both cases the wear of the waves will be neglected.

1. In the first case the water flowing off over the surface (the run-off) will descend equally in all directions. It will constitute a continuous sheet of surface-water, and both its volume and its velocity will be the same at all points equally distant from the summit. Erosion accomplished by sheets of running water, as distinct from streams, is sheet (or sheet-flood) erosion.[26] Since the material of the surface is homogeneous, the wear effected by the water will be equal at all points where its velocity and volume are equal. For obvious reasons the depth of the run-off will increase from summit to base. The gradient (slope) also increases in the same direction, and the increase of volume and of gradient conspire to augment the velocity of the water, and therefore of the wear effected by it. If the thin sheet of water starting from the top of the island with relatively low velocity be able to wash off even a little fine material from the surface, the thicker sheet farther down the slope, moving with greater velocity, will be able to carry away more of the same sort of material, and the increase will be progressive from summit to base. It follows, therefore, that the surface will be worn equally at points equally distant from the summit, but unequally at points unequally distant from it. The first shower which falls on the island may be conceived to wash off from its surface a very thin sheet of material, but a sheet which increases in thickness from top to bottom. The run-off will not be stopped immediately on reaching the sea, but will displace the sea-water to some slight depth, and wear the surface some trivial distance below the normal level of the sea. The result of successive showers working in the same way through a long period of time will be to diminish the area of the island and to steepen its slopes. The results of a considerable period of erosion under these conditions are shown 60diagrammatically in Fig. 37, which illustrates both the diminution in area which the island has suffered, and the increase in the angle of its slopes. Immediately about it, at the stage represented by aa, Fig. 37, there is a narrow marginal platform, or submerged terrace, in place of the land area which has been worn away at or just below the level of the sea. Long successions of rains working in the same way will give the island steeper slopes, a smaller area, and a wider marginal terrace. Successive stages are shown by the lines bb and cc, Fig. 37.

If rain falls on such an island until it completes the work which it is possible for running water to do, the island will be reduced essentially to the level of the sea, and in its place there will be a plain, the area of which will be equal to that of the original island. Its central point will be at the level of the sea, and its borders a trifling distance below it (Fig. 38). The island is gone, and in its place there is a plain as low as running water can wear it. Other agencies might come in to defeat the result just outlined, but if the island did not rise or sink after its formation, rain falling upon it would, under the conditions specified, finally bring about the result which has been sketched. The plain (Fig. 38) which succeeds the island is a base-level of erosion, though this term is also used in other ways. Under these conditions the slope of the land would remain convex at all stages, but the convex erosion profile of the land would meet a nearly straight line just below sea-level. The relative lengths of these two elements of the profile, the curve above and the straight line below, vary as erosion progresses, the convex portion becoming shorter and the other longer, The two parts of the profile taken together are concave upward at the lower end all the time, and for a greater distance from its lower end in all the advanced stages of erosion (Fig. 37).

In the destruction of the land under these conditions neither valleys nor hills would be developed, nor would the topography of the land be61 fashioned to correspond with the surfaces with which we are familiar.

It is to be distinctly borne in mind that the foregoing is a hypothetical case; it is not probable that such an island ever existed, or ever will; but that does not diminish the value of the illustration, since the principles involved are operating on every land mass, though in less simple relations.

2. The second case differs from the first in that the sediment washed down from the land is deposited about its borders. This results in the building up of a marginal platform, as shown in Figs. 39–41. As erosion goes on more sediment is washed down and deposited, partly on the narrow marginal shelf which has already been developed, and partly on its outer slope, as shown in the figures. The marginal flat is thus extended beyond the original shores of the island on the one hand, and toward its center on the other. As it develops, its inner portion, and indeed all except its outer edge (ab, Figs. 40 and 41), will be gradually built up above the level of the water. This marginal lowland is developed at a level as low as running water, under the conditions then and there present, can reduce the land. Such a surface may be said to be at grade, since running water neither wears it down nor builds it up. Its angle of slope is a function of (1) the volume of the water running over it, and (2) of the load which the water carries.

Since the marginal plain of the above illustration extends beyond the original shore of the island, the area of land is increased, though both its average elevation and its mass (above water) are reduced. In case destructive processes did not operate on the marginal graded62 plain the spreading and lowering suggested by Figs. 39 and 40 would go on until the central mass of the island was brought down to a gradient in harmony with that of the gently sloping border, as shown in Fig. 41. When this had been accomplished there would be a relatively large land area with low slopes (Fig. 41) in place of the smaller area with steeper ones (compare Figs. 39 and 40). The basal part of the larger island from the center to the original margin would be made up of the original material in its original position (unshaded part of Fig. 41). Its surface would be covered, least deeply near its center and most deeply near the original margin, with débris gradually shifted from higher levels, as shown in Fig. 41.

Were such an island as that shown in Fig. 41 once formed, the rain falling on it, and flowing off over its surface, would carry off its surface soil and spread it about the shores. Though the surface of the marginal flat of Fig. 40 was as low as running water could bring it at the time it was developed, the conditions of erosion have changed by the time the land reaches the conditions shown in Fig. 41, and the same amount of rainfall may now be effective in erosion. In the first case (Fig. 40) the water descending from the higher part of the land brought down sediment and started across the flat with a load. Its energy was consumed in transporting what it had, not in getting new material. In the second case (Fig. 41) the water flowing over the gently sloping surface has no initial load, and its energy is therefore available for erosion. Under continued rainfall, the area of the land shown in Fig. 41 would be increased as before by successive marginal deposits (see Fig. 42), and at the same time its average height would be reduced. The lowering and enlarging of the island would continue until the whole surface was brought so nearly to the level of the sea that water would cease to run over it with sufficient velocity to carry away even the fine material of its surface. Such a surface, brought down as low as running water can degrade it, is also (see p. 57) a base-level. It will be seen from the foregoing illustrations that a graded surface may pass into a base-level, with no sharper line of demarkation63 than that which separates a mature man from an old one. In this case, as in the preceding, the island has been base-leveled, but still without the formation of valleys or hills.

Both the preceding hypothetical cases make it clear that, from the point of view of erosion, every drop of water which runs off over the surface of the land has for its mission the getting of the land into the sea. Under ordinary conditions surface drainage must fail to bring a land area altogether to sea-level, the absolute base-level of subaërial forces; but it is not simply the water which runs off over the surface which degrades the land. That which sinks beneath the surface contributes to the same end by slowly dissolving mineral matter below the surface, and finally carrying it to the sea. In this way the reduction of land areas to sea-level may be completed.

The rain-water which evaporates from the surface without sinking beneath it does not effect much wear; but the water thus evaporated is subject to reprecipitation, so that, in the long run, it may assist in the work which has been sketched. Thus it is not simply the waters which run off over the surface of the land, but all which fall upon it, which unite to compass its destruction.

The Development of Valleys.

By the growth of gullies.—Had the slopes of the hypothetical island not been absolutely uniform the processes of erosion would have been different. Let the departure from uniformity be supposed to consist of a single slight meridional depression near the base of the island (Fig. 43). As the rain falls it will no longer run off equally in all directions. A greater volume will flow through the depression than over other parts of the surface having the same altitude, and the greater volume of water along this line will give greater velocity, greater velocity will occasion greater erosion, and greater erosion will deepen the depression. The immediate result is a gully or wash (Fig. 44). So soon as the gully is started it tends still further to concentrate drainage in itself, and is thereby enlarged. The water which enters it64 from the sides widens it; that which enters at its head lengthens it by causing its upper end to recede; and all which flows through it, so long as its bottom is above base-level, deepens it. The enlarged gully will gather more water to itself, and, as before, increased volume means increased velocity, and increased velocity increased erosion. As the gully grows, therefore, its increased size becomes the occasion of still further enlargement.

Continued growth transforms the gully into a ravine, though between a gully and a ravine there is no distinct line of demarkation. But growth does not stop with ravine-hood. Water from every shower gathers in the ravine, and, flowing through it, increases its length, width, and depth, until it reaches such proportions that the term ravine is laid aside, as childhood names are, and the depression becomes a valley.

It was assumed in the preceding paragraphs that the single depression in the slope was meridional and low on the slope, but almost any sort of depression in almost any position would bring about a similar result, since it would lead to concentration of the run-off. Had the original surface been interrupted by ridges instead of depressions, the effect on valley development would have been much the same, for a ridge, like a depression, would, in almost any position, occasion the concentration of the run-off, and so the development of valleys. Under the conditions represented in Fig. 44 the lengthening of the drainage depression is effected chiefly at its upper end, the head of the valley working its way farther and farther back into the land. This method of elongation is known as head erosion. But the lengthening of the valley is not always wholly by head erosion. The gully normally begins where concentration of run-off begins, and if this were not at sea-level, the gully might be lengthening at both ends at the same time. This would have been the case, for example, had the original depression of Fig. 43 been half-way up the slope of the island.

If while the slopes of the island were absolutely uniform its surface material failed of homogeneity, the result would be much the same as if the slopes were unequal. If the material lying along a65 certain meridian of the island be slightly softer than that over the rest of the surface, the run-off, which would at the outset be equal on all sides, would effect more erosion along the line of the less resistant material than elsewhere. The result would be a depression along this line, and, once started, the depression would be a cause of its own growth. If the soft material were disposed in any way other than that indicated, the final result would be much the same, for it would quickly give origin to a depression which would lead to the concentration of the surface-waters, and this is the condition for the development of a gully, a ravine, and finally a valley.

In the presence of sufficient rainfall, either heterogeneity of slope or of material will therefore occasion the development of valleys. If the lack of uniformity appears at but a single point there will be but a single valley. If it appears at many points the number of valleys will be large. Since it is incredible that a land mass of perfectly homogeneous material and of absolutely uniform slopes ever existed, it is believed that every land mass, affected for any considerable length of time by rain, has had valleys developed in it. The degree of heterogeneity of material and slope is usually so great as to lead to the development of many valleys, even on areas which are not large; but for the sake of emphasizing the simpler elements of the complex processes of stream work, the hypothetical case of an island with but a single valley, and that without tributaries, may first be studied. Under these conditions two cases may be considered, the one where there is no deposition about the island, and the other where deposition takes place.

1. If all the material eroded from the surface of such an island, both in and out of the valley, were carried well beyond the borders of the land before being deposited, the edge of the island would recede66 from its original position toward the center, as illustrated by Figs. 37 and 45; but the recession would be most rapid where the valley joins the sea (Fig. 45). At this point therefore a reëntrant would be developed (a, Fig. 45), and the island would lose its circular outline. Continued erosion would cause the shore-line to retreat on all sides, but fastest at the lower end of the valley, and the final result would be a base-level differing from that developed under the conditions specified on p. 60, in that the last part to be brought low would not be the center of the original island.

Under the foregoing conditions the profile of that part of the valley which is above sea-level (cb) would be convex, following the analogy of sheet erosion on a hypothetical island of uniform slopes and homogeneous material with no marginal deposition. Its side slopes, likewise developed under the influence of running water augmented in volume from top to bottom, would also be convex.

2. If the sediment washed down from the land is deposited about its borders, both the outline of the island and the profile of the valley will be altered. Deposition at the debouchure of the valley follows the same principles as deposition elsewhere; but if all the sediment brought to the sea be deposited at the shore, the seaward extension of the land by deposition would be more rapid opposite the valley than elsewhere, and the island would lose its circular outline, and develop some such form as is shown in Fig. 46. In this case the profile of the upper end of the valley, and the upper parts of its side slopes, as well as the upper parts of the extra-valley slopes of the island, are convex (compare Figs. 39 and 40); but the convexity above is exchanged for concavity below, the change beginning at the point where downward erosion of the descending waters is checked. As a valley lengthens, the larger part of its67 profile becomes concave (compare the profiles of Figs. 39 to 41), but the extreme upper end still remains convex. Since the side slopes of a valley are much shorter than its lengthwise slope, a larger proportion remains convex. Under the conditions here discussed the change from convexity above to concavity below would begin at about the point where deposition begins.[27]

The deposition at the debouchure of the valley, and later above the debouchure, will follow the same course as about the island under the conditions already discussed (pp. 61, 62).

Limits of growth.—In all cases there are limits in depth, length, and width, beyond which a valley may not grow. In depth it may reach base-level. At the coast, base-level is sea-level,[28] but inland it rises by a gentle gradient. In length, the valley will grow as long as its head continues to work inland. In the case represented by Figs. 45 and 46 the head of the valley would not cease to advance when the center of the island was reached, though beyond that point head erosion would not be more rapid than lateral erosion on either side. If but a single valley affected a land area the limit toward which it would tend, 68and beyond which it could never pass, would be the length of the land area in the direction of the valley’s axis. In width, a valley is increased69 by the side cutting of the stream, by the wash of the rain which falls on its slopes, and by the action of gravity which tends to carry down to the bottom of the slope the material which is loosened above by any process whatsoever. If there be but one valley in a land area its limiting width is scarcely less than the width of the land itself.

Had there been several initial meridional depressions instead of one in the island, or had there been several meridional belts where the material of the surface was less resistant than elsewhere, several valleys would have been developed, converging toward the center (Fig. 47). If the conditions were such as to allow of the equal development of valleys on all sides of the island, each would be lengthened by head erosion until it reached the center of the island, where the permanent divide between their heads would be established. Each would be widened by all the processes which widen valleys, and their widening would narrow the intervening areas (Figs. 48 and 49). Under conditions of equal erosion the limits of width for each valley would be the centers of the ridges on either side, and here the divides between them would be permanently established. Though erosion would continue even after the crest of the ridge had been narrowed to a line, the permanence of the divide would follow from the fact that erosion would be equal on both sides of this line, and its effect would be to lower the divide, but not to shift it horizontally (Fig. 50). The limits in length and width are therefore not the same where there are several valleys as where there is but one. The limit in depth, however, remains the same, and the final result of erosion, proceeding along these lines, would be the base-leveling of the land, leaving a plain but slightly above sea-level. The plain would not be absolutely flat, though its relief would be very slight, and the higher parts would be along the lines of the divides between the streams (Fig. 51. Compare also Fig. 42). Many valleys would occasion more rapid degradation than few, and the period of base-leveling would be correspondingly shortened.

Had the initial depressions which gave origin to the valleys had positions other than meridional, the valleys would have had other and less regularly radial courses, but the final result of their development would have been the same.

70

It is not to be inferred that the method of valley development which has been sketched is the only one. The processes of valley development are complex, and the history of some valleys has run a different course; yet the processes outlined above are in operation in all cases, and they were probably the most important ones in the development of the first drainage system on any land surface. As will be seen in the sequel the history of valleys is subject to serious accidents, and they are often of such a nature as to mask the simplicity of the more normal processes.

The permanent stream.—From the foregoing discussion, it is seen that a valley may be developed by the run-off of successive showers. If supplied only from this source surface streams would cease to flow soon after the rain ceased to fall, and a valley might attain considerable size without possessing a permanent stream. How does the valley developed by the run-off of successive showers come to have a permanent stream? The answer to this question involves a brief consideration of that part of the rainfall which sinks beneath the surface.

71

If wells be sunk in a flat region of uniform structure and composition the water in them is generally found to stand at a nearly common level. The meaning of this fact is not far to seek. If a hole 60 feet deep fills with water up to a point 20 feet from the surface, it is because the material in which the well is sunk is full of water up to that level. When the well is dug the water leaks into it, filling the hole up to the level to which the rock (or subsoil) is itself full. This level, below which the rock and subsoil (down to unknown depths) are full of water, is known as the ground-water level, ground-water surface, or water-table.

The ground-water level fluctuates. In a wet season it rises, because more water has fallen and sunk beneath the soil; but several processes at once conspire to bring it down again. Where there is growing vegetation its roots draw up water from beneath, and evaporation also goes on independently of vegetation. The water is drawn out through wells and runs out through openings. It may also flow underground from one region to another where the ground-water surface is lower. All these processes depress the ground-water surface.

A well sunk to such a level as to be supplied with abundant water in a wet season may go dry during a period of drought because the ground-water level is depressed below its bottom. Thus either well shown in Fig. 52 will have water during a wet season when the water-level is at a; but well No. 1 will go dry when the water surface is depressed to b.

The principles applicable to wells are applicable to valleys. When a valley has been deepened until its bottom reaches below the ground-water level, water seeps or flows into it from the sides. The valley is then no longer dependent on the run-off of showers for a stream. It will be readily seen that at some stage in its development, the bottom of a valley may be below the ground-water level of a wet season without being below that of a dry one. Thus the valley represented in cross-section by the line 2–2, in Fig. 53, will have a stream when the ground-water level is at aa, but none when this level is depressed to bb. If the rainfall of the year were concentrated in a single wet season, the intermittent stream would flow not only during that season, but for so72 long a time afterward as the ground-water level remained well above the valley bottom. In regions subject to frequent and short periods of heavy precipitation, alternating with droughts, the periods of intermittent flow may be many and short. Since the precipitation of many regions varies greatly from year to year, it follows that a stream may flow continuously one year and be intermittent the next. Many valleys in various parts of the earth are now in the stage of development where their streams are intermittent.

As a valley containing an intermittent stream becomes deeper, the periods when it is dry become shorter, and when it has been sunk below the lowest ground-water level, it will have a permanent stream (3, Fig. 53). Since a valley normally develops headward, its lower and older portion is likely to acquire a permanent stream, while its upper and younger part has only an intermittent one (Fig. 47 and Fig. 1, Pl. III, near Anthony, Kan. The intermittent part of the stream is indicated by the dotted blue line). For the same reason the head of a stream is likely to be farther up the valley in wet weather than in dry. So soon as a valley gets a permanent stream, the process of enlargement goes on without the interruption to which it was subject when the supply of water was intermittent.

In general a permanent stream at one point in a valley means a continuous stream from that point to the sea or lake which the valley joins; but to this rule there are many exceptions. They are likely to arise where a stream heads in a region of abundant precipitation, and flows thence through an arid tract where the ground-water level is low, and evaporation great. In such cases, evaporation and absorption may dissipate the water gathered above, and the stream disappears (Fig. 2, Pl. III, near Paradise, Nev.).

U. S. Geol. Surv.

Scale, 2+ mile per inch.

U. S. Geol. Surv.

Scale, 4+ mile per inch.

PLATE IV.

U. S. Geol. Surv.

Scale, 1+ mile per inch.

U. S. Geol. Surv.

Scale, 2+ mile per inch.

73

Other modes of valley development.—If as a new area of land emerges from the sea its surface has a depression without an outlet, and such an assumption is by no means improbable, the depression would be filled with sea-water. The inflowing water from the surrounding land might fill the basin to overflowing, and the outflow, finding exit at the lowest point in the rim of the basin, would flow thence toward the sea. Such a stream would develop a valley, the history of which would be somewhat different from that which has been sketched. Instead of developing headward from the sea, the valley would be in process of excavation all the way from the initial basin to the sea at the same time (Fig. 54). The upper end of the valley might ultimately be cut to the level of the bottom of the basin, when the lake would disappear. The head of the74 valley might then work back across the former site of the lake into the territory beyond. Valleys might have developed above the lake before it was drained, and after this event, such valleys would make connections with the valley below (Fig. 55). A valley developed in this manner is not simply a gully grown big by head erosion, and the valley would not precede the stream.

If a surface of land were notably irregular before valleys were developed in it, there might be many lakes, and the flow from a higher lake might pass to a lower. If the lakes were ultimately drained, the several sections of the valley would be joined to one another without intervening basins. In certain regions, especially those which have been affected by continental ice-sheets, this has been a common method of valley development in post-glacial time. In this case also the stream precedes75 its valley, and not the valley its stream. Many post-glacial valleys, on the other hand, antedated their permanent streams, as in the cases first described.

If the gradient of a slope on which valleys are to develop is notably unequal, though without basins, the development of valleys may follow somewhat different lines. If on emergence the seaward part of a new land area assumes the form of a plain, bordered landward by a steeper slope (Fig. 56), the most notable early growth of the valleys would be on the latter. The run-off would develop gullies on the steep slope, but on reaching the plain below the velocity of the water would be checked, and it would drop much of the detritus washed down from above. This deposition would build up (aggrade) the surface, and much or even all the water might sink into and seep through the débris thus deposited, and disappear altogether from the surface, as at b, Fig. 56. This would be most likely to occur where the débris is abundant and coarse, and the precipitation slight. If the water disappears at the base of the mountain (see Fig. 2, Pl. III), the early growth of the valley may be confined to76 the steep slope remote from the sea (ab, Fig. 56); but on the slope where the valley is growing there will be headward lengthening, as in the general case already considered. If the surface drainage does not disappear at the base of the steep slope, the run-off will find its way over the plain along the lowest accessible route to the sea (de, Fig. 56). In this case the valley may be growing throughout its length at the same time.

The conditions represented by ab, Fig. 56, may be no more than temporary. Sooner or later a valley developing headward across the plain (hi, Fig. 56) may provide a channel for the water descending from the higher land beyond. In this case the valley develops in sections, the one on the slope above, the other on the plain below, and their union (compare fghi, Fig. 56, with Fig. 57) results from their growth.

The principles here sketched have been in operation wherever land areas were so elevated as to give rise to unequal slopes, and this has perhaps been the rule rather than the exception. The results effected by the operation of these principles would of course be dependent on the varieties of slope, on the abruptness with which a slope of one gradient gave place to another, on the texture of the rock, the amount and distribution of precipitation, etc., etc.

In the preceding paragraphs the lengthening of a valley at its upper end by head erosion has been repeatedly referred to. If all valleys began their development at the sea and lengthened headward, it might seem that their seaward ends should be their oldest parts; but since the development of valleys is begun somewhat promptly after the land appears above the sea, and since the emergence is generally gradual, that part of a valley which is at the seashore at one time may be far inland a little later, because the land has been extended seaward. On an emerging land area77 therefore the normal growth of a valley involves its lengthening at its lower end as well as at its upper. The lengthening of a valley, or at least the lengthening of a stream, also takes place at its lower end if the land in which it lies is being extended seaward by deposition.

Structural valleys.—In mountain regions valleys are sometimes formed by the uplift of parallel mountain folds, leaving a depression between (Fig. 58). Drainage will appropriate such a valley so that it becomes in some sense a river valley. But it is not a river valley in the sense in which the term has been used in the preceding pages. It is rather a structural valley. In its bottom a river valley may be developed (a, Fig. 58).

The foregoing illustrations by no means exhaust the list of conditions under which valleys develop, but they suffice for the present.

The courses of valleys.—River valleys are rarely straight. To understand why they are crooked it is only necessary to understand the methods by which they grow. In so far as a river valley is a gully grown big, that is, in so far as its length is the result of head erosion, its course was determined by the course of the antecedent gully. If in the case shown in Fig. 59 the slope of the surface above the head of the gully is uniform, its material homogeneous, and the rainfall everywhere equal, more water will come into the gully from the direction a than from any other. In this case there would be more wear in the direct line of its extension than elsewhere, and the head would advance in a straight line. But if there be inequalities of slope about the head of a gully at any stage of its development more water may come in78 from some direction other than that in the direct line of its extension. In Fig. 59, for example, more water may enter from the direction of b than from that of a. Since most wear is likely to be affected along the line of greatest inflow, the head of the gully will be turned in that direction (Fig. 60). Started in this course it will continue in the new direction so long as erosion in this line is greater than that elsewhere; but whenever the configuration of the surface causes more water to enter the head of the gully from some direction other than that in which it is headed, the line of axial growth is again changed, as toward c, Fig. 60. Since new land surfaces are probably more or less undulatory, crookedness should be the rule among valleys developed from gullies by head erosion. Streams and valleys the courses of which are determined by the original slope of the land are said to be consequent.

Inequalities of material, leading to unequal rates of erosion, effect the same result, in the absence of inequalities of slope. If at any stage of a valley’s development erosion were equal in two directions at its head, and at the same time greater than at points between, two gullies would result (Fig. 61) diverging from the point in question.

In the case of a valley developed by overflow from a lake its course is determined by the lowest line of flow to which the water has access. If this line be straight the valley will be straight; if it is crooked, as it generally is, the valley is crooked also.

The development of tributaries.—Thus far valleys leading immediately to the sea have been considered, and no account taken of tributaries. As a matter of fact most considerable valleys have numerous tributaries. It is now in order to inquire into their mode of development.

So soon as a gully is started, the water flowing into it from either side wears back the slopes. The least inequality of slope, or the least variation in the character of the material, is sufficient to make the lateral erosion unequal at different points, and unequal erosion in the slopes results in the development of tributary gullies. The oldest tributaries may be nearly as old as the main which they join, and from which they developed, for the possibilities of unequal side erosion exist as soon as a79 gully is opened. While the main gully is developing into a ravine, and the ravine into a valley, the tributary gullies are likewise developing into maturer stages. Tributary to a young valley, therefore, there may be gullies near its head, ravines in its middle course, and small valleys along its oldest portion. It is not to be understood, however, that the oldest tributaries are necessarily the largest, for because of more favorable conditions for growth the younger tributaries often outstrip the older.

The position of tributaries with reference to their mains is worthy of note. The water flowing down a slope follows the line of steepest descent. A gully is usually wider at its lower end, and narrower at its upper. Wherever this is true the line of steepest descent down its side is not a line perpendicular to its axis, but a line slightly oblique to it (ef, Fig. 62), and oblique in such a direction that it meets the axis with an obtuse angle below and an acute angle above. It is in the direction corresponding to this line that tributary gullies tend to develop. Thus at the inception of its history a tributary gully is likely to join its main with an angle slightly acute on the up-stream side. If the tributary did not begin until after its main was farther advanced this tendency would be less and less pronounced. Inequalities of material or slope would often counteract this tendency, which, at best, would cause the courses of tributaries to depart but little from perpendicularity to their mains.

After the head of a tributary has worked back from the immediate slope of its main every condition which determines the course of a gully is likely to affect it, and it is by no means certain that it will continue to lengthen in the direction in which it started. Since the general slope of the surface into which the tributary works is likely to be seaward, more water is likely to enter from the landward than from the seaward side of its head, so that, except where there are notable irregularities of slope, its tendency will be to turn more and more toward the direction of its main (efg, Fig. 62).

In depth the tributary is always limited by its main. The principles which determine the length and width of a main valley determine also the length and width of a tributary (see p. 67 et seq.).

80

A CYCLE OF EROSION. ITS STAGES.

From what has preceded it is clear that the topography of a region undergoing erosion will change greatly from time to time. The first effect of erosion is to roughen the surface by cutting out valleys, leaving ridges and hills. The final effect is to make it smooth again by cutting the ridges and hills down to the level of the valleys.

The base-level of erosion has already been defined; but the mode of its development may now be illustrated in the light of the preceding discussion. Suppose a land surface affected by a series of parallel young valleys without tributaries (Fig. 63). Between them there is a series of upland plateaus. The profile of the surface between two adjacent valleys is represented in section by the uppermost line in Fig. 64. As the valleys are widened from 1–1 and 1,′–1′, to 2–2 and 2′–2′, the intervening plateau is correspondingly narrowed. When the valleys have attained the form represented by 3–3 and 3′–3′, the intervening upland has been narrowed to a ridge, a, and the valley flats have become wide. With continued erosion the ridge will be lowered (to b and below), and81 in time the surface will approach a plain. In this condition it is known as a peneplain (an “almost-plain”). Finally, when running water has done its utmost, the ridges will be essentially obliterated and a base-leveled plain (e, e′, e″) results. The figure expresses the fact that the base-level develops laterally from the axis of the valley. It also develops headward from the seaward end of the valley. Similarly, taking into account all the valleys which affect it, the seaward margin of a base-leveled plain is developed first, and thence it extends itself inland.

Tributaries are tolerably sure to develop along each main valley. The heads of the tributaries work back across the uplands between the main valleys, dissecting them into secondary ridges (Fig. 65). Tributaries will develop on the tributaries, and these tertiary valleys dissect the secondary ridges into those of a lower order. This process of tributary development goes on until drainage lines of the fourth, fifth, sixth, and higher orders are formed (Fig. 66). Since the process of valley development under such circumstances is also the process of ridge dissection, a stage is presently reached where the ridges are cut into such short sections that they cease to be ridges, and become hills instead. Even then the processes of erosion do not stop, for the rain-water falling on the hills washes the loose material from their surfaces, and starts it on its seaward journey. Thus the “everlasting hills” themselves are lowered, and, given time enough, will be carried to the sea. Under these conditions, as under those already discussed,82 the final result of stream erosion is the reduction of the land to base-level. The base-leveled surface, as before, would not be absolutely flat. The area reduced by each stream will have a slight gradient down-stream, and from each lateral divide toward the axis of the valley. The crests of the scarcely perceptible elevations which remain will be in the position of the former divides, and these will be highest where most distant from the sea by the course which this part of the drainage took. Even the insensible divides between streams flowing in a common direction may disappear, for when valleys have reached their limits in depth, their streams do not cease to cut laterally. Meandering in their flat-bottomed valleys, they often reach and undercut the divides (Pl. VII), whether they be high or low. By lateral planation, therefore, the divides between streams may be entirely eaten away.

