An aqueous solution of salts of a rather constant composition of elements whose presence determines the climate and makes life possible on the Earth and which constitutes the oceans, the mediterranean seas and their embayments. The physical, chemical, biological and geological events therein are the studies that are grouped as oceanography. Water is most often found in nature as seawater (about 98%). The rest is ice, water vapor and freshwater. The basic properties of seawater, their distribution, the interchange of properties between sea and atmosphere or land, the transmission of energy within the sea and the geochemical laws governing the composition of seawater and sediments are the fundamentals of oceanography. See also: Hydrosphere; Oceanography
Water is arguably the most important chemical substance in the terrestrial system, its large specific heat (the highest of almost any substance) is of extreme import to weather, climate and the ocean. Its physical properties include a large dielectric constant, elevated boiling and freezing points, high osmotic pressure and strong dissolving power—properties that are modified by the dissolved salts and the other materials found in the ocean.
The major chemical constituents of seawater are cations (positive ions) and anions (negative ions) [Table 1]. In addition, seawater contains the suspended solids, organic substances and dissolved gases found in all natural waters. A standard salinity of 35 practical salinity units (psu; formerly parts per thousand, or ‰) has been assumed. While salinity does vary appreciably in oceanic waters, the fractional composition of salts is remarkably constant throughout the world's oceans. In addition to the dissolved salts, natural seawater contains particulates in the form of plankton and their detritus, sediments and dissolved organic matter, all of which lend additional coloration beyond the blue coming from Rayleigh scattering by the water molecules. Almost every known natural substance is found in the ocean, mostly in minute concentrations. See also: Scattering of electromagnetic radiation
|Positive ions||Amount, g/kg||Negative ions||Amount, g/kg|
|Sodium (Na+)||10.752||Chloride (Cl)||19.345|
|Magnesium (Mg2+)||1.295||Bromide (Br−)||0.066|
|Potassium (K+)||0.390||Fluoride (F−)||0.0013|
|Calcium (Ca2+)||0.416||Sulfate (SO4−)||2.701|
|Strontium (Sr2+)||0.013||Bicarbonate (HCO3−)||0.145|
|Boron hydroxide [B(OH)3−]||0.027|
*Water, 965 psu: psu dissolved materials, 35 psu.
The density of seawater is an important quantity that enters sensitively into many dynamical and thermodynamical processes. An internationally accepted equation of state has been derived that allows the accurate computation of density, given measurements of salinity, temperature and pressure. A high accuracy is required to compute the small density variations that set up horizontal pressure gradients which, together with the Earth's Coriolis force, balance large-scale geostrophic flows such as the Gulf Stream. Salinity is usually measured via electrical conductivity. See also: Coriolis acceleration; Gulf Stream
Seawater is slightly compressible, implying that sea level is approximately 30 m (100 ft) lower than it would be if it water were incompressible. A parameter known as the thermodynamic space is quite convenient for displaying measurements of temperature and salinity taken during vertical probes of the water column and allows the identification of the origins of various water masses via a technique called isentropic analysis.
The density of sea ice drops to the range of 857–924 kg/m3, depending on the amount of brine. These effects of salinity are related to the very large osmotic pressure of seawater.
Various thermodynamic quantities define other properties of seawater (Table 2). All of these coefficients are functions of salinity, temperature and density and some of their variations with pressure, for example, lead to subtle but profound effects in the sea.
|Isothermal-isohaline compressibility||ap||4.27 × 10−10/Pa|
|Specific heat at constant volume and salinity||Cαs||3939 J/kg-°C|
|Specific heat at constant pressure and salinity||Cps||3994 J/kg-°C|
|Isobaric-isohaline thermal expansion coefficient||aT||2.41 × 10−4/°C|
|Isobaric-isothermal saline contraction coefficient||as||7.45 × 10−4/psu|
|Speed of sound||c||1520 m/s (3280 mi/h)|
|Thermal conductivity||κq||0.596 W/m-°C|
|Latent heat of vaporization||Lv||2.453 × 106 J/kg|
|Latent heat of fusion||Lf||0.335 × 106 J/kg|
*All quantities are evaluated at s = 35 psu and T = 20°C (68°F), except Lf , for which T = 0°C (32°F).
Because of the complex temperature and salinity structure of the upper ocean, highly variable sound propagation paths occur there. The oceanic mixed layer often serves as a leaky waveguide for a sound source located in that layer; in the absence of a vertical temperature gradient, sound propagates along the mixed layer by a combination of upward refraction and surface reflection–forward scattering. However, for a source located beneath the thermocline, the colder water initially refracts propagation paths downward as the sound speed decreases. This is partially offset by the slow increase in speed due to pressure at depth, with the result that there develops a minimum in overall sound speed, the so-called deep sound channel, a secondary waveguide. In tropical and temperate seas, this minimum lies near 1200 m (4000 ft) depth with the result that detectable levels of sound may be heard across entire ocean basins. In a typical sound speed profile for a midlatitude ocean, the major variability occurs in the seasonal mixed layer.
