Ocean basins Figure 7: Major features of the ocean basins. 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 (see Figure 7). 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 highabout 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 lighti.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 187276, 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 frequencies3.5 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 1020 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. Figure 8: Gravity map of the world's ocean basins, compiled from Seasat satellite data (see text). 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 (Figure 8) 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-northwestvery close to the direction in which the Pacific Plate moves over the mantle below. Chemical and physical properties of seawater Composition of seawater The chemical composition of seawater is influenced by a wide variety of chemical transport mechanisms. Rivers add dissolved and particulate chemicals to the oceanic margins. Wind-borne particulates are carried to mid-ocean regions thousands of kilometres from their continental source areas. Hydrothermal solutions that have circulated through crustal materials beneath the seafloor add both dissolved and particulate materials to the deep ocean. Organisms in the upper ocean convert dissolved materials to solids, which eventually settle to greater oceanic depths. Particulates in transit to the seafloor, as well as materials both on and within the seafloor, undergo chemical exchange with surrounding solutions. Through these local and regional chemical input and removal mechanisms, each element in the oceans tends to exhibit spatial and temporal concentration variations. Physical mixing in the oceans (thermohaline and wind-driven circulation; see below Circulation of the ocean waters) tends to homogenize the chemical composition of seawater. The opposing influences of physical mixing and of biogeochemical input and removal mechanisms result in a substantial variety of chemical distributions in the oceans. Dissolved inorganic substances In contrast to the behaviour of most oceanic substances, the concentrations of the principal inorganic constituents of the oceans (Table 2) are remarkably constant. For 98 percent of the oceans' volume, the concentrations of the constituents shown in the Table vary by less than 3 percent from the values given in columns 2 and 3. Furthermore, with the exception of inorganic carbon, the principal constituents shown in the Table have very nearly fixed ion concentration ratios (column 4). Calculations indicate that, for the main constituents of seawater, the time required for thorough oceanic mixing is quite short compared to the time that would be required for input or removal processes to significantly change a constituent's concentration. The concentrations of the principal constituents of the oceans vary primarily in response to a comparatively rapid exchange of water (precipitation and evaporation), with relative concentrations remaining nearly constant. Salinity is used by oceanographers as a measure of the total salt content of seawater. Practical salinity, symbol S, is determined through measurements of a ratio between the electrical conductivity of seawater and the electrical conductivity of a standard solution. Practical salinity can be used to calculate precisely the density of seawater samples. Because of the constant relative proportions of the principal constituents, salinity can also be used to directly calculate the concentrations of the major ions in seawater. Using the relative concentrations shown in column 4 of Table 2, ionic concentrations are calculated as 0.015577 mole per kilogram multiplied by salinity multiplied by relative concentration. The measure of practical salinity was originally developed to provide an approximate measure of the total mass of salt in one kilogram of seawater. Seawater with S equal to 35 contains approximately 35 grams of salt and 965 grams of water. Although the 11 constituents shown in Table 2 account for more than 99.5 percent of the dissolved solids in seawater, many other constituents are of great importance to the biogeochemistry of the oceans. Such chemicals as inorganic phosphorus (HPO2-/4 and PO3-/4) and inorganic nitrogen (NO-/3, NO-/2, and NH+/4) are essential to the growth of marine organisms. Nitrogen and phosphorus are incorporated into the tissues of marine organisms in approximately a 16:1 ratio and are eventually returned to solution in approximately the same proportion. As a consequence, in much of the oceanic waters dissolved inorganic phosphorus and nitrogen exhibit a close covariance. Dissolved inorganic phosphorus distributions in the Pacific Ocean strongly bear the imprint of phosphorus incorporation by organisms in the surface waters of the ocean and of the return of the phosphorus to solution via a rain of biological debris remineralized in the deep ocean. Inorganic phosphate concentrations in the western Pacific range from somewhat less than 0.1 micromole per kilogram (1 10-7 mole per kilogram) at the surface to approximately 3 micromoles/kg (3 10-6 mole/kg) at depth. Inorganic nitrogen ranges between somewhat less than 1 micromole/kg and 45 micromoles/kg along the same section of ocean and exhibits a striking covariance with phosphate. A variety of elements essential to the growth of marine organisms, as well as some elements that have no known biological function, exhibit nutrient-like behaviour broadly similar to nitrate and phosphate. Silicate is incorporated into the hard structural parts of certain types of marine organisms (diatoms and radiolarians) that are abundant in the upper ocean. Dissolved silicate concentrations range between less than 1 micromole/kg (1 10-6 mole/kg) in surface waters to approximately 180 micromoles/kg (1.8 10-4 mole/kg) in the deep North Pacific. The concentration of zinc, a metal essential to a variety of biological functions, ranges between approximately 0.05 nanomole/kg (5 10-11 mole/kg) in the surface ocean to as much as 6 nanomoles/kg (6 10-9 mole/kg) in the deep Pacific. The distribution of zinc in the oceans is observed to generally parallel silicate distributions. Cadmium, though having no known biological function, generally exhibits distributions that are covariant with phosphate and concentrations that are even lower than those of zinc. Many elements, including the essential trace metals iron, cobalt, and copper, show surface depletions but in general exhibit behaviour more complex than that of phosphate, nitrate, and silicate. Some of the complexities observed in elemental oceanic distributions are attributable to the adsorption of elements on the surface of sinking particles. Adsorptive processes, either exclusive of or in addition to biological uptake, serve to remove elements from the upper ocean and deliver them to greater depths. The distribution patterns of a number of trace elements are complicated by their participation in oxidation-reduction (electron-exchange) reactions. In