Marine sediment
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Sediments |
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Marine sediment, or ocean sediment, or seafloor sediment, are deposits of insoluble particles that have accumulated on the
Except within a few kilometres of a
Rates of sediment accumulation are relatively slow throughout most of the ocean, in many cases taking thousands of years for any significant deposits to form. Sediment transported from the land accumulates the fastest, on the order of one metre or more per thousand years for coarser particles. However, sedimentation rates near the mouths of large rivers with high discharge can be orders of magnitude higher. Biogenous oozes accumulate at a rate of about one centimetre per thousand years, while small clay particles are deposited in the deep ocean at around one millimetre per thousand years.
Sediments from the land are deposited on the
Overview
Except within a few kilometres of a
The various sources of seafloor sediment can be summarized as follows: [2]
- Terrigenous sediment is derived from continental sources transported by rivers, wind, ocean currents, and glaciers. It is dominated by quartz, feldspar, clay minerals, iron oxides, and terrestrial organic matter.
- Pelagic carbonate sediment is derived from organisms (e.g., foraminifera) living in the ocean water (at various depths, but mostly near surface) that make their shells (a.k.a. tests) out of carbonate minerals such as calcite.
- Pelagic silica sediment is derived from marine organisms (e.g., diatoms and radiolaria) that make their tests out of silica (microcrystalline quartz).
- Volcanic ash and other volcanic materials are derived from both terrestrial and submarine eruptions.
- Iron and manganese nodules form as direct precipitates from ocean-bottom water.
The distributions of some of these materials around the seas are shown in the diagram at the start of this article ↑. Terrigenous sediments predominate near the continents and within inland seas and large lakes. These sediments tend to be relatively coarse, typically containing sand and silt, but in some cases even pebbles and cobbles. Clay settles slowly in nearshore environments, but much of the clay is dispersed far from its source areas by ocean currents. Clay minerals are predominant over wide areas in the deepest parts of the ocean, and most of this clay is terrestrial in origin. Siliceous oozes (derived from radiolaria and diatoms) are common in the south polar region, along the equator in the Pacific, south of the Aleutian Islands, and within large parts of the Indian Ocean. Carbonate oozes are widely distributed in all of the oceans within equatorial and mid-latitude regions. In fact, clay settles everywhere in the oceans, but in areas where silica- and carbonate-producing organisms are prolific, they produce enough silica or carbonate sediment to dominate over clay.[2]
Carbonate sediments are derived from a wide range of near-surface pelagic organisms that make their shells out of carbonate. These tiny shells, and the even tinier fragments that form when they break into pieces, settle slowly through the water column, but they don't necessarily make it to the bottom. While calcite is insoluble in surface water, its solubility increases with depth (and pressure) and at around 4,000 m, the carbonate fragments dissolve. This depth, which varies with latitude and water temperature, is known as the carbonate compensation depth. As a result, carbonate oozes are absent from the deepest parts of the ocean (deeper than 4,000 m), but they are common in shallower areas such as the mid-Atlantic ridge, the East Pacific Rise (west of South America), along the trend of the Hawaiian/Emperor Seamounts (in the northern Pacific), and on the tops of many isolated seamounts.[2]
Texture
Sediment texture can be examined in several ways. The first way is
A third way to describe marine sediment texture is its maturity, or how long its particles have been transported by water. One way which can indicate maturity is how round the particles are. The more mature a sediment the rounder the particles will be, as a result of being abraded over time. A high degree of sorting can also indicate maturity, because over time the smaller particles will be washed away, and a given amount of energy will move particles of a similar size over the same distance. Lastly, the older and more mature a sediment the higher the quartz content, at least in sediments derived from rock particles. Quartz is a common mineral in terrestrial rocks, and it is very hard and resistant to abrasion. Over time, particles made from other materials are worn away, leaving only quartz behind. Beach sand is a very mature sediment; it is composed primarily of quartz, and the particles are rounded and of similar size (well-sorted).[1]
Origins
Marine sediments can also classified by their source of origin. There are four types: [3][1]
- Lithogenous sediments, also called terrigenous sediments, are derived from preexisting rock and come from land via rivers, ice, wind and other processes. They are referred to as terrigenous sediments since most comes from the land.
