=Flood-plain terraces due to other cause_Geology, Vol. 1 [of 3]

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 any 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).

[Illustration: +Fig.+ 196.—Terraces partly of rock and partly of

alluvium. Such terraces indicate successive uplifts, or some other

change which had a similar effect on the stream which made the

valley.]

=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 toward 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.

[Illustration: +Fig.+ 197.—Diagram to show why certain terraces are

longer lived than others.]

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.

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. These relations can rarely be seen

because of the destruction of the older terraces, and because of the

erosion by waves along shore.

[Illustration: +Fig.+ 198.—Diagram looking up the valley, showing two

terraces below, one in the middle, and none above. The relations are

purely diagrammatic.]

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.

CHAPTER IV.

THE WORK OF GROUND- (UNDERGROUND) WATER.

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

greater 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 stratification 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.

[Illustration: +Fig.+ 199.—Diagram illustrating the general point that

ground-water is not dependent entirely upon local supply.]

[Illustration: +Fig.+ 200.—Diagram illustrating the position of the

ground-water surface (the dotted line) in a region of undulatory

topography.]

=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

topography 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.

[Illustration: +Fig.+ 201.—Diagram to illustrate the relations of

ground-water to the surface.]

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 borings 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.

Depth at which Equivalent

Rate of Increase Temperature of 212° Pressure in

of Temperature. would be reached. Atmospheres.

1° for 50 feet 8,100 feet 238 (approximately)

1° for 60 „ 9,720 „ 285 „

1° for 70 „ 11,340 „ 333 „

With an initial temperature of 80°, corresponding to that of the warmer

parts of the earth’s surface, instead of 50°, the table would be as

follows:

1° for 50 feet 6,600 feet 194 (approximately)

1° for 60 „ 7,920 „ 214 „

1° for 70 „ 9,240 „ 272 „

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:

Rate of Increase Temperature at a

Initial Temperature. of Temperature. Depth of 6,800 Feet.

50° 1° for 50 feet 186° Fahr.

50° 1° for 60 „ 163° „

50° 1° for 70 „ 147° „

80° 1° for 50 „ 216° „

80° 1° for 60 „ 193° „

80° 1° for 70 „ 177° „

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:

Rate of Increase Depth of Critical

Initial Temperature. of Temperature. Temperature.

50° 1° for 50 feet 28,000 to 29,250 feet

50° 1° for 60 „ 33,600 to 35,100 „

50° 1° for 70 „ 39,200 to 40,950 „

80° 1° for 50 „ 26,500 to 27,750 „

80° 1° for 60 „ 31,800 to 33,300 „

80° 1° for 70 „ 37,100 to 38,850 „

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

sufficiently 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:

Initial Rate of Increase Temperature at

Temperature. of Temperature. Depth of Six Miles.

50° 1° for 50 feet 683° Fahr.

50° 1° for 60 „ 578° „

50° 1° for 70 „ 502° „

80° 1° for 50 „ 713° „

80° 1° for 60 „ 608° „

80° 1° for 70 „ 532° „

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 in 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 much 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 mineral 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 extracted

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.

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 that

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 flow, 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.

[Illustration: +Fig.+ 202.—Diagram to illustrate the general form

and relations of caves developed by solution. The black portions

represent the cavern spaces. Some limestone sinks are represented on

the surface above.]

RESULTS OF THE WORK OF GROUND-WATER.

[Illustration: +Fig.+ 203.—Ground-plan of Wyandotte Cave. The unshaded

areas represent the passageways. (21st Ann. Rept., Ind. Geol. Surv.)]

=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 is 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.

[Illustration: +Fig.+ 204.—Deposits in Wyandotte (Ind.) Cave. (Hains.)]

[Illustration: +Fig.+ 205.—Deposits in Wyandotte Cave. (Hains.)]

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.

[Illustration: +Fig.+ 206.—A limestone sink-hole, east-northeast of

Cambria, Wyo., exceptional for its steep sides. Minnekahta limestone.

(Darton, U. S. Geol. Surv.)]

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 site 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.)

[Illustration: +Fig.+ 207.—A fresh landslide near Medicine Lake, Mont.

The bare space shows the position from which the slide started.

(Whitney.)]

[Illustration: +Fig.+ 208.—Landslide topography. The protruding mass on

the right has slumped down from the mountain to the left. South face

of Landslip Mountain, Colo., seen from the west; Rico quadrangle.

(Cross, U. S. Geol. Surv.)]

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).

[Illustration: +Fig.+ 209.—Landslide topography on Badger Mountain,

Washington. The slumping material in this case is basalt.]

=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 bank 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.

[Illustration: PLATE XVII.

Fig. 1. HUNTERDON COUNTY, NEW JERSEY.

U. S. Geol. Surv.

Fig. 2. NEAR PIKEVILLE, TENNESSEE.

U. S. Geol. Surv.]

[Illustration: PLATE XVIII.

Fig. 1. WASHINGTON.

U. S. Geol. Surv.

Fig. 2. CALIFORNIA.

U. S. Geol. Surv.]

[Illustration: +Fig.+ 210.—A “talus glacier,” in Silver Basin, near

Silverton, Colo. (Cross, U. S. Geol. Surv.)]

[Illustration: +Fig.+ 211.—A huge mass of rock settling into the Canyon

of the Colorado. A result of gravity action. (Atwood.)]

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.

[Illustration: +Fig.+ 212.—Diagram illustrating positions, _a_ and _b_,

favorable for springs.]

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. 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

Yellowstone 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.

ANALYSES OF AMERICAN SPRING-WATERS.[109]

[+Reduced to Parts per 1000 by Dr H. J. Van Hoesen.+]

+------------------------------+--------------------+-------------------+--------------------+------------------+----------------------+--------------------+------------------+------------------+------------------+--------------------+

| | | | | | | | | | | |

| Date | | 1872 | Feb. 1876 | | | | | | | |

| | | | | | | | | | | |

| | | | | | | | | | | |

| | 189 | | | III, p. 618 | Ter., 1878, p. 393 | 1876, p. 195 | | | | p. 27 |

| Sodium, Na | .09227 | 4.72640 | 2.0398 | .45164 | .4615 | .10424 | .7743 | .3554 | .3040 | .6116 |

| | | | | | | | | | | |

| Potassium, K | .00919 | .35806 | .1285 | .05980 | ...... | Trace | .0669 | .0191 | .0094 | .0630 |

| | | | | | | | | | | |

| Calcium, Ca | .02136 | .94050 | 1.0739 | .44400 | .0344 } | | .0305 | .0367 | .0121 | .0589 |

| | | | | | } | .50600[111] | | | | |

| Magnesium, Mg | .01805 | .53470 | .2352 | .05860 | ...... } | | .0010 | .0034 | .0004 | .0604 |

