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Deepsea Sediments

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    Deepsea Sediments & PaleoceanographyIntroduction

    eep-sea sediments, those found at depths greater than about 500 m, cover roughly two-thirds of the Earth. Not

    surprisingly, there are many kinds of deep-sea sediments. Fortunately, for someone learning about them, the

    predominant deep sediment is carbonate ooze, which covers nearly half the ocean floor. Even more fortunate for the

    marine geology student, by understanding a few simple concepts about the processes of deep-sea sedimentation, one can

    predict with a high degree of accuracy the kind of sediment found in any part of the ocean .

    The basic principles to understand are

    source, means of transport, rate of supply,

    and potential for dissolution or change on

    the sea floor. The basic sources of the

    sediments found in the deep sea are

    erosion from land , eruption of volcanoes,

    production by pelagic organisms , and

    cosmic fallout. Means of transport, which applies mostly to sediments eroded from land, refers to whether the sediments

    were dispersed out over the oceans by wind, were transported to the deep sea by gravity flows, were conveyed far from

    shore by surface currents before settling out of suspension, or were carried and dropped by melting ice. Rates of supply

    for sediments eroded from land or erupted by volcanoes declines with distance from a source. Rates and types of

    production by pelagic organisms vary with nutrient supplies and temperature in the surface waters of the ocean. Potential

    for dissolution or change depends upon the chemistry of the water in the deep sea and in the deep-sea sediments

    themselves.

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    determined that stable oxygen isotopes fractionate at different temperatures during precipitation and evaporation of water.

    Emiliani, who is credited with founding "paleoceanography" as a field of research, recognized that the fractionation of

    isotopes of oxygen that are incorporated into the CaCO2 shells of marine organisms should record the oceanic temperature

    at which the shells formed. He proceeded to demolish the idea that the deep ocean environment has been constant through

    Earth history. Using 18O/16O isotope ratios in shells of benthic foraminifera, he showed that bottom water temperatures in

    the mid Cenozoic were several degrees warmer than at present.

    The Tectonic Revolution

    y the 1950's, the basic tools were available for marine geologists to begin the most important revolution in scientific

    thinking since Darwin's Theory of Evolution. Fortunately, research funding was also available, thanks in part to Cold

    War concerns about submarine warfare. The leadership and scientific vision of geoscientists Revelle of the Scripps

    Institution of Oceanography and Ewing of the Lamont Geological Observatory were instrumental in directing interest and

    resources to deep-sea geology. Ewing initiated and Heezen and Tharp developed and published the widely-used, detailed

    maps of the ocean floors . Soviet scientists were also actively involved in mapping ocean-floor features and sediment

    distributions.

    The tectonic revolution in the

    Earth sciences really began in

    1961 when Dietz of the U.S.

    Coast and Geodetic Survey

    proposed a theory of how the sea

    floor is created and destroyed,

    which he called "sea-floor

    spreading." A year later Hess of

    Princeton University proposed

    that plate formation and

    continental movement is driven

    by convection currents within the

    mantle. Hess postulated that

    ocean crust forms volcanically at

    ocean ridge crests, cools and

    subsides with distance from the

    ridge, and ultimately is dragged

    downward into oceanic trenches.

    The theory of plate tectonics

    developed quickly in subsequent

    years.

    A Brief History of Ocean

    Drilling

    he composition, distribution

    and age of ocean sediments

    could be studied without thecontext of the Theory of Plate

    Tectonics, but understanding and

    interpreting processes and

    patterns would be much more

    difficult. Also, without the

    stimulus to support or disprove

    this new theory, the resources

    that have been dedicated to ocean drilling over the past 35 years would not have been allocated. The contributions made

    by the scientists and administrators that developed and pursued the idea of drilling in the deep ocean and by the political

    leaders who made the financial resources available must also be recognized.

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    The first scientific drilling operations in the deep sea began in 1961 in 945 m water depth off southern California, drilling

    1,315 m into the sea floor. Immediately thereafter, a second site in 3,558 m of water, known as the Experimental Mohole,

    was drilled off Baha California. This hole penetrated 183 m of sediment and 13 m of basalt, failing to reach the Mohole

    but demonstrating the feasibility of recovering scientifically valuable cores at depths well beyond the reach of the

    Kullenberg piston corer.

    In 1964, four United States oceanographic institutions joined together as JOIDES, Joint Oceanographic Institutions for

    Deep Earth Sampling, proposing that the U. S. National Science Foundation (NSF) support drilling off Jacksonville,

    Florida. Six sites were continuously cored to sub-bottom depths of more than 1 km, revealing significant oceanographicchanges on the east Florida margin since the Late Cretaceous. Well preserved planktic and benthic microfossils from the

    cores were instrumental in developing the biostratigraphic zonation schemes used today.

    JOIDES then initiated the Deep Sea Drilling Project (DSDP), which originally proposed 18-months of ocean drilling in

    the Atlantic and Pacific Oceans. The NSF funded modification of a drilling vessel under construction; it was modified

    specifically for scientific ocean

    drilling, core recovery and

    analysis. The resulting Glomar

    Challenger spent 15 years

    drilling the ocean basins and

    providing geologic data to

    solidify the theory of platetectonics, to develop the

    discipline of

    paleoceanography, and togreatly advance scientific

    understanding of Earth history

    and processes.

    In the 1970's, other U.S. and

    international institutions joined

    JOIDES. In 1985, the Ocean

    Drilling Project (ODP)

    succeeded DSDP withdedication of a larger, more

    sophisticated drillship, the JOIDES Resolution . The ODP continues past its original 10-year mission. The scientific

    discoveries of DSDP and ODP have affected everything from oil and mineral exploration to predicting earthquakes and

    global-climate fluctuations. Yet those discoveries would not have been possible without such astonishing engineering

    feats as hole re-entry cones, advanced piston corers, and stabilization techniques that allow drilling in stormy Antarctic

    seas, which is further testimony to the interdisciplinary nature of the Earth sciences. Furthermore, these discoveries would

    not have been possible if the United States, Germany, France, Canada, Japan, the United Kingdom, and the European

    Science Foundation had not dedicated the monetary resources needed to undertake this level of scientific research.

    Terrigenous Sediments

    errigenous sediments are derived from land. On land, rocks are broken down by physical and chemical weathering

    processes. Physical weathering breaks rocks into pieces ranging from massive boulders to clay-sized flakes of rock

    flour. Chemical weathering alters the chemistry of the source material as rocks are converted to sediments. Some of the

    rock material is literally dissolved away, which is the source of dissolved ions in seawater. The types and degrees of

    weathering reflect the climate of the source region, also known as sedimentary provenance. For example, rocks on

    Antarctica are predominantly broken down by physical processes. In deep-sea sediments around Antarctica, the textures

    of the sediments, shapes of the grains, and chemical composition of the clay minerals all reflect physical weathering. In

    contrast, deep-sea sediments off the Congo River in Africa reflect intense chemical weathering.

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    Rivers are the major source of sediments supplied to the oceans. Muds (silt- and clay-sized sediments) are carried in

    suspension by moving water and begin to settle out soon after the river water meets the ocean, though finer clay particles

    remain in suspension for years, allowing them to be conveyed far out into the ocean before settling to the bottom. The

    dissolved load is also an important contributor to deep-sea sedimentation, for it contains PO4---, NO3

    -, and other nutrients

    needed for plant growth, as well as Ca++, HCO3-, and H4SiO4, from which pelagic organisms build their shells and

    skeletons.

    Most of the sediment particles transported by rivers are deposited relatively near their mouths. Thus, by examining a map

    of the world, one can predict with substantial accuracy where most river-borne sediments are found on continental shelves

    and margins and in the deep sea. An overview of plate tectonics allows one to further predict the distributions ofterrigenous sediments in the deep sea. Submarine canyons along the trailing margins of the North and South Atlantic and

    Northern Indian Ocean deliver great quantities of terrigenous sediments to deep sea fans and abyssal plains. On the other

    hand, the deep basins and deep-sea trenches that border much of the Pacific Ocean capture most terrigenous sediments

    before they reach the deep sea.

    Gravity-Driven Sediment Transport

    Marine transport of most terrigenous

    sediment to the deep sea is by a variety of

    gravity-driven forms of movement

    including sliding, slumping, and sediment-

    gravity flows. All are produced by gravity-

    induced slope instability, usually resulting

    from the accumulation of large volumes of

    sediments in deltas or on continental

    margins. Movement can be triggered by an

    earthquake, hurricane, or simply by over-

    accumulation upslope. Gravity-driven

    movement is a key factor in shaping

    continental margins, for such flows both

    transport and erode. Slides are movements

    of large blocks of material along well-

    defined slippage planes. Sediments within a

    slide are often transported downslope with

    relatively little internal deformation. Slumps

    are also downslope movement of relatively

    large sediment parcels that move along

    discrete shear planes. Strata within a slump

    are usually deformed and normally dip back

    towards the slope. Large-scale slumping is

    most common at the transition from the

    gentle, upper continental slope and the

    steep, lower continental slope. Both slumps

    and slides can trigger sediment gravity flows.

