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S330 Oceanography

Block 6: Waves, Tides and Shallow-Water Processes




Types of waves
Wind-generated waves in teh ocean
The fully developed sea
Wave height and wave steepness
Motion of water particles
Wave speed
Wave speed in deep and in shallow water
Assumptions made in surface wave theory
Propogation of wave energy
Attenuation of wave energy
Uses of wave energy
Wave refraction
Waves breaking upon the shore
Waves and currents
Giant waves
Satellite observations of waves


Variations in the lunar-induced tides
Interaction of solar and lunar tides
Prediction of tides by the harmonic method
Tides and tidal currents in shallow seas
Storm surges
Tides in rivers and estuaries
Tidal power


The supply of sediments to shelf seas and oceans
Variations in supply and distribution of sediments over time


Frictional forces and the boundary layer
Cohesive and non-cohesive sediments
Erosion of cohesive sediments and yield strength
The concept of shear velocity
The viscous sublayer
Velocity profiles in the sea
Shear velocity and the behaviour of non-cohesive sediments
Rates of sediment transport
Deposition of the bedload
Deposition of the suspended load


Beach profiles in relation to grain size and wave steepness
Orbital velocities and bed shear stress
Sediment movement by waves
Longshore sediment transport by wave-generated currents
Rip currents


Aggregation of sediment in estuaries
The estuarine continuum
Regions of freshwater influence
Sedimentation in estuaries
Estuaries in low latitudes
The Dynamic balance of estuaries


The deltaic continuum


Coastal and ocean currents
The effects of waves and of bioturbation

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Idealized waves of sinusoidal form have wavelength (length between successive crests), height (vertical difference between trough and crest), steepness (ratio of height to length), amplitude (half the wave height), period (length of time between successive waves passing a fixed point) and frequency (reciprocal of period). Water waves show cyclical variations in water level (displacement), from -a (amplitude) in the trough to +a at the crest. Displacement varies not only i space (one wavelength beween successive crests) but also in time (one period between crests at one location). Steeper waves depart from the simple sinusoidal model, and more closely resemble a trochoidal wave form.

Waves transfer energy across/through material without significant overall motion of the material itself, but individual particles are displaced from and return to, equilibrium positions as ech wave passes. Surface waves occur at interfaces bewen fluids, either because of relative movement between the fluids, or because the fluids are disturbed by an external force (eg wind). Waves occurring at interfaces between oceanic water layers are called internal waves. Wind-gnerated waves, once initiated, are maintained by surface tension and gravity, although only the latter is significant for water waves over 1.7 cm wavelength.

Most sea-surface waves are wind-generated. The stronger the wind, the larger the wave, so variable winds produce a range of wave sizes. A constant wind speed produces a fully developed sea, with waves of H1/3 (average height of highest 33% of the waves) characteristic of that wind speed. The Beaufort Scale relates sea state and H1/3 to the causative wind speed.

Water particles in waves in deep water follow almost circular paths, but with a small net forward drift. Path diameters at the surface correspond to wave heights, but decrease exponentially with depth. In shallow water, the orbits become flattened near the sea-bed. For waves in water deeper than 1/2 wavelength, wave speed equals wavelength/period (c = L/T) and is proportional to the square root of the wavelength (c = √gL/2π); it is unaffected by depth. For waves in water shallower than 1/20 wavelength, wave speed is proportional to the square root of the depth (c = √gd) and does not depend upon the wavelength. For idealized water waves, the three characteristics, c, L and T, are related by the equation c = L/T. In addition, each can be expressed in terms of each of the other two. For example, c = 1.56T and L = 1.56T2,

Waves of different wavelengths become dispersed, because those with greater wavelengths and longer periods travel faster then smaller waves. If two wave trains of similar wavelength and amplitude travel over the same sea area, they interact. Where they are in phase, displacement is doubled, whereas where they are out of phase, displacement is zero. A single wave train results, travelling as a series of wave groups, each separated from adjacent groups by an almost wave-free region. Wave group speed in deep water is half the wave (phase) speed. In shallowing water, wave speed approaches group speed, until the two coincide at depths less than 1/20 of the wavelength, where c = √gd.

Wave energy is proportional to the square of the wave height, and travels at the group speed. Wave power is rate of supply of wave energy and so it is wave energy multiplied by wave (or group) speed, ie it is wave energy propogated per second per unit length of wave crest (or wave speed multiped by wave energy per unit area). Total wave power is conserved, so waves entering shallowing water and/or funnelled into a bay or estuary (see also below) increase in height as their group speed falls. Wave energy has been successfully harnessed on a small scale, but large-scale utilization involves enrironmental and navigational problems, and huge capital outlay.

