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Deepsea Sediments & Paleoceanography

Introduction

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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Development of the Basic Tools

rancis Bacon, in 1620, noted the geographic evidence for continental

movement, that is, the way the continents on either side of the Atlantic appear to fit together. Wegner, between 1912 and 1930, assembled evidence from fossils, rock types and structures indicating that during the Triassic, the continents were united into a single supercontinent that he called Pangaea. Unfortunately, Wegener lacked a credible geophysical mechanism to explain how and why the continents moved apart. As a result, his synthesis was discredited for more than 30 years. Studies of successions of invertebrate fossils played a major role in the development of the science of geology in the 19th century. The terms Paleozoic, Mesozoic and Cenozoic refer to "ancient", "middle" and "recent" life. Deep-sea oozes recovered by the Challenger expedition (1872-76) contained abundant shells and skeletons of foraminifera and radiolaria, but these microfossils were not considered useful for correlation because it was assumed that deep-sea environments were unchanging through time. However, in the oil fields of the United States, Russia, and elsewhere in the early 1900's, geologists quickly recognized that fossil foraminifera are extraordinarily useful in determining the relative ages of rocks and in correlating rock formations from place to place. Thanks to their tiny size, hundreds of such microfossils are often found in small pieces of limestone or chalk. Laying of the trans-Atlantic telegraph cables, which began in the mid-1800's, required knowledge of ocean depths. Depth measurements, tediously carried out using lead-weighted lines, resulted in the discovery of the mid-Atlantic ridge and other bathymetric features. During the first scientific expedition to survey the deep ocean, the Challenger expedition (1872-1876), the basic types of deep-sea sediments were described and classified by Sir John Murray. However, it was advances in engineering and geophysics in the first half of the 20th century that provided the technology necessary for the explosion in paleoceanographic research in the second half of the century. Physical and chemical studies of radioactive and stable isotopes, including their detection and measurement, during WWII also provided breakthroughs that were directly applicable to paleoceanographic research. Rates of decay of radioactive isotopes were recognized as potential radiometric clocks for determining the ages of rocks and sediments. Urey

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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 the context 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 oceanographic changes 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 plate tectonics, to develop the discipline of paleoceanography, and to greatly 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 with dedication 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 of terrigenous 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. Examples include 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 by upward 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 are supported 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 and water 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 erosive force 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 are more 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 µ fraction clays make up 50-70% of the total oceanic sediment. In the open ocean, particles less than 0.5 µm 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 µ 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 or siliceous 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, in turn, 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 makes them 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 margins and 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 CO2 is 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 and more 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 shallow depths in low latitudes are often dominated by pteropods shells, at intermediate depths by foraminiferal tests, below the lysocline and above the CCD by coccoliths, 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 is left 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 and AABW 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 increase corrosiveness 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 the south 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 of opalline silica.

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

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-sea sediments. 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 65o N to 80o 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 the northwest 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 authigenic sediments. 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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Pyroclastic material is generally diverse and poorly sorted; though, size sorting by the wind may occur with distance from the source. In more explosive eruptions , material ejected to heights of 5-12 km into the troposphere may extend several thousand kilometers from the source region; ash may remain in the atmosphere for up to a month. Global ashfalls occur after extremely explosive eruptions inject ash into the stratosphere, where extremely fine-grained ash may remain for a year or more. These events are an important source of sediments to the deep sea. This ash is extremely fine, usually < 1 (m in diameter. Sedimentation rates of volcanic ejecta range from meters per thousand years locally to roughly 1 mm/1000 years in the deep sea. Volcanic sediments react with seawater to produce a unique suite of clay minerals, the most common being the montmorillonite/smectite clays. A layer of volcanic ash in the deep sea can be altered to smectite in 20 million years. Palagonite is a dark reddish brown to yellow glass that is formed by hydration of lavas as they cool. Palagonite is subsequently transformed to zeolites, which are white or colorless hydrous aluminum silicates similar in composition to feldspars. Zeolites are commonly associated with brown clays in the deep sea.

