readily be seen that the bottom of a glacier may be charged with
various sorts of material. There may be (1) bowlders which the ice has
picked up from the surface, or which it has broken off from projecting
points of rock over which it has passed; (2) smaller pieces of rock
of the size of cobbles, pebbles, etc., either picked up by the ice
from its bed or broken off from larger masses; (3) the fine products
(rock-flour) produced by the grinding of the débris in the ice on the
rock-bed over which it passes, and similar products resulting from the
rubbing of stones in the ice against one another; and (4) sand, clay,
soil, vegetation, etc., derived from the surface overridden. Thus
the materials which the ice carries (_drift_) are of all grades of
coarseness and fineness, from large bowlders to fine clay. The coarser
material may be angular or round at the outset, and its form may be
changed and its surface striated as it is moved forward. Whether one
sort of material or another predominates, depends primarily on the
nature of the surface overridden.
[Illustration: PLATE XIX.
PART OF THE BIGHORN MOUNTAIN RANGE, WYOMING.
U. S. Geol. Surv.]
[Illustration: PLATE XX.
A SECTION OF THE CALIFORNIA COAST NEAR SAN MATEO, CALIFORNIA.
U. S. Geol. Surv.]
[Illustration: +Fig.+ 257.—A glacial cirque. The lake occupies a rock
basin, produced by glacier erosion. Head of Little Timber Creek,
Montana.]
=The topographic effects of glacial erosion.=—In passing through its
valley, an alpine glacier deepens and widens its bottom and smooths
its slopes up to the upper limit of the ice. It tends to change a
V-shaped valley (Fig. 258) into a U-shaped one (Fig. 259). The change
in topography at the upper limit of glaciation is often marked (Figs.
260 and 261).
[Illustration: +Fig.+ 258.—The Valley of the American Fork. A V-shaped
non-glaciated valley in the Wasatch Mountains of Utah. Compare Fig.
259. (Church.)]
[Illustration: +Fig.+ 259.—U-shaped valley resulting from glaciation.
Little Cottonwood Canyon, Wasatch Mountains. (Church.)]
[Illustration: +Fig.+ 260.—Contrast between glaciated topography below
and non-glaciated topography above. The minarets in the Sierras, Cal.]
[Illustration: +Fig.+ 261.—Contrast between glaciated topography below,
and non-glaciated topography above. Needles Mountains, from slope
west of Hidden Lake. (Cross, U. S. Geol. Surv.)]
The deepening of a valley by glacial erosion may throw its tributaries
out of topographic adjustment. Thus if a main valley is lowered 100
feet by glacial erosion while its tributary is not deepened, the
lower end of the latter will be 100 feet above the former when the
ice disappears. Such a valley is called a _hanging valley_ (Figs. 262
and 263). Such valleys are of common occurrence in regions recently
glaciated, but now ice-free. Examples are common in the western
mountains of North America and elsewhere.
Ice-caps which overspread the surface irrespective of valleys and
hills, tend to reduce the angularities of the surface. Hills and ridges
are cut down and smoothed (Figs. 264 and 265); but since valleys
parallel to the direction of movement are deepened at the same time,
it is doubtful if the relief of the surface is commonly reduced by the
erosion of an ice-cap.
[Illustration: +Fig.+ 262.—A hanging valley. East side of Lake
Kootenai, B. C. All except the highest summits glaciated. (Atwood.)]
=Fiords.=—A glacier descending into the head of a narrow bay may gouge
out the bay to a very considerable depth, causing its head to recede.
When the ice finally melts, the bay may be a fiord. Thus have arisen
the glacial features of many of the fiords of high-latitude coasts, and
many of the glaciers of those coasts are now making fiords (Fig. 266).
Fiords also arise in other ways. Coasts indented by fiords are likely
to be bordered by islands.
=The positions in which débris is carried.=—As a result of the methods
by which a glacier becomes loaded, there are three positions in which
the débris is carried: (1) the basal or subglacial, (2) the englacial,
and (3) the superglacial. The material picked up or rubbed off from
the surface over which the ice moves is normally carried forward in
the base of the ice; while that which falls on the surface is usually
carried in the form of surface moraines. In the former position the
drift is basal; in the latter, superglacial. It is doubtful if much
débris is moved along beneath (that is, strictly below the bottom of)
the ice, though the movement of the latter would have a tendency to
drag or urge along with it the loose material of its bed. If drift were
carried forward in such positions, it would be strictly _subglacial_.
[Illustration: +Fig.+ 263.—A hanging valley. The water falls (Bridal
Veil) from a hanging valley. (Wineman.)]
The basal load of a glacier is constantly being mixed with new
accessions derived from ground over which the ice is passing, and
this admixture tells the story of the work done by the bottom of the
ice. The englacial and superglacial material, on the other hand, is
normally borne from the place of origin to the place of deposition
without such intermixture. It is a case of “local” versus “through”
transportation.
=Transfers of load.=—While the origin of the load usually determines
its position, exceptions and complications arise from the transfer
of load from one position to another, and from the gradation of one
horizon into another.
[Illustration: +Fig.+ 264.—A non-glaciated hill. Dalrymple Island.
North Greenland.]
[Illustration: +Fig.+ 265.—A glaciated hill. Southeastern Carey Island.
About 30 miles west-northwest of Dalrymple Island.]
Most of the débris gathered by ice is acquired at its bottom. While
such material is basal _at the outset_, some of it may find itself
above the bottom a little later. Thus when ice passes over a hill (Fig.
267) the bottom of the ice rends débris from the top of the hill. When
it descends from one level to another there is a similar result (Fig.
268). To the lee of the hill the ice from either side may close in
under that which came over the top, in which case the débris derived
from the top of the hill by the bottom of the overriding ice will be
well up in the ice. It has passed from an initial basal to a subsequent
englacial position. The change does not usually involve an actual rise
of the material, but rather a decline. If carried upward at all, the
upward movement is temporary only, and incident to the passage of the
ice over the hill, or to other local causes. The englacial débris may
be little or much above the basal zone according to the height of the
elevation overridden.
[Illustration: +Fig.+ 266.—Alaskan fiords. The shaded areas represent
land. (From charts of the C. & G. Surv.)]
[Illustration: +Fig.+ 267.—Diagram to illustrate the taking of débris
from a hill-top. It also illustrates how englacial débris may become
superglacial as the result of surface ablation.]
[Illustration: +Fig.+ 268.—Taking débris from a protuberance of the
bed.]
Superglacial débris may obviously become englacial by falling into
crevasses or by being carried down by descending waters. Either
superglacial or englacial débris may become basal by the same means.
From their form and position, there is less ice-free land in immediate
association with ice-caps than with valley glaciers. Furthermore, the
ice-free land about the borders of an ice-cap is less likely to be in
the form of cliffs above it. As a result, the surfaces of ice-caps are
comparatively clean, except at their edges where the ice is thin.
[Illustration: +Fig.+ 269.—Side view of end of a glacier on the south
side of Olriks Bay, North Greenland.]
[Illustration: +Fig.+ 270.—Closer view of a part of the ice shown in
Fig. 269.]
Englacial material may become superglacial by surface ablation. In this
case the drift does not rise, but melting brings the surface of the ice
down to its level. This occurs chiefly at the end or edge of the ice,
where the surface melting is greatest. Englacial débris, especially
that near the bottom, may also become basal by the melting of the
bottom of the ice.
Englacial material plucked or rasped from an elevation over which the
ice has passed is liable to be disposed in a longitudinal belt in
the ice in the lee of the elevation itself. By surface ablation this
material may reach the surface at some point below its source, and
be disposed as a medial moraine. Such a moraine has an origin very
different from that of a medial moraine formed by the junction of two
lateral moraines of superglacial origin.
Much less in the natural order of things is the transfer of material
from a basal to an englacial and from an englacial to a superglacial
position by upward movement of the débris itself. Such transfer is
remarkable because the specific gravity of rock is from two and a half
to three times as great as that of ice, so that its normal tendency is
to sink.
[Illustration: +Fig.+ 271.—Surface terminal moraines due to upturning.
Edge of the ice-sheet, North Greenland.]
In arctic glaciers, and probably in others, some material which has
been basal becomes englacial by being sheared forward over ice in front
of it. So far as observed this takes place chiefly where the ice in
front of the plane of shearing lies at a lower level than that behind,
as where the surface of an upland falls off into a valley, or where
a boss of rock shelters the ice in its lee from the thrust of the
overriding ice (Fig. 268).
[Illustration: +Fig.+ 272.—Diagram illustrating the upturning of the
layers of ice at the end of an arctic glacier as seen in end-section.
The bottom line represents sea level.]
At the borders of arctic glaciers the lower layers are not infrequently
upturned, as shown in Figs. 269 to 272. Where the layers turn up at
the end of a glacier (Figs. 269 and 270), basal and englacial débris
is carried to the surface by actual upward movement, and a terminal
moraine or a series of terminal moraines sometimes aggregated _where
the upturned layers of ice outcrop at the surface_ (Fig. 271). That
the material of these moraines was originally basal is abundantly
demonstrated by the bruised and scratched condition of the bowlders
and pebbles, and sometimes by the nature of the material itself. For
example, in two cases in North Greenland where glaciers descend into
the heads of shallow bays and move forward on their bottoms, moraines
formed by the upturning of the layers were seen to contain _abundant
molluscan shells derived from the bottom of the bay_. The upturning
sometimes affects the side-edges of ice-tongues (Fig. 272) as well as
their ends, and the material thus brought to the surface gives origin
to lateral moraines altogether different in origin from the lateral
moraines formed by the falling of débris upon the glaciers. Sometimes
also there is an upturning of the ice along a longitudinal zone well
back from the lateral margins (Fig. 273), and the material so borne
to the surface in such a zone gives rise to a moraine resembling the
medial moraine formed by the union of lateral moraines, but of wholly
different origin.
[Illustration: +Fig.+ 273.—Diagram illustrating the same point as 272,
where the structure is more complex. The bottom line of the figure
represents sea level.]
The phenomenon of upturning here referred to has been observed only at
or near the terminus of the ice, and is perhaps due in most part to the
resistance of frozen morainic or other material beneath and in front of
the edge. To this should probably be added the effect of the increased
rigidity of the ice at its borders, due to the low external temperature
during the larger part of the year, while the interior, with its
higher temperature, remains more fluent. But even this probably leaves
the explanation inadequate. In not a few instances the upturning is
associated with a notable _thickening_ of the layers toward their
edges (Fig. 274). This suggests that perhaps there is an exceptional
growth of the granular crystals of the ice near the edge of the
layers, owing to the penetration of the surface-waters which are much
more abundant at the borders than elsewhere, and which in the arctic
glaciers probably do not penetrate deeply before they reach a freezing
temperature.
=Wear of drift in transit.=—Drift carried at the bottom of the ice is
subject to notable wear. The materials in transportation abrade one
another and are abraded by the bed over which they pass. Englacial
drift is subject to less wear because it is commonly more scattered.
Superglacial drift is worn little or none while it lies on the surface
of the ice; but in so far as superglacial or englacial drift is derived
from the basal load, it may show the same evidences of wear as the
basal drift itself. Superglacial drift often reveals its history in
this way.
[Illustration: +Fig.+ 274.—Thickening of the upturned layers of ice.]
_Deposition of the Drift._
1. =Beneath the body of the ice.=—During the advance of a glacier,
deposition may take place both beneath the body of the ice and beneath
its end and edges. Deposition beneath the body of the ice is liable to
take place wherever the topography favors lodgment, or wherever the ice
is overloaded. The topography favoring deposition is much the same as
that favoring erosion, but the two processes are not favored at the
same point. Erosion is greatest on the “stoss” side of an obstruction
(the side against which the ice advances), and deposition on the lee
side. The ice is likely to be overloaded (1) just beyond a place where
conditions have favored the gathering of a heavy load, and (2) where
the ice is rapidly thinning. On the whole, however, the deposition of
material beneath the main body of a glacier is much more than balanced
by erosion in the same position.
[Illustration: +Fig.+ 275.—Glacier building an embankment. Southeast
side of McCormick Bay, North Greenland.]
2. =At ends and edges of glaciers.=—At and near the end of a glacier
the conditions of deposition are somewhat different. Here deposition
beneath the ice goes on faster than elsewhere, chiefly because of the
more rapid melting and the more rapid thinning and weakening of the
ice. If the end of the glacier be stationary in position, drift is
being continually brought to it and left there, for though the _end_ is
stationary, the _ice_ continues to move. If the glacier moves forward
500 feet per year, and if its end is melted at the same rate, all the
débris in the 500 feet of ice which has been melted has been deposited,
and all except that which has been washed away has been deposited
at and beneath the end of the glacier. If the end of the glacier
is retreating, the retreat means that the waste at the end exceeds
the forward movement. If the ice advances 300 feet per year, and is
melted back 500 feet in the same time, all the débris carried by the
500 feet which has been melted has been deposited, and largely in the
narrow zone (200 feet) from which the ice has receded. Even in this
case, therefore, there is a notable tendency to marginal accumulation.
