Oceanography also called oceanology or marine science, is the branch of Earth science that studies the ocean. It covers a wide range of topics, including marine organisms and ecosystem dynamics; ocean currents, waves, and geophysical fluid dynamics; plate tectonics and the geology of the sea floor; and fluxes of various chemical substances and physical properties within the ocean and across its boundaries. These diverse topics reflect multiple disciplines that oceanographers blend to further knowledge of the world ocean and understanding of processes within it: biology, chemistry, geology, meteorology, and physics.
Oceanography is a scientific discipline concerned with all aspects of the world's oceans and seas, including their physical and chemical properties, their origin and geologic framework, and the life forms that inhabit the marine environment.
Traditionally, oceanography has been divided into four separate but related branches: physical oceanography, chemical oceanography, marine geology, and marine ecology.
Physical oceanography deals with the properties of seawater (temperature, density, pressure, and so on), its movement (waves, currents, and tides), and the interactions between the ocean waters and the atmosphere.
Chemical oceanography has to do with the composition of seawater and the biogeochemical cycles that affect it. Marine geology focuses on the structure, features, and evolution of the ocean basins.
Marine ecology, also called biological oceanography, involves the study of the plants and animals of the sea, including life cycles and food production.
Oceanography is the sum of these several branches. Oceanographic research entails the sampling of seawater and marine life for close study, the remote sensing of oceanic processes with aircraft and Earth-orbiting satellites, and the exploration of the seafloor by means of deep-sea drilling and seismic profiling of the terrestrial crust below the ocean bottom.
Greater knowledge of the world's oceans enables scientists to more accurately predict, for example, long-term weather and climatic changes and also leads to more efficient exploitation of the Earth's resources.
Oceanography also is vital to understanding the effect of pollutants on ocean waters and to the preservation of the quality of the oceans' waters in the face of increasing human demands made on them.
The first major undersea survey was undertaken during the 1870s, but it was not until the last half of the 20th century that scientists began to learn what lies beneath the ocean surface in any detail. It has been determined that the ocean basins, which hold the vast quantity of water that covers nearly three-quarters of the Earth's surface, have an average depth of almost four kilometres. A number of major features of the basins depart from this average, as, for example, the mountainous ocean ridges, deep-sea trenches, and jagged, linear fracture zones. Other significant features of the ocean floor include aseismic ridges, abyssal hills, and seamounts and guyots. The basins also contain a variable amount of sedimentary fill that is thinnest on the ocean ridges and usually thickest near the continental margins.
While the ocean basins lie much lower than sea level, the continents stand high--about one kilometre above sea level. The physical explanation for this condition is that the continental crust is light and thick, whereas the oceanic crust is dense and thin. Both the continental and oceanic crust lie over a more uniform layer called the mantle. As an analogy, one can think of a thick piece of styrofoam and a thin piece of wood floating in a tub of water. The styrofoam rises higher out of the water than the wood.
The ocean basins are transient features over geologic time, changing shape and depth while the process of plate tectonics proceeds. The surface layer of the Earth, the lithosphere, consists of a number of rigid plates that are in continual motion. The boundaries between the lithospheric plates form the principal relief features of the ocean basins: the crests of oceanic ridges are spreading centres where two plates move apart from each other at a rate of several centimetres per year.
Molten rock material wells up from the underlying mantle into the gap between the diverging plates and solidifies into oceanic crust, thereby creating new ocean floor. At the deep-sea trenches, two plates converge, with one plate sliding down under the other into the mantle where it is melted. Thus, for each segment of new ocean floor created at the ridges, an equal amount of old oceanic crust is destroyed at the trenches, or so-called subduction zones (see below Deep-sea trenches and also the article plate tectonics). It is for this reason that the oldest segment of ocean floor, found in the far western Pacific, is apparently only about 200 million years old, even though the age of the Earth is estimated to be at least 4.6 billion years.
The dominant factors that govern seafloor relief and topography are the thermal properties of the oceanic plates, tensional forces in the plates, volcanic activity, and sedimentation. In brief, the oceanic ridges rise about two kilometres above the seafloor because the plates near these spreading centres are warm and thermally expanded. In contrast, plates in the subduction zones are generally cooler. Tensional forces resulting in plate divergence at the spreading centres also create block-faulted mountains and abyssal hills, which trend parallel to the oceanic ridges. Seamounts and guyots, as well as abyssal hills and most aseismic ridges, are produced by volcanism. Continuing sedimentation throughout the ocean basin serves to blanket and bury many of the faulted mountains and abyssal hills with time. Erosion plays a relatively minor role in shaping the face of the deep seafloor, in contrast to the continents. This is because deep ocean currents are generally slow (they flow at less than 50 centimetres per second) and lack sufficient power.
Exploration of the ocean basins
Mapping the characteristics of the ocean basin has been difficult for several reasons. First, the oceans are not easy to travel over; second, until recent times navigation has been extremely crude, so that individual observations have been only loosely correlated with one another; and, finally, the oceans are opaque to light--i.e., the deep seafloor cannot be seen from the ocean surface. Modern technology has given rise to customized research vessels, satellite and electronic navigation, and sophisticated acoustic instruments that have mitigated some of these problems.
The Challenger Expedition, mounted by the British in 1872-76, provided the first systematic view of a few of the major features of the seafloor. Scientists aboard the HMS Challenger determined ocean depths by means of wire-line soundings and discovered the Mid-Atlantic Ridge. Dredges brought up samples of rocks and sediments off the seafloor. The main advance in mapping, however, did not occur until sonar was developed in the early 20th century. This system for detecting the presence of objects underwater by acoustic echo provided marine researchers with a highly useful tool, since sound can be detected over several thousands of kilometres in the ocean (visible light, by comparison, can only penetrate 100 metres or so of water).
Modern sonar systems include the Seabeam multibeam echo sounder and the GLORIA scanning sonar (see undersea exploration: Methodology and instrumentation: Exploration of the seafloor and the Earth's crust). They operate on the principle that the depth (or distance) of the seafloor can be determined by multiplying one-half the elapsed time between a downgoing acoustic pulse and its echo by the speed of sound in seawater (about 1,500 metres per second). Such multifrequency sonar systems permit the use of different pulse frequencies to meet different scientific objectives.
Acoustic pulses of 12 kilohertz (kHz), for example, are normally employed to measure ocean depth, while lower frequencies--3.5 kHz to less than 100 hertz (Hz)--are used to map the thickness of sediments in the ocean basins. Very high frequencies of 100 kHz or more are employed in side-scanning sonar to measure the texture of the seafloor. The acoustic pulses are normally generated by piezoelectric transducers. For determining subbottom structure, low-frequency acoustic pulses are produced by explosives, compressed air, or water-jet implosion.
Near-bottom sonar systems, such as the Deep Tow of the Scripps Institution of Oceanography (in La Jolla, Calif., U.S.), produce even more detailed images of the seafloor and subbottom structure. The Deep Tow package contains both echo sounders and side-scanning sonars, along with associated geophysical instruments, and is towed behind a ship at slow speed 10 to 100 metres above the seafloor. It yields very precise measurements of even finer-scale features than are resolvable with Seabeam and other comparable systems.
Another notable instrument system is ANGUS, a deep-towed camera sled that can take thousands of high-resolution photographs of the seafloor during a single day. It has been successfully used in the detection of hydrothermal vents at spreading centres (see below Oceanic ridges). Overlapping photographic images make it possible to construct photomosaic strips about 10-20 metres wide that reveal details on the order of centimetres.
Three major navigation systems are in use in modern marine geology. These include electromagnetic systems such as loran and Earth-orbiting satellites (see undersea exploration: Basic elements of undersea exploration: Navigation). Acoustic transponder arrays of two or more stations placed on the seafloor a few kilometres apart are used to navigate deeply towed instruments, submersibles, and occasionally surface research vessels when detailed mapping is conducted in small areas. These systems measure the distance between the instrument package and the transponder sites and, using simple geometry, compute fixes accurate to a few metres. Although the individual transponders can be used to determine positions relative to the array with great accuracy, the preciseness of the position of the array itself depends on which system is employed to locate it.
Such Earth-orbiting satellites as SEASAT and GEOSAT have uncovered some significant topographic features of the ocean basins. SEASAT, launched in 1978, carried a radar altimeter into orbit. This device was used to measure the distance between the satellite path and the surfaces of the ocean and continents to 0.1 metre. The measurements revealed that the shape of the ocean surface is warped by seafloor features: massive seamounts cause the surface to bulge over them owing to gravitational attraction. Similarly, the ocean surface downwarps occur over trenches. Using these satellite measurements of the ocean surface, William F. Haxby computed the gravity field there.
