oceanography.jpg - 6879 Bytes

S330 Oceanography

Block 2: Seawater: Its Composition, Properties and Behaviour


Why is the ocean salty?

Variations in a salty ocean


Gases in seawater

Ocean carbon cycle modelling



The effect of dissolved salts
Water in the atmosphere
Ice in the oceans


The transfer of heat and water across the air-sea interfaces


Changes due to local conditions
Salts from seawater
Distribution of salinity with depth
Distribution of surface salinity
Chemical methods of salinity measurement
Physical methods of salinity measurement
The formal definition of salinity


Adiabatic temperature changes
Using σt
σθ and vertical stability
The use of T-S diagrams
Conservative and non-conservative properties
Molecular and turbulent diffusion
Stratification and microstructure


Illumination and vision
Underwater visibility; seeing and being seen
Colour in the sea
Electromagnetic radiation and remote sensing of the oceans
The main characteristics of sound waves in the oceans
The speed of sound: refraction and sound channels
Uses of acoustic energy in the oceans


classification of dissolved constituents
The nutrients
Dissolved gases
Dissolved gases as tracers
Comparison of seawater with other natural waters
Seawater and river water
Origin of the chloride
The sodium balance
Chemical fluxes and residence times
Interactions between dissolved species
The carbonate system, alkalinity and control of pH
Non-biological controls on minor and trace element concentrations
Biological controls on minor and trace element concentrations
Biological activity as a sink for trace elements


The special case of CO2
Climate and the Earth's orbit

Back to OU

inthegreeny.jpg - 1215 Bytes



The special properties of water - in particular, its anamalously high melting and boiling points, specific and latent heats, powerful solvent properties, and maximum density at 4 C - result from the polar structure of the water molecule. Dissolved salts increase the density of water and depress both the temperature of maximum density and the freezing point.

The oceans contain 97% of the water that circulates in the hydrological cycle. The residence time of water in the oceans is measured in thousands of years; in the atmosphere, it is measured in days.

Air is saturated with water vapour when evaporation is balanced by condensation. Clouds and fog are dondensed water vapour. Fog may form when air is cooled to its condensation temperature, either by radiation from the land, or by advection of warm humid air over a cool land or water surface.

Sea-ice is less saline than the seawater from which it freezes, so its formation increases the salt content of the remaining seawater, thus furthr depressing its freezing point and increasing its density. Icebergs in the Northern Hemisphere are formed when valley glaciers on lands surrounding the Arctic Ocean reach the sea; those in the Southern Hemisphere break off from the thick ice-shelf that surrounds the Antarctic continent.


The Earth's surface temperature is mainly determined by the amount of solar radiation it receives. On average, about 70% of incoming solar radiation reaches the surface, directly or indirectly. The proportion varies with latitude, season and time of day, and the amount absorbed depends on the albedo of the surface. The oceans have a large thermal capacity because of the high specific and latent heat of water, and they act as a temperature buffer for the surface of the Earth as a whole. Annual insolation is greatest in low laitudes and leat at the poles, mainly because of the angle that the Sun's rays make with the Earth's surface: the higher the latitude, the lower the angle.

Conduction, convection and especially evaporation/precipitation are the principal means by which heat and water are exchanged across the air-sea interface. The oceanic evaporation/precipitation cycle contributes about one-quarter of the global heat budget. Aerosol production at the sea-surface is another important mechanism for the transfer of water (and salts) into the atmosphere.

Solar radiation penetrates no more than a few hundred metres into the oceans, and most is absorbed within the topmost 10m. Downward transfer of heat is mainly by mixing, as conduction is very slow (water is a very poor conductor of heat). Mixing by winds, waves and currents produces a mixed surface layer which can be 200-300 m thick or more in winter in mid-latitudes. Below this lies the permanent thermocline, across which temperature declines to about 5C and below which temperature decreases gradually to the bottom (typically between 0 C and 3 C). In mid-latitudes, as seasonal thermocline can develop during summer, above the permanent thermocline. There may also be diurnal thermoclines, at depths of 10-15m.

