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
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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
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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.
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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
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