| | | | | | | | | | | |

| Barium, Ba | ...... | .01848 | Trace | | | | | | | |

| | | | | | | | | | | |

| | | | | | | | | | | |

| Lithium, Li | Trace | .01078 | .0003 | .00039 | ...... | Trace | | | | |

| | | | | | | | | | | |

| Iron, Fe | Trace[113] | .00341 | .0027 | Trace[114] | ...... | Trace | | | | |

| | | | | | | | | | | |

| Manganese, Mn | ...... | ...... | .0009 | ...... | ...... | Trace | | | | |

| | | | | | | | | | | |

| Chlorine, Cl | .07465 | 7.47400 | 4.2730 | .24850 | .7496 | Trace[115] | .9697 | .2396 | .2070 | .2272 |

| | | | | | | | | | | |

| Bromine, Br | | .04661 | .0025 | | | | | | | |

| | | | | | | | | | | |

| | | | | | | | | | | |

| | | | | | | | | | | |

| Carbonic acid, CO₂ | .12160 | 5.82603 | .1792 | 1.11001 | ...... | .03516 | ...... | Trace | ...... | .5787 |

| | | | | | | | | | | |

| Sulphuric acid, H₂SO₄ | .03218 | .00234 | 2.0318 | .20696 | .0325 | .16140 | .3555 | .3901 | .3492 | .3131 |

| | | | | | | | | | | |

| Phosphoric acid, H₃PO₄ | Trace | .00005 | .0004 | ...... | ...... | Trace | | | | |

| | | | | | | | | | | |

| | | | | | | | | | | |

| Alumina, Al₂O₃ | ...... | .00770 | .0022 | ...... | ...... | Trace | .0010 | ...... | ...... | .0018 |

| | | | | | | | | | | |

| Silica, SiO₂ | .00940 | .01174 | .0080 | .02010 | .7680 | Trace | .2788 | .1136 | .1310 | .1220 |

| | | | | | | | | | | |

| Hydrogen in bicarbonates, H | ...... | .09713 | .0030 | | | | | | | |

| | | | | | | | | | | |

| | | | | | | | | | | |

| Oxygen, O | ...... | ...... | ...... | ...... | ...... | ...... | .0194[117] | .0255[117] | .0080[117] | .0325[117] |

| |--------------------+-------------------+--------------------+------------------+----------------------+--------------------+------------------+------------------+------------------+--------------------+

| | .37870 | 20.05910 | 9.9814 | 2.60000 | 2.0460 | .80680 | 2.4953 | 1.1834 | 1.0211 | 2.0692 |

| | | | | | | | | | | |

| Carbonic acid, CO₂ | ...... | 2.015[118] | ...... | ...... | ...... | In excess | | | Trace | |

| | | | | | | | | | | |

| Sulphuretted hydrogen, H₂S | ...... | ...... | ...... | ...... | ...... | 0.5000 | | | | |

[Illustration: +Fig.+ 213.—“Old Faithful” in eruption.]

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]

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).

[Illustration: +Fig.+ 214.—“Giant” Geyser, Yellowstone Park, in

eruption. Shows also the cone. (Wineman.)]

[Illustration: +Fig.+ 215.—Cone (or crater) of Castle Geyser,

Yellowstone Park. (Detroit Photo. Co.)]

[Illustration: +Fig.+ 216.—Cone (or crater) of Grotto Geyser,

Yellowstone Park. (Detroit Photo. Co.)]

[Illustration: +Fig.+ 217.—Cone of Giant Geyser, Yellowstone Park.

(Wineman.)]

The heating of geyser and hot-spring water must cool the lava or

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.

[Illustration: +Fig.+ 218.—Hot springs deposits. Terraces of Mammoth

Hot Springs, Yellowstone Park.]

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 under

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.

[Illustration: +Fig.+ 219.—Hot springs and geysers. Norris Geyser

basin, Yellowstone Park.]

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 stratum 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.

[Illustration: +Fig.+ 220.—Figure illustrating the principle of

artesian wells in drift.]

CHAPTER V.

THE WORK OF SNOW AND ICE.

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.

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.

[Illustration: +Fig.+ 221.—An alpine snow-field.]

[Illustration: +Fig.+ 222.—Map of Greenland. The borders only are free

from ice. (Stieler.)]

The only condition necessary for a snow-field is an excess of snowfall

over snow waste. The lower edge of a snow-field, the _snow-line_, is

dependent chiefly on temperature and snowfall. In general it does not

depart much from the summer isotherm of 32°, though it may be well

above this isotherm where the snowfall is light. That the snow-line

is not a function of temperature only is shown by its position in

various places. In the equatorial portion of the Andes, for example,

the snow-line has an altitude of about 16,000 feet on the east side of

the mountains, where the precipitation is heavier, and of about 18,500

feet on the west side, where it is lighter. For the same reason the

snow-line in the Himalayas is 3000 or 4000 feet lower on the _south_

side than on the north.

While in equatorial regions the snow-line has an altitude of 15,000 to

18,000 feet, it approaches or even reaches sea-level in the latitude

of Antarctica and North Greenland. In intermediate latitudes it has an

intermediate position.

While temperature and snowfall are the most important factors

controlling the position of the snow-line, both humidity and the

movements of the air are of some importance, since both affect the rate

of evaporation of snow and ice.

=The passage of snow into névé and ice.=—The snow does not lie on the

surface long before it undergoes obvious change. The light flakes

soon begin to be transformed into granules, and the snow becomes

“coarse-grained.” The granular character, so pronounced in the snow of

the last banks which remain in the spring in temperate latitudes, is

even more distinct in perennial snow-fields, either at the surface or

just beneath it. This granular snow is called _névé_, or _firn_. Still

deeper beneath the surface, where the thickness of the snow is great,

the névé becomes more compact and finally coherent, and grades into

porous ice. This gradation is accomplished at no great depth, though

the thicknesses of snow and névé are by no means constant.

=Structure of the ice.=—Ice formed beneath a snow-field is in some

sense stratified. It is made up of successive falls of snow which

tend to retain the form of layers. This follows from two or three

conditions. The snow of one season, or of one period of precipitation,

may have been considerably changed before the succeeding fall of snow.

So also the surface of the snow-field at the end of the melting season

is often covered with a visible amount of earthy matter, some of which

was blown up and dropped on the surface during the melting season, and

some of which was concentrated in that position by the melting of the

snow in which it was originally imbedded. The amount of earthy matter

is often sufficient to define snows of successive years, or perhaps of

minor periods of precipitation, and makes distinct the stratification

which would otherwise remain obscure. The snowfall of successive years

has been estimated by this means[122] where the snow is exposed in

crevices in the snow-field.

In addition to its rude stratification, the ice of the deeper portions

often acquires a stratiform structure which may perhaps best be called

_foliation_ to distinguish it from the _stratification_ which arises

from deposition. The foliation appears to result mainly from the

shearing of one part over another in the course of the movements to

which the ice is subjected, as will be illustrated presently.