The resulting gravity map provides comprehensive coverage of the ocean surface on a 5' by 5' grid (five nautical miles on each side at the equator). Coverage as complete as this is not available from echo soundings made from ships. Because the gravity field at the ocean surface is a highly sensitive indicator of marine topography, this map reveals various previously uncharted features, including seamounts, ridges, and fracture zones, while improving the detail on other known features. In addition, the gravity map shows a linear pattern of gravity anomalies that cut obliquely across the grain of the topography. These anomalies are most pronounced in the Pacific basin; they are apparently about 100 kilometres across and some 1,000 kilometres long. They have an amplitude of approximately 10 milligals (0.001 percent of the Earth's gravity attraction) and are aligned west-northwest--very close to the direction in which the Pacific Plate moves over the mantle below.
Structure and composition
The oceanic crust differs from the continental crust in several ways: it is thinner, denser, younger, of different chemical composition, and created in a different plate-tectonic setting. The oceanic crust is formed at spreading centres on the oceanic ridges, whereas continental crust is formed above the subduction zones. The oceanic crust is about six kilometres thick. It is composed of several layers, not including the overlying sediment. The topmost layer, about 500 metres thick, includes lavas of basaltic composition (i.e., rock material consisting largely of plagioclase [feldspar] and pyroxene).
The lavas are generally of two types: pillow lavas and sheet flows. Pillow lavas appear to be shaped exactly as the name implies--like large overstuffed pillows about one metre in cross section and one to several metres long. They commonly form small hills tens of metres high at the spreading centres. Sheet flows have the appearance of wrinkled bed sheets. They commonly are thin (only about 10 centimetres thick) and cover a broader area than pillow lavas. There is evidence that sheet flows are erupted at higher temperatures than those of the pillow variety. On the East Pacific Rise at 8 S latitude, a series of sheet flow eruptions (possibly since the mid-1960s) have covered more than 220 square kilometres of seafloor to an average depth of 70 metres.
Below the lava is a layer composed of feeder, or sheeted, dikes that measures more than one kilometre thick. Dikes are fractures that serve as the plumbing system for transporting magmas (molten rock material) to the seafloor to produce lavas. They are about one metre wide, subvertical, and elongate along the trend of the spreading centre where they formed, and they abut one another's sides--hence the term sheeted. These dikes are also of basaltic composition. There are two layers below the dikes totaling about 4.5 kilometres in thickness. Both of these include gabbros, which are essentially basalts with coarser mineral grains.
These gabbro layers are thought to represent the magma chambers, or pockets of lava, that ultimately erupt on the seafloor. The upper gabbro layer is isotropic (uniform) in structure.
In some places, this layer includes pods of plagiogranite, a differentiated rock richer in silica than gabbro. The lower gabbro layer has a stratified structure and evidently represents the floor or sides of the magma chamber. This layered structure is called cumulate, meaning that the layers (which measure up to several metres thick) result from the sedimentation of minerals out of the liquid magma. The layers in the cumulate gabbro have less silica but are richer in iron and magnesium than the upper portions of the crust. Olivine, an iron-magnesium silicate, is a common mineral in the lower gabbro layer.
The oceanic crust lies atop the Earth's mantle, as does the continental crust. Mantle rock is composed mostly of peridotite, which consists primarily of the mineral olivine with small amounts of pyroxene and amphibole.
Investigations of the oceanic crust
Knowledge of the structure and composition of the oceanic crust comes from several sources. Bottom sampling during early exploration brought up all varieties of the above-mentioned rocks, but the structure of the crust and the abundance of the constituent rocks were unclear. Simultaneously, seismic refraction experiments enabled researchers to determine the layered nature of the oceanic crust.
These experiments involve measuring the travel times of seismic waves generated by explosions (e.g., dynamite blasts) set off over distances of several tens of kilometres. The results of early refraction experiments revealed the existence of two layers beneath the sediment cover. More sophisticated experiments and analyses led to dividing these layers into two parts, each with a different seismic wave velocity, which increases with depth. The seismic velocity is a kind of fingerprint that can be attributed to a limited number of rock types. Sampled rock data and seismic results were combined to yield a model for the structure and composition of the crust.
Great strides in understanding the oceanic crust were made by the study of ophiolites. These are slices of the ocean floor that have been thrust above sea level by the action of plate tectonics. In various places in the world, the entire sequence of oceanic crust and upper mantle is exposed. These areas include, among others, Newfoundland and the Pacific Coast Ranges of California, the island of Cyprus in the Mediterranean Sea, and the mountains in Oman on the southeastern tip of the Arabian Peninsula.
Ophiolites reveal the structure and composition of the oceanic crust in astonishing detail. Also, the process of crustal formation and hydrothermal circulation, as well as the origin of marine magnetic anomalies (see below), can be studied with comparative clarity. Although it is clear that ophiolites are of marine origin, there is some controversy as to whether they represent typical oceanic crust or crust formed in settings other than an oceanic spreading centre--behind island arcs, for example.
The age of the oceanic crust does not go back farther than about 200 million years. Such crust is being formed today at oceanic spreading centres. Many ophiolites are much older than the oldest oceanic crust, demonstrating continuity of the formation processes over hundreds of millions of years. Methods that may be used to determine the age of the crustal material include direct dating of rock samples by radiometric dating (measuring the relative abundances of a particular radioactive isotope and its daughter isotopes in the samples) or by the analyses of fossil evidence, marine magnetic anomalies, or ocean depth. Of these, magnetic anomalies deserve special attention.
A marine magnetic anomaly is a variation in strength of the Earth's magnetic field caused by magnetism in rocks of the ocean floor. Marine magnetic anomalies typically represent 1 percent of the total geomagnetic field strength. They can be stronger ("positive") or weaker ("negative") than the average total field. Also, the magnetic anomalies occur in long bands that run parallel to spreading centres for hundreds of kilometres and may reach up to a few tens of kilometres in width.
Marine magnetic anomalies were first discovered off the coast of the western United States in the late 1950s and completely baffled scientists. The anomalies were charted from southern California to northern Washington and out several hundred kilometres. Victor Vacquier, a geophysicist, noticed that these linear anomalies ended at the fracture zones mapped in this area. In addition, he noticed that they had unique shapes, occurred in a predictable sequence across their trends, and could be correlated across the fracture zones.
Soon thereafter, linear magnetic anomalies were mapped over the Reykjanes Ridge south of Iceland. They were found to occur on both sides of the ridge crest and parallel to it. Simultaneously, Alan Cox and several other American geophysicists documented evidence that the Earth's magnetic field had reversed in the past: the north magnetic pole had been the south magnetic pole about 700,000 years ago, and there were reasons to believe older reversals existed. Also at this time, Robert S. Dietz and Harry H. Hess were formulating the theory of seafloor spreading--the hypothesis that oceanic crust is created at the crests of the oceanic ridges and consumed in the deep-sea trenches.
It remained for Frederick J. Vine and Drummond H. Matthews of Great Britain and Lawrence W. Morley of Canada to put these observations together in a theory that explained marine magnetic anomalies. The theory rests on three assumptions: (1) that the Earth's magnetic field periodically reverses polarity; (2) that seafloor spreading occurs; and (3) that the oceanic crust is permanently magnetized as it forms and cools at spreading centres.
The theory expresses the assumptions--namely, that the oceanic crust records reversals of the Earth's field as it is formed during seafloor spreading. Positive anomalies result when the crust is magnetized in a "normal" polarity parallel to the ambient field of the Earth, and negative anomalies result when the crust is "reversely" magnetized in an opposite sense. As the magnetized crust moves down the flanks of a ridge away from the spreading centre, it remains permanently magnetized and "carries" the magnetic anomalies along with it. (For further details about paleomagnetism and seafloor spreading, see plate tectonics: Historical overview: Renewed interest in continental drift.)
A brilliant leap in understanding was now possible. If the age of the field reversals were known, the age of the ocean crust could be predicted by mapping the corresponding anomaly. By the mid-1960s, Cox and his colleagues had put together a schedule of reversals for the last four or five million years by studying the ages and magnetic polarities of lava flows found on land. Vine and the Canadian geologist J. Tuzo Wilson applied the time scale to marine magnetic anomalies mapped over the Juan de Fuca Ridge, a spreading centre off the northwest United States.
They thus dated the crust there and also computed the first seafloor spreading rate of about 30 millimetres per year. The rate is computed by dividing the distance of an anomaly from the ridge crest by the age of the anomaly twice. Thus the oceanic crust at the Juan de Fuca Ridge is moving at about 15 millimetres per year away from the ridge crest and at about 60 millimetres per year away from the crustal segment on the opposite side of the crest.
During the 1960s and '70s marine magnetic anomalies were mapped over wide areas of the ocean basins. By using estimates of the ages of oceanic crust obtained from core samples by deep-sea drilling, a magnetic anomaly time scale was constructed, and at the same time the spreading history for the ocean basins covering the last 200 million years or so was proposed.
It is thought that the most important contributor to marine magnetic anomalies is the layer of lavas in the upper oceanic crust. A secondary contribution originates in the upper layer of gabbros. The dike layer is essentially demagnetized by the action of hydrothermal waters at the spreading centres.