Sediments of Extraterrestrial Sources

minor component of deep-sea sediments is of extraterrestrial origin, either from impacting comets and asteroids or from interplanetary dust particles derived from large, extraterrestrial bodies in the vicinity of the Earth. The most

common particles are ultramafic, fine-grained chondritic aggregates, probably originating from comets. Chondritic particles, which originated from chondritic meteorites, may be silicate rich or Fe-S-Ni rich. Important trace constituents include the rare earths and various Pt-group metals, including iridium (Ir). Iridium concentration in deep sea sediments is one of the most sensitive indicators of extraterrestrial matter, because Iridium is roughly 103 times more concentrated in extraterrestrial material than in the Earth's crustal rocks. Excess concentrations in marine sediments are considered to be derived from vaporized or dissolved meteoric material. Layers of anomalously high Iridium concentrations have been interpreted as caused by extraterrestrial impact events. The best known Iridium spike occurred about 64.5 million years ago at the Cretaceous/Tertiary boundary and is thought by many to represent a meteor impact event that occurred at about the same time as oceanic plankton and nekton underwent mass extinctions and the last of the dinosaurs disappeared from the continents. Microtectites are small (0.03-1 mm in diameter), glassy, round or teardrop shaped particles of extraterrestrial origin. They are commonly yellow to brown in color and may be smooth or pitted. Their rate of accumulation is usually very slow, less than 0.002 mm/1000 years, so they are only noticeable where sedimentation rates are very low or in rare concentrations where their presence is apparently related to terrestrial tectite strewn fields indicating meteor impact. Tectites are 2-4 cm glass bodies, black to green in color, which are thought to result from high-velocity impacts of meteorites with the Earth's surface. On continents, concentrations of tectites occur in strewn fields covering hundreds to thousands of kilometers, indicating geologically instantaneous events. There are four main strewn fields that have been recognized: the Australian (including Australia, Indonesia, and the Philippines), the African Ivory Coast, the Czechoslovakian, and the North American fields. Microtectites occur in deep-sea deposits adjacent to all except the Czechoslovakian field.

Post-Depositional Processes in the Deep Sea

ost-depositional processes that can change or even remove deep-sea sediments include physical transport, disturbances by organisms, chemical alteration, dissolution and diagenesis. Some of these processes have already been

discussed, including as gravity-driven sediment transport mechanisms, carbonate dissolution, and authigenic mineral precipitation. Post-depositional processes also often interact, i.e., burrowing by organisms alters the chemical environment and can alter susceptibility to physical transport. Hiatuses or gaps in the sedimentary record are produced where erosion or dissolution removes sediment layers. Removal of sediments by currents results in the redistribution and redeposition downcurrent. Hardgrounds are produced where entire sedimentary layers are transformed chemically and the original sediment is lost to dissolution, reprecipitation or

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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 sediment

grain-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 abrupt fluctuations 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 dating available 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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Plate Tectonics as a Context for Interpretation

aleogeographic reconstructions based upon the theory of plate tectonics provide a context for interpreting the paleoceanographic record. For example, sequences of tropical limestones in the Emperor Seamounts in the subpolar

