رسوب شناسی-Sedimentology
Sedimentology
The study of natural
sediments, both lithified (sedimentary rocks) and unlithified, and of the
processes by which they are formed. Branches of this discipline have been known
as sedimentation or sedimentary petrology, or they have been included in
stratigraphy. Sedimentology includes all those processes that give rise to
sediment or modify it after deposition: weathering, which breaks up or dissolves
preexisting rocks so that sediment may form from them; mechanical
transportation; deposition; and diagenesis, which modifies sediment after
deposition and burial within a sedimentary basin and converts it into
sedimentary rock. Sediments deposited by mechanical processes (gravels, sands,
muds) are known as clastic sediments, and those deposited predominantly by
chemical or biological processes (limestones, dolomites, rock salt, chert) are
known as chemical sediments. Figure 1 shows a diagram of the processes that are
involved in the sedimentary cycle. See also: Stratigraphy; Weathering processes
Fig. 1 Diagram showing processes that
constitute the sedimentary cycle. Heavy arrows indicate mass fluxes of material.
Light arrows indicate where deposition is taking
place.
Sedimentary
cycle
The raw materials of
sedimentation are the products of weathering of previously formed igneous,
metamorphic, or sedimentary rocks. In the present geological era, 66% of the
continents and almost all of the ocean basins are covered by sedimentary rocks.
Therefore, most of the sediment now forming has been derived by recycling
previously formed sediment. Identification of the oldest rocks in the Earth's
crust, formed more than 3 × 109 years ago, has shown that this process has been
going on at least since then. Old sedimentary rocks tend to be eroded away or
converted into metamorphic rocks, so that very ancient sedimentary rocks are
seen at only a few places on Earth. The area (or volume) of sedimentary rocks
declines exponentially with age, and the half-life of a sedimentary rock is
about 2 × 108 years. See also: Earth crust; Igneous rocks; Metamorphic rocks
Major
controls
The major controls on
the sedimentary cycle are tectonics, climate, worldwide (eustatic) changes in
sea level, the evolution of environments with geological time, and the effect of
rare events.
Tectonics
These are the
large-scale motions (both horizontal and vertical) of the Earth's crust.
Tectonics are driven by forces within the interior of the Earth but have a large
effect on sedimentation. These crustal movements largely determine which areas
of the Earth's crust undergo uplift and erosion, thus acting as sources of
sediment, and which areas undergo prolonged subsidence, thus acting as
sedimentary basins. Rates of uplift may be very high (over 10 m or 33 ft per
1000 years) locally, but probably such rates prevail only for short periods of
time. Over millions of years, uplift even in mountainous regions is about 1 m
(3.3 ft) per 1000 years, and it is closely balanced by rates of erosion. Rates
of erosion, estimated from measured rates of sediment transport in rivers and
from various other techniques, range from a few meters per 1000 years in
mountainous areas to a few millimeters per 1000 years averaged over entire
continents. Though these rates are very slow by human standards, over periods of
millions of years they produce immense quantities of sediment; all the
continents would soon be reduced to sea level if it were not for the renewed
uplift produced by tectonics. It is estimated that the total volume of all
sediment and sedimentary rocks is over 6.5 × 1010 km3 (1.55 × 108 mi3), a volume
that would cover the Earth to a depth of 1.3 km (0.82 mi) if spread evenly over
its surface. Recent studies of the Moon and planets show that the Earth is very
active tectonically compared with other bodies of similar size in the solar
system. See also: Basin; Plate tectonics
Climate
This plays a secondary
but important role in controlling the rate of weathering and sediment
production. The more humid the climate, the higher these rates are; evidence
exists, however, that in some continental climates (for example, in the
southwestern
Sea
level
Tectonics and climate
together control the relative level of the sea. In cold periods, water is stored
as ice at the poles, which can produce a worldwide (eustatic) lowering of sea
level by more than 100 m (330 ft). Changes in the rate of sea-floor spreading
can produce large alterations in sea level by changing the average depth of the
ocean basins. Widespread, but not worldwide, uplift or subsidence can also
result in emergence or flooding of substantial areas of the continents. Changes
in sea level, whether local or worldwide, strongly influence sedimentation in
shallow seas and along coastlines; sea-level changes also affect sedimentation
in rivers by changing the base level below which a stream cannot erode its bed.