- Biogenous sediments are composed of the remains of marine organisms, and come from organisms like plankton when their exoskeletons break down
- Hydrogenous sediments come from chemical reactions in the water, and are formed when materials that are dissolved in water precipitate out and form solid particles.
- Cosmogenous sediments are derived from extraterrestrial sources, coming from space, filtering in through the atmosphere or carried to Earth on meteorites.[3][1]
Lithogenous
Lithogenous or terrigenous sediment is primarily composed of small fragments of preexisting rocks that have made their way into the ocean. These sediments can contain the entire range of particle sizes, from microscopic clays to large boulders, and they are found almost everywhere on the ocean floor. Lithogenous sediments are created on land through the process of weathering, where rocks and minerals are broken down into smaller particles through the action of wind, rain, water flow, temperature- or ice-induced cracking, and other erosive processes. These small eroded particles are then transported to the oceans through a variety of mechanisms: [1]
Streams and rivers: Various forms of runoff deposit large amounts of sediment into the oceans, mostly in the form of finer-grained particles. About 90% of the lithogenous sediment in the oceans is thought to have come from river discharge, particularly from Asia. Most of this sediment, especially the larger particles, will be deposited and remain fairly close to the coastline, however, smaller clay particles may remain suspended in the water column for long periods of time and may be transported great distances from the source.[1]
Wind:
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The face of blue glacial ice melting into the sea
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River discharge in the Yukon Delta, Alaska. The pale color demonstrates the large amounts of sediment released into the ocean via the rivers.
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A plume of wind-borne particles from Sudan (left) blow over the Red Sea
Glaciers and ice rafting: As glaciers grind their way over land, they pick up lots of soil and rock particles, including very large boulders, that get carried by the ice. When the glacier meets the ocean and begins to break apart or melt, these particles get deposited. Most of the deposition will happen close to where the glacier meets the water, but a small amount of material is also transported longer distances by rafting, where larger pieces of ice drift far from the glacier before releasing their sediment.[1]
Gravity: Landslides, mudslides, avalanches, and other gravity-driven events can deposit large amounts of material into the ocean when they happen close to shore.[1]
Waves: Wave action along a coastline will erode rocks and will pull loose particles from beaches and shorelines into the water.[1]
Volcanoes: Volcanic eruptions emit vast amounts of ash and other debris into the atmosphere, where it can then be transported by wind to eventually get deposited in the oceans.[1]
Gastroliths: Another, relatively minor, means of transporting lithogenous sediment to the ocean are gastroliths. Gastrolith means "stomach stone". Many animals, including seabirds, pinnipeds, and some crocodiles deliberately swallow stones and regurgitate them latter. Stones swallowed on land can be regurgitated at sea. The stones can help grind food in the stomach or act as ballast regulating buoyancy. Mostly these processes deposit lithogenous sediment close to shore. Sediment particles can then be transported farther by waves and currents, and may eventually escape the continental shelf and reach the deep ocean floor.[1]
- Composition
Lithogenous sediments usually reflect the composition of whatever materials they were derived from, so they are dominated by the major minerals that make up most terrestrial rock. This includes quartz, feldspar, clay minerals, iron oxides, and terrestrial organic matter. Quartz (silicon dioxide, the main component of glass) is one of the most common minerals found in nearly all rocks, and it is very resistant to abrasion, so it is a dominant component of lithogenous sediments, including sand.[1]
Biogenous
Macroscopic sediments contain large remains, such as skeletons, teeth, or shells of larger organisms. This type of sediment is fairly rare over most of the ocean, as large organisms do not die in enough of a concentrated abundance to allow these remains to accumulate. One exception is around coral reefs; here there is a great abundance of organisms that leave behind their remains, in particular the fragments of the stony skeletons of corals that make up a large percentage of tropical sand.[1]