A temperature difference across the permanent thermocline can be utilized to generate electricity, using the principles upon which the domestic refrigerator is based. The main problem in this application is that of scale.

The long-term stability of the distribution of temperature within the ocean means that sections and profiles of average temperature do not change significantly from year to year. This stable thermal structure is maintained by the continuous three-dimensional motion of the global system of surface and deep currents.


The average salinity of seawater is close to 35 parts per thousand (0/00) by weight. Eleven major ions make up 99.9 per cent of the dissolved constituents: Cl-, Na+, SO42-, Mg2+, Ca2+, K+, HCO3-, Br-, H2BO3-, Sr2_ and F-, in that order. The relative proportions of elements in solution in seawater differ greatly from the proportions in crustal rocks, because of their different solubilities in the solutions formed during terrestrial weathering and sea-floor hydrothermal activity.

Salinity varies from place to place in the oceans, but the relative proportions of most major dissolved constituents (their ionic ratios) remain virtually constant. Evaporation and precipitation change the total salinity, but do not affect the constancy of composition.

Minor departures from constancy of composition in the open oceans result mainly from the intervention of biological processes, affecting principally Ca2+, and HCO3-. Major departures are the result of local conditions, chiefly in shallow nearshore waters and under anoxic conditions, but also where hydrothermal activity occurs. Some dissolved constituents are extracted commercially from seawater.

As in the case of temperature, the vertical and lateral distributions of salinity in the oceans do not change significantly from year to year, but the waters themselves are continually moving in a three-dimensional system of surface and deep currents. Surface salinities in the open oceans are greatest (up to 38) in tropical and subtropical latitudes, where evaporation exceeds precipitation. They are lower near the Equator (c. 35) and in high latitudes (c. 33-34), because of greater rainfall and melting ice and snowfall. In middle and low latitudes, there is a halocline from the base of the mixed surface layer to about 1 000 m depth, below which salinities are generally between 34.5 and 35.

Gravimetric measurements of salinity is difficult because of decomposition of some salts on heating to evaporation. Chemical measurements of salinity, based on titration to determine chlorinity, were standard until the 1960s, but have been almost entirely superseded by electrical conductivity methods. An empirically determined formula is used to convert conductivities, measured against a standard, into salinities.


Water masses are analagous to air masses. They can be identified by characteristic combinations of temperature and salinity and other properties. The boundaries of major upper water masses correspond approximately to the major wind-driven surface current systems. Subsurface water masses have comparatively narrow ranges of temperature and salinity, inerited from surface conditions in the source regions where they form and sink by virtue of their increased density. The movement of subsurface water masses in density-driven; this is the thermohaline circulation.

Temperature and salinity together control density, but pressure is also an important factor. Pressure ncreases almost linearly with depth in the oceans, because water is virtually incompressible. A pressure of about 1 atmosphere (105N m-2 or 1 000 mbar) is exerted by a 10 m water column. Air cools adiabatically as it rises, due to expansion as pressure falls. Water is heated adiabatically as a result of increased pressure and slight compression with depth. The potential temperature (θ) of a water sample is its measured in situ temperature after correction for adiabatic heating.

Sigma-t (σt) represents the density of seawater samples at atmospheric pressure, based on salinity and in situ temperature. Sigma-θ (σθ) represents the density of seawater samples at atmospheric pressure, based on salinity and potential temperature θ. T-S diagrams are contoured in values of σt and are used to identify water masses and to determine the textend of mixing between them. θ-S diagrams are contoured in values of σθ and are used in exactly the same way. Pycnoclines re regions where density increases rapidly with depth, and the main pycnocline coincides approximately with the permanent thermocline.