If the end of the glacier is advancing 500 feet per year while it is
being melted but 300 feet, all the drift in the 300 feet melted has
been deposited, and chiefly at or beneath the immediate margin of the
ice. To the marginal and sub-marginal accumulations made in this way,
the material carried on the ice is added whenever the ice is melted
from beneath it. This addition is sometimes considerable and sometimes
meagre. If the edge of the ice is without much fluctuation in position,
the material dumped over its end may take the form of a narrow ridge or
bowlder-wall (_Geschiebe-wall_). If a glacier pushes material in front
of it, this, too, becomes a part of the general terminal aggregation of
drift.
[Illustration: +Fig.+ 276.—Embankment completed. Near the last.]
TYPES OF MORAINES.
=The terminal moraine.=—The thick accumulation of drift made at the
end of a glacier or at the edge of an ice sheet, especially where its
end or edge is stationary, or nearly stationary, for a considerable
time, is the _terminal moraine_. That part of the aggregation deposited
beneath the ice is sometimes called the _lodge_ moraine (Figs. 275 and
276; see also Fig. 235); that carried on the ice and dropped at its
edge, the _dump_ moraine; and that pushed before the ice, the _push_
moraine. Many moraines marginal to the ice appear to be push moraines,
when they are really lodge moraines from which the ice has withdrawn
(Fig. 277). The push moraine can rarely be distinguished, and the dump
moraine by no means always, after the disappearance of the ice.
[Illustration: +Fig.+ 277.—End of a glacier a few miles west of
Kaslo on Lake Kootenai, B. C. A lodge moraine from which the ice has
withdrawn, giving it the appearance of a push moraine. It is possible
that the lodge moraine material has been pushed up a little by
re-advance of the ice. (Atwood.)]
=The ground moraine.=—When a glacier disappears by melting, all its
débris is deposited. All the drift deposited beneath the advancing
ice and all deposited from the base of the ice during its dissolution
constitutes the _ground moraine_. The thickness of the ground moraine
is notably unequal. In general, it is thicker toward the terminus of
the glacier and thinner toward its source, but considerable portions of
a glacier’s bed are often left without débris when the ice melts. In
general, the terminal moraine is not only thicker, but more irregularly
disposed than the ground moraine.
=The lateral moraines.=—The surface lateral moraines of valley glaciers
are let down on the surface beneath when the ice melts out from under
them; but the lateral moraines in a valley from which the ice has
melted are not merely the lateral moraines which were on the glacier
at a given time. They are often far more massive than any which ever
existed on the ice itself at any one time (Fig. 278). As a glacier
retreats, its lateral moraine material is more or less bunched. Thus if
the ice advances 200 feet while its end is being melted back 300 feet,
the lateral moraines on the 300 feet melted are concentrated into 100
feet, as they are delivered on to the land by the melting of the ice
from beneath. If the retreat of the end of a glacier be very slow, the
bunching may be great. But even this cannot explain the massiveness of
some lateral moraines. Furthermore, the materials of which many lateral
moraines are composed are nearly as well worn as those of the ground
moraine. The massive lateral moraines of which this is true are often
made up chiefly of the drift accumulated beneath the lateral margins
of the glaciers. This accumulation is the result of the lateral motion
of the ice from center to side. Such sublateral accumulations are akin
to terminal moraines. Some of the lateral moraines of ancient valley
glaciers, such as those of the Uinta, Wasatch, and Bighorn mountains
are several hundred feet high, and in one case about 1000 feet. In
northern Italy lateral moraines are said to be 1500 to 2000 feet
high.[132]
[Illustration: +Fig.+ 278.—A lateral moraine from which the ice has
retreated. Bighorn Mountains, Wyo. (Blackwelder.)]
[Illustration: +Fig.+ 279.—Glacial drift, coarse and fine together.
(Geol. Surv. of N. J.)]
Most of the material which was englacial during the transportation
becomes either subglacial or superglacial before deposition, for it
ordinarily reaches the bottom or the top of the ice before being
deposited. Where the ends or edges of a glacier are vertical or nearly
so, as in the high arctic regions, deposition may take place from the
englacial position directly.
=Distinctive nature of glacial deposits.=—The deposits made by
glaciers are distinctive. In the first place the ice does not assort
its material, and bowlders, cobbles, pebbles, sand, and clay are
confusedly commingled (Fig. 279). In this respect, the deposits of ice
differ notably from those of water. Furthermore, many stones of the
drift show the peculiar type of wear which glaciers inflict. They are
not rounded as the stones carried by rivers, though they are notably
worn. Many of them have subangular forms with planed and beveled
faces, the planes being striated and bruised (Fig. 254). The absence
of stratification, the physical heterogeneity, and the striation
of at least a part of the stones are among the most distinctive
characteristics of glacial drift. A not less real though less obvious
characteristic is the constitution of the fine material, for it is in
general not the product of rock decay, but of rock grinding. The fine
material handled by streams (except glacial streams) on the other hand,
is usually the product of rock decay.
[Illustration: +Fig.+ 280.—Roche moutonnée, Victoria Harbor, B. C.]
=Glaciated rock surfaces.=—Another distinctive mark which a glacier
leaves behind it is the character of the surface of the rock on which
the drift rests. This is generally smoothed by the severe abrasion to
which it has been subjected, and the smoothed surfaces are marked by
grooves and striæ, similar to those on the stones of the drift (Figs.
255 and 256). Other distinctive features of a glaciated area are the
rounded bosses of rock (_roches moutonnées_, Fig. 280; see also surface
about the lakes, Fig. 261), the rock basins, the lakes (Fig. 261),
ponds, and marshes, and the peculiar topographies resulting from the
unequal erosion, and the still more unequal deposition of the drift.
Surface bowlders, often unlike the underlying formations of rock, and
sometimes in peculiar and apparently unstable positions, are still
another mark of a glaciated area.
GLACIO-FLUVIAL WORK.
The constant but unequal waste of glaciers has already been referred
to. The streams to which this gives rise are usually laden with
gravel, sand and silt derived from the ice. Since the mud is often
light-colored, the streams are sometimes described as “milky.” Where
the amount of material carried is great, much of it is dropped at
a slight distance from the ice, the coarsest being dropped first.
Glacial streams are, as a rule, aggrading streams, and therefore
develop alluvial plains, called _valley trains_ (Fig. 281 and 282),
or where they enter lakes (Fig. 283), bays, or other streams, deltas.
In its transportation, the river-borne drift is assorted; after its
deposition, it is stratified. True glacial deposits in the upper
part of a mountain valley are, therefore, often continued below by
glacio-fluvial deposits derived from the same source.
[Illustration: +Fig.+ 281.—Alluviation by glacial stream: Nicolai
Creek, Alaska. (Schrader, U. S. Geol. Surv.)]
The most common form of such deposit is a _valley train_ (Fig. 281) of
glacial wash stretching indefinitely down the valley. The silt, sand,
and gravel of such trains can usually be distinguished from valley
deposits of non-glacial origin by the character of the material, as
much of it is the product of grinding, crushing, and fracture, rather
than of ordinary surface decay. Its materials are, therefore, fresh
and often include rock material which, if long exposed at the surface,
would be decomposed or dissolved.
[Illustration: +Fig.+ 282.—Alluviation by glacial stream below Hidden
glacier, Alaska. (Gilbert, U. S. Geol. Surv.)]
Where an ice sheet ends in a broad face, as did the ancient continental
glaciers, numerous streams flow from it and spread their débris in
front of the terminal moraine, forming a broad fringing sheet or
“apron” (_outwash plain_) along it. Where streams of considerable size
form tunnels under or in the ice, these may become more or less filled
with wash, and when the ice melts the aggraded channels appear as long
ridges of gravel and sand known as _eskers_ (_osars_ and _serpentine
kames_ and _kames_ of authors. See chapter on glacial period). It has
been thought that similar ridges are sometimes formed in valleys cut
in the ice from top to bottom, and even that they arise from gravel
and sand lodged in superglacial channels. The latter at least is
probably rare, as the surface streams usually have high gradients,
swift currents, and smooth bottoms, and hence give little opportunity
for lodgment. In the case of ice-sheets, too, in connection with which
eskers are chiefly developed, there is usually no surface material
except at the immediate edge, where the ice is thin and its layers
upturned.
At the mouths of ice-tunnels or ice-channels, especially where they
end against terminal moraines, sands and gravels are liable to be
bunched in quantity, giving rise, after the adjacent ice has melted,
to peculiar hills and hollows of the knob-and-basin type. The hills
and short ridges are known as _kames_ (see glacial period). Subglacial
streams may leave washed and assorted material in their tracks under
the ice, and this is sometimes buried under deposits made by the ice
itself, so that glacio-fluvial and glacial deposits are interbedded.
[Illustration: +Fig.+ 283.—Delta at Isola, Lake of Sils, Engadine,
Switzerland. (Reid.)]
ICEBERGS.
When glaciers advance into water, the depth of which approaches their
thickness, their ends are broken off (Fig. 284), and the detached
masses float away as icebergs (Fig. 285). Many of the bergs are
overturned, or at least tilted, as they set sail. If this does not
happen at the outset, it is likely to occur later as the result of the
melting and wave-cutting which disturb their equilibrium. The great
majority of bergs do not travel far before losing all trace of stony
and earthy débris, but the finding of glacial material in dredgings far
south of all glaciers shows that they occasionally carry stones far
from land.
[Illustration: +Fig.+ 284.—End of Muir glacier, Alaska. (Reid.)]
THE INTIMATE STRUCTURE AND THE MOVEMENT OF GLACIERS.
With the preceding account of glaciers in mind, we may return to a
closer study of their origin, their intimate structure, and their mode
of motion. The key to this study is the thesis that a glacier is a mass
of crystalline rock—the purest and simplest type of crystalline rock
known—since it is made up of a single mineral of simple composition
and rare purity, which never appears in a solid state except in the
crystalline form.
=The growth and constitution of a glacier.=—The origin and history
of a glacier is little more than the origin and aggregate history of
the crystals that compose it. The fundamental conception of a glacier
is therefore best obtained by tracing the growth of its constituent
crystals. A basal fact ever to be kept in mind is that water in
the solid form is always controlled by crystalline forces. When it
solidifies from the vapor of the atmosphere it takes the form of
separate crystals (Figs. 286–291). Perfect forms are developed only
when the flakes fall quietly through a saturated atmosphere which
allows them to grow as they descend. Under other conditions, the
crystals are imperfect in growth and are mutilated by impact. But
however modified, they are always crystals. The molecules are arranged
on the hexagonal plan, and, as the expansive power of freezing water
shows, the arrangement is controlled by a strong force. Once the
definite crystalline arrangement is established, the molecules can be
displaced only by relatively great force.
[Illustration: +Fig.+ 285.—An iceberg, west coast of Greenland.]
Snow crystals often continue to grow so long as they are in the
atmosphere; but if they pass through an under-saturated stratum of
air or a stratum whose temperature is above 32° Fahr., they suffer
from evaporation or melting. When they reach the ground, the processes
of growth and decadence continue, and the crystals grow or diminish
according to circumstances.
A glacier is a colossal aggregation of crystals grown from snowflakes
to granules of much greater sizes. The microscopic study of new-fallen
snow reveals the mode of change from flakes to granules. The slender
points and angles of the former yield to melting and evaporation more
than the more massive central portions, and this change probably
illustrates a law of vital importance. It may often be seen that the
water melted from the periphery of a flake gathers about its center,
and if the temperature be right, it freezes there. This is a first
step toward the pronounced granulation of snow which has lain for some
time on the ground. If measured systematically from day to day, the
larger granules taken from beneath the surface of this coarse-grained
snow are found to be growing. In a series of experiments[133] to
determine the law of growth it was found that when the temperature of
the atmosphere was above the melting-point the growth was appreciably
more rapid than when the air was colder, but there was, on the average,
_an increase under all conditions of temperature_. A portion of this
average increase of the larger granules appears to come from the
diminution and destruction of the smaller ones, for the total number
of granules steadily diminishes. A portion of the growth doubtless
comes from the moisture of the atmosphere which penetrates the snow
and another portion from the moisture derived from surface melting;
but beneath the surface of a large body of snow the growth of the
large granules is probably chiefly at the expense of the small ones.
To follow the process it should be noted that the free surface of
every granule is constantly throwing off particles of water-vapor
(evaporation); that the rate at which the particles are thrown off is
dependent, among other things, on the curvature of the surface, being
greater the sharper the curve; that the surfaces of the granules are
at the same time liable to receive and retain molecules thrown from
other granules, and that, other things being equal, the retention
of particles also depends on the curvature of the surface, the less
curved surface retaining more than the sharply curved one. Under these
laws, it is obvious that the larger granules of smaller curvature will
lose less and gain more, on the average, than the smaller granules of
greater curvature. It follows that the larger granules will grow at the
expense of the smaller. It is also to be noted that, other things being
equal, small granules melt more readily than large ones, and that where
the temperature is nicely adjusted between melting and freezing, the
smaller may lose while the larger gain.
[Illustration: +Figs.+ 286–91.—Snowflakes. (Photographed by W. A.
Bentley.)]