Dissipation of wave energy (attenuation of waves) results from white-capping, friction between water molecules, air-resistance, and non-linear wave-wave interaction (exchange of energy between waves of differing frequencies). Most attenuation takes place in and near the storm area. Swell waves are storm-generated waves that have travelled far from their place of origin, and are little affected by wind or by shorter, high-frequency waves. The wave energy associated with a given length of wave crest decreases with increasing distance from the storm, as the wave energy is spread over an ever-increasing length of wave front.

Waves in shallow water may be refracted. Variations in depth cause variations in speed of different parts of the wave crest; the resulting refraction causes wave crests to become increasingly parallel with bottom contours. The energy of refracted waves is conserved, so converging waves tend to increase, and diverging waves to diminish, in height. Waves in shallow water disipate energy by frictional interaction with the sea-bed, and by breaking. In general, the steeper the wave and the shallower the beach, the further offshore dissipation begins. Breakers from a continuous series from steep spilling types to long-period surging breakers.

Waves propagating with a current have diminished heights, whereas a counter-current increases wave height, unles current speed exceeds half the wave group speed. If so, waves no longer propagate, but increase in height until they become unstable and break. Tsunamis are caused by earthquakes or by slumping of sediments, and their great wavelength means their speed is always governed by the ocean depth. Wave height is small in the open ocean, but can become destructively large near the shore. Seiches (standing waves) are socillations of water bodies, such that at antinodes there are great variations of water level but little lateral water movement, whereas at nodes the converse is true. The period of oscillation is proportional to basin length and inversely proportional to the square root of the depth. A seiche is readily established when the wavelength of incoming waves is four times the length of the basin.

Waves are measured by a variety of methods, eg pressure guages on the sea-floor, accelerometers in buoys on the sea-surface, and via remote-sensing from satellites.


Tides are long-period waves, generated by gravitational forces exerted by the Moon and Sun upon the oceans. They behave as shallow-water waves because of their very long wavelengths. Tidal currents are the horizontal water movements corresponding to the rise and fall (flood and ebb) of the tide.

A centrifugal force, directed away from the Moon, results from the Earth's (eccentric) rotation (period 27.3 days) around the Earth-Moon centre of mass, which is within the Earth. This centrifugal force is exactly balanced in total bu the gravitational force exerted on teh Earth by the Moon. However, gravitational force exceeds centrifugal force on the 'Moon-side' of Earth, resulting in tide-producing forces directed towards the Moon, whereas on the other side of the Earth centrifugal force exceeds gravitational force, resulting in tide-producing forces directed away from the Moon.

Tractive forces (horizontal components of tide-producing forces) are maximal on two small circles either side of the Earth, and produce two (theoretical) equilibrium tidal bulges - one directed towards the Moon, and the other directed away from it. As the Earth rotates with respect to the Moon 9with a period of 24 hours 50 minutes), the equilibrium tidal bulges would need to travel in the opposite direction (relative to the surface of the rotating Earth) in order to maintain their positions relative to teh Moon. The elliptical orbit of the Moon about the Earth causes variation in the tide-producing forces of up to 20% from the mean value.

With the Moon overhead at the Equator, the equilibrium tidal bulges would be in the same plane as the Equator, and at all points the two bulges would theoretically cause two equal high tides daily (equatorial tides). The Moon has a declination of up to 28.5deg either side of the Equator, and when the plane of the tidal bulges is offset with respect to the Equator, there are two unequal, or tropic, tides daily. The declination varies over a 27.2 day cycle.

The Sun also produces tides which show inequalities related to the Sun's declination (up to 23.4 deg either side of the Equator), and vary in magnitude due to the elliptical orbit of the Earth around the Sun. The Sun's tide-producing force has about 46% of the strength of the Moon's Solar tides combine with and interact with lunar tides. When Sun and Moon are in syzygy, the effect is additive, giving large-ranging spring tides; but when Sun and Moon are in quadrature, tidal ranges are small (neap tides). The full cycle (a lunar month), includes two neap andn two spring tides, and takes 29.5 days.