=Texture.=—The ice derived from the snow is formed of interlocking

crystalline grains. The crystalline character is present from the

beginning, for it is assumed by the snowflakes when they form, and

the subsequent changes seem only to modify the original crystals

by building up some and destroying others. By the time the snow is

converted into névé, the granules have become coarse, and wherever the

ice derived from the névé has been examined, the granular crystalline

texture is present. The individual crystals in the ice are usually

larger than those of the névé, and more closely grown together. In the

fresh unexposed ice the crystals are so intimately interlocked that

they are not readily seen except under a polarizing microscope, but

when the ice has been honeycombed by partial melting, the granules

become partially separated and may be easily seen. Fig. 223 shows

quantities of them which have been washed down from the surface,

and disposed as cones at its base. While a given mass of snow in a

great snow- and ice-field cannot be followed consecutively through

its whole history, yet since (1) the granular texture is pronounced

in the névé stage where the granules show evidences of growth, and

since (2) the same texture is also pronounced in the last stages of

the ice when it is undergoing dissolution, as well as at all observed

intermediate stages, and since (3) the crystals are, on the average,

larger in proportion to the lateness of the stage of their history,

while (4) experiment has shown that granules grow under the conditions

which exist in snow-fields, and (5) that they persist under very

considerable pressure, it is legitimate to assume that a granular

crystalline condition persists throughout all stages, and is a feature

of progressive growth.

[Illustration: +Fig.+ 223.—Figure showing cones of granules of ice

which have been washed down the front of the glacier by streamlets,

and accumulated after the manner of talus or alluvial cones. North

Greenland.]

=Inauguration of movement.=—Eventually the increase in depth of

snow and ice in a snow-field gives rise to motion. The exact nature

of the motion has not yet been demonstrated to the satisfaction of

all investigators. Brittle and resistant as ice seems, it exhibits,

under proper conditions, some of the outward characteristics of a

plastic substance. Thus it may be made to change its form, and may

even be moulded into almost any desired shape if carefully subjected

to sufficient pressure, steadily applied through long intervals

of time.[123] These changes may be brought about without visible

fracture, and have been thought to point to a viscous condition of the

ice. There is much reason, however, as will be seen later, to question

this interpretation of the ultimate nature of the movement. Whatever

this may be, the _mass result_ of the movement in a field of ice is

comparable, in a superficial way at least, to that which would be

brought about if the ice were capable of moving like a viscous liquid,

the motion taking place with extreme slowness. This slow motion of ice

in an ice-field is _glacier motion_, and ice thus moving is _glacier

ice_.

[Illustration: +Fig.+ 224.—Ice-caps of small size. The figure also

shows some valley glaciers extending out from the main ice-sheet and

from the local ice-caps. A portion of the North Greenland coast,

north of Inglefield Gulf. Lat. about 78°. (Peary.)]

[Illustration: +Fig.+ 225.—Small ice-caps in the northwestern part of

Iceland. (Thoraddsen’s geological map of Iceland.)]

[Illustration: +Fig.+ 226.—A glacial lobe, midway between an ice-cap

and a valley glacier. A protrusion from a local ice-cap east of Cape

York, Greenland.]

If both the surface on which the ice-sheet develops and its

surroundings be essentially plane, as may happen in high latitudes,

and if the snow- and ice-field be symmetrical in shape, the outward

movement will be approximately equal in all directions, and the area

covered by the spreading ice-field will remain more or less circular.

If the ice-field rests on a steeply inclined surface, like a mountain

slope, the movement becomes one-sided in conformity to the slope. If

the surface, otherwise plane, be affected by valleys parallel to the

direction of movement, the ice in the valleys will be deeper than

that on the divides between them, and its movement stronger. In the

valleys, therefore, the ice will advance farther than elsewhere before

being melted, and the outline of the ice will become lobate, the lobes

occupying the depressions. These general relations are shown in Figs.

224 and 225. If the depressions be wide and shallow, the lobes will be

broad and short (Fig. 226); if the depressions be narrow and deep, the

lobes will be relatively narrow and long. If the snow and ice rest on

a surface consisting chiefly of steep valleys and sharp ridges, as is

common on mountain slopes, the snow and ice are chiefly gathered in the

valleys, and take a linear form.

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_.

[Illustration: +Fig.+ 227.—Characteristic end of a North Greenland

(Bryant) glacier.]

[Illustration: +Fig.+ 228.—The Rhône glacier. (Reid.)]

[Illustration: +Fig.+ 229.—Characteristic end of a North Greenland

glacier. North side of Herbert Island, Inglefield Gulf.]

[Illustration: +Fig.+ 230.—The end of an alpine (Forno, Switzerland)

glacier. (Reid.)]

[Illustration: +Fig.+ 231.—Deploying end of a North Greenland glacier.]

Where ice-caps are developed on plateaus whose borders are trenched

by valleys, ice-tongues from the edge of the ice-cap often extend

down into the valleys and give rise to one type of _valley glacier_

(Figs. 224 and 227). A second and more familiar type of valley glacier

occupies mountain valleys, and is the offspring of mountain snow-fields

(Fig. 228). The former are confined chiefly to high latitudes, and are

distinguished as _polar_ or _high-latitude_ glaciers (Figs. 227 and

229); the latter are known as _alpine_ glaciers (Figs. 228 and 230).

The end and side slopes of polar glaciers are, as a rule, much steeper

than those of alpine glaciers. When a valley glacier descends through

its valley to the plain beyond, its end deploys, forming a fan (Fig.

231). The deploying ends of adjacent glaciers sometimes merge, and the

resulting body of ice constitutes a _piedmont glacier_ (Fig. 232). At

the present time, piedmont glaciers are confined to high latitudes. In

some cases the snow-field that gives rise to a glacier is restricted

to a relatively small depression in the side of a mountain, or in the

escarpment of a plateau. In such cases the snow-field and glacier are

hardly distinguishable, and the latter descends but little below

the snow-line. In many cases it does not even enter the narrow valley

which leads out from the depression occupied by the snow-field. Such

a glacier is nestled in the face of a cliff, and may therefore be

called a _cliff glacier_[124] (Figs. 233 and 234). The snow-field of a

cliff glacier is sometimes no more than a great snowdrift, accumulated

through successive years. Cliff glaciers are often as wide as long,

and are always small, and between them and valley glaciers there are

all gradations (Fig. 235). Occasionally the end of a valley glacier,

or the edge of an ice-sheet reaches a precipitous cliff, and the end

or edge of the ice breaks off and accumulates like talus below. The

ice fragments may then again become a coherent mass by regelation, and

the whole may resume motion. Such a glacier is called a _reconstructed

glacier_. The precipitous cliffs of the Greenland coast furnish

illustrations.

[Illustration: +Fig.+ 232.—The Malaspina glacier, Alaska; the best

known example of a piedmont glacier. (Russell.)]

[Illustration: +Fig.+ 233.—A cliff glacier. North Greenland type. North

side of Herbert Island, Inglefield Gulf. The lower half of the white

area is snow, and snow talus. So also are the white patches to the

right. The height of the cliff is perhaps 2000 feet. The water in the

foreground is the sea.]

[Illustration: +Fig.+ 234.—Chancy glacier; a cliff glacier of the

Montana type. (Shepard.)]