Another factor that enters into the process is that of pressure and
tension. The granules are compressed at the points of contact and put
under tension at points not in contact, and the pressure and tension
are, on the average, likely to be relatively greatest for the smallest
granules. Tension increases the tendency to evaporation and adds its
effects to curvature, and the capillary spaces adjoining the points
of contact probably favor condensation. Ice expands in crystallizing
and pressure reduces the melting-point, while tension raises it. The
effect of this is slight (p. 276), and it probably plays little part in
glacial action, but it is to be correlated with the much more important
fact that _compression produces heat_ which may raise the temperature
of the ice to the melting-point, while tension may reduce the
temperature to or below freezing. There is therefore a tendency for the
ice to melt at the points of contact and compression, and for the water
so produced to refreeze at adjacent points where the surface is under
tension. This process becomes effective beneath a considerable body of
snow, and here the granules gradually lose the spheroidal form assumed
in the early stages of granulation and become irregular polyhedrons
interlocked into a more or less solid mass.
A third factor is also to be recognized, though its effectiveness is
unknown. Under severe wind pressure, air penetrates porous bodies
with appreciable facility. The “breathing” of soils and the curious
phenomena of “blowing-wells” and “blowing-caves” teach us of the
effective penetration and extrusion of the air under variations of
barometric pressure. In the snow-fields, and in the more granular
portions of glaciers near their heads, the porosity is doubtless
sufficient to allow of the appreciable penetration of the atmosphere.
During a part of the time, the probable effect is the condensation
within the ice of moisture from the air, and during another part,
evaporation from the ice. These alternating processes are attended by
oscillations of temperature. While the balance between loss and gain
of substance may be immaterial, the oscillating nature of the process
and the fluctuations of temperature are probably favorable to granular
change.
Whether these processes furnish an adequate explanation of the changes
or not, the observed fact is that there are all gradations from
snowflakes and pellets into granular névé, and thence into glacier
granules (_Gletscherkörner_), varying in size up to that of filberts
and walnuts, and even beyond. In coherence, these aggregations may
vary from the early slightly coherent granular stage, where the grains
are small and spheroidal, to the ice stage, where the cohesion has
become strong through the interlocking growths of the large granules.
Even when the mass has become seemingly solid ice, sufficient space is
usually left between the granules to give the dispersive reflection to
light which imparts to glacier ice its distinctive whitish color.
=The arrangement of the crystal axes.=—The most radical difference
between glacier ice and ice formed directly from water is in the
arrangement (orientation) of the crystals. In the ice formed on
undisturbed water, the bases of the crystals are at the surface and
their principal axes are vertical, as shown by Mügge.[134] As they
grow, the crystal prisms extend downwards. This gives a columnar or
prismatic structure to the ice, well seen when it is “honeycombed”
by partial melting. In the glacier, on the other hand, the crystals,
starting from snowflakes, have their axes turned in various directions
according to the accidents of their fall; and as the snow develops
into ice, the principal axes of the crystals continue to lie in all
directions. Hence glacier ice, unlike pond ice, cannot usually be split
along definite planes, except where cleavage planes are subsequently
developed by extraneous agencies.
[Illustration: +Fig.+ 291_a_. +Fig.+ 291_b_. +Fig.+ 291_c_.
Figures to illustrate the method of deformation of ice crystals.]
While the crystals of a glacier usually have their principal axes in
various directions, there appears to be a tendency for them to approach
parallelism in certain positions, especially in the basal parts of a
glacier near its terminus. Observations on this point are not so full
and critical as could be desired, but it is probable that the parallel
orientation is partly general, and due to the vertical pressure of the
ice, and partly special and local, and connected with the shearing
planes and foliation.
The bearing of this partial parallelism of the crystals on shearing and
foliation is supposed to reside in the fact that a crystal of ice is
made up of a series of plates arranged at right angles to the principal
axis of the crystal. These plates may be likened to a pile of cards,
the principal axis being represented by a line vertical to them. If
a cube be cut from a large crystal of ice, it will behave much like
a cube cut from the pile of cards. If the cube be so placed that its
plates are horizontal (Fig. 291_a_), and if it be rested on supports at
two edges and heavily weighted in the middle, it will sag, the plates
sliding slightly over one another so as to give oblique ends, but in
this case the cube offers considerable resistance to deformation. If
the cube be so placed that the plates stand on edge, each reaching from
support to support (Fig. 291_b_), it will offer very great resistance
to deformation; but if the plates be vertical and transverse to the
line joining the supports, as in Fig. 291_c_, the middle portion will
sag under very moderate weighting by the sliding of the plates on one
another, and in a comparatively short time the middle portion may be
pushed entirely out, dividing the cube. These properties have been
demonstrated by McConnel[135] and Mügge, and they appear to throw light
on certain phases of the action of glaciers that are most pronounced in
their basal parts, and are best illustrated in arctic glaciers.
_The Probable Fundamental Element in Glacial Motion._
=Melting and freezing.=—It has already been shown (p. 279) that the
initial or fundamental cause of glacial motion must be operative at the
heads of glaciers where the temperature is lowest and the material
most loosely granular. In this condition, there is reason to believe
that motion takes place between the grains, rather than by their
distortion through the displacement of their laminæ. The fact that the
granular structure is not destroyed, as it would be by the indefinite
sliding of the crystal plates over each other, sustains this view. The
inference is that the gliding planes play a notable rôle in glacial
movement only in the basal parts of the lower ends of glaciers, where
the greatest thrusts are developed, and where the granules have become
largest and most completely interlocked. At the heads of glaciers,
where motion is initiated, there may be great downward pressure, but
not vigorous thrusts from behind, and probably only moderate thrusts
developed within the body itself. There seems therefore no escape from
the conclusion that the primal cause of glacial motion is one which
may operate even under the relatively low temperatures, the relatively
dry conditions, and the relatively granular textures which affect the
heads of glaciers. These considerations lead to the view that movement
takes place by the minute individual movements of the grains upon one
another. While they are in the spheroidal form, as in the névé, this
would not seem to be at all difficult. They may rotate and slide over
each other as the weight of the snow increases; but as they become
interlocked by growth, both rotation and sliding must apparently
encounter more resistance. The amount of rotary motion required of an
individual granule is, however, surprisingly small, and the meltings
and refreezings incident to shifting pressures and tensions, and to
the growth of the granules, seem adequate to meet the requirements.
In order to account for a movement of three feet per day in a glacier
six miles long, the mean motion of the average granule relative to its
neighbor would be, roundly, ¹⁄₁₀₀₀₀ of its own diameter per day, or one
diameter in 10,000 days; in other words, it would _change its relations
to its neighbors_ to the extent of its diameter in about thirty
years. A change of so great slowness under the conditions of granular
alteration can scarcely be thought incredible, or even improbable, in
spite of the interlocking which the granules may develop. The movement
is supposed to be permitted chiefly by the temporary passage of minute
portions of the granules into the fluid form at the points of greatest
compression, the transfer of the moisture to adjoining points, and its
resolidification. The points of greatest compression are obviously
just those whose yielding most promotes motion, and a successive
yielding of the points that come in succession to oppose motion most
(and thus to receive the greatest stresses) permits continuous motion.
It is merely necessary to assume that the gravity of the accumulated
mass is sufficient to produce the minute temporary liquefaction at the
points of greatest stress, the result being accomplished not so much
by the lowering of the melting-point as by the development of heat by
pressure.
[Illustration: +Fig.+ 292.—Portion of the east face of Bowdoin glacier,
North Greenland, showing oblique upward thrust, with shear.]
This conception of glacial “flowage” involves only the _momentary
liquefaction of minute portions of the mass_, while the ice as a whole
remains rigid, as its crystalline nature requires. Instead of assigning
a slow viscous fluidity like that of asphalt to the _whole_ mass, which
seems inconsistent with its crystalline character, it assigns a free
fluidity to a succession of particles that form only a minute fraction
of the whole at any instant.
This conception is consistent with the retention of the granular
condition of the ice, with the heterogeneous (in the main) orientation
of the crystals, with the rigidity and brittleness of the ice, and with
its strictly crystalline character, a character which a viscous liquid
does not possess however much its high viscosity may make it resemble a
rigid body.
=Accumulated motion in the terminal part of a glacier.=—However slight
the relative motion of one granule on its neighbor, the granules in any
part of a glacier partake in the accumulated motion of all parts nearer
the source, and hence all are thrust forward. Herein appears to lie the
distinctive nature of glacial movement. Each part of a stream of water
feels the hydrostatic pressure of neighboring parts (theoretically
equal in _all_ directions) and the momentum of motion, but not the
rigid thrust of the mass behind. Lava streams are good types of viscous
fluids flowing in masses comparable to those of glaciers and on similar
slopes, and, in their last stages, at similar rates, but their special
modes of flow and their effects on the sides and bottoms of their paths
are radically different from those of glaciers. Forceful abrasion, and
particularly the rigid holding of imbedded stones while they score and
groove the rock beneath, is unknown in lava streams and is scarcely
conceivable. There is, so far as we know, no experimental or natural
evidence that any typical viscous body in flowing over a rugose bottom
detaches and picks up fragments and holds them as graving tools in its
base so fixedly as to cut deep, long, straight grooves in the hard
bottom over which it flows. It would seem that competency to do this
peculiar class of work, which is distinctive of glaciers, should be
demonstrated before the viscous theory of glacial movement is accepted
as even a good working hypothesis. Somewhat in contrast with viscous
movement, it is conceived that a glacier is thrust forward rigidly by
internal elongation, shears forcibly over its sides and bottoms, and
leaves its distinctive marks upon them.
[Illustration: +Fig.+ 293.—Shearing plane well defined. A Spitsbergen
glacier. (Hamberg.)]
_Auxiliary Elements._
=Shearing.=—In the lower portion of a glacier where normally the
thrusts are greatest, the granules fewest, and their interlocking
most intimate, shearing takes place within the ice itself. This is
illustrated by the accompanying Figs., 292–295. The shearing results
in the foliation of the ice and in the forcing of débris between the
sheared layers. Thus the ice becomes loaded in a special englacial or
baso-englacial fashion, as previously mentioned and illustrated in Fig.
268.
Within the zone of shearing, it is probable that the gliding planes
of the crystals come into effective function. It is thought that the
combined effect of the vertical pressure, the forward thrust, and the
basal drag of the ice, may be to increase the number of granules whose
gliding planes are parallel to the glacier’s bottom. At any rate,
Drygalski reports[136] that there is a tendency to such an arrangement
in the basal portion of the Greenland glaciers at their borders. It
is conceived that where strong thrusts are brought to bear upon such
a mass of granules, those whose gliding planes are parallel to the
direction of thrust are strained with sufficient intensity to cause the
plates to slide over each other, while those which are not parallel
to the direction of thrust are either rotated into parallelism—when
they also yield—or are pressed aside out of the plane of shear. As
previously noted, shearing is observed to occur chiefly where the ice
below the plane of shearing is protected more or less from the force
of the thrust. It perhaps also occurs where the basal ice becomes so
overloaded with débris that it is incapable of ready movement.
[Illustration: +Fig.+ 294.—Portion of the lateral margin of a North
Greenland glacier. Shows upturning of the layers at the base, the
cleanness of the ice above the bottom, and, possibly, shearing.]
It is also probable that sharp differential strain and shearing are
developed at the level where the surface-water of the warm season,
descending into the ice, reaches the zone of freezing. The expanding of
the freezing water at the upper limit of the cold zone may cause the
layer expanded by it to shear over that below. As the level of freezing
is lowered with the advance of the warm season, the zone of shearing
also sinks. This may be regarded as an auxiliary agency of shearing, of
application to a special horizon.
=High temperature and water.=—In the zone of waste, a higher
temperature and more water lend their aid to the fundamental
agencies of movement, and there is need for these aids to promote a
proportionate movement, for here the granules are more intimately
interlocked and the ice more compact and inherently more solid and
rigid. The average temperature is, however, near the melting-point (p.
276), and during the warm season the ice is bathed in water so that
the necessary changes in the crystals are facilitated, and movement
apparently takes place even more readily than in the more open granular
portion of lower temperature and dryer state. The extraordinary
movements of certain tongues of ice in some of the great fiords of
Greenland are probably due to the convergence of very thick slow-moving
ice from the interior into basins leading down to the fiords. Into the
same basins a large amount of surface-water is concentrated at the same
time, with the result that the thick ice, bathed with water and having
a high gradient, develops unusual velocity during the warm season.
[Illustration: +Fig.+ 295.—Lateral margin of a North Greenland glacier,
Inglefield Gulf region. The overhanging edges of the successive
layers are not altogether the result of shear. They are due in part
at least to differential melting along the lines where débris comes
to the surface. The débris planes may be shear planes.]
=Applications.=—By a studious consideration of the coöperation of the
auxiliary agencies with the fundamental ones, the peculiarities of
glacial movement may apparently be explained. In regions of intense
cold, where a dry state and low temperature prevail, as in the heart
of Greenland, the snow-ice mass may accumulate to extraordinary
thicknesses, for the burden of movement seems to be thrown almost
wholly upon compression, with the slight aid of molecular changes due
to internal evaporation and allied inefficient processes. Since the
temperature in the upper part of the ice is very adverse (see p. 277),
the compression must be great before it becomes effective in melting
the ice, and hence the great thickness of the mass antecedent to much
motion. Similar conditions more or less affect the heads of alpine
glaciers, though here the high gradients favor motion with lesser
thicknesses of ice; but in the lower reaches of alpine glaciers, where
the temperatures are near the melting-point, and the ice is bathed in
water, movement may take place in ice which is thin and compact.
