over them than strata in any other position. In Fig. 103 the stream
which has the profile ad crosses bedding-planes at _b_ and _c_. In
Fig. 104, where the beds dip up-stream, many more division-planes
are crossed in the same distance. Since bedding-planes are planes of
weakness, it follows that horizontal and nearly horizontal strata are
not, under ordinary conditions of erosion, in a position favorable for
most rapid wear. When strata are horizontal, it makes no difference
which way the stream runs, for the current sustains the same relation
to the cleavage-planes whatever its course.
In the case of incoherent material the position of the beds, or even
their existence, has little influence on the rate of erosion. Such
formations are weak in all directions, not simply along bedding-planes.
[Illustration: +Fig.+ 105.—Diagram to illustrate the various relations
a stream may sustain to the outcrop of vertical layers of rock.]
[Illustration: +Fig.+ 106.—Diagram to illustrate the various relations
a stream may sustain to the outcrops of inclined layers of rock.]
When the strata are vertical, three distinct cases may arise (Fig.
105). The stream may flow (1) with the strike (_aa_); (2) at right
angles to the strike (_bb_); or (3) oblique to it (_cc_) at any angle
whatsoever. It is perhaps not possible to say which of these positions
is most favorable for erosion, for the character of the rock, the
thickness of its layers, its ability to stand with steep slopes, and
the strength of the currents concerned, would influence the result. A
stream which flows at right angles to the strike (_bb_, Fig. 105) would
cross more cleavage-planes in a given distance than a stream flowing in
any other direction, and would strike the outcropping edges of layers
at the angle of greatest advantage. A stream flowing along the strike
(_aa_), on the other hand, has better opportunity to sink its channel
on cleavage-planes, and the current oblique to the strike (_cc_), has
some of the advantages of each of the others.
When the strata are inclined five cases may arise. (1) The stream
may be parallel to the strike (_aa_, Fig. 106), when it makes no
difference which way the current flows; it may be at right angles to
the strike (_bb′_), and (2) flowing with the dip (toward _b′_), or
(3) against it (toward _b_); it may be oblique to the strike, and
flowing (4) in the general direction of dip (toward _c′_); or (5) in
the opposite direction (toward _c_). As before, the stream flowing at
right angles to the strike would cross the largest number of layers
in a given distance, and so have an opportunity to take advantage of
more cleavage-planes than a stream in any other position. But in the
case of inclined strata a new element enters into the problem. When the
stream flows parallel to the strike, the valley which is in process of
deepening is not sunk vertically, but is shifted more or less in the
direction of the dip (Fig. 107). This is called _monoclinal shifting_.
The result is that there is a constant tendency to undermine (sap) the
valley bluff on the down-dip side, and this process of sapping will,
according to its rate, accelerate the growth of the valley, especially
in width. Monoclinal shifting is favored by the presence of a hard
layer (_H_), as shown in Fig. 107, if this stratum is the bed of the
stream.
[Illustration: Fig. 107.—Diagram to illustrate monoclinal shifting. The
valley _abc_, as seen in cross-section, becomes _deb_, as the stream
lowers its channel.]
In the second and third cases mentioned above, the only difference is
in the angle at which the current strikes the outcropping edges of
layers and laminæ. The mechanical advantage is with the stream which
flows with the dip. In the fourth and fifth cases something will
depend on the angle which the stream’s course makes with the strike.
In all these cases, as in those where the strata are vertical, much
will depend on the thickness and resistance of the layers and on the
strength of the currents concerned.
_The Influence of Climate._
Climate has both a direct and an indirect effect on erosion. Its
direct influence is through precipitation, evaporation, changes of
temperature, and wind; its indirect, through vegetation. Like declivity
and rock structure, climate does not affect all elements of erosion
equally.
The chief elements of climate are temperature, moisture, and
atmospheric movements; the principal factors which influence it are
latitude, altitude, distance from the sea, direction of prevailing
winds, and topographic relations.
The effects of variations in temperature on rock weathering have
already been discussed (p. 43). They are chiefly mechanical, and are
seen at their best where the daily range is great.
High temperature favors chemical action, and the weathering of rock
by decomposition is at its best in the presence of abundant moisture
in regions where the temperature is uniformly high. Furthermore, a
warm moist climate favors the growth of vegetation, the decay of which
supplies the water with organic acids which greatly increase its
solvent power. The climatic conditions favoring mechanical weathering
are therefore different from those favoring chemical weathering. High
temperature and abundant moisture and vegetation are found in many
tropical regions, and here the rock is often decomposed to greater
depths, on the whole, than in high latitudes. How far this is the
result of rapid weathering, and how far of slow removal, due in part
to the protective influence of the plants, cannot be affirmed. If the
weathered material is not removed, it will presently become a mantle
thick enough to retard the processes which brought it into existence.
So long as the water of the surface and that in the soil remains
unfrozen, temperature affects neither corrasion nor transportation.
But in middle and high latitudes the surface is frozen for some part
of each year. During this time corrasion is at a minimum, for although
the streams continue to flow there is relatively little water running
over the surface outside the drainage channels, and that little is
relatively ineffective. Under some conditions, therefore, temperature
affects both corrasion and transportation.
The humidity of the atmosphere has an influence even more important
than that of temperature on the rate of erosion, and its influence
is exerted on each of the elements of that complex process. A moist
atmosphere favors oxidation, carbonation, hydration, and the growth of
vegetation, all of which promote certain phases of rock weathering.
On the other hand, humidity tends to prevent sudden and considerable
variations in temperature, thus checking the weathering effected
by this means. Precipitation, the most important single factor in
determining the rate of erosion, is dependent on atmospheric humidity.
Its amount, its kind (rain or snow), and its distribution in time, are
the elements which determine its effectiveness in any given place.
Other things being equal the greater the amount of precipitation
the more rapid the corrasion and transportation. Much, however,
depends on its distribution in time. A given amount of rainfall may
be distributed equally through the year, or it may fall during a
wet season only. The maximum inequality of distribution would occur
if all the rainfall of a given period were concentrated in a single
shower. With such concentration the volume of water flowing off over
the surface immediately after the down-pour would be greater than
under any other conditions of precipitation, and since velocity is
increased with volume, and erosive power with velocity, it follows
that the erosive power of a given amount of water would be greater
under these circumstances than under any other. Furthermore, a larger
proportion of the precipitation would run off over the surface under
these circumstances than under any other, for less of it would sink
beneath the surface and less would be evaporated. If erosive power
and rate of erosion were equal terms, this would therefore be the
condition for greatest erosion; but erosive power and rate of erosion
do not always correspond. If the water falling in this way could
get hold of all the material it could carry, extreme concentration
of precipitation would be the condition favorable for most rapid
erosion. But if the amount of available material for transportation
is slight, a large part of the force of the water could not be
utilized in erosion. It follows that if there were a large amount of
disintegrated material on the surface, erosion would be greater the
greater the concentration of precipitation. If, on the other hand,
there were but little disintegrated material on the surface, frequent
showers, with intervening periods when conditions were favorable for
weathering, that is, for preparing material for transportation, might
be more favorable for rapid erosion. While the total energy of running
water available for erosion under these conditions would be less than
before, there might in the long run be more material for transport;
for weathering in the presence of moisture, and all that goes with
it, might be more effective in preparing material for transportation,
than weathering during the long periods of drought which would occur
if the precipitation were concentrated to its maximum. Temperature
favoring, the uniform distribution of moisture through the year would
allow the growth of vegetation, which, although favoring some processes
of weathering, retards erosion in general. While therefore it is not
possible to say what distribution of rainfall favors most rapid erosion
without knowing the nature of the surface on which it is to fall,
enough has been said to show that the problem is by no means a simple
one. Some of the most striking phases of topography developed by
erosion, such as those of the Bad Lands (Figs. 75 to 78, and 108), are
developed where the rainfall is unequally distributed in time, and too
slight or too infrequent to support abundant vegetation.
[Illustration: +Fig.+ 108.—Bad-land topography developed under
conditions of aridity and unequal distribution of rainfall. Slope of
Pinal Mountains, Ariz. (Ransome, U. S. Geol. Surv.)]
During its fall, and immediately after, rain is more effective than
an equal amount of snow; but the snow may be accumulated through
a considerable period of the year, and then melted rapidly, when
it has an effect comparable to that which would be produced by the
concentration of the rainfall into a limited period of the year. If
the ground beneath be frozen when the snow melts (and this is often
the case) the erosion accomplished by the resulting water will be
diminished.
Except in dry regions, where wind-work sometimes exceeds water-work,
the movements of the atmosphere are of less importance directly
than precipitation in determining the rate of erosion. But even in
regions which are not arid the winds have much to do with the rate of
evaporation and the distribution of rainfall, so that their indirect
effect is great. Even their direct effects in moist climates are not to
be lost sight of, for even here the surface is sometimes dry enough to
yield dust and sand, and the uprooting of trees so disturbs the surface
as to make earthy débris more accessible to wind and water. Where trees
gain precarious footholds on steep slopes, as they often do, they are
likely to be overturned as soon as they are large enough to offer
considerable resistance to the wind, and in the overturning, large
quantities of rock are sometimes loosened and carried down the slope
by gravity. This phase of destructive work is seen at its best on the
walls of gorges, where trees often flourish until their tops project
above the rim of the valley.
Through vegetation, climate influences erosion in ways which are
easily defined qualitatively, but not quantitatively. Both by its
growth (wedge-work of roots) and by its decay (supplying CO₂, etc.,
to descending waters) it favors certain phases of weathering; but, on
the other hand, it retards corrasion and transportation both by wind
and water. This is well shown along the banks of streams and on the
faces of cliffs, in clay, sand, etc. Its aggregate effect is probably
unfavorable to erosion by mechanical means, and favorable to that by
chemical processes.
[Illustration: +Fig.+ 109.—Characteristic cliffs of high arid regions.
Right wall of Snake River canyon, nearly opposite the mouth of Salmon
River, Id. Two spring-formed coves, with “Castle Rock” between.
(Russell, U. S. Geol. Surv.)]
Erosion in high arid regions differs from that in regions of abundant
rainfall in several ways. It is obvious that the valleys will develop
more slowly in the former, that they will remain young longer, that the
period necessary for the dissection of the surface is greater, that
the watercourses will be less numerous, and that fewer of them will
have permanent streams. There are certain other differences which are
less obvious. If the arid region be high and composed of heterogeneous
strata, the topography which erosion develops is more angular (Fig.
83) than that of the humid region. This is because there is less
rock decay, and less vegetation to hold the products of decay. The
more resistant beds of rock therefore come into greater prominence,
especially on slopes, where they develop cliffs (Figs. 109 and 110).
These general principles find abundant illustration in the plateaus of
the western part of the United States,[48] where the cliffs are by no
means confined to the immediate valleys of the streams (Fig. 1, Pl.
XII).
[Illustration: +Fig.+ 110.—A Butte. A characteristic feature of the
arid plateau region of the West. (Dutton, Mono. II, U. S. Geol.
Surv.)]
EFFECTS OF UNEQUAL HARDNESS.
In the preceding pages incidental reference has been made to the
results of inequalities of rock resistance. This topic will now be
considered more fully.
=Rapids and falls.=—Returning for a moment to the hypothetical island
with which our study of erosion began, let a horizontal layer of hard
rock be assumed to run through it (H, Fig. 111). As the rain falls
on the land and runs off over it, wear will be less rapid where the
hard layer comes to the surface than at the higher or lower levels. As
a result, the slope will become steeper at and below the outcrop of
the hard layer, and less steep immediately above it, as shown by _ab_
in Fig. 111. Under these conditions the water passing over the hard
ledge constitutes _rapids_. The increased erosion which accompanies
the increased velocity makes the rapids more rapid. The process may
continue until the water _falls_, rather than _flows_ over the hard
layer (_cd_, Fig. 111). With continued rainfall the edges of the hard
layer, together with the slopes above and below, would continue to
recede toward the center of the island. Under conditions of absolute
homogeneity of material, save for the hard layer specified, no valley
would be developed, and therefore no stream.
If the surface was so changed as to allow of the development of a
valley (p. 63) the same principles would be applicable. As an active
stream passes from a hard layer to one less resistant, the greater
wear on the latter gives origin to rapids. At first the rapids would
be slight (_a_, Fig. 112), but would become more considerable (_b_) as
time and erosion go on. When the bed of the rapids becomes sufficiently
steep, the rapids become _falls_[49] (_cd_). When the water falls
rather than flows over the rock surface below the hard layer, erosion
assumes a new phase. The hard layer is then undermined, and the
undermining causes the falls to recede. This phase of erosion is
sometimes called _sapping_.
[Illustration: +Fig.+ 111.—Diagram representing a horizontal layer of
hard rock in an island, and its effects on erosion.]
[Illustration: +Fig.+ 112.—Diagram illustrating the development of a
fall where the hard layer dips gently up-stream.]
[Illustration: +Fig.+ 113.—Diagram illustrating the conditions which
exist at Niagara Falls. (Gilbert.)]
If the hard layer which occasions a fall dips up-stream (Fig. 112), its
outcrop in the stream’s bed becomes lower as the fall recedes (_e_).