Tidal speed is limited to about 230 ms-1 in the open oceans (less in shallower seas) and land masses constrain tidal flow. Water masses have intertia and experience friction with coasts and sea=bed, so they do not repond instantaneously to tractive forces. The Coriolis force, and constraining effects of land masses, combine to impose amphidromic systems upon tides. High tidal crests circulate (as Kelvin waves) around amphidromic points which show no change in tidal level, ie, tidal range increases with distance from an amphidromic point. Amphidromic systems tend to rotate in the opposite direction to the deflection caused by the Coriolis force.

The actual tide is made up of many constituents (partial tides), each corresponding to the period of a particular astronomical motion involveing Earth, Sun or Moon. Partial tides can be determined from tidal measurements made over a long time at individual locations, and the results used to computer future tides. Actual tides are classified by the ratio (F) of the summed amplitudes of the two main diurnal constituents to the summed amplitedes of the two main semi-diurnal constituents.

Tidal rise an dfall are produced by lateral water movements called tidal currents. Tidal current vectors typically display 'tidal ellipses' rather than simple to-and-from motions.

Areas of low atmospheric pressure cause elevated sea-levels, whereas high pressure depresses sea-level. A strong wind can hold back a high tide or reinforce it. Storm surges are caused by large changes in atmospheric pressure and the associated strong winds. Positive storm surges may result in catastrophic flooding.

In estuaries, the tidal crest travels faster than the tidal trough because speed of propagation depends upon water depth; hence the low water to high water interval is shorter than that from high water to low water. Tidal bores develop where tides are constrained by narrowing estuaries and the wave-front is forced by the rising tide to travel faster than the depth-determined spped of a shallow-water wave. Where tidal ranges are large and the water can be trapped by dams, the resultant heads of water can be used for hydro-electric power generation.


The sediments of most continental shelves are predominantly terrigeonous and consist of rock fragments, quartz sands and clay-rich muds. Where terrigenous sediments are scarce or absent, carbonate sediments of biogenic or inorganic origin may occur, especially in low latitudes.

River transport is the most important means of bringing terrigenous sediments to the ocean margins, but ice-transport, wind transport and volcanic eruptions may be important locally.

Periodic falls in sea-level during the Quaternary resulted in the deposition of river and glacial sedimetns on areas of the continental shelf which are today covered by the sea. These relict sediments are now being reworked by waves and tidal currents. Changes of sea-level lead to changes in, and migrations of, shallow-water environments on time-scales of hundreds to thousands of years.


Friction between flowing water and the sea-bed generates a boundary layer in which turbulent flow dominates, except very close to the bed. Movement of sediment (erosion) occurs when the shear stress generated by the frictional force of water flowing over the sediment overcomes the force of gravity acting on the sediment grains and the friction between the grains and the underlying bed. Shear stress is proportional to the square of the mean current speed (and to the density of the water). Movement of grains of a given size begins when the shear stress at the bed reaches a critical value (critical shear stress). Particles larger than about 0.1-0.2 mm in diameter are initially moved as bedload, then lifted intermittently into suspension, and as current speeds increase further they are finally lifted permanently into suspension. Grains smaller than obout 0.1 mm are lifted directly into suspecsion as soon as the critical shear stress is reached.

Cohesive sediments contain a high proportion of fine-grained clay minerals and are more diffcult to erode than non-cohesive sediments, which consist mostly of qua\rtz grains. For cohesive sediments, the smaller the particle size, the greater the current speed required to erode them. The resistance of muds to erosion is assessed by their yield strength. Once in suspension, clay particles can be transported for long distances by currents that would be much too weak to erode them.

Shear stress is proportional also to the velocity gradient in the boundary layer and to the viscosity of the water. When current speed is plotted against the height above se-bed (as the vertical axis) on a log-linear graph, the inverse velocity gradient d log z/du is linear. The slope of the line is used to calculate the shear velocity and the intercept of the line with the depth axis gives a measure of the bed roughness length (z0) which increases as the sediment grain size increases; roughness length will also be greater if there are bed forms such as sand ripples.

When water flows over a smooth (very fine-grained) bed, the lowermost wter layer appears to flow in a laminar fasion, forming a viscous sublayer only a few millimetres thick, which decreases in thickness as flow speed increases (and as viscosity decreases with increasing temperature). suspended particles which settle into the sublayer are subjected only to laminar flow and are soon deposited. When water flows over a coarse-grained bed, or at high current speeds, grains protrude through the viscous sublayer, breaking it down, and turbulent flow extends to the bed, giving rise to greater potential for sediment movement. In the marine envrionment, values of shear velocity (and hence of bed shear stress) may be overestimated or underestimated because of: decelerating or accelerating tidal currents (respectively); the time taken for a velocity profile to adjust to a new bed roughtness; and the concentration of suspended sediment close to the bed. Erosion and transport of sediment is initiated mainly by cycles of downward sweeps and upward bursts that result from turbulent motions in the boundary layer.