The dominant mechanism of permanent magnetization is the thermoremanent magnetization (or TRM) of iron-titanium oxide minerals. These minerals lock in a TRM as they cool below 200 to 300 C in the presence of the Earth's magnetic field. Although several processes are capable of altering the TRM, including reheating and oxidation at the seafloor, it is remarkably robust, as is evidenced by magnetic anomalies as old as 165 million years in the far western equatorial Pacific.
The largest features of the ocean basin are the oceanic ridges. In the past these features were referred to as mid-ocean ridges, but, as will be seen, the largest oceanic ridge, the East Pacific Rise, is far from a mid-ocean location, and the nomenclature is thus inaccurate. Oceanic ridges are not to be confused with aseismic ridges, which have an entirely different origin (see below).
Principal characteristics
Oceanic ridges are linear mountain chains comprising the largest features on Earth. They are found in every ocean basin and appear to girdle the Earth. The ridges rise from depths near 5 kilometres to an essentially uniform depth of about 2.6 kilometres and are roughly symmetrical in cross section. They can be thousands of kilometres wide. In places, the crests of the ridges are offset across transform faults, or fracture zones, which can be followed down the flanks of the ridges. (Transform faults are those along which lateral movement occurs.) The flanks are marked by sets of mountains and hills that are elongate and parallel to the ridge trend.
New oceanic crust (and part of the upper mantle, which, together with the crust, makes up the lithosphere) is formed at seafloor spreading centres at the crests of the oceanic ridges. Because of this, certain unique geologic features are found there. Fresh basaltic lavas are exposed on the seafloor at the ridge crests. These lavas are progressively buried by sediments as the seafloor spreads away from the site. The flow of heat out of the crust is many times greater at the crests than elsewhere in the world. Earthquakes are common along the crests and in the transform faults that join the offset ridge segments. Analysis of earthquakes occurring at the ridge crests indicates that the oceanic crust is under tension there. A high-amplitude magnetic anomaly is centred over the crests because fresh lavas at the crests are being magnetized in the direction of the present geomagnetic field.
The depths over the oceanic ridges are rather precisely correlated with the age of the ocean crust; specifically, it has been demonstrated that the ocean depth is proportional to the square root of crustal age. The theory explaining this relationship holds that the increase in depth with age is due to the thermal contraction of the oceanic crust and upper mantle as they are carried away from the seafloor spreading centre in an oceanic plate. Because such a plate is ultimately about 100 kilometres thick, contraction of only a few percent predicts the entire relief of an oceanic ridge. It then follows that the width of a ridge can be defined as twice the distance from the crest to the point where the plate has cooled to a steady thermal state.
Most of the cooling takes place within 70 or 80 million years, by which time the ocean depth is about 5 to 5.5 kilometres. Because this cooling is a function of age, slow-spreading ridges, such as the Mid-Atlantic Ridge, are narrower than faster-spreading ridges, like the East Pacific Rise (see below). Further, a correlation has been found between global spreading rates and the transgression and regression of ocean waters onto the continents. During the Early Cretaceous period about 100 million years ago, when global spreading rates were uniformly high, oceanic ridges occupied comparatively more of the ocean basins, causing the ocean waters to transgress (spill over) onto the continents, leaving marine sediments in areas now well away from coastlines.
Besides ridge width, other features appear to be a function of spreading rate. Global spreading rates range from 10 millimetres per year (mm/yr total rate) or less up to 160 mm/yr. Oceanic ridges can be classified as slow (up to 50 mm/yr), intermediate (up to 90 mm/yr), and fast (up to 160 mm/yr). Slow-spreading ridges are characterized by a rift valley at the crest. Such a valley is fault-controlled. It is typically 1.4 kilometres deep and 20 to 40 kilometres wide. Faster-spreading ridges lack rift valleys. At intermediate rates, the crest regions are broad highs with occasional fault-bounded valleys no deeper than 200 metres. At fast rates, an axial high is present at the crest. The slow-spreading rifted ridges have rough faulted topography on their flanks, while the faster-spreading ridges have much smoother flanks.
Distribution of major ridges and spreading centres
Oceanic spreading centres are found in all the ocean basins. In the Arctic Ocean a slow-rate spreading centre is located near the eastern side in the Eurasian basin. It can be followed south, offset by transform faults, to Iceland. Iceland has been created by a hot spot (see below) located directly below an oceanic spreading centre. The ridge leading south from Iceland is named the Reykjanes Ridge, and, although it spreads at 20 mm/yr or less, it lacks a rift valley. This is thought to be the result of the influence of the hot spot.
The Mid-Atlantic Ridge extends from south of Iceland to the extreme South Atlantic Ocean near 60 S latitude. It bisects the Atlantic Ocean basin, which led to the earlier designation of mid-ocean ridge for features of this type. The Mid-Atlantic Ridge became known in a rudimentary fashion during the 19th century. In 1855 Matthew Fontaine Maury of the U.S. Navy prepared a chart of the Atlantic in which he identified it as a shallow "middle ground." During the 1950s the American oceanographers Bruce Heezen and Maurice Ewing proposed that it was a continuous mountain range.
In the North Atlantic the ridge spreads slowly and displays a rift valley and mountainous flanks. In the South Atlantic spreading rates are between slow and intermediate, and rift valleys are generally absent, as they occur only near transform faults.
A very slow oceanic ridge, the Southwest Indian Ridge, bisects the ocean between Africa and Antarctica. It joins the Mid-Indian and Southeast Indian ridges east of Madagascar. The Carlsberg Ridge is found at the north end of the Mid-Indian Ridge. It continues north to join spreading centres in the Gulf of Aden and Red Sea. Spreading is very slow at this point but approaches intermediate rates on the Carlsberg and Mid-Indian ridges. The Southeast Indian Ridge spreads at intermediate rates. This ridge continues from the western Indian Ocean in a southeasterly direction, bisecting the ocean between Australia and Antarctica. Rifted crests and rugged mountainous flanks are characteristic of the Southwest Indian Ridge. The Mid-Indian Ridge has fewer features of this kind, and the Southeast Indian Ridge has generally smoother topography. The latter also displays distinct asymmetric seafloor spreading south of Australia. Analysis of magnetic anomalies shows that rates on opposite sides of the spreading centre have been unequal at many times over the past 50 or 60 million years.
The Pacific-Antarctic Ridge can be followed from a point midway between New Zealand and Antarctica northeast to where it joins the East Pacific Rise off the margin of South America. The former spreads at intermediate to fast rates.
The East Pacific Rise extends from this site northward to the Gulf of California, where it joins the transform zone of the Pacific-North American plate boundary. Offshore from Chile and Peru, the East Pacific Rise is currently spreading at fast rates of 159 mm/yr or more. Rates decrease to about 60 mm/yr at the mouth of the Gulf of California. The crest of the ridge displays a low topographic rise along its length rather than a rift valley.
The East Pacific Rise was first detected during the Challenger Expedition of the 1870s. It was described in its gross form during the 1950s and '60s by oceanographers, including Heezen, Ewing, and Henry W. Menard. During the 1980s, Kenneth C. Macdonald, Paul J. Fox, and Peter F. Lonsdale discovered that the main spreading centre appears to be interrupted and offset a few kilometres to one side at various places along the crest of the East Pacific Rise. However, the ends of the offset spreading centres overlap each other by several kilometres. These were identified as a new type of geologic feature of oceanic spreading centres and designated overlapping spreading centres. Such centres are thought to result from interruptions of the magma supply to the crest along its length and define a fundamental segmentation of the ridge on a scale of tens to hundreds of kilometres.
Many smaller spreading centres branch off the major ones or are found behind island arcs. In the western Pacific, spreading centres occur on the Fiji Plateau between the New Hebrides and Fiji Islands and in the Woodlark Basin between New Guinea and the Solomon Islands. A series of spreading centres and transform faults lie between the East Pacific Rise and South America near 40 to 50 S latitude. The Scotia Sea between South America and the Antarctic Peninsula contains a spreading centre. The Galápagos spreading centre trends east-west between the East Pacific Rise and South America near the equator.
Three short spreading centres are found a few hundred kilometres off the shore of the Pacific Northwest. These are the Gorda Ridges off northern California, the Juan de Fuca Ridge off Oregon and Washington, and the Explorer Ridge off Vancouver Island. In a careful study of the seafloor spreading history of the Galápagos and the Juan de Fuca spreading centres, the American geophysicist Richard N. Hey developed the idea of the propagating rift. In this phenomenon, one branch of a spreading centre ending in a transform fault lengthens at the expense of the spreading centre across the fault. The rift and fault propagate at one to five times the spreading rate and create chevron patterns in magnetic anomalies and the grain of the seafloor topography resembling the wake of a boat.
From the 1970s highly detailed studies of spreading centres using deeply towed instruments, photography, and manned submersibles have resulted in new revelations about the processes of seafloor spreading. The most profound discoveries have been of deep-sea hydrothermal vents (see below) and previously unknown biological communities.