When it has become so low that the water passing over it no longer
reacts effectively against the less resistant material beneath (_f_),
sapping ceases, and the point of greatest erosion may be shifted from
the soft material beneath the fall to the hard layer itself. The actual
rate of erosion at this point may be no greater than before, though
the relative rate is. Under these circumstances the vertical edge of
the hard layer will presently be converted into an incline (_f_), and
as this takes place the fall becomes rapids. The conversion of the
falls into the rapids begins about the time the lower edge of the hard
stratum in the channel reaches grade. By continuation of the process
which transformed the falls into rapids, the rapids become less rapid,
and when the upper edge of the hard layer has been brought to grade,
the rapids disappear (_h_, Fig. 112). The history of rapids which
succeed falls is the reverse of that which preceded. The later rapids
are steepest at the beginning of their history, the earlier at their
end. Stated in other terms, rapids are steepest when nearest falls
in time. Slight differences in hardness in successive layers often
occasion successive falls or rapids (Fig. 114).
If the hard layer which occasions the falls be horizontal, instead of
dipping up-stream, the general result would be the same; but, other
things being equal, the duration of the falls developed under these
conditions would be greater, since they must recede farther before
becoming rapids.
If the layers of unequal hardness in a stream’s bed be vertical and the
course of the stream at right angles to the strike, rapids, and perhaps
falls, will develop (Fig. 115). The chances for falls are greater,
the greater the difference in hardness. Falls developed under these
conditions, as well as the rapids preceding and following, would remain
constant in position until the resistant layer was brought to grade,
but they would ultimately disappear as in the preceding cases. Falls
are not likely to develop where the strata of the stream’s bed dip
down-stream, though they may develop even under these conditions if the
gradient of the stream is greater than the dip of the strata (Fig. 116).
[Illustration: +Fig.+ 114.—Falls in Utica shale, Canajoharie, N. Y.
(Darton, U. S. Geol. Surv.)]
The inequality of resistance in the rock which occasions a fall may
be original or secondary. In the case of Niagara Falls[50] (Fig.
113) relatively resistant limestone overlies relatively weak shale.
At the Falls of St. Anthony (Minneapolis) limestone overlies friable
sandstone. The falls of the Yellowstone and the Shoshone Falls of
the Snake River (Idaho), are in igneous rock. In the former case the
unequal resistance is occasioned by unequal decay of the rock, due
perhaps to the rise of hot vapors which have decomposed the rock along
the lines of their ascent; in the latter, a more resistant sort of
igneous rock overlies a less resistant.
Structural features, such as jointing, sometimes give rise to falls,
or determine their distinctive features (Fig. 117), even where the
formations involved are of uniform hardness. A joint plane has the
effect of a weak vertical or highly inclined bed. If an open joint
is discovered in a stream’s bed, the water enters it. If it finds an
outlet below, a channel is worn along the new line of flow, with rapids
or falls where the water descends. Rock originally homogeneous may be
much fractured in some parts, while it remains unbroken in others.
Where a stream passes from the solid to the broken portion rapids, or
even falls, may develop.
[Illustration: +Fig.+ 115.—Diagram illustrating the development of
falls over a vertical hard layer.]
[Illustration: +Fig.+ 116.—Diagram illustrating the possibility of
falls where the beds dip down-stream.]
Falls may originate in still other ways. If for any reason a stream is
forced out of its valley, it may in its flow find entrance to another
valley, or to another part of its own valley, over a steep slope. If
the structure of the slope favors, a fall may speedily develop. The
Falls of St. Anthony are an example, the Mississippi having been turned
out of its earlier course by deposits of glacial drift. Again, if an
obstruction of any sort, such as a flow of lava, dams a stream, rapids
or falls are developed where the water overflows the dam. When a main
valley is notably deepened by glaciation the drainage from tributary
valleys may fall into it, if the tributaries were not equally deepened.
Falls which originated in this way are common in the western mountains
of the United States, as well as in most mountain regions recently
affected by local glaciers (Fig. 118).
One waterfall often breeds others. Thus where a fall recedes beyond
the mouth of a tributary stream, the tributary falls. The Falls of
Minnehaha, on a small tributary to the Mississippi, near Minneapolis,
may serve as an illustration. In such cases the falls may not develop
from rapids. Once in existence, the fall of a tributary follows the
same history as that of a main stream.
[Illustration: +Fig.+ 117.—Kepler’s Cascade, in the Yellowstone Park.
The jointed and fractured character of the igneous rocks occasions a
series of falls and rapids. (Iddings, U. S. Geol. Surv.)]
Streams which have falls are relatively clear.[51] If a stream
favorably situated for the development of a fall carried a heavy
load, deposition would take place below the rapids, and the tendency
would be to aggrade the channel at that point and so to prevent
the development of the fall. Falls occur only on streams which have
relatively high gradients. This means that the streams which have falls
are well above base-level, and streams well above base-level are young.
Falls therefore are a mark of topographic youth.
[Illustration: +Fig.+ 118.—The Upper Yosemite Falls.]
The fall of the Niagara[52] (Pl. IX) is one of the most remarkable
known, both because of its large volume of water and its great descent,
between 160 and 170 feet. The rate at which the fall is receding is a
matter of interest not only in itself, but because, once determined,
it may be made to serve as a unit of measurement for certain important
events in geological history. It was formerly conjectured that this
fall was receding at the rate of one to three feet per century, but
it was not until recent years that its actual rate of recession was
approximately fixed. By surveys executed in 1842 and 1890 it has been
determined that its average rate of recession between those dates was
something like 4½ feet per year, or about 150 times as great as the
highest estimate stated above. It is to be noted that this is the
_average_ rate of recession, for all parts of the ledge over which
the water falls are not receding at the same rate. The point of the
“Horseshoe” has, during the same time, gone back at more than twice
this rate.[53]
[Illustration: +Fig.+ 119.—A group of pot-holes. (Turner, U. S. Geol.
Surv.)]
Rapids and falls sometimes occasion the development of _pot-holes_
(Fig. 119), a peculiar rather than important erosion feature. The holes
are excavated in part by the falling and eddying of silt-charged water,
but chiefly by stones which the eddies move. Pot-holes which are not
now in immediate association with rapids or falls often point to the
former existence of rapids or falls.
=Rock terraces.=—The tendency to sapping shown in many waterfalls is
also shown in the weathering and erosion of the sides of a valley where
a hard layer outcrops above the bottom, and the profile of the side
slopes of the valley simulates that of the stream; that is, the slope
becomes gentle just above the hard layer, and steep, or even vertical,
at and below its outcrop. This is illustrated by Fig. 120, where the
hard layer through which the stream has sunk its valley stands out as
a rock terrace on either side of the valley. Such terraces are not
rare and are popularly believed to be old “water-lines”; that is,
to represent the height at which the water once stood. In one sense
this interpretation is correct, since a river has stood at all levels
between that of the surface in which its valley started, and its
present channel, but the shelf of hard rock does not mean that the
river, after attaining its present channel, was ever so large as to
fill the valley to the level of the terrace. Rock terraces may also
result from changes of level.
[Illustration: +Fig.+ 120.—Diagram to illustrate the development of
rock terraces.]
=Narrows.=—Inequalities in hardness occasion another peculiarity common
to valleys. If a stream crosses vertical or highly inclined strata of
unequal hardness, its valley is usually constricted at the crossing
of the harder layers. If such a constriction be notable it is called
a _narrows_, or sometimes a _water-gap_ (Figs. 121, 159, and Fig. 2,
Pl. XII). The Appalachian Mountains afford numerous examples. The
constriction arises because the processes which widen the valley are
less effective on the hard layer than on the less resistant ones on
either hand. Though most narrows are due to the superior resistance
of the rock where they occur, they are sometimes the result of other
causes.
[Illustration: +Fig.+ 121.—Lower narrows of the Baraboo River, Wis. The
even-crested ridge is Huronian quartzite. The surroundings are of
Cambrian sandstone. (Atwood.)]
[Illustration: +Fig.+ 122.—A hog-back, Jura-Trias. Colorado City, Colo.
(Russell, U. S. Geol. Surv.)]
Narrows are much more conspicuous in certain stages of erosion than
in others. While a valley is still so young as to be narrow at all
points, no narrows will be conspicuous; but at a later stage in
its history, when the valley is otherwise wide, narrows are more
pronounced. At a still later stage, when the hard strata themselves
approach base-level, the narrows again become inconspicuous.
From what has preceded it is clear that rapids or falls are likely to
occur at narrows, especially in the early part of their history.
=Other effects on topography.=—Inequalities in the hardness of rock
develop certain peculiarities of topography other than those of
valleys. The less resistant portions of a land area more or less
distant from streams are worn down more readily than those which are
more resistant. If great areas of high land be capped with hard rock
they are likely to remain as plateaus after surrounding areas of
less resistance are brought low. If the hard capping affects a small
area instead of a large one, the elevation is a butte, a hill, or a
mountain, instead of a plateau (Fig. 110). Many buttes and small mesas
are but remnants of former plateaus (Mesa Lauriano, N. M., Fig. 1, Pl.
XII). A feature of buttes and mesas capped by hard rock is the steep
slope or cliff corresponding to the edge of the hard bed (Figs. 78 and
109).
[Illustration: +Fig.+ 123.—A ridge due to the outcropping edge of hard
Jurassic rock. Wyoming.]
If the rock of a region be stratified and the layers tilted, the
removal of the softer beds leaves the harder ones projecting above
the general level in the form of ridges or “hog-backs” (Figs. 122
and 123). Dikes of igneous rock, harder than the beds which they
intersect, likewise become ridges after the degradation of their
surroundings. The plugs of old volcanic vents and other igneous
intrusions of limited area often constitute conspicuous hills or
mountains after erosion has removed their less resistant surroundings
(Fig. 124). Inequalities of hardness are therefore responsible for many
hills and ridges. In the isolation of the hills and ridges picturesque
coves are developed, where the attitude and distribution of the weak
and strong rocks are propitious. The bottoms of the coves are located
on the weak rocks, and above them rise the precipitous slopes of
the resistant ones. Round valley (Fig. 1, Pl. XVII, High Bridge, N.
J., quadrangle, U. S. Geol. Surv.) and the coves about the head of
Hiawassee River (Dahlonega, Ga., quadrangle) are examples.
[Illustration: +Fig.+ 124.—Matteo tepee, Wyo. Mass of igneous rock
exposed by erosion, and preserved because of its superior resistance.
(Detroit Photo. Co.)]
Ridges and hills resulting from the unequal degradation of unequally
resistant terranes are not equally prominent at all stages in an
erosion cycle. In early youth the material surrounding the hard bodies
of rock has not been removed; in early maturity considerable portions
of their surroundings still remain about them; but in late maturity
or early old age the outcropping masses of hard rock have been more
perfectly isolated and are most conspicuous. Most of the even-crested
ridges of the Appalachian system, as well as many others which might be
mentioned, became ridges in this way. In the final stages of an erosion
cycle the ridges of hard rock are themselves brought low. Isolated
remnants of hard rock which remain distinctly above their surroundings
in the late stages of an erosion cycle (Fig. 124) are known as
_Monadnocks_, the name being derived from Mount Monadnock, N. H., an
elevation of this sort developed in a cycle antedating the present.
[Illustration: +Figs.+ 125–27.—Diagrams illustrating piracy, where the
stream which does not flow over rock of superior hardness captures
those which do. Fig. 126 represents a further development of the
drainage shown in Fig. 125, and Fig. 127 represents a still later
stage.]
[Illustration: +Figs.+ 128–30.—Diagrams to illustrate piracy, where
the competing streams all cross a hard layer. The diagrams represent
successive stages of development.]
=Adjustment of streams to rock structures.=—Valleys (gullies) locate
themselves at the outset without immediate regard to the hardness and
softness of their beds. It is primarily the slope about the head of a
gully which determines its line of growth, though relative hardness
often determines the details of slope, even in the early stages of an
erosion cycle. Once established, streams tend to hold their courses,
even if this involves the crossing of resistant layers.
While a region where more and less resistant layers of rock come to the
surface is in a youthful stage of erosion, some of the valleys (and
therefore the streams) are likely to be located on the less resistant
rock, some on the more resistant, and some partly on the one and partly
on the other. The streams on the weaker rock will deepen their valleys
more rapidly than the others, and those which flow across stronger
and weaker rocks alternately will deepen their valleys more rapidly
than those which run on hard rock all the time. The former conclusion
is self-evident. The latter appears from the fact that rapids will be
likely to develop at the crossing of each hard layer, thus accelerating
erosion at those points. Such a stream therefore not only has less hard
rock to erode than one which flows on resistant rock all the time, but
it erodes that which it does cross much faster.
[Illustration: +Figs.+ 131, 132.—The capture of the head of Beaverdam
Creek by the Shenandoah Va.-W. Va. (After Willis.)]
Streams which do not cross hard layers therefore have an advantage
over those which do, and the tributaries to such streams, since they
join deeper mains, have an advantage over the tributaries to the
others. The valleys of the former may lengthen until their heads reach
the latter, and capture their streams. This sequence of events is
illustrated in the accompanying diagrams (Figs. 125–27). Even where
several streams cross the same resistant bed, piracy is likely to take
place among them, for some are sure to deepen their valleys faster than
others, because of inequalities of volume, load, or hardness. This is
illustrated by Figs. 128–30. An actual case is shown in Figs. 131,
132. Though piracy may take place when streams do not flow over rock
of unequal hardness (p. 103), it is much more common where unequal
resistance of the rock puts one stream at a disadvantage as compared
with another.