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. In length they range from a fraction of a mile to many

miles, and though their width is usually much less than their length,

the reverse is sometimes the case (Figs. 233, 234, and 235). Their

thickness is usually measured by hundreds of feet rather than by

denominations of other orders, but the variation is great, and exact

measurements are almost wholly wanting. The minimum thickness is that

necessary to cause movement, and this varies with the slope, the

temperature, and other conditions. There is also much variation in the

thickness in different parts of a glacier. As a rule, it is thinnest in

its terminal portion, and thickest at some point intermediate between

this and its source, but nearer the latter than the former. Cliff and

reconstructed glaciers are comparable in size to the smaller valley

glaciers. Piedmont glaciers may attain greater size.

[Illustration: +Fig.+ 235.—A glacier in the Cascades near Cascade Pass,

Wash. A glacier intermediate between a cliff glacier and a valley

glacier. (Willis, U. S. Geol. Surv.)]

An ice-cap is theoretically thickest at its center and thins away to

its borders, but its actual dimensions are influenced by the topography

on which it is developed. The Greenland ice-cap is known to rise about

9000 feet above the sea, and it probably reaches considerably higher

than this in the unexplored center of its broad dome. The height of the

land surface beneath is unknown, but it is unlikely that it averages

half this amount, and hence the ice is probably 5000 feet or more thick

in the center. There is reason to think that it is much thicker in

Antarctica.

=Limits.=—The ice of a glacier is always moving forward (neglecting

temporary halts), but the _end_ of a glacier may be retreating,

advancing or remaining stationary, according as the rate of wastage is

greater, less, or just equal to the forward movement of the ice. The

position of the lower end of the glacier is therefore determined by the

ratio of movement to wastage. Its upper end is generally ill-defined.

In a superficial sense, it is the point where the ice emerges from the

snow-field; but the lower limit of the snow-field is often ill-defined,

and in any case is not the true upper limit of the glacier, since there

must be movement from the granular mass of ice beneath the snow to

make up for the waste below, and the moving ice beneath the snow-field

which feeds the tongue of ice in the valley is just as really a part

of the glacier as the more consolidated portion in the valley below.

If a definite upper limit for an alpine glacier is to be named, it

should probably be the _Bergschrund_, a gaping crevasse, or series of

crevasses which sometimes open near the precipitous slope of the peak

or cliff where the snow-field lies. The _Bergschrund_ is formed by the

moving of the lower part of the snow-field away from the portion above.

The lower end of a glacier is usually free from snow and névé in

summer, but, traced toward its source, it first becomes covered with

névé, then with snow, and finally merges into the snow-field without

having ceased to be a glacier. The term glacier is, however, commonly

used to mean merely the more solid portion outside (below) the névé.

=Movement.=—The fact of glacier movement is established in various

ways, the most obvious being by the advance of its lower end. Such

advance is too slow to be seen from day to day, and is only detected

when the lower end of the glacier overrides or overturns objects in

front of it, or moves out over ground previously unoccupied. But even

when the end of a glacier is not advancing, the movement of the ice may

be established by means of stakes or other marks set on the surface.

If the positions of these marks relative to fixed points on the sides

of the valley be determined, they are found after a time to have moved

down the valley. Rows of stakes or lines of stones set across a glacier

in the upper, middle and lower portions have revealed many facts

concerning the movement of the ice.

Generally speaking, the middle of a valley glacier moves more rapidly

than its sides (Fig. 236), but in some cases, especially in large

glaciers, there are found to be two or more main lines of movement,

with belts of lesser movement between. The top of a glacier moves, on

the whole, more rapidly than the bottom, though the observations made

do not show that the rate of movement diminishes regularly downward,

and it probably does not so diminish in many cases. In Switzerland,

where the glaciers have been studied more carefully than elsewhere,

the determined rates of movement range from one or two inches to four

feet or more per day. Some of the larger glaciers in other regions

move more rapidly, but it does not follow that large glaciers always

move faster than small ones. The Muir glacier of Alaska has been found

to move seven feet or more per day,[126] and some of the glaciers of

Greenland have been found to move, in the summer time, 50, 60, or even

more feet per day. A single estimate as high as 100 feet per day has

been made; but these high rates have been observed only where the ice

of a large inland area crowds down into a comparatively narrow fjord,

and debouches into the sea, and then only in the summer. In the case of

the glacier with the highest recorded rate of summer movement, 100 feet

per day, the advance was only 34 feet at about the same place in April.

[Illustration: +Fig.+ 236.—Diagram to show the rate of movement of the

Rhone Glacier at various points in its course at centre and sides. It

also shows the fluctuations in the positions of the end of the glacier

between 1874 and 1882, and the profile of the ice. (Heim.)]

The _average_ movement of the border of the inland ice of Greenland is

very small. Rink says that “between 62° and 68° 30′, the edge of the

inland ice is almost stationary for a remarkably long distance.”[127]

The observations of the authors between 77° and 78° were of like

import. Probably the average movement of the border of the Greenland

ice-cap is less than one foot a week.

=Conditions affecting rate of movement.=—The rate of glacier movement

appears to depend on (1) the depth of the moving ice; (2) the slope of

the surface over which it moves; (3) the slope of the upper surface

of the ice; (4) the topography of the bed over which it passes; (5)

the temperature; and (6) the amount of water which falls upon it or

is carried to it by the drainage of its surroundings, in addition to

that produced by the melting of the glacier itself. Great thickness,

a steep slope, much water, smoothness of bed, and a high (for ice)

temperature favor rapid movement. Since some of these conditions,

notably temperature and amount of water, vary with the season, the rate

of movement for any given glacier is not constant throughout the year.

Other conditions, especially the first of those mentioned above, vary

through longer periods of time, and occasion periodic variations in the

rate of movement.

Since the volume of ice concerned influences the rate of movement,

anything which changes the volume affects the rate. An excess of

snowfall with favorable conditions for its preservation for a period

of years, increases the volume of ice, and tends to accelerate its

movement. A deficiency in snowfall, or in its preservation, as from

high average temperature or from aridity, diminishes the quantity of

ice, and so retards the movement. An acceleration of velocity causes

the ice to move down the valley farther before being melted, that

is, causes the end of the glacier to advance, while a decrease of

velocity produces the opposite effect. As a matter of fact, the lower

ends of glaciers advance for a period of years and then retreat, to

advance again at a later time.[128] Observation has shown that the

periods of advance follow a succession of years when the snowfall

has been heavy and the temperature low, while the periods of retreat

follow a succession of years when the snowfall has been light and the

temperature above the average. The periods of advance and retreat lag

behind the periods of heavy and light snowfall respectively, by some

years, and a long glacier responds less promptly than a short one.

Present knowledge seems to point to a period of 35 to 40 years as the

time within which a cycle of fluctuation, that is, an advance and a

retreat, takes place.

A declining upper surface is essential to glacier motion. There are

short stretches where this is not the case, and indeed there are

particular places where the upper surface slopes backward.[129] This

may occur where the ice is pushed up over a swell in its bed, or is

crowded up against any considerable obstacle; but such cases are no

more than local exceptions, and do not militate against the truth of

the general statement that the upper surface of a glacier declines in

the direction of motion. A declining _lower_ surface is less necessary.

