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S330 Oceanography

Block 4: Marine Biogeochemical Cycles




Biogenic sediments
Terrigenous sediments


The steady-state ocean
Particulate organic matter
Particulate organic matter and the nutrient cycle
Particulate organic matter and the scavenging cycle
A classification of the elements in seawater
The two-box model and phosphate
Estimating residence times
Transfer of gases
Transfer of liquids and solids
Oxygen in seawater


The preservation of deep-sea siliceous remains
The preservation of deep-sea carbonates
Turbidity currents and other gravity flows


Sea-surface temperature
Ocean circulation
The carbonate system


Authigenic clay minerals
Manganese nodules
The diagenetic sequence
Gas hydrates

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Deep-sea sediments can be classified broadly as terrigenous (land-derived) and biogenic (formed as a result of biological activity), with minor volcanogenic and cosmic contributions. Pelagic sediments include all those sediments deposited from the water columnin the deep ocean basins beyond the influence of continental margin processes. Factors such as submarine topography and climatic patterns influence the type and abundance of sediments that accumulate in a particular region.

Pelagic biogenic sediments are composed mostly of the remains of calcareous (carbonate) and siliceaous planktonic organisms, principally coccolithophores, foraminiferans, pteropod, diatoms and radiolarians.The preservation of these depends on a number of factors such as water depth and chemistry, the shape of the skeletal remains, and the presence or absence of an organic membrane.

Wind-blown dusts consist mostly of clay minerals derived from weathering of terrestrial rocks. Coarse debris in deep-sea clays comprise clay minerals derived from a number of different sources. Four main types of clay minerals are recognized, each characteristic of particular weathering regimes; kaolinite, chlorite, illite and montmorillonite. Aeolian inputs are a major contribution to clays in many regions.

Clays predominate in the deepest parts of the ocean basins where they are not diluted by biogenic material. Calcareous sediments are largely confined to the shallowest regions of the open oceans, such as the mid-ocean ridges, while siliceous sediments predominate below regions of upwelling where there is little input from other material (ie terrigenous or calcareous material).

Accumulation rates of pelagic sediments are very slow, a few millimetres per thousand years. Sediments sccumulating along continental margins do so much more rapidly, usually >1cm per thousand years.


Most naturally occurring elements have been detected in solution in seawater. Variations in salinity do not affect the overall constancy of composition of seawater with respect to major constituents, most of which behave conservatively. Most minor and trace constituents (and a few major ones, notably carbon, and calcium to a small extent) are non-conservative, because their concentrations are affected by biological processes in the oceans.

The oceans are in a steady state. There is an overall balance between the rate of supply of dissolved constituents (including excess volatiles and cyclic salts) and their rate of removal from solution. Residence times range from several tens of millions of years to a few hundred years or less, but most residence times are long compred with the oceanic stirring time. While in the oceans, dissolved constituents participate repeatedly in (mainly) builogical cycles before being ultimately removed from solution in seawater into sediments and rocks at the sea-bed. Eventually, these sediments and rocks are accreted to continental margins or returned to the Earth's mantle by subduction at ocean trenches.

Much organic matter produced by primary production in surface waters is recycled there, but a proportion sinks out of the photic zone towards teh sea-bed. Particulate organic matter ranges from micrometres to centimetres in size. Larger praticles are mainly marine snow and faecal materials, while the smallest particles are mostly bacteria. In suface waters, the average molar ratio of carbon to the two principal nutrient elements in organic matter - nitrogen and phosphorus - is close to 106:16:1 (C:N:P); this is the basic Redfield ratio. The compostion of sinking particles changes with depth: skeletal material (carbonate and silica) dissolves only slowly, but soft tissue is consumed and decomposed by animals and bacteria, and the residue becomes more refractory with depth as nutrients are extracted - the proportion of C relative to N and P in particulate mattr increases. Information as to the composition and distribution of particulate matter has come largely from sediment traps.

Particulare organic carbon contributes only a small proportion of the organic carbon in seawater (c 0.05-0.1 mg C l -1 on average). Most organic carbon is in the form of dissolved organic compounds (c 0.5 - 1 mg C l -1) on average), but these are largely refractory and thus not available for microbial consumption.