At even greater depths, the pressure continues to increase the sound speed and to bend rays upward. Near-surface rays that are launched beneath the mixed layer at shallow angles will initially penetrate into the deep ocean while undergoing refraction, until near depths of 5000 m (16,500 ft) the rays will have been refracted through the horizontal; subsequently they will bend upward toward the surface again to intersect it at distances near 60–65 km (35–40 mi) from the source. Thus, there is a focusing of sound intensity into an annular region at the surface centered on the source, the so-called first convergence zone; sound levels there may be as much as 20 decibels above those expected from nominal spreading and attenuation. Within the annular region at the surface, reflection and scattering again launch acoustic energy downward, which then goes through the same processes once again; as a result, a series of high-intensity convergence zones are developed in the form of circular annuli surrounding the source, which have radii that are integral multiples of the basic 60–65 km (35–40 mi) and widths of perhaps 5 km (3 mi).
In addition to these intensity variations due to spreading and refraction losses, there exists attenuation from absorption and volume scattering in the seawater. While the attenuation coefficient nominally varies as the square of the frequency, there are significant modifications from dissolved chemicals and acoustic scattering.
Underwater noise levels limit the detection of sound in the sea; they vary enormously with frequency, wind, rain and shipping traffic, ranging from approximately 40 dB at 10 Hz to −30 dB at 10 kHz. At higher frequencies, receiver thermal noise is often limiting. See also: Acoustics; Sound; Underwater sound
The sea absorbs roughly 98% of the visible energy incident on it at high sun angles and converts it to heat, thereby warming the upper layers of the ocean. The solar spectral irradiance is attenuated as sunlight penetrates beneath the sea surface. Under the influences of wind wave stress, turbulence and convective overturning of the upper layers, the thermalized solar energy that is deposited into the upper layers is mixed down to depths ranging perhaps 30–150 m (100–500 ft), depending on location and season. Because of the high specific heat of seawater, this mixed-layer storage represents by far the largest heat capacity on the face of the Earth. Subsequent exchanges across the interface are governed by wind stress and air-sea temperature differences. When the air is cooler than the ocean, as occurs during cold-air outbreaks or at night, significant evaporation sets in. This results in turbulence in the atmospheric boundary layer, which causes rougher seas than when the air-sea temperature difference is positive. Values of air-sea temperature differences in excess of roughly −1°C (−1.8°F) lead to strong interchanges during such unstable conditions. When the air and sea temperatures are nearly the same, the exchanges are greatly reduced. See also: Wind stress
Evaporation of water from the sea results in rising moist air, cloud formation and, upon condensation into clouds at altitude, the release of the latent heat of evaporation to the atmosphere, warming the air. Because the elevated warm air is farther from the Earth's axis of rotation than the air was at sea level, its changed angular momentum is transformed into winds via the Coriolis force. Thus the combination of latent heat release and Coriolis effect initiates the global atmospheric circulation, the wind stress on the ocean and ultimately both the wind-driven and density-driven currents in the sea. Starting in the tropics then, the general circulation of the atmosphere is initiated via air-sea interaction. See also: Atmospheric general circulation
The great poleward-directed systems (such as the Gulf Stream) that are found in every ocean basin transport heat out of the tropics and into temperate and polar regions. During their journey, the warm currents surrender to the atmosphere large amounts of latent and sensible heat, moisture and momentum via the processes of evaporation, turbulent fluctuations, radiation and conduction, in decreasing order of importance. In midlatitudes, for example, this warming of the atmosphere is responsible for maintaining the western parts of Europe at significantly higher average temperatures than in the same latitudes in North America. Over longer times, these large-scale air-sea exchanges constitute one of the most important processes governing weather and climate dynamics.
Turbulence in the planetary boundary layer and the oceanic mixed layer is largely responsible for many of the secondary processes conditioning the air-sea interface. Surface wave generation, mixing and exchanges of heat, momentum and moisture all rely on turbulent diffusion for their efficacy. See also: Heat balance of the Earth; Ocean circulation
Beyond waves and advection by currents, properties such as heat, salt and momentum are transported by diffusive effects in the sea. Molecular diffusion coefficients are quite small; but eddylike fluctuations lead to turbulence and eddy diffusion coefficients which are orders of magnitude greater than their molecular counterparts and which depend on the scale of the fluctuations. The coefficients also differ for horizontal diffusion as compared with vertical diffusion, because buoyancy greatly inhibits transport in the vertical. The energy going into waves, currents and diffusion of properties is ultimately removed from the system at scales near the dissipation length, that is, at about 1 mm. This fluid friction, along with bottom friction, acts to slow down both the atmosphere and ocean, but their large-scale dynamics are maintained by a continuing but unsteady input of energy from the Sun and the atmosphere. See also: Diffusion
Electromagnetic and optical properties of seawater include both surface characteristics and bulk attributes.