It has now been seen that by whatever method erosion by running water proceeds, whether there be many valleys, or few or none, the final result of subaërial erosion must be the production of a base-level. It has also been seen that the base-level is first developed at the lower ends of the main streams, and that it extends itself systematically up the main valleys and up all tributaries. The time involved in the reduction of a land area to base-level is a cycle of erosion.

It will have been evident from the preceding pages that the terms “grade,” “graded plain,” and “base-level” and “base-leveled plain,” are somewhat variously, and therefore somewhat confusingly, used. “Grade is a condition of essential balance between corrasion and deposition.”83[29] A graded valley is one in which deposition and corrasion are, in the vertical sense, balanced. Its angle of slope is most variable, and is dependent on the capacity of the stream for work, and on the work it has to do. A weak river must have a higher gradient than a strong one; a stream with much sediment must have a higher gradient than one with little, and a stream with a load of coarse material must have a higher gradient than one with a load of fine. Thus the graded valley of the lower Mississippi has an inappreciable angle of slope, but the graded valleys of many of its tributaries have slopes of hundreds of feet per mile. Since both the size of the stream and the amount and coarseness of its load at a given place vary from time to time, it is clear that the inclination of a graded valley must vary also, and further, that it must be in process of continual readjustment. With the changing conditions of advancing years the slope of a graded valley normally decreases. The same principles apply to graded surfaces outside of valleys.

In the continual readjustment of grades incident to a river’s normal history the land is brought nearer and nearer to sea-level without ceasing to be at grade. When the inclination of a graded surface becomes so low that it is sensibly flat, the surface may be said to be at base-level, although this does not mean that the surface can never be degraded further. If the term be used in this way, it is clear that there is no sharp line of distinction between a graded surface and a base-leveled surface, and as the terms are now commonly applied no such distinction exists.

If the term base-level were made synonymous with sea-level, as has been proposed,[30] the term might as well be discarded, for sea-level could always be used in its stead. Furthermore, streams often erode below sea-level. The bottom of the channel of the Mississippi is below sea-level for some 400 miles above its debouchure, and locally (Fort Jackson) it is nearly 250 feet below. This deep channel is the result of the erosive activity of the stream, not of subsidence. Again, the sea-level is itself inconstant. The extent of its changes cannot now be measured, but they have probably been more considerable in the course of geological history than has been commonly recognized. It is true that they take place slowly, as far as known, but it is also true that the duration 84of an erosion cycle is sufficiently long for even very slow changes to reach great magnitude. The sea-level, therefore, can hardly be accepted as the absolute base-level, unless (1) the absolute base-level is a variable, and unless (2) the absolute base-level be a surface below which rivers may cut to the extent of at least 250 feet.

The ocean may be looked upon as a barrier which in a general way limits the down-cutting of running water; for only very large streams cut much below its level. Other barriers, such as lakes, and the outcrops of hard rock in a stream’s bed, have a comparable, though more temporary, effect on the development of valley plains above. Plains thus developed have been called temporary base-levels. They differ from other graded plains in being controlled primarily by a barrier below, rather than by conditions which exist above.

Since river valleys have a beginning and pass through various stages of development before the country they drain is base-leveled, it is important to recognize their various stages of advancement. Nor is this difficult. An old valley and a young one have different characteristics, and the one would no more be mistaken for the other by those who have learned to interpret them, than the face of an aged man would be mistaken for that of a child.

The cycle begins with the beginning of valley development, and at that stage drainage is in its infancy. The type of the infant valley is the gully or ravine (Figs. 67 and 68). It has steep85 slopes and a narrow bottom. Fig. 1 of Plate IV represents similar, or rather older, ravines in contour (shore of Lake Michigan, just north of Chicago). With age, the valley widens, lengthens, and deepens, and passes from infancy to youth. In this stage also the valleys are relatively narrow, and the divides between them broad. They may be86 deep or shallow, according to the height of the land in which they are cut, and the fall of the water flowing through them; but in any case the streams flowing through them have done but a small part of the work they are to do before the country they drain is base-leveled. Figs. 69 and 70, respectively, represent youthful valleys in regions of moderate and great relief. Fig. 2, Plate IV, shows a youthful valley in a region of slight relief (near Casselton, N. D., Lat. 46° 40′, Long. 97° 25′). The uppermost line in Fig. 64 likewise represents topographic youth, as shown in cross-profile.

Not only are narrow valleys said to be young, but the territory affected by them is said to be in its topographic youth, since but a small part of the time necessary to reduce it to base-level has elapsed. An area is in its topographic youth when considerable portions of it are still unaffected by valleys. Thus the areas (as a whole), as well as the valleys, represented on Plate IV, are in their topographic youth. It is often convenient to recognize various sub-stages, such as early, middle, and late, within the youthful stage of valleys or topographies. The different parts of the areas shown on Plate IV, for example, represent different stages of youth.

Youthful streams, as well as youthful topographies, have their distinctive characteristics. They are usually swift; their cutting is mainly at the bottom rather than at the sides, and their courses are often marked by rapids and falls.

As valleys approach base-level they develop flats. As the valleys and their flats widen, and as their tributaries increase in numbers and size, a stage of erosion is presently reached where but little of the original upland surface remains. The country is largely reduced to slopes. In this condition the drainage and the topography which it has determined are said to be mature. Mature topography is shown in contours in the figures of Plate V, and in the northern part of Plate VI, where slopes, rather than upland or valley flats, predominate. Fig. 1 of Plate V represents an area in southeastern Kentucky (Lat. 37° 12′, Long. 83° 10′); Fig. 2, an area in western Virginia. Plate VI represents an area in southern California, somewhat west of San Bernardino. The three areas are alike in representing mature drainage, though not of equal stages of advancement. The striking differences of topography of the three areas are the result of differences in rock structure and altitude, and will be considered later. Mature topography is also88 shown in Fig. 71, where the relief is moderate, and in Figs. 72 and 73, where it is great. Figs. 72 and 73 illustrate clearly the universal tendency of rivers in regions of notable relief to develop new flats well below the old surface of the region. At the same time that these low-lying flats are developing, tributary drainage is dissecting and roughening the upper surfaces. This process is well shown in Fig. 73. In both Figs. 72 and 73 the summits of the mountains on either side of the valleys appear to have had about the same elevation. The new flat is therefore developed at the expense of the old flat. As will be seen in the sequel, the first flat which a stream develops along its course is usually somewhat above base-level. It is a graded flat.

89

PLATE V.

U. S. Geol. Surv.

Scale, 2+ mile per inch.

U. S. Geol. Surv.

Scale, 2+ mile per inch.

PLATE VI.

U. S. Geol. Surv.

Scale, 1+ mile per inch.

The same processes which have made young valleys mature will in time work further changes. When the gradients of the valleys have become low and their bottoms wide, and when the intervening ridges and hills have become narrow and small, the drainage and the drainage topography have reached old age, and the streams are in a condition of senility. This is illustrated by Fig. 1, Plate VII (central Kansas), and in section by the third and lower lines in Fig. 64. Topographic old age sometimes has a different expression; this is shown in Fig. 74, where most of the surface has been brought low. The elevations which rise above the general plain are small in area, but have abrupt slopes. This phase of old-age topography is usually the result of the unequal resistance of the rock degraded. The effects of unequal rock-resistance will be considered later.

The marks of old streams are as characteristic as those of young ones. They have low gradients and are sluggish. Instead of lowering their channels steadily they cut them down in flood, and fill them up when their currents are not swollen. They meander widely in90 their flat-bottomed valleys (Fig. 1, Pl. VII, Central Kansas) and their erosion, except in time of flood, is largely lateral.

If the processes of degradation were to continue until the land surface was brought to sea-level, and this might be done by solution though not by mechanical erosion of running water, the rivers would no longer flow, and the drainage system would have reached the end of its history—death.

Not only do valleys normally pass from birth to youth, from youth to maturity, and from maturity to old age, but a single river system may show these various stages of development in its various parts. Thus in the area shown in Fig. 2, Plate VII (north central Kansas), there is a tract (extreme southwest) where the erosion history is scarcely begun. The zone of land a little farther northeast, and just reached by the heads of the valleys (same figure), is in its youth. The well-drained and uneven tract southwest of the flat of the Solomon River is in maturity, while the flat of the main valley has the general characteristics of old age.

The age of valleys in terms of erosion is also expressed more or less perfectly by their cross-sections. The line 1–1 (and 1′–1′) of Fig. 64 represents in cross-section a narrow V-shaped valley. Such a section is always indicative of youth. The stream which developed it cut chiefly at its bottom, not at its sides. It was therefore rapid, and rapid streams are young. The line 2–2, (2′–2′) (Fig. 64) shows the same valley at a later and maturer stage when downward cutting has nearly ceased. The widening of the valley by slope wash has91 become relatively more important than before, and the stream has so far lost velocity as the result of diminished gradient as to be unable to carry away all the detritus washed down from the sides. As a result of deposition at the bases of the side slopes, a concave curve has been developed. Up the valley from the point where such a section as is represented by 2–2 occurs, the valley may still have a section similar to that represented by 1–1.

PLATE VII.

U. S. Geol. Surv.

Scale, 2+ mile per inch.

U. S. Geol. Surv.

Scale, 2+ mile per inch.

PLATE VIII.

U. S. Geol. Surv.

Scale, 1+ mile per inch.

Still later stages of development are represented by the cross-sections 3–3 and 4–4. Not only has the valley become larger, but the stream has deposited detritus (not shown in the figure) in the bottom of its valley, developing an alluvial flat. On this flat the stream meanders, and the valley may be widened by the undercutting of the bluffs wherever the stream in its wanderings reaches them (Pl. VIII, near St. Louis). A valley might possess the characteristics shown by the cross-sections 3–3, 2–2, and 1–1, Fig. 64, in its lower, middle, and upper courses, respectively.

The preceding discussion, and the illustrations which accompany it, give some idea of the topography which characterizes an area in various stages of its erosion history. Whether the valleys are deep or shallow, and the intervening ridges high or low, depends on the original height of the land and its distance from the sea. The higher the land, and the nearer it is to the sea, the greater the relief developed by erosion. A plateau near the sea may become mountainous in the mature stage of its erosion history, while a plain in the same situation would only become hilly. A plateau in the heart of a continent would have less relief in its maturity than one of equal elevation near the sea, since the grade-plain in the former position is higher than in the latter. Plates IV and IX show youthful topography where the relief is relatively slight, and Plate X shows youthful topography where the relief is great. Similarly, Plates V and VI show mature topography where the relief is great, and Fig. 1, Plate III, shows mature topography where the relief is relatively slight.

Topographic youth, topographic maturity, and topographic old age are also indicated in other ways, and especially by the presence of features which rivers tend to destroy. If, for example, the surface of the land, well above the valley bottoms, is marked by numerous ponds and marshes, it is clear that drainage has not yet progressed beyond its early stages, for, unless the lakes be very deep, valleys working back into the land will find and drain them before topographic92 maturity has been reached. Their presence is evidence that the region where they occur has not yet been thoroughly dissected by erosion lines, and therefore has not reached maturity. Still other marks of topographic youth, such as rapids, falls, etc., as well as marks of topographic maturity and old age, will be mentioned in the following pages.

GENERAL CHARACTERISTICS OF TOPOGRAPHIES DEVELOPED BY RIVER EROSION.

With the characteristics of river valleys and the methods by which they grow clearly in mind it is easy to say whether rivers have been the chief agents in the development of a given topography. River valleys are distinguished from other depressions on land surfaces by their linear form and, leaving out of consideration the relatively insignificant inequalities in a stream’s channel, by the fact that any point in the bottom of a river valley is lower than any other point farther up the stream in the same valley, and higher than any point farther down the stream. The second point might be otherwise stated by saying that every valley excavated by erosion leads to a lower valley, or to the sea, or an inland basin. Streams which dry up, or otherwise disappear as they flow, constitute partial exceptions. If, therefore, the depressions on a land surface are linear, lead to other and deeper valleys, and finally to an inland basin, or the sea, and if the elevations between these valleys are such as might have been left by the excavation of the valleys, it is generally clear that rain and rivers have been the chief factors in the development of the topography. If, on the other hand, a surface is characterized by topographic features which streams cannot develop, such as enclosed depressions, or hills and ridges whose arrangement is independent of drainage lines, other agents besides rain and surface streams have been concerned in its development.

SPECIAL FEATURES RESULTING FROM SPECIAL CONDITIONS OF EROSION.

Many striking topographic and scenic features result from rain and river erosion. Some of them depend primarily on the conditions of erosion, such as climate, altitude, etc., while others depend largely93 on the structure and resistance of the rock. Between these two classes there is no sharp line of demarkation. Illustrations of two types, dependent largely but by no means wholly on conditions independent of the rock, are cited at this point. Others will be mentioned in other connections.

Bad-land topography.—To a type of topography developed in early maturity in certain high regions where the rock is but slightly, though unequally, resistant, a special name is sometimes given. Such regions are termed bad lands. Some idea of bad-land topography is gained from Figs. 75 to 78. Bad-land topography is found in various localities in the West, but especially in western Nebraska and Wyoming, and the western parts of the Dakotas. The formations here are often beds of sandstone or shale, alternating with unindurated beds of clay. Climatic factors are also concerned in the development of bad-land topography. A semi-arid climate, where the precipitation is much94 concentrated, seems to be most favorable for its development. The bad-land topography is most striking in early maturity.

Special forms of valleys; canyons.—Various conditions influence the size and shape of valleys, especially in the early stage of their development. If the altitude of the land be great, the gradient of the streams at this stage will be high. A high gradient means a swift stream, and a swift stream erodes chiefly at its bottom. High altitudes therefore favor the development of deep valleys. Such valleys will be narrow if the conditions which determine widening are absent or unfavorable. Since slope wash is one of the main factors in the widening of valleys, an arid climate favors the development of narrow valleys, if there be sufficient water to maintain a vigorous stream. Narrowness and steepness of slopes will also be favored if the valley is cut in rock which is capable of standing with steep faces. Thus 98a stream may develop a narrow valley in indurated rock where it would not do so in loose gravel, and, other things being equal, it will develop a narrower valley in rock which is horizontally bedded than in rock the beds of which are inclined. Aridity, high altitude, and the proper sort of rock structure therefore favor the development of canyons, and many of the young valleys in the western part of the United States where these conditions prevail, belong to this class.

While all canyons are valleys, most valleys are not canyons. The distinction between a canyon and a valley which is not a canyon is not sharp. The canyon depends for its distinctive character on the relation of depth, width, and angle of slope to one another; but any definition of the depth, width, and angle of slope necessary to constitute a valley a canyon is arbitrary.[31] In popular usage the rule seems to be that if a valley is sufficiently deep, narrow, and steep-sided to be distinctly striking, it is called a canyon in regions where that term is in use. Whether a valley is deep, narrow, and steep-sided enough to be striking clearly depends on the observer. The Colorado Canyon (Figs. 79 and 80) is the greatest canyon known, but it is rarely more than a mile deep, and where its depth approaches this figure it is often eight, ten, or even twelve miles wide from rim to rim. Its width at bottom is little more than the width of the stream; that is, a few hundred feet. Its cross-profile throughout much of its course is therefore not in keeping with the conventional idea of a canyon. With a depth of one mile and a width of eight, the slope, if uniform, would have an angle of less than 15°. Such a valley is represented in Fig. 81. As a matter of fact the slopes of a canyon are not commonly uniform. The slopes represented in Fig. 82 correspond more nearly than those of Fig. 81, to the actual slopes of the Colorado Canyon. The inequalities of slope are occasioned by the inequalities of hardness. It is perhaps needless to say that to an observer on the rim of the canyon the slopes seem several times as steep as those shown in the diagrams.

Like all valleys which are narrow relative to their depth, the Colorado Canyon, great as it is, is a young valley; for it represents but a 99small part of the work which the stream must do to bring its drainage basin to base-level.

While aridity and altitude are conditions which favor the development of canyons, as shown by the fact that most canyons are high and dry regions, they are not indispensable. Niagara River has a canyon below its falls (Pl. IX), and the surrounding region is neither high nor arid. The narrow part of the valley has been developed by the recession of the falls, and is so young that side erosion has not yet widened the valley or lowered its angle of slope to such an extent100 as to destroy its canyon character. This canyon is often called a gorge, a term frequently applied to small valleys of the canyon type.

Plate X shows portions of the canyons of the Yellowstone and the Colorado rivers respectively. In the first the contour interval is 100 feet, and in the second, 250 feet. The horizontal scale is ¹⁄₁₂₅₀₀₀ (about 2 miles to the inch) in the first, and ¹⁄₂₅₀₀₀₀ in the second. These scales should be borne in mind in interpreting the map.

Falls, rapids, narrows, and other peculiar features, due primarily to inequalities in the hardness of the rock affected by erosion, will be considered later.

THE STRUGGLE FOR EXISTENCE AMONG VALLEYS AND STREAMS.

It is not to be inferred that every gully becomes a valley, nor that every small valley becomes a large one. Among valleys, as among living things, there is a struggle for existence, and fitness determines 101growth and survival. At an early stage of its erosion history the number of small valleys in a given area is often great, while at a later stage the number is less and the size of the survivors greater.

PLATE IX.

U. S. Geol. Surv.

Scale, 1+ mile per inch.

PLATE X.

U. S. Geol. Surv.

Scale, 2+ mile per inch.

U. S. Geol. Surv.

Scale, 2+ mile per inch.

One phase of the struggle for existence is often well illustrated on a freshly exposed slope of clay. The number of miniature gullies which develop on such a slope, even in a single shower, may be very large (Fig. 84); but the history of many of them is ephemeral. If two adjacent ones are of unequal depth the widening of the deeper narrows and finally eliminates the divide between them, and the two become one (Fig. 85).

Another phase of the struggle for existence is shown in other situations. Examination of a good map of the north shore of Lake Superior or the west shore of Lake Michigan shows a large number of small streams and gullies (Fig. 1, Pl. IV). The valleys are short and narrow, and between and beyond them are considerable areas untouched by erosion. The drainage near the lake is therefore young, and each of the small valleys is growing. This condition of things is perhaps typical of that which has been, is, or will be along the average coast at a certain stage in its erosion history. No equal stretch of coast-line where erosion is far advanced can boast of a number of large rivers comparable to that of the many small ones along the coasts mentioned. It therefore seems evident that of these many small streams a few only will attain considerable size.

Some of the methods by which the growth of the many is arrested are easily understood. Some of the young valleys on a given coast will work their heads back into the land faster than others because of inequalities of slope and material. This will be true of the tributaries no less than of their mains. If valleys develop in ways other than by head erosion (see p. 73) the chances are also against their equality of growth. If two streams, such as a and c, Fig. 86, develop faster than the intermediate stream b, it is clear that their tributaries may work back into the territory which at the outset drained into b, so as to cut off the supply of water from the latter stream (compare a′b′c′, Fig. 87). As a result, the growth of b will be checked, and ultimately stopped. Similarly other valleys, such as f, will get the better of their neighbors, and many of the competitors, as b, d, e, and g will102 soon drop out of the race. Between the stronger streams competition still goes on. If a′ and f′ develop faster than c′ its prospective drainage territory will be preëmpted by its rivals (compare Figs. 87 and 88). Thus as the result of the unequal rate at which valleys are lengthened, the larger number of those which come into existence are arrested in their development. As a result of growth in the manner indicated, the basins of even the large streams remain narrow at their lower ends while they expand above. This is the usual form of a drainage basin the development of which has been normal.

103

Did valleys grow in length only, competition would not destroy the small ones; it would simply limit them. But valleys widen as well as lengthen, and by widening, adjacent valleys may eliminate the divide between them and become one. The elimination of the intervening ridge may be by lateral planation (p. 82), or, if the valleys be of unequal depth, by slope wash (see Fig. 85). By these and other processes many young valleys are dwarfed, and many others are destroyed.

Piracy.—Streams do not always hold the courses which they establish for themselves at the outset. If the valley occupied by the stream a, Fig. 89, is deepened more rapidly than the valley occupied by b, a tributary from the former, c, may work back across the inter-stream area to e and steal the head waters of that stream (Fig. 90). The tributary which does the stealing is known as a pirate. Stream f104 (Fig. 90) is said to be beheaded, and its upper portion, de, diverted. The beheaded stream is diminished in volume; or if its total supply of water came in above the point of tapping it would disappear altogether.

The process may not end even here. If after the diversion of de the point in the channel to the left is lowered faster than the channel of the beheaded stream f, the divide between dg and the head of f (Fig. 90) will be shifted down the valley of the latter, as shown in Fig. 91. The shifting will go on until the divide reaches a position of stability, that is, until erosion on its opposite sides is equal.

The foregoing case may be called foreign piracy because the valleys of different systems are concerned. Domestic piracy may also take place, as illustrated in the accompanying diagrams (Figs. 92 and 93). Here a tributary to a crooked river may develop, working back until it taps the main at a higher point, thus straightening the course of the stream. The change takes place only when the highest point in the tributary valley is brought below the surface of the water in the main stream at the point where the tapping takes place. This would be likely to occur only after the main stream had attained a low gradient, for so long as it is deepening its channel notably, the small amount of water flowing through the tributary valley would not be likely to bring it down to the level of the main. In any case the flow of water from the main stream through the new valley would be likely to be started during flood, and at such time the erosion in the new channel105 would be great. The complete and final diversion of the stream through the new channel might be a slow process.

Piracy may occur where the material in which the valleys are cut is homogeneous; but, as will be seen later, heterogeneity of material, by determining unequal rates of erosion, stimulates the piratical proclivities of streams.

An actual case of piracy is shown on Plate XI. North and South Lakes formerly drained westward to the Schoharie Creek, the present head of which is in the extreme northwest corner of the map. The head of Kaaterskill Creek, which had a much higher gradient, worked back and captured the head of the westward-flowing stream, diverting the drainage from North and South Lakes to itself. Schoharie Creek was thus beheaded.

Plaatekill Creek, near the south limit of the map, appears to have beheaded the creek flowing west and northwest, similarly diverting its head waters. The Dells, Wis., quadrangle (U. S. Geol. Surv.) affords an illustration of domestic piracy.

RATE OF DEGRADATION.

The amount of mechanical sediment which the Mississippi River carries to the Gulf of Mexico is estimated to represent a rate of degradation for the Mississippi basin of about one foot in 5000 years. But the mechanical sediment carried to the Gulf does not really represent the total degradation of the basin, for the water which sinks beneath the surface is dissolving more or less rock substance, especially lime carbonate. This material is carried to the sea in solution, and does not appear in the sediment on which the above estimate is based. Taking into account the matter dissolved by the water and carried to106 the sea in solution, the average rate of degradation for the Mississippi basin is estimated at one foot in 3000 to 4000 years.

It is not to be inferred that this rate is uniform, or even that erosion at any rate whatsoever is taking place in all parts of the basin. Such is not the fact. On the whole the rate of erosion is doubtless greatest toward the margins of the basins where the land is in its topographic youth or early maturity. It is notably less in the middle courses of the valleys, and erosion is locally exceeded by deposition along the lower courses of the Mississippi and some of its main tributaries.

The average elevation of North America is not accurately known, but it is probably not far from 2000 feet. If the present rate of degradation, say one foot in 3500 years, were to continue, it would take something like 7,000,000 years to bring the continent to sea-level. But this rate of degradation could not continue to the end, for as the continent became lower streams would become sluggish and erosion less rapid. Long before the continent reached base-level the rate of degradation, so far as dependent on mechanical erosion, would become so slow that the time necessary to bring the continent to sea-level would be almost inconceivably prolonged. Furthermore, it is quite possible that the land is suffering, or is liable to suffer, uplift, relative or absolute. If the rate of rise were equal to the rate of degradation the average height of the continent would of course not be affected.

The amount of sediment carried by streams in suspension varies notably according to the stage of the water. During a year when the stream was under careful study the Mississippi at Carrollton (Miss.) was found to carry ¹⁄₆₈₁ of its weight of sediment during the high-water stage of June, and ¹⁄₆₃₈₃ during the low-water of October, the average for the year being ¹⁄₁₈₀₈. The average of a greater number of records gives about ¹⁄₁₅₀₀ as the average ratio between the weight of the sediment and the weight of the water. This corresponds to about ¹⁄₂₉₀₀ by volume, the average specific gravity being about 1.9. The amount of material carried in the upper part of the water was notably less than that carried at greater depths, but that carried midway between top and bottom was about the same as that carried at the bottom.[32]

The discharge of the Mississippi River is about 19,500,000,000,000 cubic feet of water per year, and the sediment it carries in suspension is estimated to weigh about 812,500,000,000 pounds. This is equivalent to about 6,714,694,400 cubic feet. It is estimated that about 750,000,000 cubic feet of sediment is rolled along the bottom, giving a total of 7,468,694,400 cubic feet as the aggregate annual load carried to the Gulf by the river. This would be adequate to cover an area one square mile in extent to the depth of 268 feet per year.

PLATE XI.

U. S. Geol. Surv.

Scale, 2+ mile per inch.

PLATE XII.

U. S. Geol. Surv.

Scale, 2+ mile per inch.

U. S. Geol. Surv.

Scale, 1+ mile per inch.

107

The following table[36] gives the percentage of material carried in suspension by various rivers:

The composition of rain-water falling near London, as determined by analysis, was as follows:[37]

A comparison of the composition of rain-water with that of springs and rivers gives some idea of the solvent work of water. From a study of the water of nineteen of the principal rivers of the world Murray has compiled the following table[38] showing the amount of mineral matter in average river water:

Murray also estimates that the aggregate amount of water flowing into the sea annually is about 6528 cubic miles, which, on the above basis, would carry about 4,975,000,000 tons of mineral matter in solution.

A large number of analyses of waters of rivers from the United States and Canada give an average of about .15,044 part in a thousand of mineral matter in solution, more than one-third being CaCO₃. The average amount of mineral matter in solution in 48 European streams cited by Bischoff[40] is .2127 part in a thousand, of which CaCO₃ is rather more than half. The average mineral matter in solution in 36 rivers cited by Roth[41] (including some of those tabulated by Bischoff) is .2033 part in a thousand, of which CaCO₃ is slightly less than one-half.

An average for American and European rivers, so far as determinable from data at hand, is about .1888 part in a thousand in solution, of which CaCO₃ is slightly less than one-half. These last figures are probably not very far from an average for river water in general.

The following table shows the total amount of solids carried in solution by the rivers indicated:[42]

ECONOMIC CONSIDERATIONS.

Certain considerations of human interest in connection with river erosion are worthy of note. When a drainage system has reached its 109mature stage its basin has the roughest topography which it will have at any time during that cycle of erosion. At that stage, therefore, road construction is relatively difficult. If the relief be great, roads must follow the valleys, or the crests of the ridges between them, if they would avoid heavy grades. In such regions roads are usually few and crooked.

The stage of development of valleys has an influence on the navigability of their streams. Streams well advanced in life are much more readily navigable than young ones, because their grades are lower and their volumes of water greater. Old streams, on the other hand, are sometimes depositing sand or silt along their lower courses to such an extent as to interfere with navigation.

At certain stages of their development the power of streams is more easily utilized than at others. Young streams, depending as they do for their supply on the rainfall of a limited area, are likely to be fitful in their flow, and therefore unreliable as a source of power. This is especially true where the precipitation is unequally distributed, and where the slopes are steep and free from forests. Because of their great volume, old and large streams, though sluggish, have great power, but it is less easily controlled. Where streams are large enough to be navigable industrial considerations often prevent the utilization of their power, the streams being more serviceable as highways than as sources of power. Other things being equal, it follows that streams are most available for water-power when they are large enough to have a moderately steady flow, and not so large as to be beyond ready control, or to be valuable for purposes of navigation.

Streams are subject to more disastrous floods in some stages of their development than in others. Floods resulting from heavy rains are likely to be greatest where the slopes above the drainage lines are on the whole greatest, for this is the condition under which the water is most quickly gathered into the drainage channels. The most disastrous floods, humanly speaking, are those which affect wide-bottomed valleys, where the flats are settled. In such cases a relatively slight rise may flood very extensive areas. In such valleys the most disastrous floods are generally in the spring, when the waters from the melting snows of the preceding winter are being discharged.[43] Many other 110considerations enter into the problem of floods. The presence of forests and other forms of vegetation on the slopes retards the flow of water into the valleys, and so tends to prevent floods, or at any rate to make them less severe. Porous soil and subsoil, or in their absence porous rock, absorb the rainfall, and prevent its prompt descent into the valleys and so tends to prevent or diminish floods.