If the views here presented are correct, there is also, near the end
or edge of a glacier, the coöperation of rigid thrust from behind with
the tendency of the mass to move on its own account. The latter is
controlled by gravity, and conforms in its results to laws of liquid
flow. The former is a derived factor, and is a mechanical thrust. This
thrust is different from the pressure of the upper part of a liquid
stream on the lower part, because it is transmitted through a body
whose rigidity is effective, while the latter is transmitted on the
hydrostatic principle of equal pressure in all directions.
_Corroborative Phenomena._
The conception of the glacier and its movement here presented explains
some of the anomalies that otherwise seem paradoxical. While a glacier
in a sense flows over a surface, it often cuts long, deep furrows in
firm rock. It is difficult to explain this if the ice be so yielding
as to flow under its own weight on a surface which is almost flat. If
the mass is really viscous, its hold on its imbedded débris should
also be viscous, and a bowlder in the bottom should be rotated in the
yielding mass when its lower point catches on the rock beneath, instead
of being firmly held while a deep groove is cut. This is more to the
point since viscous fluids flow by a partially rotary movement. If,
on the other hand, the ice is always a rigid body which yields only
as its interlocking granules change their form by loss and gain, a
rigid hold on the imbedded rock at some times, and a yielding hold at
others, is intelligible, for on this view the nature of its hold is
dependent on the temperature and dryness of the ice. Stones in the base
of a glacier may be held with very great rigidity when the ice is dry,
scoring the bottom with much force, while they may be rotated with
relative ease when the ice is wet. In short, the relation of the ice to
the bowlders in its bottom varies radically according to its dryness
and temperature. _A dry glacier is a rigid glacier. A dry glacier is
necessarily cold, and a cold glacier is necessarily dry._
On the view here presented, a glacier should be more rigid in winter
than in summer, and the whole thickness of the glacier should
experience this rigidity chiefly at the ends and edges, where the
relative thinness of the ice permits the low temperature to reach its
bottom. The motion in these parts during the winter is, therefore, very
small.
In this view may also be found an explanation of the movement of
glaciers for considerable distances on upward slopes, even when the
_surface_ as well as the base is inclined backwards. So far does
this go that superglacial streams sometimes run for some distance
_backwards_, i.e. toward the heads of the glaciers, while in other
places surface-waters are collected into ponds and lakelets. Such a
slope of the surface of ice is not difficult to understand if the
movement be due to thrust from behind, or if it be occasioned by
internal crystalline changes acting upon a rigid body; but it must be
regarded as very remarkable if the movement be that of a fluid body,
no matter how viscous, for the length of the acclivity is sometimes
several times the thickness of the ice. Crevassing and other evidences
of brittleness and rigidity find a ready elucidation under the view
that the ice is a really solid body at all times, and that its apparent
fluency is due to the momentary fluidity of small portions of the mass
assumed in succession as compression demands.
In addition to the considerations already adduced, it may be urged that
a glacier does not flow as a stiff liquid because its granules are not
habitually drawn out into elongated forms, as are cavities in lavas and
plastic lumps in viscous bodies. Flowage lines comparable to those in
lavas are unknown in glaciers.
All this is strictly consistent with our primary thesis, that a glacier
is a crystalline rock of the purest and simplest type, and that it
never has other than the crystalline state. This strictly crystalline
character is incompatible with viscous liquidity.
_Other Views of Glacier Motion._[137]
While these views of glacial motion seem to us to best accord with the
known facts, they are not to be regarded as established in scientific
opinion, or as the views most commonly held. The mode of glacial
motion has long been a mooted question, and is still so regarded. The
main alternative interpretations that have been entertained are the
following:
(1) In the early days of glacial studies De Saussure thought that
glaciers slid bodily on their beds;
(2) Charpentier and Agassiz referred the movement to the expansion of
descending water freezing within the glacier;
(3) Rendu and Forbes, followed by many, perhaps most, modern writers,
believed ice to be viscous, and that in sufficiently large masses it
flows under the influence of its own weight, like pitch or asphalt;
(4) Others, realizing the fundamental difference between crystalline
ice and a true viscous body, have fallen back on a vague notion of
plasticity which scarcely amounts to a definite hypothesis at all;
(5) Tyndall urged that the movement was accomplished by minute repeated
fracturing and regelation, appealing to the fact that broken pieces of
ice slightly pressed together at melting temperatures freeze together,
but neglecting the fact that this would destroy the integrity of the
crystals;
(6) Moseley assigned the movement to a bodily expansion and
contraction of the glacier, analogous to the creeping of a mass of
lead on a roof;
(7) James Thompson demonstrated that pressure lowers the melting-point,
and while this effect is so small as probably to be ineffectual, it
is correlated with the very important fact that compression may cause
melting, which is not the case in most other rocks. He recognized that
under pressure partial liquefaction took place, that the water so
liberated might be refrozen as it escaped from pressure, and appears to
have regarded this as a vital factor;
(8) Croll held that the movement was due to a consecutive series of
molecular changes somewhat like the chain of chemical combinations in
electrolysis;
(9) Hugi, Eli de Beaumont, Bertin, Forel, and others thought that the
growth of the granules was the leading factor in the ice movement;
(10) McConnel and Mügge have made the gliding planes of the ice
crystals serve an important function in glacial movement.
It will be seen that the principle of partial liquefaction for which
Thompson laid the basis, the crystallization of descending water, urged
by Charpentier and Agassiz, and the granular growth on which Hugi,
Beaumont, Forel, and others founded their hypotheses, are incorporated
in the view already presented. Probably the agencies on which some of
the other views are based may also be participants in producing glacial
motion, sometimes as incidental factors, and sometimes perhaps as
important ones, for under different conditions, different agencies may
play rôles of varying importance. For example, in going over the brinks
of precipices of sufficient height, glaciers break into fragments
which are re-cemented below, and the “reconstructed” glacier moves on
as before. Here fracture and regelation are evident. The movement of
the gliding planes of the ice crystals over each other, which has been
looked upon as a special kind of viscoid movement, probably plays a
large part in the shearing movements in certain cases. But neither of
these is probably a large factor in ordinary glacial movement, and it
seems highly improbable that any of them are essential factors in the
primary movements in the snow-fields where glacial action begins.
CHAPTER VI.
THE WORK OF THE OCEAN.
The general facts concerning the depth of the ocean and the
distribution of its water have been given on a preceding page (p. 8),
and the origin of the ocean and the ocean basins is discussed in the
second volume. This chapter has to do primarily with the processes
now going on in the sea and its borders, in so far as they are of
importance in the interpretation of geologic history. The study of
these processes is prefaced by a few words concerning the amount and
composition of the sea-water, the life of the ocean, and the topography
of its bed.[138]
=Volume and composition.=—Every 1000 parts of sea-water contain about
34.40 parts by weight of mineral matter in solution. The principal
solids, acids, and bases, combined according to the principles laid
down by Dittmar, are shown in the following table:[2]
Chloride of sodium 77.758
Chloride of magnesium 10.878
Sulphate of magnesium 4.737
Sulphate of calcium 3.600
Sulphate of potassium 2.465
Bromide of magnesium 0.217
Carbonate of calcium 0.345
-------
Total salts 100.000
Expressed in terms of tons per cubic mile of sea-water, the composition
is as follows:[139]
Tons per Cubic Mile.
Chloride of sodium (NaCl) 117,434,000
Chloride of magnesium (MgCl₂) 16,428,000
Sulphate of magnesium (MgSO₄) 7,154,000
Sulphate of calcium (CaSO₄) 5,437,000
Sulphate of potassium (K₂SO₄) 3,723,000
Bromide of magnesium (MgBr₂) 328,000
Carbonate of calcium (CaCO₃) 521,000
-----------
For sea-water, total dissolved matter 151,025,000
Aside from the ingredients shown in the above tables, the presence of
the following has been proved: iodine, fluorine, phosphorus, silicon,
boron, silver, lead, copper, zinc, cobalt, nickel, iron, manganese,
aluminum, barium, strontium, arsenic, lithium, cæsium, rubidium, and
gold. Oxygen, nitrogen, and carbonic acid gas are also present in
quantity. The amount of carbonic acid is estimated to be 18 times as
great as in the atmosphere.[140]
The amount of sea-water is estimated by Murray at 323,722,150 cubic
miles,[141] or about 15 times the volume of the land above sea-level.
The volume and composition of the sea-water being known, the amount of
mineral matter which it contains may be readily calculated. Assuming
the average specific gravity of the mineral matter in solution to be
2.5, the 3.5% by weight becomes 1.4% by volume, and 1.4% of 323,722,150
cubic miles is 4,532,110 cubic miles. This then represents the
aggregate volume of mineral matter in the sea if it were precipitated
and compacted so as to have an average specific gravity of 2.5.
Assuming the average depth of the sea to be 2076 fathoms (12,456 feet),
as given by Murray, the mineral matter in solution, if precipitated,
would cover the ocean bottom to a depth of about 175 feet. Assuming the
area of the land to be to that of the sea as 28 to 72, this amount of
mineral matter would make a layer about 450 feet deep over the land.
Its amount is equal to about 20% of that of all lands above sea-level,
and it falls but little short of that in all lands below 600 feet in
altitude. If it were precipitated and concentrated in the shallow
waters about the borders of the lands, it would fill the sea out to
the depth of about 4000 feet, and would diminish its area by some
19,000,000 to 20,000,000 square miles, an area which is more than ⅓
of the present land surface. In other words, if the mineral matter in
the sea-water were precipitated and concentrated in the shallow waters
about the lands, it would restore the continental shelves to the land
areas, and add an almost equal area beyond.
These comparisons may perhaps help to give some idea of the amount of
mineral matter in solution in the sea, but they give no more than a
hint of the importance of the solvent power of water in the general
processes of rock decay, for most of the substances carried to the sea
in solution by rivers are extracted from the water about as rapidly
as they are supplied. Thus calcium carbonate is about twenty times as
abundant as sodium chloride in river-water,[142] but is only ¹⁄₁₂₅ as
abundant in sea-water.
The total river discharge into the sea is estimated at 6524 cubic miles
of water per year.[143] This water is estimated to carry to the sea
annually about half a cubic mile of mineral matter in solution. At this
rate it would take about 9,000,000 years for the streams to bring to
the sea an amount of mineral matter equal to that it now contains, but
the proportions of the ingredients would be very different.
The sodium chloride makes up about 2.4% of the mineral matter in
river-water and nearly 78% of the mineral matter of the sea. At this
rate it would take nearly 300,000,000 years for the salt of the sea
to have been contributed by the rivers. It is not to be understood,
however, that this figure indicates the age of the ocean. The salt is
not all brought in by the rivers; the rivers have probably not always
contributed at the present rate; and much salt once in the sea has been
precipitated. Nevertheless the above figure gives some suggestion as to
the order of magnitude of the figures which represent the age of the
ocean.
In contrast with the salt, the amount of calcium carbonate in the sea
is so small that at their present rate of contribution, it would be
brought to the sea by rivers in about 62,000 years.
=Topography of bed.=—The general relations of ocean basins to
continents are suggested by Fig. 296. The borders of the continental
platforms are covered by the _epicontinental sea_, while the _abysmal
sea_ occupies the ocean basins proper. From the figure it is seen that
an ocean basin is pronouncedly convex upward, and so departs as widely
as may be from the current notion of the homely utensil from which
it is named. Only when it is remembered that a level surface (on the
earth) is one which has the mean curvature of the earth, and that the
deeper parts of the ocean basin are well below the mean sphere level,
does the current name seem justified.[144] The figure also shows that
the depth of an ocean basin is slight compared with the radius of the
earth.
The bed of the ocean, like the face of the land, is affected by
elevations and depressions, and its deepest points are about as far
below its surface as the highest mountains are above it. There are
areas of the sea bottom which, as a whole, may be compared to the
plains of the land, and others which may be likened to plateaus, and
the lines of gradation between them are as indistinct as they often
are on the land. There are mountain peaks, chiefly of volcanic origin,
and depressions comparable to the great basins on the land. But apart
from these general features, there is little in common between the
topography of the sea bottom and that of the land. Mountain systems
are, for the most part, absent, though certain islands, like Cuba and
some of its associates, may be regarded as the crests of systems which
are chiefly submerged. If the water were drawn off from the ocean’s
bed so that it could be seen as the land is, its most impressive
feature would be its monotony. The familiar hills and valleys which,
in all their multitudinous forms, give the land surface its most
characteristic features are essentially absent. A large part of its
surface would be found to be so nearly flat that the eye would not
detect its departure from planeness.
[Illustration: +Fig.+ 296.—General relations of ocean basins to the
lithosphere. Lat. 20° S. Depth of the water (black) and height of
land exaggerated ten times. (Data from Murray, Scot. Geogr. Mag.,
Vol. XV, 1899.)]
The reason for this profound difference is readily found. On the land,
the dominant processes which shape the details of the surface are
degradational, and though the final result of degradation is flatness
(base-level), the immediate result is relief, and, most commonly,
relief of the hill-and-valley type. In the sea, the dominant processes
are aggradational, and tend to monotonous planeness.
=Distribution of marine life.=—Marine life has been of such importance
in the history of the earth that the elementary facts concerning its
distribution and the principles which control it are here recalled.
