Internal tide

Internal tides are generated as the surface tides move stratified water up and down sloping topography, which produces a wave in the ocean interior. So internal tides are internal waves at a tidal frequency. The other major source of internal waves is the wind which produces internal waves near the inertial frequency. When a small water parcel is displaced from its equilibrium position, it will return either downwards due to gravity or upwards due to buoyancy. The water parcel will overshoot its original equilibrium position and this disturbance will set off an internal gravity wave. Munk (1981) notes, "Gravity waves in the ocean's interior are as common as waves at the sea surface-perhaps even more so, for no one has ever reported an interior calm." [1]

Simple explanation

Figure 1: Water parcels in the whole water column move together with the surface tide (top), while shallow and deep waters move in opposite directions in an internal tide (bottom). The surface displacement and interface displacement are the same for a surface wave (top), while for an internal wave the surface displacements are very small, while the interface displacements are large (bottom). This figure is a modified version of one appearing in Gill (1982). [2]

The surface tide propagates as a wave, in which water parcels in the whole water column oscillate in the same direction at a given phase (i.e., in the trough or at the crest, Fig. 1, top). At the simplest level, an internal wave can be thought of as an interfacial wave (Fig. 1, bottom). If there are two levels in the ocean, such as a warm surface layer and cold deep layer separated by a thermocline,then motions on the interface are possible. The interface movement is large compared to surface movement. The restoring force for internal waves and tides is still gravity but its effect is reduced because the densities of the 2 layers are relatively similar compared to the large density difference at the air-sea interface. Thus larger displacements are possible inside the ocean than at the sea surface.

Tides occur mainly at diurnal and semidiurnal periods. The principal lunar semidiurnal constituent is known as M2 and generally has the largest amplitudes. (See external links for more information.)


The largest internal tides are generated at steep, midocean topography such as the Hawaiian Ridge, Tahiti, the Macquarie Ridge, and submarine ridges in the Luzon Strait. [3] Continental slopes such as the Australian North West Shelf also generate large internal tides. [4] These internal tide may propagate onshore and dissipate much like surface waves. Or internal tides may propagate away from the topography into the open ocean. For tall, steep, midocean topography, such as the Hawaiian Ridge, it is estimated that about 85% of the energy in the internal tide propagates away into the deep ocean with about 15% of its energy being lost within about 50 km of the generation site. The lost energy contributes to turbulence and mixing near the generation sites. [5] [6] It is not clear where the energy that leaves the generation site is dissipated, but there are 3 possible processes: 1) the internal tides scatter and/or break at distant midocean topography, 2) interactions with other internal waves remove energy from the internal tide, or 3) the internal tides shoal and break on continental shelves.

Propagation and dissipation

Figure 2: The internal tide sea surface elevation that is in phase with the surface tide (i.e., crests occur in a certain spot at a certain time that are both the same relative to the surface tide) can be detected by satellite (top). (The satellite track is repeated about every 10 days and so M2 tidal signals are shifted to longer periods due to aliasing.) The longest internal tide wavelengths are about 150 km near Hawaii and the next longest waves are about 75 km long. The surface displacements due to the internal tide are plotted as wiggly red lines with amplitudes plotted perpendicular to the satellite groundtracks (black lines). Figure is adapted from Johnston et al. (2003).

Briscoe (1975)noted that “We cannot yet answer satisfactorily the questions: ‘where does the internal wave energy come from, where does it go, and what happens to it along the way?’” [7] Although technological advances in instrumentation and modeling have produced greater knowledge of internal tide and near-inertial wave generation, Garrett and Kunze (2007) observed 33 years later that “The fate of the radiated [large-scale internal tides] is still uncertain. They may scatter into [smaller scale waves] on further encounter with islands[8] [9] or the rough seafloor [10] , or transfer their energy to smaller-scale internal waves in the ocean interior [11] ” or “break on distant continental slopes [12]”. [13] It is now known that most of the internal tide energy generated at tall, steep midocean topography radiates away as large-scale internal waves. This radiated internal tide energy is one of the main sources of energy into the deep ocean, roughly half of the wind energy input .[14] Broader interest in internal tides is spurred by their impact on the magnitude and spatial inhomogeneity of mixing, which in turn has first order effect on the meridional overturning circulation [3] [14] .[15]

The internal tidal energy in one tidal period going through an area perpendicular to the direction of propagation is called the energy flux and is measured in Watts/m. The energy flux at one point can be summed over depth- this is the depth-integrated energy flux and is measured in Watts/m. The Hawaiian Ridge produces depth-integrated energy fluxes as large as 10 kW/m. The longest wavelength waves are the fastest and thus carry most of the energy flux. Near Hawaii, the typical wavelength of the longest internal tide is about 150 km while the next longest is about 75 km. These waves are called mode 1 and mode 2, respectively. Although Fig. 1 shows there is no sea surface expression of the internal tide, there actually is a displacement of a few centimeters. These sea surface expressions of the internal tide at different wavelengths can be detected with the Topex/Poseidon or Jason-1 satellites (Fig. 2). [9] Near 15 N, 175 W on the Line Islands Ridge, the mode-1 internal tides scatter off the topography, possibly creating turbulence and mixing, and producing smaller wavelength mode 2 internal tides. [9]

The inescapable conclusion is that energy is lost from the surface tide to the internal tide at midocean topography and continental shelves, but the energy in the internal tide is not necessarily lost in the same place. Internal tides may propagate thousands of kilometers or more before breaking and mixing the abyssal ocean.