The changes in the courses of streams, by means of which they come to
sustain definite and stable relations to the rock structure beneath,
are known as processes of _adjustment_.[54] Since streams and valleys
adjust themselves to other conditions as well, this phase of adjustment
may be called _structural adjustment_. Structural adjustment is not
uncommon among rivers flowing over strata which are vertical or highly
inclined, since in these positions the hard and soft strata are most
likely to come to the surface in frequent alternation. The smaller
streams suffer capture and adjustment first, since, as a rule, they
have shallower valleys. It often happens that main streams, because of
their deeper valleys, hold courses not in adjustment with structure
(the Delaware, the Susquehanna, etc.), while tributary streams are
captured, diverted, and adjusted. The capture of a tributary, however,
leads both to the diminution of its main and to the increase of its
captor, and the weakened stream may ultimately fall a prey to the one
which is strengthened.
The processes of adjustment go on until the streams flow as much as
possible on the weaker beds, and as little as possible on the stronger,
when adjustment is complete. This amounts to the same thing as saying
that the outcrops of the hard layers tend to become divides. In many
cases an area is so situated that there is no escape for its drainage
except across resistant rock. In this case its drainage is completely
adjusted when as few streams as possible cross the resistant rock, and
these by the shortest routes.
Adjustment has been carried to a high degree of perfection in most
parts of the Appalachian system. Here, as in all other mountains of
similar structure, strata of unequal hardness were folded into ridges.
In this case, the folds have been truncated by erosion, exposing
the more and the less resistant beds (_H_ and _S_ respectively) in
alternate belts along the flanks of the truncated folds (_ab_ and
_cd_, Fig. 133). The streams, especially the lesser ones, now flow
along the strike of the softer beds much more commonly than elsewhere,
and where they cross the hard layers it is usually at right angles to
the strike. This is shown in Fig. 134, where the arrows indicate the
direction of strike. In the history of these rivers, however, a factor
is involved which has not yet been considered, and these streams will
be referred to later.
[Illustration: +Fig.+ 133.—Diagram showing the outcrops of hard layers
on the flanks of a truncated fold. _cd_ represents the present
surface; dotted lines above, earlier surfaces.]
[Illustration: +Fig.+ 134.—Example of adjusted drainage in a region of
folded rocks, Va.-W. Va.]
[Illustration: +Fig.+ 135.—Diagram to illustrate readjustment of
drainage, as base-level is approached.]
[Illustration: +Fig.+ 136.—Diagram to illustrate superimposition. The
consequent stream on the upper formation is superimposed on the
underlying structures when the upper bed has been cut through.]
As base-level is approached, the outcrops of hard rock are brought
low. When they have been reduced to the level of their surroundings,
the streams may flow without regard to the resistance of the rock
beneath, for downward cutting has ceased. As this stage of erosion is
approached, a readjustment of the drainage may take place, and the
waters which had taken long and circuitous courses to avoid hard rock,
may change their courses to more direct ones (compare Figs. 130 and
135). Adjustment is, therefore, a relative term, and streams which
are adjusted at one stage of erosion, are not necessarily adjusted at
another.
It sometimes happens that rocks of unequal resistance are covered
by beds of uniform hardness. A consequent stream developed on the
latter may find itself out of structural adjustment when it has cut
its channel down to the level of the heterogeneous beds below. Such
a stream is said to be _superimposed_ (Fig. 136) on the underlying
structure. Structural adjustment is likely to follow.
INFLUENCE OF JOINTS AND FOLDS.
=Joints.=—Various structural features of rock other than hardness
influence its erosion. Apart from the stratification planes, most rock
formations are affected by joints or fissures. The joints are often,
but not always, nearly vertical. Two sets are generally present, and
sometimes more. If but two, they usually meet at a large angle; if more
than two, two are likely to be nearly perpendicular to each other,
while the third and fourth sets have such directions as to cut the
others at large angles. These joints allow the ingress of water, roots,
etc., which help to weather and disrupt rocks. Occasionally there is
notable sag of the beds of rock along joint planes, but this effect
is usually superficial only (Fig. 137). Where the jointage planes are
frequent and open, the columns bounded by them sometimes topple over on
cliff faces, either by undercutting, or by the wedge-work of roots or
ice.
The effect of joints on erosion may often be seen along a stream
which flows in a rock gorge. In such situations, the outlines of the
banks are sometimes angular, and sometimes crenate (Fig. 138), the
reëntrants being located at the joints. By working into and widening
joints, running water sometimes isolates masses of rock as islands
(Fig. 139). In a region free from mantle rock, or where the mantle rock
is meagre, joints often determine the courses of valleys by directing
the course of surface drainage. This is shown in many parts of the
arid west. In regions where the rocks are notably faulted, the courses
of the streams are sometimes controlled by the courses of the fault
planes. This is the case, for example, in central Washington.[55]
[Illustration: +Fig.+ 137.—Shows the sagging of beds along joints. The
disturbance does not extend far below the surface. Cook’s quarry
(Niagara limestone) near La Salle, Niagara Co., N. Y. (Gilbert, U. S.
Geol. Surv.)]
The jointing of rocks often shows itself distinctly in the weathered
faces of cliffs (Figs. 140 and 141), especially in arid and semi-arid
regions, or where the slope is too steep for the accumulation of soil
and rock-waste on its surface.
If a stream flowing over jointed rock has falls, the conditions are
sometimes afforded for the development of an exceptional and striking
scenic feature. If above Niagara Falls, for example, there were an
open joint in the bed of the stream (as at _b_, Fig. 142), some portion
of the water would descend through it. After reaching a lower level
it might find or make a passage through the rock to the river below
the falls. If even a little water took such a course, the flow would
enlarge its channel, making a passageway between the joint through
which the water descended and the valley below the falls (_bcde_, Fig.
142). This passageway might become large enough to accommodate all the
water of the river. In this case, the entire fall would be transferred
from the position which it previously occupied (_f_) to the position of
the enlarged joint (_b_). The fall would then recede. The underground
channel between the old falls and the new would be bridged by rock
(_bf″_ and _f‴_, Fig. 143), making a _natural bridge_. The natural
bridge near Lexington, Va. (Fig. 144), almost 200 feet above the stream
which flows beneath it, is believed to have been developed in this
way. A similar bridge is now in process of development in Two Medicine
River in northwestern Montana (Fig. 145). Once in existence, a natural
bridge will slowly weather away.
[Illustration: +Fig.+ 138.—Figure showing crenate river bank, the
reëntrants being determined by joints. Dells of the Wisconsin River,
near Kilbourn, Wis. (Atwood.)]
[Illustration: +Fig.+ 139.—Lone Rock. An island isolated by the notable
widening of a series of joints. The joints in the rock of the island
have themselves been so widened that a rowboat may be taken through
it in two directions. Lower Dells of the Wisconsin. (Meyers.)]
[Illustration: +Fig.+ 140.—Effect of columnar structure on weathering.
Material unconsolidated. Spur of south end of Sheep Mountain.
(Lippincott, U. S. Geol. Surv.)]
It is not to be understood that all natural bridges have had this
history. They are sometimes developed from underground caves when parts
of their roofs are destroyed, as well as in various other ways.
[Illustration: +Fig.+ 141.—Effect of columnar structure on weathering.
Big Bad Lands, S. D. (Darton, U. S. Geol. Surv.)]
=Folds.=—The erosion of folded strata (anticlines and synclines) leads
to the development of distinctive topographic features. So soon as a
fold begins to be lifted, it is, by reason of its position, subject
to more rapid erosion than its surroundings. For the same reason the
crest of the fold is likely to be degraded more rapidly than its lower
slopes, and must suffer more degradation before it is brought to
base-level. Folds are usually composed of beds of unequal resistance,
and as the degradation of a fold proceeds, successive layers are worn
from the top, and the alternating hard and soft layers composing it are
exposed. So soon as this is accomplished, adjustment of the streams
is likely to begin, and the watercourses, and later the valley plains,
come to be located on the outcrops of the less resistant layers, while
the outcrops of the harder beds become ridges.
If the axis of an eroded anticline were horizontal, a given hard layer,
the arch of which has been cut off, would, after erosion, outcrop on
both sides of the axis. When the topography was mature these outcrops
would constitute parallel ridges, or parallel lines of hills; when the
region had been base-leveled, the outcrops would be in parallel belts,
though no longer ridges or hills. The lower the plane of truncation,
the farther apart would the outcrops be in the anticline, and the
nearer together in the syncline (compare _ab_ and _cd_, Fig. 133).
[Illustration: +Fig.+ 142.—A natural bridge in process of development;
longitudinal section at the left; transverse section, looking toward
_e_, at the right.]
If, on the other hand, the axis of the anticline or syncline to be
eroded was not horizontal, that is, if it _plunged_, the topographic
result would be somewhat different. Suppose a plunging anticline to
be truncated at base-level. If either end of the fold plunged below
the plane of truncation, the outcrops of a given layer on opposite
sides of the axis would converge in the direction of plunge, and come
together at the end. At a stage of erosion antedating planation (say
late maturity) there would have been a ridge, or a succession of hills,
in the position corresponding to the outcrop of a hard layer, with a
canoe-shaped valley within. If two hard layers were involved, instead
of one, there would be two encircling ridges, with a curved valley
between them, and a canoe-shaped valley within the innermost (Fig.
146). If the anticline plunged both ways, the valley enclosed by
the hard-layer ridge would be canoe-shaped at both ends (Fig. 147).
In such a case there would be likely to be a low gap (water-gap) in
the rim of the valley through which the drainage which degraded the
surface escaped, but there would be likely to be but one, for if
two or more streams had drained the area of the valley at an early
stage of erosion, one would be likely to have captured the others
(see p. 138) before late maturity. A succession of doubly-plunging
anticlines and synclines might give rise to a very complex series of
ridges and valleys. Illustrations of the above phenomena are found at
various points in the Appalachian Mountains, especially in eastern
Pennsylvania.[56]
[Illustration: +Fig.+ 143.—The same as Fig. 142 at a later stage of
development.]
[Illustration: +Fig.+ 144.—The Natural Bridge of Virginia, from the
southeast (Walcott, U. S. Geol. Surv.)]
[Illustration: +Fig.+ 145.—A natural bridge in development. Two
Medicine River, Mont. Corresponds to the stage represented by Fig.
142, and the view corresponds to that shown diagrammatically at the
right-hand end of the figure. (Whitney.)]
In the structural adjustment which goes with the erosion of
folds, it often happens that the valleys come to be located on the
anticlines, while the outcrops of the hard layers on the flanks of the
anticlines, or even in the original synclines, become the mountains.
The adjustments by which valleys come to be located on anticlines
are somewhat as follows:[57] Fig. 148 represents two doubly-plunging
anticlines with a syncline between, the relative elevations being shown
by contour lines. At the outset, the drainage of such a region must
have followed the structural valley, and its initial course, consequent
on the slope, must have been down the axial trough. Drainage from the
anticlines into the synclines would have promptly developed valleys,
and the valleys would soon have acquired streams.
[Illustration: +Fig.+ 146.—A canoe-shaped valley bordered by a ridge
formed by the outcrop of a hard layer in a plunging syncline. The
ridge bounding the canoe-valley is separated from an outer ridge by a
curved valley underlain by relatively weak rock. (After Willis.)]
[Illustration: +Fig.+ 147.—A diagram to illustrate the effects of
erosion on a doubly-plunging anticline made up of beds of unequal
hardness.]
[Illustration: +Figs.+ 148–51.—Diagrams to illustrate the shifting of
rivers from a synclinal to an anticlinal position. (After Davis.)]
The anticlines and synclines under consideration are assumed to have a
thick hard layer at the surface, and softer beds below. This is shown
in the cross-section introduced in the figure, the upper hard stratum
(_m_) being indicated by the dots, while the softer one (_n_) is white.
The line _oo_ represents base-level, which is below the hard layer both
in the syncline and anticline, but much farther below in the latter
position than in the former. Because of their higher gradients, and
because of the greater fracturing to which the region they drain was
presumably subject at the time of folding, the tributary streams might
cut through the hard layer sooner than the main stream which they join.
This done, they would enlarge their valleys rapidly in the softer rock
beneath, and secondary tributaries would be developed (Fig. 149). When
the condition of things represented in Fig. 149 is reached, the streams
_c_ and _d_, tributary to the synclinal stream, come into competition.
The former has the advantage over the latter, because it joins the
main stream at a lower level. Stream _c_ will therefore be likely to
capture _d_. The incipient stages of the capture are stealthy, and the
later bold. At first the divide between their head waters is shifted
northward inch by inch, because the gradient toward _g_ is higher than
that toward _e_. The capture of the head waters of _e_ is as slow as
the migration of the divide, until the divide reaches the point where
_e_ joins _f_. The stream _f_ is then diverted promptly into the valley
of _g_, and is at once led away to _c_ (see Fig. 150). Strengthened
by its increased volume, the stream _c_ (Fig. 150) lowers its valley
across the hard layer more rapidly than before, and so holds the
advantage it has gained. Not only this, but the beheaded stream _d_
(Fig. 150), because of its diminished volume, sinks its valley into the
hard layer less rapidly than before, and its decrease in power also
works to the advantage of the stream leading to _c_. The result is that
the divide between _fg_ and _d_ does not remain constant, but is driven
back step by step toward _a_.
Similarly a tributary to the main stream at _b_ (Fig. 150), may by
means of its tributary _h_, capture the waters of _fg_, and lead them
to the synclinal valley at _b_ (compare Figs. 150 and 151). Deprived
of its main source of supply (at _c_) the synclinal stream is greatly
diminished above _b_, and cuts more and more slowly, while the stream
_fgh_ (Fig. 151), having greater volume and working mainly in softer
rock, sinks its channel faster than the stream in the synclinal axis.