Microscopic sediment consists of the hard parts of microscopic organisms, particularly their shells, or tests. Although very small, these organisms are highly abundant and as they die by the billions every day their tests sink to the bottom to create biogenous sediments. Sediments composed of microscopic tests are far more abundant than sediments from macroscopic particles, and because of their small size they create fine-grained, mushy sediment layers. If the sediment layer consists of at least 30% microscopic biogenous material, it is classified as a biogenous ooze. The remainder of the sediment is often made up of clay.[1]
through sediment analysis
Biogenous sediments can allow the reconstruction of past climate history from oxygen isotope ratios. Oxygen atoms exist in three forms, or isotopes, in ocean water: O16, O17 and O18 (the number refers to the atomic masses of the isotopes). O16 is the most common form, followed by O18 (O17 is rare). O16 is lighter than O18, so it evaporates more easily, leading to water vapor that has a higher proportion of O16. During periods of cooler climate, water vapor condenses into rain and snow, which forms glacial ice that has a high proportion of O16. The remaining seawater therefore has a relatively higher proportion of O18. Marine organisms which incorporate dissolved oxygen into their shells as calcium carbonate will have shells with a higher proportion of O18 isotope. This means the ratio of O16:O18 in shells is low during periods of colder climate. When climate warms, glacial ice melts releasing O16 from the ice and returning it to the oceans, increasing the O16:O18 ratio in the water. When organisms incorporate oxygen into their shells, the shells will contain a higher O16:O18 ratio. Scientists can therefore examine biogenous sediments, calculate the O16:O18 ratios for samples of known ages, and from those ratios, infer the climate conditions under which those shells were formed. The same types of measurements can also be taken from ice cores; a decrease of 1 ppm O18 between ice samples represents a decrease in temperature of 1.5°C.[1]
The primary sources of microscopic biogenous sediments are unicellular algaes and protozoans (single-celled amoeba-like creatures) that secrete tests of either
Diatoms are particularly important members of the phytoplankton, functioning as small, drifting algal photosynthesizers. A diatom consists of a single algal cell surrounded by an elaborate silica shell that it secretes for itself. Diatoms come in a range of shapes, from elongated, pennate forms, to round, or centric shapes that often have two halves, like a Petri dish. In areas where diatoms are abundant, the underlying sediment is rich in silica diatom tests, and is called diatomaceous earth.[1]
Radiolarians are planktonic protozoans (making them part of the zooplankton), that like diatoms, secrete a silica test. The test surrounds the cell and can include an array of small openings through which the radiolarian can extend an amoeba-like "arm" or pseudopod. Radiolarian tests often display a number of rays protruding from their shells which aid in buoyancy. Oozes that are dominated by diatom or radiolarian tests are called siliceous oozes.[1]
Like the siliceous sediments, the calcium carbonate, or calcareous sediments are also produced from the tests of microscopic algae and protozoans; in this case the coccolithophores and foraminiferans. Coccolithophores are single-celled planktonic algae about 100 times smaller than diatoms. Their tests are composed of a number of interlocking CaCO3 plates (coccoliths) that form a sphere surrounding the cell. When coccolithophores die the individual plates sink out and form an ooze. Over time, the coccolithophore ooze lithifies to becomes chalk. The White Cliffs of Dover in England are composed of coccolithophore-rich ooze that turned into chalk deposits.[1]
Foraminiferans (also referred to as forams) are protozoans whose tests are often chambered, similar to the shells of snails. As the organism grows, is secretes new, larger chambers in which to reside. Most foraminiferans are benthic, living on or in the sediment, but there are some planktonic species living higher in the water column. When coccolithophores and foraminiferans die, they form