Conservative properties of seawater are those that are changed only by mixing, once the water has been removed from contact with the atmosphere and other external influences. Non-conservative properties are those that are changed by processes other than mixing. Temperature (potential temperature) and salinity are conservative properties; dissolved oxygen and nutrient concentrations are non-conservative properties.

Mixing occurs by both molecular diffusion and turbulent diffusion, the second of which is by far the more important; turbulent diffusion is much more rapid than molecular diffusion. The scale of horizontal mixin gis greater than that of vertical mixing in the oceans, partly because of their great width:depth ratio, and partly because density stratification inhibits vertical mixing.

In many parts of the ocean, there is a well-defined and graitationally stable microstructure, consisting of layers of water with fairly uniform T and S characteristics, separated by steep gradients of temperature and salinity. Small-scale processes that may operate to form and maintain the stratification are salt fingering, which results from double diffusion of heat and salt; and breaking of internal waves, the result of velocity shears along density interfaces.

Fronts are gently inclined boundaries which separate water of contrasted characteristics, typically well stratified on one side, mixed and hence more uniform on the other. They are common in shallow continental shelf waters, over the continental shelf and along continental margins; and are associated with oceanic current systems. They are characterized by sloping isopycnal surfaces (surfaces of constant density). Near-surface water can sink to greater depths along sloping isopycnals. Major fronts are normally tens of km across and slope down beneath the warmer and more stratified water, often at very small angles.

Eddies can develop wherever there is velocity shear, and are commonly associated with fronts and currents. Mesoscale eddies which form along major current systems (eg the Gulf Stream) are an important agent of large scale mixing in the oceans.


Light and all other forms of electromagnetic radiation travel at a speed of 3 x 108m s-1 in a vacuum (about 2.2 x 108m s-1 in seawater). Light travelling through water is subject to absorption and scattering, and its intensity decreases exponentially with distance from the source. Sunlight sufficient for photosynthesis cannot penetrate to more than about 200 m depth, and this defines the limit of the photic (or euphotic) zone, within which photosynthetic primary production can occur. The aphotic zone extends from the bottom of the photic zone to the sea-bed. Sunlight penetrates through only about the upper 1 000m of the aphotic zone; below that, the oceans are peermanently dark. The downwelling irradiance from sunlight or moonlight provides the non-directional (diffuse) light required for underwater illumination. Underwater vision requires directional light: light must travel direct from object to eye for a coherent image to be formed. Directional light is subject to greater attenuation than non-directional light.

Underwater visibility depends on contrast, which is a function partly of object brightness or relectivity and partly of attenuation with distance. Below depths of a few tens of metres, underwater light becomes virtually monochromatic, so contrast is mostly a matter of differences of light intensity rather than of colour. In lower parts of the aphotic zone, where many fish have bioluminescent organs (photphores), light is used in the same way as colour is used on land - for inter- and intraspecific recognition, camouflage, deterring predators, and so on.

Beam transmissometers are used to determine the attenuation coefficient (C) of directional light, and irradiance meters are used to determine the diffuse attenuation coefficient (K) of the non-directional downwelling irradiance. Mephelometers measure scattering and can be used to determine concentrations of particulate matter in the water. The Secchi disc is a simple piece of equipment for measurement of water clarity. By applying simple empirical equations, the measurements can be used to estimate visibility, attenuation coefficients, and the depth of the photic zone.

Water preferentially absorbs longer wavelengths of the electromagnetic specturm, which is why water appears blue. 'Yellow substances' and suspended particles absorb shorter wavelengths, so turbid water tends to look yellow, while productive ocean waters have the green colour of cholorophyll. In clear water, about 35% of incident blue-green light penetrates to 10 m depth. In turbid water, about 2% of yellow-green light penetrates to 10 m depth. Photosynthesis is inhibited in turbid waters.