Under these circumstances, the stream at _f_ may cut its valley below
the valley in the synclinal axis _a_ (Fig. 150). In this event, the
divide between _f_ and _a_ (Fig. 150) may be pushed back until the
synclinal stream is beheaded at _a_ and carried out of the syncline and
over into the anticlinal valley (Fig. 151). Thus, the old anticlinal
axis comes to be the course of the main stream. Similarly the stream
entering the syncline at _b_ (Fig. 151) might later be captured by _i_,
thus lengthening its anticlinal course.
It is not to be understood that this sequence of events will take
place in the degradation of every anticline, but the principles here
set forth will always be operative. The result specified will be
accomplished wherever hard and soft layers have the relations indicated
in the diagrams; that is, where the stream in the syncline finds itself
on a resistant layer as it approaches base-level, while at the same
time the (original) tributary streams are working in softer beds. It is
not to be understood, therefore, that streams migrate from synclines to
anticlines for the sake of getting out of the former positions into the
latter. If they shift their courses it is to find easier ones.
That these changes are not fanciful is shown by the fact that the
adjustment described corresponds with that shown in many parts of the
Appalachian Mountains, and in other mountains of similar structure.
If in a later stage of its history, the new main stream, _fh_, were
to cut its bed down to a lower hard layer, while the original stream,
_ab_, reached a softer bed beneath the hard one above, the latter would
again have an advantage, and a new series of adjustments would be
inaugurated which might result in re-establishing the main stream in
its original synclinal position.
EFFECT OF CHANGES OF LEVEL.
=Rise.=—If after being base-leveled, or notably reduced by erosion, a
region is uplifted so as to increase the gradients and therefore the
velocities of the streams which drain it, the streams are said to be
_rejuvenated_, and a new cycle of erosion is begun. If the rise of the
area were equal everywhere, while the coast line remained constant
in position, there would be an immediate increase in velocity only
at the debouchures of the streams flowing directly into the sea.
At the debouchures of such streams there would be rapids or falls.
Each rapids or falls would promptly recede, and with the recession,
the acceleration of velocity resulting from the uplift would be felt
farther and farther up-stream, and ultimately to its source. The
rejuvenated streams would cut new valleys in the bottoms of their
old ones (Figs. 152 and 153). The new valleys would begin where the
increase in velocity was first felt, and they would be lengthened by
head erosion just as valleys of the first cycle were lengthened.
[Illustration: +Fig.+ 152.—Cross-section of a wide valley, _ab_, in the
bottom of which a younger valley, _cd_, has been excavated as the
result of uplift.]
[Illustration: +Fig.+ 153.—Diagram to illustrate in ground plan an
ideal case of rejuvenation as the result of uplift.]
When the head of the new part of a valley of a rejuvenated stream
recedes past the mouth of a tributary adjusted[58] to the gradient of
the main stream before rejuvenation, the velocity of the tributary
is accelerated at its debouchure, and it begins to excavate a new
valley in the bottom of its old one. The new valley commences at the
lower end of the old one, and develops headward (_a_ and _b_, Fig.
153). Good illustrations are furnished by the streams in the west
central part of New Jersey. The Delaware has here a sharply defined
valley, and its tributaries are essentially as deep as their main at
the point of junction. Above this point they have high gradients for a
short distance (three to six miles), beyond which they wind sluggishly
in wide valleys with low gradients across a relatively high plateau.
Their profiles are illustrated by Fig. 154. The flat, though high,
surface in which their upper courses lie, appears to have been nearly
base-leveled in an earlier cycle, and then to have been elevated.
The date of the elevation is fixed, in terms of erosion, by the time
necessary for the excavation of the Delaware gorge, and the narrow
gorges along the lower courses of its tributaries. It was so recent
that the effects of rejuvenation, proceeding from the debouchures of
the tributaries toward their heads, have not yet advanced far from the
Delaware. Similar relations are found elsewhere (Fig. 1, Pl. XIII,
s. c. Col.). Another peculiarity of rejuvenated drainage is shown in
Fig. 2, Plate XIII (s. Kan.). Here Elm Creek flows at a level 200 feet
below that of Sand Creek, 4 miles distant. The valley of the former
appears to have entered upon a new cycle as the result of uplift,
while that of the latter, _in the area shown on the map_, is still
unrejuvenated. Farther down-stream, the valley of Sand Creek shows
signs of rejuvenation. It may be noted that a tributary of Amber Creek
has good opportunity to capture Sand Creek, for the latter flows about
25 miles before reaching the level of Amber Creek at its junction with
Elm Creek.
[Illustration: PLATE XIII.
Fig. 1. COLORADO.
U. S. Geol. Surv.
Fig. 2. KANSAS.
U. S. Geol. Surv.]
[Illustration: PLATE XIV.
Fig. 1. PENNSYLVANIA.
U. S. Geol. Surv.
Fig. 2. CALIFORNIA.
U. S. Geol. Surv.]
[Illustration: +Fig.+ 154.—Profile of a rejuvenated stream. The
Lockatong River (N. J.) to the head of Mud Run.]
Should the lower end of a tributary valley fail to be degraded as fast
as the valley of the main at the point of junction, the tributary is
out of _topographic adjustment_ with its main. Falls or rapids may
result. When the lower end of a tributary valley is distinctly above
the level of its main, the former is called a _hanging valley_. Hanging
valleys developed by stream erosion alone are not common except just
after the recession of a falls past the mouth of a tributary. Hanging
valleys, as well as the characters and relations illustrated by Figs.
152–154 are criteria of rejuvenation, but they must be applied with
discretion. Such profiles, for example, as that shown in Fig. 154 may
be developed when the rock of a stream’s bed is unequally resistant,
and hanging valleys are generally a result of glaciation (see Chapter
V).
Rejuvenated streams sometimes inherit certain peculiarities from their
aged ancestors. Thus a rejuvenated stream may intrench the meanders
possessed by the old stream which preceded (see Fig. 1, Pl. XIV, near
Harrisburg, Pa.), and intrenched meanders are one of the marks of
rejuvenated streams. They are not uncommon in the Appalachian Mountain
regions, and are known in other parts of the world. The Seine and the
Moselle furnish further illustrations.[59]
The history of the new cycle of erosion inaugurated by the uplift would
differ from that of the preceding cycle in that the new one would begin
with a drainage system already developed. Other things being equal,
therefore, the reduction of the land would proceed more rapidly in a
subsequent cycle than in the first.
The recognition of different cycles of erosion, separated by uplifts,
is often easy. The principles involved are illustrated by Fig. 155
which represents an ideal profile of considerable length (say 50
miles). The points _a_, _a′_, and _a″_ reach a common level. Below
them there are areas _b_, _b′_, and _b″_ which have a nearly common
elevation, below which are the sharp valleys _d_, _d′_, and _d″_. The
points _a_, _a′_, and _a″_ represent the cross-sections of ridges
formed by the outcrops of layers of hard rock. If the crests of the
ridges are level, the points _a_, _a′_, and _a″_ must represent
remnants of an old base-level, _since at no time after a ridge of hard
rock becomes deeply notched does it acquire an even crest, until it is
base-leveled_.[60] At all earlier stages its crest is uneven. After the
cycle represented by the remnants _a_, _a′_, and _a″_ was completed,
the region suffered uplift. A new cycle represented by the plain _b_,
_b′_, and _b″_ was well advanced, though not completed, when the region
was again elevated, and the rejuvenated streams began to cut their
valleys _d_, _d′_, and _d″_ in the plain of the previous incomplete
cycle. The elevations, _c_ and _c′_ (intermediate in elevation between
_a_, _a′_, and _a″_, and _b_, _b′_, and _b″_) may represent either
remnants of the first base-level plain which were lowered, but not
obliterated, while the plane _b, b′, b″_ was developing; or they may
represent a cycle intermediate between that during which _a, a′, a″_
and _b, b′, b″_ were developed. If the intermediate elevations (_c,
c′_) have a common height and level crests, the presumption would be in
favor of the latter interpretation. If they be numerous and of varying
heights, as is possible, they may in the field obscure the planes (_a,
a′, a″_ and _b, b′, b″_) developed in the different cycles, which, in
the figure, are distinct.
[Illustration: +Fig.+ 155.—Diagram to illustrate cycles of erosion
where the beds are tilted.]
If the strata involved be horizontal the determination of cycles is
sometimes less easy. Thus in Fig. 156, it is not possible to say
whether _a_ and _a′_ represent remnants of an old base-level, or
whether they represent the original surface from which degradation
started. So, too, the various benches below _a_, such as _b_, _b′_, and
_b″_ may readily be the result of the superior hardness of the beds at
this level. For the determination of successive uplifts in the field it
is necessary to consider areas of considerable size, and to eliminate
the topographic effects of inequalities of hardness, and of certain
other factors to be mentioned presently.
[Illustration: +Fig.+ 156.—Diagram to illustrate cycles of erosion
where the beds are horizontal.]
The inequalities in the depths of the young valleys in Figs. 155 and
156 may be explained on the supposition that the deeper ones belong to
main streams, and the shallower ones to tributaries. Such a valley as
that shown at _e_, Fig. 155, suggests rejuvenation at this point; but
farther up the stream which occupies this valley, rejuvenation might
not be apparent. In this case, the main streams might be flowing in new
valleys, _d, d′_, etc., while the heads of their tributaries are still
flowing in the older valleys of the preceding cycle (compare Fig. 154
and Fig. 1, Pl. XIII).
It is by the application of the preceding principles that it is known
that the Appalachian Mountains, after being folded, were reduced to
a peneplain (p. 76), throughout their whole extent from the Hudson
River to Alabama. The peneplain level is indicated by the level
crests of the Appalachian ridges, shown in cross profile by the high
points of Fig. 157. The system was then uplifted, and in the cycle
of erosion which followed, broad plains were developed at a new and
lower level, corresponding in a general way to the plains _b_, _b′_,
and _b″_ of Fig. 155. The plains were located, for the most part,
where the less resistant strata come to the surface. Above them rose
even-crested ridges, the outcrops of the resistant layers, which
had been isolated by the degradation of the softer beds between.
They constitute the present mountain ridges (the high points of Fig.
157). The evenness of their crests, testifying to the completeness
of the first peneplanation, is shown in Fig. 158, which represents,
diagrammatically, a longitudinal profile of an Appalachian Mountain
ridge. The evenness of the crest is interrupted by (1) notches (_b,
c_, etc., Fig. 158) cut by the streams in later cycles, and (2) by
occasional elevations above the common level (monadnocks, _a, a′_, Fig.
158). The monadnocks are generally rather inconspicuous, but there is a
notable group of them in North Carolina and Tennessee. Mount Mitchell
and Roane Mountain are examples. When long distances are considered,
the ridge crests depart somewhat from horizontality. This is believed
to be due, in part at least, to deformations of the old peneplain
during the uplift which inaugurated the second cycle of erosion.
[Illustration: +Fig.+ 157.—Cross-section of a portion of the
Appalachian Mountains to illustrate the phenomena of erosion cycles.
(After Rogers.)]
[Illustration: +Fig.+ 158.—A diagrammatic longitudinal profile of an
Appalachian Mountain ridge.]
The extent to which the second cycle of erosion recorded in the
present topography had proceeded before its interruption by uplift,
is indicated by the extent of the valley plains (Fig. 157) below the
mountain ridges. While these plains were being developed on the weak
rocks, narrow valleys only (Fig. 158) were cut in the resistant rocks
which now stood out as ridges. In Fig. 158 some of these valleys are
shallow (_c, c′, c″_, etc.), and but one of them deep. The former may
be either (1) the valleys of streams which crossed the hard layer
at the beginning of the cycle, and which were diverted before their
valleys became deep; or (2) they may represent the heads of valleys
now working back into the ridges. The deep valley (_b_) represents the
work of a stream which has held its course across the hard layer while
the latter was being isolated as a mountain ridge (compare Figs. 131
and 132). Deep narrows of this sort are often called _water-gaps_.
Similar valleys, whether shallow or deep, from which drainage has
been diverted, are sometimes called _wind-gaps_. The second cycle of
erosion, while still far from complete, was interrupted by uplift
(relative or absolute), and a new cycle inaugurated. This event was so
recent that the new (third) cycle has not yet advanced far.
[Illustration: +Fig.+ 159.—The Kittatinny Mountains and Delaware
Water-Gap from Manunka Chunk. (N. J. Geol. Surv.)]
Recently it has been urged that another cycle, intermediate between the
first and second, is to be recognized.[61]
Some of the features just described are illustrated by Fig. 159. The
even mountain crest in the background is the Kittatinny Mountain of
New Jersey and its continuation in Pennsylvania. In common with other
corresponding crests it represents the oldest recorded base-level (or
peneplain) of the region. The great gap in the mountain is the Delaware
Water-Gap. Below the mountain crest there is another plain, developed
in a subsequent cycle of erosion, while the valley plain in the
foreground represents the work of a still later cycle.
[Illustration: +Fig.+ 160.—Showing certain peculiarities of Appalachian
drainage. 1 = the Susquehanna; 2 = the Potomac; 3 = the James; 4 =
the Roanoke; 5= the Coosa; 6 = the Tennessee; 7 = the Kanawha; 8 =
head of New River; 9 = head of the French Broad.]
