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S330 Oceanography

Block 1: The Ocean Basins: Their Structure And Evolution


Sea-floor Mapping Technology


Red Sea - Juvenile

Research Beneath the Sea


Spreading Ocean Ridges

Global Plate Tectonics



Depth measurement
Bathymetry from satellites


Aseismic continental margins
Seismic continental margins and island arcs
Ridge topography
Age-depth relationships across ridges
Abyssal plains
The distribution of submarine volcanoes
Aseismic ridges


The Red Sea
The Mediterranean


Pillow lavas: the top of the oceanic crust
Why a median valley?
Formation of the volcanic layer: two case studies
Second- and third-order segmentation of fast-spreading axes
Second- and third-order segmentation of slow-spreading axes
A plausible model for lithospheric growth
Changes in spreading pattern
Crustal abnormalities


Heat flow, convection and permeability
Changes in the rocks
Changes in seawater
Variability in hydrothermal systems
Black smokers, white smokets and warm-water vents
The lifetimes of hydrothermal systems
Anatomy of a vent field
The biological significance of hydrothermal vent systems
Hydrothermal plumes
Event plumes
Off-axis hydrothermal circulation
The extent of hydrothermal meamorphism


Sediments and palaeoceanography
Different time-scales in sea-level changes
Using satellites to monitor sea-level changes
The post-glacial rise in sea-level
Measuring Quarternary changes in sea-level
The growth of an ice-sheet: Antarctica
The salinity crisis in the Mediterranean
The migration of climatic belts
The effect of plate-tectonic processes on sea-level
Major transgressions and regressions


Changes in components of the cycle
Some effects of short-term changes
The steady-state ocean


Back to OU

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Accurate navigation at sea became possible only with the development of accurate and reliable chronometers in the eighteenth century, which enabled longitude to be determined. Nowadays, there are many radio-navigational aids, both surface and satellite-mounted.

Accurate depth determinations became possible only with the development of echo-sounding early in this century. For detailed survey work, swath bathymetry is now favoured, using multiple narrow beams directed to either side of the ship's track.

Direct imaging of the sea-floor, as distinct from bathymetric mapping, is achieved by side-scan sonar instruments. These image strips of sea-floor as much as about 10km wide from a single ship track. Deep-tow devices allow narrower swaths to be imaged at higher resolution.

Satellite alimetry uses radar to measure very precisely the average height of the sea-surface, which follows a surface of equal gravitational potential - the geoid. After transient effects (tides, currents and atmospheric pressure changes) have been averaged out, the mean height of the sea-surface can be determined. Short-wavelength irregularities in the geoid (geoid anomalies) correspond to topographic features on the ocean floor, and the height of the sea-surface can therefore provide a measure of bathymetry. Imaging radar (SAR) can show up bathymetric features in shallow water.

Detailed exploration of oceanic crustal structure and composition began in the 1960s, with the DSDP, succeeded in 1985 by the ODP. Much has been learned with the help of new technologies for imaging and through submersible operations.


Aseismic (passive, Atlantic-type) continental mrgins develop where the continent and adjacent ocean bais belong to the same plate. They are underlain by stretched and thinned continental crust, upon which sediments accumulate to build up the continental crust, upon which sediments accumulate to build up the continental shelf (covered by shelf seas), slop and rise. Microcontinents are areas of continental crust split off from larger continental masses, which form either islands (crust of near-normal thickness) or submerged plateaux (thinned crust).

Seismic (active, Pacific-type) continental margins occur where a continent and the adjacent ocean basin belong to different plates, and oceanic lithosphere is being subducted beneath continental lithosphere. These are destructive plate margins. The continental shelf is typically narrower and the slope typically steeper than at aseismic margins, and a trench at the foot of the continental slope generally replaces the continental rise found at aseismic margins. Sediments accumulating in the trenches are partly scraped off to build the inner trench wall oceanwards, and partly subducted. Island arcs occur at another kind of seismic (destructive) margin. These are formed of volcanoes that lie above a subduction zone where one oceanic plate descends beneath another. They are usually backed by small basins., floored by oceanic crust that has formed by spreading somehow related to the nearby subduction.