In the case of valley glaciers, the bed does, as a rule, decline in

the direction of motion, but that there are local exceptions is shown

by the deep basins in rock which such glaciers often leave behind them

when they retreat. In the great continental glaciers of recent geologic

times, the ice frequently moved up slopes for scores, and even hundreds

of miles; but in all such cases, the upper surface must have declined

in the direction of movement. With a given thickness of ice, the

greater the decline of its lower surface in the direction of motion,

the more rapid its progress. A rough bed, or a crooked course retards

the motion of a glacier, while a smooth bottom and a straight course

facilitate it.

Slope, roughness of bed, and volume affect the movement of glaciers

somewhat as they affect the movement of rivers. The temperature of

the water, on the other hand, has little effect on the flow of a

river so long as it remains unfrozen; but the effect of temperature

on the motion of ice is most important. In many cases, indeed, the

temperature, together with the water that is incidental to it, seems

to be the chief factor in determining the rate of movement. The way in

which its effects are felt will be discussed later.

=Likenesses and unlikenesses of glaciers and rivers.=—Many of the

characteristics of a valley glacier may be understood from the study of

the accompanying figure (Fig. 237) of the White (Alaska) glacier. From

this figure it will be seen that the glacier is an elongate river-like

body, following the curves of the valley in stream-like fashion. It

has its origin in the snows collected on the mountain heights seen

in the distance, and it works its way down the valley in a manner

which, in the aggregate, is similar to the movement of a stiff liquid.

The likeness to a river extends to many details. Not only does the

center move faster than the sides, and the upper part faster than the

bottom, as in the case of streams, but the movement is more rapid in

constricted portions of the valley and slower in the broader parts.

These and other likenesses, some of which are apparent rather than

real, have given origin to the view that glacier ice moves like a stiff

viscous liquid.

[Illustration: +Fig.+ 237.—White glacier (central background) joining a

larger glacier (foreground), Alaska. (Reid.)]

But while the points of likeness between glaciers and rivers are

several, their differences are at least equally numerous and

significant. The trains of débris on the surface (the dark bands in the

illustration), like the central currents of streams, pass nearer the

projecting points of the valley walls and farther from the receding

bends; but beyond this point the analogy fails, for the trains of

débris on the ice do not conform in detail to the courses of the

currents of a winding stream, nor is there evidence of the rotatory

motion that characterizes river water. Furthermore, the glacier is

readily fractured, as the numerous gaping crevices on many glaciers

show. The crevasses are sometimes longitudinal, sometimes transverse,

and sometimes oblique. In the case of Arctic glaciers, longitudinal

crevassing is especially conspicuous.

[Illustration: +Fig.+ 238.—Cracking of glacier due to change in grade

of bed. A North Greenland glacier overriding a mound of

moraine-stuff.]

_Crevasses_ appear to be developed wherever there is appreciable

tension, and the causes of this tension are many. An obvious cause is

an abrupt increase of gradient in the bed (Fig. 238). If the change

of gradient be considerable, an ice-fall or cascade results, and the

ice may be greatly riven (Fig. 228). Below the cascade, the surface

may bristle with wedges and pinnacles of ice (séracs, Fig. 239).

Transverse crevices at the margin sometimes appear to be the result

of the tension developed on a curve. Oblique crevices on the surface

near the sides are commonly ascribed to the tension between the

faster-moving center and the slower-moving margins, and in like manner

crevasses that rise obliquely from the bottoms are attributed to the

tension between the faster-moving portions above and the slower-moving

portions below. All these crevasses indicate strains to which a liquid,

whose pressures are equal in all directions, does not offer a close

analogy. Longitudinal crevasses may affect both the river-like part

of a glacier and its deploying end, and are the result of tension

developed by movement within the ice itself, to which, again, rivers

offer no analogy. Somewhat similar cracks develop in the outer crust

of asphalt, when a mass of it is allowed to stand and spread; but in

this case there is evaporation of the volatile ingredients, giving to

the outer part relative rigidity and brittleness, while the inner part

remains more fluent. The analogy is therefore not perfect and probably

not really illustrative. The crevices may be narrow or wide, and both

narrow and wide may be found in the same glacier. The narrow crevices

that never open much are the most significant, as they show that very

little stretching is needed to satisfy the tension. The opening of a

gaping crevice is sometimes the work of weeks, and, in the slow-moving

glaciers of high latitudes, sometimes the work of successive seasons.

All this shows that the glacier is a very brittle body, incapable of

resisting even very moderate strains brought to bear upon it very

slowly. Had the ice even moderate ductility, it would adapt itself to

tension brought to bear upon it so slowly as are many of the tensions

which produce crevassing. In its behavior under tension therefore a

glacier is notably unlike a river.

[Illustration: +Fig.+ 239.—Séracs of glacier. (Reid.)]

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 middle. There may be several medial

moraines on one glacier, in which case some of them may be far from

the center. In alpine glaciers, the _surface terminal_ moraine is less

well-defined; in polar glaciers it often connects two lateral moraines,

making a loop roughly concentric with the terminus of the glacier.

[Illustration: +Fig.+ 240.—End of Mount Dana glacier, Cal. Shows

irregularities of surface due to crevassing farther up the glacier,

and to unequal melting.]

[Illustration: +Fig.+ 241.—Shows irregularities due to unequal melting

of veined ice. End of small glacier south of Forno hut, Engadine,

Switzerland. (Reid.)]

Besides the surface moraines, which represent belted aggregations of

débris, there may be scattered bowlders and bits of rock of various

sizes on the ice, and, in addition to the coarse material, there is

often some dust which has been blown upon the ice.

=Relief due to surface débris.=—The débris on the ice affects its

topography by influencing the melting of the subjacent and adjacent

ice. The rock débris absorbs heat more readily than the ice. A small

and thin piece of stone lying on the ice is warmed through by the

sun’s rays, and, melting the ice beneath, sinks, just as a piece of

black cloth on snow will sink because of the increased melting beneath

it. Though a good absorber of heat, rock is a poor conductor, and so

the lower surface of a large mass of stone is not notably warmed. The

ice beneath it is protected from the direct rays of the sun, and is

therefore melted more slowly than that around it. The result is that

the bowlder presently stands on a protuberance of ice (Fig. 244). When

its pedestal becomes high, the oblique rays of the sun and the warm air

surrounding it cause it to waste away, and the capping bowlder falls.

In high latitudes, the great obliquity of the rays sometimes allows

them to strike under isolated bowlders. In this case, they are warmed

from below, and thus aid rather than hinder the melting of the ice.

[Illustration: +Fig.+ 242.—Irregular surface due to uneven bottom.

Bowdoin glacier, Inglefield Gulf, North Greenland. The dark patches

near the left margin of the glacier are lakelets in basins produced by

the upward bending of the ice as it overrides an elevation in its bed.

The figure also shows a depressed medial moraine.]

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.

[Illustration: +Fig.+ 243.—A Swiss glacier, showing surface moraines,

characteristic profile, etc.]