general, electron-exchange reactions lead to profound changes in the solubility and reactivity of trace metals in seawater. Such reactions are important to the oceanic behaviour of a variety of elements, including iron, manganese, copper, cobalt, chromium, and cerium. The processes that deliver dissolved, particulate, and gaseous materials to the oceans ensure that they contain, at some concentration, very nearly every element that is found in the Earth's crust and atmosphere. The principal components of the atmosphere, nitrogen (78.1 percent), oxygen (21.0 percent), argon (0.93 percent), and carbon dioxide (0.035 percent), occur in seawater in variable proportions, depending on their solubilities and oceanic chemical reactions. In equilibrium with the atmosphere, the concentrations of the unreactive gases, nitrogen and argon, in seawater (0 C, salinity 35) are 616 micromoles/kg and 17 micromoles/kg, respectively. For seawater at 35 C, these concentrations would decrease by approximately a factor of two. The solubility behaviours of argon and oxygen are quite similar. For seawater in equilibrium with the atmosphere, the ratio of oxygen and argon concentrations is approximately 20.45. Since oxygen is a reactive gas essential to life, oxygen concentrations in seawater that are not in direct equilibrium with the atmosphere are quite variable. Although oxygen is produced by photosynthetic organisms at shallow, sunlit ocean depths, oxygen concentrations in near-surface waters are established primarily by exchange with the atmosphere. Oxygen concentrations in the oceans generally exhibit minimum values at intermediate depths and relatively high values in deep waters. This distribution pattern results from a combination of biological oxygen utilization and physical mixing of the ocean waters. Estimates of the extent of oxygen utilization in the oceans can be obtained by comparing concentrations of oxygen with those of argon, since the latter are only influenced by physical processes. The physical processes that influence oxygen distributions include, in particular, the large-scale replenishment of oceanic bottom waters with cold, dense, oxygen-rich waters sinking toward the bottom from high latitudes. Due to the release of nutrients that accompanies the consumption of oxygen by biological debris, dissolved oxygen concentrations generally appear as a mirror image of dissolved nutrient concentrations. While the atmosphere is a vast repository of oxygen compared to the oceans, the total carbon dioxide content of the oceans is very large compared to that of the atmosphere. Carbon dioxide reacts with water in seawater to form carbonic acid (H2CO3), bicarbonate ions (HCO-/3), and carbonate ions (CO2-/3). Approximately 90 percent of the total organic carbon in seawater is present as bicarbonate ions. The formation of bicarbonate and carbonate ions from carbon dioxide is accompanied by the liberation of hydrogen ions (H+). Reactions between hydrogen ions and the various forms of inorganic carbon buffer the acidity of seawater. The relatively high concentrations of both total inorganic carbon and boronas B(OH)3 and B(OH)-/4in seawater (see Table 2) are sufficient to maintain the pH of seawater between 7.4 and 8.3. (The term pH is defined as the negative logarithm of the hydrogen ion concentration in moles per kilogram. Thus, a pH equal to 8 is equivalent to 1 10-8 mole of H+ ions per kilogram of seawater.) This is quite important because the extent and rate of many reactions in seawater are highly pH-dependent. Carbon dioxide produced by the combination of oxygen and organic carbon generally produces an acidity maximum (pH minimum) near the depth of the oxygen minimum in seawater. In addition to exchange with the atmosphere and, through respiration, with the biosphere, dissolved inorganic carbon concentrations in seawater are influenced by the formation and dissolution of the calcareous shells (CaCO3) of organisms (foraminiferans, coccolithophores, and pteropods) abundant in the upper ocean. Circulation of the ocean waters General observations The general circulation of the oceans defines the average movement of seawater, which, like the atmosphere, follows a specific pattern. Superimposed on this pattern are oscillations of tides and waves, which are not considered part of the general circulation. There also are meanders and eddies that represent temporal variations of the general circulation. The ocean circulation pattern exchanges water of varying characteristics, such as temperature and salinity, within the interconnected network of oceans and is an important part of the heat and freshwater fluxes of the global climate. Horizontal movements are called currents, which range in magnitude from a few centimetres per second to as much as 4 metres per second. A characteristic surface speed is about 5 to 50 centimetres per second. Currents diminish in intensity with increasing depth. Vertical movements, often referred to as upwelling and downwelling, exhibit much lower speeds, amounting to only a few metres per month. As seawater is nearly incompressible, vertical movements are associated with regions of convergence and divergence in the horizontal flow patterns. Ocean circulation derives its energy at the sea surface from two sources that define two circulation types: (1) wind-driven circulation forced by wind stress on the sea surface, inducing a momentum exchange, and (2) thermohaline circulation driven by the variations in water density imposed at the sea surface by exchange of ocean heat and water with the atmosphere, inducing a buoyancy exchange. These two circulation types are not fully independent, since the sea-air buoyancy and momentum exchange are dependent on wind speed. The wind-driven circulation is the more vigorous of the two and is configured as large gyres that dominate an ocean region. The wind-driven circulation is strongest in the surface layer. The thermohaline circulation is more sluggish, with a typical speed of one centimetre per second, but this flow extends to the seafloor and forms circulation patterns that envelop the global ocean. Distribution of ocean currents Maps of the general circulation at the sea surface are constructed from a vast amount of data obtained from inspecting the residual drift of ships after course direction and speed are accounted for in a process called dead reckoning. This information is amplified by satellite-tracked drifters at sea. The pattern is nearly entirely that of wind-driven circulation. Deep-ocean circulation consists mainly of thermohaline circulation. The currents are inferred from the distribution of seawater properties, which trace the spreading of specific water masses. The distribution of density or field of mass is also used to estimate the deep currents. Direct observations of subsurface currents are made by deploying current meters from bottom-anchored moorings and by setting out neutral buoyant instruments whose drift at depth is tracked acoustically.
Meaning of OCEAN BASINS in English
Britannica English vocabulary. Английский словарь Британика. 2012