Spreading centres are divided into several geologic zones. The neovolcanic zone is at the very axis. It is 1 to 2 kilometres wide and is the site of recent and active volcanism and of the hydrothermal vents. It is marked by chains of small volcanoes or volcanic ridges. Adjacent to the neovolcanic zone is one marked by fissures in the seafloor. This may be 1 to 2 kilometres wide. Beyond this point occurs a zone of active faulting. Here, fissures develop into normal faults with vertical offsets.
This zone may be 10 or more kilometres wide. At slow spreading rates the faults have offsets of hundreds of metres, creating rift valleys and rift mountains. At faster rates the vertical offsets are 50 metres or less. A deep rift valley is not formed because the vertical uplifts are cancelled out by faults that downdrop uplifted blocks. This results in linear, fault-bounded abyssal hills and valleys trending parallel to the spreading centre.
Warm springs emanating from the seafloor in the neovolcanic zone were first found on the Galápagos spreading centre. These waters were measured to have temperatures about 20 C above the ambient temperature. In 1979 hydrothermal vents with temperatures near 350 C were discovered on the East Pacific Rise off Mexico. Since then, similar vents have been found on the spreading centres off the Pacific Northwest coast of the United States, on the south end of the northern Mid-Atlantic Ridge, and at many locations on the East Pacific Rise.
Hydrothermal vents are localized discharges of heated seawater. They result from cold seawater percolating down into the hot oceanic crust through the zone of fissures and returning to the seafloor in a pipelike flow at the axis of the neovolcanic zone. The heated waters often carry sulfide minerals of zinc, iron, and copper leached from the crust. Outflow of these heated waters probably accounts for 20 percent of the Earth's heat loss. Exotic biological communities exist around the hydrothermal vents. These ecosystems are totally independent of energy from the Sun. They are not dependent on photosynthesis but rather on chemosynthesis by sulfur-fixing bacteria. The sulfide minerals precipitated in the neovolcanic zone can accumulate in substantial amounts and are sometimes buried by lava flows at a later time. Such deposits are mined as commercial ores in ophiolites on Cyprus and in Oman.
Magma chambers have been detected beneath the crest of the East Pacific Rise by seismic experiments. (The principle underlying the experiments is that partially molten or molten rock slows the travel of seismic waves and also strongly reflects them.) The depth to the top of the chambers is about two kilometres below the seafloor. The width is more difficult to ascertain, but is probably one to four kilometres. Their thickness seems to be about two to six kilometres based on studies of ophiolites. The chambers have been mapped along the trend of the crest between 9 and 13 N latitude. The top is relatively continuous, but is apparently interrupted by offsets of transform faults and overlapping spreading centres.
Fracture zones
As was noted above, oceanic ridges (and their associated spreading centres) are offset along their trend by fracture zones. These are ridges and valleys on the order of tens of kilometres wide that cut across the crests of the ridges at approximately right angles and offset their trend (Figure 9). Typically, a regional depth offset is present across a fracture zone, owing to the juxtaposition of crust of different ages (and, therefore, depth) across it. In the Atlantic, on the slow spreading Mid-Atlantic Ridge, fracture zones are numerous and occur every 55 kilometres on average along the trend of the ridge. They offset the crest between 5 and 40 kilometres.
Some of the larger fracture zones in the North Atlantic are the Gibbs at 52 N, the Atlantis at 30 N, and the Vema at 11 N. These and others can be followed across both flanks of the ridge for some 3,000 kilometres. The Vema Fracture Zone offsets the Mid-Atlantic Ridge 320 kilometres to the left. It is marked by a sediment-filled valley more than 5 kilometres deep and 10 to 20 kilometres wide and is flanked by mountains 3,500 metres high. Basalts, gabbros, and serpentinized peridotites (i.e., those peridotites that have been altered in varying degrees to serpentine) of the oceanic crust and mantle have been recovered from the mountain flanks.
Fracture zones occur less frequently on the East Pacific Rise, but they offset the ridge by a greater amount. More than a dozen can be found between 20 N and 30 S. Typical offsets are roughly 100 kilometres. Several fracture zones more than 3,000 kilometres long are found off the shore of western North America. These include the Mendocino, Murray, Molokai, and Clarion fracture zones. They are not associated with a ridge crest. Rather, they occur on the west flank of the defunct Pacific-Farallon oceanic ridge. The Farallon Plate has all but disappeared down a subduction zone that extended along the entire coast of California and Baja California until about 25 to 30 million years ago. Subduction now occurs north of the Mendocino Fracture Zone. These fracture zones off western North America were among the first mapped. Menard has traced them almost 10,000 kilometres westward across the Pacific. The continental margin of northern California is displaced to the right where the Mendocino Fracture Zone and its transform portion, the Gorda Escarpment, intersect it.
Transform faults
The portion of a fracture zone between different offset spreading centres constitutes a transform fault. Transform faults also connect spreading centres to subduction zones (deep-sea trenches). Faults of this kind are the only segments of fracture zones that are seismically active. J. Tuzo Wilson recognized this and other features and explained the phenomenon as a transfer of motion from one spreading centre to another. The American geologist W. Jason Morgan, one of the several outstanding pioneers in plate tectonics, recognized that transform faults are zones where opposing lithospheric plates slip past one another. Morgan proposed that opposing plates along an oceanic ridge crest offset by fracture zones are divided by the spreading centres and transform faults. The inactive portions of the fracture zone on the ridge flanks are scars on the ocean floor created in the transform faults.
This theory made a very dramatic prediction: namely, that the direction of motion on the transform faults was opposite to the offsets of the ridge crests. For example, if a ridge crest was offset to the left by a transform fault, implying leftward movement on a fault joining the offset crests, the movement across the transform fault was instead to the right (Figure 9). This is clear when it is realized that the plate boundaries are confined to the spreading centres and transform faults, not to the inactive part of the fracture zone. Seismic studies of earthquakes from transform faults soon revealed that the motion was opposite, as predicted.
Not everywhere in the ocean basins are plate motions exactly parallel to transform faults. In places where a component of opening motion occurs across the transform, volcanic activity results, and the fracture zone is termed a leaky transform fault. South of New Zealand, between it and the Pacific-Antarctic Ridge, a component of shortening is occurring across a transform called the Macquarie Ridge. Here, subduction may be taking place at a slow rate.
Types
Although the term trench has been applied to many deep, long linear troughs in the ocean floor, the most common and accurate usage relates it to subduction zones. According to plate tectonic theory, subduction zones are locations where a lithospheric plate bearing oceanic crust slides down into the upper mantle under the force of gravity. The result is a topographic depression where the oceanic plate comes in contact with the overriding plate, which may be either oceanic or continental. If the overriding plate is oceanic, an island arc develops (Figure 10). The trench forms an arc in plan view, and islands with explosive volcanoes develop on the overriding plate.
If the overriding plate is continental, a marginal trench forms where the topographic depression appears to follow the outline of the continental margin. Explosive volcanoes are found here too. Both types of subduction zones are associated with large earthquakes that originate at a depth of as much as 700 kilometres. The deep earthquakes below subduction zones occur in a plane that dips 30 or more under the overriding plate. Typical trench depths are 8 to 10 kilometres. The longest trench is the Peru-Chile Trench, which extends some 5,900 kilometres along the west side of South America. Trenches are relatively narrow, usually less than 100 kilometres wide.
The Pacific basin is rimmed by trenches of both marginal and island arc varieties. Marginal trenches bound the west side of Central and South America from the Gulf of California to southern Chile. Although they are deeply buried in sediment, trenches are found along the western North American continental margin from Cape Mendocino (in northern California) to the Canadian border. The Aleutian Trench extends from the northernmost point in the Gulf of Alaska west to the Kamchatka Peninsula in the Soviet Union. It can be classified as a marginal trench in the east but is more properly termed an island arc west of Alaska.
In the western Pacific, the trenches are associated with island arcs. These include the Kuril, Japan, Bonin, Mariana, Ryukyu, and Philippine trenches that extend from Kamchatka to near the equator. A complex pattern of island arcs is found in Indonesia. The major island arc here is the Java Trench extending from northern Australia to the northwestern end of Sumatra in the northeast Indian Ocean. The region of New Guinea and the Solomon Islands includes the New Britain and Solomon trenches, the latter of which joins the New Hebrides Trench directly to the south. East of this area the Tonga and Kermadec trenches extend south from the Fiji Islands to New Zealand.
Two island arcs occur in the Atlantic Ocean. The South Sandwich Trench is located west of the Mid-Atlantic Ridge between South America and Antarctica. The Puerto Rico Trench joins the Lesser Antilles Island arc in the eastern Caribbean. Some seafloor features bear the name trench and are deep linear troughs but are not subduction zones. The Vema Trench on the Mid-Indian Ridge is a fracture zone. The Vityaz Trench northwest of Fiji is an aseismic (inactive) feature of unknown origin. The Diamantina trench (Diamantina Fracture Zone) extends westward from the southwest coast of Australia. It is a rift valley that was formed when Australia separated from Antarctica between 60 and 50 million years ago.