The ocean is a moderately good conductor of electricity. Electrical conductivity, like the other material properties, depends not only on salinity but on temperature and pressure as well, although the dependence on pressure is quite weak. In seawater, conductivity is conventionally measured and an algorithm is then used to convert it to salinity; a method for computing the inverse relationship is available as well. Not only is conductivity important for determination of the oceanic density field, but along with the dielectric constant, it is relevant to problems involving remote sensing of the sea surface by microwave methods.
Scattering and emission
Microwave power reflection coefficients (or reflectivities) are dependent on polarization and incidence angle. Waves on the sea modify the reflection coefficients significantly. The customary measure of reflection-scattering from a rough surface is the normalized radar cross section, relative to 1 m2 of ocean surface. This is a function of frequency, incidence, azimuth and reflection angles; polarization; wind speed–sea state; and surfactant coverage. Emission of thermal energy at microwave frequencies is used to determine thermodynamic temperatures, wind speed and, at lower frequencies, salinity of the sea surface. See also: Absorption of electromagnetic radiation; Microwave; Polarized light
The optical characteristics of seawater derive from a combination of physical, chemical, biological and geological properties. Central to the subject are the questions of light transmission in the sea (for studies of biological productivity and underwater signaling), the color of seawater and the use of variations in water color as a means for flow visualization in optical images of the sea.
Surface reflectivities are due to a combination of intrinsic Fresnel reflection and rough surface scattering. The smooth-surface reflectivity is polarization sensitive and is similar to that for microwave frequencies. Since direct sunlight is unpolarized, the variation of reflectivity with incidence angle is given by the average of vertical and horizontal values. As the wind speed increases, reflection falls off.
Index of refraction
Over the nominal range of variation of wavelength, temperature, salinity and pressure, the index of refraction does not depart from 1.34 by more than 0.015.
Volume absorption and scattering
While seawater is a reasonable conductor at microwave frequencies and below, at optical frequencies clearest seawater is a remarkably low-loss dielectric with a broad transparency band in the visible. The spectral beam absorption coefficient is a measure of the fractional absorption of a collimated beam of light in traversing 1 m of water. The addition of photosynthetic biological material profoundly alters the adsorption coefficient; chlorophyll a and phaeophytin a are the major absorbers, along with so-called yellow substance, or dissolved tanninlike compounds. Beyond absorption, the process of scattering is important and this is described by a beam-scattering coefficient which is large in the forward direction but much smaller in the backscatter direction. Thus, the total beam attenuation coefficient is due to absorption and scattering.
Remote sensing with imaging optical spectrometers on satellites has yielded the first global estimates of near-surface chlorophyll and sediment concentrations by utilizing the concepts of ocean optics in the interpretation of multispectral images from satellites and aircraft. See also: Optical detectors; Remote sensing; Scientific and applications satellites
The complex mixture of salts and other substances found dissolved in seawater is quantitatively dominated by chloride and sodium ions, with somewhat lesser amounts of sulfate, magnesium, calcium and other ions. The total salt content of seawater is expressed by the term salinity. This measure of the salt concentration has been rigorously defined in several ways; since 1978 it has been defined and measured according to the ratio of the electrical conductivity of a seawater sample to that of a standard solution of potassium chloride. For both technical and historical reasons, the salinity is not exactly the total mass of salts in solution, but the two values differ by only about 0.5%.
The average salinity of the world ocean is 34.73‰ (parts per thousand by weight, or grams per kilogram of seawater) and is remarkably constant. Only about one-tenth of all the water in the ocean departs by more than 1% from this mean value, so measurement techniques of the most exquisite sensitivity must be used to discern the real and important, differences in salinity from place to place within the oceans. The salinity departs from this relative constancy in estuaries and other near-coastal regions where seawater is diluted with river water, in polar regions where ice melts in the spring and in coastal lagoons or small seas where evaporation may sometimes exceed the input of freshwater by rivers and rain. Since evaporation and rainfall generally change only the total salt content and not the relative proportions of the salts, it is conventional to express the concentrations of substances in the ocean on the basis of a standard ocean water with a salinity of 35‰ (Tables 3 and 4).
|Sodium ion (Na+)||10.781||468.96|
|Potassium ion (K+)||0.399||10.21|
|Magnesium ion (Mg2+)||1.284||52.83|
|Calcium ion (Ca2+)||0.4119||10.28|
|Chloride ion (Cl−)||19.353||545.88|
|Sulfate ion (SO42−)||2.712||28.23|
|Bicarbonate ion (HCO3−)||0.126||2.06|
|Bromide ion (Br−)||0.0673||0.844|
|Boron hydroxide [B(OH)3−]||0.0263||0.425|
|Fluoride ion (F−)||0.00130||0.086|
*Total mass concentration of above constituents, 35.170; total concentration of all other dissolved substances, ∼0.03; water, ∼964.80.