The acreage of arable land within a given area stands in some relation to its drainage development. At an early stage in its erosion history, before an upland has been dissected by valleys, nearly all of it may be arable. Later, when drainage is at its maturity, and when hillsides and ridge slopes constitute a large part of the area, there is probably the least acreage of arable land. This is especially true if the slopes are so steep as to allow the soil to be readily washed away. At a still later stage, when the valley bottoms have become wide and the slopes of the ridges and hills so reduced as to be available, the area of cultivable land is again increased.

Marshes, ponds, and lakes have some bearing on the resources and industries of a region, and they stand in a more or less definite relation to the stage of erosion in which a region finds itself. In its youth ponds and lakes may occupy much of the surface; in its maturity they will have been largely drained.

These suggestions are sufficient to show that the topography of a region, even in so far as shaped by erosion, touches human interests at many points.

ANALYSIS OF EROSION.[44]

Erosion is the term applied to all the processes by which earthy matter or rock is loosened and removed from one place to another. It consists of three sub-processes, namely, weathering, transportation, and corrasion.

Weathering.

The term weathering is applied to nearly all those natural processes which tend to loosen or change the exposed surfaces of rock. The lettering of inscriptions on exposed marble becomes fainter and fainter as time goes by, and finally disappears, because the rock in 111which the letters were cut has weathered away. Some of it has crumbled off as the result of the expansion and contraction induced by changes of temperature, and some of it has been dissolved by the rain which has fallen upon it. In this case the weathering is effected partly by the atmosphere and partly by water. These are the chief, but not the only agents concerned in the general processes of weathering. Those phases of weathering which are the result of the activities of the atmosphere, whether physical or chemical, have been discussed in connection with the atmosphere (pp. 42 and 54).

The rain which falls upon the surface of exposed rock, and that which sinks through the soil to the solid rock below, dissolves, even if slowly, some of the rock constituents. Each constituent of a rock composed of several minerals may be looked upon as a binding material for the others. When one is dissolved the rock crumbles, much as mortar does when the lime which cements the sand is dissolved.

The solution of mineral matter by ground water, as well as the other chemical changes it effects, is greatly augmented by the impurities, especially carbonic and other organic gases, dissolved by the water from the atmosphere and the soil. The commonest chemical changes effected by the joint action of water and air, oxidation and carbonation, have been referred to in Chapter II. Hydration is more exclusively the work of water, and is one of the commonest processes of rock change, and often of rock disintegration. Numerous other less simple chemical changes resulting from the activities of ground water are constantly in progress, and in so far as they lead to the disintegration of rock are processes of weathering. Many chemical changes involve notable changes in volume of the mineral matter concerned. Merrill has calculated that in the conversion of the granitic rock of the vicinity of Washington, D. C., into soil, its volume has been increased 88 percent., largely as the result of hydration.[45] Even when the chemical changes do not themselves directly involve the disintegration of the rock, the accompanying increase of volume is sometimes sufficient to cause its physical disruption. This also may be regarded as a phase of weathering.

The weathering accomplished by water, or under its influence, proceeds at rates which vary with the composition of the rock, the amount and composition of the water, the temperature, and certain 112other factors less susceptible of brief statement. The weathering effected by ground water has a wider range both in area and depth than that due to changes of temperature, for while the latter is effective only where temperature changes are considerable, and where coherent material lies at the surface (p. 45), the former is operative to all depths to which water sinks.

There are other processes of weathering not due directly either to the atmosphere or to water. The roots of trees and smaller plants frequently grow into cracks of rocks, and, increasing in size, act much like freezing water (p. 45) in similar situations. This wedge-work of roots is a phase of weathering.

From the faces of steep cliffs masses of rock frequently fall. However dislodged, their descent is effected by gravity. The quantities of débris at the bases of many cliffs, forming slopes of talus (Fig. 94), testify to the importance of the action of gravity in getting material from higher to lower levels. Another phase of gravity-work is shown in Fig. 95. Here, under the influence of gravity and expansion and contraction, due to freezing and thawing and wetting and drying, the surface material is creeping down slope. In the process the rock113 is being broken. The process illustrated by the figure involves weathering as well as other factors.

The foregoing are among the commoner processes of weathering, although they do not exhaust the list. The more active and tangible processes by which surface rocks are broken up, such as wave wear, river wear and glacier wear, are processes of corrasion. The mechanical wear effected by wind-driven sand might be considered either as corrasion or as weathering. It is more likely to be regarded as corrasion if the amount of wear is considerable enough to be obvious. Rock is sometimes decomposed by the chemical action of hot vapors, gases, and waters rising to the surface from considerable depths. This is often seen in volcanic regions. A conspicuous illustration is seen in the canyon of the Yellowstone in the National Park. Decay of this sort is perhaps not properly weathering, but is not always readily distinguished from it.

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The importance of weathering in the general processes of erosion is shown in many ways. In regions where the mantle rock is the product of the decay of the solid rock beneath, and such regions constitute a large portion of the earth’s surface, the soil and subsoil represent the excess of weathering over transportation. Since most of the earth’s surface is covered with soil to a greater or less depth, it is clear that, on the whole, weathering keeps ahead of transportation. Again, it is clear that the loosening of rock by weathering greatly increases the erosion which a given amount of moving water can accomplish. Not only this, but weathering plays a much more important rôle in the development of valleys than is commonly realized. This is best illustrated by the valleys of young swift streams. The valley which is not at its top ten times as wide as its stream is rare. The stream which has such a canyon has been cutting chiefly at its bottom. Ignoring its lateral corrasion, which is slight, the valley which it would cut would have a width equal to its own. This is illustrated by Fig. 96. Weathering in its broadest sense is largely responsible for the width of such a valley, in so far as it exceeds the width of the stream. The work of weathering, slope wash, etc., has been to get the material which originally lay between a, b, and c down to the stream. The stream has then carried it away. The above illustration would not apply to old and sluggish streams, for they, by their meandering, widen their valleys independently of weathering.

Weathering is a part of erosion, but only a part. In so far as it is effected by solution the process involves the transportation of that which is dissolved to some other point. Transportation is also involved to some extent in the other processes of weathering, but the central idea of the processes embraced under this term is the loosening and disrupting of rock by which it is prepared for transportation.

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Transportation.

The second element of erosion is transportation. The transportation of mechanical sediment is to be distinguished from the transportation of materials in solution. In so far as mineral matter is dissolved it becomes, so far as flowage is concerned, a part of the stream. If the quantity dissolved were large it might influence the mobility of the water, but the amount is usually too slight to influence the flow sensibly.

The sediment transported by a stream is either rolled along its bottom or carried in suspension at some higher level. The coarser materials (gravel and sand) are carried chiefly in the former position, and the finer (silt and mud) largely in the latter.

Transporting power and velocity.—The transporting power of running water depends on its velocity. The formula expressing the relations between them is as follows: Transporting power, t, varies as the sixth power of velocity, v, (tαv⁶); that is, doubling the velocity of the stream increases its transporting power 64-fold. Strictly speaking, this means that if a stream of given velocity is just able to move a stone of a given size, a stream with double that velocity will be just able to move a stone of the same shape 64 times as large as the first. This may be graphically illustrated as follows: Let a current be supposed just able to move the cube a (Fig. 97). If the current be doubled, twice as much water will strike the same surface with twice the force in the same time; that is, the force exerted on the cube a will be quadrupled. It will, therefore, be able not only to move the one cube, but it will be able to move three other cubes (b, c, and d) besides (Fig. 98). The same current against any other equal surface would also be able to move four small cubes, and there are sixteen such surfaces on the face of the large cube (Fig. 99). It follows that the dimension of the cube which the stream with the doubled velocity can move is four times as great as that of the cube which the original current could move, and the cubical contents of such a cube is 64 times as great as that of the first (64 = 2⁶) (Fig. 99). Swift streams, therefore, have enormously greater power of transportation than sluggish ones. It does not necessarily follow that transportation keeps pace with transporting power; that depends on the accessibility of materials suitable for transportation. A stream of great transporting power, like the Niagara at its rapids, may carry little sediment, because there is little to be had.

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The velocity of a stream depends chiefly on three elements—its gradient, its volume, and its load, (i.e., the sediment it is moving). The higher the gradient the greater the volume, and the less the load the greater the velocity. The relation between gradient and velocity is evident; that between volume and velocity is illustrated by every stream in time of flood, when its rate of flow is greatly increased. The relation between velocity and load is less obvious, but none the less definite. Every particle of sediment carried by a stream makes a draught on its energy, and energy expended in this way reduces the velocity. The draught on a stream’s energy of a particle carried in suspension is measured by its mass into the distance it would fall in a unit of time in still water. It follows that a large particle makes a stronger draught on a stream’s energy than the same amount of material in smaller pieces. It follows also that the comminution of sediment facilitates transportation in much more than a simple ratio, for not only can a given amount of energy carry more fine material than coarse, but a larger proportion of a stream’s energy can be utilized in the transportation of the fine.

How sediment is carried.—Coarse materials, such as gravel stones, are rolled along the bottoms of the swift streams which carry them. Their movement is effected by the impact of water. The same is true to a large extent of sand grains, especially if they be coarse. So far as concerns the material rolled along the bottom it is to be noted that a stream’s transporting power is dependent on the velocity of the water at its bottom. This is much less than the surface, or even the average velocity. The particles of fine sediments, such as silt and mud, are frequently carried by streams quite above their bottoms, as shown by the roiliness of many streams. A particle of mud is usually117 a small bit of mineral matter, the specific gravity of which is two or three times that of water. Why does it not sink through the water and come to rest at the bottom of the stream, or suffer transportation as the gravel does?

A particle of sediment in running water is obviously subject to two forces, that of the current which tends to move it nearly horizontally down-stream, and that of gravity which tends to carry it to the bed of the stream. In Fig. 100, the arrows ab and ac represent respectively the relative force of gravity and a current of 5 miles per hour. As a result of these two forces the particle would tend to descend in the general direction of ad, a line which represents the resultant of these forces, though not the exact path which a particle acted on by them would take in water. If a river were the simple straightforward current which it is popularly thought to be, a particle in suspension would reach its bottom in the time it would take to sink through an equal depth of still water, for the descent would be none the less certain and none the less prompt because of the forward movement of the water. The current would simply be a factor in determining the position of the particle when it reached the bottom, not the time of reaching it. Very fine particles, like those of clay, though having the same specific gravity as grains of sand, would sink less readily than coarser ones, because they expose larger surfaces, relative to their mass, to the water through which they sink. But even such particles, unless of extraordinary fineness, would presently reach the bottom if acted on only by a horizontal current and gravity. Since even sediment which is not of exceeding fineness is kept in suspension it is clear that some other factor is involved. This is found, in part at least, in the subordinate upward currents in a stream.

Where a bowlder occurs in the bed of a stream (Fig. 101) the water which strikes it is in part forced up over it. If there be many bowlders the process is frequently repeated, and the number of upward currents118 is great. Any roughness will serve the same purpose, and every stream’s bed is rough to a greater or less extent. Where there are roughnesses at the sides of a channel, currents are started which flow from them toward the center. The varying velocities of the different parts of a stream serve a similar purpose. The curves in a river tend to give the water a rotatory movement. A river is therefore to be looked upon not as a single straightforward current, but as a multitude of currents, some rising from the bottom toward the top, some descending from top to bottom, some diverging from the center toward the sides, and some converging from the sides toward the center. The existence of these subordinate currents is often evident from the boiling and eddying readily seen in many streams. It is, of course, true that the sum of the upward currents is always less than the sum of the downward, so that the aggregate motion of the water is down slope; but it is also true that minor upward currents are common. Sediment in suspension is held up chiefly by such currents, which, locally and temporarily, overcome the effect of gravity. The particles in suspension are constantly tending to fall, and frequently falling; but before they reach the bottom many of them are seized and carried upward by the subordinate currents, only to sink and be carried up again. Even if they reach the bottom, as they frequently do, they may be picked up again. It is probable that every particle of sediment of such size that it would sink readily in still water is dropped and picked up many times in the course of any long river journey, and its periods of rest often exceed its periods of movement.

Independently of the subordinate currents, the different velocities of the different parts of a stream tend to keep materials in suspension by exerting different pressures on the different sides of suspended particles.[46]

River ice sometimes facilitates the transportation of débris which the water alone could not carry. The ice freezes to bowlders in the 119banks of the streams, to those which are partially submerged, and sometimes to those altogether submerged beneath slight depths of water. When the ice breaks up in the spring such bowlders, buoyed up by the ice, may be floated far down the stream. The influence of ice in this connection is most considerable in high latitudes, but it is of consequence as far south as Virginia, where the river deposits sometimes contain bowlders which the unaided streams could not have carried. Ground ice sometimes forms about bowlders in the bottoms of streams, especially in the quiet pools of turbulent rivers, and floats them to the surface before the surface itself is frozen.[47] In the floods of spring rivers often spread their ice widely over their flood-plains. It is sometimes massed in constricted portions of valleys so as to form great dams, the breaking of which is attended with great destruction.

Corrasion.

Abrasion.—The wear effected by running water is corrasion. So long as the materials to be carried away are incoherent it is easy to see how running water picks them up and carries them forward. The water which gathers in the depressions on the slope of a cultivated field gathers earthy matter from the surface over which it passes, even before it is concentrated into rills, and the rills continue the process. Thus the loose materials of the surface are gathered at the very sources of the streams, and the amount of sediment in the water after a heavy shower, even at the head of the stream, may be great. The run-off from the slopes of any valley in any part of its course likewise brings sediment to the stream, which gathers more from its bed whereever it flows with sufficient velocity over incoherent material. Streams also undercut their banks, and receive new load from the fall of the overhanging material.

By far the larger part of the sediment acquired by a normal stream is made up of material loosened in advance by the processes of weathering. The stream, or the waters which get together to make the stream, find them ready-made; but rivers frequently wear rock which is not weathered, for the principal valleys of the earth’s surface are cut in solid rock, and many of them in rock of exceeding hardness. How does the stream wear the solid rock?

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When a stream flows over a rock bed, the wear which it accomplishes depends chiefly on the character of the rock, the velocity of the stream, and the load it carries. If the rock be stratified and in thin layers, and if these thin layers be broken by numerous joints at high angles to the stratification planes, the impact of the water of a clear stream of even moderate strength may be effective in dislodging bits of the rock. This condition of things is often seen where streams run on beds of shale or slate. If the rock be hard and without bedding-planes and joints, or if its layers be thick and its joints few, clear water will be much less effective. If the surface of the rock be rough, the mechanical action of a swift stream of clear water might still produce some effect on it; but if massive hard rock presents a smooth surface to a clear stream, the mechanical effect of even a swift current is slight.

This general principle is illustrated by the Niagara River. Just above the falls the current is swift. When the river is essentially free from sediment, the surface of the limestone near the bank beneath it is sometimes distinctly green from the presence of the one-celled plants (fresh-water algæ) which grow upon it. The whole force of the mighty torrent is not able to sweep them from their moorings. Were the stream supplied with a tithe of the sand which it is capable of carrying, it would not take many hours, and perhaps not many minutes, to remove the last trace of vegetation. This illustration furnishes a clue to the method by which the erosion of solid rock in a stream’s bed is effected.

It has been seen that the ingathering waters which make a stream often have abundant sediment before they reach well-defined stream channels, and that the streams continue to gather sediment whereever their beds are composed of material which is readily detached. The sediments which the stream carries are the tools with which it works. Without them it is relatively impotent, so far as the abrasion of solid rock is concerned; with them, it may wear any rock over which it passes (Fig. 102).

We have next to inquire the methods by which running water uses its tools in the excavation of valleys. When gravel is rolled along in the channel of a stream there is friction between it and the bed over which it moves. If the pebbles be as hard as the bed over which they are rolled their movement must result in its wear, and even if they be softer more or less wear takes place. As the moving stones wear the121 rock of the stream’s bed they are themselves worn by impact with it and with one another. In all cases the softer material suffers the more rapid wear. The first effect of wear on materials in transportation is the reduction of their rugosities of surface. The projecting points and sharp angles are worn off, and the stones are reduced to rounded water-worn forms. The particles broken off make grains of sand, or, if very fine, particles of silt or mud. Even after a stone has been rounded it is subject to further wear and reduction, and in the course of time may be literally worn out.

The sediment carried in suspension, as well as that rolled along the bottom, may wear the rock bed of a stream. When a grain of sand in suspension escapes from an upward moving current it may not sink quietly. If it be caught by a downward current it may be made to strike a blow on the bed of the stream, and the effect of the blow is to wear the surface which receives it. The larger the grain and the stronger the current the greater the wear.

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The ceaseless repetition of the blows struck by the material in suspension, or rolled on its bottom, hour after hour, day after day, and year after year, will accomplish sensible results. In the long course of the ages this process has excavated deep valleys. Concomitant processes are largely concerned in making valleys wide, but the depth of valleys cut in solid rock is chiefly the result of the impact and friction of the sediment in transportation.

The wear effected in this way is not proportional to the number of blows struck. Since every pebble and every grain of sand carried diminishes the velocity of a stream, and since with diminished velocity the force of the blows struck is diminished, it follows that the blows may become so weak, as the result of their multiplication, as to be ineffective. The larger the load, therefore, which the stream carries, the more the tools with which it has to work, but the less effectively can it use them; and the load may be so far increased as to destroy its corrasive power altogether. On the other hand, the smaller the load of the stream the greater its velocity and the more effectively will its tools be used; but their number may be so far reduced that their aggregate effect is slight. To accomplish the greatest results on a bed of solid rock a stream must have tools to work with, but must not be so heavily burdened as to interfere with its effective use of them.

Whatever the cause of their unequal velocities swift and slow streams corrade their valleys differently. The erosion of a swift stream is chiefly at the bottom of its channel. The sluggish stream lowers its channel less rapidly, while lateral erosion is relatively more important. The result is that slow streams increase the width of their valleys more than the depth, while swift streams increase the depth more than the width. It follows that slow streams develop flats, while swift ones do not. Not only is a slow stream more likely to have a flat, and therefore a better chance to meander, but it is more likely to take advantage of opportunities in this line, for a slow stream gets out of the way for such obstacles as it may encounter, while a swift stream is much more likely to get obstacles out of its way.

Special phases of corrasion are introduced where waterfalls and other peculiarities dependent on inequalities of rock resistance occur.

Solution.—In most cases the solution effected by a stream is much less important than its mechanical work. Only when conditions are unfavorable to the latter, is solution the chief factor in the excavation of a valley. This may be the case where a stream’s bed is over soluble123 rock, such as limestone, and where the stream is clear, or its gradient so low that its current is sluggish. The solvent power of water is not influenced by the presence of sediment, though the presence of sediment offers the water a greater surface on which to work.

CONDITIONS AFFECTING THE RATE OF EROSION.

In considering the rate of erosion, both the work of the stream in its valley and that of the general run-off are to be considered. The conditions which favor the most rapid erosion in a stream’s channel are not necessarily those which determine most rapid degradation in the basin outside of the valley.

The Influence of Declivity.

In general the greater the declivity the more rapid the rate of erosion, whether in the stream’s channel or on the slopes above it. The truth of this conclusion is illustrated by the great erosive power of swift streams as compared with slow ones.

It does not follow, however, that high declivity favors each element of erosion. The effect of declivity on weathering is far from simple. For example, great declivity, by allowing more of the rainfall to flow off over the surface, and by causing it to flow off more promptly, restricts the work of solution, and therefore of decomposition, both at the surface and beneath it. High declivity is also unfavorable to the growth of vegetation, and so to the wedge-work of roots. On the other hand, a given amount of wedge-work of roots and ice is more effective where the slope is steep than where it is gentle, for such materials as are loosened descend the slopes more readily. The prompt removal of weathered materials, by exposing fresh surfaces of rock, accelerates weathering. The total amount of weathering may therefore not be diminished by the increase of slope, even though certain of its processes are hindered.

The effect of high declivity on transportation, the second element of erosion, is too patent to need explanation.

Corrasion likewise is favored by high declivity, for the abrasive power of a stream increases as the square of its velocity. With corrasive power increased, corrasion will also be increased if the water124 has tools to work with. Since high declivity greatly increases both the transporting and the corrasive power of running water, and favors certain elements of weathering, it is clear that the aggregate effect of high declivity is to favor erosion, whether in the channel of the stream or on the general surface of its drainage basin.

The Influence of Rock.

The physical constitution, the chemical composition, and the stratigraphy of a rock formation, influence the rate at which it may be broken up and carried away. Clastic or fragmental rocks are usually stratified and made up of cemented pebbles (conglomerate), sand grains (sandstone), or particles of mud (shale). Igneous rocks, such as granite, are massive instead of stratified, and are usually made up of great numbers of interlocking crystals which bind one another together. Some crystalline rocks, such as schists, though not stratified, possess cleavage, which has much the effect of stratification, so far as erosion is concerned. All rocks are affected by systems of more or less nearly vertical cracks called joints. All these structures have their influence upon the rate of degradation.

Physical constitution.—Clastic rocks may be firmly cemented, or their constituents may be loosely bound together. The less the coherence the more ready the disintegration, and the finer the particles the more easily are they carried away. When the particles in transportation are angular they effect more wear on the bed over which they move, and on one another, than when they are round. The difference is great where the particles are large, and little where they are very small. If the materials carried be harder than the bed over which they pass, corrasion of the latter is favored.

Chemical composition.—Something also depends on the chemical composition of the rock, since this affects its solubility, and therefore its rate of decomposition. The more soluble the rock the larger the proportion of it which will be taken away in solution; but it does not follow that the most soluble rock will be most rapidly eroded, since the rate of erosion depends on abrasion as well as solution, and a rock which is readily soluble, as rocks go, may be less easily abraded than a rock which is made of discrete and insoluble particles bound together by a soluble cement. In such rocks, for example a sandstone in which the grains are cemented together by lime carbonate,125 the solution of the cement sets free a considerable quantity of sand, so that a small amount of solution prepares a large amount of sediment for removal. A stream might cut its valley much more rapidly in such a sandstone than in a compact limestone, though the latter is, as a whole, the more soluble. The constituent minerals of crystalline rocks resist solution and decay unequally, and when any one is dissolved or decomposed the rock crumbles and the less soluble constituents are ready for removal by mechanical means. So long as the material loosened by disintegration is removed, chemical heterogeneity favors erosion; but if the loosened débris is not removed erosion is not favored by chemical heterogeneity. In such a case erosion would be most rapid where the rock was most soluble.

Structure.—The structure of the rock has much to do with the rate of its erosion. Other things being equal, stratified rock is more readily eroded than massive rock, since stratification-planes are planes of cleavage, and therefore of weakness. Taking advantage of these planes the water has less breaking to perform to reduce the material to a transportable condition. For the same reason a thin-bedded formation is more easily eroded than a thick-bedded one.

The beds of stratified rock may be horizontal, vertical, or inclined, and inclined strata may stand at any angle between horizontality and verticality. In indurated formations the rate of erosion is influenced both by the position of the strata and by the relation of the direction of the flowing water to their dip and strike. On the whole the strata which are horizontal, or but slightly inclined, are probably less favorable for rapid erosion than those which are vertical or inclined at considerable angles. This is at least true where the layers are of uniform hardness and the joints infrequent.

Horizontal strata expose fewer cleavage planes to the water flowing over them than strata in any other position. In Fig. 103 the stream which has the profile ad crosses bedding-planes at b and c. In Fig. 104,126 where the beds dip up-stream, many more division-planes are crossed in the same distance. Since bedding-planes are planes of weakness, it follows that horizontal and nearly horizontal strata are not, under ordinary conditions of erosion, in a position favorable for most rapid wear. When strata are horizontal, it makes no difference which way the stream runs, for the current sustains the same relation to the cleavage-planes whatever its course.

In the case of incoherent material the position of the beds, or even their existence, has little influence on the rate of erosion. Such formations are weak in all directions, not simply along bedding-planes.

When the strata are vertical, three distinct cases may arise (Fig. 105). The stream may flow (1) with the strike (aa); (2) at right angles to the strike (bb); or (3) oblique to it (cc) at any angle whatsoever. It is perhaps not possible to say which of these positions is most favorable for erosion, for the character of the rock, the thickness of its layers, its ability to stand with steep slopes, and the strength of the currents concerned, would influence the result. A stream which flows at right angles to the strike (bb, Fig. 105) would cross more cleavage-planes in a given distance than a stream flowing in any other direction, and would strike the outcropping edges of layers at the angle of greatest advantage. A stream flowing along the strike (aa), on the other hand, has better opportunity to sink its channel on cleavage-planes, and the current oblique to the strike (cc), has some of the advantages of each of the others.

When the strata are inclined five cases may arise. (1) The stream may be parallel to the strike (aa, Fig. 106), when it makes no difference which way the current flows; it may be at right angles to the strike (bb′), and (2) flowing with the dip (toward b′), or (3) against it (toward b); it may be oblique to the strike, and flowing (4) in the general direction of dip (toward c′); or (5) in the opposite direction (toward c). As before, the stream flowing at right angles to the strike would cross the largest number of layers in a given distance, and so have an opportunity to take advantage of more cleavage-planes than127 a stream in any other position. But in the case of inclined strata a new element enters into the problem. When the stream flows parallel to the strike, the valley which is in process of deepening is not sunk vertically, but is shifted more or less in the direction of the dip (Fig. 107). This is called monoclinal shifting. The result is that there is a constant tendency to undermine (sap) the valley bluff on the down-dip side, and this process of sapping will, according to its rate, accelerate the growth of the valley, especially in width. Monoclinal shifting is favored by the presence of a hard layer (H), as shown in Fig. 107, if this stratum is the bed of the stream.

In the second and third cases mentioned above, the only difference is in the angle at which the current strikes the outcropping edges of layers and laminæ. The mechanical advantage is with the stream which flows with the dip. In the fourth and fifth cases something will depend on the angle which the stream’s course makes with the strike. In all these cases, as in those where the strata are vertical, much will depend on the thickness and resistance of the layers and on the strength of the currents concerned.

The Influence of Climate.

Climate has both a direct and an indirect effect on erosion. Its direct influence is through precipitation, evaporation, changes of temperature, and wind; its indirect, through vegetation. Like declivity and rock structure, climate does not affect all elements of erosion equally.

The chief elements of climate are temperature, moisture, and atmospheric movements; the principal factors which influence it are latitude, altitude, distance from the sea, direction of prevailing winds, and topographic relations.

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The effects of variations in temperature on rock weathering have already been discussed (p. 43). They are chiefly mechanical, and are seen at their best where the daily range is great.

High temperature favors chemical action, and the weathering of rock by decomposition is at its best in the presence of abundant moisture in regions where the temperature is uniformly high. Furthermore, a warm moist climate favors the growth of vegetation, the decay of which supplies the water with organic acids which greatly increase its solvent power. The climatic conditions favoring mechanical weathering are therefore different from those favoring chemical weathering. High temperature and abundant moisture and vegetation are found in many tropical regions, and here the rock is often decomposed to greater depths, on the whole, than in high latitudes. How far this is the result of rapid weathering, and how far of slow removal, due in part to the protective influence of the plants, cannot be affirmed. If the weathered material is not removed, it will presently become a mantle thick enough to retard the processes which brought it into existence.

So long as the water of the surface and that in the soil remains unfrozen, temperature affects neither corrasion nor transportation. But in middle and high latitudes the surface is frozen for some part of each year. During this time corrasion is at a minimum, for although the streams continue to flow there is relatively little water running over the surface outside the drainage channels, and that little is relatively ineffective. Under some conditions, therefore, temperature affects both corrasion and transportation.

The humidity of the atmosphere has an influence even more important than that of temperature on the rate of erosion, and its influence is exerted on each of the elements of that complex process. A moist atmosphere favors oxidation, carbonation, hydration, and the growth of vegetation, all of which promote certain phases of rock weathering. On the other hand, humidity tends to prevent sudden and considerable variations in temperature, thus checking the weathering effected by this means. Precipitation, the most important single factor in determining the rate of erosion, is dependent on atmospheric humidity. Its amount, its kind (rain or snow), and its distribution in time, are the elements which determine its effectiveness in any given place.