=Dust-wells.=—The wind-blown dust sometimes gives rise to peculiar

topographic features of small size. The dust is not distributed by the

wind with absolute equality, and the surface drainage of the ice tends

to aggregate it. Every dust particle acts like a small stone, and where

aggregations of dust occur, they melt their way down into the ice,

developing holes or “dust-wells” (Fig. 245). These wells rarely reach a

depth of more than a few inches, but they may be so numerous that the

pedestrian is obliged to watch his steps. This is especially true near

the edge of the large ice-caps. It is evident that the depth of these

dust wells must be slight, for so soon as they are deep enough to cut

off the sun’s rays from the dust at the bottom, the deepening ceases.

Other things being equal, they are deeper in low latitudes than in high.

[Illustration: +Fig.+ 244.—Bowlder on ice pinnacle. Forno glacier,

Switzerland. (Reid.)]

[Illustration: +Fig.+ 245.—Dust-wells. Igloodahomyne glacier, North

Greenland.]

[Illustration: +Fig.+ 246.—Disposition of débris in ice. North

Greenland glacier. (Libbey.)]

[Illustration: +Fig.+ 247.—Profile of the lower part of the lateral

margin of a glacier. Southeast side of McCormick Bay, North

Greenland.]

=Débris below the surface.=—The lower part of a glacier, as well as

the upper, carries rock débris. This débris is sometimes so abundant,

especially near the ends and edges of the ice, that it is difficult to

locate the bottom of the glacier; for between the moving ice which is

full of débris, and the stationary débris which is full of ice, there

seems to be a nearly complete gradation. The débris in the lower part

of arctic glaciers, and to some extent of others, is often disposed in

thin sheets sandwiched in between layers of clean ice. These débris

sheets are often numerous and usually discontinuous, though groups

of such sheets often persist for considerable distances. Débris also

occurs to some extent in the ice well above its base. It is sometimes

in belts, as seen in section, and sometimes in bunches. These various

relations are illustrated by Figs. 227, 229, and 246–249.

Another characteristic of the basal débris-laden part of some glaciers

is the _foliation of the ice_ (Figs. 248, 249, etc.). This is

especially well shown in the arctic glaciers, the ends and sides of

which have steep or vertical faces. The foliation is best developed

in the débris zone, though often shown above. The foliation is

sometimes minute, consisting of layers of clean ice, an inch or less

in thickness, separated by mere films of earthy matter. In extreme

cases there are a score or more of laminæ within a foot. Locally, and

especially where débris is abundant, the laminæ are much contorted.

This is seen both in section (Figs. 248 and 249) and on the surface

(Fig. 250).

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.

[Illustration: +Fig.+ 248.—Side view of end of glacier. Southeast side

of McCormick Bay, North Greenland. Shows structure of ice as well as

position of débris.]

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.

[Illustration: +Fig.+ 249.—Side view of a North Greenland glacier (East

glacier), showing position of débris and structure of the ice.]

=The winter wave.=—There are but few observations on the internal

temperatures of glaciers during the winter season, but it seems certain

that the winter wave diminishes rapidly downward and dies out below,

much as does the winter wave which affects land surfaces not covered

with ice. Conduction alone considered, the temperature of the ice where

the cold wave dies out, should correspond, approximately, to the mean

annual temperature of the region, provided that temperature is below

the melting point of ice.

Assuming that in the high altitudes and high latitudes where glaciers

abound, the temperature of the surface may average about -12° Fahr.

(about -25° C.) for the winter half of the year, which is about the

case for north Greenland, Spitzbergen, and Franz Josef Land, and that

the conductivity of the ice in the C. G. S.[130] system is .005,

the temperature would be lowered appreciably only about 40 feet

below the surface at the close of the winter period, conduction only

being considered. How far the internal temperature may be influenced

by air forced through the ice by winds and by variations of the

barometer is not known and cannot well be estimated. The wave of low

temperature descending from the surface in winter would probably become

inappreciable before reaching a depth of 60 feet. At this depth the

temperature should be about 15° Fahr.—the mean annual temperature of

the region.

[Illustration: +Fig.+ 250.—Contorted lamination shown at the surface. A

small glacier south of Forno hut, Engadine, Switzerland. (Reid.)]

=The summer wave.=—The warm wave follows the analogy of the summer wave

of ice-free land surfaces much less closely. This is because of the

low melting temperature of ice as compared with other forms of solid

earth-matter. On this account the summer wave is bi-fold. The one part

travels downward by conduction, the other by the descent of water; the

one has to do primarily with the temperature before the melting-point

of ice is reached; the other, with the temperature after that point

is reached; the first conforms measurably to the warm wave affecting

other solid earth-matter, while the second is governed by special

laws. After the surface portion of the ice is brought to the melting

temperature, the additional heat which it receives melts the ice and

is transformed from _sensible_ into _potential_ heat. Ice charged with

water is _potentially_, but not _sensibly_, warmer than ice which has

just reached the melting temperature.

The warm wave of conduction dies out below like the cold wave. The

warm wave descending by the flow of water stops where the freezing

temperature of water is reached. In regions where the average

temperature is below freezing, the water-wave does not descend so far

as the wave of conduction, since the latter descends below the zone

where the melting temperature is found.

The foregoing considerations warrant the generalization that glaciers

normally consist of two zones (1) an outer or upper zone of fluctuating

temperature, and (2) an under zone of nearly constant temperature. The

under zone obviously does not exist where the thickness of the ice is

less than the thickness of the zone of fluctuating temperature. This

may be the case in very thin glaciers in very cold regions, and in the

thin ends and edges of all glaciers.

=The temperature of the bottom.=—The internal heat of the earth is

slowly conducted to the base of a glacier where it melts the ice at

the estimated average rate of about one-fourth of an inch per year.

The temperature of melting is a little below 32° Fahr. since pressure

lowers the melting-point at the rate of .0133° Fahr. (.0075° C.) for

one atmosphere of pressure. At the bottom of a mile of ice therefore

the melting temperature is about 30.2° Fahr. (-1° C.) It is probable

that in all thick glaciers the temperature of the bottom is constantly

maintained at the melting-point. This may be indicated by the streams

which issue from beneath glaciers during the winter, though this

criterion is hardly decisive since the issuing waters may be derived

partly or wholly from the rock beneath. In glaciers or in parts of

glaciers so thin as to lie wholly within the zone of fluctuating

temperature, the temperature of the bottom is obviously not constant.

=Temperature of the interior of the ice.=—The variation of temperature

of the surface of a glacier has already been shown to lie between a

maximum of 32° Fahr. and the minimum temperature of the region where

the glacier occurs. Lower, in the zone of fluctuating temperature, the

variation is less, and where the zone of fluctuating temperature passes

into the zone of constant temperature, variation ceases. The thickness

of the zone of fluctuating temperature varies with the temperature of

the region where the glacier occurs, being greatest where the winters

are coldest. In the case of all glaciers except thin ones in very cold

regions, the temperatures within the zone of constant temperature range

from the mean annual temperature of the region at the top of the zone

(provided this is not above the melting-point of ice at this depth) to

the melting temperature of the ice at the bottom. Within these limits

the range may be great or slight.