northwest Pacific are readily interpretable if the seafloor/conveyor belt, upon which these seamounts sit, has moved from low to high latitudes over the past 100 million years. Equally important is understanding the depth history of the ocean floor. There is a relatively simple relationship between age of the sea floor and its depth, caused by crustal cooling with distance from the spreading center. Sea floor depth (D) can be estimated as D = 2,500 + 350 t0.5 where t is age of the floor in millions of years. This depth history provides a context for interpreting a sediment sequence of, for example, basalt crust overlain by calcareous ooze grading upward into siliceous ooze and finally into brown clay, that might be seen in a deep-sea sediment core taken in the northwest Pacific, where subsidence has occurred. Perhaps most importantly, recognition of oceanic plate creation and subduction, and accompanying breakup and movement of continental plates, provides a context for interpreting major changes in oceanic circulation and global climatic patterns that have been recognized and will be summarized in the final section of this chapter. On the other hand, paleoceanographic evidence, including changes in sediment composition and texture, biogenic constituents and geochemical characteristics, have contributed immeasurable to interpreting the timing and effects of key tectonic events. Tools and Techniques for Paleoceanographic Interpretations The history of the oceans is recorded in the sediments and rocks of the deep sea and ocean margins. We read that history by studying the characteristics of those sediments and rocks by seismic profiling, sediment sampling and ocean drilling. The most basic data from a sediment core is the lithologic sequence, i.e., what are the types of sediments and rocks recovered in the core and in what order. As indicated in the chapter on deep sea sediments, the composition of the sediments has been influenced by the climatic conditions under which the terrigenous sediments eroded and in which the organisms that produced the biogenic sediments lived. Physical and geochemical conditions of the bottom-water masses and interstitial waters determined whether the sediment particles that reached the sea floor were preserved or altered. There are three important long-term, unidirectional phenomenon that influence Earth history and the paleoceanographic record.

• as the Sun has aged, its solar output has increased about 40% • the segregation of the Earth into the heaviest components in the core and the lightest components in the

atmosphere, driven by the escape of heat from the Earth's interior and the gradual cooling of the Earth • the evolution of life, which has profoundly changed the Earth's atmosphere from a CO2-rich, reducing atmosphere

to a CO2-poor, oxidizing atmosphere.

A Paleoceanographic Summary

here are two basic approaches to paleoceanography. • to define "paleoceanography" as the study of ocean systems through Earth history from whatever sedimentary

records are available (used by Schoff). • to consider the realm of paleoceanography to be that which can be studied from the deep sea record (used by

Kennett). • Since the oldest areas of sea floor are Jurassic and Cretaceous in age, that approach emphasizes the study of the

development of modern glacial-interglacial ocean systems from the "Greenhouse World" of the Late Mesozoic.

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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 the Himalayas) 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 region of 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 and redbeds 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 are common 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 least a 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 towards substantially 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 Kimmeridgian limestones 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 "cretacé" means "chalk"; "Terrain Cretacé" (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 northern boreal (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 of cooled 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 conditions are 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 shales . Even where deep waters were oxygenated, they had much longer to accumulate CO2 and therefore were more corrosive to calcareous sediments, resulting in relatively shallow carbonate compensation depths. Thus, despite extensive pelagic production of calcareous particles by planktic forams and coccolithophorids, there was strong fractionation of carbonate sediments, with widespread chalks and limestones representing shallow to upper slope depths. Deep sea sediments were predominantly organic-rich clays in the lower to middle Cretaceous and multicolored clays in the late Cretaceous. For example, in the expanding North Atlantic basins, Hauterivian chalks were replaced by Aptian-Cenomanian black bituminous shale, followed by multicolored and red clays during the Cenomanian. Because the western North Atlantic was bordered by higher continental terrain than the eastern North Atlantic, terrigenous sedimentation rates were higher in the west. In fact, much of what is now northwestern Europe was shallow epieric sea, accumulating chalks and rudistid limestones. During the latest Cretaceous (middle Maestrchtian), the CCD deepened to in excess of 5 km in the North Atlantic and deposition of calcareous oozes commenced throughout the North Atlantic. The opening of the South Atlantic occurred from south to north during the early Cretaceous. The Cape and Argentine Basins opened first, followed later by the Brazil and Angola Basins; the Rio Grande Rise-Walvis Ridge separated the southern basins from those to the north. The Brazil and Angola Basins accumulated thick evaporite sequences during the early Cretaceous, while the Cape and Argentine were characterized by terrigenous and black shale deposition. Isolation of the northern basins continued until the late Cretaceous, though sedimentation shifted from evaporites to black shales as the northern basins expanded and deepened. Sedimentation in the southern basins was more similar to that of the North Atlantic, oxygenated terrigenous sediments and clays deposited under a shallow CCD. Eroding land masses on either side of the expanding South Atlantic delivered terrigenous sediments into these basins throughout most of the Cretaceous. Permanent connection between the North and South Atlantic commenced about 90 Ma, establishing open ocean conditions throughout the Atlantic . Nevertheless, the deep sea topography created by the mid-Atlantic Ridge, the Rio

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Grande Ridge and the Walvis Rise promoted subtle interbasinal differences in the CCD and deep-sea sedimentation, even in modern times.