*At 18°C (64°F) and salinity of 35% and at pressure of one standard atmosphere.
Major and minor elements
It is also conventional to consider that the major dissolved constituents of seawater consist of 11 substances (Table 3), all present in concentrations of greater than 1 mg/kg or 1 part per million (ppm). In addition, however, the dissolved gases nitrogen and oxygen, while variable in concentration, generally amount to a total of about 10–30 mg/kg and dissolved organic matter is present in variable concentrations of about 1 to several milligrams per kilogram.
The total concentration of salts in seawater, about 35 g in a kilogram, is great enough that the chemical activity of each ion is strongly affected by the presence of the others and in addition numerous ion pairs and other complexes are present. These effects must be taken into account in any quantitative evaluation of the behavior of substances in seawater.
With the exceptions of oxygen and helium, the dissolved gases in seawater (Table 4) are not subject to any processes that can significantly change their concentrations once these concentrations have been established by coming to equilibrium with the atmosphere. They will not change during the circulation of that water into deeper parts of the ocean, except by mixing with other parcels. The concept of attaining equilibrium with the atmosphere is, however, a little fuzzy. It is common to assume standard atmospheric pressure, but atmospheric pressure varies. During windy conditions, breaking waves cause the entrainment of bubbles to some depth in the water; as these dissolve, they increase the concentrations, sometimes up to several percent above saturation. Observation shows that all ocean water has been affected to a measurable extent by these effects, so that they have to be taken into account when making careful calculations. Once below the surface, however, the concentrations of nitrogen and the noble gases (with the exception of helium) are not subject to further change. Helium, especially the rare isotope helium-3, is increased in deep water by its release from the mantle in places where fresh hot basalt comes in contact with seawater. Nitrogen is used and produced by biological processes, but not in amounts that can easily be measured in most of the ocean. Oxygen is both produced and consumed in biological reactions and measurements of the concentration of oxygen are thus an important tracer of oceanic processes.
The minor elements exhibit a great range of concentrations. Most are present in only tiny fractions of a part per billion, providing a great challenge to the analyst. In many cases the limitation has been the great difficulty of collecting seawater from ships without contamination being far greater than the concentrations in the water, as well as the difficulty of obtaining reagents and equipment that do not contain more of the trace elements than the water samples themselves. It is possible to measure such low concentrations by combining great attention to ultraclean sampling and processing with modern highly sensitive instrumentation.
Unlike the major elements and the conservative minor elements, the important plant nutrients (phosphate, nitrate and silicate) are not uniformly distributed in the ocean according to the salinity of the water. Instead, they exhibit marked changes from place to place, both vertically and horizontally. Characteristically, the nutrients are depleted in surface waters and are enriched at depth (Fig. 1). Similar and related profiles from much of the ocean have provided the general understanding of the interaction of ocean circulation and biogeochemical processes.