If we consider only the effects of the external seasonal temperatures

and the internal heat of the earth, it appears that all the ice in

the zone of constant temperature in the lower end of a typical alpine

glacier should have a constant melting temperature, for the average

temperature of regions where the ends of such glaciers occur is usually

above 32° Fahr., and this determines a temperature of 32° Fahr. (or a

little less) at the top of the zone, while a melting temperature is

maintained at the bottom by the earth’s interior heat. In thin glaciers

of very cold regions, where the zone of constant temperature has

relatively slight thickness, the low temperature descending from the

surface may so far overcome the effect of internal heat as to keep the

bottom of the ice at a freezing temperature. In all other cases, the

ice at the bottom of the under zone has a melting temperature, while

that above is probably colder.

In the higher altitudes and in the polar latitudes where glaciers

are chiefly generated, the mean annual temperature of the surface is

usually below the melting-point of ice. Here the temperature of the

ice between the top and bottom of the zone of constant temperature

must, on the average, be below the melting-point, unless heat enough

is generated in the interior of the ice to offset the effect of the

temperature above. For example, where the mean annual temperature is

20° Fahr. or lower, as in middle Greenland, Spitzbergen, and Franz

Josef Land, and at certain high altitudes in more southerly latitudes,

the mean temperature in the zone of constant temperature should range

from 20° Fahr. at the top to 32° Fahr. (or a little less) below;

i.e., it should average about 6° below the melting-point. Under these

conditions, all the ice in the zone of constant temperature, except

that at its bottom, must be permanently below the melting-point, but it

is perhaps worthy of especial note that much of it is but little below.

In alpine glaciers the part of the ice affected by this constant low

temperature (below freezing) is presumed to be chiefly that which lies

beneath the snow-fields. In polar glaciers the low temperature probably

prevails beneath the surface, not only throughout the great ice-caps,

but also in the marginal glaciers which descend from them.

From these theoretical considerations we may deduce the generalization

that in the zone of constant temperature _within the area of glacial

growth_, the temperature of the ice is generally _below_ the

melting-point, while _within the area of wastage_, the temperature of

the corresponding zone is generally _at_ the melting-point.

=Compression and friction as causes of heat.=—The foregoing conclusions

are somewhat modified by dynamic sources of heat. The compression

arising from gravity, and the friction developed where there is motion,

are causes of heat. Since friction occurs only when motion takes place,

the heat which it generates is secondary and may, for present purposes,

be neglected. Compression not only lowers the melting-point slightly,

but it _produces heat at the point of compression_. Where the ice is

granular, the compression, due to weight, takes place at the contacts

of the grains. At intermediate points the pressure tends to cause

them to bulge, and this has the effect of lowering the temperature

of the bulging points. If therefore the compression be considerable,

the granules may be warmed to the melting-point where they press each

other, while at other points their temperature may be lower. In this

case melting will take place at the points of compression, and the

moisture so produced will be transferred to the adjacent parts of the

granule and immediately refrozen. Melting at the points of compression

would result in some yielding of the mass, and in some shifting of

the pressure to new points where compression and melting would again

take place. Thus the melting, the refreezing, and the attendant

movement might go on until the limits of the power of gravity in this

direction were reached. From considerations already adduced, it appears

that the temperature in some parts of every considerable body of ice

must be such as to permit these changes. The heat due to depression

and friction may modify the theoretical deductions drawn above from

atmospheric and internal influences.

=Summary.=—If the foregoing generalizations be correct, (1) the

surface of a glacier is likely to be melted during the summer, (2) its

immediate bottom is slowly melting all the time (unless the thickness

of the ice be less than the thickness of the zone of annual variation

or of permanent freezing temperature); (3) its subsurface portion _in

the zone of waste_ is generally melting, owing to descending water,

compression, and friction; while (4) its subsurface portion _in the

zone of growth_ is probably below the melting-point except as locally

brought to that temperature by compression, friction, and descending

water, and at the bottom by conduction from the rock beneath.

=Movement under low temperature.=—Glacier motion will not be discussed

at this point, but one of the bearings of the preceding conclusions

on glacier motion may be pointed out. Since there must be motion

in the area of growth to supply the loss in the area of waste, the

_fundamental_ cause of motion must be operative in bodies of ice the

mean temperature of which is below the melting-point, _unless the

dynamic sources of heat are considerable_. This fundamental cause does

not exclude the _coöperation_ of causes that work only (1) at the

melting temperature, or (2) where the ice is bathed with water, or (3)

in the plane of contact between wet ice above and dry ice below. These

may be _auxiliary_ causes which abet the fundamental one in producing

the more rapid movement of warm seasons, or in bringing about the

especially rapid movement in situations where there is abundant water,

or in inducing the shearing which is such a remarkable feature of

arctic glaciers.

=Evaporation.=—The ice wastes by evaporation as well as by melting, and

while the former process is far less important than the latter, its

results are probably larger than is commonly apprehended. One of the

most remarkable features of some of the deposits of ancient glaciers is

the slight evidence they afford of escaping waters. The most plausible

explanation seems to lie in the supposition that the ice was largely

wasted by evaporation. This conclusion finds support in many places

in the presence of a mantle of fine silt over the drift, the silt

being apparently composed of dust blown upon the ice. It is supposed

to imply aridity in the region about the ice. If a sufficient mantle

of dust were spread over the border zone of the ice, and if the air

were very dry, nearly all the water melted on the surface of the ice

might be held back by the dust-wells until the water was evaporated or

absorbed.

[Illustration: +Fig.+ 251.—Spouting stream. Glacier south side of

Olriks Bay, North Greenland.]

=Drainage.=—Some of the water produced by surface melting forms little

streams on the ice. Sooner or later they plunge into crevasses or over

the sides and ends of the glacier. In the former case, they may melt

or wear out well-like passages (moulins) in the ice, and even in the

rock beneath. Much of the surface water sinks into the ice. Its ready

penetration is aided by the “dust-wells” which mark the surface of

many glaciers. In north Greenland wells which contain six or eight

inches of water at the end of a warm day are often dry in the morning.

The water has leaked out and passed to lower levels. From these and

other harmonious observations it is inferred that the superficial

part of a glacier at least is readily penetrated by water. The depth

to which surface water penetrates is undetermined, but it doubtless

varies greatly, not only in different glaciers, but in different parts

of the same glacier, and in the same part at different times. Above

the line of perennial snow there is little water either from melting

or from rain, and hence relatively slight penetration. Below the line

of perennial snow there is much melting and much rain, and here it is

probable that the water sometimes, perhaps usually, penetrates to the

bottom of the ice during the melting season, even independently of

crevasses.

Once within the glacier, the course of the water is variable.

Exceptionally it follows definite englacial channels, as shown by

springs or streams issuing from the ice at some point above its bottom

(Fig. 251). Oftener it descends or moves forward through the irregular

openings which the accidents of motion have developed. If it reaches a

level where the temperature is below its freezing-point, it congeals.