    Sediment gravity flows occur when sediment is transported under the influence of gravity and sediment motion moves the

    accompanying interstitial fluid. Sediments are transported by a variety of mechanisms including suspension, saltation,

    traction, upward granular flow, direct interaction between grains, and the support of grains by a cohesive fluid. There are

    four main types of sediment gravity flows, in increasing order of importance:

    Grain flows occur when the sediment is supported and moved by direct grain to grain interactions. Examplesinclude downslope sand movement in submarine canyons that result in well-sorted sands or gravels deposited in

    channels of submarine fans.

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    Fluidized sediment flows are liquified, cohesionless particle movement in which the sediment is supported byupward flow of fluid escaping from between the grains as the grains settle by gravity. Such flow typically occurs

    in loosely packed sand which can move downslope as a traction carpet.

    Debris flows are downslope movements of mixtures of coarse and fine debris and water in which larger grains aresupported by a mixture of interstitial fluid and fine sediment. Deposits are typically massive and very poorly

    sorted. Sediments can be transported for tens, even hundreds of kilometers by debris flows.

    Turbidity currents are powerful, short-lived, gravity-driven currents consisting of dilute mixtures of sediment andwater having a density greater than the surrounding water.

    The sediments are supported mainly by the upward

    component of fluid turbulence. Turbidity currents are

    the major mechanism of transport of shallow-water

    sediments to deep abyssal plains.

    The incredible speed and power of turbidity currents

    was revealed by submarine cable breaks following an

    earthquake at Grand Banks, off Nova Scotia, Canada,

    on November 19, 1929. The quake triggered a

    turbidity current which progressively broke several

    telegraph cables over a 13-hour period, as the current

    traveled down the continental slope and continental

    rise, and out across the abyssal plain to more than 720

    km from its source. On the continental slope, velocity

    of the turbidity current exceeded 40 km/hr. After the

    cable break, a turbidite layer up to 1 m thick covered

    an area of at least 100,000 km2.

    Turbidites, which are the distinctive sediment deposits

    left by turbidity currents, are characterized by graded

    bedding, moderate sorting and well-developed primary sedimentary structures, as first described by Bouma. Pelagic

    sediment layers typically lie between individual turbidites. However, because the coarser sands settle out first while the

    finest muds travel farthest, the texture, sedimentary structures, and thickness of an individual turbidite changes from near

    the source to its periphery. Proximal turbidites resemble debris flows in that they are massive, with poorly developed

    sedimentary structures, weak grading, and little interbedded pelagic sediment or terrigenous mud, because the erosiveforce of the proximal turbidity flow removed previously deposited finer sediments. Classical turbidites, showing complete

    Bouma sequences , are typically intermediate in distance from the source. Distal turbidites, which are most distant from

    the source area, consist of thin; fine-grained layers that often exhibit well developed cross-lamination.

    Submarine canyons are the major conduits for movement of terrigenous sediments from river deltas and continental

    shelves down the continental margin to the deep sea. Submarine canyons themselves have been cut and sculpted by the

    erosive power of submarine gravity flows. During glacial advances when sea level was as much as 100 m or more lower,

    rivers delivered more sediment directly to the continental margins, so submarine canyons undoubtedly transported more

    sediment and eroded more rapidly.

    Grain flows are probably the most common mechanism of downslope transport in submarine canyons and result in

    massive, relatively well-sorted channel deposits in the deep-sea fans at the mouths of these canyons. Turbidity currents aremore sporadic events, but they carry much larger volumes of sediments and spread them far beyond the submarine fans

    onto the abyssal plains.

    Major river deltas on continental margins typically merge downslope into massive abyssal cones, where sedimentation

    rates can be meters to 10's of meters per 1000 years, depending upon sea level and denudation rates in the source region.

    The Atlantic has seven major abyssal cones off the St. Lawrence, Hudson, Mississippi, Amazon, Orange, Congo, and

    Niger Rivers. The largest cones in the world have been built by the Amazon, Ganges-Bramaputra and Mississippi Rivers.

    The most massive of these is the Bengal Cone, which is 3000 km long, up to 1000 km wide and up to 12 km in thickness.

    The Bengal Cone is produced by redistribution of sediment from the Ganges and Bramaputra Rivers, whose source waters

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    are in the Himalayas. The present rates of sediment influx into the Bay of Bengal indicate denudation rates in the

    Himalayas of up to 70 cm per 1000 years.

    Abyssal cones generally grade seaward into extensive abyssal plains, which are formed by accumulations of turbidites up

    to 1 km thick. The vast abyssal plains of the North and South Atlantic Oceans, the Aleutian Abyssal Plain in the northeast

    Pacific, and others are built of layer upon layer of turbidites, often interbedded with pelagic sediments. Sedimentation

    events are sporadic, but averaged over time; accumulation rates on abyssal plains may be 10's of cm to more than a meter

    per 1000 years. Interestingly, abyssal plains are not extensive in the northern Indian Ocean, despite voluminous sediment

    supplies, because of topographic restriction.

    Marine Clays

    he terrigenous sediments most likely to reach the deep sea are the clays, which arrive at the ocean margins in

    suspension, either in the air over the oceans or in surface waters, and may be transported by wind and ocean currents

    thousands of kilometers from their terrestrial source. In modern oceans, the less than 2 fraction clays make up 50-70%

    of the total oceanic sediment. In the open ocean, particles less than 0.5 m may stay in suspension for a hundred years or

    more before settling to the bottom. The settling process is accelerated by flocculation of clay aggregates and by

    incorporation into fecal pellets by pelagic organisms.

    Sediments that drape upper and middle continental slopes around the world are known as hemipelagic sediments. They

    grade from predominantly terrigenous muds into biogenic oozes. Even where biogenic constituents predominate,

    hemipelagic sediments typically have a dark color, which is imparted by the terrigenous component. The composition of

    the terrigenous muds reflects weathering intensity in the sedimentary provenance. The terrigenous muds, which were

    delivered to the ocean by rivers or by direct runoff from land, remained in suspension and were carried out to the

    continental margin by surface currents of by sediment-gravity flows.

    Accumulation rates of hemipelagic sediments can be quite high, up to 10-30 cm/1000 years. Two factors account for these

    rates, proximity to terrigenous sediment sources and proximity to terrestrial nutrient sources. Nutrients stimulate

    biological productivity, including either carbonate or siliceous sediment production.

    Clay minerals are aluminum silicates of varying complexities and stabilities. They occur as platy, lath-shaped or needle-

    like crystals, usually less than 4 in diameter. Their most striking property is cohesion, the tendency for constituent

    particles to stick together. Freshly deposited clay sediments contain much water and resemble cream or are jelly-like.

    Under pressure, clay sediments loose water and behave plastically, flowing under moderate stresses. Under very high

    pressure, clay sediments become sedimentary rocks such as shales that contain negligible water and are impermeable to

    fluids.

    Mineralogies of clays often reflect their origin to a substantial degree. There are four major classes of clay minerals in

    marine sediments; three reflect the relative degree of chemical weathering in the source region, while the fourth indicates

    volcanic origin.

    Clay-sized particles that have been primarily mechanically broken down and transported by ice, wind or very cold water

    have their cation suites relatively intact, including quite reactive cations such as Fe++. The most common of these unstable

    clay minerals is chlorite, which is found in high concentrations only at high latitudes where weathering processes are

    predominantly physical. Only 13% of the clay minerals in the oceans are chlorite.

    Illite is the most common clay mineral, often composing more than 50 percent of the clay-mineral suite in the deep sea.

    Illites are indicative of mechanical rather than chemical weathering, but are more stable than mica minerals. Illites are

    characteristic of weathering in temperate climates or in high altitudes in the tropics, and typically reach the ocean via

    rivers and wind transport.

    Kaolinites are recrystallization products of intense chemical weathering, and therefore are mostly found in low latitudes.

    Kaolinite is common throughout the equatorial Atlantic, but less so in the Pacific for lack of source. Maximum

    concentrations of kaolinite in deep-sea sediments are found off equatorial West Africa. High concentrations in the eastern

    Indian Ocean result from wind weathering of extensive "fossil" kaolinite-rich laterites in arid western Australia. These

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    laterites formed under wetter paleoclimatic conditions. Like chlorites, kaolinites make up only about 13% of the clay

    minerals in the deep sea.

    The fourth major group of clay minerals are the montmorillonites or smectites, which are chemical alteration products of

    volcanic material. Smectites are most common in areas where sedimentation rates are low and volcanic sources are

    nearby. Source material can be either windblown volcanic ash or volcanic glass on the sea floor. Smectites are most

    common in the South Pacific where they make up about 50 percent of the clay-mineral suite.