Evolution of
environments
One of the major
conclusions from the study of ancient sediments has been that the general nature
and rates of sedimentation have been essentially unchanged during the last
billion years of geological history. However, this conclusion,
uniformitarianism, must be qualified to take into account progressive changes in
the Earth's environment through geological time, and the operation of rare but
locally or even globally important catastrophic events. The most important
progressive changes have been in tectonics and atmospheric chemistry early in
Precambrian times, and in the nature of life on the Earth, particularly since
the beginning of the Cambrian. See also: Cambrian
Rare
events
Throughout geological
time, events that are rare by human standards but common on a geological time
scale, such as earthquakes, volcanic eruptions, and storms, produced widespread
sediment deposits. There is increasing evidence for a few truly rare but
significant events, such as the rapid drying up of large seas (parts of the
Sediment
movement and deposition
Sediment is moved either
by gravity acting on the sediment particles or by the motions of fluids (air,
water, flowing ice), which are themselves produced by gravity. Deposition takes
place when the rate of sediment movement decreases in the direction of sediment
movement; deposition may be so abrupt that an entire moving mass of sediment and
fluid comes to a halt (mass deposition, for example, by a debris flow), or so
slow that the moving fluid (which may contain only a few parts per thousand of
sediment) leaves only a few grains of sediment behind.
Fluids transport
sediment mainly by traction (sliding, rolling, bouncing on the bed) and
suspension caused by turbulence. Though the irregular, turbulent motion of
fluids such as air and water are the most important way that sediment is held up
within a mass of flowing fluid, other mechanisms such as buoyancy and impacts
between the grains may also operate. See also: Turbulent flow
A fundamental dynamic
property of sediment grains is their settling velocity. This property is the
constant velocity of fall that a grain attains in a fluid at rest, when the
force of gravity is balanced by the forces of buoyancy and fluid resistance. The
settling velocity depends on the density and viscosity of the fluid, as well as
on the size, shape, and density of the grains. As measured under standard
conditions (in water at 20°C or 68°F), it can be used directly to characterize
sediment or be converted into an equivalent diameter (the diameter of a quartz
sphere with the same settling velocity as the measured value). Settling
velocities in water vary from less than 0.1 mm (0.004 in.) per second for silt,
to 5–10 cm (2–4 in.) per second for average sand, to meters per second for large
boulders.
Transport and
deposition by flowing water
The different ways in
which sediment is transported by water flows can be illustrated by describing
the changes that are observed in a laboratory channel (flume) as the discharge
of water is gradually increased over a flat bed of sand. The flow strength just
necessary to start movement of grains on the bed defines the competence of the
flow for a particular type (size, shape, density) of sediment. The first grain
motion is by sliding or rolling of the grains in contact with the bed, and this
is soon followed by a leaping motion known as saltation. In saltation, grains
rise up at a steep angle from the bed, are caught by the flow, and drop back to
the bed; they follow a so-called ballistic trajectory. Sliding, rolling, and
saltation generally are grouped together as traction. Sediment moved in this way
is described as bed load. At low flow strengths, the grains do not rise far
above the bed (only a few grain diameters in water) and they move much more
slowly than the fluid. Rates of sediment movement (as measured by the sediment
discharge, or mass rate of transport of sediment past a particular cross section
of the channel) depend on a high power of flow strength. Bed load is supported
by direct contact with the bed (and by buoyancy). Its motion responds to the
fluctuating shear stress on the bed produced by turbulence in the flow rate, but
it is otherwise not much affected by turbulent fluid motions (Fig. 2).
Fig. 2 Diagram of types of load carried by a
stream. (After R. M. Garrels, A Textbook of Geology, Harper and Row,
1951)
Fig. 3 Bedforms a–d produced by flow over a
sand bed. (After H. Blatt, G. V. Middleton, and R. Murray, Origin of Sedimentary
Rocks, 2d ed., Prentice-Hall, 1980)
As flow strength is
increased, however, the upward velocity components of turbulent eddies become
strong enough to overcome the settling velocity of the sediment, and the grains
are taken into suspension. A part of the sediment load is still moved by
traction, but a larger and larger part is moved by suspension (suspended load).