The rate of bedload transport is proportional to the cube of he shear velocity (see equation p 124) and hence also to the cube of the average current speed, but it is difficult to measure directly in the sea. The rate of suspended load transport can be calculated form the product of the current velocity and the sediment concentration.

Deposition of bedload begins when the current speed falls so that the shear stress at the bed is only a little below the critical shear stress required to start the sediment moving. The rate of deposition of the bedload is proportional to the reduction in the cube of the average current velocity. The rate of deposition of the suspended load varies according to sediment grain size. The settling lag of fine suspended particles means that they may reach the bed well after they began to settle from suspension. The rate of deposition of grains of a given size in the suspended sediment load depends both on the vertical distribution of sediment concentration above the bed, and on the settling velocity of th grains.

Raised sediment features on the sea-bed are called bed forms. Bed forms produced by waves are symmetrical, those formed by currents are asymmetrical. The type of bed form depends mainly on current speed and ripples form at relatively slow current speeds and where sediment is finer than about 0.6mm grain diameter. As current speeds increase or where sediments are coarser-grained, larger-scale megaripples are formed; and sand-waves develop where sand is abundant. At current speeds of more than about 1 ms-1, bed forms such as sand-bandks and sand ribbons develop parallel, rather than transverse to, the current flow. At still higher current speeds, erosional features such as furrows and scour hollows develop.


Beach profiles are controlled by the influence of waves, tidal range, and sediment particle size. The wave zones are the swash zone, the surf zone and the breaker zone. Steep beaches are characterized by berms, shallow ones by swash bars and runnels. Longshore bars may develop seawards of the intertidal zone. Coarse-grained sediments lead to steep beaches becazuse water is readily lost through percolation and the abckwash is too weak to move much of the sediment that has been transported up the beach face by the swash. Conversely, fine-grained sediments lead to shallow beaches. Small gentle waves and swell waves tend to build up beaches and steep storm waves tend to tear them down and glatten them.

Water particles in shallow-water waves follow orbitgal paths which become progressively flattened towards the sea-bed. I shallow water, the maximum horizontal orbital velocity and shear stress at the bed increase as the wave height increases and as water depth decreases. The conditions which determine sediment movement for a given grain size may be achieved from many different combinations of wave height, wave period and water depth. The orbital velocity necessary to intiate sediment movement (threshold velocity) for a given grain size increases as the wave period increases.

Beneath a wave, sediment is moved landwards as the crest passes and seawards as the trough passes. Strong shoreward velocities move both coarser sediment (as bedload) and finer sediment (as suspended load) landwards. Weaker seaward velocities, of longer duration, move only the finer bedload and suspended load seawards. Coupled with the effects of percolation, this leads to a net movement of coarse sediment landwards and fine sediment seawards.

Straight-crested, symmetrical ripples form as a result of the oscillatory water movement beneath waves. Rhomboid patterns are formed by fast-flowing backwash. Larfer sedimentary structures on some beaches include cusps formed at the high tide mark.

Wave-induced longshore currents are generated when waves break obliquely to the shoreline. These currents, and the zig-zag movement of swa\sh and backwash on steep beahces, move sediment along the shoreline, and can also lead to the geneation of rip currents. Rip currents develop also as a consequence of horizontal pressure gradients between regions having wave set-up of different heights. Convergences of resulting longshore currents lead to the return of water seawards in narrow fast-flowing (rip) currents.

The wave power available for longshore sediment transport can be calculated from the wave group speed, average wave height and the angle the wave crest makes with the shoreline. The rate of sediment transport along the shoreline can be estimated using the wave power. Most net sediment transport occurs when movement by currents is enhanced by wave action. Wave action lifts sediment into suspension whee it is transported by currents which, by themselves, may be unable to lift sediment off the sea-bed.

Drawing up sediment budgets and identifying coastal sediment transport cells can help to quantify the dynamic equilibrium of the coastal zone, ie whether there is net erosion or net deposition (accretion) of a coastline. Such studies are useful to assess teh likely or acual impace of coastal engineering or construction works. Attempts to alter the hynamic equailibrium along one stretch of coastline (eg by erosion-prevention measures) are likely to disturb the equilibrium elsewhere, resulting in accelerated (and unwanted) erosion and/or accretion elsewhere. It is probably wiser to let Nature take its course.