    Clays are present in virtually all marine sediments, though their proportions may be minor. In open ocean regions, remote

    from terrigenous sources, accumulation rates of deep-sea clays are on the order of a few mm per 1000 years. In pelagic

    sediments where clay minerals are the dominant constituent, sediments are typically bright red to chocolate brown in color

    and are known as red or brown clays. The color results from coatings of iron oxide on the sediment particles. The red

    clays were first described and mapped during the Challenger expedition (1872-1876). Accessory constituents include silt-

    or clay-sized grains of quartz, feldspar and pyroxene minerals, meteoric and volcanic dust, fish bones and teeth, whale ear

    bones, and manganese micro-nodules.

    Windblown Sediments

    he fine-grained sediments that reach the deep sea in regions remote from direct terrigenous sources are predominantly

    windblown. These include volcanic ash, terrigenous silts and clays, and some biogenic material such as freshwater

    diatoms, spores and pollen. Particles from each of these sources can tell something about the provenance from which they

    came. The chemical composition and particle size of the volcanic ash tells something of source, intensity and time of the

    eruption.

    Changes in the size distribution of quartz grains that reach the deep sea can reflect changes in intensity of high-altitude

    winds that transported the eolian dust. Composition of the clay minerals, as well as the types of biogenic material reflect

    climatic conditions of the source region. Biotic constituents may also indicate relative age. Deep-sea sediments in both the

    North Atlantic and the North Pacific contain substantial proportions of windblown sediments; clay minerals in both

    regions are predominantly illites. Accumulation rates of windblown sediments in the deep sea are typically up to a few

    mm/1000 years.

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    Glacial Marine Sediments

    or 20 million years, the Earth has hosted permanent ice sheets on Antarctica. Over the past two million years, ice

    sheets have been common in both polar regions. At the present time, there are immense continental glaciers on

    Antarctica and a smaller one on Greenland. Continental glaciers merge into ice shelves that generate icebergs laden with

    sediment . On a smaller scale, high latitude, higher elevation areas have montaine glaciers, which, if they reach the

    coastline, also calve off icebergs laden with ice-borne sediments. Even the lowlands of the Arctic tundra yield ice-rafted

    sediments via river pack ice.

    Sediment is scoured from

    land by the mechanical

    action of ice; 1-2% of the

    volume of this ice is typically

    sediment. The composition

    of the rock material is

    relatively unaltered as it is

    transported by ice and

    ultimately dropped as the ice

    melts. Thus, "drop stones"

    indicate both source and

    distance transported. Around

    Antarctica, most icebergs

    form at the inner margins of

    the Ross and Weddell Seas,

    and are carried into the

    Circumpolar Current system.

    North of the Antarctic

    Convergence, where water

    temperatures warm above 0o

    C, icebergs melt, and so ice-rafted sediments seldom reach beyond 40o S. In the North Atlantic, the iceberg limit is

    roughly the boundary between very cold polar waters and temperate waters. The extent of ice rafting was much greater

    during glacial advances, particularly in the North Atlantic.

    Glacial marine sediments include coarse, poorly-sorted

    debris and a silt fraction composed of rock flour; they

    typically contain little or no carbonate or biogenic

    material. Around Antarctica, there is a zonal

    distribution of sediment facies. Along the inner

    continental shelf, deposits are subglacial till, gravels,

    and sands, with some biogenic material. The outer

    continental shelf deposits are similar, but more

    characterized by sands and silts that grade into the

    pelagic clays of the abyssal regions. These clays

    contain occasional ice-rafted detritus. The pelagic

    clays grade northward into siliceous biogenic oozes.

    Glacial-marine sedimentation rates are low around

    Antarctica, in part because the climate is so cold and

    dry that the dry-base glaciers carry minimal sediment

    loads. In addition, the very cold, slowly accumulating

    and slowly moving permanent ice cover on the

    Antarctic continent seems to protect the continent

    from erosion more than it erodes.

    Glacial-marine sedimentation rates vary widely,

    depending upon climate in the source region. The North Atlantic Ocean, south of Iceland, receives about 60 percent of

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    global ice-rafted deposition. Higher snowfall and warmer, faster-moving glaciers on Greenland result in sediment delivery

    rates nearly 30 times faster than those of Antarctic glaciers. For similar reasons, ice-rafted sedimentation in the Norwegian

    Sea is volumetrically comparable to that of the circum-Antarctic, despite the huge difference in source areas. The North

    Pacific and the Arctic Ocean together receive roughly similar volumes of glacial marine sediments as the Norwegian Sea

    and the Antarctic region individually. Arctic glacial

    sediments tend to be silts and clays, reflecting

    eroded permafrost soils that are carried in river pack

    ice into the Arctic Ocean.

    Biogenic Sediments

    iogenic sediments, which are defined as

    containing at least 30% skeletal remains of

    marine organisms, cover approximately 62% of the

    deep ocean floor. Clay minerals make up most of the

    non-biogenic constituents of these sediments. While

    a vast array of plants and animals contribute to the

    organic matter that accumulates in marine

    sediments, a relatively limited group of organisms

    contribute significantly to the production of biogenic

    deep-sea sediments, which are either calcareous orsiliceous oozes.

    Distributions and accumulation rates of biogenic oozes in oceanic sediments depend on three major factors:

    rates of production of biogenic particles in the surface waters, dissolution rates of those particles in the water column and after they reach the bottom, and rates of dilution by terrigenous sediments.

    The abundances and distributions of the organisms that produce biogenic sediments depend upon such environmental

    factors as nutrient supplies and temperature in the oceanic waters in which the organisms live. Dissolution rates are

    dependent upon the chemistry of the deep ocean waters through which the skeletal remains settle and of the bottom and

    interstitial waters in contact with the remains as they accumulate and are buried. The chemistry of deep-sea waters, is, inturn, influenced by the rate of supply of both skeletal and organic remains of organisms from surface waters. It is also

    heavily dependent upon the rates of deep ocean circulation and the length of time that the bottom water has been

    accumulating CO2 and other byproducts of biotic activities.

    Carbonate Oozes

    ost carbonate or calcareous oozes are produced by

    the two different groups of organisms. The major

    constituents of nanofossil or coccolith ooze are tiny (less

    than 10 microns) calcareous plates produced by

    phytoplankton of the marine algal group, the

    Coccolithophoridae or by an extinct group called

    discoasters. Foraminiferal ooze is dominated by the tests

    (shells) of planktic protists belonging to the

    Foraminiferida. Most foraminiferal tests are sand-sized

    (>61 mm in diameter), so many foraminiferal oozes are

    bimodal in particle-size distribution, because they are made up of sand-sized foraminiferal tests and mud-sized coccolith

    plates.

    Discoasters, coccoliths and foraminiferal tests are all made of the mineral calcite. Pteropod ooze is produced by the

    accumulation of shells of pteropods and heteropods, which are small planktic mollusks. As these shells are composed of

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    the mineral aragonite, pteropod oozes are more easily dissolved, so are restricted to relatively shallow depths (less than

    3,000 m) in tropical areas.

    Carbonate oozes are the most widespread shell deposits on

    earth. Nearly half the pelagic sediment in the world's

    oceans is carbonate ooze . Furthermore, foraminifera and

    coccolithophorids have been major producers of pelagic

    sediment for the past 200 million years. As a result, these

    are arguably among the most important and scientifically

    useful organisms on Earth. Because their larger size makesthem easier to identify and work with, this is particularly

    true for the foraminifera. Their fossils provide the single

    most important record of Earth history over the past 200

    million years. That history is recorded not only by the

    evolution of species and higher taxa through that time, but

    is also preserved in the chemistry of the fossils themselves

    The field of Paleoceanography owns much of its existence

    to biostratigraphy, isotope stratigraphy and paleoenvironmental analyses that utilize fossil foraminifera.

    The distributions and abundances of living planktic foraminifera and coccolithophorids in the upper few hundred meters

    of the ocean depends in large part on nutrient supply and temperature. Coccolithophorids, because they are marine algae,

    require sunlight and inorganic nutrients (fixed N, P, and trace nutrients) for growth. However, most coccolithophorid

    species grow well with very limited supplies of nutrients and do not compete effectively with diatoms and dinoflagellates

    when nutrients are plentiful. Furthermore, both high nutrient supplies and cold temperatures inhibit calcium carbonate

    production to some degree. For these reasons, diversities (number of different kinds) of coccolithophorids are high and

    production rates of coccoliths are moderate even in the most nutrient-poor regions of the subtropical oceans, the

    subtropical gyres. Production of coccoliths is higher in equatorial upwelling zones and often along continental marginsand in temperate latitudes where nutrient supplies are higher, though diversities decline. In very high nutrient areas, such

    as upwelling zones in the eastern tropical oceans (i.e., meridional upwelling), polar divergences and near river mouths,

    production of coccoliths is minimal.

    Even though planktic foraminifera are protozoans rather than algae, their distributions, diversities, and carbonate

    productivity are quite similar to those of coccolithophorids. Many planktic foraminifera, especially the spinose species

    that live in the upper 100 m of temperate to tropical oceans host dinoflagellate symbionts which aid the foraminifera by

    providing energy and enhancing calcification. Having algal symbionts is highly advantageous in oceanic waters where

    inorganic nutrients and food are scarce, so a diverse assemblage of planktic foraminifera thrives along with the

    coccolithophorids in the nutrient-poor subtropical gyres. Greater abundances of fewer species thrive in equatorial

    upwelling zones and along continental margins, so rates of carbonate shell production are higher. And similar to

    coccolithophorids, few planktic foraminifera live in very high nutrient areas, such as upwelling zones in the eastern

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    tropical oceans, polar divergences and near river mouths, so production of carbonate sediments is minimal in these areas.