The deepest water on Earth (11,034 metres) is located in the southern end of the Mariana Trench near Guam. A few trenches are partially filled with sediments derived from the bordering continents. The Aleutian Trench is effectively buried east of Kodiak Island in the Gulf of Alaska. Here, the ocean floor is smooth and flat. To the west farther from the sediment supply on Alaska, the trench reaches depths of more than seven kilometres. The Lesser Antilles trench in the eastern Caribbean also is buried by sediments originating from South America.
Structure
Oceanward of trenches the seafloor is usually bulged upward in an outer ridge or rise of up to 1,000 metres relief. This condition is thought to be the elastic response of the oceanic plate bending down into a subduction zone. The landward or island-arc slope of the trench is often interrupted by a submarine ridge, which sometimes breaks the ocean surface, as in the case of the Java Trench. Such a ridge is constructed from deformed sediments scraped off the top of the descending oceanic plate and is termed an accretionary prism. A line of explosive volcanoes, extruding (erupting) a lava that forms the volcanic rock andesite, is found on the overriding plate usually 100 kilometres or so from the trench. In marginal trenches these volcanoes form mountain chains, such as the Cascades in the Pacific Northwest or the great volcanoes of the Andes. In island arcs they form active volcanic island chains, such as the Mariana Islands.
Behind the volcanic line of island arcs are sometimes found young, narrow ocean basins. These basins are bounded on the opposite side by submarine ridges. Such interarc, or backarc, basins are sites of seafloor spreading directly caused by the dynamics of subduction. They originate at the volcanic line, so that the outer bounding submarine ridge, or third arc, represents an older portion of the volcanic line that has spread away. These backarc basins bear many of the features characteristic of oceanic spreading centres. Well-studied examples of these features are found in the Lau Basin of the Tonga arc and also west of the Mariana Islands. The Sea of Japan originated from backarc spreading behind the Japanese arc that began some 30 million years ago. At least two backarc basins have opened behind the Mariana arc, creating seafloor in two phases from about 30 to 17 million years ago in the western Parece Vela Basin and from 5 million years ago in the Mariana Trough next to the islands.
Aseismic ridges
In some oceans the basin floors are crossed by long, linear and mountainous aseismic ridges. The term aseismic distinguishes these ridges from oceanic spreading centres because the former lack earthquakes. Most aseismic ridges are constructed by volcanism from a hot spot and are composed of coalescing volcanoes of various sizes. A hot spot is a magma-generating centre fixed in the Earth's deep mantle and leaves a trail of volcanic outpourings on the seafloor as an oceanic plate travels over it. This form of volcanism is not associated with the volcanism at spreading centres and is distinct from it chemically in that the magma extruded onto the surface has a higher alkali composition. (For additional information on hot spots, see volcano: Volcanism and tectonic activity: Intraplate volcanism.)
The Hawaiian-Emperor chain is the best displayed aseismic ridge. Earthquakes do occur here, but only at the end of the ridge where volcanism is current--in this case, on the island of Hawaii (commonly known as the Big Island) to the southeast end of the island chain. Taking into account the relief of the island of Hawaii above the seafloor, it is the largest volcanic edifice on Earth. The Hawaiian-Emperor chain stretches from the Big Island to the intersection of the Kuril and Aleutian trenches in the northwest Pacific.
There are roughly 18 volcanoes or seamounts per 1,000 kilometres along the Hawaiian segment and 13 per 1,000 kilometres on the Emperor portion beyond the bend. The Hawaiian Islands are a part of the chain--the young part--that rises above sea level.
The Hawaiian-Emperor chain has two main trends:
(2) from this point to the Aleutian Trench, the trend of the Emperor segment is north-northwest.
Other prominent aseismic ridges include the Ninetyeast Ridge and the Chagos-Laccadive Plateau in the Indian Ocean and the Walvis Ridge and Rio Grande Rise in the South Atlantic. The Ninetyeast Ridge is thought to have originated from hot spot volcanic activity now located at the Kerguelen Islands near Antarctica. These islands lie atop the Kerguelen Plateau, which also originated from volcanism at this hot spot. The Ninetyeast Ridge stretches parallel to 90 E longitude in a long, linear chain of seamounts and volcanic ridges from the Andaman Islands in the Bay of Bengal more than 4,500 kilometres to the south where it intersects Broken Ridge at 30 S latitude. Broken Ridge is an aseismic ridge and was once part of the Kerguelen Plateau. It was split away from the plateau as Australia separated from Antarctica.
Core samples of the seafloor along the Ninetyeast Ridge have been retrieved through deep-sea drilling. Analyses of the samples show that the ridge is slightly less than 30 million years old in the south and about 80 million years old in the north. Additionally, sediments on the ridge indicate that parts of it were above sea level while it was being built near a spreading centre. The ridge then subsided as it rode north on the Indian Plate.
The Walvis Ridge and Rio Grande Rise originated from hot spot volcanism now occurring at the islands of Tristan da Cunha 300 kilometres east of the crest of the Mid-Atlantic Ridge. The Walvis Ridge trends northeast from this location to the African margin. The Rio Grande Rise trends roughly southeast from the South American margin toward the Mid-Atlantic Ridge. Both the Walvis Ridge and Rio Grande Rise began forming from the same hot spot near the spreading centre as the South Atlantic was in its initial opening stages 100 to 80 million years ago. The spreading centre shifted west of the hot spot about 80 million years ago, ending construction of the Rio Grande Rise but continuing to build the Walvis Ridge. Volcanic activity has since diminished, resulting in the younger part of the latter ridge being smaller. The findings of ocean drilling on the Rio Grande Rise show that it was once a volcanic island some two kilometres high.
Seamounts are submarine volcanoes with more than 1,000 metres of relief. Aseismic ridges are built by chains of overlapping seamounts. A seamount is akin to a subaerial shield volcano in that it also has gently sloping sides (5 to 15) and is constructed by nonexplosive eruptions of alkaline basalt lavas that are thought to originate from depths of roughly 150 kilometres. About 2,000 seamounts are known; they are most common in the Pacific and on fast-spreading ridges. Like the Hawaiian-Emperor chain, the lines of seamounts and islands trending northwest-southeast in the central and south Pacific (Marshall Islands, Line Islands, Tuamotu Archipelago, and Cook and Austral Islands) may be due to hot spot volcanism. Isolated seamounts also occur, and many of these are located in the far western Pacific. Another group of smaller seamounts is found in the northeastern Pacific.
Flat-topped seamounts are called guyots. They are particularly abundant in the western Pacific and along the Emperor seamount chain. Bottom samples and drill cores of shallow-water sediments and fossils capping guyots have been retrieved. The presence of such geologic materials suggest that guyots are seamounts that have had their peaks planed off by wave action and have since subsided below sea level. The western Pacific guyots are capped by drowned coral atolls and reefs.
These reefs are generally of Late Cretaceous age (about 95 million years old). The cause of the subsidence is attributed to the sinking of the seafloor as it moves down the flanks of an oceanic ridge. However, the reason for the demise of the coral reefs on the Cretaceous guyots is less clear. Under normal conditions, coral growth can easily keep up with sinking due to seafloor spreading. The Cretaceous guyots may have resulted from the northward drift of seamounts and reefs on the Pacific Plate away from the tropical zone of favourable growth. Another hypothesis is that the reefs were killed by unusually anoxic (oxygen-depleted) conditions that developed suddenly, a situation possibly related to intense seafloor volcanism in the Pacific at this time.
Abyssal hills are low-relief (less than 1 kilometre) features usually 1 to 10 kilometres wide and elongate parallel to spreading centres or to marine magnetic anomalies located in the vicinity of the latter. The tops of the hills are often flat, in which case they have steep sides. Gently sloping sides, however, are equally common. Abyssal hills are extremely numerous, so much so that Menard declared them "the most widespread physiographic forms of the face of the earth." Abyssal hills are most common in the Pacific basin, where they cover 80 to 85 percent of the seafloor.
Because they cover the entire flanks and crests of the oceanic ridges, such hills are thought to form during crustal accretion at spreading centres. They are commonly associated with intermediate- and fast-spreading ridges. On slow-spreading ridges, such as the Mid-Atlantic, the topographic features are much larger and have steeper sides. Bottom-sampling and seismic reflection studies reveal that abyssal hills are relief features on the top of the oceanic crust; they are not constructed from ocean-bottom sediments. In areas such as the abyssal plains (see below), abyssal hills are buried by sediments.
Apparently the hills are constructed by two processes: volcanism and block faulting. The relative contribution of each may depend on the spreading rate. At slower rates, faulting of the oceanic crust is a dominant factor in forming the relief, and the relief of the hills is greater as the rate is slower. At the crest of a spreading centre, volcanism in the neovolcanic zone initiates the construction of volcanic hills. The zone of active faulting is where they form or are modified by block faulting. The existence of discrete and separate volcanic hills indicates that volcanism at a spreading centre is episodic.