Other things being equal the greater the amount of precipitation the more rapid the corrasion and transportation. Much, however,129 depends on its distribution in time. A given amount of rainfall may be distributed equally through the year, or it may fall during a wet season only. The maximum inequality of distribution would occur if all the rainfall of a given period were concentrated in a single shower. With such concentration the volume of water flowing off over the surface immediately after the down-pour would be greater than under any other conditions of precipitation, and since velocity is increased with volume, and erosive power with velocity, it follows that the erosive power of a given amount of water would be greater under these circumstances than under any other. Furthermore, a larger proportion of the precipitation would run off over the surface under these circumstances than under any other, for less of it would sink beneath the surface and less would be evaporated. If erosive power and rate of erosion were equal terms, this would therefore be the condition for greatest erosion; but erosive power and rate of erosion do not always correspond. If the water falling in this way could get hold of all the material it could carry, extreme concentration of precipitation would be the condition favorable for most rapid erosion. But if the amount of available material for transportation is slight, a large part of the force of the water could not be utilized in erosion. It follows that if there were a large amount of disintegrated material on the surface, erosion would be greater the greater the concentration of precipitation. If, on the other hand, there were but little disintegrated material on the surface, frequent showers, with intervening periods when conditions were favorable for weathering, that is, for preparing material for transportation, might be more favorable for rapid erosion. While the total energy of running water available for erosion under these conditions would be less than before, there might in the long run be more material for transport; for weathering in the presence of moisture, and all that goes with it, might be more effective in preparing material for transportation, than weathering during the long periods of drought which would occur if the precipitation were concentrated to its maximum. Temperature favoring, the uniform distribution of moisture through the year would allow the growth of vegetation, which, although favoring some processes of weathering, retards erosion in general. While therefore it is not possible to say what distribution of rainfall favors most rapid erosion without knowing the nature of the surface on which it is to fall, enough has been said to show that the problem is by no means a simple one. Some of the most striking phases of topography130 developed by erosion, such as those of the Bad Lands (Figs. 75 to 78, and 108), are developed where the rainfall is unequally distributed in time, and too slight or too infrequent to support abundant vegetation.

During its fall, and immediately after, rain is more effective than an equal amount of snow; but the snow may be accumulated through a considerable period of the year, and then melted rapidly, when it has an effect comparable to that which would be produced by the concentration of the rainfall into a limited period of the year. If the ground beneath be frozen when the snow melts (and this is often the case) the erosion accomplished by the resulting water will be diminished.

Except in dry regions, where wind-work sometimes exceeds water-work, the movements of the atmosphere are of less importance directly than precipitation in determining the rate of erosion. But even in regions which are not arid the winds have much to do with the rate of evaporation and the distribution of rainfall, so that their indirect effect is great. Even their direct effects in moist climates are not to be lost sight of, for even here the surface is sometimes dry enough to yield dust and sand, and the uprooting of trees so disturbs the surface as to make earthy débris more accessible to wind and water. Where trees gain precarious footholds on steep slopes, as they often do, they are likely to be overturned as soon as they are large enough to offer131 considerable resistance to the wind, and in the overturning, large quantities of rock are sometimes loosened and carried down the slope by gravity. This phase of destructive work is seen at its best on the walls of gorges, where trees often flourish until their tops project above the rim of the valley.

Through vegetation, climate influences erosion in ways which are easily defined qualitatively, but not quantitatively. Both by its growth (wedge-work of roots) and by its decay (supplying CO₂, etc., to descending waters) it favors certain phases of weathering; but, on the other hand, it retards corrasion and transportation both by wind and water. This is well shown along the banks of streams and on the faces of cliffs, in clay, sand, etc. Its aggregate effect is probably unfavorable to erosion by mechanical means, and favorable to that by chemical processes.

Erosion in high arid regions differs from that in regions of abundant rainfall in several ways. It is obvious that the valleys will develop more slowly in the former, that they will remain young longer, that the period necessary for the dissection of the surface is greater, that the watercourses will be less numerous, and that fewer of them will have permanent streams. There are certain other differences which are less obvious. If the arid region be high and composed of heterogeneous strata, the topography which erosion develops is more angular (Fig. 83) than that of the humid region. This is because there is132 less rock decay, and less vegetation to hold the products of decay. The more resistant beds of rock therefore come into greater prominence, especially on slopes, where they develop cliffs (Figs. 109 and 110). These general principles find abundant illustration in the plateaus of the western part of the United States,[48] where the cliffs are by no means confined to the immediate valleys of the streams (Fig. 1, Pl. XII).

EFFECTS OF UNEQUAL HARDNESS.

In the preceding pages incidental reference has been made to the results of inequalities of rock resistance. This topic will now be considered more fully.

Rapids and falls.—Returning for a moment to the hypothetical island with which our study of erosion began, let a horizontal layer of hard rock be assumed to run through it (H, Fig. 111). As the rain 133falls on the land and runs off over it, wear will be less rapid where the hard layer comes to the surface than at the higher or lower levels. As a result, the slope will become steeper at and below the outcrop of the hard layer, and less steep immediately above it, as shown by ab in Fig. 111. Under these conditions the water passing over the hard ledge constitutes rapids. The increased erosion which accompanies the increased velocity makes the rapids more rapid. The process may continue until the water falls, rather than flows over the hard layer (cd, Fig. 111). With continued rainfall the edges of the hard layer, together with the slopes above and below, would continue to recede toward the center of the island. Under conditions of absolute homogeneity of material, save for the hard layer specified, no valley would be developed, and therefore no stream.

If the surface was so changed as to allow of the development of a valley (p. 63) the same principles would be applicable. As an active stream passes from a hard layer to one less resistant, the greater wear on the latter gives origin to rapids. At first the rapids would be slight (a, Fig. 112), but would become more considerable (b) as time and erosion go on. When the bed of the rapids becomes sufficiently steep, the rapids become falls[49] (cd). When the water falls rather than flows over the rock surface below the hard layer, erosion assumes a new phase. The hard layer is then undermined, and the undermining causes the falls to recede. This phase of erosion is sometimes called sapping.

If the hard layer which occasions a fall dips up-stream (Fig. 112), its outcrop in the stream’s bed becomes lower as the fall recedes (e). When it has become so low that the water passing over it no longer 134reacts effectively against the less resistant material beneath (f), sapping ceases, and the point of greatest erosion may be shifted from the soft material beneath the fall to the hard layer itself. The actual rate of erosion at this point may be no greater than before, though the relative rate is. Under these circumstances the vertical edge of the hard layer will presently be converted into an incline (f), and as this takes place the fall becomes rapids. The conversion of the falls into the rapids begins about the time the lower edge of the hard stratum in the channel reaches grade. By continuation of the process which transformed the falls into rapids, the rapids become less rapid, and when the upper edge of the hard layer has been brought to grade, the rapids disappear (h, Fig. 112). The history of rapids which succeed falls is the reverse of that which preceded. The later rapids are steepest at the beginning of their history, the earlier at their end. Stated in other terms, rapids are steepest when nearest falls in time. Slight differences in hardness in successive layers often occasion successive falls or rapids (Fig. 114).

If the hard layer which occasions the falls be horizontal, instead of dipping up-stream, the general result would be the same; but, other things being equal, the duration of the falls developed under these conditions would be greater, since they must recede farther before becoming rapids.

If the layers of unequal hardness in a stream’s bed be vertical and the course of the stream at right angles to the strike, rapids, and perhaps falls, will develop (Fig. 115). The chances for falls are greater, the greater the difference in hardness. Falls developed under135 these conditions, as well as the rapids preceding and following, would remain constant in position until the resistant layer was brought to grade, but they would ultimately disappear as in the preceding cases. Falls are not likely to develop where the strata of the stream’s bed dip down-stream, though they may develop even under these conditions if the gradient of the stream is greater than the dip of the strata (Fig. 116).

The inequality of resistance in the rock which occasions a fall may be original or secondary. In the case of Niagara Falls[50] (Fig. 113) relatively resistant limestone overlies relatively weak shale. At the Falls of St. Anthony (Minneapolis) limestone overlies friable sandstone. The falls of the Yellowstone and the Shoshone Falls of the Snake River (Idaho), are in igneous rock. In the former case the unequal 136resistance is occasioned by unequal decay of the rock, due perhaps to the rise of hot vapors which have decomposed the rock along the lines of their ascent; in the latter, a more resistant sort of igneous rock overlies a less resistant.

Structural features, such as jointing, sometimes give rise to falls, or determine their distinctive features (Fig. 117), even where the formations involved are of uniform hardness. A joint plane has the effect of a weak vertical or highly inclined bed. If an open joint is discovered in a stream’s bed, the water enters it. If it finds an outlet below, a channel is worn along the new line of flow, with rapids or falls where the water descends. Rock originally homogeneous may be much fractured in some parts, while it remains unbroken in others. Where a stream passes from the solid to the broken portion rapids, or even falls, may develop.

Falls may originate in still other ways. If for any reason a stream is forced out of its valley, it may in its flow find entrance to another valley, or to another part of its own valley, over a steep slope. If the structure of the slope favors, a fall may speedily develop. The Falls of St. Anthony are an example, the Mississippi having been turned out of its earlier course by deposits of glacial drift. Again, if an obstruction of any sort, such as a flow of lava, dams a stream, rapids or falls are developed where the water overflows the dam. When a main valley is notably deepened by glaciation the drainage from tributary valleys may fall into it, if the tributaries were not equally deepened. Falls which originated in this way are common in the western mountains of the United States, as well as in most mountain regions recently affected by local glaciers (Fig. 118).

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One waterfall often breeds others. Thus where a fall recedes beyond the mouth of a tributary stream, the tributary falls. The Falls of Minnehaha, on a small tributary to the Mississippi, near Minneapolis, may serve as an illustration. In such cases the falls may not develop from rapids. Once in existence, the fall of a tributary follows the same history as that of a main stream.

Streams which have falls are relatively clear.[51] If a stream favorably situated for the development of a fall carried a heavy load, deposition would take place below the rapids, and the tendency would 138be to aggrade the channel at that point and so to prevent the development of the fall. Falls occur only on streams which have relatively high gradients. This means that the streams which have falls are well above base-level, and streams well above base-level are young. Falls therefore are a mark of topographic youth.

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The fall of the Niagara[52] (Pl. IX) is one of the most remarkable known, both because of its large volume of water and its great descent, between 160 and 170 feet. The rate at which the fall is receding is a matter of interest not only in itself, but because, once determined, it may be made to serve as a unit of measurement for certain important events in geological history. It was formerly conjectured that this fall was receding at the rate of one to three feet per century, but it was not until recent years that its actual rate of recession was approximately140 fixed. By surveys executed in 1842 and 1890 it has been determined that its average rate of recession between those dates was something like 4½ feet per year, or about 150 times as great as the highest estimate stated above. It is to be noted that this is the average rate of recession, for all parts of the ledge over which the water falls are not receding at the same rate. The point of the “Horseshoe” has, during the same time, gone back at more than twice this rate.[53]

Rapids and falls sometimes occasion the development of pot-holes (Fig. 119), a peculiar rather than important erosion feature. The holes are excavated in part by the falling and eddying of silt-charged water, but chiefly by stones which the eddies move. Pot-holes which are not now in immediate association with rapids or falls often point to the former existence of rapids or falls.

Rock terraces.—The tendency to sapping shown in many waterfalls is also shown in the weathering and erosion of the sides of a valley where a hard layer outcrops above the bottom, and the profile of the side slopes of the valley simulates that of the stream; that is, the slope becomes gentle just above the hard layer, and steep, or even vertical, at and below its outcrop. This is illustrated by Fig. 120, where the hard layer through which the stream has sunk its valley stands out as a rock terrace on either side of the valley. Such terraces are not rare and are popularly believed to be old “water-lines”; that is, to represent the height at which the water once stood. In one sense this interpretation is correct, since a river has stood at all levels between that of the surface in which its valley started, and its present channel, but the shelf of hard rock does not mean that the river, after attaining its present channel, was ever so large as to fill the valley to the level of the terrace. Rock terraces may also result from changes of level.

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Narrows.—Inequalities in hardness occasion another peculiarity common to valleys. If a stream crosses vertical or highly inclined strata of unequal hardness, its valley is usually constricted at the crossing of the harder layers. If such a constriction be notable it is called a narrows, or sometimes a water-gap (Figs. 121, 159, and Fig. 2, Pl. XII). The Appalachian Mountains afford numerous examples. The constriction arises because the processes which widen the valley are less effective on the hard layer than on the less resistant ones on either hand. Though most narrows are due to the superior resistance of the rock where they occur, they are sometimes the result of other causes.

Narrows are much more conspicuous in certain stages of erosion than in others. While a valley is still so young as to be narrow at142 all points, no narrows will be conspicuous; but at a later stage in its history, when the valley is otherwise wide, narrows are more pronounced. At a still later stage, when the hard strata themselves approach base-level, the narrows again become inconspicuous.

From what has preceded it is clear that rapids or falls are likely to occur at narrows, especially in the early part of their history.

Other effects on topography.—Inequalities in the hardness of rock develop certain peculiarities of topography other than those of valleys. The less resistant portions of a land area more or less distant from streams are worn down more readily than those which are more resistant. If great areas of high land be capped with hard rock they are likely to remain as plateaus after surrounding areas of less resistance are brought low. If the hard capping affects a small area instead of a large one, the elevation is a butte, a hill, or a mountain, instead of a plateau (Fig. 110). Many buttes and small mesas are but remnants of former plateaus (Mesa Lauriano, N. M., Fig. 1, Pl. XII). A feature of buttes and mesas capped by hard rock is the steep slope or cliff corresponding to the edge of the hard bed (Figs. 78 and 109).

If the rock of a region be stratified and the layers tilted, the removal of the softer beds leaves the harder ones projecting above the general level in the form of ridges or “hog-backs” (Figs. 122 and 123).143 Dikes of igneous rock, harder than the beds which they intersect, likewise become ridges after the degradation of their surroundings. The plugs of old volcanic vents and other igneous intrusions of limited area often constitute conspicuous hills or mountains after erosion has removed their less resistant surroundings (Fig. 124). Inequalities of hardness are therefore responsible for many hills and ridges. In the isolation of the hills and ridges picturesque coves are developed, where the attitude and distribution of the weak and strong rocks are propitious. The bottoms of the coves are located on the weak rocks, and above them rise the precipitous slopes of the resistant ones. Round valley (Fig. 1, Pl. XVII, High Bridge, N. J., quadrangle, U. S. Geol. Surv.) and the coves about the head of Hiawassee River (Dahlonega, Ga., quadrangle) are examples.

Ridges and hills resulting from the unequal degradation of unequally resistant terranes are not equally prominent at all stages in an erosion cycle. In early youth the material surrounding the hard bodies of rock has not been removed; in early maturity considerable portions of their surroundings still remain about them; but in late maturity or early old age the outcropping masses of hard rock144 have been more perfectly isolated and are most conspicuous. Most of the even-crested ridges of the Appalachian system, as well as many others which might be mentioned, became ridges in this way. In the final stages of an erosion cycle the ridges of hard rock are themselves brought low. Isolated remnants of hard rock which remain distinctly145 above their surroundings in the late stages of an erosion cycle (Fig. 124) are known as Monadnocks, the name being derived from Mount Monadnock, N. H., an elevation of this sort developed in a cycle antedating the present.

146

Adjustment of streams to rock structures.—Valleys (gullies) locate themselves at the outset without immediate regard to the hardness and softness of their beds. It is primarily the slope about the head of a gully which determines its line of growth, though relative hardness often determines the details of slope, even in the early stages of an erosion cycle. Once established, streams tend to hold their courses, even if this involves the crossing of resistant layers.

While a region where more and less resistant layers of rock come to the surface is in a youthful stage of erosion, some of the valleys (and therefore the streams) are likely to be located on the less resistant rock, some on the more resistant, and some partly on the one and partly on the other. The streams on the weaker rock will deepen their valleys more rapidly than the others, and those which flow across stronger and weaker rocks alternately will deepen their valleys more rapidly than those which run on hard rock all the time. The former conclusion is self-evident. The latter appears from the fact that rapids will be likely to develop at the crossing of each hard layer, thus accelerating erosion at those points. Such a stream therefore not only has less hard rock to erode than one which flows on resistant rock all the time, but it erodes that which it does cross much faster.

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Streams which do not cross hard layers therefore have an advantage over those which do, and the tributaries to such streams, since they join deeper mains, have an advantage over the tributaries to the others. The valleys of the former may lengthen until their heads reach the latter, and capture their streams. This sequence of events is illustrated in the accompanying diagrams (Figs. 125–27). Even where several streams cross the same resistant bed, piracy is likely to take place among them, for some are sure to deepen their valleys faster than others, because of inequalities of volume, load, or hardness. This is illustrated by Figs. 128–30. An actual case is shown in Figs. 131, 132. Though piracy may take place when streams do not flow over rock of unequal hardness (p. 103), it is much more common where unequal resistance of the rock puts one stream at a disadvantage as compared with another.

The changes in the courses of streams, by means of which they come to sustain definite and stable relations to the rock structure beneath, are known as processes of adjustment.[54] Since streams and valleys adjust themselves to other conditions as well, this phase of adjustment may be called structural adjustment. Structural adjustment is not uncommon among rivers flowing over strata which are vertical or highly inclined, since in these positions the hard and soft strata are most likely to come to the surface in frequent alternation. The smaller streams suffer capture and adjustment first, since, as a rule, they have shallower valleys. It often happens that main streams, because of their deeper valleys, hold courses not in adjustment with structure (the Delaware, the Susquehanna, etc.), while tributary streams are captured, diverted, and adjusted. The capture of a tributary, however, leads both to the diminution of its main and to the increase of its captor, and the weakened stream may ultimately fall a prey to the one which is strengthened.

The processes of adjustment go on until the streams flow as much as possible on the weaker beds, and as little as possible on the stronger, when adjustment is complete. This amounts to the same thing as saying that the outcrops of the hard layers tend to become divides. In many cases an area is so situated that there is no escape for its drainage except across resistant rock. In this case its drainage is completely adjusted when as few streams as possible cross the resistant rock, and these by the shortest routes.

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Adjustment has been carried to a high degree of perfection in most parts of the Appalachian system. Here, as in all other mountains of similar structure, strata of unequal hardness were folded into ridges. In this case, the folds have been truncated by erosion, exposing the more and the less resistant beds (H and S respectively) in alternate149 belts along the flanks of the truncated folds (ab and cd, Fig. 133). The streams, especially the lesser ones, now flow along the strike of the softer beds much more commonly than elsewhere, and where they cross the hard layers it is usually at right angles to the strike. This is shown in Fig. 134, where the arrows indicate the direction of strike.150 In the history of these rivers, however, a factor is involved which has not yet been considered, and these streams will be referred to later.

As base-level is approached, the outcrops of hard rock are brought low. When they have been reduced to the level of their surroundings, the streams may flow without regard to the resistance of the rock beneath, for downward cutting has ceased. As this stage of erosion is approached, a readjustment of the drainage may take place, and the waters which had taken long and circuitous courses to avoid hard rock, may change their courses to more direct ones (compare Figs. 130 and 135). Adjustment is, therefore, a relative term, and streams which are adjusted at one stage of erosion, are not necessarily adjusted at another.

It sometimes happens that rocks of unequal resistance are covered by beds of uniform hardness. A consequent stream developed on the latter may find itself out of structural adjustment when it has cut its channel down to the level of the heterogeneous beds below. Such a stream is said to be superimposed (Fig. 136) on the underlying structure. Structural adjustment is likely to follow.

INFLUENCE OF JOINTS AND FOLDS.

Joints.—Various structural features of rock other than hardness influence its erosion. Apart from the stratification planes, most rock formations are affected by joints or fissures. The joints are often, but not always, nearly vertical. Two sets are generally present, and sometimes more. If but two, they usually meet at a large angle; if more than two, two are likely to be nearly perpendicular to each other, while the third and fourth sets have such directions as to cut the others at large angles. These joints allow the ingress of water, roots, etc., which help to weather and disrupt rocks. Occasionally there is notable sag of the beds of rock along joint planes, but this effect is usually superficial only (Fig. 137). Where the jointage planes are frequent and open, the columns bounded by them sometimes topple over on cliff faces, either by undercutting, or by the wedge-work of roots or ice.

The effect of joints on erosion may often be seen along a stream which flows in a rock gorge. In such situations, the outlines of the banks are sometimes angular, and sometimes crenate (Fig. 138), the151 reëntrants being located at the joints. By working into and widening joints, running water sometimes isolates masses of rock as islands (Fig. 139). In a region free from mantle rock, or where the mantle rock is meagre, joints often determine the courses of valleys by directing the course of surface drainage. This is shown in many parts of the arid west. In regions where the rocks are notably faulted, the courses of the streams are sometimes controlled by the courses of the fault planes. This is the case, for example, in central Washington.[55]

The jointing of rocks often shows itself distinctly in the weathered faces of cliffs (Figs. 140 and 141), especially in arid and semi-arid regions, or where the slope is too steep for the accumulation of soil and rock-waste on its surface.

If a stream flowing over jointed rock has falls, the conditions are sometimes afforded for the development of an exceptional and striking scenic feature. If above Niagara Falls, for example, there were 153an open joint in the bed of the stream (as at b, Fig. 142), some portion of the water would descend through it. After reaching a lower level it might find or make a passage through the rock to the river below the falls. If even a little water took such a course, the flow would enlarge its channel, making a passageway between the joint through which the water descended and the valley below the falls (bcde, Fig. 142). This passageway might become large enough to accommodate all the water of the river. In this case, the entire fall would be transferred from the position which it previously occupied (f) to the position of the enlarged joint (b). The fall would then recede. The underground channel between the old falls and the new would be bridged by rock (bf″ and f‴, Fig. 143), making a natural bridge. The natural bridge near Lexington, Va. (Fig. 144), almost 200 feet above the stream which flows beneath it, is believed to have been developed154 in this way. A similar bridge is now in process of development in Two Medicine River in northwestern Montana (Fig. 145). Once in existence, a natural bridge will slowly weather away.

It is not to be understood that all natural bridges have had this history. They are sometimes developed from underground caves when parts of their roofs are destroyed, as well as in various other ways.

Folds.—The erosion of folded strata (anticlines and synclines) leads to the development of distinctive topographic features. So soon as a fold begins to be lifted, it is, by reason of its position, subject to more rapid erosion than its surroundings. For the same reason the crest of the fold is likely to be degraded more rapidly than its lower slopes, and must suffer more degradation before it is brought to base-level. Folds are usually composed of beds of unequal resistance, and as the degradation of a fold proceeds, successive layers are worn from the top, and the alternating hard and soft layers composing it are exposed. So soon as this is accomplished, adjustment of the streams155 is likely to begin, and the watercourses, and later the valley plains, come to be located on the outcrops of the less resistant layers, while the outcrops of the harder beds become ridges.

If the axis of an eroded anticline were horizontal, a given hard layer, the arch of which has been cut off, would, after erosion, outcrop on both sides of the axis. When the topography was mature these outcrops would constitute parallel ridges, or parallel lines of hills; when the region had been base-leveled, the outcrops would be in parallel belts, though no longer ridges or hills. The lower the plane of truncation, the farther apart would the outcrops be in the anticline, and the nearer together in the syncline (compare ab and cd, Fig. 133).

If, on the other hand, the axis of the anticline or syncline to be eroded was not horizontal, that is, if it plunged, the topographic result would be somewhat different. Suppose a plunging anticline to be truncated at base-level. If either end of the fold plunged below the plane of truncation, the outcrops of a given layer on opposite sides of the axis would converge in the direction of plunge, and come together at the end. At a stage of erosion antedating planation (say late maturity) there would have been a ridge, or a succession of hills, in the position corresponding to the outcrop of a hard layer, with a canoe-shaped valley within. If two hard layers were involved, instead of one, there would be two encircling ridges, with a curved valley between them, and a canoe-shaped valley within the innermost (Fig. 146). If157 the anticline plunged both ways, the valley enclosed by the hard-layer ridge would be canoe-shaped at both ends (Fig. 147). In such a case there would be likely to be a low gap (water-gap) in the rim of the valley through which the drainage which degraded the surface escaped, but there would be likely to be but one, for if two or more streams had drained the area of the valley at an early stage of erosion, one would be likely to have captured the others (see p. 138) before late maturity. A succession of doubly-plunging anticlines and synclines might give rise to a very complex series of ridges and valleys. Illustrations of the above phenomena are found at various points in the Appalachian Mountains, especially in eastern Pennsylvania.[56]

158

In the structural adjustment which goes with the erosion of folds, it often happens that the valleys come to be located on the anticlines, while the outcrops of the hard layers on the flanks of the anticlines, or even in the original synclines, become the mountains. The adjustments159 by which valleys come to be located on anticlines are somewhat as follows:[57] Fig. 148 represents two doubly-plunging anticlines with a syncline between, the relative elevations being shown by contour lines. At the outset, the drainage of such a region must have followed the structural valley, and its initial course, consequent on the slope, must have been down the axial trough. Drainage from the anticlines into the synclines would have promptly developed valleys, and the valleys would soon have acquired streams.

The anticlines and synclines under consideration are assumed to have a thick hard layer at the surface, and softer beds below. This is shown in the cross-section introduced in the figure, the upper hard stratum (m) being indicated by the dots, while the softer one (n) is white. The line oo represents base-level, which is below the hard layer both in the syncline and anticline, but much farther below in the latter position than in the former. Because of their higher gradients, and because of the greater fracturing to which the region they drain 160was presumably subject at the time of folding, the tributary streams might cut through the hard layer sooner than the main stream which they join. This done, they would enlarge their valleys rapidly in the softer rock beneath, and secondary tributaries would be developed (Fig. 149). When the condition of things represented in Fig. 149 is reached, the streams c and d, tributary to the synclinal stream, come into competition. The former has the advantage over the latter, because it joins the main stream at a lower level. Stream c will therefore be likely to capture d. The incipient stages of the capture are stealthy, and the later bold. At first the divide between their head waters is shifted northward inch by inch, because the gradient toward g is higher than that toward e. The capture of the head waters of e is as slow as the migration of the divide, until the divide reaches the point where e joins f. The stream f is then diverted promptly into the valley of g, and is at once led away to c (see Fig. 150). Strengthened by its increased volume, the stream c (Fig. 150) lowers its valley across the hard layer more rapidly than before, and so holds the advantage it has gained. Not only this, but the beheaded stream d (Fig. 150), because of its diminished volume, sinks its valley into the hard layer less rapidly than before, and its decrease in power also works to the advantage of the stream leading to c. The result is that the divide between fg and d does not remain constant, but is driven back step by step toward a.

Similarly a tributary to the main stream at b (Fig. 150), may by means of its tributary h, capture the waters of fg, and lead them to the synclinal valley at b (compare Figs. 150 and 151). Deprived of its main source of supply (at c) the synclinal stream is greatly diminished above b, and cuts more and more slowly, while the stream fgh (Fig. 151), having greater volume and working mainly in softer rock, sinks its channel faster than the stream in the synclinal axis. Under these circumstances, the stream at f may cut its valley below the valley in the synclinal axis a (Fig. 150). In this event, the divide between f and a (Fig. 150) may be pushed back until the synclinal stream is beheaded at a and carried out of the syncline and over into the anticlinal valley (Fig. 151). Thus, the old anticlinal axis comes to be the course of the main stream. Similarly the stream entering the syncline at b (Fig. 151) might later be captured by i, thus lengthening its anticlinal course.

161

It is not to be understood that this sequence of events will take place in the degradation of every anticline, but the principles here set forth will always be operative. The result specified will be accomplished wherever hard and soft layers have the relations indicated in the diagrams; that is, where the stream in the syncline finds itself on a resistant layer as it approaches base-level, while at the same time the (original) tributary streams are working in softer beds. It is not to be understood, therefore, that streams migrate from synclines to anticlines for the sake of getting out of the former positions into the latter. If they shift their courses it is to find easier ones.

That these changes are not fanciful is shown by the fact that the adjustment described corresponds with that shown in many parts of the Appalachian Mountains, and in other mountains of similar structure.

If in a later stage of its history, the new main stream, fh, were to cut its bed down to a lower hard layer, while the original stream, ab, reached a softer bed beneath the hard one above, the latter would again have an advantage, and a new series of adjustments would be inaugurated which might result in re-establishing the main stream in its original synclinal position.

EFFECT OF CHANGES OF LEVEL.

Rise.—If after being base-leveled, or notably reduced by erosion, a region is uplifted so as to increase the gradients and therefore the velocities of the streams which drain it, the streams are said to be rejuvenated, and a new cycle of erosion is begun. If the rise of the area were equal everywhere, while the coast line remained constant in position, there would be an immediate increase in velocity only at the debouchures of the streams flowing directly into the sea. At the debouchures of such streams there would be rapids or falls. Each162 rapids or falls would promptly recede, and with the recession, the acceleration of velocity resulting from the uplift would be felt farther and farther up-stream, and ultimately to its source. The rejuvenated streams would cut new valleys in the bottoms of their old ones (Figs. 152 and 153). The new valleys would begin where the increase in velocity was first felt, and they would be lengthened by head erosion just as valleys of the first cycle were lengthened.