    Finally, planktic foraminifera require deep oceanic waters to complete their life cycles, which they cannot do in neretic

    waters over continental shelves.

    Cool temperatures work together with higher nutrient supplies to reduce diversities of coccolithophorids and planktic

    foraminifera, and ultimately to shift the ecological community to organisms that do not produce carbonate sediments. A

    10o C drop in temperature is physiologically similar to doubling nutrient supply, which is why the pelagic community in

    an equatorial upwelling zone resembles that of a temperate oceanic region, while the pelagic community of an intensive

    meridional upwelling zone resembles subpolar to polar communities.

    If surface production was the only factor controlling accumulation rates of carbonate oozes, deep-sea sediment patterns

    would be quite simple. Carbonate oozes would cover the seafloor everywhere except

    beneath intensive meridional upwelling zones, :beneath polar seas, and where they are overwhelmed by terrigenous sedimentation.

    Rates of accumulation would be on the order of 3-5 cm/1000 years in the open ocean and 10-20 cm/year beneath

    equatorial upwelling zones and along most continental margins.

    Dissolution

    ver much of the ocean floor, carbonate accumulation rates are controlled more by dissolution in bottom waters than

    by production in surface waters. Dissolution of calcium carbonate in seawater is influenced by three major factors:

    temperature, pressure and partial pressure of carbon dioxide (CO2). The easiest way to understand calcium carbonate

    (CaCO3) dissolution is to recognize that it is controlled, in large part, by the solubility of CO2:

    CaCO3 + H20 + CO2 Ca++ + 2HCO3

    -

    The more CO2 that can be held in solution, the more CaCO3 that will dissolve. Since more CO2 can be held in solution at

    higher pressures and cooler temperatures, CaCO3 is more soluble in the deep ocean than in surface waters. Finally, as CO2is added to the water, more CaCO3 can dissolve. The result is that, as more CO2 is added to deep ocean water by the

    respiration of organisms, the more corrosive the bottom water becomes to calcareous shells.

    The rain of organic matter from surface waters through time increases the partial pressure of CO2 in bottom water, so the

    longer the bottom water has been out of contact with the surface, the higher its partial pressure of CO2. Beneath high-

    nutrient surface waters, primary production exceeds what is utilized in the surface mixed layer. Excess organic matter

    falling through the water column accumulates on the bottom, where organisms feed upon it and oxidize it to CO2.

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    The depth at which surface production of CaCO3 equals dissolution is called the calcium carbonate compensation depth

    (CCD). Above this depth, carbonate oozes can accumulate, below the CCD only terrigenous sediments, oceanic clays, or

    siliceous oozes can accumulate. The calcium carbonate compensation depth beneath the temperate and tropical Atlantic is

    approximately 5,000 m deep, while in the Pacific, it is shallower, about 4,200-4,500 m, except beneath the equatorial

    upwelling zone, where the CCD is about 5,000 m. The CCD in the Indian Ocean is intermediate between the Atlantic and

    the Pacific. The CCD is relatively shallow in high latitudes.

    Surface waters of the ocean tend to be saturated with respect to CaCO3; low latitude surface waters are usually

    supersaturated. At shallow to intermediate seafloor depths (less than 3000 m), foraminiferal tests and coccolith plates tend

    to be well preserved in bottom sediments. However, at depths approaching the CCD, preservation declines as smaller andmore fragile foraminiferal tests show signs of dissolution. The boundary zone between well preserved and poorly

    preserved foraminiferal assemblages is known as the lysocline.

    The preservation potential of the various kinds of

    carbonate shells and skeletons differs. Pteropod shells

    are aragonite, a less stable form of CaCO3. Pteropod

    shells dissolve at depths greater than 3,000 m in the

    Atlantic Ocean and below a few hundred meters in the

    Pacific. Calcitic planktic foraminiferal tests, especially

    small tests of juvenile spinose foraminifera, dissolve

    more readily than coccoliths, which are also made of

    calcite. Pelagic sediments from relatively shallowdepths in low latitudes are often dominated by

    pteropods shells, at intermediate depths by foraminiferal

    tests, below the lysocline and above the CCD bycoccoliths, and below the CCD by red clays.

    Regional changes in the depths of the lysocline and

    CCD result, in part, from changes in CO2 content of

    bottom waters as they "age". In modern oceans, deep

    ocean circulation is driven by formation of bottom

    waters during the freezing of sea ice. Seawater, due to

    its salt content, can cool below -1o C before ice begins

    to form. When sea ice forms, the salt is excluded and isleft behind in the seawater. Water in the vicinity of the

    freezing sea ice becomes more saline and therefore

    more dense. As a result, large-scale sea ice formation

    creates very dense water masses that sink to the bottom

    of the ocean to form deep bottom water. During the

    Antarctic winter, the freezing of sea ice in the Weddell

    Sea produces Antarctic Bottom Water (AABW), which

    sinks to the sea bottom and spreads northward into the

    South Atlantic. During the Arctic winter, sea ice

    formation in the Norwegian and Greenland Seas

    produce North Atlantic Deep Water (NADW), which sinks to the bottom of the North Atlantic and flows southward.

    AABW is slightly more dense than NADW, so when they meet, AABW flows beneath NADW. As the NADW andAABW spread eastward into the Indian and Pacific Oceans, they mix to become Deep Pacific Common Water (DPCW).

    The "youngest" bottom waters are in the Atlantic, the "oldest" are in the North Pacific.

    When seawater is at the surface, it equilibrates with the atmosphere with respect to O2 and CO2. From the time a water

    mass sinks from the surface until it comes back to the surface, respiration by organisms in the water column and on the

    bottom use up O2 and add CO2. As a result, the longer bottom water is away from the surface, the more corrosive it is to

    CaCO3.

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    Carbonate Sedimentation Worldwide

    he depth of the CCD and the pattern of carbonate sedimentation in any part of the world's ocean reflects the

    influences of surface production of organic matter, surface production of carbonates, and the corrosiveness of the

    bottom water to CaCO3.

    Because coccolithophorids and planktic foraminifera thrive in temperate to subtropical oceans where surface nutrient

    supplies are very limited, these organisms produce a continual rain of CaCO3 to the sea floor. In equatorial upwelling

    zones, organic productivity is elevated enough to stimulate higher rates of production of calcareous and siliceous skeletal

    remains, but not enough to export excess organic matter to the deep ocean where its respiration would increasecorrosiveness of bottom waters to CaCO3.

    In more intensive upwelling zones, especially in the eastern tropical Pacific and the Antarctic divergence, and off major

    river deltas, high nutrient supplies stimulate high rates of organic productivity by diatoms and dinoflagellates, often to the

    exclusion of coccolithophorids and planktic foraminifera, which reduces CaCO3 production. At the same time, the rain of

    organic matter to the ocean floor supports abundant deep-sea life whose respiration adds significantly to the CO2 in

    bottom waters. The result is substantial shoaling of the lysocline and CCD in these regions. The greater corrosiveness of

    AABW compared to NADW at approximately the same "age" is caused by upwelling-induced high organic productivity

    at the Antarctic divergence, which exports excess of organic matter into AABW.

    Pelagic sediments in the Atlantic and Indian Oceans are predominantly calcareous oozes. In the Pacific Ocean, where the

    CCD is deeper, red clays dominate, especially in the North Pacific. Carbonate oozes delineate shallower regions in thesouth Pacific, including the East Pacific Rise and the complex topography to the southwest.

    Siliceous Oozes

    Biogenic siliceous oozes have two major and two minor contributors.

    Golden-brown algae known as diatoms (Bacillariophyceae) construct a type of shell called a frustule out ofopalline silica.

    The radiolaria , a large group of marine protists distantly related to the foraminifera, also construct opalline silicaskeletons.

    Silicoflagellates are a minor group of marine algae that also construct opalline silica skeletons. Sponge spicules are also

    an important biogenic source of opalline silica in neretic waters, but are of minor importance in the deep sea.

    Silica is undersaturated throughout most of the

    world's oceans. As a result, extraction of silica from

    seawater for production of silica shells or skeletons

    requires substantial energy. Furthermore, for

    siliceous sediments to be preserved, they must be

    deposited in waters close to saturation with respect

    to silica and they must be buried quickly. Young

    seawaters that are highly undersaturated with respect

    to H4SiO4 are far more corrosive to SiO2 than are old

    seawaters that have been dissolving and

    accumulating H4SiO4 over hundreds to thousands of

    years.