At about the same time as South America and Africa were separating, India began to rift away from Australia-Antarctica. Subsequently, in the Late Cretaceous, India began to separate from Madagascar, as it continued its northward movement toward Asia. Despite the relatively shallow Cretaceous CCD, extensive areas of the newly developing Indian Ocean were sufficiently shallow and ridges common that carbonate sedimentation was extensive. Terrigenous and volcanic sediments have also been recorded, often organic rich, indicating deposition under anoxic conditions. Deep sea sedimentation in the North Pacific reflects the interplay of plate motion, the CCD and equatorial productivity. DSDP cores from this region characteristically show initial deposition of calcareous ooze on the newly formed seafloor at mid-ocean ridge depths. As the site moved off the ridge and deepened, deposition changed from carbonates to pelagic clays or biogenic silica, depending upon location relative to the equator. Biogenic silica, which lithified to chert, was deposited as the site traversed beneath the higher productivity surface waters of the equatorial upwelling zone. Sedimentation of residual pelagic clays proceeded as the site moved beneath the North Pacific gyre.

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Interestingly, even the guyots of the North Pacific reflect this geographic passage through equatorial waters. Scientists on ODP found that flat-topped seamounts, which are capped by sequences of shallow-water carbonates followed by pelagic carbonates, "drowned" as they crossed the equator. That is, shallow-water carbonate sedimentation on the banktops failed to keep pace with subsidence and relative sea-level rise during the transit through equatorial latitudes, so the former oceanic reefs became seamounts. One explanation for this curious phenomenon is the shift of benthic communities in response to elevated nutrient availability, as predicted by Hallock and Schlager. 1986 That is, as the banks moved into the equatorial upwelling zone, benthic communities shifted from predominantly carbonate sediment producing organisms like rudistid bivalves and calcareous algae, to predominantly fleshy algae and bivalves lacking reef-building potential and shallow-water carbonate sedimentation failed to keep pace with subsidence. Cretaceous-Tertiary Boundary

he extinction event at the Cretaceous-Tertiary boundary, while not the largest in the geologic record, certainly altered the course of geologic and evolutionary history. Approximately 75% of fossilizable plant and animal species became

extinct. Besides terminating the "Age of the Dinosaurs", this event especially affected biogenic producers of pelagic and neritic carbonate sediments. The cause of the extinction event is still controversial. Although there is strong evidence for a major impact by an extraterrestrial body (bolide) at the boundary, proponents of more "Earthly" causes cite obvious changes in paleoenvironments and declining biotas in the Late Maastrictian. as well as evidence for other large bolide impacts in the geologic record that did not perpetrate catastrophic extinctions. Clearly global environmental conditions were changing, stressing certain biotas. Some scientists are suggesting that both scenarios have merit; changing environmental conditions may have been eliminating some kinds of habitats and stressing the inhabitants of others, causing observed late Maastrictian declines in, for example, the rudistid bivalves. A massive bolide impact under these changing environmental conditions may therefore have perpetrated higher rates of extinctions than had a similar event occurred during more optimal conditions. Tethyan biotas, particularly marine organisms with calcareous shells and skeletons, suffered the highest rates of extinctions. Rudistid bivalves were completely eliminated. Scleractinian corals were substantially reduced. Coccolithophorids and planktic foraminifera lost all but a few species, as did larger benthic foraminifera. Both pelagic and neritic carbonate sedimentation ceased briefly, then resumed with depauperate assemblages. Paleocene-Eocene Summary

he continuity of Greenhouse World paleoenvironments through the Paleocene and early Eocene supports the hypothesis that a catastrophic event occurred at the Cretaceous-Tertiary boundary. Corfield suggested that, had there