When the head of the new part of a valley of a rejuvenated stream recedes past the mouth of a tributary adjusted[58] to the gradient of the main stream before rejuvenation, the velocity of the tributary is 163accelerated at its debouchure, and it begins to excavate a new valley in the bottom of its old one. The new valley commences at the lower end of the old one, and develops headward (a and b, Fig. 153). Good illustrations are furnished by the streams in the west central part of New Jersey. The Delaware has here a sharply defined valley, and its tributaries are essentially as deep as their main at the point of junction. Above this point they have high gradients for a short distance (three to six miles), beyond which they wind sluggishly in wide valleys with low gradients across a relatively high plateau. Their profiles are illustrated by Fig. 154. The flat, though high, surface in which their upper courses lie, appears to have been nearly base-leveled in an earlier cycle, and then to have been elevated. The date of the elevation is fixed, in terms of erosion, by the time necessary for the excavation of the Delaware gorge, and the narrow gorges along the lower courses of its tributaries. It was so recent that the effects of rejuvenation, proceeding from the debouchures of the tributaries toward their heads, have not yet advanced far from the Delaware. Similar relations are found elsewhere (Fig. 1, Pl. XIII, s. c. Col.). Another peculiarity of rejuvenated drainage is shown in Fig. 2, Plate XIII (s. Kan.). Here Elm Creek flows at a level 200 feet below that of Sand Creek, 4 miles distant. The valley of the former appears to have entered upon a new cycle as the result of uplift, while that of the latter, in the area shown on the map, is still unrejuvenated. Farther down-stream, the valley of Sand Creek shows signs of rejuvenation. It may be noted that a tributary of Amber Creek has good opportunity to capture Sand Creek, for the latter flows about 25 miles before reaching the level of Amber Creek at its junction with Elm Creek.

PLATE XIII.

U. S. Geol. Surv.

Scale, 2+ mile per inch.

U. S. Geol. Surv.

Scale, 2+ mile per inch.

PLATE XIV.

U. S. Geol. Surv.

Scale, 1+ mile per inch.

U. S. Geol. Surv.

Scale, 1+ mile per inch.

Should the lower end of a tributary valley fail to be degraded as fast as the valley of the main at the point of junction, the tributary is out of topographic adjustment with its main. Falls or rapids may result. When the lower end of a tributary valley is distinctly164 above the level of its main, the former is called a hanging valley. Hanging valleys developed by stream erosion alone are not common except just after the recession of a falls past the mouth of a tributary. Hanging valleys, as well as the characters and relations illustrated by Figs. 152–154 are criteria of rejuvenation, but they must be applied with discretion. Such profiles, for example, as that shown in Fig. 154 may be developed when the rock of a stream’s bed is unequally resistant, and hanging valleys are generally a result of glaciation (see Chapter V).

Rejuvenated streams sometimes inherit certain peculiarities from their aged ancestors. Thus a rejuvenated stream may intrench the meanders possessed by the old stream which preceded (see Fig. 1, Pl. XIV, near Harrisburg, Pa.), and intrenched meanders are one of the marks of rejuvenated streams. They are not uncommon in the Appalachian Mountain regions, and are known in other parts of the world. The Seine and the Moselle furnish further illustrations.[59]

The history of the new cycle of erosion inaugurated by the uplift would differ from that of the preceding cycle in that the new one would begin with a drainage system already developed. Other things being equal, therefore, the reduction of the land would proceed more rapidly in a subsequent cycle than in the first.

The recognition of different cycles of erosion, separated by uplifts, is often easy. The principles involved are illustrated by Fig. 155 which represents an ideal profile of considerable length (say 50 miles). The points a, a′, and a″ reach a common level. Below them there are areas b, b′, and b″ which have a nearly common elevation, below which are the sharp valleys d, d′, and d″. The points a, a′, and a″ represent the cross-sections of ridges formed by the outcrops of layers of hard rock. If the crests of the ridges are level, the points a, a′, and a″ must represent remnants of an old base-level, since at no time after a ridge of hard rock becomes deeply notched does it acquire an even crest, until165 it is base-leveled.[60] At all earlier stages its crest is uneven. After the cycle represented by the remnants a, a′, and a″ was completed, the region suffered uplift. A new cycle represented by the plain b, b′, and b″ was well advanced, though not completed, when the region was again elevated, and the rejuvenated streams began to cut their valleys d, d′, and d″ in the plain of the previous incomplete cycle. The elevations, c and c′ (intermediate in elevation between a, a′, and a″, and b, b′, and b″) may represent either remnants of the first base-level plain which were lowered, but not obliterated, while the plane b, b′, b″ was developing; or they may represent a cycle intermediate between that during which a, a′, a″ and b, b′, b″ were developed. If the intermediate elevations (c, c′) have a common height and level crests, the presumption would be in favor of the latter interpretation. If they be numerous and of varying heights, as is possible, they may in the field obscure the planes (a, a′, a″ and b, b′, b″) developed in the different cycles, which, in the figure, are distinct.

If the strata involved be horizontal the determination of cycles is sometimes less easy. Thus in Fig. 156, it is not possible to say whether a and a′ represent remnants of an old base-level, or whether they represent the original surface from which degradation started. So, too, the various benches below a, such as b, b′, and b″ may readily be the result of the superior hardness of the beds at this level. For the determination of successive uplifts in the field it is necessary to consider areas of considerable size, and to eliminate the topographic effects of inequalities of hardness, and of certain other factors to be mentioned presently.

The inequalities in the depths of the young valleys in Figs. 155 and 156 may be explained on the supposition that the deeper ones belong to main streams, and the shallower ones to tributaries. Such a valley as that shown at e, Fig. 155, suggests rejuvenation at this point; but farther up the stream which occupies this valley, rejuvenation 166might not be apparent. In this case, the main streams might be flowing in new valleys, d, d′, etc., while the heads of their tributaries are still flowing in the older valleys of the preceding cycle (compare Fig. 154 and Fig. 1, Pl. XIII).

It is by the application of the preceding principles that it is known that the Appalachian Mountains, after being folded, were reduced to a peneplain (p. 76), throughout their whole extent from the Hudson River to Alabama. The peneplain level is indicated by the level crests of the Appalachian ridges, shown in cross profile by the high points of Fig. 157. The system was then uplifted, and in the cycle of erosion which followed, broad plains were developed at a new and lower level, corresponding in a general way to the plains b, b′, and b″ of Fig. 155. The plains were located, for the most part, where the less resistant strata come to the surface. Above them rose even-crested ridges, the outcrops of the resistant layers, which had been isolated by the degradation of the softer beds between. They constitute the present mountain ridges (the high points of Fig. 157). The evenness of their crests, testifying to the completeness of the first peneplanation, is shown in Fig. 158, which represents, diagrammatically, a longitudinal profile of an Appalachian Mountain ridge. The evenness of the crest is interrupted by (1) notches (b, c, etc., Fig. 158) cut by the streams in later cycles, and (2) by occasional elevations above the common level (monadnocks, a, a′, Fig. 158). The monadnocks are generally rather inconspicuous, but there is a notable group of them in North Carolina and Tennessee. Mount Mitchell and Roane Mountain are examples. When long distances are considered, the ridge crests depart somewhat from horizontality. This is believed to be due, in part at least, to deformations of the old peneplain during the uplift which inaugurated the second cycle of erosion.

167

The extent to which the second cycle of erosion recorded in the present topography had proceeded before its interruption by uplift, is indicated by the extent of the valley plains (Fig. 157) below the mountain ridges. While these plains were being developed on the weak rocks, narrow valleys only (Fig. 158) were cut in the resistant rocks which now stood out as ridges. In Fig. 158 some of these valleys are shallow (c, c′, c″, etc.), and but one of them deep. The former may be either (1) the valleys of streams which crossed the hard layer at the beginning of the cycle, and which were diverted before their valleys became deep; or (2) they may represent the heads of valleys now working back into the ridges. The deep valley (b) represents the work of a stream which has held its course across the hard layer while the latter was being isolated as a mountain ridge (compare Figs. 131 and 132). Deep narrows of this sort are often called water-gaps. Similar valleys, whether shallow or deep, from which drainage has been diverted, are sometimes called wind-gaps. The second cycle of erosion, while still far from complete, was interrupted by uplift (relative or absolute), and a new cycle inaugurated. This event was so recent that the new (third) cycle has not yet advanced far.

Recently it has been urged that another cycle, intermediate between the first and second, is to be recognized.[61]

Some of the features just described are illustrated by Fig. 159. The even mountain crest in the background is the Kittatinny Mountain of New Jersey and its continuation in Pennsylvania. In common 168with other corresponding crests it represents the oldest recorded base-level (or peneplain) of the region. The great gap in the mountain is the Delaware Water-Gap. Below the mountain crest there is another plain, developed in a subsequent cycle of erosion, while the valley plain in the foreground represents the work of a still later cycle.

The oldest erosion plain of the Appalachian Mountains, the results of which are seen in the even-crested ridges so characteristic of the system, is sometimes called the Kittatinny base-level.[62] It was completed169 early in the Cretaceous period, and hence is sometimes known as the Cretaceous base-level. The next lower plain, imperfectly developed, has been called the Shenandoah Plain,[62a] from the Shenandoah Valley where it is well seen (Fig. 132 and Fig. 2, Pl. XII). It is to be noted that the terms base-level and peneplain have both been used in connection with these old plains. Graded plain is equally applicable. The truth is that the topographic types represented by these three terms grade into one another. It may be questioned whether definitions should be insisted on which differentiate these types more sharply than Nature has.

Many of the peculiarities of the drainage of the Appalachian Mountain system are intimately connected with the history just outlined. Thus three great rivers, the Delaware, the Susquehanna, and the Potomac, have their sources west of the Appalachians proper, cross the system in apparent disregard of the structure, and flow into the Atlantic. The James and Roanoke head far to the west, although not beyond the mountain system, and flow eastward, while the New River (leading to the Kanawha) farther south, heads east of the mountain-folds, and flows northwestward across the alternating hard and soft beds of the whole Appalachian system, to the Ohio (Fig. 160). The French Broad, a tributary to the Tennessee, has a similar course. Such streams are clearly not in structural adjustment, and afford good opportunities for piracy. Their courses were apparently assumed during the time of the Kittatinny base-level, when the streams had so low a gradient as not to be affected by the structure (p. 150). Elevation rejuvenated them, and they have held their courses in succeeding cycles across beds of unequal resistance, though smaller streams have become somewhat thoroughly adjusted. Crustal deformations have also helped them to hold their courses, for the Cretaceous peneplain seems to have been tilted to the southeast at its northern end, and to the southwest at its southern, when the succeeding cycle began.

Streams which hold their early courses in spite of changes which have taken place since their courses were assumed are said to be antecedent. They antedate the crustal movements which, but for pre-existent streams, would have given origin to a different arrangement of river courses. As a result of crustal movements, therefore, a consequent stream may become antecedent. Master streams are more 170likely to hold their courses, and therefore to become antecedent, than subordinate ones.

The uplift of base-leveled beds, especially if the beds are tilted so as to bring layers of unequal resistance to the surface at frequent intervals, affords conditions favorable for extensive adjustment. The numerous wind-gaps in the mountain ridges, representing the abandoned courses of minor streams, and the less numerous water-gaps, which indicate the resistance of large streams to structural adjustment, are instructive witnesses of the extent to which adjustment has gone. So extensive has been the adjustment among the streams of the Appalachian Mountains that there is probably no considerable stream in the whole system which has not gained or lost through its own or its neighbors’ piracy. The history of the rivers of the Appalachian Mountains has been further complicated by a considerable amount of warping during the periods of uplift.[63]

Sinking.—The land on which a river system is developed may be depressed relative to sea-level. In this case the sea would occupy the lower ends of valleys, converting them into bays and estuaries. A stream in this condition is said to be drowned. Of drowned rivers there are many examples along the Atlantic coast. Thus the St. 171Lawrence River is drowned up to Montreal, and the Hudson up to Albany. If the drowned portion of the latter valley were not so narrow, it would be a bay. Delaware and Chesapeake Bays, as well as many smaller ones, both north and south, are likewise the drowned ends of river valleys (see figures, Chapter VI). If all parts of a drainage basin sank equally, the velocities of the streams above the limit of drowning would not be changed, for the gradients would remain the same as before. The fact that a river’s channel is below sea-level is not to be taken as proof that the valley is drowned. Thus the bottom of the channel of the Mississippi is as much as 100 feet below the level of the Gulf, some 20 miles above New Orleans.[64]

Differential movement. Warping.—Where a land surface on which a river system is established suffers warping, some parts going up and others down, the opposite movements being either absolute or relative, various phenomena would result. This may be illustrated by the accompanying diagrams (Figs. 161 and 162), where the profiles of the streams are represented as warped from the positions represented by the dotted lines, to the positions shown by the full lines. The velocity will be accelerated below the points of differential elevation (between a and b, Fig. 161, and between a and b, and c and d, Fig. 162), but checked above (above a, and between b and c, Fig. 162). Above an elevation which notably checks its flow, a stream is ponded. If the ponding is slight, a marsh may develop above the obstruction; if more considerable, a lake is formed. Lakes of this class are likely to be short-lived, since the ponded waters are likely to soon overflow and lower their outlet so as to drain the lake. The elevation which ponds the stream may be great enough and rapid enough so that the resulting lake finds an outlet by some course other than that originally followed by the stream. Where a stream holds its course across an uplift athwart its valley, either with or without ponding, it becomes an antecedent stream (see p. 169), since it has a course assumed before the latest deformation of the crust and in apparent disregard of present surface configuration. Thus the Columbia River holds 172its antecedent course across areas which have been uplifted (differentially) hundreds and even thousands of feet.[65] Some of the striking scenic features of this noble valley are the result of these changes in the country through which it flows. A lesser stream would have been 173diverted, as many of its tributaries have been. Even its course across the Cascade ranges is believed to be antecedent.[66]

Another peculiarity of valleys and streams resulting from changes of level is illustrated in Fig. 2, Pl. XIV (southern California). The main valleys of this part of the coast were developed when the land stood considerably higher than now. Later the subsidence of the coast converted the lower ends of the valleys into bays or fiords. The bays were then transformed into lagoons by deposition. Subsequent rise of the land or depression of the sea allowed the drainage from the old lagoons to cut across the deposits which had converted the bays into lagoons. The result is an old, wide valley above, suggested by a young one below.

If the warpings were considerable, much more decisive changes in drainage would result. Suppose the drainage of a given region to be represented by the streams in Fig. 163. If there is uplift along the axis 1–2, that part of ac above the axis of uplift would be ponded, or at least have its velocity checked, while the flow of some of the tributaries of d would be accelerated, and might work back and capture the other stream (Fig. 164).

Crustal warping was one of the conditions under which the Tennessee achieved its present anomalous course, and its history[67] is illustrative of the complex changes which drainage suffers when warping affects the area where the rock structures are of unequal resistance. At the close of the Cretaceous cycle of erosion, when the Appalachian Mountains had been reduced to a peneplain, the waters falling in the area now drained by the upper course of the Tennessee flowed south-south-west to the Gulf in a stream (the Appalachian River, a, Fig. 165) the lower part of which had the general position of the Coosa and the Alabama.

To the west of the Appalachian River, shorter streams flowed west and southwest into the Mississippi embayment (Fig. 165) by courses which are not now definitely known. The succeeding cycle of erosion was inaugurated by uplift and deformation of the peneplain. The axis 175of greatest elevation (AB, Fig. 166) was nearly parallel to the Appalachian River, and the effect of the differential uplift was to impose a greater task on this river (a, Fig. 166), which flowed along the axis of uplift, than upon the rivers which flowed westward and southwestward to the Mississippi embayment. The result was that the strongest of the176 southwesterly flowing streams worked its head back into the drainage basin of the Appalachian River, and captured, one by one, the head-waters of its westerly tributaries, establishing some such drainage relations as are shown in Fig. 166. Still later, after the land area of 177the region had been considerably extended by the withdrawal of the sea, the Appalachian River itself was reached by the invading stream, and its waters carried away to the Mississippi Bay by a course the lower part of which is thought to have corresponded approximately with the course of the present Black River (b, Fig. 167).

Still later there was further deformation which caused additional changes in the drainage. The whole region was uplifted, relatively if not absolutely, but the uplift was differential, being greatest along the axis represented by AB, Fig. 167. The effect of the deformation was to stimulate the tributaries of the Ohio flowing north from this axis. Their growth was further accelerated by the weakness of the strata over which they ran. At the same time, the uplift to the south led the southwesterly flowing stream (b, Fig. 167) to discover relatively hard beds of rock in its lower course, and these beds retarded its down-cutting. The result was that a tributary of the Ohio (a, Fig. 167) finally tapped the main stream flowing to the southwest (b, Fig. 167) and carried its upper part over to the Ohio (Fig. 168). This was the beginning of the present Tennessee.

THE AGGRADATIONAL WORK OF RUNNING WATER.

Principles involved.—Since deposition results from the failure of transportation, the factors which control transportation also influence deposition. Transportation by streams is determined largely by velocity, and the most important factors influencing velocity are slope, volume, and load (p. 115). Of these the first two are usually of greater importance than the third.

A stream is said to be loaded when it has all the sediment it can carry; it is loaded with fine material when it has all the fine material it can carry, and with coarse material when it has all the coarse it can transport. A stream loaded with coarse material flows more swiftly than one loaded with fine, for a larger percentage of a stream’s energy can be utilized in carrying fine material than coarse, and hence a larger percentage of the energy of a stream which carries a load of the latter will express itself in velocity.

Deposition takes place whenever a stream finds itself with more load than it can carry, and is an expression of the stream’s refusal178 to remain overloaded. A stream may become overloaded in various ways. It might at first seem unnecessary to inquire whether a stream may be overloaded at its source, but the question is not necessarily to be answered in the negative. The source of a stream is not always a definite point. In a general way it may be said that the source of the normal stream is at that point in its valley where the bottom is as low as the ground-water level of the region. But since the ground-water level is not constant (p. 71) the source of a stream is likely to be farther up its valley in a wet season than in a dry one (p. 72). After a heavy shower, the run-off descends to the axis of the valley from the slopes on all sides, and temporarily the stream begins above the point which marks even its wet-season source. If under such circumstances the slopes about the head of the valley are notably steeper than the slope of the valley itself, as they frequently are, the water flowing down them may gather an amount of material which it cannot carry after it reaches the bottom of the valley. This may be the case at, or even above, the point which marks the source of the permanent stream. It is, therefore, possible for a stream to be overloaded at its source, if we take the source to be the point whence the water permanently flows. Deposition may, therefore, be taking place in a valley at the head of its permanent stream, or temporarily even in the valley above it.

Streams issuing from glaciers sometimes have more load than they can carry after they escape from the ice. If the stream be regarded as beginning at the point where it issues from beneath the ice it may be overloaded at its source.[69]

Under certain circumstances, a stream may overload itself. Thus if a stream loaded with coarse detritus reaches a portion of its valley where fine material is accessible in abundance, some of the velocity which is helping to carry the coarse may be used in picking up and carrying the fine. This reduces the velocity, and since the stream already had all the coarse material it could carry, reduction of velocity must result in deposition. It follows that when a stream fully loaded 179with coarse material picks up fine, it becomes overloaded, so far as the coarse material is concerned.

Again, tributaries may overload their mains. While tributaries are usually smaller than their mains, they frequently have higher gradients, and the smaller stream of higher gradient may bring to the larger stream of lower gradient more material than the latter can carry away. Thus deposition may take place at the point of junction of tributaries with their mains. This may go so far as to pond the latter enough to cause its expansion into a river-lake. Lake Pepin, in the Mississippi River at the mouth of the Chippewa (in Wis.), is an example.

Streams may become overloaded by losing velocity or volume, or both. Decrease in velocity is brought about either by decrease in declivity or in volume. In general, streams have lower gradients and greater volumes in their lower courses than in their upper, and these two elements affect velocity in different ways. If the increase in volume be not enough to counterbalance the decrease in declivity, as is often the case, a stream which is loaded in its upper course will deposit in its lower. The decrease of velocity at the debouchure of a stream almost always leads to deposition.

Decrease in velocity as the result of decrease in volume is less common. When decrease in volume occurs, it may be the result of (1) evaporation, (2) the absorption of water into the bed of the stream, or (3) branching—the giving off of distributaries. While evaporation is going on everywhere, the diminution of a stream by this means is usually more than balanced by the increase from tributaries, rainfall, and springs; but in arid regions a very different condition of things sometimes exists. If mountains in an arid region be capped with snow, its melting supplies the streams during the melting season. As the streams flow out from the mountains through dry regions, they receive little or no increment from rainfall, tributaries, or springs, and evaporation reduces the volume of water, or even dissipates it altogether. Absorption of water into the bed of the stream often accompanies evaporation. Reduction of volume by evaporation and by absorption is especially common in arid regions. Wherever loaded streams are reduced in volume, whether by evaporation or absorption, deposition takes place.

180

The third way by which velocity is decreased as the result of decreasing volume is illustrated at the debouchures of many streams. Near the Gulf, for example, the Mississippi branches repeatedly (see Fig. 190). The same phenomena are often seen where one stream joins another (Fig. 169). Individually the distributaries are much smaller than the main stream before they separated from it, and because they are smaller their combined surfaces are greater, and the amount of energy consumed in the friction of flow is increased. The velocity of the water and its carrying power are, therefore, reduced. Thus the branching of streams gives rise to deposition, and where deposition takes place the gradient of the stream is reduced, and this occasions still further deposition. The sediment which fills up the channel and checks the flow finally compels the stream, or some part of it, to transgress its banks. Deposition, therefore, favors the development of distributaries, and the development of distributaries in turn favors deposition.

The foregoing statements make it clear that a stream may be eroding in one part of its valley while it is depositing in another, and that erosion may alternate with deposition in the same place, on account of fluctuations in volume, and, therefore, in velocity of the stream. It will be seen in the sequel that erosion and deposition may be taking place at the same time in the same part of the valley. The activities181 of a river are so nicely balanced that slight disturbance at one point causes disturbance at all points below.

The deposits.

Types.—Turning from the principles which underlie river deposition to the deposits themselves, they are found to occur in various situations. Running water usually descends from steeper slopes above to gentler slopes below, and ends at the sea, or in a lake or inland basin. Wherever there is a sudden decrease in its gradient, as at the base of a hill, ridge, or mountain, running water is likely to leave a large part of its load, building an alluvial fan or cone (Figs. 67, 68, and Pl. VI). Even where there is no sudden decrease in the gradient of a stream, there is likely to be a gradual one, and in spite of the fact that the increased volume of a stream in its lower course tends to overcome the effect of diminished gradient on velocity, deposition is likely to take place as the gradient is reduced. Deposits occasioned by the gradual reduction of a stream’s velocity often have great extent in the direction of a stream’s flow. They cover the flood plains of streams, making them alluvial plains (Fig. 73). When a stream reaches the sea or a lake its current is destroyed and its load dropped, unless taken in charge by the waves and currents of the standing water. Sediment accumulated in quantity at the debouchures of streams gives rise to deltas (Figs. 169, 187). Alluvial cones and fans, alluvial plains, and deltas, are the principal types of river deposits. Apart from these well-defined types there are bars in the channels of depositing streams, and much ill-defined alluvium which does not allow of ready classification.

Alluvial fans and cones.—The only distinction between the alluvial fan and the alluvial cone is one of slope, the cones (they are but half-cones at best) being steeper than the fans. Alluvial fans and cones have their most striking development where temporary torrents, occasioned by showers or the rapid melting of snow, issue from mountain ravines. Such streams usually carry heavy loads of detritus, the coarser part of which is likely to be deposited at the base of the mountain slope. Cones and fans built by such streams have a periodic rather than a steady growth.

At the beginning of its development the material of the alluvial cone is deposited much as in a talus cone (compare Fig. 170 with Figs. 67 and 68). Its deposition chokes the channel of the stream, and182 some of the water then seeks new courses to right and left of the apex of the deposit. This expands the area of deposition to right and left, while the water which flows over it lengthens it in the direction of flow.

The course and behavior of the water after reaching an alluvial cone is instructive. As its velocity is checked, deposition often takes place in the channel, diminishing its capacity. As the channel is filled up, the water tends to overflow on either side. The overflowing water, being shallow, has so little velocity that much of its load is dropped on either margin of the channel, building up levees. The water ever and anon breaks through the levees, giving rise to distributary streams, each of which aggrades its channel and builds its own miniature levees (Fig. 171). Not rarely this process of channel-filling and levee-building goes on until the channels of the little rivulets are above the general level of the cone on which they rest. The rivulet then runs in a groove on the crest of a little ridge. The channels on the surfaces of fans and cones are fewest and deepest at their heads, and more numerous and shallower below. In some cases the surface-water disappears altogether before the outer border of the fan is reached, by sinking into the débris.

Alluvial fans and cones have various forms, and often attain considerable dimensions. Their angles of slope depend on the amount of reduction of velocity which the depositing water suffers, and the amount and kind of load which it carries. The maximum slope of the cone is the angle at which the loose material involved will lie. The minimum slope of the fan, on the other hand, approaches horizontality.183 If many alluvial fans develop in proximity to one another, as at the base of a mountain range, they may expand laterally until they merge. A long succession of them may thus give rise to an extensive alluvial piedmont plain, or a compound alluvial fan. The lower edge of such a fan is often somewhat lobate. Such plains exist along the bases of many mountain ranges (Pl. VI), and may be seen in miniature even along low ridges.

A permanent stream, as well as a temporary one, may develop an alluvial fan at the base of a mountain slope; but since the mountain course of the former is likely to be less steep than that of the latter, its waters suffer a correspondingly less reduction of velocity at any one point. The fan of the permanent stream is therefore likely to be relatively flat, and to stretch far down the valley. Such fans grade into valley plains. From the general principles already discussed, it is clear that well-developed fans go with relatively youthful stages of erosion, and belong normally to the upper parts of drainage lines.

Ill-defined alluvium.—There is a widespread mantle of alluvial material deposited by running water which was not organized into distinct streams. The water which runs down smooth slopes in sheets during showers carries fine earthy matter, as well as some that is184 coarser. These materials are largely deposited at the bases of the slopes, forming basal accumulations of greater or less extent, comparable in origin to alluvial fans. A relatively small amount of the slope wash is carried far out from the base of the declivities. It is not easy to realize the extent to which this process is taking place. There is hardly a slope without loose material, and there is hardly an acre of low land below a slope on which running water has not deposited sediment washed down from above. When it is remembered that this is as true of gentle slopes and their surroundings as of steep slopes, though perhaps not to the same extent, and that a very large part of the earth’s surface is made up of sensible slopes, or of flats at their bases, some idea of the aggregate effect may be gained.

There is another way of looking at the same question. Earthy matter is being continually transferred from land to sea, and chiefly from high land. Rarely does it start from any point distant from the shore and move uninterruptedly to it. It is transported a short distance and lodged, to be again picked up, carried forward another step in its journey, and lodged again. For a very large part of the earth’s surface it would be true to say that its mantle rock is material in transit from higher land to the sea.

Alluvial plains.—Most streams, whether heading in mountains or not, have gentler gradients in their lower courses than in their upper, and in spite of increasing volume are usually unable to carry to their debouchures all the material gathered above. The excess of load is dropped chiefly on the flood-plains of the streams and constitutes them alluvial plains.

The making of an alluvial plain usually involves both erosion and deposition. When a stream has cut its channel to grade, downward erosion ceases, or more exactly, downward cutting is, on the average, counterbalanced by deposition. So long as a stream is cutting downward rapidly, it carries away whatever débris descends the side slopes. When it approaches grade, the débris which descends the side slopes tends to accumulate at their bases, and the V-shaped cross-section of the valley becomes U-shaped (see Fig. 172). At about the same time the stream begins to meander, for, having lost something of its former velocity, it is more easily turned from side to side. As it begins to meander, it widens the bottom of its valley. This is the initial stage in the development of the valley flat (2 and 3, Fig. 172). In its meandering185 the stream encroaches on the talus accumulations at the bases of its valley’s slopes. The side-cutting may remove all the loose débris and even undercut the bluff as at a, Fig. 173. The stream’s meanders shift their positions from time to time so that the valley flat is successively widened at different points. By lateral planation, therefore, a stream tends to develop a flat as soon as it reaches grade. This is the initial part of erosion in the making of a river flat, but a flat developed by erosion alone is not an alluvial plain.

So soon as the flat developed by a stream exceeds the width of its channel, the water (except in times of flood) does not cover it all at the same time. On any part which it temporarily abandons, some débris (alluvium) is likely to be left. This deposit of alluvium constitutes the valley flat an alluvial plain (Fig. 174). It will be seen that the valley flat is commonly an alluvial plain from the beginning.

Once the valley flat and alluvial plain are begun, their further development is easily followed. The stream in flood overflows the banks of its channel. The velocity of the overflowing water is reduced, and if it has much load a part of it will be dropped and the plain aggraded. Meantime meandering and lateral planation continue. Thus the flood-plain is widened by erosion, and aggraded by alluviation, the two processes going on simultaneously.

186

Flood-plains, chiefly the result of planation, but partly of aggradation, are a normal feature of river valleys, after a certain stage of development has been reached. This stage is that at which downward erosion becomes slight in comparison with lateral erosion. It follows that an alluvial plain normally begins its development where the valley is first brought to grade, that is, in its lower course. As the development of the valley goes on, the head of the flood-plain advances up-stream, and at the same time its older parts become wider.