Small plants (the mostly unicellular photosynthetic plankton) in the sunlit surface waters use these nutrients (along with carbon and smaller amounts of other nutrients) to form their body structures; these in turn provide the base of the food chain from zooplankton all the way up to the largest fish and whales. Particulate matter, including plant cells, feces and other debris, sinks below the surface and is eaten by deeper-living plankton, fish, bacteria and other organisms throughout the water column and on the bottom. The metabolism of these deeper living organisms results in the release and enrichment of the deeper water with carbon dioxide, phosphate and the other nutrients and the consumption of oxygen. See also: Food web; Marine ecology; Phytoplankton
The difference in the concentrations of the nutrients between the Mediterranean and the Atlantic is explained by the circulation of seawater. Surface water from the Atlantic flows into the Mediterranean and is already depleted in nutrients. The return flow through the Strait of Gibraltar is somewhat deeper, underneath the surface inflow, so that the small accumulation of nutrients in the deeper water tends to be removed. The Pacific Ocean provides the reverse example. Deep water from the Atlantic flows into the Pacific, carrying its load of nutrients; in the Pacific it upwells and the return flow to the Atlantic includes a considerable amount of surface water, depleted in nutrients through the activities of the biota that caused the transport of the nutrients back to the deep water in the Pacific. The nutrients therefore tend to be retained and to build up in the Pacific; the same is also true of the Indian Ocean. In many cases, more detailed analyses also provide analogous explanations of the structure seen in the vertical profiles. See also: Atlantic Ocean; Indian Ocean; Mediterranean Sea; Pacific Ocean
Surface water concentrations of nutrients vary considerably throughout the ocean, depending on the detailed physical processes that bring water from some depth to the surface. These upwelling regions are especially notable, for example, off the western shores of North and South America, off the southwestern shore of Africa and in the complex current systems flowing along the Equator. Accordingly, these are regions of high biological productivity and also regions where there is a very strong transport of nutrients from the surface downward as the debris sinks. Shallow water regions tend also to be very productive, for two reasons. First, the interaction of ocean currents with these shallow regions causes turbulence that enhances the mixing of surface and deeper nutrient-rich water. Second, the sinking debris cannot go very deep and much is metabolized on the shallow bottom, so that the resulting nutrients are easily mixed into the surface water again. See also: Seawater fertility; Upwelling
In general, an inverse relationship prevails between the concentration of oxygen (O2) in the deep water of the ocean and the concentrations of phosphorus (P) as phosphate [and nitrogen (N) as nitrate] and carbon dioxide (CO2). The reason is that, on average, the debris that sinks to depth has a generally uniform composition and the nutrients are released in proportion to the oxygen used in metabolism of the debris. The quantitative relationship is expressed in the Redfield ratio: 175O2:122C:16N:1P; that is, for every 175 molecules of oxygen that are utilized, 122 atoms of carbon (as carbon dioxide), 16 atoms of nitrogen (as nitrate) and 1 atom of phosphorus (as phosphate) are released. Variations in this ratio from place to place appear to be only a few percent. Additional carbon dioxide, in the form of carbonate ion (CO32−), is released by particles of calcium carbonate in places where these dissolve.
In addition to phosphorus, nitrogen and carbon, organisms need a number of other substances in small amounts. For example, some metals are needed, such as iron, copper, molybdenum, zinc and cobalt, with iron in the greatest amount. Iron is quite insoluble in seawater, however and is much depleted in the surface water of the open ocean. There are a number of places in the ocean, generally far from land, where the other nutrients are available, but the growth of plankton is limited by the availability of iron. In these cases the surface depletions of phosphate and nitrate are less evident.
Many trace elements show vertical distributions that, to varying degrees, resemble the vertical distributions of phosphate. For those that are known to be plant nutrients, these distributions are easily explicable, as they must be carried down with the sinking debris from life at the surface. In other cases it appears that some elements may be simply adsorbed to the surfaces of sinking particles, carried down and released at depth. The vertical profiles of concentration of some elements exhibit quite different shapes; in these cases other processes must be called upon to explain them.
The chemistry of carbon dioxide in seawater is complicated and important, not least because the oceans are taking up a considerable part of the carbon dioxide produced by the burning of fossil fuels and by other human activities. When carbon dioxide dissolves in seawater, it enters into a number of chemical equilibrium reactions. The equilibrium constants of all such reactions are affected by temperature and other parameters. Of particular importance are those reactions that are strongly affected by the temperature and salinity of the water and by hydrostatic pressure and accurate knowledge of all these parameters is important in assessing the chemical state of the carbonate system. Because of the participation of the hydrogen ion (H+) in some reactions, the concentration of this ion (expressed usually as the pH) is an important indicator of the state of the system. When the concentration of carbon dioxide increases, the pH drops; and when carbon dioxide decreases, the pH rises. Accordingly, as anthropogenic carbon enters the ocean it causes a small decrease in the pH. The effect is, so far, much smaller than the changes commonly caused by the growth of plants (increase in pH) and by the respiration of all living things (decrease in pH).