Older calcareous sediment layers contain the remains of another type of organism, the discoasters; single-celled algae related to the coccolithophores that also produced calcium carbonate tests. Discoaster tests were star-shaped, and reached sizes of 5-40 µm across. Discoasters went extinct approximately 2 million years ago, but their tests remain in deep, tropical sediments that predate their extinction.[1]
Because of their small size, these tests sink very slowly; a single microscopic test may take about 10–50 years to sink to the bottom! Given that slow descent, a current of only 1 cm/sec could carry the test as much as 15,000 km away from its point of origin before it reaches the bottom. Despite this, the sediments in a particular location are well-matched to the types of organisms and degree of productivity that occurs in the water overhead. This means the sediment particles must be sinking to the bottom at a much faster rate, so they accumulate below their point of origin before the currents can disperse them. Most of the tests do not sink as individual particles; about 99% of them are first consumed by some other organism, and are then aggregated and expelled as large
Hydrogenous
Seawater contains many different dissolved substances. Occasionally chemical reactions occur that cause these substances to precipitate out as solid particles, which then accumulate as hydrogenous sediment. These reactions are usually triggered by a change in conditions, such as a change in temperature, pressure, or pH, which reduces the amount of a substance that can remain in a dissolved state. There is not a lot of hydrogenous sediment in the ocean compared to lithogenous or biogenous sediments, but there are some interesting forms.[1]
In hydrothermal vents seawater percolates into the seafloor where it becomes superheated by magma before being expelled by the vent. This superheated water contains many dissolved substances, and when it encounters the cold seawater after leaving the vent, these particles precipitate out, mostly as metal sulfides. These particles make up the "smoke" that flows from a vent, and may eventually settle on the bottom as hydrogenous sediment.[1] Hydrothermal vents are distributed along the Earth's plate boundaries, although they may also be found at intra-plate locations such as hotspot volcanoes. Currently there are about 500 known active submarine hydrothermal vent fields, about half visually observed at the seafloor and the other half suspected from water column indicators and/or seafloor deposits.[4]
Manganese nodules are rounded lumps of manganese and other metals that form on the seafloor, generally ranging between 3–10 cm in diameter, although they may sometimes reach up to 30 cm. The nodules form in a manner similar to pearls; there is a central object around which concentric layers are slowly deposited, causing the nodule to grow over time. The composition of the nodules can vary somewhat depending on their location and the conditions of their formation, but they are usually dominated by manganese- and iron oxides. They may also contain smaller amounts of other metals such as copper, nickel and cobalt. The precipitation of manganese nodules is one of the slowest geological processes known; they grow on the order of a few millimetres per million years. For that reason, they only form in areas where there are low rates of lithogenous or biogenous sediment accumulation, because any other sediment deposition would quickly cover the nodules and prevent further nodule growth. Therefore, manganese nodules are usually limited to areas in the central ocean, far from significant lithogenous or biogenous inputs, where they can sometimes accumulate in large numbers on the seafloor (Figure 12.4.2 right). Because the nodules contain a number of commercially valuable metals, there has been significant interest in mining the nodules over the last several decades, although most of the efforts have thus far remained at the exploratory stage. A number of factors have prevented large-scale extraction of nodules, including the high costs of deep sea mining operations, political issues over mining rights, and environmental concerns surrounding the extraction of these non-renewable resources.[1]
Oolites are small, rounded grains formed from concentric layers of precipitation of material around a suspended particle. They are usually composed of calcium carbonate, but they may also from phosphates and other materials. Accumulation of oolites results in oolitic sand, which is found in its greatest abundance in the Bahamas.[1]
Cosmogenous
Cosmogenous sediment is derived from extraterrestrial sources, and comes in two primary forms; microscopic spherules and larger meteor debris.