Suspended sediment, at first concentrated near the bed, becomes more uniformly
distributed in the flow as the flow strength increases. The total rate at which
sediment can be moved (capacity) depends upon the availability of erodible
sediment in the banks or bed and upon the sediment properties and flow strength.
Sediment discharge continues to increase with flow strength, but not as rapidly
as in the early stages of sediment transport. There is almost no limit to the
amount of fine sediment that can be carried in suspension: concentrations as
high as 30% by weight have been recorded. As concentration increases, so does
the likelihood of collisions between the coarser grains. It has been suggested
that such collisions become an important support mechanism (dispersive pressure)
for sand and gravel at volume concentrations higher than about 10%, and such
concentrations are certainly present close to the bed at high flow strengths.
As the rates of sediment
transport increase, the nature of the bed also changes. Even if the bed is flat
at first, it develops asymmetrical waves or bedforms migrating downstream (Fig.
3). Small current ripples (10–30 cm or 4–12 in. long) form soon after sediment
movement begins, but at higher flow strengths they are replaced by large dunes
on the bed; in deep flows, these may reach heights of several meters and lengths
of tens of meters. At very high flow strengths, however, the dunes are washed
out and the bed becomes almost flat again. If the flow becomes supercritical,
nearly symmetrical bedforms (antidunes) migrating upstream may be formed. See
also: Depositional systems and environments; Froude number; Stream transport and
deposition
Transport by
waves
Water waves produce an
oscillating motion of water close to the bed. Therefore, movement of sediment by
waves differs in detail from that produced by unidirectional flows. Net sediment
movement generally still results, however, because of the asymmetrical nature of
the oscillating flow close to the bed. Waves produce stages of sediment
transport similar to those resulting from unidirectional flows. Suspension
replaces traction at high levels of wave activity. Bedforms show a transition
from small to larger ripples, and a plane bed reappears at the highest wave
intensities. Breaking of waves greatly enhances the ability of waves to carry
sediment into suspension. After breaking, waves form a bore (swash) that washes
sediment up the beach, and down again in the backwash. The approach of waves is
generally oblique to the shore and produces a movement of sand along the beach
that is known as longshore drift. See also: Nearshore processes
Transport by
wind
Sediment transport by
wind is similar to transport by water, but different forms of transport
predominate because of the low density of air compared with water, and the great
thickness and high speed attained by moving air masses. Whereas sand is easily
transported in suspension by water flowing at speeds of the order of 1 m (3.3
ft) per second, it can be suspended only rarely by winds. Most sand is moved by
wind in saltation, which is a far more effective mechanism of sediment transport
in air than in water. Saltating sand grains rise to heights of several
centimeters and return to the bed with enough momentum to drive forward other
grains, a process known as surface creep. Ripples produced by saltating grains
are smaller and more regular than those produced by flow of water, but dunes
develop to much larger sizes and show a wide variety of forms because of the
thickness of the wind boundary layer and the variability of wind directions. See
also: Air mass
Only clay and silt are
easily carried into suspension by the wind, but once suspended, these grains may
be transported for hundreds of kilometers. Fine sediment (ash) is produced by
explosive volcanic eruptions and carried high into the upper atmosphere, where
it is transported great distances by the wind. Deposits of silt (loess) several
meters thick were formed at the end of the last ice age by winds that blew
sediment derived from rivers that drained the ice sheets. See also: Dune; Dust
storm; Loess; Wind
Transport by
ice
Glaciers and ice sheets
incorporate sediment that falls onto the ice surface or is eroded by grinding or
plucking from material beneath the ice. Such material is transported en masse to
the point where the glacier melts. Sediment deposited by melting directly out of
ice is almost completely unsorted and is known as till. Some ice sheets melt
directly into lakes or the sea, producing debris that is modified as it settles
to the bottom. Much of the sediment that is derived from glaciers is reworked by
mass flow (flow till) or by meltwater (outwash). Some of the sediment eroded by
glaciers consists of finely ground material known as rock flour. The abundance
of this material in lakes fed by glaciers is at least partly responsible for
their intense green color. See also: Glaciology; Till
Sediment gravity
flows
These are flows whose
motion results from gravity acting on the sediment grains rather than on the
fluid in the flow (indeed, no fluid is necessary for certain types of gravity
flow, and so these flows can operate in fluid-free environments, such as the
Moon). Four classes of sediment gravity flow have been recognized,
differentiated by the mechanisms that support sediment above the bed. If the
mechanism is grain impacts, the flow is a grain flow; if it is turbulence, the
flow is a turbidity current; if it is the upward escape of interstitial fluid,