Formation and decomposition of particulate organic matter is a major regulator of seawater composition, for minor and trace constituents in particular. Elements classified as recycled have concentration profiles resembling those for the major nutrients (nitrate, phosphate, silica). They are taken up by organisms during growth and released back into solution as the organisms are consumed and decomposed on sinking into deeper waters after death. Recycled elements can also be subdivided into biolimiting (concentrations near-zero in surface waters), and bio-intermediate (concentrations only somewhat reduced in surface waters). Which of these categories an element falls into depends greately on the amounts used in biological production relative to their total cocnentrations in seawater. Scavenged elements have profiles showing depletion at depth, the result of adsorption onto surfaces of particles (mainly bacteria) and scavenging from the water column by larger sinking particles. The solubility of elements with more than one oxidation state can vary according to whether conditions are oxidizing or reducing.

The two-box model enables first-approximation estimates to be made of the relative amounts of dissolved constituents that are removed into sediments and recycled in the water column. It can only be applied to elements in the recycled category. Only a minute fraction of the particulate material containing the biolimited constituents reaches sediments on the sea-bed; the remainder is recycled, mainly above the permanent thermocline.

In the deep oceans, there are marked lateral variations in concentrations of biologically active constituents. The pattern of thermohaline circulation results in an overall enrichment of nutrients in the deep waters of the North Pacific relative to the North Atlantic.

Transfer of gases, liquids and solids can take place across the air-sea interface. Gas exchange occurs by molecular diffusion and, as wind speed increases, vertical turbulence. Gas exchange is continuous, but at equlibrium there is no net flux in either direction. There is approximate equilibrium between atmosphere and ocean for the major gases. Minor gases produced by organisms in surface wters have a net flux from sea to air. Both gases and solids are also transferred across the air-sea interface in precipitation. In this case, the air-sea flux depends on the nature and intensity of precipitation and on the washout ratio; and the sea-air flux depends on the extent to which aerosols are produced.

The deep oceans are well supplied with oxygen by deep water masses formed at high latitudes. Dissolved oxygen concentrations decrease as water masses 'age' as they move away from their source regions and marine organisms use the oxygen in respiration/decomposition. The excess of abstraction over replenishment of oxygen reaches a maximum at about 1 km depth, and sub-oxic conditions can develop. High levels of biological production can cause some coastal waters to become sub-oxic or anoxic. In the absence of oxygen, bacteria use oxidizing agents such as sulphate to decompose organic matter.


Biogenic sediments are most abundant where productivity in surface waters is high and terrigenous sediments are scarce. The geographical separation of carbonate and siliceous sediments is related to the preservation potential of the planktonic organisms and the chemistry of the water column. Biological aggregation is important in the transport of planktonic skeletal remains to the sea-bed.

Seawater at all depths is undersaturated with respect to silica, and the preservation of siliceous sediments depends on the survival of siliceous skeletal debris as it descends the water column. The likelihood of preservation is greater once the remains have survived descent through surface waters. The chances of siliceous sediemtns accumulating are greatest where surface productivity is high, and where water depths are great, so that dilution by terrigenous or calcareous material is low. Siliceous sediments are most abundant at high latitudes in the Pacific Ocean, in equatorial regions of both the Pacific and Indian Oceans, and in coastal upwelling areas.

Carbon dioxide gas dissolves in seawater to form dissolved CO2 (commonly referred to as carbonic acid, H2CO3) and hydrogen carbonate (bicarbonate) and carbonate ions. Hydrogen carbonate and carbonate ions are also supplied to the oceans by rivers. Total dissolved inorganic carbon (ΣCO2) increases in deep water chiefly because of the decomposition of organic matter in the water column (liberating CO2, which goes into solution to form carbonic acid). The greater the concentration of ΣCO2, the more carbonic acid in the water so the more likely calcium carbonate is to dissolve.

The solubility of calcium carbonate increases with depth in the oceans. Surface waters are supersaturated with respect to calcium carbonate, and deep waters are undersaturated.

The depth at which dissolution of calcium carbonate in sediments is observed to commence is called the lysocline. It lies at or below the depth where the water is shown from measurements to be sautrated with respect to calcium carbonate (the saturation horizon). The carbonate compensation depth (CCD) is defined as the depth at which the carbonate content of the sediments in 20% or less. It tends to be depressed (deeper) beneath areas of high biological productivity in the open oceans. Both the lysocline and CCD are shallower for aragonite than for calcite.

Terrigenous sediments are the products of physical and chemical weathering at the land surface. They are transported to the continental shelves and redistributed by waves and currents. Fine sediment is resuspended at the shlf edge by waves and currents. It escapes down the continental slope in low-density suspension (lutite flows) and by cascading. Where the sediments at the shelf edge are unstable, slides, slumps or debris flows occur (along with turbidity currents), carrying large quantities of coarse and fine sediment down to the continental rise.