Cosmogenous sediment is fairly rare in the ocean and it does not usually accumulate in large deposits. However, it is constantly being added to through space dust that continuously rains down on Earth. About 90% of incoming cosmogenous debris is vaporized as it enters the atmosphere, but it is estimated that 5 to 300 tons of space dust land on the Earth's surface each day.[1]
Composition
Siliceous ooze
mineral forms |
protist involved |
name of skeleton | typical size | ||||
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diatom | frustule | 0.002 to 0.2 mm [8] | diatom microfossil from 40 million years ago | ||||
radiolarian
|
test or shell | 0.1 to 0.2 mm | elaborate silica shell of a radiolarian |
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diatoms
(click 3X to fully magnify)
Calcareous ooze
The term calcareous can be applied to a fossil, sediment, or sedimentary rock which is formed from, or contains a high proportion of,
Calcareous ooze
| |||||||
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mineral forms |
protist involved |
name of skeleton | typical size | ||||
CaCO3 calcite aragonite limestone marble chalk |
foraminiferan
|
test or shell | under 1 mm | ||||
coccolithophore | coccoliths | under 0.1 mm [10] | Coccolithophores are the largest global source of biogenic calcium carbonate, and significantly contribute to the global white cliffs of Dover .
|
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Calcareous microfossils from marine sediment consisting mainly of star-shaped discoaster with a sprinkling of coccoliths
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Illustration of a Globigerina ooze
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Shells (tests), usually made of calcium carbonate, from a foraminiferal ooze on the deep ocean floor
Lithified sediments
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Opal can contain protist microfossils of diatoms, radiolarians, silicoflagellates and ebridians [13]
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Carbonate-silicate cycle
Distribution
Where and how sediments accumulate will depend on the amount of material coming from a source, the distance from the source, the amount of time that sediment has had to accumulate, how well the sediments are preserved, and the amounts of other types of sediments that are also being added to the system.[1]
Rates of sediment accumulation are relatively slow throughout most of the ocean, in many cases taking thousands of years for any significant deposits to form. Lithogenous sediment accumulates the fastest, on the order of one metre or more per thousand years for coarser particles. However, sedimentation rates near the mouths of large rivers with high discharge can be orders of magnitude higher.[1]
Biogenous oozes accumulate at a rate of about 1 cm per thousand years, while small clay particles are deposited in the deep ocean at around one millimetre per thousand years. As described above, manganese nodules have an incredibly slow rate of accumulation, gaining 0.001 millimetres per thousand years.[1]
Marine sediments are thickest near the continental margins where they can be over 10 km thick. This is because the crust near passive continental margins is often very old, allowing for a long period of accumulation, and because there is a large amount of terrigenous sediment input coming from the continents. Near mid-ocean ridge systems where new oceanic crust is being formed, sediments are thinner, as they have had less time to accumulate on the younger crust.[1]
As distance increases from a ridge spreading center the sediments get progressively thicker, increasing by approximately 100–200 m of sediment for every 1000 km distance from the ridge axis. With a seafloor spreading rate of about 20–40 km/million years, this represents a sediment accumulation rate of approximately 100–200 m every 25–50 million years.[1]
The diagram at the start of this article ↑ shows the distribution of the major types of sediment on the ocean floor. Cosmogenous sediments could potentially end up in any part of the ocean, but they accumulate in such small abundances that they are overwhelmed by other sediment types and thus are not dominant in any location. Similarly, hydrogenous sediments can have high concentrations in specific locations, but these regions are very small on a global scale. So cosmogenous and hydrogenous sediments can mostly be ignored in the discussion of global sediment patterns.[1]
Coarse lithogenous/terrigenous sediments are dominant near the continental margins as
Coarse lithogenous sediments are less common in the central ocean, as these areas are too far from the sources for these sediments to accumulate. Very small clay particles are the exception, and as described below, they can accumulate in areas that other lithogenous sediment will not reach.[1]
The distribution of biogenous sediments depends on their rates of production, dissolution, and dilution by other sediments. Coastal areas display very high primary production, so abundant biogenous deposits might be expected in these regions. However, sediment must be >30% biogenous to be considered a biogenous ooze, and even in productive coastal areas there is so much lithogenous input that it swamps the biogenous materials, and that 30% threshold is not reached. So coastal areas remain dominated by lithogenous sediment, and biogenous sediments will be more abundant in pelagic environments where there is little lithogenous input.[1]