    Seawaters around volcanic islands and island arcs

    tend to have higher concentrations of H4SiO4 in

    solution and therefore are more conducive to silica

    production in surface waters and silica preservation in sediments. Siliceous sediments are most common beneath

    upwelling zones and near high latitude island arcs, particularly in the Pacific and Antarctic. More than 75% of all oceanic

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    silica accumulates on the sea floor between the Antarctic convergence and the Antarctic glacial marine sedimentation

    zone. Accumulation rates of siliceous oozes can reach 4-5 cm/1,000 years in these areas.

    Conditions favoring deposition of silica or calcium carbonate are different . Silica solubility increases with decreasing

    pressure and increasing temperature. Silica is undersaturated in the oceans, but it is less undersaturated in deep water.

    Carbonate solubility increases with depth, and bottom waters become more undersaturated in calcium carbonate. The

    patterns of carbonate and silica deposits reflect different processes of formation and preservation, resulting in carbonate

    oozes that are poor in biogenic silica and vice versa.

    The diatoms are extremely important primary producers that benefit physiologically from rich supplies of dissolved

    inorganic nutrients. Under such conditions, their growth rates far exceed other phytoplankton and they can rapidly

    produce both organic matter and siliceous sediments. They thrive in areas of intensive upwelling and near terrestrial

    sources of dissolved nutrients, including silica. Silicoflagellates show similar distributions. On the other hand, because

    both groups require substantial nutrient resources for growth, they are never abundant where nutrients are scarce, and so

    are insignificant primary and sediment producers in subtropical gyres. Diatom oozes, which contain more than 30%

    diatom frustules, are found beneath the Antarctic divergence, off the Aleutian island arc in the far North Pacific, and

    beneath areas of intensive meridional upwelling such as the eastern tropical Pacific. These oozes contain a significant

    percentage of radiolarian and silicoflagellate skeletons as well. Diatom-rich muds are common on continental shelves and

    margins where runoff from land contributes terrigenous muds as well as nutrients that stimulate diatom production.

    Radiolaria, being protists, are slightly less dependent on the most nutrient-rich areas of the oceans. They are important

    contributors to siliceous oozes around the Antarctic, but radiolarian oozes (> 30% radiolarian skeletons) are primarily in

    the tropical Pacific beneath the equatorial upwelling zone and below the CCD. Above the CCD in this region, the

    sediments are calcareous with a significant siliceous component.

    After burial, most siliceous oozes remain unconsolidated, but a fraction dissolve and reprecipitate as chert beds or

    nodules. Chert is cryptocrystalline and microcrystalline quartz, which is very hard and impermeable. Chert beds are very

    difficult to drill, which has frustrated ocean drillers since the early days of the Deep Sea Drilling Project (DSDP). The

    abundance and widespread distribution of chert beds of Eocene age, discovered by the DSDP, indicate important changes

    in deep-sea chemistry over the past 50 million years.

    Authigenic Sediments

    substantial number of authigenic minerals are precipitated in situ on the sea floor, but only a few common examples

    will be discussed. Formation of these minerals depends on local geochemical conditions, including elemental

    abundances, water characteristics, proximity of hydrothermal sources, and rate of sediment accumulation. Precipitation of

    minerals on or within the sediments of the sea floor generally results from supersaturation of the element or compound

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    required to form the mineral. Supersaturation may occur as the result of change in oxidation state of an element from a

    soluble, reduced state to a lower solubility oxidized state, resulting in precipitation of a hydrogenous phase, such as iron

    and manganese crusts. Because authigenic mineral accumulation rates are often less than 1 mm/1000 years, resulting

    sediments are common only where terrigenous and biogenic accumulation rates are nearly zero. In many cases, crusts of

    authigenic minerals form where bottom currents prevent the accumulation of other sediments.

    Barite

    arite (BaSO4) occurs in crystalline or microcrystalline phases or as replacement material in fecal pellets in deep-seasediments. Barite concentrations average 1% in deep sea sediments, but can make up as much as 10% by weight of

    the carbonate-free fraction on the East Pacific Rise, where it is associated with hydrogenous iron oxide. Most (80%) of the

    elemental barite in the oceans enters through rivers, about 20% comes from hydrothermal vents. A major conduit of

    barium to ocean sediments is secretion by a group of deep-sea protozoans, the xenophyophorans that produce barite

    crystals in large quantities. Elemental barite is found in biogenic sediments and has been attributed to production by these

    organisms or by concentration in organic matter following the death of the organism. Deep-sea sediments tend to be richer

    in barite than slope-depth deposits. Sediment pore waters in the deep sea are saturated with respect to barite; preservation

    potential is estimated at 30% in oxidized sediment and much lower in anoxic sediments.

    In the Pacific, barite is found in radiolarian oozes beneath the equatorial upwelling zone. In the Atlantic, elevated barite

    concentrations are found on the mid-ocean ridges in areas of low sedimentation rates and where there is an abundance of

    ferromanganese or iron oxide from hydrothermal sources.

    Glauconite

    lauconite is a well-ordered K- and Fe-rich mica-structure clay mineral. It occurs as flakes or pellets, and may occur

    as infilling in foraminiferal shells and sponge spicules. It may occur in fissures in feldspars, as crusts on phosphorite

    nodules, and as replacement mineral in coproliths. The color is usually blue-green, but this depends on the original clay-

    type and chemical composition. For example, dark-green illitic clays alter to dark-green glauconite, while yellowish

    smectite clays alter to yellowish glauconite. It is usually associated with organic residues, indicating that organic matter

    plays a role in formation of the mineral. Bacterial activity may promote glauconite formation by producing micro-reducing conditions in the sediment.

    Glauconite deposits occur from 65

    o

    N to 80

    o

    N, but are most common on lower latitude outer shelves and slopes from 20-700 m water depth. Glauconite forms from micaceous minerals or muds of high iron content where sedimentation rates

    are relatively low. Associated sediments are mainly calcareous, with a high proportion of fecal pellets.

    Marine Phosphates

    hosphate concentrations are typically very low within the euphotic zone of the oceans because phytoplanktons extract

    phosphate nutrients to photosynthesize organic matter. Vertebrates also concentrate phosphate into apatite, from

    which their bones are constructed. Vertically migrating fish and invertebrates feed on phytoplankton and zooplankton in

    surface waters at night and retreat to the shelter of darker subsurface waters during the day. Excretion of wastes in

    subsurface waters, along with decay of organic matter settling through the water column, concentrates inorganic

    phosphate ions and compounds below the euphotic zone, especially within thermocline depths. Both organic matter and

    skeletal remains accumulate on the sea floor, where decay and dissolution return phosphate to solution in bottom waters.Where the seafloor is at thermocline depths, especially beneath upwelling surface waters, that promotes export of organic

    matter to the bottom and phosphate ions may become sufficiently concentrated to precipitate phosphatic nodules or crusts.

    The most important phosphatic mineral is microcrystalline carbonate fluorapatite. Phosphatic nodules and crusts typically

    form along continental shelves, upper continental slopes and on oceanic plateaus beneath upwelling surface waters and

    where bottom currents limit accumulation of detrital sediments. Typical areas of phosphatic deposition are the continental

    margins of Peru, Chile, and southwest and northwest Africa. Phosphorite nodules or crusts average 18% phosphate.

    Conglomerates of phosphatized limestone pebbles and megafossils in a matrix of glauconite may have up to 15%

    phosphate.

    B

    G

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    Marine phosphates and phosphorite deposits are also found associated with anoxic sediments. Phosphorite may form by

    replacement of carbonate by phosphate. Upwelling occurs in the southern Caribbean in the surface waters above the

    Cariaco Basin, resulting in export of organic matter to bottom sediments. Phosphate precipitation is occurring along the

    rim of the basin where anoxic water from the trench mix with oxygenated waters from above. Phosphate may also be

    adsorbed by hydrous iron minerals, aluminum oxides and clay minerals. This accounts for phosphate concentrations of 1-

    2% in some iron-rich, clay or zeolite sediments in the deep sea.

    Heavy Metals

    ron oxides are an important constituent in slowly accumulating deep-sea clays where they occur as amorphous or

    poorly crystalline reddish-brown coatings on clays and other minerals and as minute globules in the sediments. Iron-

    rich basal deposits are found in oxidizing environments on the crests and flanks of actively spreading ocean ridges. Here,

    brownish-stained carbonate oozes may contain up to 14% Fe2O3. Iron-manganese minerals in these sediments are

    commonly attributed to hydrothermal activity associated with ocean-ridge volcanism. These associations result from

    penetration of seawater into hot volcanic rock, where the seawater is heated and becomes acidic and reducing, by

    geochemically reacting with fresh lava. As the hot solution mixes with cold seawater, sulfides precipitate first. With

    further mixing, iron and manganous

    oxides precipitate, producing iron-rich

    basal sediments.