Estuaries are tidal inlets at the mouths of rivers wshere mixing of freshwater and seawater occurs. They are ephemeral features on geological time-scales, and most are now slowly being infilled with sediment. They are characterized by channels and intertidal flats. There is a progression of sediment grain size towards the estuary shore; from sands in the channels, through sands and silts (with some muds) on the main intertidal flats, to muds on the high tidal flats, which are only submerged when tidal currens are weak at slack water. Accretion of tidal mud flats is promoted by the cohesive nature of muddy sediments, by settling lag, and by colonization of the mud flats by algae and eventually by land plants, leading to the formation of salt-marshes at mid-to high latitudes and mangrove swamps at low latitudes.

Find sediment is deposited through aggregation into larger particles with higher settling velocites. Aggregation occurs mainly by flocculation in saline water, aided by turbulence in the water column, and also by biological processes (formation of faecal pellets and 'fluffy' aggregates of organic material). Cation exchange reactions take place between seawater and clay minerals, which can also adsorb heavy metals from solution in contaminated waters.

Estuaries range from strongly stratified to well-mixed, depending upon the relative magnitudes of tidal curents and river flow in the main channels. Salt-wedge (well-stratified) estuaries develop in virtually tideless seas, and are dominated by seaward flow of freshwater at the surface, with only minor landward movement (residual flow) of salt water at the bed. Current shear at the halocline leads to entrainment of salt water up into the freshwaater layer. Partially mixed (moderately stratified) estuaries develop where there is a moderate tidal range. Greater mixing of fresh and salt water occurs becuase of turbulence, both at the be and at the freshwater-seawater interface, and there is significant movement of water both seawards at the surface and landwards at the bed. Well-mixed (unstratified) estuaries develop where the tidal range is high. There is very little variation in salinity with depth, though in wide estuaries (especially if they are well-mixed) there can be lateral salinity gradients because river and tidal flows are on opposite sides of the estuary (as a result of the Coriolis effect) and ther is a horizontal residual circulation. Even so, the mean velocity is seawards at all depths.

An estuary can exhibit different degrees of stratification and mixing between spring and neap tides and/or as a consequence of changes in river low: high river discharge promotes stratification, low discharge promotes mixing. The upstream limit of the landward movement of salt water near the bed is called the null point. It occurs at salinities of between about 0.1 and 5 (parts per thousand), depending upon circumstances, and moves up and down the main channel with the tides.

The plume of brackish water that flows from the estuary mouth can affect offshore waters over consierable areas, and regions of freshwater influence (ROFIs) can extend up to hundreds of kilometres from the estuary mouth, depending upon the magnitude of the river discharge. Seawater is mixed and entrained into the plume at th ebase and along the margins, where there are convergent fronts. Where tidal ranges are large, tidal intrusion fronts form in the rising tide at the mouths of some smaller estuaries; and in some well-mixed estuaries longitudinal fronts are observed, the result of transverse surface water movements caused by cross-estuary gradients of salinity (and hence density).

A turbidity maximum develops near the null point, because sediment is carried into it both by th eriver flow and by the landward flow of salt water near the bed, aided by flocculation near the null point. The turbidity maximum also moves up and down the river with the tides, and is the source of most of the muds deposited on the high tidal flats. It tends to be also enhanced during spring tides and/or at times of low river discharge, but is reduced during neap tides and/or at times of high river discharge. In some esutaries, high concentrations of fluid mud may form near the bed during neap tides, to be susequently dispersed by the spring tides. Most estuaries are net accumulators of sediment since they are supplied with material from both the river and the sea. The landward movement of sediment is aided by the asymmetry of tidal flows in estuaries.

Negative estuarine circulation can develop in arid regions, where very high evaporation rates at the head of the estuary lead to sinking of dense hypersaline water, and a landward flow of seawate of normal salinity at the surface to replace it. In such estuaries, sands may be deposited from the bedload at the head of the estuary, while the fine sediments are carried seawards in suspension by the hypersaline flow at the bed.

Lagoons commonly form in the shelter of sand or gravel spits formed by longshore transport. Most are shallow and well mixed, and have only narrow outlets to the sea, so that tidal influences and wave activity are relatively weak. Breaching of spits by wave action, or isolation of longshore bars by rising sea-level, can lead to the formation of barrier islands, behind which wave action is limited. If the tidal range is moderate to large, tidal flats (similar to those occurring in estuaries) can form behind barrier islands. In low latitudes, colonisation of tidal flat sediments by carbonate-secreting algae leads to accumulations of layers of algal mats. Along arid coastlines, where evaporation is high, salt flats (sabkhas) develop, and where terrigenous sediment input is negligible, carbonate muds can accumulate.