Flood-plains due to alluviation only.—Exceptionally, an alluvial plain is developed by deposition only. Thus if a stream becomes overloaded while its valley is still narrow, as sometimes happens, deposition follows, and, as aggradation proceeds, the narrow valley acquires a progressively wider bottom (Fig. 175). Wide valley plains are sometimes developed in this way. Flood plains developed wholly by alluviation are sometimes formed under conditions which are independent of the stage of a valley’s development. Thus if a stream suddenly acquires an exceptional supply of detritus in its upper course, the development of an alluvial plain begins immediately below the point of overloading.

The overload might be acquired in various ways. (1) If a stream taps another (piracy) which carries a large quantity of sediment, carrying off both water and sediment to a channel with a lower gradient, deposition may take place where, under the earlier conditions, there was none. (2) Again, when a stream cuts through a barrier near its head waters, its velocity, and, therefore, its eroding power, may be so increased in its upper course that sediment enough is acquired to occasion deposition below, where none took place before. (3) In working back through formations of varying degrees of resistance, a stream’s head may presently reach a formation or a region which yields abundant sediment, even though there was no especial barrier below. (4) If an advancing glacier should reach the head waters of a stream, its discharge to the stream would greatly increase the load of the latter, and, although its volume would be augmented at the same time, deposition might result. As a matter of fact, streams carrying glacial drainage187 are usually aggrading streams. In general, anything which greatly increases the load of a stream near its head is likely to cause deposition, and so the development of a flood plain, at some point farther down the valley.

Streams which are actively aggrading their valleys are likely to anastomose (Figs. 176, 177). This results from the filling of the channels until they are too small to accommodate all the water. The latter then breaks out of the channel at few or many points. The new channels thus established suffer the same fate.

188

Flood-plains due to obstructions.—Again, any obstacle in a stream’s course is likely to cause deposition above. Thus dams built across rivers entail the deposition of sediment above. Where a stream flows over the outcropping edges of strata of different strength, the more resistant serve, in some sense, as dams. Above them the stream cuts its bed to a low gradient, and, becoming sluggish, drops more or less of the detritus brought down from above. Obstacles of any sort across a stream’s channel, therefore, favor the development of alluvial plains.

Levees.—As the stream in flood escapes its channel and overspreads its plain, its immediate banks are the site of active deposition, for it is here that the velocity of the overflowing water is first notably checked. On the banks of the channel, therefore, low alluvial ridges, called natural levees, are built up (Fig. 178, and Pl. XV). They may be narrow, or hundreds of feet in width, and are often several feet above the plains behind them, giving the latter a slope away from the channel of the stream. They are sometimes high enough to control the courses of tributary streams, as shown by numerous tributaries to the Mississippi below the Ohio. The Yazoo, for example, flows some 200 miles on the flood-plain of the Mississippi before it joins that river near Vicksburg. The levees even become divides, directing drainage away from the streams they guard (Pl. XV). Streams sometimes build levees faster than their tributaries aggrade their channels. The latter are then ponded, giving rise to lakes. The lakes on the lower courses of the tributaries to the Red River of Louisiana are examples.[70] They are sometimes built up above their natural level and kept in repair by human agency so as to confine the streams in time of flood. This is a source of danger unless they be steadily maintained, for the breaking of such levees often occasions great destruction. A case in point is the breaking of the levees of the Mississippi near New Orleans in 1890. The water broke through the levees at the Nita and Martinez crevasses (Fig. 187) and flowed eastward (from the former) with a current of 15 miles per hour, spreading destruction in its 190path. The water flowed eastward through Lakes Pontchartrain and Borgne, and entered Mobile Bay with such volume, velocity, and load of mud, as to destroy for a time the oyster and fish industries of that locality.[71]

PLATE XV.

U. S. Geol. Surv.

Scale, 1+ mile per inch.

U. S. Geol. Surv.

Scale, 2+ miles per inch.

PLATE XVI.

U. S. Geol. Surv.

Scale, 2 miles per inch.

U. S. Geol. Surv.

Scale, 2+ miles per inch.

Flood-plain meanders. Cut-and-fill.—A stream with an alluvial plain is likely to meander widely (Pl. XVI). In general terms this may be said to be the result of low velocity, which allows it to be easily turned aside. Were the course of such a stream made straight, it would soon become crooked again. The manner of change is illustrated by Figs. 180 and 181. If the banks be less resistant at some points than at others, as is always the case, the stream will cut in at those points. If the configuration of the channel is such as to direct a current against a given point, a (Fig. 180), the result is the same, even without inequality of material. Once a curve in the bank is started, it is increased by the current which is directed into it. Furthermore, as the current issues from the curve, it impinges against the opposite bank and develops a curve at that point. The water issuing from this curve develops another, and so on.

Once started, the curves or meanders tend to become more and more pronounced (compare Figs. 180 and 181). In the case represented by Fig. 1, Plate XVI (Missouri River near Brunswick, Mo.) the narrow neck of land between curves is almost cut through. When this is accomplished, the stream will abandon its wide curve. A later stage in the process is shown in Fig. 2, Plate XVI (the Osage River near Schell, Mo.).

191

The straightening of the channel is often accomplished in another way. Even before the meanders reach the stage represented by Fig. 1, Plate XVI, the position of the channel becomes unstable. In time of flood, the whole flat is covered with flowing water. The greater depth of water in the channel tends to give it a velocity greater than that of the water on the flat outside. But the distance from a to c via b (Fig. 181) is much greater than that in a direct line. It follows that the slope from a to c direct is greater than that by way of b. If the current between a and c in time of flood be strong enough to erode, it may deepen its bed, and thereby increase the volume of water following this course. The increased volume gives increased velocity, and the result may be the opening of a channel between a and c direct. The channel may be worn so deep that when the flood subsides, the stream will follow it. So long as the abandoned channel-curve remains unfilled with sediment, it is often called a cut-off. If it contains standing192 water and has the proper form, it is called an ox-bow lake (Fig. 182), or sometimes a bayou. The water-filled portions are not always bows (Fig. 183, Osage River, near Butler, Mo.). Cut-offs, with or without standing water, are of common occurrence along most rivers with wide plains. Meandering is not confined to streams which are near sea-level.193 Even small creeks at high altitudes may meander, if so situated as to have slight velocity. Trout Creek in the Yellowstone Park (Fig. 184) is an example.

There seems to be some relation between the width of the belt within which a stream meanders, and the width of the stream itself. Recently it has been estimated that the ratio between them is 18:1.[72]

During the development of the meanders it is to be noted that lateral planation on the one side of a stream is accompanied by deposition on the other. This is cut-and-fill. The sediment eroded from the curve which is concave toward the stream is shifted down-stream, while that deposited in the curve which is convex toward the stream is brought down from above. Thus even in the development of meanders, the material which is dislodged is shifted down-stream. Since the current directed against the down-stream side of a growing meander is on the average stronger than that directed against the opposite side, the meander itself has a tendency to migrate down-stream (Fig. 182).

In their evolution, the curves of a stream’s channel often reach and undermine the valley bluff (Pl. VII). Since the meanders are, on the average, shifted down-stream individually, and since meanders are frequently developed in new places, it follows that a meandering stream tends to widen its valley throughout. Widening is also effected in other ways, for a stream with a flood-plain sometimes abandons its channel altogether for miles at a stretch, and the new course chosen may be against one of the bluffs of the valley. Such changes are most likely to take place where deposition along channel and levees has 194brought the part of the flood-plain (though not necessarily the bottom of the channel) adjacent to the stream above the level of that farther from it (Fig. 178). The change is likely to be effected in time of flood.

Flood-plains often attain great size. That of the Mississippi below the Ohio (Fig. 179) has a width ranging from rather more than 20 miles at Helena (Ark.), to something like 80 miles in the latitude of Greenville (Miss.).[73] Below the Ohio its area is something like 30,000 square miles, and its entire area has been estimated at about 50,000 square miles.[74]

Theoretically, the rotation of the earth should affect the erosion of streams, increasing it on the one bank (the right in the northern hemisphere and the left in the southern) and decreasing it on the other.[75] The streams doubtless accommodate themselves to the rotation of the earth in the original development of their gradation-plains and flood-plains, and the later effects of rotation are usually inconspicuous.

Scour-and-fill.—It has already been shown that aggrading streams cut laterally at the same time that they build up their plains. It is 195now to be added that they periodically deepen their channels to a notable extent, and that the deepening of the channel takes place at the very time when the flood-plain is being aggraded. In other words, the stream in flood aggrades its plain, and degrades its channel. This follows from the fact that the current is sluggish in the former position, where the water is shallow, and rapid in the latter, where it is deep. When the flood subsides, the channel, deepened while the current was torrential, is filled again by the feebler current which follows. This alternate deepening and filling is known as scour-and-fill. It is well illustrated by the Missouri River. At Nebraska City, scour is believed to occasionally reach depths of 70 to 90 feet.[76] At Blair, about 25 miles above Omaha, the same river is believed to cut to bed-rock (about 40 feet below the bottom of the channel in low water) twice a year, that is, during floods.[77] Fig. 185 shows the changes recorded in the channel of the river at this point during the year 1883. It shows that the scour-and-fill during this year amounted to almost 40 feet. All 196streams similarly situated do a like work. The material thus eroded is shifted down-stream, some of it for short distances only, and some of it to the sea. Even an aggrading stream therefore is not without erosive activity; it is a stream whose fill exceeds its scour, not one which has ceased to erode.

Materials of the flood-plain.—As a result of its varying velocities in flood and low water, a stream may deposit coarse material at one time and fine at another. A similar sequence of deposits takes place in the flood-plain of a meandering stream, irrespective of floods. Flood-plain deposits are often therefore very heterogeneous, as shown in Fig. 186, which represents the constitution of the alluvium of the Missouri River at Omaha. The deposits of the streams range from the finest clay, through sand to gravel, and even bowlders. In general they become finer down-stream. In a given plain, they are usually coarser below and finer above.

Topography of the flood-plain.—The flood-plain is nearly, but not altogether, flat. It has a gentle slope down-stream, and often for a distance from the sides toward the center (Fig. 174). This latter slope is the result of deposition by waters descending to the plain from the sides. It is destroyed wherever a meandering stream reaches its bluffs. When levees are well developed, there is a slope from them197 toward the sides of the valley (Fig. 178), but it rarely continues to the limiting bluffs. Since a stream with a well-developed flat frequently shifts its course, old levees and abandoned channels lend variety to the topography of the flood-plain.

The topographic adjustment of tributaries.[78]—The meandering and shifting of a main stream affects its tributaries. If a main stream swings against the bluff through which a tributary enters, the latter brings its channel into topographic adjustment by lowering its end to the level of the main. If now the main stream opposite the tributary swings to the other side of its valley, the tributary must make its way across the flat with a very low gradient. Not only this, but the flat of the main valley through which the tributary must flow is likely to be aggraded by the main in time of flood. The result is that the tributary stream becomes an aggrading stream at its debouchure, and topographic adjustment is not established until it has filled up the lower end of its valley to some notable extent. The filling of the lower end of the tributary likewise affects the lower ends of its lower tributaries.

198

If the main stream again swings over to the point where the tributary issues from its valley, the tributary stream and all its affected tributaries again become eroding streams. Thus scour-and-fill are not confined to the valley of the main stream.

River-lakes.—While rivers are in general hostile to lakes, they sometimes give origin to them. Oxbow lakes (Fig. 182 and Pl. XVI), due to the cut-offs of meandering streams, have already been referred to. Lakes formed in the same way have other forms (Pl. XVI and Fig. 183). Rivers also give rise to lakes through the deposits they make. If a main stream obstructs its tributaries by deposition at their debouchures, their lower courses are ponded and converted into lakes. The lakes along the tributaries to the Red River of Louisiana have already been cited as examples. If a tributary brings more load to its main than the latter can carry away, the detritus constitutes a partial dam, ponding the river and causing it to expand into a lake above. Such is the origin of Lake Pepin already referred to. In mountain regions, the alluvial cones of tributary valleys sometimes pond their mains.

Rivers may be dammed in other ways, as by lava flows, by landslides, by glacial drift, etc. In all such cases, lakes may come into existence, but they are not due primarily to the activity of the river itself.

Deltas.[79]—Where a stream enters standing water, or a slower stream, a special form of plain, the delta, is sometimes built up (Figs. 169, 187, and 188). Deltas and alluvial fans have much in common, and their only notable differences are those imposed by the differences in the conditions of deposition. The current of the stream is checked, but not altogether stopped, at its immediate debouchure. If it carries abundant sediment, much of it will be promptly dropped where the decrease in velocity is first felt. Such flow as there is beyond the debouchure 199is not confined to a definite channel, and the deposits made are therefore spread more or less on either side of the line which represents the continuation of the stream’s course.

As the depth of the water into which the stream flows increases, the current diminishes. Out to the point where the depth of the standing water is less than the depth of the current, the latter affects the bottom, and the surface of the deposits made slopes gently seaward; but where the depth of the standing water is such that the projected stream current is ineffective at the bottom, all the load rolled along the bottom is dropped, and a depositional slope is established (Fig. 188), its upper edge being below sea-level by an amount corresponding roughly to the depth of the current which brings the detritus. The outer slope is relatively steep and well-defined where the detritus is coarse, and relatively gentle and ill-defined where it is fine. Thus the stream tends to construct a sort of platform in the water just beyond its debouchure. The successive deposits on the outer abrupt slope will dip conformably with its surface (Fig. 189). The finest sediment will be carried beyond the steep slope, and conform to the topography of the bottom beyond (c, Fig. 189).

At the beginning, the top of the delta platform is at the level of the bottom of the stream’s channel at the point of debouchure, but it is gradually aggraded as water continues to flow over it. Its landward margin is presently built up to sea-level and then above it, and as the delta grows the delta-land is extended seaward (compare Figs. 188 and 189). At the same time the channel of the stream above the original head of the delta is aggraded, for the current there is checked by the aggradation of the delta. Thus alluvial deposits continuous with the delta are extended landward.

The projection of the direction of the lower end of the stream may be said to be the axis of the extra-debouchure current. From this200 axis, where the flow is strongest, the movement diverges more or less to right and left. Since the velocity of the diverging water is reduced more rapidly than that of the water which follows the axis of flow, deposition is likely to take place faster on either side of the axis than along the axis itself. The result is that the extra-debouchure current tends to build up levee-like ridges on either side, making a sort of sluice-way for itself. This sluice-way is gradually extended seaward, and at the same time gradually filled. As its capacity is reduced, more and more of the water flows over its sides. Presently the escape of the water over the little side-levees will develop a break at some point, and a line of distinct flow then diverges from the main current. This distributary repeats the history of its main. Thus the processes of levee-building, channel-filling, and levee-breaking follow one another,201 until some such system of currents as shown in Fig. 190 is developed. The result is that a delta’s growth is not simply in the line of extension of the main stream, but in a more or less semicircular area, the center of the circle being a point slightly below the position of the debouchure of the stream when the delta began. At any stage in its development the margin of the delta is more or less crenate (Fig. 191), or characterized by delta fingers (Fig. 190), the projections corresponding to the positions of the debouchures of the latest streams flowing across it. The extreme ends of the delta lobes (Fig. 190), and of groups of the delta fingers, often have something of the shape of the Greek letter from which the name originated, but the resemblance in form between a well-developed delta and the Greek letter is not striking. Deltas are sometimes built in bays, and in such cases their forms are predetermined on all sides but one. The head of a delta is sometimes arbitrarily located at the point where the first distributaries are given off. Since this point shifts widely with time, the definition can hardly be accepted. On this basis the head of the Mississippi delta is about 200 miles above its lower end. In reality it is much farther north.

The structure of a delta, shown in Fig. 189, shows its history. At any stage in its growth the river discharges its sediment across that part of the platform already built. The sediment rolled at the bottom of the current is dumped on reaching the steep slope, and constitutes202 the inclined fore-set beds shown in Fig. 189. The material in suspension is carried farther, settles more gradually, and constitutes the bottom-set beds (c, Fig. 189). In time the bottom-set beds, originally deposited some distance beyond the debouchure, may come to be overlain by the fore-set beds, deposited at a later time. While the fore-set beds are being deposited on the steep slopes of the delta, and the bottom-set beds beyond, deposition is also taking place on the top of the delta. These top-set beds are laid down in a nearly horizontal position, and their seaward margin is gradually extended. Thus the delta comes to have the threefold structure shown in Fig. 189.

That part of the delta which is above the abrupt slope of its front corresponds in all essentials to an alluvial fan; but the delta as a whole differs from the fan in its abrupt and crenate or digitate margin.

It is to be noted that the delta is not wholly the product of a stream’s activity. The stream supplies the material, but the lake or sea renders at least passive assistance in its disposition. Not all rivers opening into the sea build deltas, and their failure is often the result of waves or shore currents which carry off the river sediment. Deltas are, however, sometimes formed in tidal seas, as at the debouchures of the Yukon; the Mackenzie, where the tidal range is three feet; the Niger, where the range is four feet; the Hoang-Ho, where the range is eight feet; and the Brahmaputra and Ganges, where the range is sixteen feet.[80] Since lakes, bays, gulfs, and inland seas have weaker waves and currents than the open sea, they are more favorable than the latter for the growth of deltas. Hence occur such deltas as those of the Mississippi, the Nile, the Po, and the Danube.

Deltas are likely to be absent, or confined to the heads of bays, on coasts which have recently sunk. Their general absence on the Atlantic coast of the United States is a case in point.

The following figures give some idea of the extent of deltas, and of their importance in land building. The Mississippi delta is advancing into the Gulf at the rate of about 100 yards per year, or a mile in 16 or 17 years. Its length is more than 200 miles, its area more than 12,000 square miles, and its depth at New Orleans has been estimated at 700[81] to 1000[82] feet. This great depth is believed to be the result of subsidence,203 and so of the superposition of one delta on another.[83] The delta of the Yukon has a sea margin of 70 miles, and extends more than 100 miles inland. The delta of the Rhône has also had a remarkable growth, considering the size and the history of the stream. Arles, near the debouchure of the stream, was 14 to 16 miles inland in the fourth century b.c., and is now 30 miles inland.[84] The Rhône has also built a great delta in Lake Geneva, and its lower delta is built of sediment gathered below the lake. The Po has built a delta 14 miles beyond Adria, the port which gave its name to the Adriatic Sea. The extension of this delta has been at the average rate of about 50 feet per year, but recently, on account of artificial embankments, the rate has been much more rapid.[85] The Ganges and Brahmaputra together have made a delta of great size. Its area is sometimes estimated to be as high as 50,000 or 60,000 square miles, and its head is more than 200 miles from the sea.[86] The head of the Nile delta is 90 miles from the sea, and it has a coastal border of 180 miles. The head of the delta of the Hoang-Ho is about 300 miles from the coast, and its seaward border has a length of about 400 miles, though with some highland interruptions.[87]

After a delta has been built into a lake, the lake may disappear, leaving the delta out of water. Such “fossil” deltas, if so recently exposed that erosion has not destroyed their distinctive features, are readily recognized by their flat tops, their abrupt and lobate fronts, and their characteristic structure. They are often a means of determining the former existence of extinct lakes,[88] or the former higher levels of lakes which still exist.[89] Elevated deltas on seashores show either a rise of the land or a depression of the sea-level.

The material which is carried along the coasts or shores from the mouths of rivers may take on various and peculiar forms, according 204to the strength, direction, and relations of waves and currents. The consideration of these forms belongs more properly to the work of the sea than to that of rivers, since rivers are not concerned in their construction except in supplying material.

Delta lakes.—Delta-building streams sometimes help to form lakes by throwing their deposits around an area which fails to be aggraded to sea-level. Lake Pontchartrain, and other lakes in the delta of the Mississippi are examples (Fig. 187).

STREAM TERRACES.

Stream terraces[90] are bench-like flats or narrow plains along the sides of valleys (Fig. 192). They are usually narrow, but sometimes have great length in the direction of the axis of the valley. They originate in various ways.

Due to inequalities of hardness.—Reference has already been made (p. 140) to the effect of hard horizontal layers in the development of terraces and terraciform projections on the sides of valleys (Fig. 120). Such terraces are the result of differential degradation, and the upper 205surface of the hard layer marks the lower limit of the terrace, which commonly has a distinct slope toward the stream. Except where interrupted by tributary valleys, such terraces are likely to be continuous in a valley so long as the structure remains the same and the stream sustains the same relation to it. Such terraces would first show themselves in the older part of the valley. The effect of inclination of the hard stratum on the development of such terraces will be readily inferred. Terraces and benches of this sort are not equally distinct at all stages of a valley’s history. For great distinctness, the hard layer should have been exposed long enough to allow the general processes of erosion to have effected considerable differential wear, but not long enough to allow the topographic effects of unequal resistance to be obliterated.

Normal flood-plain terraces.—It has been seen that deposition in a river valley stands in more or less definite relationship to the stage of its development, and that the deposition which leads to the development of an alluvial plain is likely to take place where the higher gradient of the upper course gives place to the gentler gradient of the lower. It has also been seen that as a stream’s history advances, the stretch where the gradient is high recedes up-stream, and that the point which marks the head of active deposition follows. It follows that a river flat or flood-plain normally begins in the lower part of a valley, and works progressively headward, its upper end following, at some considerable distance, the head of the valley itself.

The commoner river terraces are remnants of former flood-plains, below which the streams which made them have cut their channels. It has already been pointed out (p. 184) that processes of erosion and deposition work together in the development of flood-plains, and that some flood-plains have but little alluvium (Fig. 174), while others owe their origin wholly to stream deposits (Fig. 175). It follows that terraces developed from flood-plains may be of rock, of alluvium, or of rock covered with alluvium.

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The amount which a river channel must be deepened in order to change the remnants of its flood-plain to terraces cannot be definitely stated. When a channel is so deep that the remnants of a former flood-plain are no longer flooded, they would be called terraces, especially if a lower flood-plain has been developed. Even though not above the reach of floods, they are often called terraces if they are notably above the channel and separated from it by a lower plain. Thus the flat at b, Fig. 193, would be called a terrace, even though covered by water in exceptional floods; but the flat at c, but slightly above the channel, would hardly be called a terrace.

The question now arises why a stream, having once developed a flood-plain, should sink its channel to a lower level, leaving parts of the old flood-plain as terraces. This may be brought about by the operation of various causes.

(1) In the first place, the head of the valley-plain where the first notable deposition takes place normally advances up-stream. After the advance has been considerable, the descending stream may, on reaching the head of its valley-plain, lose so much of its load as to be able to sink its channel into the flood-plain farther down the valley (Fig. 194).

(2) Ordinarily a stream does not drop all its load at the head of its plain, but only its excess; but it will always drop coarse sediment to take fine, if fine be available. For a relatively small amount of coarse material dropped, a relatively large amount of fine may be taken up (p. 179). Other things being equal, it follows that when a stream207 drops coarse material to take fine, its channel is degraded unless there is at the same time a great reduction in the stream’s energy. Such reduction is likely to go with the decreasing declivity down-stream; but this is partly, or sometimes wholly, counterbalanced by the increasing volume of water. By the exchange of load, therefore, a stream may ultimately sink its channel below the flood-plain which the earlier and perhaps smaller stream had developed.

(3) Again, so long as a stream is actively eroding at its head, there is likely to be some aggradation below. At a later stage in the stream’s history, when active erosion at the head has ceased because of the reduction of the surface, less material will be carried from the upper part of the valley, and the stream on the flood-plain below, formerly loaded with material from up the valley, is now free to take up and carry away material temporarily left on the flood-plain. The result is a deepening of the channel.

(4) Any stream which has reached the flood-plain stage is likely to meander. After the flood-plain has become wide, the width of the belt within which the stream meanders is less than the width of its plain. In the Lower Mississippi, for example, the meander belt is often no more than a third to a tenth of the width of the flood-plain. It has already been pointed out that the meanders migrate down the valley. In so doing they depress the meander belt, the tendency being to reduce it to the level of the channel, and, therefore, below the level of the flood-plain.208 As the meander belt widens, the depression which it develops becomes more and more capacious. Presently it may attain such dimensions as to hold the water of ordinary floods. At this stage, or even before, such parts of the earlier flood-plain as remain, are terraces.

These several tendencies conspire to partially destroy river flood-plains, and to transform such parts as remain into terraces in the normal course of a river’s history. They appear first in the lower part of the valley, and migrate headward, following the course of nearly every other phase of activity in a stream’s history. The heads of the terraces follow, at a respectful distance, the head of the flood-plain, just as the head of the flood-plain follows at a distance the head of the valley. The second and subsequent flood-plains and the terraces to which they give origin follow the same course.

Terraces developed by the normal activities of a stream are always low, and it is improbable that they would ordinarily be conspicuous. The vertical distance between the first (highest) and second is greater than that between the second and third. The principles developed on page 65 et seq., in connection with the erosion of the hypothetical island, are applicable here.

Flood-plain terraces due to other causes.—Certain other causes, accidental rather than normal to a stream, result in the development of terraces from flood-plains. (1) If there be uplift in a region where the rivers have flats, the streams are rejuvenated, and the remnants of their former flood-plains become terraces. (2) If an alluvial flood-plain has been built as the result of an excessive supply of sediment (p. 186), the exhaustion or withdrawal of the excessive supply would leave the stream again relatively clear, and free to erode where it had been depositing. It would forthwith set to work to carry away the material which it had temporarily unloaded on the plain. The plains built up in many valleys in the northern part of our continent during the glacial period, when the drainage from the ice coursed through them, have subsequently been partially destroyed by erosion, and their remnants have become terraces. A notable reduction in the amount of available sediment, even when the earlier supply was not excessive, produces a similar result. (3) A notable increase in the volume of a stream, without corresponding increase in load, as when one stream captures another, may occasion the development of terraces by allowing the stream to deepen its channel. (4) Above any209 barrier which dams a stream, a flood-plain is likely to be developed. When the barrier is removed the stream will cut more or less deeply into the plain above, leaving terraces. (5) The recession of a falls through a flood-plain converts such parts of it as remain, into terraces.

In conclusion, it may be stated that many river terraces, mostly very low, are normal features of valley development, coming into existence at definite stages in a valley’s history. They are generally composed, in large part, of river alluvium. Others result from more or less accidental causes, working singly or in conjunction, and to this class belong all of the more conspicuous terraces developed from flood-plains. The structure of a terrace often affords some clue to its origin (Fig. 196).

Discontinuity of terraces.—When a stream sinks its channel into its flood-plain, it does not follow that a terrace remains on each side. Where the stream’s deepened channel is in the middle of its flood-plain, there is, temporarily, a terrace on either side; but wherever the deepened channel is at one margin of its flood-plain, a terrace remains on the other side only. Even where continuous at the outset, terraces soon become discontinuous, for all processes of subaërial erosion conspire to destroy them. A stream is likely to meander on its second and later flood-plains, as on its first and highest one. Wherever the meanders on its second flood-plain reach the borders of the first flood-plain, the terrace at that point disappears, and since the meanders are continually migrating, terraces are continually disappearing. The same would be true of the second terrace, if a second were developed. The removal of portions of a terrace by the sweep of meanders is likely to leave the remnants cuspate toward the stream.[91] Again, tributary streams, in bringing their channels into topographic adjustment with their mains, cut through the terraces of the latter. New gullies develop on the faces of the terraces and their heads work back across them, dissecting them still further. At the same time, sheet erosion and other phases of slope wash tend to drive the scarps of the terraces back 210toward the bluff beyond. By the time a second series of terraces is well developed, no more than meagre remnants of the first may remain.

From the foregoing considerations it is clear that the extent to which river terraces once developed, now remain, is dependent in part on the length of time which has elapsed since the river sank its channel below them. Other things being equal, the greater their age the more meagre their remnants.

Terraces developed from river plains formed chiefly by alluviation stand a better chance of long life than most other alluvial terraces. This is because of the configuration of the original valley, the aggradation of which gave origin to the plain. The principle involved is illustrated by Fig. 197. In developing the second flood-plain the river encounters the rock wall of the valley. This greatly retards lateral erosion, and the terrace above, defended[92] by the rock, is likely to be long-lived.

Alluvial terraces, like rock shelves, are popularly thought to mark “old levels of the river.” In one sense this is true, but not in the sense in which the expression is commonly used. Every level, from the crest of the bounding bluffs to the bottom of a valley, is a level at which water ran for a longer or shorter time; but the terrace does not mean that the river was once so much larger than now as to fill the valley from its present channel to the level of the terraces.

Termini of terraces.—From the mode of development of terraces it will be seen that, traced up-stream, each terrace should theoretically grade into a flood-plain at its upper end (Fig. 194), and that the upper end of the second (from the top) terrace, where there are two, would not be so far up-stream as the upper end of the first (highest). This is represented diagrammatically in Fig. 198.