not been a bolide impact, the Mesozoic might have ended 30 million years later, with the terminal Eocene event, when global environments made the first major step in the change from the Greenhouse World to the Icehouse (glacial) World of the late Cenozoic. Although warm-water biotas extended to latitudes as high as 60o at times during the late Paleocene, and particularly during the early Eocene, subantarctic conditions were generally cooling, from 20o C in the early Eocene, to 12-14o by the middle Eocene, to 10o C by the late Eocene. The paleogeographic changes that perpetrated paleoclimatic changes on Antarctica from ice-free, relatively temperate climates to continental glaciation did not involve any significant movement of Antarctica itself. The continent moved into polar latitudes in the Cretaceous, and has remained there for roughly 100 million years. Rather, it was the breakaway of Australia and South America, combined with events in the Tethys and North Atlantic that generated climatic deterioration and glaciation of Antarctica. Although these paleogeographic changes were fairly gradual, rifting in two high-latitude regions and collision in the eastern Tethys began the gradual changes that would eventually lead to profoundly different global climates and deep ocean circulation (terminal Eocene Event). In the North Atlantic, the Norwegian and Greenland Seas were opening during the Paleocene and Eocene, permitting surface exchange with the Arctic and likely influencing abyssal circulation. The collision of India into Asia commenced in the early Eocene, disrupting circum-equatorial circulation and initiating extensive terrigenous sedimentation in the northern Indian Ocean. Most importantly, in the middle Eocene, Australia began to move northward, altering surface circulation patterns in the high southern latitudes and triggering the onset of cooler deep water formation in the Late Eocene, reflected in cooler bottom water temperatures down to 7o C.

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New assemblages of warm-water coccolithophorids and planktic foraminifera evolved from survivors of the K-T extinctions, but relatively shallow CCDs throughout the oceans limited pelagic carbonate sedimentation to depths less than 3,500 m throughout most of the oceans. In both the North and South Atlantic, depositional hiatuses are widespread, so Paleocene sediments are poorly represented. A particularly widespread hiatus occurs at the Paleocene-Eocene boundary. Siliceous diatoms and radiolaria survived the K/T boundary with relatively few extinctions. The pattern of biogenic silica beneath Pacific equatorial waters continued, and, in the Eocene, cherts become more common in the Atlantic as well. A widespread chert horizon is common in Eocene cores from the North Atlantic, and where the chert is missing, there is typically a depositional hiatus, indicating erosive bottom-water movement. Carbonate sedimentation on shallow shelves and seas was dominated by coralline algae and smaller foraminifera in the early Paleocene, though larger foraminifera quickly rediversified and became major contributors to Paleocene and Eocene shallow-water carbonate buildups. Scleractinian corals survivde but were reduced in diversity to about thirty genera. The Paleocene reefs had corals that were survivors from the Cretaceous. Most reef-building families were present through the Eocene and diversification occurred. With the Eocene, there was a new radiation of zooxanthellate corals. The terminal Eocene event appears to hve been the onset of deep-sea thermohaline circulation, driven by large-scale sea ice formation in the Ross Sea of Antarctica. Opening of the Tasmanian Seaway, which isolated Antarctica from Australia, is thought to be the trigger for this event. Before the opening, the Ross Sea was warmed by the East Australian current, but formation of the seaway effectively isolated Antarctica from its influence. Bottom water temperatures dropped globally by 4-5oC in as little as 105 years. The terminal Eocene event provides an excellent example of how a relatively simple, regional tectonic change can have profound global influence. Both deep-sea benthic and planktic communities were substantially changed. Deep-sea faunas were rather suddenly exposed to colder waters carrying significantly higher concentrations of O2 and lower concentrations of CO2. The CCD abruptly deepened by 1,000 m or more in lower latitudes. "Warm-water" assemblages of planktic foraminifera and coccolithophorids were rapidly replaced by more temperate species, while very low diversity assemblages of planktic foraminifera with simple morphologies replaced temperate species in the high latitudes. Terrestrial biotas were also abruptly changed in mid to high latitudes, as mean annual temperature range increased from a few degrees to as much as 25o C in the Oligocene.