In order for biogenous sediments to accumulate their rate of production must be greater than the rate at which the tests dissolve. Silica is undersaturated throughout the ocean and will dissolve in seawater, but it dissolves more readily in warmer water and lower pressures; that is, it dissolves faster near the surface than in deep water. Silica sediments will therefore only accumulate in cooler regions of high productivity where they accumulate faster than they dissolve. This includes upwelling regions near the equator and at high latitudes where there are abundant nutrients and cooler water.[1]
Oozes formed near the equatorial regions are usually dominated by radiolarians, while diatoms are more common in the polar oozes. Once the silica tests have settled on the bottom and are covered by subsequent layers, they are no longer subject to dissolution and the sediment will accumulate. Approximately 15% of the seafloor is covered by siliceous oozes.[1]
Biogenous calcium carbonate sediments also require production to exceed dissolution for sediments to accumulate, but the processes involved are a little different than for silica. Calcium carbonate dissolves more readily in more acidic water. Cold seawater contains more dissolved CO2 and is slightly more acidic than warmer water. So calcium carbonate tests are more likely to dissolve in colder, deeper, polar water than in warmer, tropical, surface water. At the poles the water is uniformly cold, so calcium carbonate readily dissolves at all depths, and carbonate sediments do not accumulate. In temperate and tropical regions calcium carbonate dissolves more readily as it sinks into deeper water.[1]
The depth at which calcium carbonate dissolves as fast as it accumulates is called the calcium carbonate compensation depth or
The CCD is deeper in the Atlantic than in the Pacific since the Pacific contains more CO2, making the water more acidic and calcium carbonate more soluble. This, along with the fact that the Pacific is deeper, means that the Atlantic contains more calcareous sediment than the Pacific. All told, about 48% of the seafloor is dominated by calcareous oozes.[1]
Much of the rest of the deep ocean floor (about 38%) is dominated by abyssal clays. This is not so much a result of an abundance of clay formation, but rather the lack of any other types of sediment input. The clay particles are mostly of terrestrial origin, but because they are so small they are easily dispersed by wind and currents, and can reach areas inaccessible to other sediment types. Clays dominate in the central North Pacific, for example. This area is too far from land for coarse lithogenous sediment to reach, it is not productive enough for biogenous tests to accumulate, and it is too deep for calcareous materials to reach the bottom before dissolving.[1]
Because clay particles accumulate so slowly, the clay-dominated deep ocean floor is often home to hydrogenous sediments like manganese nodules. If any other type of sediment was produced here it would accumulate much more quickly and would bury the nodules before they had a chance to grow.[1]
Coastal sediments
The sediment itself is often composed of limestone, which forms readily in shallow, warm calm waters. The shallow marine environments are not exclusively composed of siliciclastic or carbonaceous sediments. While they cannot always coexist, it is possible to have a shallow marine environment composed solely of carbonaceous sediment or one that is composed completely of siliciclastic sediment. Shallow water marine sediment is made up of larger grain sizes because smaller grains have been washed out to deeper water. Within sedimentary rocks composed of carbonaceous sediment, there may also be evaporite minerals.[16] The most common evaporite minerals found within modern and ancient deposits are gypsum, anhydrite, and halite; they can occur as crystalline layers, isolated crystals or clusters of crystals.[16]
In terms of geologic time, it is said that most Phanerozoic sedimentary rock was deposited in shallow marine environments as about 75% of the sedimentary carapace is made up of shallow marine sediments; it is then assumed that Precambrian sedimentary rocks were too, deposited in shallow marine waters, unless it is specifically identified otherwise.[17] This trend is seen in the North American and Caribbean region.[18] Also, as a result of supercontinent breakup and other shifting tectonic plate processes, shallow marine sediment displays large variations in terms of quantity in the geologic time.[18]
Bioturbation
Bioturbators are ecosystem engineers because they alter resource availability to other species through the physical changes they make to their environments.[22] This type of ecosystem change affects the evolution of cohabitating species and the environment,[22] which is evident in trace fossils left in marine and terrestrial sediments. Other bioturbation effects include altering the texture of sediments (diagenesis), bioirrigation, and displacement of microorganisms and non-living particles. Bioturbation is sometimes confused with the process of bioirrigation, however these processes differ in what they are mixing; bioirrigation refers to the mixing of water and solutes in sediments and is an effect of bioturbation[20]
Walruses and salmon are examples of large bioturbators.[23][24][25] Although the activities of these large macrofaunal bioturbators are more conspicuous, the dominant bioturbators are small invertebrates, such as polychaetes, ghost shrimp and mud shrimp.[20][26] The activities of these small invertebrates, which include burrowing and ingestion and defecation of sediment grains, contribute to mixing and the alteration of sediment structure.