the flow is a liquefied flow; and if it is the strength of the matrix between
the grains (for example, a mud matrix), the flow is a plastic debris flow. Such
flows are generally short-lived, but they may be very large. Some recorded
turbidity currents have transported as much as 100 km3 (24 mi3) of sediment and
have flowed for several hundred kilometers over the ocean floor. See also:
Turbidity current
Modification by
transport
Sediment is modified
during transport by chemical action (continuation of the weathering process
begun in the source) and by physical breakage, abrasion, and sorting. Physical
impact is most effective on the larger grains, and it has little effect on
grains small enough to be transported in suspension. Boulders and pebbles are
rapidly rounded by abrasion of the corners and edges of the grain, and prolonged
transport in rivers may lead to a nearly spherical shape. The more friable
minerals found in sand (for example, feldspar) may be broken, but the main
mineral, quartz, is very resistant to abrasion. Nevertheless, studies using the
scanning electron microscope have shown evidence of rounding and submicroscopic
impact marks, and reduction in size by flaking of small chips from the surface.
Transport by wind is far more effective than transport by water in rounding
quartz grains, and reworking by waves on beaches is more effective than
transport down rivers—even when the transport distance is hundreds of kilometers
down major rivers.
Sediment weathered from
an igneous or metamorphic source rock, or produced from it by mechanical
breakage, tends to have a wide range of grain sizes that follow a statistical
distribution known as Rosin's distribution. Though a similar distribution may be
detected in some cases after several kilometers of river transport, generally it
is rapidly modified by hydraulic sorting, which tends to segregate the variously
sized fractions by their different mechanisms of transport. In a river, for
example, the coarsest grains are confined to the deepest part of the channel;
the sand is confined to bars within the channel or to the natural levees; and
either the mud is transported in suspension rapidly through the system to the
sea, or it is deposited during floods in overbank deposits on the floodplain.
Because each of these deposits consists mainly on only one fraction of the
original source material, they are better sorted and the size distribution is
changed. Most naturally occurring size distributions appear to be composed of
mixtures of more than one (lognormal) distribution, resulting from different
sources or mechanisms of deposition. See also: Fluvial sediments
Flocculation
Gravel and sand grains
are generally chemically inert and noncohesive. This is not true of the finest,
clay-sized particles. Colloid-sized particles (finer than 0.2 micrometer)
generally have electrically charged surfaces, even if composed of inert
minerals. Many clay minerals, such as montmorillonite, readily react with
dissolved ions. The surface charge depends in part on the nature and
concentration of ions in the surrounding solution. Though the grains of clay
minerals are very small (less than 1 μm), electrochemical effects may cause them
to aggregate into large flocs whose settling velocity (though still relatively
low, because of the low bulk density of the floc) may be many times higher than
that of the individual grains. Flocculation is very common as clay is delivered
by rivers into the sea; thus it is an important process in estuaries and deltas.
See also: Clay minerals; Delta; Estuarine oceanography; Gravel; Sand
Chemical
sedimentation
Chemical weathering
dissolves rock materials and delivers ions in solution to lakes and the ocean.
The concentrations of ions in river and ocean water are quite different, showing
that some ions must be removed by sedimentation. Comparison of the modern rate
of delivery of ions to the ocean, with their concentration in the oceans, shows
that some are removed very rapidly (residence times of only a few thousand
years) whereas others, such as chlorine and sodium, are removed very slowly
(residence times of hundreds of millions of years). The main mechanisms by which
ions are removed from solution are biochemical precipitation, evaporation,
adsorption or exchange with clays and other rock materials (including glass in
fine volcanic ash), and reaction with hot rock in hydrothermal systems
associated with mid-oceanic ridges. The most important chemical sediments are
calcareous (limestones and dolomites), followed by siliceous chert deposits and
evaporites, but there are a very large variety of chemical sediments whose
origins are not yet fully understood. Chemical dissolution and precipitation are
not restricted to weathering and primary sedimentation from lake water or
seawater, but are equally important in diagenesis, that is, in transforming
sediment after deposition into sedimentary rock. See also: Marine sediments;
Mid-Oceanic Ridge
Carbonate
sediments
Organisms began to
secrete skeletons composed of calcium carbonate during the Cambrian Period, and
since then most carbonate sediment has been precipitated biochemically. Though
the warm surface waters of the oceans are known to be supersaturated with
respect to calcite (the thermodynamically stable form of calcium carbonate in
most natural waters), direct chemical precipitation of calcite is known in only
a few environments (for example, limestone caves). The common carbonate minerals
in modern sediments are biochemically precipitated aragonite and calcite.
