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.
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
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.
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
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
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
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
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 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
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
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
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
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
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
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