Deltas are coastal accumulations of river-borne sediments which accrete when sediment discharge is too large to be dispersed by tidal currents and wave action. The main components of the delta are: the delta plain above sea-level; the shallow water and shoreline region of the delta front, dminated by sands; and the deeper water prodelta, dominated by silts and clays. The shapes of most deltas are controlled by the interaction of fluvial, tidal and wave processes, and a deltaic continuum can be identified on the basis of the relative importance of these three processes, analogous to the estuarine continuum of chapter 6.

Where tidal range is small and wave action is weak, the river discharge plume spreads seawards as a thin layer over the denser seawater, establishing a density stratification. Mixing and entrainment of salt water into the base of the plume gnerates a landward flow i the salt wedge beneath it. Bedload deposition of sediment forms a delta bar and near-parallel subaqueous levees. When the river discharge is high, vigourous turbulent mixing disrupts the stratification, and increased sediment transport results in rapid advance of the delta front.

With increasing tidal range, there is stronger turbulent mixing along the sides and base of the plume, and the water column resembles that of partially mixed estuaries. Where the tidal range is large (>4m), strong tidal currents inhibit development of density stratification, and the wate column at distribuary mouths is well-mixed. Sediment movement occurs both up and down the distributary channels, with formation of sediment ridges within distributary mouths, parallel to the direction of tidal ebb and flow, producing a ragged outline to the delta.

Where wave energy is high, the outflowing freshwater from the delta distributaries behaves as a counter-current, slowing down and steepening the approaching waves and causing them to break in deeper wte than usual. Waves are also refracted so that wave energy is concentrated on the freshwate plume. Both processes lead to vigorous mixing of freshwater and salt water, rapid deceleration of the freshwater flow and deposition of sediments. Wave action reworks the sediments and moves the coarsser sediments landwards to form swash bars and beaches, creating a straight shoreline with only minor protuberances at the distributary mouth(s).

Human interference with deltas, such as dammin gof rivers inland, agricultural acrivity or hydrocarbon exploitation on the delta plain, can disrupt the natural patters of water flow and of sediment transport and deposition. Sediment supplies to the delta front may be reduced, and as subsidence continues there is likely to be erosion and retreat of the delta front, instead of deposition and advance. The discharge of agricultural and industrial wastes can significantly contaminate the coastal waters.


Shelf seas extend from just below low tide level out ot he shelf break, and shelves are underlain by considerable thicknesses of sediment. Both modern and relict (glacial) sediments form the sea-bed and are being reworked by waves and currents. The principal factors determining sediment distribution in shelf seas are the residual current field and the rateof supply of sediments from rivers and coastal erosion.

Coastal geostrophic currents develop where offshore pressure gradients are set up in response to winds and/or heavy rain. Longshore winds lead to Ekman transport of surface wate offshore or onshore, and hece to upwelling or downwelling respectively, influencing both biological production and sediment transport.

Sediment transport in shelf seas is determined by the residual current field, which may not be the same at the surface and the bed. In general, the stronger the currents at the sea-bed, the coarser the sediments, and reworking by waves tends to move finer sediments from shallow to deep water. Waves have a large effect during storms, and can aid sediment transport by lifting sediment into suspension to be carried away by currents. Bioturbation by burrowing animals can locally also resupend significant amounts of sediment. Bedload partings, where sediments are transported in opposite directions, occur at zones of divergence in near-bed currents. Sea-bed resources in shelf seas include aggregates (sands and gravels, and shell deposits), placers (heavy mineral concentrations) and phosphorites, all resulting mainly from reworking or chemical alteration of relict sediments. Freshwater (artesian) springs may occur at the sea-bed if the underlying geological structure is suitable.

A rough zonation can be discerned in shelf seas, between shore and shelf break. The 'zones' may not always be developed in all shelf seas and their relative widths and positions can vary with tidal range and season. Regions of freshwater influence (ROFIs) off estuaries or deltas, may be stratified or mixed depending on wind and tidal state. A tidal mixing from separates the shallower and vertically mixed marine water column seward of the ROFI from the deeper and seasonally stratified waters of teh mid-outer shelf, and its position varies according to water depth, tidal range and season. Near the shelf break, there may be localized disruption of the seasonal stratification by tidally generated internal waves.