Bioirrigation
Bioirrigation works as two different processes. These processes are known as
Pelagic sediments
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Sediment supply from terrigenous and biological sources
as well as its dispersion and settling through the water column [28]
Pelagic sediments, or pelagite, are fine-grained sediments that accumulate as the result of the settling of particles to the floor of the open ocean, far from land. These particles consist primarily of either the microscopic, calcareous or siliceous shells of phytoplankton or zooplankton; clay-size siliciclastic sediment; or some mixture of these. Trace amounts of meteoric dust and variable amounts of volcanic ash also occur within pelagic sediments. Based upon the composition of the ooze, there are three main types of pelagic sediments: siliceous oozes, calcareous oozes, and red clays.[29][30]
An extensive body of work on deep-water processes and sediments has been built over the past 150 years since the voyage of HMS Challenger (1872–1876), during which the first systematic study of seafloor sediments was made.[31][32] For many decades since that pioneering expedition, and through the first half of the twentieth century, the deep sea was considered entirely pelagic in nature.[28]
The composition of pelagic sediments is controlled by three main factors. The first factor is the distance from major landmasses, which affects their dilution by terrigenous, or land-derived, sediment. The second factor is water depth, which affects the preservation of both siliceous and calcareous biogenic particles as they settle to the ocean bottom. The final factor is ocean fertility, which controls the amount of biogenic particles produced in surface waters.[29][30]
Turbidites
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Continental margins can experience slope failures triggered by earthquakes or other geological disturbances. These can result in turbidity currents as turbid water dense with suspended sediment rushes down the slope. Chaotic motion within the sediment flow can sustain the turbidity current, and once it reaches the deep abyssal plain it can flow for hundreds of kilometres.[33]
Turbidites were first recognised in the 1950s [34] and the first facies model was developed by Bouma in 1962.[35] Since that time, turbidites have been one of the better known and most intensively studied deep-water sediment facies. They are now very well known from sediment cores recovered from modern deep-water systems, subsurface (hydrocarbon) boreholes and ancient outcrops now exposed on land. Each new study of a particular turbidite system reveals specific deposit characteristics and facies for that system. The most commonly observed facies have been variously synthesised into a range of facies schemes.[36][37][28]
Contourites
A contourite is a sedimentary deposit commonly formed on continental rise to lower slope settings, although they may occur anywhere that is below storm wave base. Countourites are produced by thermohaline-induced deepwater bottom currents and may be influenced by wind or tidal forces.[39][40] The geomorphology of contourite deposits is mainly influenced by the deepwater bottom-current velocity, sediment supply, and seafloor topography.[41]
Contourites were first identified in the early 1960s by
Hemipelagic
-
Hemipelagite facies models
Standard model showing simple cyclicity between clay-rich and biogenic-rich parts. Variations depend on component inputs.[28]
Ecology
Part of a series related to |
Benthic life |
---|
- Ancient Greekhupér 'over', live just above the sediment.