Ocean ridges (spreading axes) are the most important physiographic (bathymetric) features of the ocean basins. They are the constructive margins of the plates, where new oceanic lithosphere is continually being generated. Gradients on the flanks of slow-spreading ridges are steeper than those on fast-spreading ridges. Median rift valleys are better developed (wiser and deeper) on slow-spreading than on fast-spreading ridges.

There is a systematic relationship between the depth to the top of the oceanic crust and its age of formation. This is the result of progressive cooling and subsidence of lithosphere with distance from the constructive margin. This relationship enables the age of oceanic crust to be estimated from its depth, and vice versa, and works to a good first approximation throughout the major ocean basins.

Spreading axes are offset by transform faults that lie along arcs of small circles about the pole of relative rotation of lithospheric plates. Transform faults are seismically active, and they separate plates moving in opposite directions. Beyond the region of offset, they become fracture zones, and lie within a single plate, so seismic activity is much less. These faults and fractures result in scarps and clefts on the ocean floor. Major transform faults are also known as conservative plate mrgins, but some transform faults can have a component either of spreading (leaky transforms), or of subduction. However, the predominant sense of movement is lateral. Minor offsets (less than about 10 km) of spreading axes are not usually true transform faults, and are described as non-transform offsets.

Abyssal plains occupy large areas of deep ocean floor. They are very flat as a result of burial of the rough topography of the oceanic crust by sediments. The sediments are either supplied by turbidity currents from adjacent aseismic margins (eg Atlantic Ocean), or deposited from suspension in seawater (pelagic sediment), especially where trenches bordering seismic margins have trapped continent-derived material, and prevented it from reaching the plain (eg much of the Pacific Ocean).

Seamounts, oceanic islands and aseismic ridges are volcanic features rising from the ocean floor. Linear chains of seamounts and islands, and aseismic ridges, are thought to result from hot-spot volcanism, whereby the oceanic plate moves over an intermittently or continually active fixed source of magma rising from the deep mantle.

Satellite altimetry measurements can be used to map features on the sea-floor, because the mean surface height correlates with ocean bathymetry. Conventional bathymetric observations can be refined and extended using these data.


Oceanic crust is much younger than most continental crust. Oceanic lithosphere must have been generated at spreading axes (ridges) and destroyed at subduction zones many times since the formation of the Earth. Ocean basins form initially by stretching and splitting (rifting) of continental crust, and the rise of mantle material and magma into the crack to form new oceanic lithosphere.

The Red Sea is an embryonic ocean that appears to be opening progressively from the south, where the axial region is underlain by oceanic crust and has a rift valley. Further north are isolated deeps - with metal-rich muds - but there is less evidence of oceanic crust in the axial region. The thick evaporites bordering the axial region rest on thinned continental crust.

Among the major ocean basins, the Atlantic has the simplest pattern of ocean-floor ages. Subduction is confined to relatively small island arc systems in the Caribbean and the extreme south-west. Successive stages in the shape of the Atlantic basin are therefore fairly easy to reconstruct, by moving the continents back along a direction at 90 deg to the magnetic anomaly stripes and parallel to the transform faults.

In contrast, both the Pacific and Indian Oceans (which have major subduction zones) are characterized by changes of spreadin grate and direction and the development of new spreading axes. Because of these complications, it is difficult to work out how the shapes of these ocean basins have changed with time. The occurrence in western North American of 'exotic terranes', which in some cases are believed to have originated as microcontinents in the south-west Pacific, make the task of such reconstruction even more complicated.

The Mediterranean represents an ocean in the final stages of its life cycle, contracting as Africa pushes northwards into Europe and western Asia. However, no in situ oceanic crust older than the Cretaceous is known from the Mediterranean basin, and crust as young as 2 Ma has been found there, demonstrating that there is no inverse correlation between age of an ocean basin and the vigour of sea-floor spreading and plate-tectonic activity.