Passive remote sensing of the oceans makes use of reflected and radiated visible, infrared and microwave radiation, to determine properties such as sea-surface temperature and water colour. Active remote sensing uses microwave imaging radar techniques to obtain information about the state of the sea-surface. Electromagnetic radiation cannot penetrate far through water, so remote sensing with the electromagnetic specturm can provide direct information only about surface or near-surface waters, depending on wavelength; and radio communication is all but impossible below the surface of the ocean.

Sound travels much more slowly than light through water but can travel much further, and so is used for remote sensing and communication in the oceans. Frequencies of interest in the oceans lie approximately in the 30 Hz to 1.5 MHz range. Sound intensities decrease with distance from the source because of two processes; (a) spreading loss, due to being spread out over (i) the surface of a sphere (loss proportional to distance2), or (ii) the surface of a cylinder (loss proportional to distance)_, as in the sound channel; and (b) attenuation, due to (i) absorption by the water and reactions involving its dissolved constituents, notably the dissociation of B(OH)3 and MgSo4 (attenuation increases as frequency increases, and high frquencies are very rapidly attenuated), and (ii) scattering, ie reflection by suspended particles.

The speed of sound in seawater, c, increases as the axial modulus of seawater increases, and decreases as the density increases; it is about 1 500 m s-1. A temperature rise of 1C causes an increase of about 3 m s-1. A salinity increase o 1 causes an increase of about 1.1 m s-1. A prssure increase equivalent to an increase in depth of 100 m causes an increase of about 1.8 m s -1. The speed of sound is at a minimum both at the surface and in the sound channel.

Sonar is used for depth determination, sea-bed mapping, and the location of objects, especially fish and submarines; many marine animals also make use of the technique. The reliability of echo-sounding depends partly upon the acoustic impedance: the higher the impedance contrast between water and the material of the object sought, the better the 'target' provided.

Sofar is used for longer-range location, and also for tracking, espcially of neutrally buoyant acoustic floats within and near the sound channel. To fix the position of Sofar devices reliably, variations of the speed of sound throughout the oceans must be known as accurately as possible. The axis of the sound channel lies between about 0.5 and 1.5 km depth throughout most of the oceans, between the latitudes of about 60 N and S. Poleward of these latitudes there is no sound channel.

In any acoustic receiving system there is background noise due to ambient sounds emanating from instrumental, physical and biological sources; and reverberation due to multiple reflections - scattering - by particles, and at the ocean boundaries.

Acoustic oceanography experiments make use of the effect of temperature and other properties on the speed and attnuation of sound in seawater, to detect and monitor relatively short-term changes within and between water masses on scales ranging from microstructure to whole ocean basins.


There are eleven major dissolved constituents of seawater, with concentrations greater than 1 part per million by weight (1 in 106). The remainder are minor and trace constituents, the boundary between the two being about 1 part per billion by weight (1 in 109). Most of the known chemical elements have been found in seawater solution; it is likely that all are present and will eventually be detected. The boundary between what constitutes dissolved and particulate matter is chosen for practical reasons at 0.45μm, but this does not always separate very fine colloidal particles from material truly in solution. Particulate matter (the seston) can remain in suspension for long periods because of turbulance.

Phosphate and nitrate are minor constituents and essential nutrients. They are extracted from surface water by photsynthesizing phytoplankton to make organic tissue. They may become totally depleted in surface waters where biological production is high, and are known as biolimiting constituents - they limit production because when they are exhausted, production ceases. When the organisms are consumed or when they die and decompose, the nutrients are returned to the water column (re-mineralized). The molar ratio of N to P in both seawater and organic tissue is about 15:1. Silica is also a biolimiting nutrient, but is used only to make the hard parts of some planktonic organisms. The skeletal remains dissolve only slowly as they sink into deep water after death, and can accumulate in sediments on the sea-floor.

Carbon is essential to all life, but is so abundant in seawater that its involvement in biological production makes only a small difference to its concentration. Calcium is used to make calcium carbonate skeletons and shells, but like carbon it is so abundant that its concentration is little affected. Carbon and calcium are bio-intermediate constituents. Bio-unlimited constituents are those whose concentrations are unaffected by biological activity.