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The down-stream termini of terraces are rarely distinct. This is partly because the notable meandering of the streams in their lower courses is antagonistic to the preservation of terraces. If all terraces once developed remained, and if delta-building proceeded without interruption from waves, the relations should be somewhat as follows: Traced down-stream, the cliff between the oldest (highest) terrace and the next younger becomes gradually lower until it finally disappears, and the continuation of the two is found in a common plain. The cliff between the second and third terraces should disappear in the same way, and below its disappearance the plain representing their continuation is continuous with that representing the continuation of the first and second. The cliff between the second and third terraces may or may not continue farther down-stream than that between the first and second. The plains below the terraces finally become continuous with the lowest flood-plain and with the delta. These212 relations can rarely be seen because of the destruction of the older terraces, and because of the erosion by waves along shore.

The topography of terraces is similar to that of flood-plains, except in so far as modified by erosion. While flat in general, the terrace may slope either toward or from the valley bluff, and its surface may be marked by all the minor irregularities which characterize a flood-plain.

Many familiar facts demonstrate the general presence of abundant water beneath the surface of the land. The thousands of wells in regions peopled by civilized races, and the countless springs which issue from the sides of mountains and valleys are a sufficient proof both of the wide distribution of ground-water and of its great abundance.

Certain well-known facts make it clear that ground-water is intimately connected with rainfall. In a dry season the level of the water in wells commonly sinks, and after a heavy rain it rises (p. 71); and the amount of sinking is greater when the drought is long, and the rise is most notable when the rainfall is heavy. Many springs which discharge large quantities of water during a wet season flow with reduced volume, or cease to flow altogether in periods of drought. Furthermore, the water of springs and wells has the properties which rain-water would possess after sinking beneath the surface and dissolving mineral substances. Rain-water is seen to sink beneath the surface with every shower, and since this source seems altogether adequate for ground-water, and since no other source is known whence any considerable amount of ground-water might come, it is concluded that atmospheric precipitation is its chief source.

Water gets beneath the surface by processes which are readily seen. Wherever the soil is porous some of the rain which falls upon it is absorbed. Sinking through the soil to the solid rock it finds cracks and pores through which it descends to great depths. Nowhere are the rocks beneath the mantle rock so compact and so free from cracks, when any considerable area is considered, as to prevent the percolation of water through them.

Conditions influencing descent of rain-water.—There are several conditions which influence not only the amount of water which sinks beneath the surface in a given area, but the proportion of the precipitation which follows this course. These are as follows: (1) Amount of precipitation.—In a general way it is true that the greater214 the amount of precipitation the greater the amount of water which will sink beneath the surface. Other things being equal, a region of heavy precipitation is a region where wells are easily obtained and springs common. (2) Rate of precipitation.—A given amount of precipitation may be concentrated in a short interval, or distributed through a considerable period of time. In the latter case more of the water sinks beneath the surface; in the former, a larger proportion runs off over the surface. The reason is readily seen. Water passes through small spaces, such as those of soil, slowly, and its rate of passage decreases rapidly with decreasing size of the passageways. When rain falls rapidly on a surface of even moderately close texture, the uppermost layer of soil is promptly filled with water, and since the water passes downward slowly, the uppermost saturated part of the soil becomes virtually impervious. While in this condition, the water which falls on it will run off if there be slope, and stand if there be none. In the latter case it will sink slowly as the water in the soil passes down to lower levels. If precipitation takes place no faster than the water can sink through the soil, all the water may become ground-water. (3) The topography of the surface has much to do with determining the proportion of rainfall which becomes ground-water. If the surface be flat, more will sink in; if it be sloping, more of it will run off before it has time to sink. Other things being equal, the steeper the slope the larger the proportion of the rainfall which will run off over it. (4) The texture of the soil, or other material on which the rain falls, helps to determine what proportion of it sinks beneath the surface. If the surface materials be porous, the water sinks readily; if of close texture, it finds less ready ingress. Other things being equal, the closer the texture of the soil the less the proportion of the rainfall which will enter it. (5) The texture and structure of the rock beneath the surface have some influence on the amount of ground-water. The rock may be stratified or massive; it may be abundantly or sparsely jointed; it may be porous or compact. On the whole, stratified rock is more favorable for the entrance of water than unstratified, partly because of its greater average porosity, and partly because the planes of division between beds often allow the passage of water. If the beds of stratified rock are vertical or inclined, water finds its way into them more readily than if they are horizontal, in so far as it descends along stratification215 planes. Horizontally bedded rock, or rock which is not bedded at all, may be so much jointed, and the joints so open, as to allow the water to enter readily.

The conditions favorable to the sinking of abundant water below the surface are therefore heavy precipitation, falling slowly on a surface with little relief, a soil of open texture underlain by rock which is porous, or affected by vertical or highly inclined planes of cleavage. The annual discharge of water by rivers is estimated to be about 22 percent. of the rainfall on the land.[93]

Supply of ground-water not altogether dependent on local rainfall.—The amount of ground-water in a given region is not always entirely dependent on the local rainfall. Ground-water is in constant movement, and entering the soil or rock at one point it may, after a long subterranean journey, reach a point far from that at which it entered. Thus beneath the Great Plains of the West there is much subterranean water which fell on the eastern slopes of the mountains to the west. It has flowed beneath the surface to the Plains, where some of it is now withdrawn for the purposes of irrigation in regions where rainfall is deficient. The accompanying diagram (Fig. 199) illustrates the flow here described.

The ground-water surface. Water table.—The water table has already been defined (p. 71) as the upper surface of the ground-water. In a flat region of uniform structure the ground-water surface is essentially level, but rises and falls with the rainfall. Where the 216topography of a region is not flat, the ground-water surface is not level. As a rule it is higher, though farther below the surface, under an elevation than under surrounding lowlands, as illustrated by Fig. 200. The explanation is not far to seek. If a hill of sand be exposed to rainfall, most of the water falling on its porous surface will sink into it. If the precipitation continues long enough, as in a protracted rain, the hill of sand will be filled with water, the water occupying the interstices between the grains. If the sand of the hill could be removed, leaving the water which it contains on the same area, it would constitute a mound perhaps a third or a fourth as high as the hill itself. If unsupported, this mound of water would spread promptly in all directions until its surface was level. While the sand remains, the water in it constitutes a mound, and has a tendency to spread. It does in fact spread, but since the process involves great friction the spreading is slow. With the spreading the surface of the water in the sand sinks, and sinks fastest at the center where it is highest (b, Fig. 201). If the process were not interrupted the surface of the water in the hill would, in time, sink approximately to the level of the water in the surrounding land (d, Fig. 201); but at every stage preceding the last, the surface of the water would be higher beneath the summit of the hill than elsewhere, though farther from the surface. In regions of even moderate precipitation the water surface beneath the hills rarely sinks to the level of that in the lowlands adjacent, before being raised by further rains.

The water-level beneath the lowlands also sinks. Some of it finds its way into valleys, some of it sinks to greater depths, and some of it evaporates; but since the water surface beneath the elevation sinks more rapidly than that beneath the lowland, the two approach a common level. Their difference will be least at the end of a drought, and greatest just after heavy rains.

Depth to which ground-water sinks.—The depth to which ground-water penetrates has not been determined empirically. No borings217 or excavations of any sort have been made to such depths as to indicate that its limit was being approached, though some of them are a mile or more deep. There is a popular belief that water sinks until it reaches a temperature sufficient to convert it into steam, but except for special localities where hot lava lies near the surface, this belief is not well founded. In the first place, it is not known at what temperature water below the surface would be converted into steam. While water boils at sea-level at a temperature of 212° (Fahr.) a higher temperature would be necessary below that level.

Assuming the temperature of water sinking beneath the surface to be 50° Fahr., its temperature must be raised 162° to bring it to the temperature at which it would boil at sea-level. On the above assumption of initial temperature, the following table shows the depths at which water would reach a temperature of 212° Fahr. under various assumptions as to the rate of increase of temperature. It shows also the pressure in atmospheres which would exist at these several depths if the overlying rock were full of water.

The temperature at which water boils increases with the pressure. A pressure of about 200 atmospheres is the critical pressure for water; that is, the pressure which, if increased, will prevent boiling altogether. The depth at which a pressure of 200 atmospheres would be reached, supposing the upper rock to be full of water, is about 6800 feet. The temperature of the water at this depth, under various assumptions as to initial temperature and rate of increase of heat, is shown in the following table:

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Only one of these temperatures reaches the boiling-point of water at sea-level. It is therefore clear that at this depth water has not even closely approached the boiling temperature for this depth, and since this is the depth of the critical pressure, it is clear that it cannot boil at any greater depth. The descent of water is therefore not stopped, under normal conditions of crustal temperature, because it reaches its boiling-point. Locally, as in the vicinity of active or recently extinct volcanoes, the case may be different.

It is conceivable that water may descend until it reaches its critical temperature (somewhere between 610° and 635° Fahr.). The depth at which the critical temperature would be reached, under various assumptions, is shown in the following table:

There is good reason, in the increasing density beneath the surface, for believing that the rate of increase of temperature decreases with depth, and therefore that the rate of 1° for 50 feet for the depths concerned is too high. The greater depths of the table above are therefore believed to more nearly represent the truth than the lesser ones. (See discussion of underground temperature in Chapter XI.)

If descending water attained its critical temperature, the extent to which the resulting water-gas might be absorbed is not known. So far as limited by temperature, therefore, it is not possible to assign a limit to the descent of water under average conditions of crustal temperature.

Other considerations seem to place a limit to the descent of water. Rock, solid and unyielding as it seems, is yet plastic when under sufficiently219 great pressure. The cracks and cavities affecting it are believed to descend a distance which is but slight in comparison with the radius of the earth. Even if openings were once formed at greater depths, they could not persist, for the adjacent rock, under the pressure which there exists, would “flow” in, in effect (though perhaps not in principle) much as stiff liquid might, and close them. The outer zone of the earth where cavities may exist is known as the zone of fracture.[94] The depth of the zone of fracture differs for different rocks, but is not believed to extend below some such depth as five or six miles, even for the most resistant.[95] It is to be noted that these depths are less than those at which the critical temperature of water would be reached under most of the conditions, including all the more probable ones, specified in the above table.

Let it be assumed that water descends through openings in the rock to a depth of six miles. At this depth it would, under the various assumptions specified in the first and second columns of the following table, have the temperature indicated in the third column:

In two of these cases, namely, those in which the assumed rate of increase of temperature is highest, the temperature of the water at the assumed lower limit of the zone of fracture is above the critical temperature of water. If the assumptions involved in these two cases be correct, water might descend to the point where it would be converted into water-gas, and in this condition it might be occluded by the hot rock. In the other cases, involving the more probable assumptions, the critical temperature is not reached at a depth of six miles. If pores and cracks do not extend to greater depths, liquid water could not; and since the water at this depth has probably not reached its critical temperature, it cannot exist as water-gas. If it does not exist 220in the form of water-gas, its occlusion by the hot rock substance would not be probable. It would seem, therefore, that the descent of water under ordinary conditions is much more likely to be limited by the zone of fracture, than by temperature.

Movement of ground-water.[96]—Ground-water is in more or less continual movement. If all the water be pumped out of a well it soon fills up again to its normal level by inflow from all sides. Springs and flowing wells also demonstrate the movement of ground-water. Near the surface the movement of ground-water is primarily downward if the medium through which it passes is equally permeable in all directions; but so soon as the descending water reaches the water surface, its descent is checked and its movement is partly lateral.

The commonest sort of movement of ground-water is that exemplified as the water sinks beneath the surface, namely, slow percolation through the pores and cracks of the soil and rock. Ground-water is not generally organized into definite streams, though underground streams, mostly small, are sometimes seen in caves and crevices, and sometimes issue as springs. Most underground streams which issue as springs probably have definite channels for short distances only before they issue. It is probable that ground-water frequently flows in considerable quantity along somewhat definite planes, without having open channels. Thus every porous bed of rock is likely to serve as the pathway along which subterranean drainage passes. This is especially true where the porous bed is underlain by an impervious one. The “reservoirs” from which artesian wells draw their supply are not usually streams or lakes, but porous beds of rock through which abundant water passes. As the supply is drawn off at one point, it is renewed by water entering elsewhere. Since the freedom of movement of ground-water is notably influenced by the porosity of the rock, and since the rock is, on the average, most porous and the pores largest near the surface, the movement of ground-water is, on the average, greatest near the surface, and least at its lower limit. In general the decrease of movement is much more rapid than the decrease in the size of the pores. It follows that while the upper part of the ground-water, especially that above ground-water level, moves somewhat freely, the lower part moves 221much more slowly. It is probable, indeed, that the movement in the lower part of the subterranean hydrosphere is extremely slight.

The amount of ground-water.—The porosity of surface rocks varies widely, and the porosity of but few has been determined.[97] Such determinations as have been made are chiefly on building stones, in which the range of porosity varies from a fraction of one percent., in the case of granite, to nearly 30 percent. in the case of some sandstones. Building stone is perhaps more dense than the average surface rock. Furthermore, such tests as have been made do not take account of the larger cracks and openings of rock, for these would not appear in the specimens tested; nor of the mantle rock, which generally contains a large amount of water. From such determinations as have been made it is estimated that the average porosity of the outer part of the lithosphere is somewhere between 5 and 10 percent. If the porosity diminishes regularly to a depth of six miles, where it becomes zero, the average porosity to this depth would be half the surface porosity.[98] An average porosity of two and one-half percent. would mean that the rock contains enough water to form a layer nearly 800 feet deep. With an average porosity of 5 percent., this figure would be doubled.[99] While these figures are not to be regarded as measurements, they perhaps give some idea of the amount of ground-water. It is this sphere of ground-water which justifies the term hydrosphere, as applied to the waters of the earth.

Fate of ground-water.—Most of the water which sinks into the earth reaches the surface again after a longer or shorter journey. Some of it is evaporated from the surface directly; some of it is taken up by plants and is passed by them into the atmosphere; some of it issues in the form of springs; some of it seeps out; some of it is drawn out through wells; and much of the remainder finds its way underground to the sea or to lakes, issuing as springs beneath them. A small portion of the descending waters enters into permanent combination with 222mineral matter. Many minerals are known to take up water, being changed thereby from an anhydrous to a hydrous condition. It does not necessarily follow, however, that the total supply of water is thereby decreasing. Minerals once hydrated may be dehydrated subsequently, the water being set free. Furthermore, considerable quantities of water in the form of vapor issue from volcanoes, and volcanic vents often continue to steam long after volcanic action proper has ceased. It is probable that some, and perhaps much of the water issuing from these vents has never been at the surface before, and it is not now possible to affirm that the supply from this source does not offset, or even surpass, the depletion of the hydrosphere resulting from mineral hydration.

THE WORK OF GROUND-WATER.

Ground-water effects very considerable results in the course of its history. These results are partly chemical and partly mechanical, the former being far more important than the latter.

Chemical Work.

The results of the chemical and chemico-physical action of water may be grouped in several more or less distinct categories.

1. The simplest effect is the subtraction of soluble mineral matter. Pure water is in itself a solvent of certain minerals; but the carbonic-acid gas extracted from the atmosphere, and the products of organic decay extracted from the soil make ground-water a much more efficient solvent. Something of the results which it achieves is shown by its composition. All ground-water, whether issuing as springs or drawn out through wells, contains much more mineral matter than the water which falls as rain, and the excess is acquired in its underground course.

The subtraction of soluble matter from rock renders it porous. The amount of material dissolved from a given place may be trivial or considerable, according to the character of the rock, the readiness with which water has access to it, and the character of the water. Locally, the subtraction of mineral matter may be the chief, or even the only appreciable, effect of the ground-water.

2. It sometimes happens that ground-water with certain mineral substances in solution exchanges them for other substances extracted223 from the rock. Thus the process of substitution is effected. By this process the lime carbonate of a shell imbedded in rock may be removed, molecule by molecule, and some other substance, such as silica, left in its place. When the process is complete, the substance of the shell has been completely removed, though its form and structure are still preserved in the new material which has taken the place of the old. Buried logs are sometimes converted into stone by the substitution of mineral matter for the vegetable tissue. This is petrification. Petrification is altogether distinct from incrustation, which simply means the coating of an object with mineral matter. A bird’s nest may be incrusted with lime carbonate, but it is not thereby petrified. Solution is a necessary antecedent of substitution.

3. The materials which are subtracted from the rock at one point may be added to other rock elsewhere. Thus a third type of change, addition, is effected. Rock may at one time and place be rendered porous by the subtraction of some of its substance, and the openings thus formed may subsequently become the receptacles of deposits from solution. This is exemplified in the stalactitic deposits of many caves. Not uncommonly cracks and fissures are filled with mineral matter deposited by the waters which pass through them. Thus arise veins which, for the most part, are nothing more than cracks and crevices filled by mineral matter brought to them in solution, and precipitated on their walls. Most veins of metallic ores have originated in this way.

4. A further series of changes is effected by ground-water when it, or the mineral matter it contains, enters into combination with the mineral matter through which it passes. One of the commonest processes of this sort, hydration, has already been referred to (pp. 43 and 428); but in the development of many of the commoner hydrous minerals changes other than hydration are involved. These changes result in new mineral combinations, the new minerals being developed out of the old, usually with some additions or subtractions. In the long course of time changes of this sort may be very great, so great indeed that large bodies of rock are radically changed, both chemically and physically. Much of the old substance may remain, but it has entered into new and more stable combinations with the materials which the water has brought to it.

Quantitative importance of solution.—In general, solution is probably most effective at a relatively slight distance below the surface.224 In the outermost zone of mantle rock the materials are usually less soluble than below, for they often represent the residuum after the soluble parts of the formation from which they originated were dissolved out. Below this zone the rock contains more soluble matter, and the water, charged with organic matter in its descent through the soil, is in condition to dissolve it. At greater depths the water has become saturated to some extent, and, so far forth, less active. Here, too, the movement is less free. The increased pressure at considerable depths, on the other hand, facilitates solution, which must be understood to take place under proper circumstances in any zone reached by the water.

Calculations have been made which illustrate the quantitative importance of the solution effected by ground-water. The springs of Leuk (Switzerland) bring to the surface annually more than 2000 tons of calcium sulphate (gypsum) in solution, and in the same time the springs of Bath (England) bring up an amount of mineral matter in solution sufficient to make a column 9 feet in diameter, and 140 feet high.[100]

The amount of mineral matter in solution in streams is also significant, for while stream-water is not all derived from ground-water, much of it had such an origin. In the case of several streams, among them the Thames and the Elbe, careful estimates of the amount of dissolved mineral matter have been made. Though the Thames drains an area only about one-tenth as large as the State of New York, it is estimated to carry about 1500 tons of mineral matter in solution to the sea daily.[101]

From the uppermost 20,000 square miles of its drainage basin the Elbe is estimated to carry yearly about 1,370,000 tons of mineral matter in solution. Estimates of the amounts of material carried to the sea in solution by several rivers are given on pp. 102 and 103. Much of this matter was brought to the rivers by waters which had been underground before reaching the streams.

From these figures it is clear that we have to reckon here with a very considerable factor in the lowering of land surfaces. From the amount of lime carbonate carried by the Thames it has been estimated 225that the average amount of this material dissolved from the limestone area drained by this stream is 143[102] tons per square mile per year. It is estimated that, on the average, something like one-third as much matter is carried to the sea in solution as in the form of sediment, and that by this process alone land areas would be lowered something like one foot in 13,000 years.[103]

Deposition of mineral matter from solution.—The deposition of material from solution is effected in several ways. (1) It is sometimes deposited by evaporation. This is well shown where water seeps out on arid lands. The same process is illustrated when water is boiled. (2) Reduction of temperature often occasions deposition. In general, hot water is a better solvent of mineral matter than cold,[104] and if it issues with abundant mineral matter in solution the precipitation of some of it is likely to take place. (3) Plants sometimes cause the precipitation of mineral matter from solution. About some hot springs, even where the temperature of the water is very high small plants of low type (algæ) grow in profusion. In ways which are not perfectly understood these algæ extract the mineral matter from the hot water. They are now thought to be a chief factor in the deposits about the hot springs of the Yellowstone Park.[105] The influence of organisms on precipitation from solution is not confined to the waters of hot springs. (4) A fourth factor involved in the deposition of mineral matter from solution is pressure. Pressure increases the solvent power of water with respect to minerals directly; it produces the same effect indirectly by its effect on the solution of gases. As water charged with gas comes to the surface, the pressure is relieved and some of the gas escapes. Such mineral matter as was held in solution by the help of the gas which escapes is then precipitated. (5) Precipitation is also sometimes effected by the mingling of waters containing different mineral substances in solution. Such mingling of solutions would be most common along lines of ready subterranean 226flow, and while each portion of the water entering a crevice or porous bed may be able to keep its own mineral matter in solution, their mingling may involve chemical changes resulting in the formation of insoluble compounds, and therefore in deposition. This principle has probably been involved in the filling of many fissures and crevices, converting them into veins. (6) The escape of gases from water, whether from increase of temperature or by the disturbance of water, sometimes causes the deposition of mineral matter held in solution.

The deposition of material held in solution is most notable at two zones, one below that of most active solution, and the other at the surface, where evaporation is active. Under proper conditions, however, deposition may take place at any level reached by water.

Mechanical Work.

The mechanical work of ground-water is relatively unimportant. Wherever it is organized into definite streams, the channels through which it flows are likely to be increased by mechanical erosion as well as by solution. Either beneath the surface, or after the streams issue, the mechanical sediment carried will be deposited.

RESULTS OF THE WORK OF GROUND-WATER.

Weathering.—Where the solution effected by ground-water in any locality is slight and equally distributed, the result is to make the rock porous. If, for example, some of the cement of sandstone is dissolved, the texture of the rock becomes more open; but if all the cement be removed the rock is changed from sandstone to sand. If a complex crystalline rock contains among its many minerals some one which is227 more soluble than the others, that one may be dissolved. This has the effect of breaking up the rock, since each mineral acts as a binder for the rest. It might happen that no one of the minerals is dissolved completely, but that some one of them is decomposed by water, and certain of its constituents removed. Such change would be likely to cause the mineral so affected to crumble, and with its crumbling, if it be an important constituent of the rock, the integrity of the rock is destroyed. Where considerable chemical changes, especially subtractions, are going on, the rock is likely to crumble. The increase in volume attendant on hydration, etc., sometimes leads to the disruption of rock. These are phases of weathering. (For other phases of weathering see pp. 54 and 110.)

Caverns.[106]—Where local solution is very great results of another sort may be effected. In formations like limestone, which are relatively soluble, considerable quantities of material are frequently dissolved from a given place. Instead of making the rock porous, in the usual sense of the term, large caverns may be developed (Fig. 202). In their production, solution may be abetted by the mechanical action of the water passing through the openings which solution has developed. Considerable caves are found chiefly in limestone. They were probably developed when the surface relief was slight, and surface drainage therefore poor. Regions where caves were developed under these conditions may subsequently acquire relief, so that caves are not now confined to flat regions.

One of the best known regions of caves is in the basin of the Ohio in Kentucky and southern Indiana, where the number of caves is large, and the size of some of them, such as Mammoth and Wyandotte, very great. A ground-plan of Wyandotte (Ind.) Cave is shown in Fig. 203. The aggregate length of the passageways is about 23½ miles.

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Deposition often takes place in caves after they are formed (Figs. 204 and 205). It may even go on at the same time that the cave is being excavated. Here are formed the well-known stalactites and stalagmites. A stalactite may start from a drop of water leaking through the roof of the cave. Evaporation, or the escape of some of the carbonic gas in solution, results in the deposition of some of the lime carbonate about the margin of the drop, in the form of a ring. Successive drops make successive deposits on the lower edge of the ring, which grows downward into a hollow tube through which descending water passes, making its chief deposits at the end. Deposition in the tube may ultimately close it, while deposition on the outside, due to water trickling down in that position, may greatly enlarge it.

Underground caves sometimes give rise to topographic features which are of local importance. When the solution of material in a cavern has gone so far that its roof becomes thin and weak, it may collapse, giving rise to a sink or depression in the surface over the site231 of the original cave. This is so common that regions of limestone caves are often affected by frequent sinks formed in this way. They are a conspicuous feature of the landscape in the cave region of Kentucky, and are well known in many other limestone districts. They are known as limestone sinks. (Fig. 206 and Fig. 2, Pl. XVII.)

Under certain circumstances caves may give rise to striking features of another sort. If for any reason the roof is destroyed at the two ends of a cave, remaining intact over the middle, the latter part constitutes a natural bridge. Natural bridges also originate in other ways (pp. 151, 153).

Creep, slumps, and landslides.—When the soil and subsoil on a slope become charged with water they tend to move downward. When the movement is too slow to be sensible it is called creep. The common downward inclination of trees growing in such situations, the result of the more rapid creep of the surface as compared with the deeper part of the soil, is both an expression of the movement and of its slowness. Other factors besides ground-water are involved in creep (see p. 112).

When the movement is rapid enough to be sensible the material is said to slump or slide. This may happen when the slope on which water-charged mantle rock lies is steep (Fig. 207). Great landslides of this sort have been recorded, and some of them have done great damage. Where a stream’s banks are high, and of unindurated material, such as clay, considerable masses sometimes slump from the bank232 or bluff into the river, or settle away slowly from their former positions. This is a common phenomenon along streams which have cut valleys in drift, and along shores on which waves are encroaching. The same phenomenon is common on a larger scale on the slopes of steep mountains.[107] Considerable terraces are sometimes developed on their slopes in this way, but they are usually irregular and discontinuous (Figs. 208, and 209). The loose débris on steep slopes sometimes assumes a sort of flowing motion and descends the slope with some such form and at some such rate as a glacier. Such bodies of débris are sometimes called “talus glaciers” (Fig. 210). In many such cases, snow and ice have had some part in their development.

In creep and in landslides gravity is the force involved, and the ground-water only a condition which makes gravity effective. Gravity alone accomplishes similar results, as illustrated by Fig. 211.

Summary.

All in all, ground-water is to be looked upon as a most important geological agent. When it is remembered that a very large part of all the water which falls on the surface of the earth, either in the form of rain or snow, sinks beneath the surface; that much of it sinks to a great depth; that much of it has a long underground course before it reappears at the surface; that it is everywhere and always active, either in subtracting from the rock through which it passes, in adding to it, in effecting the substitution of one mineral substance for another, or in bringing about new chemical combinations; and when it is remembered that this process has been going on for untold millions of years, it will be seen that the total result accomplished must be stupendous. The rock formations of the earth to the depths to which ground-water penetrates are to be looked upon as a sort of chemical laboratory through which waters are circulating in all directions, charged with all sorts of mineral substances. Some of the substances in solution are deposited beneath the surface, and some are brought to the surface where the waters issue. Much of the material brought to the surface in solution is carried to the sea and utilized by marine organisms in the making of shells. Without the mineral matter brought to the sea by springs and rivers, many shell-bearing animals of great importance, geologically, would perish. Biologically, therefore, as well as geologically, ground-water is of great importance.

PLATE XVII.

U. S. Geol. Surv.

Scale, 1+ miles per inch.

U. S. Geol. Surv.

Scale, 1+ miles per inch.

PLATE XVIII.

U. S. Geol. Surv.

Scale, 2+ miles per inch.

U. S. Geol. Surv.

Scale, 2+ miles per inch.

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SPRINGS AND FLOWING WELLS.

The term spring is applied to any water which issues from beneath the surface with sufficient volume to cause a distinct current. If the water issues so slowly as to merely keep the surface moist, it is not called a spring, but seepage. The spring from which water issues with a strong current, especially if it be upward, is comparable to a flowing well, while the spring from which water issues with little force, and without upward movement, is comparable to the flow of water into a common well.

Springs often issue from the sides of valleys (Fig. 212), the bottoms of which are below ground-water level. They are especially likely to issue at the surface of relatively impervious layers, and where the valley slopes cut joints, porous beds, or other structures which allow free flow of ground-water.