In most places, oceanic crust has a seismically well-defined layered structure. Layer 1 consists mostly of sediments overlying the igneous crust of layers 2 and 3. On the simplest interpretation, layer 2 is volcanic, dominated by pillow lavas and other types of basaltic lave (above) and basaltic dykes (below), and layer 3 is gabbro and represents magma that has crystallized at depth. However, the layer 2/layer 3 boundary is sometimes found to represent an increase in the intensity ('grade') of metamorphism with depth. Layer 4 is the uppermost (lithospheric) mantle. The total thickness of normal igneous oceanic crust is about 7 km. In both chemical and mineralogical composition, vittually all the rocks of layers 2 and 3 are basaltic. Layer 4 is peridotite in composition. The principal seismic discontinuities in the oceanic crust are at the base of layer 1 and the top of layer 4 (the Moho). Seismic velocities generally increase with depth, and variations in the gradient of this velocity change enable subdivisions of the crust to be recognised.

Seismic tomography and seismic reflection studies have shown that spreading axes are underlain by zones of crystal mush, at a depth equivalent to layer 3. On fast-spreading axes, thin magma lenses have been identified at the top of this zone, but slow-spreading axes generally lack persistent magma bodies.

Several sophisticated techniques - including swath bathymetry, side-scan sonar, underwater photography, and submersible operations - have shown that volcanism in the median rift valley of ocean ridges is not continuous but episodic, at intervals of around 104-105 years. Lavas are erupted from volcanic abyssal hills and hummocky volcanic ridges a few kilometres long, rather than from continuous fissures.

Spreading axes are segmented on a variety of scales. Transorm faults define the ends of first-order segments, which may be hundreds of kilometres in length. Second-order segment boundaries are not transform faults, but are persistent enough to leave mappable bathymetric traces off-axis; on fast-spreading axes these are termed non-transform offsets. Third-order segments are bounded by deviationss from axial linearity (devals). Some third-order and all second- and first-order segment boundaries occur where there is a dimunution or cessation in magma supply. This occurs about every 10 km or so. Ridges are highest, and the crust thickest, at the midpoints of segments, where the volcanic activity is greatest. This falls off towards the ends of segments, where the ridge is topographically lower, and the crust thinner. The gabbro of layer 3 sometimes thins out to nothing close to fracture zones, and virtually the entire crust may be missing where anagmatic extension occurs.

As a model for axial processes, it seems that an elongate zone of less dense asthenosphere rising beneath denser lithosphere will form regularly spaced protrusions into the overlying layer. Each of thse protrusions feeds a crystal mush zone or small magma chamber beneath a segment of spreading axis (except for those spreading amagmatically). Each crystal mush zone or magma chamber may beed several volcanic vents in an episodic fashion.

Oceanic crust is thinner than normal in back-arc basins and thicker than normal in some other parts of the oceans, notably in large igneous provinces.

Fractures and faults can bring rocks of deeper crustal layers to the surface.

Seamounts and volcanic islands are formed by isolated submarine volcanoes, building up from the sea-bed. Only a minority reach the surface to form islands. Flat-topped seamounts (guyots) are mostly islands that have been planed off by wave action, but some may have formed with flattish tops due to volcanic processes. Dissolved gases cannot escape from lava at depths of more than about 500 m, so volcanic ash is rare in the deep oceans.


Direct evidence of hydrothermal circulation through oceanic crust was not obtained until the late 1970s, but it had been predicted from several lines of indirect evidence for at least a decade before. These included the distribution of metal-rich sediments at ridge crests, metamorphic rocks dredged from ridge crests, a major conductive heat flow deficit along the whole ocean-ridge system and laboratory experiments on basalt-seawater interactions at high temperatures and sea-floor pressure.

All hydrothermal systems require a permeable layer of rock which allows cold watr to percolate slowly downwards over a wide area, a localized heat source below the permeable layer, and channelways for the heated buoyant (low-density) plume of hot water to escape more rapidly at the surface above the heat source.