The distribution of marine life is influenced by many factors, chief
among which are _temperature_ and _depth of water_. Not only is life
more abundant in the warmer parts of the ocean than in the colder, but
the species inhabiting cold waters are different from those in warm,
and few species range through great variations. Many forms of life
are restricted to shallow water. Many more, especially those which do
not live on the bottom, swim about freely without reference to the
depth of the water beneath them, while relatively few are restricted
to great depths. Many species are also influenced by the _salinity of
the water_, which varies notably along coasts where the fresh waters
from the land are discharged; by the _character of the sediment_ at the
bottom, some species preferring mud, others sand, and others gravel;
by the _movement of the waters_, some species preferring still waters
and others rough; and some species by the _abundance and nature of the
food-supply_, and by _rival and hostile species_.
Subject to the exceptions determined by temperature, etc., plant life
abounds in shallow water out to depths of 100 fathoms or so, and is
found in abundance at the surface where the depth is much greater.
Animal life abounds in shallow water, both at the bottom and above it,
out to depths of 200 or 300 fathoms, and occurs in great profusion in
the surface-waters of temperate and tropical regions without regard to
the depth. The great body of the ocean water lying below a depth of
some few hundred fathoms is nearly tenantless, though life reappears
sparingly at the bottom, even where the depth is great. For further
discussion of this topic, see Chapter XI.
PROCESSES IN OPERATION IN THE SEA.
Within the area of the sea, as on the area of the land, three sets of
processes are at work—_diastrophism_, _vulcanism_, and _gradation_.
=Diastrophism= (p. 2) affects the sea-bottom as the land, but the
results are notably different in certain respects. So far as the
lithosphere is concerned, the sea-level may be said to be the critical
level. At and above it, many processes are in operation which do not
appear below, and below it, many which do not take place above. Changes
of level which do not involve the submergence of areas which were
land, or the emergence of areas which were under water, are relatively
unimportant, compared with those which effect such changes. The rise
of the bottom of the sea from a depth of 500 fathoms to a depth of 200
fathoms would not lead to important consequences, so far as the area
itself is concerned, while an equal rise of the bottom beneath 200
fathoms of water, or an equal subsidence of land 500 feet high, would
be attended by more striking consequences. It follows that the changes
effected by diastrophism are much more obvious along coasts than in the
deep seas. Emergence or submergence shifts the zones of aggradation
and degradation, shifts the zone of contact of ocean and land, and
changes the region concerned from one appropriate for sea life to one
appropriate for terrestrial forms, or _vice versa_.
Over the continental shelves the water is shallow and the bottom
relatively smooth. If a coastal region be elevated evenly, or if the
sea-level be drawn down, the new shore-line on the smooth surface of
the former submerged shelf will be relatively regular, even though the
coast was notably irregular before the change. Thus in Fig. 297 the
coast-line is notably irregular. A sea-withdrawal or a land-uplift of
120 feet would change the coast-line to the position of the 20-fathom
line, when it would be notably less irregular than now. If it were
shifted to the 100-fathom line, few irregularities would remain. In so
far as new coast-lines formed by the lowering of the sea (or rise of
the crust) depart from straightness, it is usually by broad, smooth
curves. Local uplifts of coastal lands, and especially uplifts along
axes normal to the trend of the coast, would give rise to projections
of land, and so to coastal irregularities; but such uplifts are rarely
so localized as to give origin to minor projections. It follows that
rising coasts, and those which have recently risen, or more likely,
coasts along which the sea-level is sinking or has recently sunk,
are likely to be regular so far as details of outline are concerned.
Subsidence of a coast-line (or rise of the sea-level) tends to the
opposite results, for in this case the sea advances on a surface which
has more or less relief, and the water takes possession of every
depression brought to its level. The lower parts of the valleys are
converted into bays, the length and width of which depend on the slope
and width of the valleys drowned. The numerous bays at the debouchures
of the streams along the Atlantic coast of the United States, from Long
Island Sound to Carolina, such as the Delaware, Chesapeake, (Fig. 297)
and numerous smaller bays, are the results of recent sinking, which
has allowed the sea to invade the lower ends of river valleys. The
ragged coast of Maine is another example, though glaciation as well as
subsidence has been operative here. From the present configuration of
coast-lines, it has been inferred that the present is, on the whole, an
era of continental depression.[145] River valleys, the lower ends of
which are embayed, are sometimes found to be continuous with submerged
valleys beyond the coast-line (Fig. 298). Submerged river valleys show
that the surface in which they lie was once land.
[Illustration: +Fig.+ 297.—Sketch of the eastern coast of the United
States from Cape May to Cape Henry, showing coastal irregularities.
The figures represent the depths of water in fathoms. (From charts of
C. and G. Surv.)]
[Illustration: +Fig.+ 298.—Sketch of Carmel Bay, Cal. The contours
below sea-level show a deep submerged channel. (From charts of C. and
G. Surv.)]
Bays may be developed by local subsidence as well as by the submerging
of valleys, though decisive examples are not readily cited. Bays may
also be produced by uplift of the surface on either side of an area
which does not change its level. For example, uplift on either side of
the Gulf of California has probably been one element, though probably
not the only one, in the development of this indentation. The general
outline of a great bay produced by coastal warping might be regular,
though it would be likely to be marked by small irregularities where
the streams enter. It is not to be understood that all, or even most,
bays are due to local diastrophism.
Diastrophism, then, as it affects the ocean-bottoms and the
ocean-borders, may make the water of any ocean shallower or deeper; it
may cause the emergence or submergence of land; it may make coast-lines
regular or irregular; it may shift the habitat of life, and through
these changes may greatly influence the processes of gradation, which
are especially active along the contact of sea and land.
=Vulcanism= affects the sea-bottom much as it affects the land. At the
volcanic centers, where the great body of extruded matter accumulates,
mounds and mountains are built up. Most of the mountain peaks of the
sea-bottom, whether their crests are islands, or whether they are
wholly submerged, have had a volcanic origin. The rock material ejected
from submarine vents is probably less widely distributed than that
from vents on land, and so far forth, the volcanic cones in the oceans
are steeper than those on land. Where volcanic cones are built up near
the surface of the sea, they often furnish a home for shallow-water
life, such as polyps. Wherever built up so as to be within the reach of
waves, gradational processes are stimulated.
The processes of vulcanism do not commonly influence coasts of
continents directly, for few volcanoes lie immediately on coasts.
In places, however, as at various points in and about Italy, the
configuration of the coast is influenced by the building of volcanoes.
Indirectly, vulcanism influences the shape of coast-lines, for the
resistance of igneous rock is often different from that of the rock
with which it is associated, and under the influences of the forces of
gradation it may come to form projecting points or reëntrants, as the
case may be.
The number of active volcanoes on islands is about 200, or about
two-thirds of all now known. Since the area of the sea is about three
times that of the land, the known active volcanoes in the sea are
rather less numerous per unit area than those on the land. The number
of active vents beneath the sea is altogether unknown. A few submarine
eruptions have been observed, and those observed are probably but a
small percentage of those which have taken place in historic time.
Slight eruptions in deep water might not manifest themselves at the
surface in an unequivocal way, even were observers stationed near them.
Volcanic cones which fail to reach the surface are known, and the forms
of many sea-bottom mountain peaks are such as to make it probable that
they are volcanic. These phenomena, as well as the numerous volcanic
islands, give some indication of the importance of submarine eruptions
in past time.
Ocean volcanoes, and especially submarine volcanoes, affect both the
temperature and the composition of the sea-water. Both the increase of
temperature and the solution of volcanic gases increase the capacity of
the water for mineral matter, and both the change in temperature and
composition affect the life of the adjacent waters. The destruction
of life during eruptions occasions the generation of the products of
organic decomposition, and these stimulate further chemical changes.
The diffusion of affected waters occasions chemical changes wherever
they go. The effects of oceanic volcanoes on the sea-water are,
therefore, appreciable, when long periods of time are considered. The
deposition of the finer parts of volcanic discharges will be considered
in connection with the deposits of the deep sea.
=Gradation.=—The gradational processes of the land and the sea are
in striking contrast. On the land, degradation predominates, and
aggradation is subordinate. In the sea, aggradation predominates, and
degradation is subordinate. On the land, degradation is, on the whole,
greatest where the land is highest, while aggradation is of consequence
only where the land is low, or where steep slopes give place to gentle
ones. In the sea, degradation is virtually confined to shallow water,
or to what might be called the highlands of the sea, while aggradation
is nearly universal, but most considerable in shallow water, or
where shallow water gives place to deep. Both the degradational and
aggradational work of the sea are greatest near its shores. Opposed as
the gradational work of the land and sea are, they yet tend to a common
end—the leveling of the surface of the lithosphere.
The gradational processes which affect the sea-bottom may be divided
into three categories: (1) Those effected by mechanical means, (2)
those effected by chemical means, and (3) those effected by organic
agencies.
The mechanical work of gradation in the sea is effected chiefly by the
movements of the water, and, very subordinately, by the movements of
the ice which the water carries. The results of these movements may be
degradational wherever the water is sufficiently shallow for the motion
to affect the bottom. Elsewhere it is aggradational.
The direct gradational work effected by chemical means is likewise
partly degradational and partly aggradational. If at any time or
place the water becomes supersaturated with any mineral substance,
precipitation takes place, and the precipitate accumulates as sediment
on the bottom. This sometimes happens in lagoons and other small
inclosures, and perhaps in open water. On the other hand, wherever
solution is effected, degradation is the result. Solution is most
important where the bottom consists of relatively soluble rock, such as
lime carbonate.
Organic agencies are, on the whole, aggradational. Accumulations of
coral, coral débris, shells, etc., help to build up the sea-bottom,
and most rapidly in shallow water where the proper forms of life are
most abundant. Here also should be mentioned the accumulations of
carbonaceous matter, especially in the form of plant bodies. In the
aggradation effected directly by organic agencies, the sea is passive.
Its only part is to support the life which gives rise to the solid
matter, and incidentally to float a part of it in its currents.
MOVEMENTS OF THE SEA-WATER.
The movements of the sea-water fall into several categories. There is
(1) a general circulation of sea-water, determined chiefly by three
factors: differences in density in the sea-water, differences of level,
and the general movements of the atmosphere; (2) periodic movements
which are not primarily circulatory, brought about by the attraction
of the sun and moon; and (3) aperiodic movements, due to occasional
causes, such as earthquakes, volcanic explosions, landslides, etc.,
which determine local and temporary movements, often of exceptional
strength.
=Differences in density and their results.=—Differences in density
result from differences in temperature and salinity. Temperature alone
considered, water would be densest where it is coldest, namely in
the polar regions. Differences in salinity result from differences
in evaporation and from inequalities in the supply of fresh water.
Evaporation alone considered, the sea-water should be densest where
evaporation is greatest; but the equatorial region, where evaporation
is greatest, is also a region where precipitation is heavy, and
precipitation, by freshening the water, opposes the effect of great
evaporation. The greatest differences in density due to the unequal
supply of fresh water are to be found near the borders of continents,
where the precipitation on the land is discharged into the sea. In
the polar regions, the great supply of fresh water, especially during
the season when the ice is melting, opposes the effect of the low
temperature, so far as the density of the water is concerned. The
result of the operation of these factors affecting the density of the
sea-water is to insure a general circulation, directed to the end
of equalizing the densities; and since the disturbing factors are
constantly in operation, equilibrium is never established, and the
movements of the water are perpetual.
The pressure gradients resulting from differences of density are so
slight that the resulting movements are scarcely more than a creep of
the waters. In general they are far too slow to be of importance in
gradational work; but the earth’s rotation deflects the creeping waters
and tends to concentrate the equator-ward movement into currents on
the east sides of the continents, and the pole-ward movement on the
west sides. In favorable situations these currents may be competent to
produce sensible mechanical results. Even where this is not the case
the circulation helps to equalize the temperatures of the sea, and so
of the air above and of the land about. Indirectly, therefore, the
circulation of the ocean-waters affects every geological process which
is sensitive to climate.
=Differences in level and their results.=—While the surface of the
ocean is the common datum plane to which elevations and depressions
are referred, it is to be remembered that the sea has “a very
complicated undulating surface in consequence of the attraction which
the heterogeneous and elevated portions of the lithosphere exercise on
the liquid hydrosphere. In the opinion of geodesists, the geoid may in
some places depart from the figure of the spheroid by 1000 feet.”[146]
These variations in level would, however, not occasion circulation. The
differences in level which determine circulation are much more trivial.
Every stream which pours fresh water into the sea tends to raise the
level of the water where it enters. The waters brought to the ocean by
the Amazon, the Mississippi, and other great rivers would appreciably
change the level of the sea at their debouchures, if the excess did
not promptly flow away. The ready mobility of the water, however,
prevents its accumulation, and the discharge of every stream generates
widespread movement. This movement is strongest at the debouchure, and
weakens with increasing distance from it, though in the case of great
streams, such as the Amazon, the movement is traceable, by means of the
sediment which the water carries, hundreds of miles out to sea.