The ocean basin floor is everywhere covered by sediments of different types and origins. The only exception are the crests of the spreading centres where new ocean floor has not existed long enough to accumulate a sediment cover. Sediment thickness in the oceans averages about 450 metres. The sediment cover in the Pacific basin ranges from 300 to 600 metres thick, and that in the Atlantic is about 1,000 metres. Generally, the thickness of sediment on the oceanic crust increases with the age of the crust. Oceanic crust adjacent to the continents can be deeply buried by several kilometres of sediment. Deep-sea sediments can reveal much about the last 200 million years of Earth history, including seafloor spreading, the history of ocean life, the behaviour of the Earth's magnetic field, and the changes in the ocean currents and climate.
The study of ocean sediments has been accomplished by several means. Bottom samplers, such as dredges and cores up to 30 metres long, have been lowered from ships by wire to retrieve samples of the upper sediment layers. Deep-sea drilling has retrieved core samples from the entire sediment layer in several hundred locations in the ocean basins. The seismic reflection method has been used to map the thickness of sediments in many parts of the oceans. Besides thickness, seismic reflection data can often reveal sediment type and the processes of sedimentation. (For more information on the equipment and techniques used by investigators to study deep-sea sediments, see undersea exploration.)
Sediment Types
Deep-sea sediments can be classified as terrigenous, originating from land; biogenic, consisting largely of the skeletal debris of microorganisms; and authigenic, formed in place on the seafloor. Pelagic sediments, either terrigenous or biogenic, are those that are deposited very slowly in the open ocean either by settling through the volume of oceanic water or by precipitation. The sinking rates of pelagic sediment grains are extremely slow because they ordinarily are no larger than several micrometres. However, fine particles are normally bundled into fecal pellets by zooplankton, which allows sinking at a rate of 40 to 400 metres per day.
Terrigenous sediments
Terrigenous sediments are transported to the oceans by rivers and wind. The sediments that reach the continental shelf are often stored in submarine canyons on the continental slope. Turbidity currents carry these sediments down into the deep sea (see above Density currents in the oceans: Turbidity currents). These currents create sedimentary deposits called turbidites, which are layers up to several metres thick composed of sediment particles that grade upward from coarser to finer sizes. The turbidites build sedimentary deep-sea fans adjacent to the base of the continental slope. Turbidites also are found below the major river deltas of the world where they build features called abyssal cones. The largest of these is the Ganges Fan (also called the Ganges Cone or Bengal Cone) in the Bay of Bengal east of the Indian subcontinent. It measures 3,000 kilometres long (north-south) by 1,000 kilometres wide (east-west) and is up to 12 kilometres thick. The Bengal Cone is forming from rock material eroded from the Himalayas and transported by the Ganges and Brahmaputra rivers.
Abyssal plains are formed by the accumulation of turbidites beyond the limits of deep-sea fans and abyssal cones in locations where there is a very large sediment supply. In contrast to fans and cones, abyssal plains are flat and featureless. They are prominent near both margins of the Atlantic and in the northeast Pacific. Tectonic and climate controls have influenced the formation of abyssal plains. The last major glaciation near the end of the Pleistocene epoch about 10,000 years ago greatly increased erosion and sediment supply to the deep sea, but deep-sea trenches interrupted the flow of turbidity currents to the ocean floor. Off the Pacific Northwest coast of the United States, however, the trenches were filled by turbidites, and subsequent turbidity currents passed beyond them to form the Alaska and Tufts abyssal plains.
Brown clays are a variety of pelagic sediment, mostly of terrigenous origin, which are composed largely of four different clay minerals: chlorite, illite, kaolinite, and montmorillonite. By definition, clays have less than 30 percent biogenic components. Quartz, volcanic ash, and micrometeorites are common as minor constituents. Brown clays are widespread in the deeper areas of the oceans below four kilometres.
They dominate the floor of the central North Pacific. Clays accumulate very slowly, averaging about one millimetre per 1,000 years. The type of clay found in a given area is a function of the source region on land and the climate. For example, chlorite is dominant in polar regions and kaolinite in the tropics. Clays are introduced into the oceans by river transport, although kaolinite is also carried by the wind from the arid regions of Africa and Australia. Montmorillonite is an alteration product of volcanic material and can form from either wind-blown volcanic ash or basaltic glass on the seafloor.
Sediments composed mostly or entirely of volcanic ash are commonly found adjacent to the island arcs and marginal trenches. These are normally deposited as turbidites. Volcanic ash that has been ejected higher than five kilometres during an eruption can be carried by wind and settle out through the atmosphere and oceans as pelagic sediment. The ocean floor encircling Antarctica is covered by glacial marine sediments. These sediments are carried by icebergs from the continent as far north as the Antarctic Convergence at 45 to 55 latitude.
Biogenic oozes are pelagic sediments that have more than 30 percent skeletal material. They can be either carbonate (or calcareous) ooze or siliceous ooze. The skeletal material in carbonate oozes is calcium carbonate usually in the form of the mineral calcite but sometimes aragonite. The most common contributors to the skeletal debris are such microorganisms as foraminiferans and coccoliths, microscopic carbonate plates that coat certain species of marine algae and protozoa. Siliceous oozes are composed of opal (amorphous, hydrated silica) that forms the skeleton of various microorganisms, including diatoms, radiolarians, siliceous sponges, and silicoflagellates. The distribution of biogenic oozes depends mainly on the supply of skeletal material, dissolution of the skeletons, and dilution by other sediment types, such as turbidites or clays.
Primary productivity in the ocean surface waters controls supply to a large extent. Productivity is high at the equator and in zones of coastal upwelling and also where oceanic divergences occur near Antarctica. Productivity is lowest in the central areas of the oceans (the gyres) in both hemispheres. Siliceous oozes are more reliable indicators of high productivity than carbonate oozes. This is because silica dissolves quickly in surface waters and carbonate dissolves in deep water; hence, high surface productivity is required to supply siliceous skeletons to the ocean floor. Carbonate oozes dominate the deep Atlantic seafloor, while siliceous oozes are most common in the Pacific; the floor of the Indian Ocean is covered by a combination of the two.
Carbonate oozes cover about half of the world's seafloor. They are present chiefly above a depth of 4,500 metres; below that they dissolve quickly. This depth is named the Calcite Compensation Depth (or CCD). It represents the level at which the rate of carbonate accumulation equals the rate of carbonate dissolution. In the Atlantic basin the CCD is 500 metres deeper than in the Pacific basin, reflecting both a high rate of supply and low rate of dissolution in comparison to the Pacific. The input of carbonate to the ocean is through rivers and deep-sea hydrothermal vents. Variation in input, productivity, and dissolution rates in the geologic past have caused the CCD to vary over 2,000 metres. The CCD intersects the flanks of the world's oceanic ridges, and as a result these are mostly blanketed by carbonate oozes.
Siliceous oozes predominate in two places in the oceans: around Antarctica and a few degrees of latitude north and south of the equator. At high latitudes the oozes include mostly the shells of diatoms. South of the Antarctic Convergence diatom oozes dominate the seafloor sediment cover and mix with glacial marine sediments closer to the continent.
Seventy-five percent of all the oceans' silica supply is being deposited in the area surrounding Antarctica. Radiolarian oozes are more common near the equator in the Pacific. Here, both siliceous oozes and calcareous oozes occur, but carbonate deposition dominates the region immediately near the equator. Siliceous oozes bracket the carbonate belt and blend with pelagic clays farther north and south. Because siliceous skeletons dissolve so quickly in seawater, only the more robust skeletal remains are found in the siliceous oozes. Thus, fossils of this kind are not completely representative of the organisms living in the waters above.
Authigenic sediments
The most significant authigenic sediments in the ocean basins today are metal-rich sediments and manganese nodules. Metal-rich sediments include those enriched by iron, manganese, copper, chromium, and lead. These sediments are common at spreading centres, indicating that processes at the centres are responsible for their formation--specifically, hydrothermal circulation is the controlling factor.
Deep-sea drill cores have revealed the presence of metal-rich sediments on top of ancient oceanic crust away from ridge crests. It can be inferred from this that the processes controlling their formation existed in the past, but with variations. Which type of enriched sediment is deposited depends on the degree of mixing between the hydrothermal water deep in the crust at a spreading centre and the cold seawater percolating down into the crust. Little mixing produces sulfides, liberal mixing yields manganese-rich crustal material, and intermediate conditions give rise to sediments enriched in iron and manganese.
Manganese nodules are pebbles or stones about the size of walnuts that are built of onionlike layers of manganese and iron oxides. Minor constituents include copper, nickel, and cobalt, making the nodules a potential ore of these valuable elements. Mining of manganese nodules has been the subject of study and experimentation since the 1950s.