    As seawater percolates into hot, volcanic

    rocks, seawater sulfate reacts with

    reduced iron. Where the hot solutions are

    forcibly expelled from the rocks ( vents

    and fumeroles ), metal sulfides precipitate

    as crusts and chimneys up to several

    meters high ridges. Localized

    accumulation rates can be a meter per

    year. Deposits rich in Fe, Mn, Cu, and Zn

    can occur where there is hydrothermal

    activity on the sea floor. One of the most

    spectacular examples of ridge-crest

    metalliferous deposits was discovered in

    the Red Sea in 1963. Rather than

    localized vents, metals are concentrated in

    deep, brine-filled basins. Manganese

    micronodules (less than 1 cm in diameter), nodules (1-10 cm in diameter) and crusts or coatings form in sediments or on

    exposed hard surfaces in the deep sea ridges. These oxides are brown-black agglomerations of manganese and iron oxides

    in fine-grained silicates or iron

    oxide-rich groundmasses in

    detrital and biogenic grains.

    Accessory metals include Ni,

    Cu, K. Ca, and Co. Elemental

    distribution patterns within

    nodules are variable and

    depend both on the

    environment of deposition and

    the nature of the mineral

    phases they contain. Where

    redox potential is lower,

    nodules are more iron rich; in

    well-oxidized deep-sea

    settings, nodules are richer in

    Mn.

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    A nodule commonly forms around a nucleus such as a shark's tooth or volcanic fragment. Nodules grow in concentric

    layers that may represent changes in seawater composition during growth. Rates of nodule growth are 1-4 mm/106 years.

    They commonly occur where sedimentation rates are less than 5 mm/1000 years. Apparently, sporadic movement by

    benthic organisms burrowing through the sediments is sufficient to keep most nodules at the sediment surface, where they

    can grow. The greatest area of manganese nodule development occurs in the Pacific, where 75% of the equatorial and

    North Pacific deep sea floor is covered with nodule patches. Fields of nodules develop in areas swept clean of fine detrital

    sediments by bottom currents. Where nodules cover 100% of the sediment surface, the area is called a manganese nodule

    pavement. In some cases, nodules join to form a solid surface. Such pavements are found on deep plateaus including the

    Blake Plateau in the western North Atlantic and the Agulas Plateau south of South Africa.

    The manganese comes from terrestrial sources by wind and water transport. In the water column, plankton extract

    manganese from solution, then carry it to the bottom. Manganese is also scavenged from seawater and deposited on the

    bottom by organic aggregates. Local deep-water sources of manganese may be interstitial waters leaching sediments rich

    in Mn and Fe near basaltic rocks. Near mid-ocean ridges, nodules may derive their Fe, Mn and accessory minerals from

    volcanic sources, as noted above.

    : Although there is economic interest in both metalliferous sulfide deposits and in manganese nodules, the costs of mining

    currently exceed the value of the minerals.

    Organic-Rich Sediments

    rganic material is measured in sediment as total organic carbon (TOC) or particulate organic matter (POC) which in

    ocean water is primarily living organisms or the remains of dead organisms. Upon the death of an organism, its

    remains are subjected to chemical and bacterial degradation processes. Detrital POC, which is produced in surface waters

    by primary production, may sink through the water column as fecal pellets or as marine snow and flocculate into what is

    called the fluffy layer. Skeletal remains, including coccoliths, diatom frustules, foraminiferal tests and radiolarian

    skeletons, as well as clay particles and volcanic ash, sink along with the organic matter. Both organic and inorganic

    particles influence to some degree the water chemistry of the waters they pass through. Within the water column, organic

    matter provides food for filter-feeding animals, which remove usable compounds and package unusable materials,

    including inorganic debris, into fecal pellets. The greater size and density of these pellets greatly increases settling rates of

    this material.

    When the organic matter reaches the sea floor, it provides food for benthic filter-feeding and detritus feeding organisms,

    reducing the concentration of POC accumulating in the sediments relative to what reaches the sea floor. In the Panama

    Basin, which is an upwelling area, depth-stratified sediment trap studies indicate that approximately 5% of the particulate

    matter reaching the bottom are POC, yet TOC concentrations in the sediments are less than 2%. Utilizable organic matter

    is known as labile organic matter. The least degradable materials, which often include terrestrial cellulose brought to the

    deep ocean in gravity flows, are called refractory solid organic matter. In typical pelagic sediments, TOC concentrations

    are less than 1%.

    Most organic carbon in sediments accumulates under conditions of high primary productivity in surface waters and low

    oxygen in bottom waters or interstitial pore waters. As a result of coastal upwelling and runoff from land that provide

    nutrients to phytoplankton communities in surface waters, combined with relatively rapid sedimentation rates in these

    regions, roughly 50% of all organic carbon burial occurs on continental shelves and margins.

    Organic-rich sediments that accumulate where bottom waters are depleted of oxygen (anoxic) are called sapropels.

    Anoxic conditions develop either because of rapid influx of POC or because of stagnation of bottom waters. Though

    limited in extent in modern oceans, sapropels occur in a variety of settings, including semi-isolated basins with restricted

    bottom circulation and portions of continental margins or slopes that lie within the mid-water oxygen minimum zone and

    below upwelling zones.

    Late Quaternary deep-water sediments in the Black Sea provide an example of restricted bottom circulation under which

    sapropels (ooze or sludge rich in organic matter) formed. From 23,000 to about 9,000 years ago, when sea level was 40 m

    or more lower than today, the Black Sea was completely isolated from the Mediterranean and was a large, freshwater lake

    which was aerobic thoughout. As sea level rose following the last glacial advance, seawater began to occasionally spill

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    over the Bosphorus Sill into the Black Sea, filling the deeper parts of the basin with dense seawater. However, river runoff

    into the Black Sea kept surface waters fresh. Because of higher evaporation rates in the Mediterranean, most of the flow

    of water through the Bosphorus was freshwater from the Black Sea to the Mediterranean. The seawater filling the basin of

    the Black Sea was isolated from air beneath a layer of low density fresh water. Primary productivity in the surface waters

    rained organic matter into the deep waters, depleting all oxygen, so that by 7,000 years ago, anoxic conditions were fully

    developed. About 3,000 years ago, two-way circulation developed with the Mediterranean, driving turnover of the deep

    waters of the Black Sea and allowing deep sea marine faunas to become established.

    Examples of modern sapropel formation within the oxygen minimum zone beneath upwelling high productivity surface

    waters can be found on the continental slope of the Arabian Peninsula and in the California borderlands. Upwelling in thenorthwest Indian Ocean provides sufficient surface productivity to provide an excess of organic matter to sediments on the

    continental slope of the Arabian Peninsula where the oxygen minimum zone intersects the slope. Off California, the

    combined effects of sluggish circulation in semi-isolated basins, continental margin depths within the oxygen minimum

    zone, and high surface water productivity all contribute to accumulation of laminated, organic-rich sediments in the Santa

    Barbara basin.

    Anoxic sediments have been widespread in the past and are of great economic importance as source rocks for

    hydrocarbon deposits. Expansion and intensification of the oceanic oxygen minimum zone, probably during times of

    reduced thermohaline circulation, is one mechanism that seems to account for many sapropels. Deep basins connected

    only by shallow connections, which resulted in restricted bottom circulation. , were especially common during early

    stages of continental rifting that formed the Atlantic basins .

    Volcanic Marine Sediments

    olcanogenic sediments are either the primary or secondary result of volcanic activity. Aerial volcanic explosions

    produce marine pyroclastic sediments. Reworked fragments of volcanic rocks produce marine epiclastic sediments,

    which may originate from altered fragments of pyroclastic sediments or from submarine volcanic flows. Sediments that

    form on the seafloor, either as a result of submarine eruptions or from hydrothermal activity are called authigenicsediments. Deep-sea volcanic sediments vary in thickness from thin ash layers to extensive tephra deposits more than a

    kilometer thick near volcanic island arcs.

    Pyroclastic and epiclastic sediments are distributed in the marine realm by the same mechanisms that disperse terrigenous

    sediments: wind, streams, submarine gravity flows, ocean currents, and sea ice. However, because of the explosive nature

    of many volcanoes, eolian transport is more important. Tephra deposits are typically thickest on the leeward side of a

    volcano and thin with distance from the source. Volcanic ash in deep-sea sediments may be in discrete layers or dispersed

    through other sediments; thinner deposits are usually more dispersed. Local ashfalls are deposited within a few hundred

    kilometers of the source.

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    transformation to another chemical or mineral state. In many cases, current removal of fine sediments promotes chemical

    precipitation on and alteration of the hiatal surface.

    Current Motion in the Deep Sea

    ntil about 20 years ago, most geologists assumed that the deep sea floor was a tranquil environment, so that beyond

    the reach of gravity-driven transport from the continental margins, sediment transport was minimal. Bottom

    photography, seismic reflection surveys and detailed analyses of sediment cores have all revealed that assumption to be

    incorrect.

    Circulation of deep bottom waters is driven by four main factors: formation in source regions, deep-sea topography, inter-

    ocean connections, and the Earth's rotation. Thermohaline circulation is density driven as the most dense waters flow

    along the bottom of the deepest parts of the ocean. The effect of the Earth's rotation, the Coriolis effect, accelerates

    currents along the western sides of basins. Current flow is also accelerated over any topographic high or through any

    constriction or passageway. While most of the deep sea floor experiences rather slow currents (less than 2 cm/sec), current

    velocities of 10-15 cm/sec and higher have been recorded in areas of current acceleration. Furthermore, current velocities

    at midwater depths can be 2-3 times those on the bottom, so current velocities over seamounts can be strongly erosional.