An additional complication is the precipitation of calcium carbonate by planktonic plants and animals (as well as by many types of organisms on coral reefs). The solubility of calcium carbonate is also strongly affected by temperature, salinity and pressure. The shells and other bits of calcium carbonate formed in the plankton also sink, directly or in fecal pellets, out of surface layers and tend to dissolve at depth. In many places they dissolve completely because of the combined effects of pressure and the acidification of the deep water by respiratory carbon dioxide. The concentration of carbon dioxide in deep water is therefore increased, relative to the surface (Fig. 2). As with the nutrients, this concentration increases from the North Atlantic to the South Atlantic, through the South Pacific to the North Pacific. See also: Chemical equilibrium; Hydrogen ion; pH
Most radioactivity in seawater is due to nuclides present since the world began (Table 5). The distributions of many of the radioactive nuclides resulting from the decay of uranium isotopes and those produced by cosmic-ray-induced processes in the atmosphere, as well as those produced artificially and deposited into the ocean after nuclear explosions in the atmosphere, have been utilized to help trace mixing and transport processes in the ocean and have helped provide quantitative understanding of the rates of these processes. See also: Radioactivity; Radioisotope; Uranium
|Uranium-234 (234U)†||Decay of 238U|
|Strontium-90 (90Sr)||Bomb explosions|
|Cesium-137 (137Cs)||Bomb explosions|
|Carbon-14 (14C)||Cosmic rays in atmosphere and bomb explosions|
|Hydrogen-3 [tritium] (3H)||Cosmic rays in atmosphere and bomb explosions|
*The total number detected is well over 40, but those not listed are present in very small concentrations.
†Nuclides of uranium that are the parents of long series of radioactive decay products.
Changes in composition
The overall chemical composition of the ocean results from the balance of numerous processes. These processes include (1) the weathering of exposed rocks on continents and islands, with the transport of the resulting materials to the oceans; (2) the uptake and release of various substances when seawater comes in contact with hot spots on the sea floor; (3) exchange with the atmosphere, mostly gases; (4) uptake and release of many substances by organisms; (5) sinking of particles from surface water into deep water; (6) chemical exchanges between seawater and sediments and chemical reactions occurring within the sediments; and (7) precipitation of complex mixtures of salts in places where seawater can evaporate to dryness, a process that has varied greatly in magnitude over the course of geological time. See also: Hot spots (geology)
It should not be assumed that all these processes are constant over the span of geologic time. Known changes in climate, rates of weathering of the continents, rates of sea-floor spreading and rearrangements of geological structures must have caused some changes in the chemistry of the ocean. Most of these changes would not likely be perceptible on the time scale of human history. However, the activities of human civilization since the beginning of the industrial revolution have resulted in detectable chemical signals throughout most of the ocean, especially in surface water.
Distribution of Properties
The distribution of physical characteristics in the ocean is principally the result of radiation (of heat), exchange with the land (of water and salt), exchange with the atmosphere (water and heat) and mixing and stirring processes within the ocean.
The ocean is heated by the sun, but very little of its radiation penetrates more than a few meters beneath the surface. Essentially all the heat lost by the ocean is across the air-sea interface by long-wave radiation, evaporation and conduction. Although the heat losses balance the heat gains over an annual cycle, they seldom balance at any one location. The surface ocean gains more heat than it loses equatorward of about 35° and it loses more heat than it gains poleward of 35°. As a consequence, the sea surface temperature is high (more than 28°C or 82°F) in equatorial regions and low (less than 1°C or 34°F) in polar regions. Below a shallow surface layer of a few hundred meters, the temperature structure of the ocean is a result of this continuing poleward transfer of the order of 1015 W of heat energy.
Various dissolved solids have entered the sea from the land and have been so mixed that their relative amounts are everywhere nearly constant. For well over 95% of the ocean, the salt content is 3.43–3.51% by weight. Oceanographers routinely measure salinity to a few thousands of a percent and salinity is reported in parts per thousand (‰). Considerable insight in oceanic processes can be achieved by examining changes in seawater salinity distributions of the order of ±0.02‰. In the middle latitudes the evaporation of water exceeds precipitation and the surface salinity is high; in low and high latitudes precipitation exceeds evaporation and dilution reduces the surface salinity. The Atlantic Ocean is more saline than the Pacific and Indian oceans.
The density of seawater varies with temperature, salinity and pressure. It can vary horizontally only in the presence of currents and hence its distribution depends closely upon the structure of the flow.
Pycnocline, thermocline and halocline
The ocean is stratified. The upper layer, in contact with the atmosphere, is warmed and cooled by exchange of heat through the surface and freshened and made more saline by rainfall and evaporation. Over the ocean this layer varies in depth from less than 100 m (330 ft) to more than 300 m (1000 ft). During summer heating, a thin cap of warmer water may be formed at the top, disappearing in fall and winter by cooling and mixing with the underlying water.
Beneath the upper layer, there is a strong increase in density, called the pycnocline. This increase may include a sharp drop in temperature (the thermocline) or a rise in salinity (the halocline). In most middle latitudes, both temperature and salinity decrease. In high latitudes, both temperature and salinity may increase downward if the vertical gradient of salinity is strong enough.
Characteristics beneath the surface
The oceans of the world are comprised of layers (Figs. 3 – 5). Because seawater is slightly compressible, it is made warmer with pressure. The potential temperature which is defined as the temperature that a parcel of water would have if it were acted upon by atmospheric pressure alone. Near 4000 m (14,000 ft), where pressures exceed 400 atm (41,000 kilopascals), the actual temperatures are a few tenths of a degree higher than the potential temperatures.