The oldest erosion plain of the Appalachian Mountains, the results of
which are seen in the even-crested ridges so characteristic of the
system, is sometimes called the _Kittatinny base-level_.[1] It was
completed early in the Cretaceous period, and hence is sometimes known
as the _Cretaceous base-level_. The next lower plain, imperfectly
developed, has been called the Shenandoah Plain,[62] from the
Shenandoah Valley where it is well seen (Fig. 132 and Fig. 2, Pl. XII).
It is to be noted that the terms _base-level_ and _peneplain_ have
both been used in connection with these old plains. _Graded plain_ is
equally applicable. The truth is that the topographic types represented
by these three terms grade into one another. It may be questioned
whether definitions should be insisted on which differentiate these
types more sharply than Nature has.
Many of the peculiarities of the drainage of the Appalachian Mountain
system are intimately connected with the history just outlined. Thus
three great rivers, the Delaware, the Susquehanna, and the Potomac,
have their sources west of the Appalachians proper, cross the system
in apparent disregard of the structure, and flow into the Atlantic.
The James and Roanoke head far to the west, although not beyond the
mountain system, and flow eastward, while the New River (leading to
the Kanawha) farther south, heads east of the mountain-folds, and
flows northwestward across the alternating hard and soft beds of the
whole Appalachian system, to the Ohio (Fig. 160). The French Broad,
a tributary to the Tennessee, has a similar course. Such streams are
clearly not in structural adjustment, and afford good opportunities
for piracy. Their courses were apparently assumed during the time of
the Kittatinny base-level, when the streams had so low a gradient as
not to be affected by the structure (p. 150). Elevation rejuvenated
them, and they have held their courses in succeeding cycles across beds
of unequal resistance, though smaller streams have become somewhat
thoroughly adjusted. Crustal deformations have also helped them to hold
their courses, for the Cretaceous peneplain seems to have been tilted
to the southeast at its northern end, and to the southwest at its
southern, when the succeeding cycle began.
Streams which hold their early courses in spite of changes which
have taken place since their courses were assumed are said to
be _antecedent_. They antedate the crustal movements which, but
for pre-existent streams, would have given origin to a different
arrangement of river courses. As a result of crustal movements,
therefore, a consequent stream may become antecedent. Master streams
are more likely to hold their courses, and therefore to become
antecedent, than subordinate ones.
The uplift of base-leveled beds, especially if the beds are tilted so
as to bring layers of unequal resistance to the surface at frequent
intervals, affords conditions favorable for extensive adjustment. The
numerous wind-gaps in the mountain ridges, representing the abandoned
courses of minor streams, and the less numerous water-gaps, which
indicate the resistance of large streams to structural adjustment, are
instructive witnesses of the extent to which adjustment has gone. So
extensive has been the adjustment among the streams of the Appalachian
Mountains that there is probably no considerable stream in the whole
system which has not gained or lost through its own or its neighbors’
piracy. The history of the rivers of the Appalachian Mountains has been
further complicated by a considerable amount of warping during the
periods of uplift.[63]
[Illustration: +Figs.+ 161, 162.—Diagrams to illustrate the effect of
crustal warping on stream erosion. The dotted lines represent the
profiles of the streams before deformation; the full lines, after.
Erosion will be stimulated between _a_ and _b_ in each case, and
between _c_ and _d_ in Fig. 162. Below _b_, Fig. 161, the stream
will be drowned, and erosion therefore stopped. Erosion will also be
stopped or retarded above _a_, between _b_ and _c_, and below _d_ in
Fig. 162.]
=Sinking.=—The land on which a river system is developed may be
depressed relative to sea-level. In this case the sea would occupy
the lower ends of valleys, converting them into bays and estuaries. A
stream in this condition is said to be _drowned_. Of drowned rivers
there are many examples along the Atlantic coast. Thus the St.
Lawrence River is drowned up to Montreal, and the Hudson up to Albany.
If the drowned portion of the latter valley were not so narrow, it
would be a bay. Delaware and Chesapeake Bays, as well as many smaller
ones, both north and south, are likewise the drowned ends of river
valleys (see figures, Chapter VI). If all parts of a drainage basin
sank equally, the velocities of the streams above the limit of drowning
would not be changed, for the gradients would remain the same as
before. The fact that a river’s channel is below sea-level is not to
be taken as proof that the valley is drowned. Thus the bottom of the
channel of the Mississippi is as much as 100 feet below the level of
the Gulf, some 20 miles above New Orleans.[64]
=Differential movement. Warping.=—Where a land surface on which a river
system is established suffers warping, some parts going up and others
down, the opposite movements being either absolute or relative, various
phenomena would result. This may be illustrated by the accompanying
diagrams (Figs. 161 and 162), where the profiles of the streams are
represented as warped from the positions represented by the dotted
lines, to the positions shown by the full lines. The velocity will be
accelerated below the points of differential elevation (between _a_ and
_b_, Fig. 161, and between _a_ and _b_, and _c_ and _d_, Fig. 162), but
checked above (above _a_, and between _b_ and _c_, Fig. 162). Above an
elevation which notably checks its flow, a stream is _ponded_. If the
ponding is slight, a marsh may develop above the obstruction; if more
considerable, a lake is formed. Lakes of this class are likely to be
short-lived, since the ponded waters are likely to soon overflow and
lower their outlet so as to drain the lake. The elevation which ponds
the stream may be great enough and rapid enough so that the resulting
lake finds an outlet by some course other than that originally followed
by the stream. Where a stream holds its course across an uplift athwart
its valley, either with or without ponding, it becomes an antecedent
stream (see p. 169), since it has a course assumed before the latest
deformation of the crust and in apparent disregard of present surface
configuration. Thus the Columbia River holds its antecedent course
across areas which have been uplifted (differentially) hundreds and
even thousands of feet.[65] Some of the striking scenic features of
this noble valley are the result of these changes in the country
through which it flows. A lesser stream would have been diverted, as
many of its tributaries have been. Even its course across the Cascade
ranges is believed to be antecedent.[66]
[Illustration: +Figs.+ 163, 164.—Piracy stimulated by warping. Uplift
along axis 1–2.]
Another peculiarity of valleys and streams resulting from changes of
level is illustrated in Fig. 2, Pl. XIV (southern California). The
main valleys of this part of the coast were developed when the land
stood considerably higher than now. Later the subsidence of the coast
converted the lower ends of the valleys into bays or fiords. The bays
were then transformed into lagoons by deposition. Subsequent rise of
the land or depression of the sea allowed the drainage from the old
lagoons to cut across the deposits which had converted the bays into
lagoons. The result is an old, wide valley above, suggested by a young
one below.
If the warpings were considerable, much more decisive changes in
drainage would result. Suppose the drainage of a given region to be
represented by the streams in Fig. 163. If there is uplift along the
axis 1–2, that part of _ac_ above the axis of uplift would be ponded,
or at least have its velocity checked, while the flow of some of the
tributaries of _d_ would be accelerated, and might work back and
capture the other stream (Fig. 164).
Crustal warping was one of the conditions under which the Tennessee
achieved its present anomalous course, and its history[67] is
illustrative of the complex changes which drainage suffers when warping
affects the area where the rock structures are of unequal resistance.
At the close of the Cretaceous cycle of erosion, when the Appalachian
Mountains had been reduced to a peneplain, the waters falling in
the area now drained by the upper course of the Tennessee flowed
south-south-west to the Gulf in a stream (the Appalachian River, _a_,
Fig. 165) the lower part of which had the general position of the Coosa
and the Alabama.
To the west of the Appalachian River, shorter streams flowed west
and southwest into the Mississippi embayment (Fig. 165) by courses
which are not now definitely known. The succeeding cycle of erosion
was inaugurated by uplift and deformation of the peneplain. The axis
of greatest elevation (_AB_, Fig. 166) was nearly parallel to the
Appalachian River, and the effect of the differential uplift was to
impose a greater task on this river (_a_, Fig. 166), which flowed
along the axis of uplift, than upon the rivers which flowed westward
and southwestward to the Mississippi embayment. The result was that
the strongest of the southwesterly flowing streams worked its head
back into the drainage basin of the Appalachian River, and captured,
one by one, the head-waters of its westerly tributaries, establishing
some such drainage relations as are shown in Fig. 166. Still later,
after the land area of the region had been considerably extended by
the withdrawal of the sea, the Appalachian River itself was reached by
the invading stream, and its waters carried away to the Mississippi Bay
by a course the lower part of which is thought to have corresponded
approximately with the course of the present Black River (_b_, Fig.
167).
[Illustration: +Fig.+ 165.—Shows the general position of the main
drainage lines in the southern Appalachians at the close of the
Cretaceous cycle of erosion. The lower part of stream _b_ is made to
follow the course of a portion of the present Tennessee.]
[Illustration: +Fig.+ 166.—Shows the general position of the main
drainage lines in the southern Appalachians, after the capture of the
westerly tributaries of the Appalachian River by stream _b_. Compare
Fig. 165.]
[Illustration: +Fig.+ 167.—A stage later than that shown in Fig. 166.
The sea is represented as having withdrawn from a considerable area
which was submerged at earlier stages (Figs. 165, 166).]
[Illustration: +Fig.+ 168.—Shows the final change which resulted in the
present course of the Tennessee. The land is represented as somewhat
higher than now.[68]]
Still later there was further deformation which caused additional
changes in the drainage. The whole region was uplifted, relatively if
not absolutely, but the uplift was differential, being greatest along
the axis represented by _AB_, Fig. 167. The effect of the deformation
was to stimulate the tributaries of the Ohio flowing north from this
axis. Their growth was further accelerated by the weakness of the
strata over which they ran. At the same time, the uplift to the south
led the southwesterly flowing stream (_b_, Fig. 167) to discover
relatively hard beds of rock in its lower course, and these beds
retarded its down-cutting. The result was that a tributary of the Ohio
(_a_, Fig. 167) finally tapped the main stream flowing to the southwest
(_b_, Fig. 167) and carried its upper part over to the Ohio (Fig. 168).
This was the beginning of the present Tennessee.
THE AGGRADATIONAL WORK OF RUNNING WATER.
=Principles involved.=—Since deposition results from the failure
of transportation, the factors which control transportation also
influence deposition. Transportation by streams is determined largely
by velocity, and the most important factors influencing velocity are
slope, volume, and load (p. 115). Of these the first two are usually of
greater importance than the third.
A stream is said to be loaded when it has all the sediment it can
carry; it is loaded with fine material when it has all the fine
material it can carry, and with coarse material when it has all the
coarse it can transport. A stream loaded with coarse material flows
more swiftly than one loaded with fine, for a larger percentage of a
stream’s energy can be utilized in carrying fine material than coarse,
and hence a larger percentage of the energy of a stream which carries a
load of the latter will express itself in velocity.
Deposition takes place whenever a stream finds itself with more load
than it can carry, and is an expression of the stream’s refusal to
remain overloaded. A stream may become overloaded in various ways.
It might at first seem unnecessary to inquire whether a stream may
be overloaded at its source, but the question is not necessarily to
be answered in the negative. The source of a stream is not always a
definite point. In a general way it may be said that the source of the
normal stream is at that point in its valley where the bottom is as low
as the ground-water level of the region. But since the ground-water
level is not constant (p. 71) the source of a stream is likely to be
farther up its valley in a wet season than in a dry one (p. 72). After
a heavy shower, the run-off descends to the axis of the valley from the
slopes on all sides, and temporarily the stream begins above the point
which marks even its wet-season source. If under such circumstances
the slopes about the head of the valley are notably steeper than the
slope of the valley itself, as they frequently are, the water flowing
down them may gather an amount of material which it cannot carry after
it reaches the bottom of the valley. This may be the case at, or even
above, the point which marks the source of the permanent stream. It is,
therefore, possible for a stream to be overloaded at its source, if we
take the source to be the point whence the water permanently flows.
Deposition may, therefore, be taking place in a valley at the head of
its permanent stream, or temporarily even in the valley above it.
Streams issuing from glaciers sometimes have more load than they can
carry after they escape from the ice. If the stream be regarded as
beginning at the point where it issues from beneath the ice it may be
overloaded at its source.[69]
Under certain circumstances, a stream may overload itself. Thus if a
stream loaded with coarse detritus reaches a portion of its valley
where fine material is accessible in abundance, some of the velocity
which is helping to carry the coarse may be used in picking up and
carrying the fine. This reduces the velocity, and since the stream
already had all the coarse material it could carry, reduction of
velocity must result in deposition. It follows that when a stream fully
loaded with coarse material picks up fine, it becomes overloaded, _so
far as the coarse material is concerned_.
Again, tributaries may overload their mains. While tributaries
are usually smaller than their mains, they frequently have higher
gradients, and the smaller stream of higher gradient may bring to the
larger stream of lower gradient more material than the latter can
carry away. Thus deposition may take place at the point of junction of
tributaries with their mains. This may go so far as to pond the latter
enough to cause its expansion into a river-lake. Lake Pepin, in the
Mississippi River at the mouth of the Chippewa (in Wis.), is an example.
Streams may become overloaded by losing velocity or volume, or both.
Decrease in velocity is brought about either by decrease in declivity
or in volume. In general, streams have lower gradients and greater
volumes in their lower courses than in their upper, and these two
elements affect velocity in different ways. If the increase in volume
be not enough to counterbalance the decrease in declivity, as is often
the case, a stream which is loaded in its upper course will deposit
in its lower. The decrease of velocity at the debouchure of a stream
almost always leads to deposition.