The nodules grow very slowly, about one to four millimetres per million years. They are found in areas of slow sedimentation, usually five millimetres per thousand years or less. The North and South Pacific hold the greatest concentration of manganese nodules; in some places, the nodules cover 90 percent of the surface of the ocean floor. Coverages this high also are found in the southernmost South Atlantic. The Indian Ocean floor is largely devoid of manganese nodules. Because seawater is supersaturated in manganese, the direct precipitation of the element onto an available surface is the most likely mode of nodule formation.
Two significant mysteries surround manganese nodules. Drilling and coring in the sediment column has shown that nodules are vastly more abundant at the seafloor than below it and that the rate of growth of nodules is 10 times slower than the lowest known sedimentation rates. If such is the case, the nodules should be quickly buried and should be common in the sediment below the seafloor. Current theories for explaining these observations propose that bottom currents keep areas of nodule growth free of sediment deposition and that burrowing organisms nudge and roll the nodules in the process of feeding, thereby keeping them at the surface of the seafloor. Observations in the deep sea support both explanations.
Sedimentation Patterns
The patterns of sedimentation in the ocean basins have not been static over geologic time. The existing basins, no more than 200 million years old, contain a highly variable sedimentary record. The major factor behind the variations is plate movements and related changes in climate and ocean water circulation. Since about 200 million years ago, a single vast ocean basin has given way to five or six smaller ones. The Pacific Ocean basin has shrunk, while the North and South Atlantic basins have been created.
The climate has changed from warm and mild to cool, stormy, and glacial. Plate movements have altered the course of surface and deep ocean currents and changed the patterns of upwelling, productivity, and biogenic sedimentation. Seaways have opened and closed. The Strait of Gibraltar, for example, was closed off about 6 million years ago, allowing the entire Mediterranean Sea to evaporate and leave thick salt deposits on its floor. Changes in seafloor spreading rates and glaciations have caused sea level to rise and fall, greatly altering the deep-sea sedimentation pattern of both terrigenous and biogenic sediments. The CCD has fluctuated more than 2,000 metres in response to changes in carbonate supply and the corrosive nature of ocean bottom waters.
Bottom currents have changed, becoming erosive or nondepositional in some regions to produce unconformities and redistributing enormous volumes of sediment to other locations. The Pacific Plate has been steadily moving northward, so that biogenic sediments of the equatorial regions are found in drill cores taken in the barren North Pacific.
Through most of geologic time, probably extending back 2 billion years, the ocean basins have both grown and been consumed as plate tectonics continued on Earth. The latest phase of ocean basin growth began just less than 200 million years ago with the breakup of the supercontinent Pangaea, the enormous landmass composed of nearly all the present-day continents. Since that time, the major developments have included a shrinking of the Pacific basin at the expense of the growing Atlantic and Arctic basins, the opening of the Tethys seaway circling the globe in tropical latitudes and its subsequent closing, and the opening of the Southern Ocean (see above General considerations) as the southern continents moved north away from Antarctica.
As was noted earlier, the oldest known oceanic crust (estimated to be about 200 million years old) is located in the far western equatorial Pacific, east of the Mariana Island arc. The Pacific ocean floor at this site was generated during seafloor spreading from a pattern of ridges and plates that had existed for some unknown period of time. At least five different seafloor spreading centres were involved.
In the Indian Ocean the oldest segment of seafloor was formed about 165 to 145 million years ago by the rifting away of Africa and South America from Gondwana, a supercontinent consisting largely of the present-day continents of the Southern Hemisphere. At this time, Africa was joined to South America, Eurasia, and North America. Today, this old seafloor is found along the east coast of Africa from the Somali Basin to the east coast of South Africa and adjacent to Queen Maud Land and Enderby Land in East Antarctica.
Close to 180 million years ago (but before 165 million years ago), North America and Eurasia, which together made up most of the large northern continent of Laurasia, began drifting away from Africa and South America, creating the first seafloor in the central region of the North Atlantic and opening the Gulf of Mexico. The Tethys seaway also opened during this rifting phase as Europe pulled away from Africa.
Shortly after this time continental fragments, including possibly Tibet, Myanmar (Burma), and Malaya, rifted away from the northwest coast of Australia and moved northward, thereby creating the oldest seafloor in the Timor Sea. During this period spreading continued in the Pacific basin with the growth of the Pacific Plate and the consumption by subduction of its bordering plates, including the Izanagi, Farallon, and Phoenix. The Pacific Plate moved northward during this phase and continues to do so today.
India and Madagascar, as a unit, rifted away from Australia and Antarctica prior to 130 million years ago and began drifting northward, creating seafloor adjacent to Western Australia and East Antarctica. Possibly simultaneously or shortly after this rifting began, South America started to separate from Africa, initiating the formation of seafloor in the South Atlantic Ocean.
Between 90 and 80 million years ago, Madagascar and India separated, and the spreading ridges in the Indian Ocean were reorganized. India began drifting northward directly toward Asia. During this same period Europe, joined to Greenland, began drifting away from North America, which resulted in the emergence of the seafloor in the Labrador Sea and the northernmost Atlantic Ocean. This spreading phase affected the passages in the Tethys seaway between Europe (Iberia) and northwest Africa, intermittently opening and closing it. In the southwest Pacific, New Zealand, along with the Lord Howe Rise and the Norfolk Ridge, rifted away from Australia and Antarctica between 80 and 60 million years ago, opening the Tasman Sea.
About 60 million years ago a new rift and oceanic ridge formed between Greenland and Europe, separating them and initiating the formation of oceanic crust in the Norwegian Sea and the Eurasian basin in the eastern Arctic Ocean. The Amerasian basin in the western Arctic Ocean had formed during an earlier spreading phase from about 130 to 110 million years ago.
Between 60 and 50 million years ago, significant events occurred in the Indian Ocean and southwest Pacific. Australia began drifting northward, away from East Antarctica, creating seafloor there. The northward movement of Australia resulted in the emergence of several subduction zones and island arcs in the southwest and equatorial Pacific. The Indian subcontinent first touched against the Asian continent about 53 million years ago, developing structures that preceded the main Himalayan orogeny (mountain-building event), which began in earnest some 40 million years ago.
Less than 30 million years ago, seafloor spreading ceased in the Labrador Sea. Along the west coast of North America, the Pacific Plate and the North American Plate converged along what is now California shortly after 30 million years ago. This resulted in the cessation of a long history of subduction in the area and the gradual conversion of this continental margin to a transform fault zone. Continued closure between Africa and Europe, which began about 100 million years ago, caused the isolation of the Mediterranean Sea, so that by 6 million years ago it had completely evaporated.
The present-day Mediterranean seafloor was formed during a complex sequence of rifting between small plates in this region, beginning with the separation of North America and Europe from Africa about 200 million years ago. In the eastern Mediterranean, the seafloor is no older than about 100 million years. West of Italy it was created during subsequent spreading between 30 and 20 million years ago.
The Caribbean Sea and the Gulf of Mexico formed as a result of the relative movement between North America and South America. The seafloor of the Gulf of Mexico began forming some 160 to 150 million years ago. A proto- or ancient Caribbean seafloor also was formed during this period but was later subducted. The present Caribbean seafloor consists of a captured piece of the Farallon Plate (from the Pacific basin) and is estimated to be for the most part of Cretaceous age (i.e., about 120 to 85 million years old).
The seafloor in the western portion of the Philippine Sea developed between 60 and 35 million years ago. In the east, it was formed by backarc spreading from 30 million years ago (see above). The origin of the older crust is not completely clear. It either was created by spreading in the Pacific basin and subsequent capture by the formation of the Bonin and Mariana arcs, or it resulted from backarc spreading behind trenches to the south.
Through knowledge of the ocean sedimentary record, the history of plate motions, glacial changes, and established relations between present sedimentation patterns and environmental factors, it is possible to reconstruct an oceanographic history for approximately the past 200 million years. This is the emerging field of paleoceanography.
Prior to the breakup of Pangaea, one enormous ocean, Panthalassa, existed on Earth. Currents in this ocean would have been simple and slow, and the Earth's climate was, in all likelihood, warmer than today. The Tethys seaway formed as Pangaea broke into Gondwana and Laurasia (see above). In the narrow ocean basins of the central North Atlantic, restricted ocean circulation favoured deposition of evaporites (halite, gypsum, anhydrite, and other less abundant salts). Evaporites also were deposited some 100 million years ago in the equatorial regions of the South Atlantic during the early opening of this ocean.
Sequences of organic-rich, black shales were deposited during the early phases of spreading in the North and South Atlantic. These sediments indicate anoxic conditions in the deep ocean waters. The oceans must have been well stratified into dense layers to prevent the overturning and mixing required to replace depleted oxygen. Black shales also were deposited in the older areas of the eastern Indian Ocean.
During the time interval between 200 and 65 million years ago, but especially from 100 to 65 million years ago, microplankton abundance and diversity increased enormously in the oceans. This resulted in increased deposition of biogenic sediments in the ocean basin.