The densest waters of the open ocean are formed in high latitudes by cooling of some of the waters that have flowed poleward from the more saline regions. The principal source of the abyssal waters is the Weddell Sea, with some contribution from the Ross Sea and some from the Norwegian-Greenland Sea passing over the shallow sill between Greenland and Iceland into the North Atlantic Ocean (Fig. 3). The dense waters formed in this fashion are cold and moderate in salinity. They flow equatorward along the western boundaries of the oceans and become progressively warmer, more saline and less dense by mixing with the overlying water.
The abyssal waters formed in the Weddell and Ross seas are not formed by convection to the bottom in the open ocean. Though these seas are at freezing temperature most of the year, the salinity in the upper layer is so low (about 34.0 parts per thousand) that even at the freezing point the surface waters are not as dense as the underlying waters. The densest waters are formed on the continental shelves, where two processes raise the salinity. Higher-salinity waters from lower latitudes extend southward beneath the surface and raise the salinity on the continental shelves. Leaching out of brine from the ice adds significant amounts of salt to the shallower layer on the shelves. At these raised salinities and the freezing point, the shelf waters become dense enough to spill down the continental slopes to abyssal depths. These waters mix with the open-ocean waters as they pour down. As they reach the bottom, their temperature is near −0.25°C (31.6°F). Their salinity is 34.66–34.7 parts per thousand, far higher than any of the open-ocean salinity values of the Antarctic Ocean.
Overturn to the bottom
Overturn to the bottom does not occur in the open ocean, but takes place only in adjacent seas. The warm and saline waters of the open North Atlantic Ocean flow into the Norwegian-Greenland Sea, where the intense cooling increases their density and convection occurs. The passages from the Norwegian-Greenland Sea to the open ocean are narrow and shallow; and where the dense Norwegian-Greenland Sea waters spill back into the Atlantic, they mix with the less dense waters and their density is reduced. The waters that pour down from the Denmark Strait, between Greenland and Iceland and still dense enough to reach the bottom in the North Atlantic Ocean, but their density is not as great as those at the bottom of the Weddell Sea. The mixture from the overflow between Scotland and Iceland is less dense and reaches only about 3000 m (10,000 ft). Their major contribution is to the middepth water.
The Mediterranean Sea is an evaporation basin. Water from the Atlantic flows in at the surface to make up the loss to the atmosphere. Evaporation raises the salinity so high (more than 38 parts per thousand) that even at a temperature of 12°C (36°F) this water is denser than any of the Atlantic waters and the Mediterranean Sea overturns to the bottom. Then the water pours out through the Strait of Gibraltar, beneath the inflowing Atlantic water and spills down the slope. Rapid mixing with the Atlantic waters and the water's low compressibility (from its high temperature) keep it from penetrating to the bottom. It provides the very high salinity and temperature that characterize the middepth Atlantic Ocean.
Between the upper layer and the dense abyssal layer, the middepth waters are also supplied from the higher latitudes. Warm and extremely saline water from the Mediterranean Sea makes the North Atlantic the warmest and saltiest of the oceans at depths of 1000–2500 m (3300–8300 ft; Figs. 3 – 5). Between Iceland and Scotland, warm and saline water from the near-surface waters of the eastern North Atlantic meets the cold and extremely dense water from the Norwegian Sea. Their mixture produces a denser layer of warm and saline water. This water flows westward south of Iceland and Greenland and into the Labrador Sea, where it turns equatorward along the western boundary.
These three North Atlantic sources (Greenland Sea, Mediterranean Sea and Norwegian Sea–Northeastern Atlantic) account for the characteristics of much of the middepth water in the world ocean. The high salinity in this layer extends southward through the Atlantic (Fig. 3) at depths of 1500–2500 m (5000–8300 ft). It rises with the Antarctic Circumpolar Current near 45 to 60°S latitude and turns westward into the South Indian Ocean (Fig. 4). It continues through the Pacific Ocean (Fig. 5), where part of it turns northward and part continues eastward to return, in a diminished form, to the Atlantic Ocean.
Layer of low salinity
The surface waters in high latitudes are made colder and less saline and dense than the adjacent midlatitude waters by exchange of heat and water with the atmosphere. As they extend equatorward, they descend below the upper layer because of their higher density. They are recognized near 800–1000 m (2700–3300 ft) as layers of low salinity from the high southern latitudes of all three oceans and also from the north in the Pacific.