Rock metamorphism during hydrothermal activity involves hydration, the addition of large amounts of magnesium, the loss of calcium and sometimes sodium and potassium, and also gains and losses of numerous minor and trace elements. The normal basaltic mineral assemblage (plagioclase, pyroxene +/- olivine, and basaltic glass) is transformed into various combinations of albite, chlorite actinolite, zeolite, quartz and other minerals, depending on the temperature and pressure conditions. Serpentinite is formed when hot seawater penetrates to seismic layer 4 and hydrates the peridotite. It behaves plastically and is of low denisty, so it can be forced upwards along fractures by pressure. Eventually, it may be exposed at the sea-bed. Metamorphic rocks formed in oceanic crust may later be exposed along scarps of faults and fractures.

During hydrothermal circulation, seawater loses all its magnesium and sulphate (which is reduced to sulphide), and gains signficant amounts of calcium and sometimes also potassium and sodium, as well as several minor elements, especially silicon, but also others (eg barium, rubidium, iron and manganese). When hydrothermal solutions mix with normal seawater at vents, some of these constituents are precipitated, in cracks and fisures within the rock, around vents to build chimneys, and in the form of particles forming the 'smoke' at black smokers and white smokers. Sulphide is precipitated as metal sulphides (chiefly iron, but also notably of copper and zinc), calcium as calcium sulphate (anhydrite) and calcium carbonate (calcite), barium is precipitated as barium sulphate (barite), and silicon is precipitated as silica (including quartz). Manganese is precipitated as an oxide, as is some of the iron, though most of this goes into sulphides.

Sea-floor weathering occurs at bottom-water temperatures and involves hydration, alteration of feldspars and glass to clay minerals and (sometimes) zeolites, as well as oxidation, especially of iron and manganese to form oxide and hydroxide coatings.

Hydrothermal solutions are more acid and reducing than seawater, and carry sulphide ions (see item 4 above). There is a continuum of types of hydrothermal vent: black smokers, emerging at temperatures of about 350 C or above and precipitating mainly sulphide particles, which build vent chimneys and form black 'smoke' plumes; white smokers, with exit temperatures of about 30-330 C and precipitating mainly sulphate particles; and warm-water vents, with exit temperatures only a few degrees above normal bottom-water temperature. The latter two types are mixtures of normal seawater (which saturates the permeable upper crust), with the high-temperature solutions that elsewhere form black smokers.

Hydrothermal convection occurs throughout the ocean-ridge system and extends out to crust up to about 70 Ma old, with exponentially decreasing intensity away from ridge crests. Heat-loss calculations show that the equivalent of the whole ocean volume is cycled through oceanic crust in the space of ten million years or less, and the resulting fluxes of some elements into and out of the curst are comparable with (or may exceed) those of river transport to the oceans.

3He is a ubiquitous constituent of hydrothermal fluids, being released from the primordial 'store' within the Earth when the upper mantle partially melts below spreading axes. It is an ideal tracer, for it has no other natural sources in the oceans; it can be used to trace the movement of hydrothermal effluents and hence to help provide information about current patterns and the whereabouts of active vents. Other gases that have been found in hydrothermal plumes are methane (the most abundant dissolved gas), hydrogen, hydrogen sulphide, carbon dioxide, carbon monoxide, and nitrous oxide. They may be of abiotic (inorganic) or microbial origin.


Sediments of oceanic seismic layer 1 thicken from a few metres maximum at ridge crests to several kilometres at continental margins. Deep-sea pelagic sediments consist of clays and biogenic sediments formed of calcareous and/or siliceous skeletal remains of mostly planktonic organisms. Calcareous sediments tend to predominate on the flanks of ocean ridges, and siliceous sediments and clays accumulate in deeper parts of the ocean basins. Metalliferous sediments resulting from hydrothermal activity are believed to form the base of the sediment sequence nearly everywhere.

Sediments and the organisms preserved in them provide much information about the nature and timing of events in the evolution of ocean basins. For example, the spread of a species of the plankgonic foraminiferan genus Guembelitria in the Southern Ocean has helped to refine our knowledge of the history of continental separation in this region and the development of the Antarctic Circumpolar Current.