During Cretaceous time (from 144 to 66.4 million years ago), sea level was often high, and shallow seas lapped onto the continents. This may have provided an environment favourable to the explosion in the numbers of species of foraminiferans, diatoms, and calcareous nannoplankton. Increased abundance of calcareous nannoplankton shifted the locus of carbonate sedimentation from shallow seas to the deep ocean. The end of Cretaceous time is marked by a sudden extinction of many life-forms on Earth, and marine organisms were no exception. Coccolithophores (calcareous nannoplankton) and planktonic foraminiferans were particularly affected, and only a few species survived. Ocean sediments were suddenly less biogenic, and clays became widespread.
After Cretaceous time the Earth underwent a gradual cooling, especially at high latitudes. Deep-sea sedimentation changed as thermohaline bottom water circulation became fully developed (see above Circulation of the ocean waters: Thermohaline circulation). The CCD rose in the Pacific and dropped in the Atlantic as a result of changes in thermohaline circulation. An event of major significance was the spreading away of Australia from Antarctica beginning about 53 million years ago. This separation initiated limited circum-Antarctic circulation, which isolated Antarctica from the warmer oceans to the north, and led to cooling, which set the stage for later major glaciation.
At the Eocene-Oligocene boundary (36.6 million years ago), Antarctic Bottom Water began to form, resulting in greatly decreased bottom-water temperatures in both the Pacific and Atlantic oceans. Bottom-living organisms were strongly affected, and the CCD suddenly dropped from about 3,500 metres to approximately 4,000 to 5,000 metres in the Pacific.
Bottom-water temperatures were generally warm, 12 to 15 C, during the time preceding this event. In a study of deep-sea sediment core material from near Antarctica, J.P. Kennett and Lowell D. Stott of the United States discovered that there was a period between roughly 50 and 35 million years ago when deep waters were very warm (20 C) and salty. The origin of these ocean waters was most likely in the low latitudes and resulted from high evaporation rates there.
The modern oceans are distinguished by very cold bottom water. The gradual changes toward this condition began 10 million years after the origination of Antarctic Bottom Water. Particularly significant among these changes was the closing of the Tethys seaway as Australia and several microcontinents moved north into the Indonesian region. Also, Australia moved far enough north that circum-Antarctic surface circulation became fully established.
The modern ocean circulation patterns and basin shapes were mostly in place by the beginning of Miocene time (nearly 24 million years ago). An exception was an ocean connection between the Pacific and Caribbean Sea in Central America that persisted until about 3 million years ago. Major and probably permanent ice sheets on Antarctica formed during Miocene time, and glacial sediments began to dominate the seafloor surrounding the continent shortly thereafter. Siliceous oozes also became widespread around Antarctica.
increased in this area at the expense of siliceous sedimentation in equatorial regions. Ocean circulation became more vigorous, global climate became cooler, and sedimentation rates in the ocean basins increased. Planktonic microorganisms were segregated into latitudinal belts. Bottom-water flow north through the Drake Passage between South America and Antarctica began in Miocene time, resulting in erosion and nondeposition of sediments in the southwest Atlantic and southeast Pacific oceans. Also during Miocene time rifting between Greenland and Europe had progressed to a point where a connection was established between the North Atlantic and the Norwegian Sea.
This resulted in the formation of North Atlantic Deep Water (see above Circulation of the ocean waters: Thermohaline circulation), which began flowing south along the continental rise of North America at this time. Sediments redistributed and deposited by this deep current are called contourites and have been extensively studied by Bruce Heezen, Charles D. Hollister, and Brian E. Tucholke, among others.
Sudden global cooling set in near the end of the Miocene some 6 million years ago. The strength of ocean circulation must have increased, as evidence of increased upwelling and biological productivity is present in ocean sediments. Diatomaceous sediments were deposited in abundance around the rim of the Pacific.
This cooling event is synchronous with a drop in sea level, thought to be about 40 or 50 metres by various authorities, and probably corresponds to the further growth of the Antarctic ice sheet. This lowered sea level, coupled with the closure of narrow seaways probably due to plate movements, isolated the Mediterranean Sea. Subsequently, the sea dried up, leaving evaporite deposits on its floor. The Swiss geologist Kenneth J. Hsüand the American oceanographer William B.F. Ryan have concluded that the Mediterranean probably dried up about 40 times as seaways opened and closed between 6 and 5 million years ago. This evaporation removed about 6 percent of the salt from the world ocean, which raised the freezing point of seawater and promoted further growth of the sea ice surrounding Antarctica.
Enormous ice sheets emerged in the Northern Hemisphere between 3 and 2 million years ago, and the succession of Quaternary glaciations began at 1.6 million years ago. The exact cause of the glacial period is unclear, but it is most likely related to the variability in solar isolation, increased mountain building, and an intensification of the Gulf Stream at 3 million years ago due to the closing off of the Pacific-Caribbean ocean connection in Central America.
The Quaternary glaciations, of which there were probably 30 episodes, left the most dramatic record in ocean sediments of any event in the previous 200 million years. Terrigenous sedimentation rates greatly increased in response to fluctuations in sea level of up to 100 metres and a more extreme climate. Biogenic sedimentation also increased and fluctuated with the glacial episodes. Deep-sea erosion began in many places as a result of intensified bottom-water circulation.
Deep-sea sediments can be classified as terrigenous, originating from land; biogenic, consisting largely of the skeletal debris of microorganisms; and authigenic, formed in place on the seafloor. Pelagic sediments, either terrigenous or biogenic, are those that are deposited very slowly in the open ocean either by settling through the volume of oceanic water or by precipitation. The sinking rates of pelagic sediment grains are extremely slow because they ordinarily are no larger than several micrometres. However, fine particles are normally bundled into fecal pellets by zooplankton, which allows sinking at a rate of 40 to 400 metres per day.
Terrigenous sediments are transported to the oceans by rivers and wind. The sediments that reach the continental shelf are often stored in submarine canyons on the continental slope. Turbidity currents carry these sediments down into the deep sea. These currents create sedimentary deposits called turbidites, which are layers up to several metres thick composed of sediment particles that grade upward from coarser to finer sizes. The turbidites build sedimentary deep-sea fans adjacent to the base of the continental slope. Turbidites also are found below the major river deltas of the world where they build features called abyssal cones. The largest of these is the Ganges Fan (also called the Ganges Cone or Bengal Cone) in the Bay of Bengal east of the Indian subcontinent. It measures 3,000 kilometres long (north-south) by 1,000 kilometres wide (east-west) and is up to 12 kilometres thick. The Bengal Cone is forming from rock material eroded from the Himalayas and transported by the Ganges and Brahmaputra rivers.
Abyssal plains are formed by the accumulation of turbidites beyond the limits of deep-sea fans and abyssal cones in locations where there is a very large sediment supply. In contrast to fans and cones, abyssal plains are flat and featureless. They are prominent near both margins of the Atlantic and in the northeast Pacific.
Tectonic and climate controls have influenced the formation of abyssal plains. The last major glaciation near the end of the Pleistocene epoch about 10,000 years ago greatly increased erosion and sediment supply to the deep sea, but deep-sea trenches interrupted the flow of turbidity currents to the ocean floor. Off the Pacific Northwest coast of the United States, however, the trenches were filled by turbidites, and subsequent turbidity currents passed beyond them to form the Alaska and Tufts abyssal plains.
Brown clays are a variety of pelagic sediment, mostly of terrigenous origin, which are composed largely of four different clay minerals: chlorite, illite, kaolinite, and montmorillonite. By definition, clays have less than 30 percent biogenic components. Quartz, volcanic ash, and micrometeorites are common as minor constituents. Brown clays are widespread in the deeper areas of the oceans below four kilometres.
They dominate the floor of the central North Pacific. Clays accumulate very slowly, averaging about one millimetre per 1,000 years. The type of clay found in a given area is a function of the source region on land and the climate. For example, chlorite is dominant in polar regions and kaolinite in the tropics. Clays are introduced into the oceans by river transport, although kaolinite is also carried by the wind from the arid regions of Africa and Australia. Montmorillonite is an alteration product of volcanic material and can form from either wind-blown volcanic ash or basaltic glass on the seafloor.
Sediments composed mostly or entirely of volcanic ash are commonly found adjacent to the island arcs and marginal trenches. These are normally deposited as turbidites. Volcanic ash that has been ejected higher than five kilometres during an eruption can be carried by wind and settle out through the atmosphere and oceans as pelagic sediment. The ocean floor encircling Antarctica is covered by glacial marine sediments. These sediments are carried by icebergs from the continent as far north as the Antarctic Convergence at 45 to 55 latitude.
Oceanography also called oceanology or marine science, is the branch of Earth Sciences that studies the ocean. It covers a wide range of topics, including marine organisms and ecosystem dynamics; ocean currents, waves, and geophysical fluid dynamics; plate tectonics and the geology of the sea floor; and fluxes of various chemical substances and physical properties within the ocean and across its boundaries. These diverse topics reflect multiple disciplines that oceanographers blend to further knowledge of the world ocean and understanding of processes within it: biology, chemistry, geology, meteorology, and physics.
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