The four principal atmospheric gases are nitrogen, oxygen, argon and carbon dioxide. Carbon dioxide is the most soluble gas in seawater, but occurs in solution mostly not as a gas but as bicarbonate and carbonate ions, in the main forms of total dissolved inorganic carbon (ΣCO2) in seawater. Concentrations of gases in surface waters are determined by their individual solubilities at the prevailing temperatuer and their atmospheric partial pressure. The solubility of gases decreases with increased temperature and salinity, and increases with pressure. Diffusion rates across the air-sea interface are increased in stormy weather, and dissolved gases are carried to deeper levels mainly by turbulent diffusion.

Oxygen is supersaturated in surface waters. The compensation depth at the base of the photic zone can be defined as the depth at which the amount of oxygen used (or carbon 'burnt' or dissipated) in respiration is equal to the amount of oxygen liberated (or carbon fixed) by photsynthesis. Below the photic zone, respiration uses up available oxygen and an oxygen minimum layer develops at a depth of a few hundred metres. Deep water is richer in oxygen because of cold well-oxygenated water sinking in polar regions.

The concentration of total dissolved inorganic carbon (ΣCO2) increases with depth because CO2 is used during photsynthesis (and formation of calcium carbonate) and released again during respiration (and dissolution of calcium carbonate); and the solubility of CO2 (and of CaCO3) is increased by increased pressure. Σ concentrations are an important factor in controlling the pH of seawater, which is mostly within the range of 7.7 +/1 0.2, being greater at the surface (less acid) that at depth (more acid). many minor gases in seawater are produced by biological activity and are supersaturated in surface layers, so they have a net flux from sea to air.

Rainwater is a dilute version of seawater, because aerosols carry marine salts into the atmosphere where they provide nuclei for rain formation. The dissolved constituents in river water result from rock weathering and are dominated by calcium and bicarbonate ions, whereas seawater and rainwater are dominated by sodium and chloride ions. The chloride and much of the sodium in river water come from recycled sea salt. Sodium balance calculations show that most of the dissolved consitutents of seawater can be accounted for by rock weathering, but some cannot, notably chloride, bromide and sulphate. These are excess volatiles whose main source is probably volcanic gases. Hydrothermal activity in the ocean basins is an important additional source of some constituents of seawater.

The ocean is generally beleived to be in a chemical steady state; rates of input and removal of dissolved constituents are in long-term balance. Residence times of dissolved constituents range from about 100 million years down to 1 000 years or less, and there is a very rough correlation between concentration and residence time. The residence time of water in the oceans is about 4 000 years and the average stirring (or mixing or turnover) time is of the order of 500 years.

Dissolved constituents in seawater are mostly in ionic form, and all ions are surrounded by a hydration sphere. Many ions form ion pairs or ionic complexes, in which the hydration spheres re more or less merged. Several of the major constituent ions for ion pairs; two of the most abundant pairs are MgSO4 and MgCO3.

Minor and trace elements in seawater generally have short residence times. Speciation is important, because several elements have more than one valency and can occur in more than one form. Different forms ahve different solubilities, and some may be removed by adsorption and scavenging, precipitated by redox reactions, or co-precipitated with insoluble complexes of other elements. Many are involved in biological processes and become greatly enriched in the tissues of marine organisms, so that organic-rich sediments may also be rich in trace elements. Some are biolimiting micro-nutrients; others show bio-intermediate behaviour.