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Oligocene Summary

he Oligocene is perhaps the most enigmatic epoch of the Cenozoic. Apparently contradictory biotic and geochemical data have yet to be completely resolved. The onset of thermohaline deep sea circulation sharply increased the rate of

ocean turnover and return of nutrients to surface waters, yet biotic diversity plunged to Cenozoic minima, clearly demonstrating the almost inverse relationship between food supply and biotic diversity. CCD depths dropped, increasing pelagic carbonate sedimentation, and global paleoclimates were relatively cool, yet low latitude corals finally began to construct significant reefs. Reconstructions of Oligocene global marine circulation patterns superimposed on palinspastic restorations of plate positions and the experimental generation of paleo-oceanic circulation models indicate a number of important relationships. During the Paleogene, with a seaway connecting the northwestern Indian Ocean with the Mediterranean across what is now the Middle East, and with an open Isthmus of Panama, there was a tropical/subtropical westward flowing surface current through the western Tethys, across the central Atlantic, and through the Caribbean into the eastern Pacific Ocean. The North American gyre was smaller than at present and positioned far enough north to have relatively little effect in diverting warm westward flowing water up the northeast coast of North America as does the modern Gulf Stream. Thus, there is evidence to contend that migration and dispersal patterns of the cosmopolitan Oligocene reef coral fauna were from east to west. Though sea level was low during the early Oligocene, the late Oligocene-Early Miocene transgression and global amelioration resulted in latitudinal expansion and diversification of warm-water biotas. Crucial tectonic events occurred during the Oligocene, altering both tropical and polar surface circulation. The eastern Tethys closed in the Late Oligocene, further restricting circumtropical circulation. The Drake Passage between Antarctica and South America opened in the latest Eocene or earliest Oligocene, and may have been involved in the terminal Eocene cooling. With the Drake Passage open to shallow water, as Australia moved north, Antarctica became progressively isolated. By the early late Oligocene, the South Tasman Rise separated from Victoria Land, Antarctica, initiating the Circum Antarctic Current. The Oligocene marked a time of world wide maximum abundance and diversity of Tertiary reefs, with their widespread development in both the Caribbean/Gulf of Mexico and the Tethys regions. The zooxanthellate corals that were largely responsible for the construction of these reefs comprise the direct ancestors of the modern Indo-Pacific hermatypic coral fauna. The closing of the Tethys seaway passage into the Mediterranean at the close of the Oligocene led to the extinction of major elements of Oligocene reef coral fauna. Miocene-Pliocene

ollowing the opening of the Drake Passage to deep water in the late Oligocene, major tectonic events shifted from the southern hemisphere to the north. Despite the thermal isolation of Antarctica and active winter sea-ice formation