Aragonite is metastable under surface conditions and is generally dissolved away
within a few thousand years (rarely millions) to be replaced by diagenetic
calcite or dolomite. See also: Limestone
The best-known
biochemical accumulations of carbonate at present are the tropical coral-algal
reefs, but a wide variety of organisms precipitate carbonate from seawater now
or they have done so in the past. Besides reef organisms, pelecypods, algae, and
foraminifera are important in cooler marine waters; and in the surface layers of
the oceans, planktonic foraminifera, pteropods, and algae (coccolithophorids)
are important today, and were so at other periods of the Cenozoic and Mesozoic.
Cretaceous chalks, for example, are composed largely of microscopic coccoliths
and foraminifera, not very different from modern deep-sea calcareous oozes
(though most ancient chalks were deposited in water only a few hundred meters
deep). In other geological periods, reefs were built by different types of
corals, by specialized pelecypods (rudists), by algae, and by now-extinct groups
such as stromatoporoids. See also: Cenozoic; Chalk; Cretaceous; Mesozoic; Reef
The most important
differences between ancient carbonate rocks and modern carbonate sediments,
however, are due to diagenesis. Modern carbonate sediments are composed mainly
of aragonite and a form of calcite containing several percent of magnesium.
Ancient carbonates consist largely of calcite low in magnesium and of dolomite,
a calcium-magnesium carbonate only rarely present in modern sediments. The
mineralogical changes necessary to produce typical ancient carbonates take place
rapidly, particularly under the action of meteoric water. Even the formation of
dolomite, which generally requires the wholesale dissolution of aragonite and
calcite and the exchange of large quantities of magnesium from water passing
through the porous rock, has been a common and frequently rapid process
throughout geological history. See also: Aragonite; Authigenic minerals;
Calcite; Carbonate minerals; Diagenesis; Dolomite; Marl
Siliceous
sediment
Sediments composed
largely of fine-grained quartz (chert) are common in the geological record,
extending back to the earliest Precambrian. Since the beginning of the Cambrian,
most cherts have been formed by biochemical precipitation by organisms that
secrete an opaline skeleton, such as radiolarians, diatoms, and some sponges,
followed by diagenetic dissolution and reprecipitation as microcrystalline
quartz and chalcedony. See also: Chalcedony; Chert
Evaporites
Evaporation of seawater
in the laboratory gives rise to a sequence of different precipitates as
evaporation proceeds. Small amounts of calcareous sediments are first
precipitated, followed by gypsum, anhydrite, halite, and finally by complex
salts containing sodium, potassium, magnesium, sulfate ions, and chloride ions.
A close match with this experimentally determined sequence is known to exist in
several ancient deposits, though the relative abundances of different minerals
vary considerably. This indicates that these deposits were formed by evaporation
of seawater, and that the composition of the ancient seawater was not very
different from that of modern seawater. Evaporites can presently be observed
forming in several shallow-water environments (lakes and marine lagoons), though
none is comparable in magnitude to the evaporite basins inferred for some
periods in the past. Formation of large evaporite deposits was probably a rare
event, but once conditions were appropriate, large volumes could be formed in
only a few tens of thousands of years. See also: Saline evaporites; Seawater
Other chemical
sediments
Other chemical sediments
are rich in iron, phosphorus, manganese, and carbon. The carbonaceous sediments
include marine muds rich in organic matter, which are important as the source
rocks of petroleum, and nonmarine peats, some of which give rise to commercial
coals. Many sedimentary rocks of unusual character are economically important;
they serve as ores of iron, copper, manganese, phosphate, gold, diamonds, and so
forth. Equally important economically are more common sedimentary rocks, which
serve as sources of building stone, crushed rock for construction, limestone for
smelting, pure quartz sands, and so forth. See also: Coal; Ore and mineral
deposits; Petroleum
Biological
effects
Many so-called chemical
sediments are actually produced by biochemical action. Much is then reworked by
waves and currents, so that the chemical sediment shows clastic textures and
consists of grains rounded and sorted by transport. Depositional and diagenetic
processes, however, are often strongly affected by organic action, no matter
what the origin of the sediment. Plants in both terrestrial and marine
environments tend to trap sediment, enhancing deposition and slowing erosion.