Carbonite, alkalinity and pH. Some marine organisms use calcium carbonate to form their hard parts, which redissolve when the organisms die and sink into deep water. The depth at which dissolution begins is called the lysocline; the depth at which little or no calcium carbonate remains is called the carbonate compensation depth or CCD. Most dissolution occurs at the sea-bed, so the CCD is a sort of 'snowline'. CaCo3 may be in the form of the less common and less stable aragonite, or the more abundant and more stable calcite. The lysocline and CCD are deeper for calcite than for aragonite. The concentration of calcium is greater in deep than in surface waters because of the dissolution of calcium carbonate. So is that of total dissolved carbon (ΣCO2), partly because of the dissolution of calcium carbonate, but chiefly because of the decomposition of organic tissue through consumption and respiration. To a good first approximation:

[(ΣCO2)] = [HCO3-] + [CO32-] (6.10)

Alkalinity (A) is the combined negative charge due to bicarbonate and carbonate ions in seawter, expressed as molar concentrations. It is determined by titration with acid. It can also be defined as the excess of total major cations over total major anions (other than bicarbonate and carbonate), in molar 'charge-equivalent' terms.


  • A is changed only by formation and dissolution of calcium caronate (CACO3)
  • A is not changed by formation and decomposition of organic matter.
  • Formation and decomposition of organic matter changes only (ΣCO2).
  • Formation and dissolution of CaCO3 change [(ΣCO2)] as well as A.

(Figure 6.15 summarizes these four points).

[(ΣCO2)] increases more with depth than does A, so from

A- [(ΣCO2)] = [CO32-] (6.11)

the concentration of carbonate ions decreases with depth; and from:

[H+] = K [HCO3-]/[CO32-] (6.15)

deep water is generally more acid (lower pH) than surface water (higher pH); and skeletal remains formed of calcium carbonate dissolve as they sink into deep water.


The oceans are maintained in a chemical steady state by the global cycling of elements through (i) weathering and solution; (ii) precipitation, re-solution and sedimentation; (iii) subduction, uplift, volcanic activity; and (iv) back to weathering. Sediments and rocks of the marine environment have changed little in composition through geological time, which suggests that seawater composition has not changed greately and that elements have been cycled through the oceans in similar porportions and amounts since early in the Earth's history.

Seawater composition may initially have been governed partly by high CO2:O2 ratios in the early atmosphere, with greater bicarbonate and lower sulphate concentrations, and greater concentrations of reduced cationic species (eg Fe2+ and Mn2+). CO2:O2 ratios decreased as carbon was fixed and oxygen released by photsynthesis.

The decrease in CO2 with time was accompanied by increased solar lunimosity. The greenhouse effect of atmospheric CO2 deceased with time, thus helping to maintain an equable surface temperature. The decline in atmospheric CO2 is progressive, but probably irregular. The typical surface environment of the Earth is one of ice-free poles and higher average temperatures than at present.

Fluctuations in the concentration of atmospheric CO2 are strongly correlated with surface temperature, at least over the past 200 000 years, and probably much longer. Cycles of eccentricity of the Earth's orbit, changes in the angle of tilt of its axis of rotation and precession of the equinoxes (c. 110 000, c. 40 000 and c. 22 000 years respectively - the Milankovitch cycles), are probably the major causes of climatic variation. Atmospheric CO2 appears to reinforce trends of changing global temperature, but is probably not a major factor initiating them.

As CO2 levels now increase again because of human activity, the enhanced greenhouse effect is predicted to lead to melting ice-caps and rising sea-levels. Some excess CO2 is taken up by terrestrial and marine photosynthetic activity. The notion that the biosphere helps to maintain the Earth's surface in a life-supporting condition is the basis of the Gaia Hypothesis.

Ice core records suggest that the climate of the past 10 000 years has been untypically table. Much of the preceding 200 000 years was characterized by phases of abrupt and rapid climatic warming or cooling ('flickers'), especially in the Northern ('Land') Hemisphere. Temperatures seem to have risen or fallen several degrees on time-scales as short as decades. Global warming caused by increased concentrations of CO2 (and other greenhouse gases) in the atmosphere could 'flip' the climate system into one of these phases of rapid change.