    A variety of sedimentary features have been observed in deep-sea sediments, including ripples , mud waves, channels,

    furrows, and even dunes. Ripples can be formed by

    contour currents, which typically flow along

    bathymetric contours along western sides of basins. In

    passages, the bottom may be scoured of sediment, that

    lies in drifts on the downstream side. Erosion can

    cause unconformities or hiatuses in sediment

    accumulation, particularly in areas where flow is likely

    to intensify. Ocean drilling has revealed widespread

    hiatuses in the deep-sea record. Furthermore, the

    sediment water interface on the deep sea floor is not

    always an abrupt surface. More commonly, the bottom

    grades from the overlying water column, through a

    cloud of sediment particles known as the nepheloid

    layers, to consolidated sediment. The nepheloid layer

    is quite mobile and can be transported over large

    distances by bottom currents.

    Sediment Stabilization and Redistribution by Organisms

    n the deep sea, organisms bind and

    alter sediments in a variety of ways.

    Growth of bacterial mats binds sediments

    and alters water chemistry locally,

    particularly the redox potential.

    Burrowing organisms stir the sediments

    and add mucus and excretory products,which alters sediment chemistry. Deep-

    sea sponges and agglutinated foraminifera

    bind the sediments in which they live.

    Bioturbation occurs when organisms

    actively or passively disturb sediments

    and sedimentary structures mechanically.

    It affects sediment by changing physical

    properties such as resistance to erosion

    and porosity. It also influences the chemistry of interstitial waters by introducing oxygen into the sediments and by mixing

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    sediments away from and towards the sediment-water interface, where there is more oxygen. A great variety of

    organisms, including benthic foraminifera, annelids and other worms, arthropods, gastropods, bivalves, echinoids,

    holothurians, brittle stars and fish, cause some degree of bioturbation. Active bioturbators crawl along the bottom or

    burrow into the sediment, moving sedimentary particles as they go. Deposit feeding benthos ingest sedimentary particles,

    which are returned as fecal material or pellets.

    Trace fossils are the physical evidence of bioturbation. They are found in most ocean sediments except anoxic ones. Trace

    fossils are preserved in the geologic record where they attest to the variety and activity of ancient benthic life.

    Bioturbation destroys layering, so only sediments deposited in the absence of burrowing organisms are laminated.

    History and Nature of Paleoceanography

    simple definition of "paleoceanography" is "the study of the development of ocean systems In a larger context, it

    involves the study of the interconnectedness of Earth systems. That interconnectedness is reflected in the history of

    paleoceanography, for it demonstrates how a multitude of human endeavors, including detailed scientific descriptions by

    individuals and by teams of researchers, brilliant syntheses by individuals and groups, visionary leadership by scientists

    and politicians, wartime technologies, and large-scale international scientific cooperation, all contributed to

    revolutionizing our understanding of the oceans. This knowledge base and recognition of the interconnectedness of Earth

    systems will be crucial to development of philosophies, in the 21st century, for local, regional and global management of

    Earth resources for future generations of both humans and other inhabitants of the planet.

    The history of the oceans is recorded in the rocks and sediments of the ocean basins and margins. Deciphering that history

    has involved observations, research and discoveries in fields as diverse as geography, paleontology, petrology, structural

    geology, engineering, geophysics, sedimentology, geochemistry, and biological and physical oceanography. Kennett

    concluded that the rapid progress in Cenozoic paleoceanography has resulted from technical and conceptual

    breakthroughs in four major areas

    :

    engineering advancements that enabled recovery of deep-sea sediment cores; development of biostratigraphic schemes that are

    chronologically calibrated, which provided a

    temporal framework for interpreting deep-sea

    cores;

    development of the concept of plate tectonics,which provided the context for interpreting

    paleogeography

    development of numerous paleontologic,geochemical and mineralogical techniques to

    interpret paleoenvironmental conditions under

    which sediments were deposited.

    Stratigraphic Time Frames

    nalysis of deep-sea cores and samples ranges from

    time-honored fossil identification and sedimentgrain-size analysis to use of the most sophisticated

    geophysical and geochemical tools. Data from high

    technology procedures are by no means more valuable

    than basic fossil and sedimentological evidence. In fact,

    fossils and sediments are the direct records of

    oceanographic processes; geochemical data can be

    irrevocably modified by diagenesis or can be

    misinterpreted because the biogeochemical processes

    that influenced a particular geochemical record might be

    poorly known or misunderstood.

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    The most important kinds of data from sediment cores are relative age dates that allow cores to be compared with one

    another. There are many ways to do this and usually several methods are used. Biostratigraphic correlations , based upon

    the makeup and changes in assemblages of planktic foraminifera, coccoliths, radiolaria and/or diatoms, are the basic

    means of comparing cores. Because these groups of microorganisms have different ecologic requirements and because

    their remains tend to be preserved under quite different deep-sea conditions, ocean-wide correlations require use of all of

    these groups. Because remains of silicious and calcareous microorganisms are scarce to absent in deep-sea clays, analyses

    of fish debris (ichthyoliths), spores and pollen are required to correlate those sediments. Evolutionary changes in plants

    and animals are unidirectional, so assemblages for any biostratigraphic zone are unique to that zone, which represents a

    relative time unit. Sedimentological, geophysical and geochemical data, with a few exceptions, provide records of abruptfluctuations or gradual changes that have occurred numerous times in Earth history. Such fluctuations can often be

    correlated and may provide greater resolution than microfossils, but require microfossil data to accurately place within the

    relative time frame.

    Paleomagnetic measurements are still among the most important geophysical data collected from deep sea cores. Most

    marine sediments contain little material of use in radiometric dating, which is the closest thing to "absolute" age datingavailable in geologic research. Thus, absolute age dates are often assigned by a three or more step process. Microfossils

    are used to determine the relative age of a sample, whose paleomagnetic signature is also determined. The known

    paleomagnetic episode from a deep sea core is correlated with its counterpart from a terrestrial volcanic event whose rocks

    have been dated radiometrically. That is how paleoceanographers estimate that a particular event occurred , for example,

    36.5 million years ago.

    Emiliani proposed that stable isotope signatures in fossiliferous sediments would provide high resolution stratigraphy, and

    that has occurred with technological advances in mass spectrometry. The highest resolution schemes are based upon the

    integrated use of biostratigraphy, magnetostratigraphy, and isotope stratigraphy. High-resolution isotope sequences are

    often interpreted in the context of Milankovitch cycles of 22,000, 41,000 and 96,000 years.

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    This summary will follow a modified Kennett approach, with the addition of a brief discussion of key events of the latest

    Paleozoic and early Mesozoic that provide a context for understanding late Mesozoic and Cenozoic events. Key topics

    that will be summarized will be

    changed paleogeographic setting sea level predominant marginal and deep-sea sediment types major sediment-producing biota

    global paleoclimatic patterns surface and deep-sea paleocirculation

    Late Paleozoic Setting

    he major continental masses came together during the late Paleozoic to form one supercontinent, Pangea, surrounded

    by a superocean, Panthalassia. Sea level was low relative to this supercontinent, in part because plate movements that

    drove the continents together reduced the global continental area relative to global oceanic area as the continents were

    sutured together. A small-scale, Cenozoic analogy is the drop in sea level that resulted from the collision of India with

    Asia to form the Himalayas.

    The continental area lost

    during the collision

    (crumpled into theHimalayas) is roughly the

    area of modern India. During

    the collision, the Earth's

    ocean area increased by

    roughly the area of modern

    India, which is an

    approximately 0.8%

    increase. Since the average

    ocean depth is about 3,800

    m, increasing the area by 0.8% reduces the depth comparably, resulting in a sea level drop of roughly 30 m. The Paleozoic

    lowering would have been much greater.

    Radiolaria were the only significant producers

    of biogenic pelagic sediments in the Paleozoic;

    calcareous producers of pelagic sediments had

    not yet evolved. Therefore, both deep sea

    sediments and oceanic biogeochemical cycles

    were quite different from those of the mid-late

    Mesozoic and Cenozoic. The relatively few

    deep-sea sediments preserved as sedimentary

    rocks were primarily of terrigenous or volcanic

    origin. Shelf carbonates are common in the

    early-mid Permian records from the

    southwestern North America, the Perm regionof Russia and elsewhere. Late Permian

    sequences are dominated by evaporites and

    redbeds, the latter being evidence for

    widespread fluvial sedimentation from the

    eroding uplands.

    The paleohistory of massive, non-structured

    limestone bodies described as reef deposits

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    extends from Cambrian to the present. These are preserved, because the environment of deposition was the shallow shelf

    areas which were part of the continental blocks. The discussion of reef development will be presented for only the

    Cretaceous to modern reefs.