World-wide (eustatic) changes in sea-level are caused by changes in the total volume of water in the oceans, and by changes in the shape and volume of the ocean basins. Sea-level is presently rising world-wide, as a result of climate warming after the last glaciation, and is beginning to be exacerbated by warming caused by an enhanced greenhouse effect because of anthropogenic CO2 emissions into the atmosphere.

Isostatic effects such as sediment loading or rebound after disappearance of ice-sheets complicate the study of sea-level changes on a local scale. However, satellites can now provide methods for global monitoring of present-day changes in sea-level and sea-surface temperature, and the value of these methods will increase as longer time-series of data are accumulated.

Sea-level changes prior to the last glaciation can be reliably determined by measuring δ18O values in the carbonate skeletal remains of marine organisms. These values are higher when the ice-caps are large, because water vapour (and hence the snow which builds the ice-caps) is relatively enriched in H216O. When ice-caps are small and oceans are comparatively enriched in H216O and so δ18O values in marine organisms are lower. It is possible to calibrate δ18O values in terms of sea-level changes, and water temperatures. This has enabled the growth of the Antarctic ice-sheet to be charted. Growth of the ice-sheet became especially rapid during the Miocene.

In the late Miocene (Messinian), a combination of tectonic forces and falling sea-levels resulting from the growth of the Antarctic ice-sheet led to the isolation of the Mediterranean from the world ocean. Seawater in the Mediterranean evaporated and deposits of evaporite salts were laid down. During the main period of evaporite deposition from about 5.5 to 4.8 Ma ago there were many alternations of flooding and desiccation, so that the total thickness of evaporites is greater than 1 km. The barrier at Gibraltar was finally submerged at about 4.8 Ma ago, and present-day conditions were established.

The growth and decay of ice-caps is accompanied by expansions and contractions of climatic belts, especially expansion of the cold polar zones. The extent of these fluctuations can be monitored by the study of organisms in deep-sea sediments.

Plate tectonic processes are major long-term influences on sea-level. Rapid sea-floor spreading and continental fragmentation and/or formation of large submarine igneous provinces lead to reduction in the total volume of the ocean basins, resulting in widespread transgressions onto continents. Continental collisions result in thickening and uplift of parts of the continental crust, causing sea-level to fall. Deposition of thick accumulations of sediment along continental margins also contributes to sea-level rise.


The oceans are an important part of the global cycle of chemical elements. Cations (positively charged ions) are supplied to the oceans mainly as a result of atmospheric weathering of terrestrial rocks, and partly from hydrothermal activity. Anions (negatively charged ions) to balance these in the freshwater environment come from the atmosphere (mainly CO2) as bicarbonate HCO3- ions in solution), but in the oceans the cations are balanced mainly by chloride and sulphate anions, whose principal source is volcanic outgassing from the Earth's interior.

A variety of biological and chemical processes are responsible for removal of dissolved constituents from seawater. The resulting sediments and metamorphosed crustal rocks may eventually be returned to the continents by some combination of accretion, subduction, collision, metamorphism, volcanism and uplift.

The compositions of the atmosphere and oceans 2.5-4.5 billion years ago were markedly different from what they are today. Even as recently as about a billion years ago, the environment was more reducing, because there was less oxygen and more CO2. Evolution of the biosphere has played a major role in decreasing the CO2:O2 ratio, and the atmosphere has had approximately its present composition for about the past 200 million years. There may have been variations in seawater composition associated with changes in the rate of hydrothermal exchange, or the rate at which weathering products have been supplied to the oceans from continents.

The oceans are believed to be in a steady-state condition when measured over periods of millions of years, the supply and removal of dissolved constituents being in overall balance.

The speed of movement of water through hydrothermal systems is orders of magnitude greater than the rate of movement of lithspheric plates. Huge volumes of water are cycled through oceanic crust and huge quantities of chemical elements are exchanged in the time that it takes for a piece of crust to move a couple of hundred km from the ridge axis. Rates of global sea-leavel rise and fall may at times be as great as those of horizontal plate movements.