Changes of level are also brought about by the winds, which pile up
water along the shore against which they blow. The level of the water
is said to have risen 24 feet at Calcutta on October 5, 1864, as the
result of a severe storm. While this is exceptional, a rise of 2
feet is not rare. This piling up of the waters along shore insures a
compensating movement (undertow, littoral currents, etc.) in some other
direction. Unequal evaporation and precipitation likewise disturb the
level of the sea and occasion movement. In the open sea the movements
generated by differences of level, like those generated by differences
of density, are chiefly slow, creeping movements, but movements which
never cease. In bays and gulfs, on the other hand, the surface of the
water may be so raised, either as the result of wind, river discharge,
or heavy precipitation, as to give rise to strong outward currents.
There is little doubt at the present time that the Gulf Stream owes its
origin primarily to the difference of level between the Gulf of Mexico
and the Atlantic.[147]
=Movements generated by winds.=—The circulation resulting from the
tendency of the winds to change the level of the sea-water has already
been mentioned, but the wind also works in other ways. Where the winds
have a somewhat constant direction and are at the same time strong,
they determine a general movement of the surface-waters in their own
direction, the surface-water being dragged along at a rate somewhat
less than that of the wind itself. The constant trades appear to be
the chief generators of the equatorial ocean-currents. Once generated,
these currents may be concentrated and their courses modified. The
currents generated by trades are turned north and south when directed
against a continent; they are modified by the configuration of the
bottom if the water be shallow, and always and everywhere, except, at
the equator, they are deflected by the rotation of the earth, in the
northern hemisphere to the right, and in the southern to the left. The
pole-ward currents generated in the equatorial region by the trades,
and directed by the winds, the lands, the configuration of the bottom,
and the rotation of the earth, determine compensating currents from
high latitudes to low, and the same influences which control the course
of the former direct the latter as well.
Since the atmospheric movements are so far constant that there is a
prevailing direction of winds in all latitudes, the winds, as well
as differences of density and differences of level, insure a general
and continual circulation of sea-water. The geological effects of
this circulation are direct and indirect; direct, by gradation of the
bottom over which they flow, and indirect, by the modifications of
climate they produce. Since rotation deflects the pole-ward currents
to the east sides of the oceans (west sides of the continents) and the
equator-ward movements to the west sides of the oceans (east sides of
the continents), the east shores of the oceans are warmer than the west
in corresponding latitudes, and the west sides of the continents are
both warmer and moister[148] than the east sides.
The most obvious disturbance of sea-water resulting from the winds is
the generation of _waves_. Waves are not primarily parts of the general
oceanic circulation. Since they are generated in other ways than by
winds, and since the gradational effects of waves are independent
of their origin, the effects of wind-waves will not be considered
separately.
=Movements generated by attraction.=—One of the movements of the
sea-water which is not primarily circulatory results from the
attraction of the moon and sun. The tide is really the result of the
_inequalities_ of the attraction of these bodies on different parts
of the earth. The lunar tide is more important than the solar, not
because the attraction of the moon is greater, for it is not, but
because its differential attraction, the result of its lesser distance,
is greater.
The distance of the moon from the earth is about 240,000 miles. If
this be taken as the distance from the center of the moon to the
center of the earth, 236,000 and 244,000 miles respectively are the
distances from the center of the moon to the nearest and most distant
points on the earth. The distance of the sun from the earth is about
93,000,000 miles. If this be taken as the distance between the centers
of these bodies, then the distances from the center of the sun to the
nearest and most distant points on the earth’s surface are 92,996,000
and 93,004,000 miles respectively. The ratio of 4000 to 236,000 or to
244,000 is much greater than the ratio of 4000 to 92,996,000 or to
93,004,000. Hence the tide-producing force of the moon is greater than
that of the sun.
The tides show themselves along shores in the form of waves which, in
shallow water, become translatory. They differ from the wind-waves in
their periodicity, and locally in their greater height. The effects of
the tidal waves on the shores of the sea, and on the bottom in shallow
water, are the same as the effect of wind-waves of equal strength, and
need not be separately considered in connection with the gradation of
the sea-bottom. In passing through narrow straits or narrow passes of
any sort, the tidal movement becomes a current which, under favorable
conditions, abrades or “scours” the bottom effectively. The tidal
currents in the narrow passes about New York harbor may serve as an
illustration.
=Aperiodic movements.=—In addition to the wind-waves which are
essentially constant and universal, and to the tidal waves, which are
periodic, there are accidental waves which are locally and temporarily
of importance. Such are earthquake-waves, which are sometimes extremely
destructive. Thus an earthquake-wave on the coast of Peru in 1746 swept
a frigate several miles inland and deluged Lima, seven miles from the
shore. The havoc of most earthquakes affecting coasts, such as that
of Lisbon in 1755, is greatly aggravated by accompanying sea-waves.
Earthquake-waves differ from ordinary waves in being translatory, and
so in being more effective on the bottom in deep water. Their greatest
force, however, is felt in shallow water and on shores. Volcanic
eruptions likewise give rise to exceptional aperiodic waves. The same
is true of landslides where they affect the coast or any part of the
sea-bottom. The fall of glacier ends and the capsizing of icebergs
likewise generate strong waves. To the category of exceptional waves
also belong those generated by the winds of exceptional storms, such as
that which devastated Galveston in 1900.[149]
=Summary.=—From the point of view of their direct geological results
in shallow water, all movements of the sea-water may be grouped into
two main classes—(1) waves, with the undertow and the littoral currents
they generate, and (2) ocean-currents.[150]
WAVES.
=Wave-motion.=[151]—The most common waves, and from the present point
of view the most important, are those generated by winds. During the
passage of a wave, each particle affected by it rises and falls, and
moves forward and backward describing an orbit in a vertical plane. If
the passing wave is a swell, the orbit of the particle is closed and is
either a circle or an ellipse; but in the case of a wind-wave the orbit
is not closed. In such a wave two things move forward, the undulation
and the water. The velocity of the undulation is relatively rapid;
that of the water, slow and rhythmic. On the crest of the wind-wave
each particle of water moves forward, and in the trough it moves less
rapidly backward, and the excess of the forward movement over the
backward gives it a slight residual advance. This residual advance is
the initiatory element of current. By virtue of it, the upper layer
of water is carried forward with reference to the layer below, in the
direction toward which the wind blows. The waves of any considerable or
long-continued wind, therefore, generate a current tending in the same
direction as the wind.
The agitation of which waves are the superficial manifestation
is not restricted to the surface, but is propagated indefinitely
downward. Near the surface the amount of motion diminishes rapidly
with increasing depth (Fig. 299), but the rate of diminution itself
diminishes, and there seems no theoretic reason for assigning any
definite limit to the downward propagation of the oscillation.
[Illustration: +Fig.+ 299.—Figure illustrating the decrease in the
amount of wave-movement with increase of depth. (Fenneman.)]
At the surface, the radius of the circular orbit which a particle of
water in a wave tends to describe is half the height of the wave. At a
depth equal to one wave-length, the radius of the circle described by
a particle is ¹⁄₅₃₅ as great as at the surface, and at a depth equal
to two wave-lengths, ¹⁄₃₀₀₀₀₀. If the height of a wave be 43 feet, the
radius of the circle described by a surface particle is 21½ feet. If
the length of the wave be 300 feet, the radius of a particle at a depth
of 300 feet is only about ⁴⁄₁₀ of an inch, and at 600 feet ¹⁄₁₂₀₀ of an
inch.[152] These figures make it clear that effective agitation of the
water does not extend to great depths.
So long as the velocity of the wind remains constant, the velocity
of the current which the wind-waves generate is less than that of
the wind, and there is always a differential movement of the water,
each layer moving faster than the one beneath. The friction is thus
distributed through the whole vertical column of the water in movement,
and is even borne in part by the sea-bottom if the movement extends
so far down. The greater the depth, the smaller the share of the
friction each layer of water is called upon to bear, and the greater
the velocity of the current generated by a given wind. But while the
wave-motion extends indefinitely downward, the lower limit of agitation
effective in erosion is soon reached. Engineering operations have shown
that submarine structures are little disturbed at depths of five
meters in the Mediterranean and eight meters in the Atlantic.[153] On
the other hand, débris as coarse as gravel, which is transported by
rolling on the bottom, is not infrequently carried out to depths of
50 feet, and sometimes even to 150 feet. Fine sediment, like silt, is
disturbed at still greater depths, for ripple-marks, which indicate
agitation of the water, are said to have been found at depths of 100
fathoms.[154]
When a wave approaches a shelving shore, its habit is changed. The
velocity of the undulation is diminished, while the velocity of the
advancing particle of water in the crest is increased; the wave-length,
measured from trough to trough, is diminished, and the wave-height is
increased; the crest becomes acute, with the front steeper than the
back, and these changes culminate in the breaking of the crest, when
the undulation proper ceases. Waves of a given height break in about
the same depth of water, and the line along which incoming waves break
is the line of _breakers_. The line of breakers is in deeper water and
farther from shore when the waves are strong than when they are weak.
Waves are reported to have broken in 100 fathoms of water,[155] but
this must be regarded as very exceptional. The return of the water
thrown forward in the crests of waves is accomplished by a current
along the bottom called the _undertow_. The undertow is sensibly normal
to the coast when uninfluenced by oblique waves, and is efficient in
removing the products of erosion.
Since the incoming wave affects water which is at the same time under
the influence of the undertow, it gives to that current a pulsating
character, for the wave-motion sometimes supports and sometimes
opposes the undertow, and thus endows it with a higher transporting
power than belongs to its mean velocity. Near the breaker-line, the
oscillations communicated by the wave may momentarily overcome and
even reverse the movement of the undertow. Inside the breaker-line,
irregular oscillation only is communicated. The broken wave-crest,
dashing forward, overcomes the undertow and throws it back, and the
water returns as a simple current descending a slope. The power of the
undertow diminishes rapidly from the breaker-line outward as the depth
of the water increases.
[Illustration: +Fig.+ 300.—Diagram showing relative directions of wave,
undertow, and shore-current.]
When waves advance on the shore obliquely, a shore-current is developed
as illustrated by Fig. 300, where _ab_ represents the direction of the
incoming wave, _bc_ the direction of the littoral current, and _bd_ the
direction of the undertow. Where they strike the borders of land, the
wind-waves, therefore, generate two other movements, the undertow and
the littoral current. Any particle of water near shore may be affected
by any two or by all three of these movements at the same moment. The
effect of littoral current and undertow is to give a particle of water
on which both are working a direction between the two, as _be_. The
effect of other combinations can be readily inferred. These various
combinations are of consequence in the transportation of débris.
WORK OF THE WAVES.
_Erosion._
The general effects of the waves and the other movements to which
they give rise along shores are (1) the wear of the shores; (2) the
transportation for greater or less distances of the products of wear;
and (3) the deposition of the transported materials.
=By waves and undertow.=—In the dash of the waves against the shore,
the chief wear is effected by the impact of the water and of the débris
which the water carries. Lesser results are accomplished in other ways.
When the land at the margin of the water consists of unconsolidated
material, or of fragmental material but slightly cemented, the impact
of the water is sufficient to displace or erode it. If weak rock be
associated with resistant rock within the zone of wave-work, the
removal of the former may lead to the disruption and fall of the
latter, especially when weak rock is washed out from beneath the
strong. The impact of the water is competent also to break up and
remove rock which was once resistant, but which has been superficially
weakened by changes of temperature. Rock affected by numerous open
joints is likewise attacked with success, for by the dash of the waves
the blocks between the joints may be loosened and literally quarried
out. It may, however, be doubted whether the dash of waves of clear
water, even when their force is many tons to the square foot, has any
appreciable power to wear rock which is thoroughly solid.
[Illustration: +Fig.+ 301.—Angular blocks of rock which have fallen
from the cliff above, as a result of undercutting by the waves.
Grand Island, Lake Champlain. The rock is Black River limestone.
Although from the shore of a lake instead of the sea, the principles
illustrated are the same. (Perry.)]
The impact of the waves is generally reinforced and made effective
by the impact of the detritus they carry. The sand, the pebbles, and
such stones as the waves can move are used as weapons of attack, being
turned against one another and against the shore. Masses of rock too
large for the waves to move (Fig. 301) are worn by the detritus
driven back and forth over them, and in time reduced to movable
dimensions (Fig. 302). They then become the tools of the waves, and in
use, are reduced to smaller and smaller size. Thus bowlders are reduced
to cobbles, cobbles to pebbles, pebbles to sand, and sand to silt.
The silt is readily held in suspension in agitated water, and thus is
carried out beyond the range of breakers, and settles in water so deep
as not to be effectively agitated to its bottom. Thus one generation
of bowlders after another is worn out, and the comminuted products are
carried out from the immediate shore and deposited in deeper water.
The effectiveness of waves, whether they work by impact of water alone,
or by impact of water and detritus, is dependent on their strength
and on the concentration of their blows.[156] The strength of waves
is dependent on the strength of the winds (or other generating cause)
and the depth and expanse of the water, and the concentration of
their blows is conditioned by the slope against which they break. On
exposed ocean-coasts the fetch of the waves is always great. The winds
are variable. For a given coast they have an average strength, but
the effectiveness of wave-erosion is determined less by the average
strength of waves than by the strength of the storm-waves. This is
often very great. On the Atlantic and North Sea coasts of Britain,
winter breakers which exert a pressure of three tons per square
foot are not infrequent.[157] So great is the force of exceptional
storm-waves that blocks of rock exceeding 100 tons in weight are
known to have been moved by them. Ground-swells, “even when no wind
is blowing, often cover the cliffs of north Scotland with sheets of
water and foam up to heights of 100 or even nearly 200 feet. During
northeasterly gales the windows of the Dunnet Head lighthouse, at
a height of upwards of 300 feet above high-water mark, are said
to be sometimes broken by stones swept up the cliffs by sheets of
sea-water.”[158] The average force of waves on the Atlantic coast of
Britain has been found to be 611 lbs. per square foot in summer, and
2086 lbs. in winter.[159]
Where deep water extends up to the shore, the force of the wave is
almost wholly expended near the water line; where shallow water borders
the land, the force of the waves is expended over a greater area.