Springs are classified in various ways, and these several classifications suggest characteristics worthy of note. They are sometimes said to be deep and shallow. The “deep” spring, as the term is ordinarily used, is one which issues with great force, and with something of upward movement, and the “shallow” spring, one which issues with little force, and without upward movement; but the spring which issues with force is not necessarily deep, nor is the one which issues with little force necessarily shallow. The idea involved in this grouping would be better expressed by strong and feeble. Springs are also classified as cold and thermal, the latter term meaning simply that the temperature is such as to make the springs seem warm or hot. The temperature of thermal springs ranges up to the boiling-point of water.235 Between deep springs and shallow ones, and between cold springs and thermal, respectively, there is no sharp line of demarkation. Again, some springs are continuous in their flow, while others are intermittent. Most intermittent springs flow after periods of precipitation, but dry up during droughts (see p. 202). Springs are also classified as mineral and common. Mineral springs, in the popular sense of the term, are of two types: (1) Those which contain an unusual amount of mineral matter, and (2) those which contain some unusual mineral. Springs are especially likely to be called mineral if the substances which they contain, have, or are supposed to have, some medicinal property. All springs which are not “mineral” are “common.” This classification is not altogether rational, for all springs contain more or less mineral matter, and many springs which are “common,” contain more mineral matter than some springs that are “mineral.” Mineral springs are themselves classified according to the kind and amount of mineral matter they contain. Thus saline springs contain salt; sulphur springs contain compounds (especially gaseous) of sulphur; chalybeate springs contain iron compounds, especially the sulphate; calcareous springs contain abundant lime carbonate, etc. These various mineral substances are extracted from the rock, sometimes by simple solution, and sometimes by solution resulting from other chemical change. The salt of saline springs is usually extracted from beds of salt beneath the surface. Lime carbonate, one of the commonest substances in solution in ground-water, is dissolved from limestone, or derived by chemical change from rocks containing other calcium compounds. Thus lime feldspars, by carbonation, give rise to lime carbonate. The chalybeate waters often arise from the oxidation of iron sulphide, a mineral which is common in many sedimentary rocks. The iron sulphate is itself subject to change in the presence of the ubiquitous lime carbonate. From this change iron carbonate results, and this is usually quickly altered to iron oxide, which, being relatively insoluble, is precipitated. About chalybeate springs, therefore, iron oxide is frequently being deposited. Medicinal springs are those which contain some substance or substances which have, or are supposed to have, curative properties.

Mineral matter in solution.—The number and variety of mineral substances in spring water is very great, and the amount of solid matter in solution varies widely. Some of the hot springs of the Yellowstone236 Park contain nearly three grams (2.8733) of mineral matter per kilogram.[108]

The composition of various spring and well waters is shown in the accompanying table, which gives some idea of the range of mineral substances commonly in solution in ground-water.

Geysers.—Geysers are intermittently eruptive hot springs. They occur only in volcanic regions (past or present) and in but few of them. Active geysers are virtually confined to the Yellowstone Park and Iceland, though they formerly existed at other places. Those of New Zealand have but recently become extinct. The great geyser region of the world is the Yellowstone Park, where there are said to be more than sixty active geysers.

The cause of the eruption is steam. The surface-water sinks down until, at some unknown depth, it comes into contact with rock sufficiently hot to boil it. The source of the heat is not open to inspection, but it is believed to be the uncooled part of an extrusive lava flow, or of an intrusive lava mass. From what was said on pp. 216 and 217 it is clear that geysers do not have their origin in water which sinks down to the zone of great heat, where the increment of heat is normal.

The water of a geyser issues through a tube of unknown length. Whether the tube is open down to the source of the heat is not determinable, but water from such a source finds its way to the tube. Water may enter the tube from all sides and at various levels from top to bottom. The heating may precede or follow its entrance into the tube, or both. So far as the water is heated after it enters the tube, the point of most rapid heating may be at the bottom of the tube or at some point above. If the temperature of the source of heat were high enough to convert the descending water into steam as fast as it enters the tube, the steam would escape continuously, though there would be no geyser; but if the rock is only hot enough to bring the water to the boiling-point after some lapse of time, and after some water has accumulated, an eruption is possible.

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The exact sequence of events which leads up to an eruption is not known, but a definite conception of the principles involved may perhaps be secured by a definite case. Suppose a geyser-tube filled with water, and heated at its lower end. As the water is heated below, convection tends to distribute the heat throughout the column of water above. If convection were free, and the tube short, the result would be a boiling spring; but if the tube is long, and especially if convection is impeded, the water at some level below the surface may be brought to the boiling-point earlier than that at the top. Under these circumstances if even a little water in the lower part of the tube is converted into steam, the steam will raise the column of water above, and it will overflow. The overflow relieves the pressure on all parts of the column of water below the surface. If before the overflow there was any considerable volume of water essentially ready to boil, the relief of pressure following the overflow might allow it to be converted into steam suddenly, and the sudden conversion of any considerable quantity of water into steam would cause the eruption of all the water above it (Figs. 213 and 214). The height to which the water would be thrown would depend upon the amount of steam, the size and straightness of the tube, etc.

It is clear that everything which impedes convection in the geyser tube will hasten the period of eruption, since impeded circulation will have the effect of holding the heat down, and so of bringing the water at some level below the top more quickly to the boiling-point. It follows that anything which chokes up the tube, or which increases the viscosity of the water, or its surface tension, would hasten an eruption.[119]

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Geysers often build up crater-like basins or cones (Figs. 214 to 217) about themselves, the cone being of material deposited from solution. In the Yellowstone Park the precipitation of the matter in solution (chiefly silica) is partly due to cooling and partly to the algæ which abound even in the boiling water, and the brilliant colors of the deposits about the springs are attributable to these plants. When the water from any geyser or hot spring ceases to flow the plants die and the colors disappear. The details of the surface of the deposits about geysers and hot springs are often complicated, and frequently very beautiful (Fig. 218).

The heating of geyser and hot-spring water must cool the lava or240 other source of heat below. As this takes place, the time between eruptions becomes longer and longer. In the course of time, therefore, the geyser must cease to be eruptive, and when this change is brought about the geyser becomes a hot spring. Within historic times several geysers have ceased to erupt and new ones have been developed. In the Yellowstone Park, where there are said to be something like 3000 vents of all sorts, hot springs which are not eruptive greatly outnumber the geysers. From many of the vents but little steam issues, and from some, little else.

A few geysers have somewhat definite periods of eruption. Of such “Old Faithful” is the type; but even this geyser, which formerly erupted at regular intervals of about an hour, is losing the reputation on which its name is based. Not only is its period of eruption lengthening, but it is becoming irregular, and the irregularity appears to be increasing. In the short time during which this geyser has been under241 observation its period has changed from a regular one of sixty minutes, or a little less, to an irregular one of seventy to ninety minutes. In the case of some geysers years elapse between eruptions, and in some the date of the last eruption is so distant that it is uncertain whether the vent should be looked upon as a geyser or merely a hot spring.

In the Yellowstone Park[120] the geysers are mainly in the bottoms of valleys (Fig. 219), but the deposits characteristic of geysers are found in not a few places well above the present bottoms. These deposits record the fact that in earlier times the geysers were at higher levels than now. It is probable that they have been, at all stages in their history, near the bottoms of the valleys, and that, as the valleys have been deepened the ground-water has found lower and lower points of issue. In this respect the geysers have probably had the same history as other springs.

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Unless new intrusions of lava occur, or unless heat is otherwise renewed at the proper points, it is probable that all existing geysers will become extinct within a time which is, geologically, short. New geyser regions may, however, develop as old ones disappear.

Artesian wells.—Originally the terms artesian wells and flowing wells were synonymous; but at the present time any notably deep well is called artesian, especially if it descends to considerable depths below the mantle rock. The artesian well which does not flow, does not differ from common wells in principle; but being deeper, the water which it affords is often more thoroughly filtered and frequently more highly mineralized than that of other wells. The flowing well is really a gushing spring, the opening of which was made by man.

Flowing wells[121] depend upon certain relations of rock structure, water supply, and elevation. Generally speaking a flowing well is possible in any place underlain by any considerable bed of porous rock, if such rock outcrops at a sufficiently higher level in a region of adequate rainfall, and is covered by a layer or bed of impervious, or relatively impervious rock. This statement involves four conditions, all of which are illustrated by Fig. 199, where a is the bed of porous rock. It is not necessary that the beds of rock form a structural basin, nor is it usually necessary to take account of the character of the rock beneath the porous bed which contains the water.

The bed of porous rock is the “reservoir” of the flowing well. Formations of sand or sandstone, and of gravel or conglomerate, most commonly serve as the reservoirs. In order that it may contain abundant water it must have some thickness, and its outcropping edge must be so situated that the water may enter freely and be replenished, chiefly by rain, as the water flows out at the well.

A relatively impervious layer of rock above the reservoir (b, Fig. 199) is most important; otherwise the water in the reservoir will leak out, and there will be little or no “head” at the well site. Thus if the rock overlying stratum a (Fig. 199) were badly broken, the fractures extending up to the surface, the conditions would be unfavorable for flowing wells. Under such conditions, wells in the positions of those shown in Fig. 199 might get abundant water, but they would not be likely to flow. If the 243stratum next below the reservoir is not impervious, some lower one probably is. No layer of rock is more impervious than one which is full of water, and the substructure of any bed which might serve as a reservoir is usually full of water, even if the rock be porous.

If the outcrop of the reservoir be notably above the site of the well, and if it be kept full by frequent rains, the “head” will be strong, though the water at the well will not rise to the level of the outcrop of the reservoir. Experience has shown that an allowance of about one foot per mile of subterranean flow should be made. Thus if the site of the well be 100 miles from the outcrop of the water-bearing stratum, and 200 feet below it, the water will rise something like 100 feet above the surface at the well. This rule is, however, not applicable everywhere. The failure of the water to rise to the level of its head is due to the adhesion and the friction of flow through the rock. The more porous the rock the less the reduction of head by friction. The height of the flow is also influenced by the number of wells drawing on the same reservoir, on the degree of imperviousness of the confining bed above, etc.

Flowing wells, often relatively shallow, are frequently obtained from unconsolidated drift. Some such relations as suggested by Fig. 220 would afford the conditions for flowing wells in such a formation.

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- AND ICE-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.

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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 general246 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 more247 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 partially248 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 be249 moulded into almost any desired shape if carefully subjected to sufficient pressure, steadily applied through long intervals of time.[123] These 250changes 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 short251 (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.

TYPES OF GLACIERS.

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.

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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,256 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.

THE GENERAL PHENOMENA OF GLACIERS.[125]

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.257 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 denominations258 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 with259 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.

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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 262average. 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 263has 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 numerous264 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. Longitudinal265 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,266 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.

SURFACE FEATURES.

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 268middle. 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,269 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.

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 the270 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.

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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 laminæ273 are much contorted. This is seen both in section (Figs. 248 and 249) and on the surface (Fig. 250).

TEMPERATURE, WASTE, AND DRAINAGE.

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.

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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.

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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.

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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 partly277 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°278 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,279 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 apparently280 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 which281 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 OF GLACIERS.

Erosion and transportation.

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.

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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 resistance283 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, 284beyond 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, favorable285 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 be286 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 of287 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.

PLATE XIX.

U. S. Geol. Surv.

Scale, 2+ miles per inch.

PLATE XX.

U. S. Geol. Surv.

Scale, 1+ miles per inch.

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).

288

289

290

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 which291 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 normally292 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.

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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,294 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.

Superglacial débris may obviously become englacial by falling into crevasses or by being carried down by descending waters. Either superglacial or englacial débris may become basal by the same means.

From their form and position, there is less ice-free land in immediate association with ice-caps than with valley glaciers. Furthermore, the ice-free land about the borders of an ice-cap is less likely to be in the295 form of cliffs above it. As a result, the surfaces of ice-caps are comparatively clean, except at their edges where the ice is thin.

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 its level. This occurs chiefly at the end or edge of the ice,296 where the surface melting is greatest. Englacial débris, especially that near the bottom, may also become basal by the melting of the bottom of the ice.

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, 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.

Much less in the natural order of things is the transfer of material from a basal to an englacial and from an englacial to a superglacial position by upward movement of the débris itself. Such transfer is remarkable because the specific gravity of rock is from two and a half to three times as great as that of ice, so that its normal tendency is to sink.

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. 268).

At the borders of arctic glaciers the lower layers are not infrequently upturned, as shown in Figs. 269 to 272. Where the layers turn up at the end of a glacier (Figs. 269 and 270), basal and englacial débris is297 carried to the surface by actual upward movement, and a terminal moraine or a series of terminal moraines sometimes aggregated where the upturned layers of ice outcrop at the surface (Fig. 271). That the material of these moraines was originally basal is abundantly demonstrated by the bruised and scratched condition of the bowlders and pebbles, and sometimes by the nature of the material itself. For example, in two cases in North Greenland where glaciers descend into the heads of shallow bays and move forward on their bottoms, moraines formed by the upturning of the layers were seen to contain abundant molluscan shells derived from the bottom of the bay. The upturning sometimes affects the side-edges of ice-tongues (Fig. 272) as well as their ends, and the material thus brought to the surface gives origin to lateral moraines altogether different in origin from the lateral moraines formed by the falling of débris upon the glaciers. Sometimes also there is an upturning of the ice along a longitudinal zone well back from the lateral margins (Fig. 273), and the material so borne to the surface in such a zone gives rise to a moraine resembling the medial moraine formed by the union of lateral moraines, but of wholly different origin.

The phenomenon of upturning here referred to has been observed only at or near the terminus of the ice, and is perhaps due in most 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 increased rigidity of the ice at its borders, 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. In not a few instances the upturning is associated with a notable thickening of the layers toward their edges (Fig. 274). This suggests that perhaps there is an exceptional growth of the granular crystals of the ice298 near the edge of the layers, owing to the penetration of the surface-waters which are much more abundant at the borders than elsewhere, and which in the arctic glaciers probably do not penetrate deeply before they reach a freezing temperature.

Wear of drift in transit.—Drift carried at the bottom of the ice is subject to notable wear. The materials in transportation 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.

Deposition of the Drift.

1. Beneath the body of the ice.—During the advance of a glacier, deposition may take place both beneath the body of the ice and beneath its end and edges. Deposition beneath the body of the ice is liable to take place wherever the topography favors lodgment, or wherever the ice is overloaded. The topography favoring deposition is much the299 same as that favoring erosion, but the two processes are not favored at the same point. Erosion is greatest on the “stoss” side of an obstruction (the side against which the ice advances), and deposition on the lee side. 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, however, the deposition of material beneath the main body of a glacier is much more than balanced by erosion in the same position.

2. At ends and edges of glaciers.—At and near the end of a glacier the conditions of deposition are somewhat different. Here deposition beneath the ice goes on faster than elsewhere, chiefly because of the more rapid melting and the more rapid thinning and weakening of the ice. If the end of the glacier be stationary in position, drift is being continually brought to it and left there, for though the end is stationary, the ice continues to move. If the glacier moves forward 500 feet per year, and if its end is melted at the same rate, all the débris in the 500 feet of ice which has been melted has been deposited, and all except that which has been washed away has been deposited at and300 beneath the end of the glacier. 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 300 feet per year, and is melted back 500 feet in the same time, all the débris carried by the 500 feet which has been melted has been deposited, and largely in the narrow zone (200 feet) from which the ice has receded. Even in this case, therefore, there is a notable tendency to marginal accumulation. If the end of the glacier is advancing 500 feet per year while it is being melted but 300 feet, all the drift in the 300 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. This addition is sometimes considerable and sometimes meagre. If the edge of the ice is without much fluctuation in position, the material dumped over its end may take the form of a narrow ridge or bowlder-wall (Geschiebe-wall). If a glacier pushes material in front of it, this, too, becomes a part of the general terminal aggregation of drift.

301

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 stationary, for a considerable time, is the terminal moraine. That part of the aggregation deposited beneath the ice is sometimes called the lodge moraine (Figs. 275 and 276; see also Fig. 235); that carried on the ice and dropped at its edge, the dump moraine; and that pushed before the ice, the push moraine. Many moraines marginal to the ice appear to be push moraines, when they are really lodge moraines from which the ice has withdrawn (Fig. 277). The push moraine can rarely be distinguished, and the dump moraine by no means always, after the disappearance of the ice.

The ground moraine.—When a glacier disappears by melting, all its débris is deposited. All the drift deposited beneath the advancing302 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 of the glacier and thinner toward its source, but considerable portions of a glacier’s bed are often left without débris when the ice melts. In general, the terminal moraine is not only thicker, but more irregularly disposed than the ground moraine.

The lateral moraines.—The surface lateral moraines of valley glaciers are let down on the surface beneath when the ice melts out from under them; 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 itself at any one time (Fig. 278). As a glacier retreats, its lateral moraine material is more or less bunched. Thus if the ice advances 200 feet while its end is being melted back 300 feet, the lateral moraines on the 300 feet melted are concentrated into 100 feet, as they are delivered on to the land by the melting of the ice from beneath. If the retreat of the end of a glacier be very slow, the bunching may be great. But even this cannot explain the massiveness of some lateral moraines. Furthermore, the materials of which many lateral moraines are composed are nearly as well worn as those of the ground moraine. The massive lateral moraines of which this is true are often made up303 chiefly of the drift accumulated beneath the lateral margins of the glaciers. This accumulation is the result of the lateral motion of the ice from center to side. Such sublateral 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 several hundred feet high, and in one case about 1000 feet. In northern Italy lateral moraines are said to be 1500 to 2000 feet high.[132]

Most of the material which was englacial during the transportation becomes either subglacial or superglacial before deposition, for it ordinarily reaches the bottom or the top of the ice before being deposited. Where the ends or edges of a glacier are vertical or nearly so, as in the high arctic regions, deposition may take place from the englacial position directly.

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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. 279). In this respect, the deposits of ice differ notably from those of water. Furthermore, many stones of the drift show the peculiar type of wear which glaciers inflict. They are not rounded as the stones carried by rivers, though they are notably worn. Many of them have subangular forms with planed and beveled faces, the planes being striated and bruised (Fig. 254). The absence of stratification, the 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 less obvious characteristic is the constitution of the fine material, for it is in general not the product of rock decay, but of rock grinding. The fine material handled by streams (except glacial streams) on the other hand, is usually the product 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 are marked by grooves and striæ, similar to those on the stones of the drift (Figs. 255 and 256). Other distinctive features of a glaciated area are the rounded bosses of rock (roches moutonnées, Fig. 280; see also surface about the lakes,305 Fig. 261), the rock basins, the lakes (Fig. 261), ponds, and marshes, and the peculiar topographies resulting from the unequal erosion, and the still more unequal deposition of the drift. 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.

GLACIO-FLUVIAL WORK.

The constant but unequal waste of glaciers has already been referred to. The streams to which this gives rise are usually laden with gravel, sand and silt derived from the ice. Since the mud is often light-colored, the streams are sometimes 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. 281 and 282), or where they enter lakes (Fig. 283), bays, or other streams, deltas. In its transportation, the river-borne drift is assorted; after its deposition, it is stratified. True glacial deposits in the upper part of a mountain valley are, therefore, often continued below by glacio-fluvial deposits derived from the same source.

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The most common form of such deposit is a valley train (Fig. 281) of glacial wash stretching indefinitely down the valley. The silt, sand, and gravel of such trains can usually be distinguished from valley deposits of non-glacial origin by the character of the material, as much of it is the product of grinding, crushing, and fracture, rather than of ordinary surface decay. Its materials are, therefore, fresh and often include rock material which, if long exposed at the surface, would be decomposed or dissolved.

Where an ice sheet ends in a broad face, as did the ancient continental glaciers, numerous streams flow from it and spread their débris in front of the terminal moraine, forming a broad fringing sheet or “apron” (outwash plain) along it. Where streams of considerable size form tunnels under or in the ice, these may become more or less filled with wash, and when the ice melts the aggraded channels appear as long ridges of gravel and sand known as eskers (osars and serpentine kames and kames of authors. See chapter on glacial period). It has been thought that similar ridges are sometimes formed in valleys cut in the ice from top to bottom, and even that they arise from gravel and sand lodged in superglacial channels. The latter at least is probably rare, as the surface streams usually have high gradients, swift currents, and smooth bottoms, and hence give little opportunity for lodgment.307 In the case of ice-sheets, too, in connection with which eskers are chiefly developed, there is usually 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, sands and gravels are liable to be bunched in quantity, giving rise, after the adjacent ice has melted, to peculiar hills and hollows of the knob-and-basin type. The hills and short ridges are known as kames (see glacial period). Subglacial streams may leave washed and assorted material in their tracks under the ice, and this is sometimes buried under deposits made by the ice itself, so that glacio-fluvial and glacial deposits are interbedded.

ICEBERGS.

When glaciers advance into water, the depth of which approaches their thickness, their ends are broken off (Fig. 284), and the detached masses float away as icebergs (Fig. 285). Many of the bergs are overturned, or at least tilted, as they set sail. If this does not happen at308 the outset, it is likely to occur later as the result of the melting and wave-cutting which disturb their equilibrium. The great majority of bergs do not travel far before losing all trace of stony and earthy débris, but the finding of glacial material in dredgings far south of all glaciers shows that they occasionally carry stones far from land.

THE INTIMATE STRUCTURE AND THE MOVEMENT OF GLACIERS.

With the preceding account of glaciers in mind, we may return to a closer study of their origin, their intimate structure, and their mode of motion. The key to this study is the thesis that a glacier is a mass of crystalline rock—the purest and simplest type of crystalline rock known—since it is made up of a single mineral of simple composition and rare purity, which never appears in a solid state except in the crystalline form.

The growth and constitution of a glacier.—The origin and history of a glacier is little more than the origin and aggregate history of the crystals that compose it. The fundamental conception of a glacier is therefore best obtained by tracing the growth of its constituent crystals. A basal fact ever to be kept in mind is that water in the solid form is always controlled by crystalline forces. When it solidifies from the vapor of the atmosphere it takes the form of separate crystals (Figs. 286–291). Perfect forms are developed only when the flakes fall quietly through a saturated atmosphere which allows them to grow as they descend. Under other conditions, the crystals are imperfect in growth and are mutilated by impact. But however modified, they are always309 crystals. The molecules are arranged on the hexagonal plan, and, as the expansive power of freezing water shows, the arrangement is controlled by a strong force. Once the definite crystalline arrangement is established, the molecules can be displaced only by relatively great force.

Snow crystals often continue to grow so long as they are in the atmosphere; but if they pass through an under-saturated stratum of air or a stratum whose temperature is above 32° Fahr., they suffer from evaporation or melting. When they reach the ground, the processes of growth and decadence continue, and the crystals grow or diminish according to circumstances.

A glacier is a colossal aggregation of crystals grown from snowflakes to granules of much greater sizes. The microscopic study of new-fallen snow reveals the mode of change from flakes to granules. The slender points and angles of the former yield to melting and evaporation more than the more massive central portions, and this change probably illustrates a law of vital importance. It may often be seen that the water melted from the periphery of a flake gathers about its center, and if the temperature be right, it freezes there. This is a first step toward the pronounced granulation of snow which has lain for some time on the ground. If measured systematically from day to day, the larger granules taken from beneath the surface of this coarse-grained snow are found to be growing. In a series of experiments[133] to determine310 the law of growth it was found that when the temperature of the atmosphere was above the melting-point the growth was appreciably more rapid than when the air was colder, but there was, on the average, an increase under all conditions of temperature. A portion of this average increase of the larger granules appears to come from the diminution and destruction of the smaller ones, for the total number of granules steadily diminishes. A portion of the growth doubtless comes from the moisture of the atmosphere which penetrates the snow and another portion from the moisture derived from surface melting; but beneath the surface of a large body of snow the growth of the large granules is probably chiefly at the expense of the small ones. To follow the process it should be noted that the free surface of every granule is constantly throwing off particles of water-vapor (evaporation); that the rate at which the particles are thrown off is dependent, among other things, on the curvature of the surface, being greater the sharper the curve; that the surfaces of the granules are at the same time liable to receive and311 retain molecules thrown from other granules, and that, other things being equal, the retention of particles also depends on the curvature of the surface, the less curved surface retaining more than the sharply curved one. Under these laws, it is obvious that the larger granules of smaller curvature will lose less and gain more, on the average, than the smaller granules of greater curvature. It follows that the larger granules will grow at the expense of the smaller. It is also to be noted that, other things being equal, small granules melt more readily than large ones, and that where the temperature is nicely adjusted between melting and freezing, the smaller may lose while the larger gain.

Another factor that enters into the process is that of pressure and tension. The granules are compressed at the points of contact and put under tension at points not in contact, and the pressure and tension are, on the average, likely to be relatively greatest for the smallest granules. Tension increases the tendency to evaporation and adds its effects to curvature, and the capillary spaces adjoining the points of contact probably favor condensation. Ice expands in crystallizing and pressure reduces the melting-point, while tension raises it. The effect of this is slight (p. 276), and it probably plays little part in glacial action, but it is to be correlated with the much more important fact that compression produces heat which may raise the temperature of the ice to the melting-point, while tension may reduce the temperature to or below freezing. 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 more or less solid mass.

A third factor is also to be recognized, though its effectiveness is unknown. Under severe wind pressure, air penetrates porous bodies with appreciable facility. The “breathing” of soils and the curious phenomena of “blowing-wells” and “blowing-caves” teach us 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 allow of the appreciable penetration of the atmosphere. During a part of the time, the probable effect is the condensation within the ice of moisture312 from the air, and during another part, evaporation from the ice. These alternating processes are attended by oscillations of temperature. While the balance between loss and gain of substance may be immaterial, the oscillating nature of the process and the fluctuations of temperature are probably favorable to granular change.

Whether these processes furnish an adequate explanation of the changes or not, the observed fact is that there are all gradations from snowflakes and pellets into granular névé, and thence into glacier granules (Gletscherkörner), varying in size up to that of filberts and walnuts, and even beyond. In coherence, these aggregations may vary from the early slightly coherent granular stage, where the grains are small and spheroidal, to the ice stage, where the cohesion has become strong through the interlocking growths of the large granules. Even when the mass has become seemingly solid ice, sufficient space is usually left between the granules to give the dispersive reflection to light which imparts to glacier ice its distinctive whitish color.

The arrangement of the crystal axes.—The most radical difference between glacier ice and ice formed directly from water is in the arrangement (orientation) of the crystals. In the ice formed on undisturbed water, the bases of the crystals are at the surface and their principal axes are vertical, as shown by Mügge.[134] As they grow, the crystal prisms extend downwards. This gives a columnar or prismatic structure to the ice, well seen when it is “honeycombed” by partial melting. In the glacier, on the other hand, 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 313crystals continue to lie in all directions. Hence glacier ice, unlike pond ice, cannot usually be split along definite planes, except where cleavage planes are subsequently developed by extraneous agencies.

While the crystals of a glacier usually have their principal axes in various directions, there appears to be a tendency for them to approach parallelism in certain positions, especially in the basal parts of a glacier near its terminus. Observations on this point are not so full and critical as could be desired, but it is probable that the parallel orientation is partly general, and due to the vertical pressure of the ice, and partly special and local, and connected with the shearing planes and foliation.

The bearing of this partial parallelism of the crystals on shearing and foliation is supposed to reside in the fact that a crystal of ice is made up of a series of plates arranged at right 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 vertical to them. If a cube be cut from a large crystal of ice, it will behave much like a cube cut from the pile of cards. If the cube be so placed that its plates are horizontal (Fig. 291a), and if it be rested on supports at two edges and heavily weighted in the middle, it will sag, the plates sliding slightly over one another so as to give oblique ends, but in this case the cube offers considerable resistance to deformation. If the cube be so placed that the plates stand on edge, each reaching from support to support (Fig. 291b), it will offer very great resistance to deformation; but if the plates be vertical and transverse to the line joining the supports, as in Fig. 291c, the middle portion will sag under very 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. These properties have been demonstrated by McConnel[135] and Mügge, and they appear to throw light on certain phases of the action of glaciers that are most pronounced in their basal parts, and are best illustrated in arctic glaciers.

The Probable Fundamental Element in Glacial Motion.

Melting and freezing.—It has already been shown (p. 279) that the initial or fundamental cause of glacial motion must be operative at the 314heads of glaciers where the temperature is lowest and the material most loosely granular. In this condition, there is reason to believe that motion takes place between the grains, rather than by their distortion through the displacement of their laminæ. The fact that the granular structure is not destroyed, as it would be by the indefinite sliding of the crystal plates over each other, sustains this view. The inference is that the gliding planes play a notable rôle in glacial movement only in the basal parts of the lower ends of glaciers, where the greatest thrusts are developed, and where the granules have become largest and most completely interlocked. At the heads of glaciers, where motion is initiated, there may be great downward pressure, but not vigorous thrusts from behind, and probably only moderate thrusts developed within the body itself. There seems therefore no escape from the conclusion that the primal cause of glacial 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 takes place by the minute individual movements of the grains upon one another. While they are in the spheroidal form, as in the névé, this would not seem to be at all difficult. They may rotate and slide over each other as the weight of the snow increases; but as they become interlocked by growth, both rotation and sliding must apparently encounter more resistance. The amount of rotary motion required of an individual granule is, however, surprisingly small, and the meltings and refreezings incident to shifting pressures and tensions, and to the growth of the granules, seem adequate to meet the requirements. 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, or one diameter in 10,000 days; in other words, it would change its relations to its neighbors to the extent of its diameter in about thirty years. A change of so great 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. The movement is supposed to be permitted 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 moisture to adjoining points, and its resolidification. The points of greatest compression315 are obviously just those whose yielding most promotes motion, and a successive yielding of the points that come in succession to oppose motion most (and thus to receive the greatest stresses) permits continuous motion. It is merely necessary to assume that the gravity of the accumulated mass is sufficient to produce the minute temporary liquefaction at the points 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.