Otherwise it remains in cavities or descends to the bottom. The water

produced by melting within the glacier probably follows a similar

course. So far as these waters descend to the bottom, they join those

produced by basal melting and issue from the glacier with them. In

alpine glaciers the waters beneath the ice often unite in a common

stream in the axis of the valley, and hollow out a tunnel. Thus the

Rhone is already a considerable stream where it issues from beneath the

Rhone glacier. In the glaciers of high latitudes, subglacial tunnels

are less common and the drainage is in streams along the sides of the

glaciers or through the débris beneath and about them.

At the end of the glacier, all waters, whether they have been

superglacial, englacial or subglacial, unite to bear away the silt,

sand, gravel, and even small bowlders set free from the ice, and

to spread them in belts along the border of the ice or in trains

stretching down the valleys below. These are the most common of the

glacio-fluvial deposits.

THE WORK 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.

=Getting load.=—If the surface on which the snow-field which is to

become a glacier accumulates be rough and covered with abundant rock

débris, as such surfaces usually are, the glacier already has a basal

load when its movement begins, for the snow covers, surrounds, and

includes such loose blocks of rock as project above the general surface

and envelops all projecting points of rock within its field. When the

snow becomes ice and the ice begins to move, it carries forward the

loose rock already imbedded in it, and tears off the weak points of

the enveloped rock-projections. It may perhaps also move some of the

soil and mantle rock of the original surface to which it is frozen. In

addition to the _subglacial_ load which the glacier thus has at the

outset, there may be a surface load which has fallen on the snow or ice

from cliffs above. This is especially true of mountain-valley glaciers.

If this has been buried by snow and ice it is _englacial_; if it lies

on the surface it is _superglacial_.

Once in movement, the ice carries away the débris to which it was

originally attached, and at the same time gathers new load from the

same area. The new load is acquired partly by the rasping effect of the

rock-shod ice on its bed, and partly by its rending power which, under

favorable circumstances, may quarry out considerable blocks of rock.

This “plucking” process is at its best where the ice passes over cliffs

of jointed rock or steep-sloped hills.

As the ice advances into new territory it acquires additional basal

load, partly by rasping, partly by plucking, partly by freezing to

it, and partly in other ways. One of these ways may be illustrated by

the sequence of events when the end of a glacier advances on a very

large bowlder. As the ice approaches it, the reflection of heat from

it melts the adjacent edge of the ice, making a slight reëntrant. With

farther advance, the ice closes in against and around the bowlder, and

finally carries it along in the bottom of the moving mass. In some

cases, especially when its advance is rapid, the ice may push débris in

front of itself. Even where this is the case, the amount of material

pushed forward is generally slight, partly because the extreme edge of

the ice often fails to rest on the land in summer, when the movement

is greatest, being melted from below by the heat of the surface over

which it is spreading (see Fig. 252), and partly because the earth in

front of the glacier is frozen during a large part of the year. In this

condition, the earthy matter has greater resistance than the ice,

and the latter rides over it. Superglacial material may be acquired

during movement by the fall of débris from cliffs, or by the descent of

avalanches.

[Illustration: +Fig.+ 252.—Diagram showing lack of contact of the edge

of the ice with its bed.]

=Conditions influencing rate of erosion.=—An obstructive attitude

of the surface toward the movement of the ice is as necessary for

effective erosion as the movement of the ice itself. Advancing over a

flat surface, ice ordinarily inflicts but little wear, since there is

little for it to get hold of. So slight is the abrasive power of ice

under these conditions that it frequently overrides and buries the soil

with more or less of its herbaceous vegetation. But while a certain

measure of roughness of surface is favorable for glacial erosion, the

topography may be so uneven as to seriously impede the ice. Erosion

is probably at its maximum, so far as influenced by topography, when

the roughness of the surface is such as to offer notable catchment for

the basal ice, but not such as to impede its motion very seriously.

The amount of relief favorable for the greatest erosion increases with

increasing thickness of the ice.

Other conditions which influence erosion by ice are (1) the amount of

loose or slightly attached débris on the surface, (2) the resistance of

the rock, (3) the slope of the surface, (4) the thickness of the ice,

(5) the rate of movement, and (6) the abundance and character of the

débris which the ice has to work with. The effect of the first five of

these conditions is evident. The effect of the last is less simple.

Clean ice passing over a smooth surface of solid rock has little effect

upon it; but a rock-shod glacier will abrade the same surface notably.

The effect of this abrasion is shown in the grooves and scratches

(_striæ_) which the stones in the bottom of the ice inflict on the

surface of the rock over which they pass (Figs. 253, 255, and 256).

At the same time the stones in the ice are themselves worn both by

abrasion with the bottom, and with one another (Fig. 254). It does not

follow, however, that the more material in the bottom of the ice the

greater the erosion it effects; for with increase of débris there may

be decrease of motion[131] and, beyond a certain point, the decrease

of motion seriously interferes with the efficiency of erosion. When any

considerable thickness of ice at the bottom of the glacier is full of

débris, this loaded basal portion may approach stagnancy, and the lower

limit of considerable movement may lie between the loaded ice below

and the relatively clean ice above. A moderate, but not an excessive

load of débris is, therefore, favorable for great erosion. Something

depends, too, on the character of the load. Coarse, hard, and angular

débris is a more effective instrument of erosion than fine, soft, or

rounded material. The adverse influence of the overloading of the ice

on its motion has been likened to the stiffening of a viscous liquid by

the addition of foreign matter, but it may better perhaps be referred

to the destruction of the granular-crystalline continuity on which

glacier motion probably depends.

[Illustration: +Fig.+ 253.—Glacial striæ and bruises. The block to the

right shows two sets of striæ: that to the left shows the peculiar

curved fractures known as _Chatter Marks_.]

[Illustration: +Fig.+ 254.—Bowlders showing glacial striation. (Drawn

by Miss Matz.)]

[Illustration: +Fig.+ 255.—Striæ on bed rock, Kingston, Des Moines Co.,

Ia. (Iowa Geol. Surv.)]

From the preceding statement, it is evident that erosion is not equally

effective at all points beneath a glacier. So far as concerns the ice

itself, erosion is not most effective at the end of a valley glacier,

or at the edge of an ice sheet, for here the strength of movement is

too slight and the load too great; nor is the most effective erosion

at the source or near it, for though the ice may here be thick, the

movement is slow and the load likely to be slight. Ice conditions only

being considered, erosion is most effective somewhere between the

source and the terminus, and probably much nearer the latter than the

former. The conditions of the surface over which the ice passes may

be such as to vary the place of greatest erosion widely. Thus in an

Alpine glacier, erosion may be most effective at the _Bergschrund_

because the slope here favors “plucking.” Here, notable amphitheatres

(_cirques_) are sometimes excavated. After the glacier disappears,

the bottom of the cirque is often seen to contain rock basins (Fig.

257). Glacial cirques abound in mountains where glaciers once existed,

but from which they have now disappeared. The cirques of the Bighorn

mountains of Wyoming (Pl. XIX) are examples.

[Illustration: +Fig.+ 256.—Striæ, grooves, etc., in a canyon tributary

to Big Cottonwood Canyon, Wasatch Mountains. (Church.)]

=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.

=Flood-plain terraces due to other causes.=—Certain other causes,