Waves are, therefore, most efficient on bold coasts bordered by broad
expanses of deep water.
The less familiar phases of wave-work are accomplished by hydraulic
pressure, compressed air, the use of ice, etc. When the water of a
wave is driven into an open joint or a cave, the hydraulic pressure
is great, and if the structure be weak, the rock may be broken. When
water is driven with force into a cave, the compression of the air
may be great if the wave be high enough to close the entrance. When
the water runs out of a cave, the air within may be greatly rarefied,
while that above exerts its normal pressure. In either case the roof
of the cave, if it be weak, may be broken. At certain seasons of the
year, especially during the spring, waves make destructive use of the
ice which is then breaking up, but it is only in high latitudes that
sea-ice is of consequence in this way. In general, the effect of its
presence in keeping down waves overbalances its effect as an agent of
erosion.
[Illustration: +Fig.+ 302.—Showing blocks similar to those of Fig. 306,
reduced and rounded by wave-action. Shore of Lake Champlain. The rock
is Utica shale. (Perry.)]
The direct effect of wave-erosion is restricted to a zone which is
narrow both horizontally and vertically. There is no impact of breakers
at levels lower than the troughs of the waves, though erosion may
extend down to the limit of effective agitation (p. 341). The efficient
impact of waves is limited upward by the level of the wave-crests,
although the dash of the water produces feebler blows at higher levels.
The rise and fall of the water during the flow and ebb of the tides
gives the waves a greater vertical range than wind-waves alone would
have. The vertical zone of direct wave-work is therefore limited above
by the level of wave-crests, and below by the depth of wave-troughs
(nearly). The indirect work of waves is limited only by the height
of the shore, for as the zone of excavation is carried landward,
masses higher up the slope are undermined and fall. The fallen rock
temporarily protects the shore against the waves, but are themselves
eventually broken up.
[Illustration: +Fig.+ 303.—Diagram illustrating high sea-cliffs. It
also shows a submerged terrace, due partly to wave-cutting (wave-cut
terrace), and partly to building (wave-built terrace). (Gilbert.)]
[Illustration: +Fig.+ 304.—Diagram showing a low sea-cliff. (Gilbert.)]
The pulsating current of the undertow (p. 341) has both an erosive and
a transporting function. It carries the detritus of the shore to and
fro, and dragging it over the bottom, continues downward the erosion
initiated by the breakers. This downward erosion is the necessary
concomitant of the shoreward progress of wave-erosion; for, if the
land were merely planed away to the level of the wave-troughs, the
incoming waves would break where shoal water was first reached, and
become ineffective at the water margin. The rate of erosion by the
undertow becomes less and less as the surface it affects is lowered.
Littoral currents do little erosive work beyond that inflicted on the
material which they transport.
The general result of wave-erosion is the advance of the sea on the
land, the rate of advance being determined chiefly by the nature of the
material attacked and the strength of the waves. Numerous as examples
are of the retreat of coast-lines before the advance of the sea, it
is not to be understood that the advance of the sea on the land is
universal or uninterrupted. Numerous instances may be cited of the
encroachment of the land on the sea. At Long Branch the advance of the
sea, in spite of elaborate breakwaters, has been so rapid in recent
years as to menace important buildings, while a few miles to the north
and south, the land is advancing in the face of the waves. The low
coast of the Middle Netherlands has retreated two miles or more in
historic times,[160] but the opposite tendency is shown at other points
in the same region. On the coast of England the sites of villages have
disappeared by the advance of the sea within historic times,[161] but
the coast of the same island affords illustrations of land advance. On
the south side of Nantucket island, the sea-cliff has been known to
retreat before the waves as much as six feet in a single year.[162]
Almost every considerable stretch of coast affords illustrations both
of the advance of the sea on the land and of land on the sea; but in
the long run, the former must exceed the latter, diastrophic movements
aside.
[Illustration: +Fig.+ 305.—Steep cliff developed by waves. Allen Point,
Grand Island, Lake Champlain. (Perry.)]
[Illustration: +Fig.+ 306.—Cliff in unconsolidated material (bowlder
clay), with lake-beach in foreground. South Manitou Island, Lake
Michigan. (Russell, U. S. Geol. Surv.)]
[Illustration: +Fig.+ 307.—Steep cliff in unconsolidated material, the
result of rapid cutting. Southeast extremity of Grove Point, Md.]
_Topographic Features Developed by Wave-erosion._
[Illustration: +Fig.+ 308.—Standing Rock. A wave-erosion monument. West
shore of Random Sound, south of Clarenville, N. F. (Walcott, U. S.
Geol. Surv.)]
=The sea-cliff.=—The action of the waves, cutting as they do along
a definite horizontal zone, has been compared to the action of a
horizontal saw. As the waves cut into the shore at and near the
water-level, the material above, being unsupported, falls, leaving a
steep face above the line of cutting. This steep face is known as _the
sea-cliff_ (Figs. 301 to 306). The same term is sometimes applied to
the cliffs of lakes. The principles involved in the development of the
sea-cliff are applicable to any broad stretch of water.
The height of the cliff depends on the height of the land on which
the sea is advancing. Its slope may be steep or gentle (compare Figs.
303 to 306), according to the nature of the material of which it is
composed and the rapidity of the cutting. Rapid cutting tends to
produce steep cliffs and slow cutting gentle ones, for in the latter
case weathering is more important relative to the cutting, and at
sea-level (low altitudes) weathering generally tends to reduce the
angle of slope. In general, the more resistant the material the steeper
the slope of the cliff. Incoherent materials, such as sand and clay,
are not likely to form steep cliffs; but if the cutting be very rapid,
bold faces may be developed even in such materials (Fig. 307). If beds
of slight resistance at sea-level underlie beds of greater resistance,
the development of steep cliffs is favored. The structure of the
cliff-rock also has an influence on the slope. The rock may be massive
or bedded. If bedded, the beds may be horizontal, or they may dip at
any angle, in any direction. The rock, whether stratified or not, may
be abundantly or sparsely jointed. All these structures influence the
slope and configuration of the sea-cliff (see Figs. 305 to 308).
[Illustration: +Fig.+ 309.—“Old Man of Hoy.” (Geikie.)]
=Chimney-rocks, etc.=—By working in along the joints of the rock,
widening them and quarrying out the intervening blocks, pillars of
rock (“chimney-rocks,” “pulpit-rocks”) or even considerable islets are
sometimes isolated by the waves. This is most readily accomplished
where the joints converge back from the shore. A well-known example
of this sort is the “Old Man of Hoy” (Fig. 309) on the coast of the
Orkneys. A pulpit-rock or other island, or any jutting point of
rock may be pierced, giving an arch or bridge. La Roche Percée, a
steep-faced isle near Gaspé Harbor, is an example.
=Sea-caves.=—Waves sometimes excavate caves at the bases of cliffs.
This is especially likely to occur where the rock is much jointed and
where the joints are not continued to the surface in a single plane.
The bottom and roof of a sea-cave usually have a pronounced inclination
landward. If the cliff be low, the cave may be extended landward until
its roof is pierced. Through such an opening in the top of the cliff
the water of the incoming waves may be forced in the form of spray.
On the New England coast such holes are sometimes known as “spouting
horns.” Similar openings may be made, as already pointed out, by the
compression or rarefaction of the air in the cave as the wave enters or
retreats. If the roof of the cave be partially destroyed, the portion
which remains may form an arch or _bridge_. Such a bridge occurs on
Santa Cruz Island, California (Fig. 310).
[Illustration: +Fig.+ 310.—An arch developed by waves. Santa Cruz
Island, Cal. (Law.)]
The cave, the “spouting horn,” the “bridge,” the “pulpit-rock,” and
other isolated islets, are all closely associated with the sea-cliff in
origin.
=The wave-cut terrace.=—The bottom of the sea-cliff is bordered by a
submerged platform over which the water is shallow. This platform,
or at any rate its landward portion, represents the area over which
the water has advanced as the result of wave-cutting, and is,
therefore, known as the _wave-cut terrace_. From the method of cliff
development it will be seen that the wave-cut terrace is its necessary
accompaniment. Such a terrace has a gentle slope to seaward, for its
outer and older edge has been degraded longer and more. Its slope is
influenced by the strength of the waves, being greater where they are
stronger. The outer edge of the wave-cut terrace is often marked by
an abrupt descent. Fig. 303 represents the wave-cut terrace in its
relation to the sea-cliff above.
[Illustration: +Fig.+ 311.—An elevated cliff above Great Salt Lake.
In this case the water-level has been lowered. (Gilbert, U. S. Geol.
Surv.)]
So long as wave-cut terraces are submerged, they do not appear on
topographic maps of the land, though they appear on the charts of the
coasts; but if a coastal tract with wave-cut terraces be elevated, or
if the sea-level be drawn down, the terraces become land. Elevated
sea-cliffs and wave-cut terraces are among the best evidences of change
of relative level between water and land (Fig. 311).
=Wave-erosion and horizontal configuration.=—The structure of the rock
along shore has as much to do with the horizontal configuration of
the wave-shaped coast, as with its relief. In general, waves develop
reëntrants in the less resistant portions of the shore, leaving the
more resistant parts as headlands (San Pedro Point and Devil’s Slide,
Pl. XX, Coast of California). It is to be noted that the resistance
of rock to wave-erosion is not determined by its hardness alone.
Every division plane, whether due to bedding, to jointing, or to
irregular fracture, is a source of weakness to the rock, and rock
of great hardness may be so broken as to offer relatively little
resistance. Inequalities of resistance, whatever their cause, give
origin to inequalities of coastal configuration where wave-erosion is
in progress. Given a coast of marked regularity and equal exposure,
but composed of unequally resistant material, the waves will make it
irregular by cutting most where the material is least resistant. A
regular coast of uniform material, but unequal exposure, will be made
irregular by the greater cutting at the points of greater exposure. A
coast of marked irregularity and homogeneous material will be made more
regular by the cutting off of the projecting points, because they are
most exposed. With a given set of conditions, waves tend to develop
a certain sort of shore-line which, so far as its horizontal form is
concerned, is relatively stable. Such a shore-line may be said to be
_mature_[163] so far as wave-erosion is concerned. Since coastal lands
are, in general, both heterogeneous and unequally exposed, a mature
coast-line is somewhat irregular. Its maturity is attained when the
lesser exposure in the reëntrants developed in the less resistant
parts, balances the superior exposure of the projections of the more
resistant portions.
Since the conditions of erosion along coasts are constantly, even if
slowly, changing, maturity is constantly being approached, but rarely
reached. Other forces and processes, such as those of aggradation,
vulcanism, and diastrophism, are in operation along coasts, and
their results are sometimes antagonistic to those of the waves. The
horizontal configuration of coasts is, therefore, the result of many
coöperating forces, of which waves are but one. It is, nevertheless,
important to note the goal to which the waves are working, even though
they are continually defeated in their attempt to reach it. Their
immediate goal is an equilibrium of erosion-rate and maturity of
configuration; their final goal is the destruction of the land and the
deposition of its substance in the sea, that is, in a position nearer
the center of gravity of the earth.
_Transportation by Waves._
The material eroded from the shore by the waves in the shaping of the
cliff and terrace is carried away by the joint action of the waves,
undertow, and shore-currents.
The in-coming wave begins to shift material where it begins to drag
bottom, that is, a little outside the line of breakers. From the line
where transportation begins, to the line of breakers, bottom detritus
is shifted shoreward by the waves, while the undertow tends to carry it
back again. Between the breakers and the shore there is also a tendency
for the on-shore movement to carry débris to the water’s edge, and for
the ebbing wave to carry it back again. The result of these opposed
tendencies is to keep sediment in transit between the shore and the
line of breakers. If the in-coming waves have a direction normal to the
shore, the advance and recoil of the water move particles toward and
from the shore, but effect no transfer along the shore; but the results
which waves normal to the shore would achieve are always modified by
other waves and by littoral currents.
If the in-coming wave is oblique to the shore, it shifts material in
its own direction. The transfer by undertow, taken alone, would be
sensibly normal to the shore, but the effect of the oblique waves is
to slightly modify this direction. There is thus a slow transportation
along shore, even in the absence of steady currents. A great amount
of transportation would be effected in this way, though it would
be carried on at a slow rate. Oblique waves also tend to develop a
definite shore-current (p. 342) which affects both the amount and
direction of the transportation. Any particle in suspension, or in
motion on the bottom as the result of the wave or undertow, is shifted
along shore by the littoral current, which affects the same water
(Fig. 300). By the coöperation of wave- and shore-current, more and
heavier material can be moved than by either alone, and the direction
of movement is more nearly parallel to the shore than that of the
wave. Similarly, by the coöperation of undertow and shore-current,
more and heavier material can be moved than by either alone. The
direction of movement is readily inferred from Fig. 300. The direction
in which débris is shifted by wave- and shore-current is modified by
the undertow, and the direction which would result from undertow and
current is modified by the wave. It is often the waves of storms,
rather than those of the prevailing winds, which determine the
direction of greatest shore transportation.