Turbidity currents are the most important means whereby terrigenous sesdiments are transported from teh continental slop out to the abyssal plains. They can travel at speeds of several tens of kilometres per hour.

The passage of most turbidity currents is marked by characteristic deposits known as turbidite sequences which build up submarine fans at the base of the continental slope. The terminal stages of turbidity currents are commonly dilute suspensions of silt and mud.


The sea-floor sediment record extrends back to ~200 Ma ago. Sediments are thinnest closest to spreading axes, where the ocean crust is young, and increase in thickness with distance (and depth) from the axis as the crust ages and more time has elapsed for sediments to accumulate.

Sediments recovered in deep-sea drill cores contain a wealth of information about past ocean environments. This is obtained from the nature of the sediments themselves, and from measurements of sediment properties (such as the isotopic composition of microfossil shells) which respond systematically to important environmental variables, such as past sea-surface temperature, which cannot be directly measured. Such sediment properties are known as 'proxies'.

The most widely used proxy for sea-surface temperature is the oxygen-isotope ratio (usually expressed as δ18O) of planktonic foraminiferans. Foraminiferans incorporate proportionally more 18O into their calcium carbonate shells at low temperatures, and more 16O at high temperatures. Planktonic foraminiferal δ18O is not the perfect proxy, however. Oxygen-isotope ratios are also affected by the 18O/16O ratio of the water in which the foraminiferans lived; this varies with ice volume, and hece sea-level. Other proxies that are useful for determining past sea-surface temperatures include planktonic foraminiferal Mg/Ca, and the alkenone unsaturation index (Uk37).

Proxies for the concentrations of the nutrient elements in the deep sea can be used to trace changes in the rate of formation of deep water, and the pattern of thermohaline circulation. Once such proxy is the Cd concentration in shells of benthic foraminiferans: Cd mimics phosphate in the deep ocean. Another proxy is the carbon-isotope ratio of shells of benthic foraminiferans. Photsynthesizing organisms take up 12C in preference to 13C, so surface waters are relatively depleted in 12C. This 12C is returned to seawater by respiration at depth. Therefore the older the water mass, the more 12C it contains, and the lower its δ13C value.

Reconsturcting past changes in the marine carbonate system is central to understanding the causes of climate change. Deep-water carbonate ion concentration can be estimated from past levels of the CCD, and by analysis of microfossil fragmentation.

Changes in biological production can be reconstructed from analyses of Cd in the shells of planktonic foraminiferans: concentrations of Cd (phosphate) in surface waters are low if photosynthetic activity is high, so the proportion of Cd incorporated into the shell decreases.

While sediment proxies have provided an abundance of information for reconstructing past climate, they may yeield conflicting information. There are a number of reasons for this; our understanding of the parameters that control proxy relationships is incomplete; the chemical signature of microfossils may be altered upon burial; and some proxies require better calibration. Palaeocoeanographers are working hard to resolve these conflicts, and to develop new proxies.


Deep-sea sediments are disturbed and mixed (bioturbated) by animals moving over and through them in search of food. Bottom currents can make bed forms such as rupple marks on the surface, and they can resuspend large amounts of sediment. In regions of the ocean where there are strong western boundary currents, abyssal storms lead to erosion of bottom sediments. In the lowermost few tens of metres, the water column is turbid with suspended sediments and is called the benthic boundary layer.

Authigenesis - the formation of new minerals at the sea-bed - includes the formation of clays (montmorillonite and phillipsite) and manganese nodules. The latter are mainly spherioidal structures growing in successive layers around a nucleus at rates of a few millimetres per million years (at least in the deep sea), and reaching average sizes of a few centimetres. They grow by precipitation both from overlying seawater and from pore waters in underlying sediment. Deep-sea nodules contain Co, Ni and Cu in combined concentrations of up to 3%, which make them commercially attractive.

Diagenesis encompasses reactions that occur between pore waters and solid phases below the sediment surface, after the sediments become buried. The distinction between authigenesis and diagenesis is not always clear-cut.

There are considerable differences in the concentration of redox-sensitive consitutents above and below the oxic-anoxic interface in marine sediments. Below the interface, different redox species tend to be utilized in succession: nitrate, manganese and iron oxides, then sulphate. Many of these redox rections are brought about by bacteria living within the sediments.

If the flux of organic carbon to the sediment is extremely high, methane (and other) hydrocarbons form. Under favourable conditions of low temperature and high pressure, large accumulations of gas hydrates may develop.