Decrease in velocity as the result of decrease in volume is less
common. When decrease in volume occurs, it may be the result of (1)
evaporation, (2) the absorption of water into the bed of the stream,
or (3) branching—the giving off of _distributaries_. While evaporation
is going on everywhere, the diminution of a stream by this means is
usually more than balanced by the increase from tributaries, rainfall,
and springs; but in arid regions a very different condition of things
sometimes exists. If mountains in an arid region be capped with
snow, its melting supplies the streams during the melting season. As
the streams flow out from the mountains through dry regions, they
receive little or no increment from rainfall, tributaries, or springs,
and evaporation reduces the volume of water, or even dissipates it
altogether. Absorption of water into the bed of the stream often
accompanies evaporation. Reduction of volume by evaporation and by
absorption is especially common in arid regions. Wherever loaded
streams are reduced in volume, whether by evaporation or absorption,
deposition takes place.
The third way by which velocity is decreased as the result of
decreasing volume is illustrated at the debouchures of many streams.
Near the Gulf, for example, the Mississippi branches repeatedly (see
Fig. 190). The same phenomena are often seen where one stream joins
another (Fig. 169). Individually the distributaries are much smaller
than the main stream before they separated from it, and because they
are smaller their combined surfaces are greater, and the amount of
energy consumed in the friction of flow is increased. The velocity of
the water and its carrying power are, therefore, reduced. Thus the
branching of streams gives rise to deposition, and where deposition
takes place the gradient of the stream is reduced, and this occasions
still further deposition. The sediment which fills up the channel and
checks the flow finally compels the stream, or some part of it, to
transgress its banks. Deposition, therefore, favors the development of
distributaries, and the development of distributaries in turn favors
deposition.
[Illustration: +Fig.+ 169.—Delta of the Chelan River at its junction
with the Columbia. Shows the tendency of streams to distribute where
active deposition is in progress. (Willis, U. S. Geol. Surv.)]
The foregoing statements make it clear that a stream may be eroding
in one part of its valley while it is depositing in another, and that
erosion may alternate with deposition in the same place, on account of
fluctuations in volume, and, therefore, in velocity of the stream. It
will be seen in the sequel that erosion and deposition may be taking
place at the same time in the same part of the valley. The activities
of a river are so nicely balanced that slight disturbance at one point
causes disturbance at all points below.
_The deposits._
=Types.=—Turning from the principles which underlie river deposition to
the deposits themselves, they are found to occur in various situations.
Running water usually descends from steeper slopes above to gentler
slopes below, and ends at the sea, or in a lake or inland basin.
Wherever there is a sudden decrease in its gradient, as at the base of
a hill, ridge, or mountain, running water is likely to leave a large
part of its load, building an _alluvial fan_ or _cone_ (Figs. 67, 68,
and Pl. VI). Even where there is no sudden decrease in the gradient
of a stream, there is likely to be a gradual one, and in spite of the
fact that the increased volume of a stream in its lower course tends to
overcome the effect of diminished gradient on velocity, deposition is
likely to take place as the gradient is reduced. Deposits occasioned by
the gradual reduction of a stream’s velocity often have great extent
in the direction of a stream’s flow. They cover the flood plains of
streams, making them _alluvial plains_ (Fig. 73). When a stream reaches
the sea or a lake its current is destroyed and its load dropped, unless
taken in charge by the waves and currents of the standing water.
Sediment accumulated in quantity at the debouchures of streams gives
rise to _deltas_ (Figs. 169, 187). Alluvial cones and fans, alluvial
plains, and deltas, are the principal types of river deposits. Apart
from these well-defined types there are _bars_ in the channels of
depositing streams, and much ill-defined alluvium which does not allow
of ready classification.
=Alluvial fans and cones.=—The only distinction between the alluvial
fan and the alluvial cone is one of slope, the cones (they are but
half-cones at best) being steeper than the fans. Alluvial fans and
cones have their most striking development where temporary torrents,
occasioned by showers or the rapid melting of snow, issue from mountain
ravines. Such streams usually carry heavy loads of detritus, the
coarser part of which is likely to be deposited at the base of the
mountain slope. Cones and fans built by such streams have a periodic
rather than a steady growth.
At the beginning of its development the material of the alluvial cone
is deposited much as in a talus cone (compare Fig. 170 with Figs. 67
and 68). Its deposition chokes the channel of the stream, and some of
the water then seeks new courses to right and left of the apex of the
deposit. This expands the area of deposition to right and left, while
the water which flows over it lengthens it in the direction of flow.
The course and behavior of the water after reaching an alluvial cone is
instructive. As its velocity is checked, deposition often takes place
in the channel, diminishing its capacity. As the channel is filled up,
the water tends to overflow on either side. The overflowing water,
being shallow, has so little velocity that much of its load is dropped
on either margin of the channel, building up levees. The water ever and
anon breaks through the levees, giving rise to distributary streams,
each of which aggrades its channel and builds its own miniature
levees (Fig. 171). Not rarely this process of channel-filling and
levee-building goes on until the channels of the little rivulets are
above the general level of the cone on which they rest. The rivulet
then runs in a groove on the crest of a little ridge. The channels on
the surfaces of fans and cones are fewest and deepest at their heads,
and more numerous and shallower below. In some cases the surface-water
disappears altogether before the outer border of the fan is reached, by
sinking into the débris.
[Illustration: +Fig.+ 170.—A talus cone. North Greenland Coast. The
talus cone reaches the sea-level. Drawn from photograph.]
Alluvial fans and cones have various forms, and often attain
considerable dimensions. Their angles of slope depend on the amount
of reduction of velocity which the depositing water suffers, and the
amount and kind of load which it carries. The maximum slope of the
cone is the angle at which the loose material involved will lie. The
minimum slope of the fan, on the other hand, approaches horizontality.
If many alluvial fans develop in proximity to one another, as at the
base of a mountain range, they may expand laterally until they merge.
A long succession of them may thus give rise to an extensive _alluvial
piedmont plain_, or a _compound alluvial fan_. The lower edge of such a
fan is often somewhat lobate. Such plains exist along the bases of many
mountain ranges (Pl. VI), and may be seen in miniature even along low
ridges.
[Illustration: +Fig.+ 171.—Miniature levees on an alluvial cone. Slope
of Gray Peak, Colo. (R. T. Chamberlin.)]
A permanent stream, as well as a temporary one, may develop an alluvial
fan at the base of a mountain slope; but since the mountain course of
the former is likely to be less steep than that of the latter, its
waters suffer a correspondingly less reduction of velocity at any
one point. The fan of the permanent stream is therefore likely to be
relatively flat, and to stretch far down the valley. Such fans grade
into valley plains. From the general principles already discussed, it
is clear that well-developed fans go with relatively youthful stages of
erosion, and belong normally to the upper parts of drainage lines.
=Ill-defined alluvium.=—There is a widespread mantle of alluvial
material deposited by running water which was not organized into
distinct streams. The water which runs down smooth slopes in sheets
during showers carries fine earthy matter, as well as some that
is coarser. These materials are largely deposited at the bases of
the slopes, forming basal accumulations of greater or less extent,
comparable in origin to alluvial fans. A relatively small amount of the
slope wash is carried far out from the base of the declivities. It is
not easy to realize the extent to which this process is taking place.
There is hardly a slope without loose material, and there is hardly an
acre of low land below a slope on which running water has not deposited
sediment washed down from above. When it is remembered that this is
as true of gentle slopes and their surroundings as of steep slopes,
though perhaps not to the same extent, and that a very large part of
the earth’s surface is made up of sensible slopes, or of flats at their
bases, some idea of the aggregate effect may be gained.
There is another way of looking at the same question. Earthy matter
is being continually transferred from land to sea, and chiefly from
high land. Rarely does it start from any point distant from the shore
and move uninterruptedly to it. It is transported a short distance and
lodged, to be again picked up, carried forward another step in its
journey, and lodged again. For a very large part of the earth’s surface
it would be true to say that its mantle rock is material in transit
from higher land to the sea.
=Alluvial plains.=—Most streams, whether heading in mountains or not,
have gentler gradients in their lower courses than in their upper, and
in spite of increasing volume are usually unable to carry to their
debouchures all the material gathered above. The excess of load is
dropped chiefly on the flood-plains of the streams and constitutes them
_alluvial plains_.
The making of an alluvial plain usually involves both erosion and
deposition. When a stream has cut its channel to grade, downward
erosion ceases, or more exactly, downward cutting is, on the average,
counterbalanced by deposition. So long as a stream is cutting downward
rapidly, it carries away whatever débris descends the side slopes.
When it approaches grade, the débris which descends the side slopes
tends to accumulate at their bases, and the V-shaped cross-section of
the valley becomes U-shaped (see Fig. 172). At about the same time the
stream begins to meander, for, having lost something of its former
velocity, it is more easily turned from side to side. As it begins to
meander, it widens the bottom of its valley. This is the initial stage
in the development of the valley flat (2 and 3, Fig. 172). In its
meandering the stream encroaches on the talus accumulations at the
bases of its valley’s slopes. The side-cutting may remove all the loose
débris and even undercut the bluff as at _a_, Fig. 173. The stream’s
meanders shift their positions from time to time so that the valley
flat is successively widened at different points. By lateral planation,
therefore, a stream tends to develop a flat as soon as it reaches
grade. This is the initial part of erosion in the making of a river
flat, but a flat developed by erosion alone is not an alluvial plain.
So soon as the flat developed by a stream exceeds the width of its
channel, the water (except in times of flood) does not cover it all at
the same time. On any part which it temporarily abandons, some débris
(alluvium) is likely to be left. This deposit of alluvium constitutes
the valley flat an alluvial plain (Fig. 174). It will be seen that the
valley flat is commonly an alluvial plain from the beginning.
[Illustration: +Fig.+ 172.—Diagram illustrating the transformation of a
V-shaped valley into a U-shaped valley.]
[Illustration: +Fig.+ 173.—Diagram to illustrate the widening of a
valley flat by erosion. Compare 3, Fig. 172.]
Once the valley flat and alluvial plain are begun, their further
development is easily followed. The stream in flood overflows the banks
of its channel. The velocity of the overflowing water is reduced,
and if it has much load a part of it will be dropped and the plain
aggraded. Meantime meandering and lateral planation continue. Thus the
flood-plain is widened by erosion, and aggraded by alluviation, the two
processes going on simultaneously.
[Illustration: +Fig.+ 174.—An alluvial plain. The diagram suggests
the relative importance of lateral planation and alluviation in the
development of the flat.]
Flood-plains, chiefly the result of planation, but partly of
aggradation, are a normal feature of river valleys, after a certain
stage of development has been reached. This stage is that at which
downward erosion becomes slight in comparison with lateral erosion. It
follows that an alluvial plain normally begins its development where
the valley is first brought to grade, that is, in its lower course.
As the development of the valley goes on, the head of the flood-plain
advances up-stream, and at the same time its older parts become wider.
[Illustration: +Fig.+ 175.—Diagrammatic representation of a flood plain
developed by alluviation only.]
=Flood-plains due to alluviation only.=—Exceptionally, an alluvial
plain is developed by deposition only. Thus if a stream becomes
overloaded while its valley is still narrow, as sometimes happens,
deposition follows, and, as aggradation proceeds, the narrow valley
acquires a progressively wider bottom (Fig. 175). Wide valley plains
are sometimes developed in this way. Flood plains developed wholly by
alluviation are sometimes formed under conditions which are independent
of the stage of a valley’s development. Thus if a stream suddenly
acquires an exceptional supply of detritus in its upper course, the
development of an alluvial plain begins immediately below the point of
overloading.
The overload might be acquired in various ways. (1) If a stream taps
another (piracy) which carries a large quantity of sediment, carrying
off both water and sediment to a channel with a lower gradient,
deposition may take place where, under the earlier conditions, there
was none. (2) Again, when a stream cuts through a barrier near its
head waters, its velocity, and, therefore, its eroding power, may be
so increased in its upper course that sediment enough is acquired
to occasion deposition below, where none took place before. (3) In
working back through formations of varying degrees of resistance, a
stream’s head may presently reach a formation or a region which yields
abundant sediment, even though there was no especial barrier below.
(4) If an advancing glacier should reach the head waters of a stream,
its discharge to the stream would greatly increase the load of the
latter, and, although its volume would be augmented at the same time,
deposition might result. As a matter of fact, streams carrying glacial
drainage are usually aggrading streams. In general, anything which
greatly increases the load of a stream near its head is likely to cause
deposition, and so the development of a flood plain, at some point
farther down the valley.
[Illustration: +Fig.+ 176.—Anastomosing of a depositing stream. Yahtse
River, Alaska. (Russell, U. S. Geol. Surv.)]
Streams which are actively aggrading their valleys are likely to
anastomose (Figs. 176, 177). This results from the filling of the
channels until they are too small to accommodate all the water. The
latter then breaks out of the channel at few or many points. The new
channels thus established suffer the same fate.
[Illustration: +Fig.+ 177.—Anastomosing of the Platte River, Dawson
Co., Neb. (U. S. Geol. Surv.)]
=Flood-plains due to obstructions.=—Again, any obstacle in a stream’s
course is likely to cause deposition above. Thus dams built across
rivers entail the deposition of sediment above. Where a stream flows
over the outcropping edges of strata of different strength, the more
resistant serve, in some sense, as dams. Above them the stream cuts its
bed to a low gradient, and, becoming sluggish, drops more or less of
the detritus brought down from above. Obstacles of any sort across a
stream’s channel, therefore, favor the development of alluvial plains.