Layers of high salinity
The highest salinity in the open ocean is found in the upper layer near the tropic circles, where evaporation exceeds precipitation. Vertical mixing through the pycnocline takes place only slowly and the high surface salinity does not penetrate very well. But these high salinities do spread laterally beneath the less dense surface waters of lower latitudes. In all three oceans (Figs. 3 – 5), shallow layers of high salinity can be seen extending equatorward from the great high-salinity cells around the tropic circles.
There is a rich spectrum of temperature variations in the ocean. Those with lateral scales of thousands of kilometers affect the horizontal density distribution and, hence, have a crucial role in determining the geostrophic currents. At the smallest scales (on the order of a centimeter), the temperature variations are usually so small that they have no effect on the density. In these cases the temperature behaves as a passive tracer that shows the stirring of water parcels. Between these extreme spatial scales, temperature can be a tracer of the motion or an active participant in the motion via its effect on the density.
Since temperature is relatively easy to measure, it has been a popular way to study the processes which stir and mix the oceans. Two general types of measurements predominate: (1) vertical profiles of temperature or its gradient, either single profiles or repeated profiles from a steaming or drifting vessel; and (2) horizontal sections of isotherm depth, derived from vertical arrays of many thermistors towed through the water. Operational problems are significant and the data are always limited by the fact that the vertical profiles are spaced horizontally and the towed arrays have fixed and finite vertical spacing between the sensors.
Temperature generally decreases with depth in the ocean. This decrease is neither steady nor monotonic. The detailed fluctuations with vertical scales on the order of meters are called fine structure. Because of the stratification of the ocean, these features are relatively thin and wide; that is, they have a large aspect ratio. Fine-structure features extend horizontally 100–1000 times (or more) further than they extend vertically.
A parcel of water that is warmer or colder than the water found both above and below must have come into the water column laterally. Known as intrusions, they are quite common (and especially dramatic) in frontal regions where waters with substantially different salinity and temperature characteristics meet. Intrusions 20 m (66 ft) thick with horizontal scales of several kilometers are observed in coastal regions. The rate of change of temperature with depth varies greatly in magnitude and even changes sign. Intrusions are driven by the horizontal pressure gradients that arise from the horizontal density differences. It is possible to have intrusions without a temperature maximum or minimum. These cannot be recognized from the temperature alone, but require examination of the salinity, the density, or some other property such as oxygen.
Another form of temperature fine structure is created by vertical mixing. As the mixing events in the ocean are rarely thorough enough to produce a truly homogeneous patch of water, the result is a relatively low gradient in the patch and relatively high gradients at the edges. A special case of fine structure due to mixing is associated with double diffusive convection and leads to very strong layering of almost isothermal water with a horizontal extent of many kilometers.
Fine structure is also created by internal waves. Waves produce differential motion with depth, producing alternating regions of horizontal convergence and divergence. Even in an ocean with a smooth temperature profile that decreases monotonically, internal waves can produce local regions of higher and lower gradient. As a result, a smooth temperature profile develops kinks but not inversions. The irregular fine-structure features in the profile can be observed to propagate vertically.
Fine structure due to internal wave motion appears and disappears with the arrival and departure of the wave energy. Hence, it is called reversible fine structure. Lateral intrusions and vertical mixing are not reversible processes; they generate irreversible fine structure. Irreversible fine structure is left behind after the processes that produced it have finished their work. Measurements indicate that in most regions free from intrusions the reversible fine structure dominates over the irreversible fine structure, which would arise from vertical mixing. In regions with intrusions, no general statement can be made.
At scales smaller than the fine structure, from less than 1 m (3.3 ft) down to a few millimeters, the temperature fluctuations are called microstructure. This microstructure is usually due to turbulent mixing processes, with temperature acting as a tracer of the motion. The aspect ratio of these features (ratio of horizontal to vertical scale) decrease with the vertical scale. It might be as large as 100 for 1-m vertical scales and shrink toward 1 at the centimeter scale. The magnitude of the temperature variations decreases with decreasing size.
The magnitude of the fluctuations is determined by the local temperature gradient as well as the size and intensity of the mixing event. Thus, there is no simple relation between the temperature variations and the motion in the water. Molecular thermal diffusivity smooths out the inhomogeneities at all scales, but this effect is most noticeable at scales of a few centimeters and smaller. At these smallest scales, the time for the temperature gradient to diffuse away is from a few seconds (1 mm) to 10 min (2 cm). Therefore, each patch of microstructure is transient, lasting only slightly longer than the processes which produce it.
The detailed features of the thermal structure portend similar features in other parameters such as salinity, density, oxygen and nutrients. The physical processes that produce fine structure and microstructure also affect the distribution of phytoplankton, small herbivores and carnivores. However, there are many other processes that affect the chemical and biological distributions in addition to the physical processes that dominate the thermal structure.