Organisms burrow through the sediment, producing distinctive structures
(bioturbation) and modifying or destroying original clastic structures and
textures. Bacteria promote decay of organic particles, changing the chemical
environment within the sediment, and producing gases such as carbon dioxide and
methane. At deeper levels of burial within a sedimentary basin, the residual
organic matter in the sediments (kerogen or sapropel) is broken down by heat to
produce liquid and gaseous hydrocarbons and carbon dioxide, which may migrate
many tens of kilometers upward along permeable strata within the basin. See
also: Kerogen; Sapropel
Sedimentary
environments and facies
Sedimentary rocks
preserve the main direct evidence about the nature of the surface environments
of the ancient Earth and the way they have changed through geological time.
Thus, besides trying to understand the basic principles of sedimentation,
sedimentologists have studied modern and ancient sediments as records of ancient
environments. For this purpose, fossils and primary sedimentary structures are
the best guide. These structures are those formed at the time of deposition, as
opposed to those formed after deposition by diagenesis, or by deformation. In
describing sequences of sedimentary rocks in the field (stratigraphic sections),
sedimentologists recognize compositional, structural, and organic aspects of
rocks that can be used to distinguish one unit of rocks from another. Such units
are known as sedimentary facies, and they can generally be interpreted as having
formed in different environments of deposition. Though there are a large number
of different sedimentary environments, they can be classified in a number of
general classes, and their characteristic facies are known from studies of
modern environments. Furthermore, the transition from one environment to another
is not a random process, because only those environments that originally existed
side by side can succeed each other without a break (Walther's law). Thus, it is
possible to produce plausible reconstructions of environments that existed many
millions of years ago. Like all historical reconstructions based on fragmentary
evidence, these depend upon the imaginative skill of the investigator, as well
as upon reliable deductions from data. See also: Facies (geology); Trace fossils
Basin
analysis
This describes studies
of the larger-scale aspects of sediment accumulation, such as the different
types of sedimentary basins, and how they form, and the types of source areas
and how they are linked to particular types of basins. Clastic sediments derived
from preexisting rocks (terrigenous sediments) have textures and mineralogy that
preserve clues about the nature of the source from which they were derived
(their provenance). Sedimentary structures and textures can be used to indicate
the direction of sediment transport at the place of deposition (paleocurrents).
By carefully dating the deposits, or at least determining which deposits are the
same age (stratigraphic correlation), by studying their provenance and
paleocurrents, and by reconstructing their depositional environments, it is
possible to reconstruct how a sedimentary basin developed and where the sediment
came from. In addition, the long-term rates of deposition can be determined, and
by studying mineralogical changes produced by diagenesis it is possible to
reconstruct the sequence of pressures and temperatures that have affected the
sediments in the basin. Tectonic theories suggest models that simulate the
patterns of deposition and diagenesis within sedimentary basins. See also:
Sedimentary rocks
G. V. Middleton
Bibliography
P. A. Allen, Earth
Surface Processes, 1997
H. Blatt, G. V.
Middleton, and R. Murray, Origin of Sedimentary Rocks, 2d ed.,
1980
M. R. Leeder,
Sedimentology: Process and Product, 1982
M. R. Leeder,
Sedimentology and Sedimentary Basins: From Turbulence to Tectonics,
1999
R. K. Matthews, Dynamic
Stratigraphy: An Introduction to Sedimentation and Stratigraphy, 2d ed.,
1984
A. D. Miall, Principles of Sedimentary Basin Analysis, 2d ed., 1990
Alifazeli=egeology.blogfa.com
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