[Illustration: +Fig.+ 178.—The levees of the Mississippi in
cross-section, 4 miles north of Donaldsonville, La. Vertical scale
×50. The horizontal line in the diagram represents sea-level. The
bottom of the channel at this point is far below sea-level.]
=Levees.=—As the stream in flood escapes its channel and overspreads
its plain, its immediate banks are the site of active deposition, for
it is here that the velocity of the overflowing water is first notably
checked. On the banks of the channel, therefore, low alluvial ridges,
called _natural levees_, are built up (Fig. 178, and Pl. XV). They may
be narrow, or hundreds of feet in width, and are often several feet
above the plains behind them, giving the latter a slope away from the
channel of the stream. They are sometimes high enough to control the
courses of tributary streams, as shown by numerous tributaries to the
Mississippi below the Ohio. The Yazoo, for example, flows some 200
miles on the flood-plain of the Mississippi before it joins that river
near Vicksburg. The levees even become divides, directing drainage away
from the streams they guard (Pl. XV). Streams sometimes build levees
faster than their tributaries aggrade their channels. The latter are
then ponded, giving rise to lakes. The lakes on the lower courses of
the tributaries to the Red River of Louisiana are examples.[70] They
are sometimes built up above their natural level and kept in repair by
human agency so as to confine the streams in time of flood. This is a
source of danger unless they be steadily maintained, for the breaking
of such levees often occasions great destruction. A case in point is
the breaking of the levees of the Mississippi near New Orleans in 1890.
The water broke through the levees at the Nita and Martinez crevasses
(Fig. 187) and flowed eastward (from the former) with a current of 15
miles per hour, spreading destruction in its path. The water flowed
eastward through Lakes Pontchartrain and Borgne, and entered Mobile Bay
with such volume, velocity, and load of mud, as to destroy for a time
the oyster and fish industries of that locality.[71]
[Illustration: PLATE XV.
NEAR HAHNVILLE, LOUISIANA.
U. S. Geol. Surv.]
[Illustration: PLATE XVI.
Fig. 1. MISSOURI.
U. S. Geol. Surv.
Fig. 2. MISSOURI.
U. S. Geol. Surv.
Fig. 3. MISSOURI.
U. S. Geol. Surv.]
[Illustration: +Fig.+ 179.—Flood-plain of the Mississippi River south
of the mouth of the Ohio. (From charts of the Miss. Riv.
Commission.)]
[Illustration: +Fig.+ 180.—Diagram illustrating an early stage in the
development of meanders. The shaded part represents the area over
which the stream has worked.]
=Flood-plain meanders. Cut-and-fill.=—A stream with an alluvial plain
is likely to meander widely (Pl. XVI). In general terms this may be
said to be the result of low velocity, which allows it to be easily
turned aside. Were the course of such a stream made straight, it would
soon become crooked again. The manner of change is illustrated by Figs.
180 and 181. If the banks be less resistant at some points than at
others, as is always the case, the stream will cut in at those points.
If the configuration of the channel is such as to direct a current
against a given point, _a_ (Fig. 180), the result is the same, even
without inequality of material. Once a curve in the bank is started,
it is increased by the current which is directed into it. Furthermore,
as the current issues from the curve, it impinges against the opposite
bank and develops a curve at that point. The water issuing from this
curve develops another, and so on.
Once started, the curves or meanders tend to become more and more
pronounced (compare Figs. 180 and 181). In the case represented by Fig.
1, Plate XVI (Missouri River near Brunswick, Mo.) the narrow neck of
land between curves is almost cut through. When this is accomplished,
the stream will abandon its wide curve. A later stage in the process is
shown in Fig. 2, Plate XVI (the Osage River near Schell, Mo.).
The straightening of the channel is often accomplished in another way.
Even before the meanders reach the stage represented by Fig. 1, Plate
XVI, the position of the channel becomes unstable. In time of flood,
the whole flat is covered with flowing water. The greater depth of
water in the channel tends to give it a velocity greater than that of
the water on the flat outside. But the distance from _a_ to _c via b_
(Fig. 181) is much greater than that in a direct line. It follows that
the slope from _a_ to _c_ direct is greater than that by way of _b_. If
the current between _a_ and _c_ in time of flood be strong enough to
erode, it may deepen its bed, and thereby increase the volume of water
following this course. The increased volume gives increased velocity,
and the result may be the opening of a channel between _a_ and _c_
direct. The channel may be worn so deep that when the flood subsides,
the stream will follow it. So long as the abandoned channel-curve
remains unfilled with sediment, it is often called a _cut-off_. If
it contains standing water and has the proper form, it is called an
_ox-bow lake_ (Fig. 182), or sometimes a _bayou_. The water-filled
portions are not always bows (Fig. 183, Osage River, near Butler, Mo.).
Cut-offs, with or without standing water, are of common occurrence
along most rivers with wide plains. Meandering is not confined to
streams which are near sea-level. Even small creeks at high altitudes
may meander, if so situated as to have slight velocity. Trout Creek in
the Yellowstone Park (Fig. 184) is an example.
[Illustration: +Fig.+ 181.—Diagram illustrating later stages in the
development of meanders.]
[Illustration: +Fig.+ 182.—Meanders and cut-offs (ox-bow lakes) in
the Mississippi Valley a little below Vicksburg. The figure also
shows the migration of meanders down-stream, and their tendency to
increase themselves. (From charts Nos. 18 and 19, Mississippi River
Commission.)]
[Illustration: +Fig.+ 183.—Bayou Lakes, Osage River, near Butler, Mo.]
There seems to be some relation between the width of the belt within
which a stream meanders, and the width of the stream itself. Recently
it has been estimated that the ratio between them is 18 : 1.[72]
During the development of the meanders it is to be noted that lateral
planation on the one side of a stream is accompanied by deposition on
the other. This is _cut-and-fill_. The sediment eroded from the curve
which is concave toward the stream is shifted down-stream, while that
deposited in the curve which is convex toward the stream is brought
down from above. Thus even in the development of meanders, the material
which is dislodged is shifted down-stream. Since the current directed
against the down-stream side of a growing meander is on the average
stronger than that directed against the opposite side, the meander
itself has a tendency to migrate down-stream (Fig. 182).
In their evolution, the curves of a stream’s channel often reach and
undermine the valley bluff (Pl. VII). Since the meanders are, on the
average, shifted down-stream individually, and since meanders are
frequently developed in new places, it follows that a meandering stream
tends to widen its valley throughout. Widening is also effected in
other ways, for a stream with a flood-plain sometimes abandons its
channel altogether for miles at a stretch, and the new course chosen
may be against one of the bluffs of the valley. Such changes are most
likely to take place where deposition along channel and levees has
brought the part of the flood-plain (though not necessarily the bottom
of the channel) adjacent to the stream above the level of that farther
from it (Fig. 178). The change is likely to be effected in time of
flood.
Flood-plains often attain great size. That of the Mississippi below the
Ohio (Fig. 179) has a width ranging from rather more than 20 miles at
Helena (Ark.), to something like 80 miles in the latitude of Greenville
(Miss.).[73] Below the Ohio its area is something like 30,000 square
miles, and its entire area has been estimated at about 50,000 square
miles.[74]
Theoretically, the rotation of the earth should affect the erosion
of streams, increasing it on the one bank (the right in the northern
hemisphere and the left in the southern) and decreasing it on the
other.[75] The streams doubtless accommodate themselves to the rotation
of the earth in the original development of their gradation-plains
and flood-plains, and the later effects of rotation are usually
inconspicuous.
[Illustration: +Fig.+ 184.—Meanders of Trout Creek, Yellowstone Park.
(Walcott, U. S. Geol. Surv.)]
=Scour-and-fill.=—It has already been shown that aggrading streams cut
laterally at the same time that they build up their plains. It is now
to be added that they periodically deepen their channels to a notable
extent, and that the deepening of the channel takes place at the very
time when the flood-plain is being aggraded. In other words, the stream
in flood aggrades its plain, and degrades its channel. This follows
from the fact that the current is sluggish in the former position,
where the water is shallow, and rapid in the latter, where it is deep.
When the flood subsides, the channel, deepened while the current was
torrential, is filled again by the feebler current which follows. This
alternate deepening and filling is known as scour-and-fill. It is well
illustrated by the Missouri River. At Nebraska City, scour is believed
to occasionally reach depths of 70 to 90 feet.[76] At Blair, about 25
miles above Omaha, the same river is believed to cut to bed-rock (about
40 feet below the bottom of the channel in low water) twice a year,
that is, during floods.[77] Fig. 185 shows the changes recorded in the
channel of the river at this point during the year 1883. It shows that
the scour-and-fill during this year amounted to almost 40 feet. All
streams similarly situated do a like work. The material thus eroded
is shifted down-stream, some of it for short distances only, and some
of it to the sea. Even an aggrading stream therefore is not without
erosive activity; it is a stream whose fill exceeds its scour, not one
which has ceased to erode.
[Illustration: +Fig.+ 185.—Diagram illustrating scour and fill in the
Missouri River. A record of soundings at Blair Bridge (near Omaha),
1883. Shows also the cross-sections of the river at various rates.
(Todd, Bull. 158, U. S. Geol. Surv.)]
=Materials of the flood-plain.=—As a result of its varying velocities
in flood and low water, a stream may deposit coarse material at one
time and fine at another. A similar sequence of deposits takes place
in the flood-plain of a meandering stream, irrespective of floods.
Flood-plain deposits are often therefore very heterogeneous, as shown
in Fig. 186, which represents the constitution of the alluvium of the
Missouri River at Omaha. The deposits of the streams range from the
finest clay, through sand to gravel, and even bowlders. In general they
become finer down-stream. In a given plain, they are usually coarser
below and finer above.
[Illustration: +Fig.+ 186.—Diagram to show the heterogeneous character
of alluvial deposits. (Todd, Bull. 158, U. S. Geol. Surv.)]
=Topography of the flood-plain.=—The flood-plain is nearly, but not
altogether, flat. It has a gentle slope down-stream, and often for a
distance from the sides toward the center (Fig. 174). This latter slope
is the result of deposition by waters descending to the plain from the
sides. It is destroyed wherever a meandering stream reaches its bluffs.
When levees are well developed, there is a slope from them toward the
sides of the valley (Fig. 178), but it rarely continues to the limiting
bluffs. Since a stream with a well-developed flat frequently shifts
its course, old levees and abandoned channels lend variety to the
topography of the flood-plain.
=The topographic adjustment of tributaries.=[78]—The meandering and
shifting of a main stream affects its tributaries. If a main stream
swings against the bluff through which a tributary enters, the latter
brings its channel into topographic adjustment by lowering its end to
the level of the main. If now the main stream opposite the tributary
swings to the other side of its valley, the tributary must make its
way across the flat with a very low gradient. Not only this, but the
flat of the main valley through which the tributary must flow is likely
to be aggraded by the main in time of flood. The result is that the
tributary stream becomes an aggrading stream at its debouchure, and
topographic adjustment is not established until it has filled up the
lower end of its valley to some notable extent. The filling of the
lower end of the tributary likewise affects the lower ends of its lower
tributaries.
[Illustration: +Fig.+ 187.—A general view of the Mississippi delta.]
If the main stream again swings over to the point where the tributary
issues from its valley, the tributary stream and all its affected
tributaries again become eroding streams. Thus scour-and-fill are not
confined to the valley of the main stream.
=River-lakes.=—While rivers are in general hostile to lakes, they
sometimes give origin to them. Oxbow lakes (Fig. 182 and Pl. XVI),
due to the cut-offs of meandering streams, have already been referred
to. Lakes formed in the same way have other forms (Pl. XVI and Fig.
183). Rivers also give rise to lakes through the deposits they make.
If a main stream obstructs its tributaries by deposition at their
debouchures, their lower courses are ponded and converted into lakes.
The lakes along the tributaries to the Red River of Louisiana have
already been cited as examples. If a tributary brings more load to its
main than the latter can carry away, the detritus constitutes a partial
dam, ponding the river and causing it to expand into a lake above. Such
is the origin of Lake Pepin already referred to. In mountain regions,
the alluvial cones of tributary valleys sometimes pond their mains.
[Illustration: +Fig.+ 188.—Longitudinal section of an incipient delta
made of coarse material.]
Rivers may be dammed in other ways, as by lava flows, by landslides, by
glacial drift, etc. In all such cases, lakes may come into existence,
but they are not due primarily to the activity of the river itself.
=Deltas.=[79]—Where a stream enters standing water, or a slower stream,
a special form of plain, _the delta_, is sometimes built up (Figs. 169,
187, and 188). Deltas and alluvial fans have much in common, and their
only notable differences are those imposed by the differences in the
conditions of deposition. The current of the stream is checked, but not
altogether stopped, at its immediate debouchure. If it carries abundant
sediment, much of it will be promptly dropped where the decrease in
velocity is first felt. Such flow as there is beyond the debouchure is
not confined to a definite channel, and the deposits made are therefore
spread more or less on either side of the line which represents the
continuation of the stream’s course.