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generating the formation of Antarctic Bottom Water, glacial accumulation on Antarctica was apparently limited, not by cold, but by moisture. Events in the northern oceans soon changed that. Three key areas were the Iceland Faeroe Ridge, the Central American Seaway, and the Straits of Gibralter. First came the submergence of the Iceland-Faeroe Ridge in the middle Miocene, completing the deepwater connection of the Norwegian and Greenland Seas to the North Atlantic Basin. Bottom waters generated by sea ice formation in these northern seas could now flow into the North Atlantic, initiating North Atlantic Deep Water (NADW) formation. NADW is not as cold and not quite as dense as AABW, so after it began to form and flow southward, it began to emerge at the surface in the Antarctic divergence. Surfacing of these slightly warmer waters increased rates of evaporation to the atmosphere, which subsequently slightly increased snowfall on the Antarctic continent, resulting in a significant increase in thickness of the continental glaciers and a lowering of sea level in the late Miocene. The onset of NADW influenced pelagic sedimentation throughout the Atlantic Ocean and increased sedimentation differences between the Atlantic and the Pacific. NADW carries low concentrations of CO2 through the deep North Atlantic so that, wherever it reached the sea floor, carbonate oozes accumulated. Upwelling NADW in the Antarctic divergence promoted biological productivity and biogenic silica production, establishing high latitude diatom oozes that halo and intermix with glacial sediments rafted from Antarctica. Appearance of many new Scleractinian genera produced an Early Miocene fauna which is transitional between the cosmopolitan aspect of the Oligocene and endemic fauna of late Miocene-Holocene Caribbean. The zooxanthellate coral assemblages became substantially different from their Indo-Pacific counterparts by Late Miocene. This difference while partially increased by the evolution of endemic corals such as Agaricia, was achieved largely by the progressive regional extinction of corals such as Goniopora, Stylophora, Pocillopora, Pavona, Goniastrea and other long-ranging lineages in the Caribbean. Stephanocoenia and Madracis appear to extend back to late Early Miocene and Agaricia species and Helioseris cucullata appeared by Late Miocene; stock probably derived from a Pavona ancestor. Porites asteroides can be traced into Late Miocene. The gradual closure of the Central American Seaway influenced Caribbean biotas in the Miocene. The extinctions of a substantial proportion of the reef-building coral fauna and larger foraminifera may have resulted from an increase in nutrient flux to surface waters of the Caribbean. When the Central American Seaway was open, there was active exchange between the Caribbean and the eastern Pacific. The easterly trade winds drove surface currents westward from the Caribbean to the Pacific. However, because evaporation rates are higher in the Atlantic than in the Pacific, sea level was slightly higher in the Pacific, forcing subsurface flow back into the Caribbean. As long as the passages were fairly deep and open, backflow of nutrient rich, eastern tropical Pacific subsurface waters had minimal influence on shallow-water biotas. But as the seaway shoaled and became more complex, topographically induced upwelling of eastern Pacific waters increased delivery of nutrient-rich waters to shallow-water communities. Pliocene compression of climatic belts and the rise of the Isthmus of Panama restricted reef growth to two distinct regions - the Atlantic-Caribbean and the Indo-Pacific. An episode of faunal turnover affected Caribbean coral during the Plio-Pleistocene. The late Miocene and early Pliocene reef coral are distinct from the Pleistocene and modern reef communities. Acropora palmata is only found in later Pleistocene reefs so the time of origin and evolution cannot be determined other than after earliest Pleistocene, but it became the dominant species on Caribbean reefs replacing Pocillopora dominance. Closure of the Central American Seaway also had global implications. As the passage closed, more and more of the Caribbean Current was diverted northward into a western boundary current via the Florida Loop Current and the Florida Current to join the Gulf Stream. Besides triggering erosion of the Nicaraguan Rise, the Florida Straits and the Blake Plateau, this enhanced western boundary current carried greater volumes of warm water to the far north Atlantic, where enhanced evaporation triggered increased snowfall. The closure of the Isthmus in the late Pliocene is thought to provide a key trigger for northern hemisphere glaciation, pushing global climates towards greater influence by cycles in the Earth's orbit (Milankovitch Cycles).

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Quaternary Paleoceanography

he colleagues of Charles Darwin, in the mid-19th century, recognized that vast ice sheets must have covered northern Europe and northern North America at some time in the not too distant past. The first long piston cores recovered by