    Triassic Summary

    he paleogeographic setting for the Triassic was similar to that of the late Paleozoic; the continents were joined into

    one supercontinent , Pangea, and sea level was low relative to the continental margins. But terrestrial flood basalts

    interbedded with evaporites andredbeds indicate the onset of rifting of

    the continents, as heat from the mantle

    began to build beneath the massive

    supercontinent. The continent of Africa

    provides something of a small-scale

    modern analogy for Triassic Pangea.

    Africa lacks extensive continental

    shelves and its Great Rift Valley is

    characterized by flood basalts, redbeds

    and evaporitic lakes.

    In terms of neritic and terrestrial biotas,

    a great extinction event marks the

    Paleozoic-Mesozoic boundary, better

    known as the Permian-Triassic

    boundary. Approximately 95% of

    fossilizable late Paleozoic species did

    not survive into the Triassic. The

    boundary is characterized by a prolonged hiatus of approximately 8 million years in neritic carbonate deposition. When

    carbonate deposition resumed in the Tethyan region during the middle Triassic, the sediment-producers were a

    depauperate biota of cyanobacteria, calcareous sponges and problematic taxa.

    Evolutionary events that occurred in the mid to late Triassic forever altered both neritic and pelagic sedimentation and

    geochemical cycles. The importance of the evolution of coccolithophorids and planktonic foraminifera cannot be

    overemphasized, for these events made possible the shift of large-scale carbonate sedimentation from shelves and shallow

    seas to the deep ocean. The series of events that altered shelf carbonate sedimentation included the appearance of

    Scleractinian corals in the mid Triassic. By the late Triassic, these corals apparently hosted algal symbionts, which

    allowed them to grow to much larger sizes and produce and trap much larger volumes of carbonate sediments as corals

    became the dominate reef-building organisms. Wood attributes the latter event to the evolution of dinoflagellates with the

    potential for entering into symbiotic relationships not only with corals, but also with planktic and benthic foraminifera and

    bivalve mollusks.

    Global paleoclimate was relatively uniform and relatively mild during the early Mesozoic. Surface circulation in

    Panthalassa was probably more symmetric between the northern and southern hemisphere than in the modern Pacific.

    North and south anticyclonic subtropical gyres were separated by an equatorial countercurrent; cyclonic subarctic gyres

    characterized the high latitudes.

    Jurassic Summary

    he Jurassic was the time of change from a supercontinent-superocean global setting to the rapidly separating

    continents of the Cretaceous. Modern analogies for the Jurassic can be found in modern rifts. In Ethiopia, the north

    end of the Africa's Great Rift Valley is periodically invaded by marine waters, accumulating thick sequences of

    evaporites. The Arabian Gulf and Red Sea provide examples of progressively later stages of rifting, the Arabian Gulf

    being characterized by shallow-water carbonates and evaporites, while the Red Sea is a deep basin connected to the Indian

    Ocean by a shallow seaway that strongly influences deep-water circulation. These rift settings provide some insight into

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    the depositional environments created by the initial breakup of Pangea as seaways and basins began to form between

    Laurasia (North America, Europe and Asia) and Gondwana (South America, Africa, India, Australia and Antarctica).

    By the early Jurassic, significant basins had begun to open in what is now the Gulf of Mexico. Thick sequences of

    evaporites were deposited as

    the deepening basin was

    alternatively joined to and

    isolated from oceanic waters.

    Those salt beds are the

    reason why salt domes arecommon around the Gulf of

    Mexico. Great quantities of

    recoverable hydrocarbons

    have been found trapped by

    these domes. Similar early to

    mid Jurassic evaporites are

    also found off eastern North

    America and west Africa. By

    approximately 190 million

    years ago, Laurasia and

    Gondwana were effectively

    separated, providing at leasta shallow-water opening for initiation of circumtropical circulation through the Tethys seaway.

    Sea level was relatively low in the early Jurassic, fluctuating throughout the period with an overall trend towardssubstantially higher levels in the Cretaceous. Factors driving sea level rise included relative increase in continental area as

    the continents were stretched, thinned and broken by rifting, subsidence as the continents moved away from spreading

    ridges, and accelerating rates of sea floor spreading. Fluctuations in sea level alternatively isolated and reconnected

    marginal seas, providing optimum conditions for the origin (in isolation) and subsequent dissemination of new taxa. With

    increasing sea floor spreading rates came increasing partial pressures of CO2 in the atmosphere, further ameliorating

    global climates.

    Pelagic sedimentation patterns are poorly known because most Jurassic seafloor has been subducted. The best known deep

    sea sediments of late Jurassic age are found in the North Atlantic. Jansa et al. recognized Oxfordian and Kimmeridgianlimestones overlain by Tithonian-Hauterivian chalk, the latter representing pelagic oozes produced primarily by planktic

    foraminifera and coccolithophorids. Along the margins and shallow seas of the Tethys, shallow water carbonates were

    widespread and diverse. Besides Scleractinian corals, major carbonate-producing organisms included coralline algae,

    sponges, and bivalves.

    Cretaceous Summary

    he Cretaceous Period is exceptional for a variety of reasons. On land, the "Age of the Dinosaurs" continued and

    concluded, while flowering plants (angiosperms) expanded in diversity and ecological importance. The appearance of

    benthic diatoms was significant, not so much for their influence in the Cretaceous, but for their future in the Cenozoic.

    The shallow marine realm was characterized by widespread carbonates. The shallowest shelves and epeiric seas of the

    expanded "Tethys" were dominated by a diverse biota of unique giant clams known as rudists . Scleractinian corals were

    common and diverse, particularly in slightly deeper waters along bank margins, but were secondary to the rudists in

    producing extensive limestone deposits. Most notable in the Cretaceous were the coccolithophorids and planktic

    foraminifera that produced widespread chalk deposits on the deeper shelves and epieric seas and in the open ocean. The

    French word "cretac" means "chalk"; "Terrain Cretac" (chalk terrains) are widespread in northern France and England,

    also in the Middle East, Australia and around the Gulf of Mexico. Cretaceous limestones and chalks are among the most

    common rocks worldwide.

    The Cretaceous was a relatively quiet time on the receding continents. The closest modern analogy is Australia, with its

    low mean elevation and extensive marginal temperate and tropical carbonate margins. It is moving northward away from

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    Antarctica; the plate collision margin is to the north of its shallow northern seas. Compare that with Cretaceous North (or

    South) America, moving westward followed the breakup of Pangea. Swampy lowlands bordered vast shallow shelves; the

    Western Interior Seaway separated the continent from the trench-island arc collision margin to the west. The major

    tectonic action was along the very actively rifting oceanic ridges, the rapidly subducting trenches, and the comparably

    rapidly accreting island arcs that surrounded the shrinking Pacific Ocean. Sea floor spreading rates of up to 10 cm/yr not

    only pushed sea level to record highs, but emissions of volcanic gases into the atmosphere from ocean ridges and island-

    arc volcanoes resulted in atmospheric CO2 concentrations 3-10 times higher than modern levels.

    The result of both high sea level and high CO2 concentrations were warm global climates, often called "Greenhouse

    World" conditions, in which polar regions were ice-free. There were only three major biogeographic regions, the northernboreal (temperate), Tethyan (tropical) and southern boreal provinces. Whether tropical climates were warmer or cooler

    than present tropics is controversial. Paleotemperature data based on stable oxygen isotopes, as well as some global

    climate models, indicate tropical ocean temperatures as much as 5o cooler than present (18-23o C), while paleontological

    interpretations indicate a core "Supertethys" several degrees warmer than the modern tropics. Because water loses as

    much energy during evaporation as it takes to heat water further, open ocean water temperatures cannot rise above 32o C,

    thereby limiting global warming.

    The globally mild climate had a profound effect on deep ocean circulation and sedimentation. Bottom water formation is

    thought to have been halothermal (driven primarily by salinity changes and secondarily by temperature changes), rather

    than the thermohaline mode in modern oceans. A modern analogy for halothermal bottom water formation is in the

    Mediterranean Sea. Evaporation exceeds freshwater input from rivers, so salinities in the Mediterranean are higher than in

    the Atlantic. Local winter cooling (to 10-14o C) of this slightly hypersaline water increases density, resulting in sinking ofcooled water masses to form Mediterranean bottom water. In the case of the Mediterranean, normal salinity surface water

    from the Atlantic flows into the Mediterranean, while hypersaline Mediterranean bottom water flows out over the

    Gibralter sill, contributing Mediterranean intermediate water to subsurface North Atlantic circulation. Similar conditionsare believed to be responsible for most bottom water formation during the Cretaceous. Cool, slightly hypersaline deep

    waters initially carried less oxygen than do near-freezing, normal salinity modern bottom waters. Furthermore, rates of

    bottom water formation are estimated to have been 1-2 orders of magnitude slower, so rates of deep-water turnover were

    on the order of 104-105 years, rather than modern rates of 102-103 years. The significance for deep sea sedimentation were

    profound. The oxygen minimum zone was greatly expanded during much of the Cretaceous, sometimes including entire

    basins, resulting in widespread deposition of anoxic black sh