The waves, the undertow, and the littoral currents work together in
assorting the detritus of the shore. The coarsest parts may be beyond
the power of all but the strongest waves. They accumulate where
agitation is great. Less coarse parts are shifted farther from the site
of greatest agitation, but no materials which are classed as coarse
are carried beyond the depth of sensible movement. The coarse material
which covers the bottom where the agitation of the water at the bottom
is effective, constitutes _shore drift_.
Shore drift is not all derived from the shore by the cutting of the
waves. A part of it is brought to the sea by streams and mingled with
that eroded from the cliffs. The material which is fine enough to be
held in suspension is measurably independent of depth. This is shown
during storms when the water becomes turbid far beyond the line of
breakers, and clears only after the waves have died away.
This sorting of shore drift, effected while it is in transportation, is
often very perfect. The conditions favoring assortment are (1) vigorous
wave-action, (2) prolonged transportation, and (3) a moderate volume of
sediment.[164] The effect of these several conditions will be readily
understood.
Extensive transportation of shore drift of a given degree of coarseness
is favored by (1) strong waves and undertow, (2) continuous currents,
and (3) shallow water, deepening but gradually off shore.
_Deposition by Waves, Undertow, and Shore-currents._[165]
[Illustration: +Fig.+ 312.—Cross-section of the beach. (Gilbert.)]
=The beach.=—The zone occupied by the shore drift in transit is the
_beach_. The lower margin is beneath the water, a little beyond the
line where the great storm-waves break. Its upper margin is at the
level reached by storm-waves, and is usually a few feet above the level
of still water. To the beach, material is brought from seaward by the
in-coming waves, and from it detritus is carried out by the undertow.
The cross-section of a beach is shown in Fig. 312. In horizontal
position the beach follows the general boundary between water and land,
though it does not conform to its minor irregularities (Fig. 313). The
beach or barrier ridge often causes the deflection of the lower courses
of streams descending to it (Pl. XXI).
[Illustration: +Fig.+ 313.—A lake-beach (barrier); Griffin’s Bay, Lake
Ontario.]
[Illustration: PLATE XXI.
NEW JERSEY
U. S. Geol. Surv.]
[Illustration: PLATE XXII.
Fig. 1. PORTION OF SOUTH COAST OF MARTHAS VINEYARD, MASSACHUSETTS.
U. S. Geol. Surv.
Fig. 2. PORTION OF THE CALIFORNIA COAST NEAR TAMALPAIS.
U S. Geol. Surv.]
[Illustration: +Fig.+ 314.—Section of a barrier. (Gilbert.)]
=The barrier.=—When the agitation of the water along shore becomes
insufficient to carry the material, it is dropped. In its deposition it
assumes various forms. Where the bottom of the lake or sea near shore
has a very gentle inclination, the in-coming waves break some distance
from the shore-line, and it is here that the most violent agitation
occurs when the waves are strong. To this line of breakers, material
is shifted from both directions: from shore by undertow, and from
seaward by the waves. Accumulating here, it builds up a low ridge.
This is a _barrier_ (Fig. 314). If it is built up above the surface of
the water by storm-waves, it may shut in a lagoon behind it, and this
may ultimately be filled by sediment washed down from the land. At one
stage in the filling, the lagoon becomes a marsh.[166] In the part
which the barrier plays in the history of a coast, it is identical with
the beach.
[Illustration: +Fig.+ 315.—A recurved spit. Dutch Point, Grand Traverse
Bay, Lake Michigan.]
[Illustration: +Fig.+ 316.—Cross-section of a bar. (Gilbert.)]
=The spit, the bar, and the loop.=—The disposition of shore-deposits
depends largely on the currents at and near shore. If the coast-line
is deeply indented, the littoral current usually fails to follow the
reëntrants. In holding its course across the mouth of a small bay,
a shore-current usually passes into deeper water. Here its velocity
is checked because its motion is communicated to the water beneath
it, and a larger amount of water being involved in the motion, the
motion of each part is diminished. If sediment was being moved along
its bottom before the current was checked, some part of it is dropped
when and where the current is slackened. It follows that deposition
commonly takes place beneath a littoral current as it crosses the
mouth of a bay. The belt of deposition is often narrow, and the result
is the construction of a ridge beneath the water in the direction of
the current. The current would never build the embankment up to the
water-level, but when its surface approaches the level of effective
agitation, the waves may begin to work on it, as on a barrier, and
may build it up to, and even above, the surface of the water. So long
as the end of such an embankment is free, it is a _spit_ (Fig. 315
and Pl. XXI). If the spit be lengthened until it crosses, or nearly
crosses, the bay, shutting it off from the open water, it becomes a
bar. Bars have shut in lakes (ponds) on the coast of Martha’s Vineyard,
Mass. (Fig. 1, Pl. XXII), and lakes and lagoons at numerous points
both on the Atlantic and the Pacific coasts (Fig. 2, Pl. XXII, Rodeo
lagoon). The same phenomena are to be seen along many lake shores.
Bars sometimes tie islands to the mainland (Pl. XXIII, Fig. 1, Nahant,
Mass.; Fig. 2, near Biddeford, Me.). The structure of a bar as seen in
cross-section is shown in Fig. 316.
The construction of a spit has been aptly compared to the construction
of a railway embankment across a depression. The material is first
carried out from the bordering upland (shallow water) and dumped where
the slope to the depression (deep water) begins. The embankment thus
begun is extended by the carrying out of new material, which is left at
the end of the dump already made.
If the bay across which the bar is built receives abundant drainage
from the land, the outflow from the bay may be sufficient to prevent
the completion of the bar (Fig. 2, Pl. XXII), for when the growth of
the spit has sufficiently narrowed the outlet of the bay, the sediment
brought to the end of the spit by the littoral current will be swept
out beyond the spit by the current setting out from the bay.
The completion of a bar may be interfered with by tidal currents, even
without land-drainage. Currents generated by the tides may sweep in
or out of the bay with increased force as the entrance is narrowed,
carrying in or out the sediment which the littoral current would have
left at the end of the spit. The scour of the tides often insures
deep entrances (inlets) to bays, and maintains definite channels or
“thorofares” in the lagoon marshes behind barriers and spits. The
sediment brought down from the land, as well as that washed in by tidal
currents and waves, tends to fill up the lagoon behind a barrier, a
spit, or a bar, converting it into land (Fig. 317).
[Illustration: +Fig.+ 317.—Sketch of a portion of the New Jersey coast.
The dotted belt next the sea is the barrier, modified by the wind.
The area marked by the diagonal lines is the mainland. In the marshy
area between, there are numerous channels or “thorofares” kept open
by the currents. The figures show the depths of water in feet. Scale
about ⅜ inch = 1 mile.]
Since spits and bars are built only where there is shore-drift in
transit, they are always built out from a beach or barrier. The distal
end of the bar may also join a beach or barrier. Traced back to
its source, the beach from which a spit leads out is often found to
terminate in the cliff from which the material of the beach and the
spit were derived (Pl. XX and Fig. 2, Pl. XXII). In such cases the
sediment of the beach has been shifted but a short distance; but in
other cases it has traveled far.
[Illustration: +Fig.+ 318.—Map of shore-terraces, largely wave-built.
Lake Bonneville. (Gilbert.)]
[Illustration: +Fig.+ 319.—A portion of the Texas coast showing the
tendency of shore-deposition to simplify the coast-line. The deposits
(the narrow necks of land parallel to the coast) shut in the bays.
(From chart of C. and G. Surv.)]
[Illustration: +Fig.+ 320.—Map showing that in the early stages of the
simplification of a shore-line the irregularities are increased. The
numbers indicate the depth of water in fathoms. (From chart of C. and
G. Surv.)]
The spit is usually either straight or in conformity with the general
course of the shore-current, but since the littoral current itself is
subject to alteration as the result of shifting winds, the spit may
depart from straightness. Winds which simply reverse the direction of
the littoral current retard its construction, but may not otherwise
affect it; but if a strong current be made to flow past the end of a
spit, it may cut away its extremity and rebuild the materials into a
smaller spit, joining the main one at an angle. This gives rise to a
_hook_ (Fig. 315). Successive storms may develop successive hooks along
the side of a growing spit. The end of a hook may be so extended as to
join the mainland, when it becomes a _loop_.
=Wave-built terraces.=—Under the influence of off-shore currents,
littoral currents may be drawn from the coast-line. If such a current
continues as a well-defined surface-current, it builds a spit, but
if it spreads, it tends to build a terrace. The accumulation then
is not at the end of a beach, as in the case of a spit, but on its
side, and the result of the deposition is to carry the beach seaward.
The undertow abets the process. The widened beach is a _wave-built
terrace_. The wave-built terrace often borders the wave-cut terrace
along its seaward margin (Figs. 303 and 318). With the help of waves,
the surface of the terrace may be built up into land by the expansion
of the crest of the beach. Terrace-cutting and terrace-building are
both involved in the development of the continental shelves.
Beach ridges, spits, bars, etc., like sea-cliffs and wave-cut terraces,
are often preserved after the relative level of sea and land has
changed. If the shore has risen, relatively or absolutely, these
features are relied on as evidences of the change. If shore features
be submerged instead of elevated, they furnish less accessible, though
not less real, evidence of the change of level. Similar features
about lakes have a like significance, but in this case it is often
demonstrable that it is the water rather than the land which has
changed its level.
_Effect of Shore-deposition on Coastal Configuration._
The tendency of shore-deposition is to cut off bays and to straighten
and simplify the shore-lines. This is abundantly illustrated along
the Atlantic and Gulf coasts of the United States (see Fig. 319 and
Pl. XXII). It is to be noted, however, that in the simplification of
the shore-line through deposition, the initial stages often result
in great irregularity (Fig. 320 and Pl. XXIII). In some cases, the
irregularities are not temporary. Thus deltas (p. 198), though
not wholly the work of sea- (or lake-) water, often constitute
irregularities of a more or less permanent nature. This is the case
where they project beyond the general trend of the coast-line. Where,
on the other hand, they are built at the heads of bays, they tend to
simplify the coast-line by obliterating the indentation. The delta at
the head of the Gulf of California is an example. So too is the delta
of the Mississippi, the real head of which is far above the present
debouchure of the stream. The form of the delta in ground-plan depends
on the horizontal configuration of the coast where it is developed, on
the strength of the waves and shore-currents, and on their relation
to the amount of detritus contributed by the stream concerned. Good
illustrations are furnished by the Gulf of Mexico where the deltas of
the Mississippi and Rio Grande are in contrast.
So far as concerns the vertical configuration of coasts, erosion and
deposition are in contrast, for while the former tends to develop
steep, irregular, and often high slopes (p. 349) from the land to
the sea, the latter tends to develop gentle, regular, and low ones.
A partial exception to the latter part of this general statement
comes about through the building of dunes, the material for which is
furnished by the waves.
SUMMARY OF COASTAL IRREGULARITIES.
The horizontal irregularities of coasts are both large and small.
Some of them, like Florida, Sandy Hook, etc., consist primarily of
projections of land into the sea; others, like Chesapeake Bay, the Gulf
of Mexico, and Puget Sound, are projections of the sea into the land;
while still others, like the Gulf of California and its associated
peninsula, cannot readily be put in either of the foregoing classes.
Some of the irregularities of the land border, such as Yucatan, are
more or less nearly normal to the general trend of the coast which they
affect, while others, such as the “beaches” along the Atlantic and
Gulf coasts of the United States (Figs. 319 and 320), are more or less
nearly parallel with it. Some of the irregularities, especially some of
the small ones, are more or less angular in their outline (Pl. XX and
parts of Fig. 2, Pl. XXII), while others are bounded by curves instead.
In many cases more than one factor has been involved in the development
of irregularities. In the case of great irregularities, diastrophism
has generally been the dominant factor. The Gulf of Mexico and the
Mediterranean Sea perhaps represent differential subsidence, while
Florida and the Iberian peninsula represent differential uplift
(relative, though perhaps not absolute). The narrow bays which indent
many coasts generally represent the subsidence of a region previously
affected by valleys (Fig. 297). Many of them, such as Narragansett,
Delaware, and Chesapeake Bays, are primarily the drowned ends of
river valleys, while others, such as Puget Sound,[167] are primarily
structural valleys (synclines). Many of the long and narrow bays or
fiords common in the high latitudes of North America and Europe (Fig.
266, p. 293) appear to be the drowned ends of valleys previously
deepened by glaciers. The drowned ends of river canyons, and the
submerged parts of valleys excavated (not sunk) beneath the sea by
glaciers, would also be fiords.
[Illustration: PLATE XXIII.
Fig. 1. MASSACHUSETTS.
U. S. Geol. Surv.
Fig. 2. MAINE.
U. S. Geol. Surv.]
[Illustration: PLATE XXIV.
PORTION OF THE COAST OF MAINE.
U. S. Geol. Surv.]