As the depth of the water into which the stream flows increases, the
current diminishes. Out to the point where the depth of the standing
water is less than the depth of the current, the latter affects the
bottom, and the surface of the deposits made slopes gently seaward; but
where the depth of the standing water is such that the projected stream
current is ineffective at the bottom, all the load rolled along the
bottom is dropped, and a depositional slope is established (Fig. 188),
its upper edge being below sea-level by an amount corresponding roughly
to the depth of the current which brings the detritus. The outer slope
is relatively steep and well-defined where the detritus is coarse, and
relatively gentle and ill-defined where it is fine. Thus the stream
tends to construct a sort of platform in the water just beyond its
debouchure. The successive deposits on the outer abrupt slope will dip
conformably with its surface (Fig. 189). The finest sediment will be
carried beyond the steep slope, and conform to the topography of the
bottom beyond (_c_, Fig. 189).
At the beginning, the top of the delta platform is at the level of the
bottom of the stream’s channel at the point of debouchure, but it is
gradually aggraded as water continues to flow over it. Its landward
margin is presently built up to sea-level and then above it, and as the
delta grows the delta-land is extended seaward (compare Figs. 188 and
189). At the same time the channel of the stream above the original
head of the delta is aggraded, for the current there is checked by the
aggradation of the delta. Thus alluvial deposits continuous with the
delta are extended landward.
[Illustration: +Fig.+ 189.—Longitudinal section of the delta at a later
stage of development.]
The projection of the direction of the lower end of the stream may be
said to be the axis of the extra-debouchure current. From this axis,
where the flow is strongest, the movement diverges more or less to
right and left. Since the velocity of the diverging water is reduced
more rapidly than that of the water which follows the axis of flow,
deposition is likely to take place faster on either side of the axis
than along the axis itself. The result is that the extra-debouchure
current tends to build up levee-like ridges on either side, making a
sort of sluice-way for itself. This sluice-way is gradually extended
seaward, and at the same time gradually filled. As its capacity is
reduced, more and more of the water flows over its sides. Presently
the escape of the water over the little side-levees will develop a
break at some point, and a line of distinct flow then diverges from the
main current. This distributary repeats the history of its main. Thus
the processes of levee-building, channel-filling, and levee-breaking
follow one another, until some such system of currents as shown in
Fig. 190 is developed. The result is that a delta’s growth is not
simply in the line of extension of the main stream, but in a more or
less semicircular area, the center of the circle being a point slightly
below the position of the debouchure of the stream when the delta
began. At any stage in its development the margin of the delta is more
or less crenate (Fig. 191), or characterized by delta fingers (Fig.
190), the projections corresponding to the positions of the debouchures
of the latest streams flowing across it. The extreme ends of the
delta lobes (Fig. 190), and of groups of the delta fingers, often
have something of the shape of the Greek letter from which the name
originated, but the resemblance in form between a well-developed delta
and the Greek letter is not striking. Deltas are sometimes built in
bays, and in such cases their forms are predetermined on all sides but
one. The head of a delta is sometimes arbitrarily located at the point
where the first distributaries are given off. Since this point shifts
widely with time, the definition can hardly be accepted. On this basis
the head of the Mississippi delta is about 200 miles above its lower
end. In reality it is much farther north.
[Illustration: +Fig.+ 190.—The terminus of the delta of the
Mississippi. (C. and G. Survey.)]
[Illustration: +Fig.+ 191.—A miniature delta.]
The structure of a delta, shown in Fig. 189, shows its history. At any
stage in its growth the river discharges its sediment across that part
of the platform already built. The sediment rolled at the bottom of
the current is dumped on reaching the steep slope, and constitutes
the inclined _fore-set_ beds shown in Fig. 189. The material in
suspension is carried farther, settles more gradually, and constitutes
the _bottom-set_ beds (_c_, Fig. 189). In time the bottom-set beds,
originally deposited some distance beyond the debouchure, may come to
be overlain by the fore-set beds, deposited at a later time. While the
fore-set beds are being deposited on the steep slopes of the delta, and
the bottom-set beds beyond, deposition is also taking place on the top
of the delta. These _top-set_ beds are laid down in a nearly horizontal
position, and their seaward margin is gradually extended. Thus the
delta comes to have the threefold structure shown in Fig. 189.
That part of the delta which is above the abrupt slope of its front
corresponds in all essentials to an alluvial fan; but the delta as a
whole differs from the fan in its abrupt and crenate or digitate margin.
It is to be noted that the delta is not wholly the product of a
stream’s activity. The stream supplies the material, but the lake
or sea renders at least passive assistance in its disposition. Not
all rivers opening into the sea build deltas, and their failure is
often the result of waves or shore currents which carry off the river
sediment. Deltas are, however, sometimes formed in tidal seas, as at
the debouchures of the Yukon; the Mackenzie, where the tidal range is
three feet; the Niger, where the range is four feet; the Hoang-Ho,
where the range is eight feet; and the Brahmaputra and Ganges, where
the range is sixteen feet.[80] Since lakes, bays, gulfs, and inland
seas have weaker waves and currents than the open sea, they are more
favorable than the latter for the growth of deltas. Hence occur such
deltas as those of the Mississippi, the Nile, the Po, and the Danube.
Deltas are likely to be absent, or confined to the heads of bays, on
coasts which have recently sunk. Their general absence on the Atlantic
coast of the United States is a case in point.
The following figures give some idea of the extent of deltas, and of
their importance in land building. The Mississippi delta is advancing
into the Gulf at the rate of about 100 yards per year, or a mile in
16 or 17 years. Its length is more than 200 miles, its area more than
12,000 square miles, and its depth at New Orleans has been estimated
at 700[81] to 1000[82] feet. This great depth is believed to be the
result of subsidence, and so of the superposition of one delta on
another.[83] The delta of the Yukon has a sea margin of 70 miles, and
extends more than 100 miles inland. The delta of the Rhône has also
had a remarkable growth, considering the size and the history of the
stream. Arles, near the debouchure of the stream, was 14 to 16 miles
inland in the fourth century +B.C.+, and is now 30 miles inland.[84]
The Rhône has also built a great delta in Lake Geneva, and its lower
delta is built of sediment gathered below the lake. The Po has built
a delta 14 miles beyond Adria, the port which gave its name to the
Adriatic Sea. The extension of this delta has been at the average rate
of about 50 feet per year, but recently, on account of artificial
embankments, the rate has been much more rapid.[85] The Ganges and
Brahmaputra together have made a delta of great size. Its area is
sometimes estimated to be as high as 50,000 or 60,000 square miles,
and its head is more than 200 miles from the sea.[86] The head of the
Nile delta is 90 miles from the sea, and it has a coastal border of
180 miles. The head of the delta of the Hoang-Ho is about 300 miles
from the coast, and its seaward border has a length of about 400 miles,
though with some highland interruptions.[87]
After a delta has been built into a lake, the lake may disappear,
leaving the delta out of water. Such “fossil” deltas, if so recently
exposed that erosion has not destroyed their distinctive features,
are readily recognized by their flat tops, their abrupt and lobate
fronts, and their characteristic structure. They are often a means
of determining the former existence of extinct lakes,[88] or the
former higher levels of lakes which still exist.[89] Elevated deltas
on seashores show either a rise of the land or a depression of the
sea-level.
The material which is carried along the coasts or shores from the
mouths of rivers may take on various and peculiar forms, according
to the strength, direction, and relations of waves and currents. The
consideration of these forms belongs more properly to the work of the
sea than to that of rivers, since rivers are not concerned in their
construction except in supplying material.
=Delta lakes.=—Delta-building streams sometimes help to form lakes by
throwing their deposits around an area which fails to be aggraded to
sea-level. Lake Pontchartrain, and other lakes in the delta of the
Mississippi are examples (Fig. 187).
[Illustration: +Fig.+ 192.—Terraces of the Frazier River at Lillooet,
B. C. (Calvin.)]
STREAM TERRACES.
Stream terraces[90] are bench-like flats or narrow plains along the
sides of valleys (Fig. 192). They are usually narrow, but sometimes
have great length in the direction of the axis of the valley. They
originate in various ways.
=Due to inequalities of hardness.=—Reference has already been made
(p. 140) to the effect of hard horizontal layers in the development
of terraces and terraciform projections on the sides of valleys (Fig.
120). Such terraces are the result of differential degradation, and
the upper surface of the hard layer marks the lower limit of the
terrace, which commonly has a distinct slope toward the stream. Except
where interrupted by tributary valleys, such terraces are likely to
be continuous in a valley so long as the structure remains the same
and the stream sustains the same relation to it. Such terraces would
first show themselves in the older part of the valley. The effect of
inclination of the hard stratum on the development of such terraces
will be readily inferred. Terraces and benches of this sort are not
equally distinct at all stages of a valley’s history. For great
distinctness, the hard layer should have been exposed long enough to
allow the general processes of erosion to have effected considerable
differential wear, but not long enough to allow the topographic effects
of unequal resistance to be obliterated.
[Illustration: +Fig.+ 193.—Diagram illustrating a distinct terrace and
a “second bottom (_b_),” which may be regarded as a low terrace.]
=Normal flood-plain terraces.=—It has been seen that deposition in
a river valley stands in more or less definite relationship to the
stage of its development, and that the deposition which leads to
the development of an alluvial plain is likely to take place where
the higher gradient of the upper course gives place to the gentler
gradient of the lower. It has also been seen that as a stream’s history
advances, the stretch where the gradient is high recedes up-stream,
and that the point which marks the head of active deposition follows.
It follows that a river flat or flood-plain normally begins in the
lower part of a valley, and works progressively headward, its upper end
following, at some considerable distance, the head of the valley itself.
The commoner river terraces are remnants of former flood-plains,
below which the streams which made them have cut their channels. It
has already been pointed out (p. 184) that processes of erosion and
deposition work together in the development of flood-plains, and that
some flood-plains have but little alluvium (Fig. 174), while others
owe their origin wholly to stream deposits (Fig. 175). It follows that
terraces developed from flood-plains may be of rock, of alluvium, or of
rock covered with alluvium.
The amount which a river channel must be deepened in order to change
the remnants of its flood-plain to terraces cannot be definitely
stated. When a channel is so deep that the remnants of a former
flood-plain are no longer flooded, they would be called terraces,
especially if a lower flood-plain has been developed. Even though not
above the reach of floods, they are often called terraces if they are
notably above the channel and separated from it by a lower plain. Thus
the flat at _b_, Fig. 193, would be called a terrace, even though
covered by water in exceptional floods; but the flat at _c_, but
slightly above the channel, would hardly be called a terrace.
[Illustration: +Fig.+ 194.—Diagram illustrating the beginning of the
development of a terrace from a flood-plain.]
The question now arises why a stream, having once developed a
flood-plain, should sink its channel to a lower level, leaving parts
of the old flood-plain as terraces. This may be brought about by the
operation of various causes.
(1) In the first place, the head of the valley-plain where the first
notable deposition takes place normally advances up-stream. After the
advance has been considerable, the descending stream may, on reaching
the head of its valley-plain, lose so much of its load as to be able
to sink its channel into the flood-plain farther down the valley (Fig.
194).
(2) Ordinarily a stream does not drop all its load at the head of its
plain, but only its excess; but it will always drop coarse sediment
to take fine, if fine be available. For a relatively small amount of
coarse material dropped, a relatively large amount of fine may be
taken up (p. 179). Other things being equal, it follows that when a
stream drops coarse material to take fine, its channel is degraded
unless there is at the same time a great reduction in the stream’s
energy. Such reduction is likely to go with the decreasing declivity
down-stream; but this is partly, or sometimes wholly, counterbalanced
by the increasing volume of water. By the exchange of load, therefore,
a stream may ultimately sink its channel below the flood-plain which
the earlier and perhaps smaller stream had developed.
[Illustration: +Fig.+ 195.—Diagram to illustrate the development of
river terraces by the widening of a channel or meander belt. The
valley flat above might not be called a terrace; but the same plain
below, where the meander belt has some width, would be called a
terrace.]
(3) Again, so long as a stream is actively eroding at its head, there
is likely to be some aggradation below. At a later stage in the
stream’s history, when active erosion at the head has ceased because
of the reduction of the surface, less material will be carried from
the upper part of the valley, and the stream on the flood-plain below,
formerly loaded with material from up the valley, is now free to take
up and carry away material temporarily left on the flood-plain. The
result is a deepening of the channel.
(4) Any stream which has reached the flood-plain stage is likely to
meander. After the flood-plain has become wide, the width of the belt
within which the stream meanders is less than the width of its plain.
In the Lower Mississippi, for example, the meander belt is often no
more than a third to a tenth of the width of the flood-plain. It has
already been pointed out that the meanders migrate down the valley. In
so doing they depress the meander belt, the tendency being to reduce
it to the level of the channel, and, therefore, below the level of
the flood-plain. As the meander belt widens, the depression which it
develops becomes more and more capacious. Presently it may attain such
dimensions as to hold the water of ordinary floods. At this stage,
or even before, such parts of the earlier flood-plain as remain, are
terraces.
These several tendencies conspire to partially destroy river
flood-plains, and to transform such parts as remain into terraces _in
the normal course of a river’s history_. They appear first in the
lower part of the valley, and migrate headward, following the course
of nearly every other phase of activity in a stream’s history. The
heads of the terraces follow, at a respectful distance, the head of the
flood-plain, just as the head of the flood-plain follows at a distance
the head of the valley. The second and subsequent flood-plains and the
terraces to which they give origin follow the same course.
Terraces developed by the normal activities of a stream are always low,
and it is improbable that they would ordinarily be conspicuous. The
vertical distance between the first (highest) and second is greater
than that between the second and third. The principles developed on
page 65 et seq., in connection with the erosion of the hypothetical
island, are applicable here.