- Ancient Greekepí 'on top of', live on top of the sediments.
- oxygenatedtop layer.
Microbenthos
Marine
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Elphidium a widespread abundant genus of benthic forams
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Heterohelix, an extinct genus of benthic forams
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Darkfield photo of a gastrotrich, 0.06-3.0 mm long, a worm-like animal living between sediment particles
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Armoured Pliciloricus enigmaticus, about 0.2 mm long, live in spaces between marine gravel
Diatoms form a (disputed) phylum containing about 100,000 recognised species of mainly unicellular algae. Diatoms generate about 20 per cent of the oxygen produced on the planet each year,[52] take in over 6.7 billion metric tons of silicon each year from the waters in which they live,[53] and contribute nearly half of the organic material found in the oceans.
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Diatoms are one of the most common types of phytoplankton
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Their protective shells (frustles) are made of silicon
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They come in many shapes and sizes
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Like diatoms, radiolarians come in many shapes
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Also like diatoms, radiolarian shells are usually made of silicate
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Howeveracantharian radiolarians have shells made from strontium sulfatecrystals
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Cutaway schematic diagram of a spherical radiolarian shell
Like radiolarians,
Both foraminifera and diatoms have
The sudden extinction event which killed the dinosaurs 66 million years ago also rendered extinct three-quarters of all other animal and plant species. However, deep-sea benthic forams flourished in the aftermath. In 2020 it was reported that researchers have examined the chemical composition of thousands of samples of these benthic forams and used their findings to build the most detailed climate record of Earth ever.[59][60]
Some endoliths have extremely long lives. In 2013 researchers reported evidence of endoliths in the ocean floor, perhaps millions of years old, with a generation time of 10,000 years.[61] These are slowly metabolizing and not in a dormant state. Some Actinomycetota found in Siberia are estimated to be half a million years old.[62][63][64]
Sediment cores
The diagram on the right shows an example of a sediment core. The sample was retrieved from the
Carbon processing
Thinking about ocean carbon and carbon sequestration has shifted in recent years from a structurally-based chemical reactivity viewpoint toward a view that includes the role of the ecosystem in organic carbon degradation rates.
Evolutionary history
Animation of PangaeariftingThe surface of the Earth has continually reshaped itself over billions of years. Continents formed and broke apart, migrating across the surface and occasionally combining to form a supercontinent. The earliest-known supercontinent Rodinia assembled about one billion years ago, and then began to break apart about 700 million years ago. The continents later recombined to form Pannotia, 600 to 540 million years ago, then finally Pangaea, which broke apart 200 million years ago.
To begin with, the Earth was molten due to extreme volcanism and frequent collisions with other bodies. Eventually, the outer layer of the planet cooled to form a solid crust and water began accumulating in the atmosphere. The Moon formed soon afterwards, possibly as a result of the impact of a planetoid with the Earth. Outgassing and volcanic activity produced the primordial atmosphere. Condensing water vapor, augmented by ice delivered from comets, produced the oceans.[71][72][73]
By the start of the
The
During the
Marine habitats |
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Coastal habitats |
Ocean surface |
|
Open ocean |
Sea floor |
Patterns or traces of bioturbation are preserved in
Important
Research history
The first major study of deep-ocean sediments occurred between 1872 and 1876 with the
Earlier theories of
See also
- Bioturbation
- Depositional environment
- Cosmic dust
- Deep biosphere
- Great Calcite Belt
- Marine clay
- Microbially induced sedimentary structure
- Oolitic aragonite sand
- Organic-rich sedimentary rocks
- Paleolimnology § Paleoclimate proxies
- Redox gradient
- Seafloor depth versus age
- Sediment-water interface
- Sedimentary rock
- Sediment transport
References
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