the Albatross Expedition in 1947-48 not only revealed that the Pleistocene sediments in the deep sea recorded alternating climate states, but showed more fluctuations than had been recognized on land. In the deep sea record, these cycles are expressed by a variety of characteristics including changes in supply and distribution of terrigenous sediments, fluctuations in biogenic sediments and in the species making up those sediments, fluctuations in oxygen-18 content in foraminiferal shells, and changes in the relative abundance of carbonate in the sediments. There were two major factors that brought about the Icehouse World of the Neogene that culminated in the very active, Milankovitch-cycle Glaciation of the Quaternary. Plate movements in key areas, including the far North Atlantic, the Tethys seaway, and particularly those that isolated Antarctica, produced changes in global climate and in surface and deep ocean circulation changes. Falling atmospheric CO2 concentrations, from several times pre-anthropogenic levels, must also have played a role in global cooling. The causes of this drop are not well defined, but certainly involve reduced rates of sea-floor spreading, but may also have involved more efficient CO2 sequestering in sediments by a variety of groups of organisms that arose in the Mesozoic. Radiations of the highly efficient primary producers, the diatom and angiosperm floras, may have sequestered organic matter, especially in terrestrial and shallow-water sediments. The roles of widespread deposition of pelagic CaCO3 by planktic forams and coccolithophorids, in addition to shallow-water carbonate sedimentation in the subtropics and tropics by Scleractinian corals, calcareous algae and larger foraminifera, are unknown. Certainly glacial-interglacial atmospheric CO2 concentrations are significant, and even summer-winter differences indicate the potential for a close feedback between primary production and CO2. While over uptake of CO2 by organisms should cause cooling (reverse greenhouse conditions), progressive cooling shuts down primary production, allowing CO2 levels to catch up. Scientists are only beginning to appreciate range and nature of feedbacks between the hydrosphere, atmosphere, cryosphere and biosphere.

Summary

ifferent sources, transport mechanisms, rates of supply, and rates of dissolution combine to provide modern deposition rates and patterns of deep-sea sedimentation. Sediment supplies near the continents allow terrigenous

sediments to dominate over biogenic sediments on abyssal cones and plains. Hemipelagic sediments are a combination of terrigenous muds and biogenic sediments. In limited areas along continental margins beneath upwelling zones, these muds are black and organic-rich. The types of sediments found in the deep ocean and on the continental slopes and rises can be summarized. The distribution of deep sea sediments is related to their settling from the water column, bottom transportation by gravity flows during settling: turbidity currents, debris flow, slumping; transportation by geostrophic bottom currents; and chemical or biochemical precipitation on the ocean floor. Calcareous oozes and clays are topographically controlled because of increased solubility of carbonate with depth. Calcareous oozes dominate the sea floor in temperate and tropical latitudes at depths shallower than the CCD, which varies with latitude and ocean. In the deep basins of the Pacific, residual red clays and manganese nodules slowly accumulate. In localized areas along mid-oceanic ridges, authigenic sediments rich in minerals are found. Siliceous oozes do not follow this elevation controlled dichotomy; they are characteristic of areas of high fertility -- ocean margins and polar regions. Glacial marine sediments are restricted to high latitudes within the iceberg limits. Paleoceanography reveals how shifts in these patterns and processes provide evidence for the evolution of the modern ocean system.

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Summaries such as presented here may give the impression that most of the paleoceanographic problems and questions have been solved. This is certainly not the case. One of the fathers of modern geology, Charles Lyell said "the greater the circle of light, the greater the circle of darkness surrounding it." Better understanding of the role of terrestrial biota on short and long-term CO2 concentrations, understanding the role of carbonate-producing organisms in CO2 dynamics and climate change, the deep sea, etc. are all critical to scientific input that can help insure the continued liveability of this planet. Deep sea sediments may be grouped into several fairly simple categories. Terrigenous slide, slump, debris flows turbidites glacial marine

Origin on land with proximity to glaciers, deserts, rivers, volcanoes or mounain belts determining the amount of terrigenous sediment available.

brown clays Pelagic brown clays biogenic carbonate oozes –

Globigerina, Pteropod, Coccolithophore

siliceous oozes – Diatom, Radiolaria

Biogenic sediments are controlled by nutrient supply, temparature, salinity, oxygen and carbon dioxide content and by the pH of the surface waters of the ocean. These same factors influence the deep waters through control of the CCD. The physical and chemical factors are influenced by positions of the continents relative to the major climatic belts of the planet.

Authigenic manganese nodules phosphate nodules zeolites

Authigenic sediments are controlled by locations and extent of black smoker activity on the ridges, by where upwelling zones are found and by super-slow processes that only produce abundant sediment if all other types are essentially non existent

Other volcanic wind-borne extraterrestrial