Abyssal mixing and meridional overturning circulation

The importance of internal tides and internal waves in general relates to their breaking, energy dissipation, and mixing of the deep ocean. If there were no mixing in the ocean, the deep ocean would be a cold stagnant pool with a thin warm surface layer. [16] While the meridional overturning circulation (also referred to as the thermohaline circulation) redistributes about 2 PW of heat from the tropics to polar regions, the energy source for this flow is the interior mixing which is comparatively much smaller- about 2 TW. [14] Sandstrom (1908) showed a fluid which is both heated and cooled at its surface cannot develop a deep overturning circulation. [17] Most global models have incorporated uniform mixing throughout the ocean because they do not include or resolve internal tidal flows.

However, models are now beginning to include spatially variable mixing related to internal tides and the rough topography where they are generated and distant topography where they may break. Wunsch and Ferrari (2004) describe the global impact of spatially inhomogeneous mixing near midocean topography: “A number of lines of evidence, none complete, suggest that the oceanic general circulation, far from being a heat engine, is almost wholly governed by the forcing of the wind field and secondarily by deep water tides... The now inescapable conclusion that over most of the ocean significant ‘vertical’ mixing is confined to topographically complex boundary areas implies a potentially radically different interior circulation than is possible with uniform mixing. Whether ocean circulation models... neither explicitly accounting for the energy input into the system nor providing for spatial variability in the mixing, have any physical relevance under changed climate conditions is at issue.” There is a limited understanding of “the sources controlling the internal wave energy in the ocean and the rate at which it is dissipated” and are only now developing some “parameterizations of the mixing generated by the interaction of internal waves, mesoscale eddies, high-frequency barotropic fluctuations, and other motions over sloping topography.”

Internal tides at the beach

Scripps internal wave T
Figure 3: The internal tide produces large vertical differences in temperature at the research pier at the Scripps Institution of Oceanography. The black line shows the surface tide elevation relative to mean lower low water (MLLW). Figure provided by Eric Terrill, Scripps Institution of Oceanography with funding from the U.S. Office of Naval Research

Internal tides may also dissipate on continental slopes and shelves [12] or even reach within 100 m of the beach (Fig. 3). Internal tides bring pulses of cold water shoreward and produce large vertical temperature differences. When surface waves break, the cold water is mixed upwards, making the water cold for surfers, swimmers, and other beachgoers. Surface waters in the surf zone can change by about 10 °C in about an hour.

Internal tides, internal mixing, and biological enhancement

Internal tides generated by tidal semidiurnal currents impinging on steep submarine ridges in island passages, ex: Mona Passage, or near the shelf edge, can enhance turbulent dissipation and internal mixing near the generation site. The development of Kelvin-Helmholtz instability during the breaking of the internal tide can explain the formation of high diffusivity patches that generate a vertical flux of nitrate (NO3) into the photic zone and can sustain new production locally. [18] [19] Another mechanism for higher nitrate flux at spring tides results from pulses of strong turbulent dissipation associated with high frequency internal soliton packets. [20] Some internal soliton packets are the result of the nonlinear evolution of the internal tide.

See also


  1. ^ Munk, W. (1981). B. A. Warren; C. Wunsch (eds.). "Internal Waves and Small-Scale Processes". Evolution of Physical Oceanography. MIT Press: 264–291.
  2. ^ Gill, A. E. (1982). Atmosphere-ocean dynamics. Academic. p. 662. ISBN 978-0-12-283522-3.
  3. ^ a b Simmons, H. L.; B. K. Arbic & R. W. Hallberg (2004). "Internal wave generation in a global baroclinic tide model". Deep-Sea Research Part II. 51 (25–26): 3043–3068. Bibcode:2004DSR....51.3043S. CiteSeerX doi:10.1016/j.dsr2.2004.09.015.
  4. ^ Holloway, P. E. (2001). "A regional model of the semidiurnal tide on the Australian North West Shelf". J. Geophys. Res. 106 (C9): 19, 625–19, 638. Bibcode:2001JGR...10619625H. doi:10.1029/2000jc000675.
  5. ^ Carter, G. S.; Y. L. Firing; M. A. Merrifield; J. M. Becker; K. Katsumata; M. C. Gregg; D. S. Luther; M. D. Levine & T. J. Boyd (2008). "Energetics of M2 Barotropic-to-Baroclinic Tidal Conversion at the Hawaiian Islands". J. Phys. Oceanogr. 38 (10): 2205–2223. Bibcode:2008JPO....38.2205C. doi:10.1175/2008JPO3860.1.
  6. ^ Klymak, J. M.; M. C. Gregg; J. N. Moum; J. D. Nash; E. Kunze; J. B. Girton; G. S. Carter; C. M. Lee & T. B. Sanford (2006). "An Estimate of Tidal Energy Lost to Turbulence at the Hawaiian Ridge". J. Phys. Oceanogr. 36 (6): 1148–1164. Bibcode:2006JPO....36.1148K. doi:10.1175/JPO2885.1.
  7. ^ Briscoe, M. (1975). "Introduction to a collection of papers on oceanographic internal waves". J. Geophys. Res. 80 (3): 289–290. Bibcode:1975JGR....80..289B. doi:10.1029/JC080i003p00289.
  8. ^ Johnston, T. M. S.; M. A. Merrifield (2003). "Internal tide scattering at seamounts, ridges and islands". J. Geophys. Res. 108. (C6) 3126 (C6): 3180. Bibcode:2003JGRC..108.3180J. doi:10.1029/2002JC001528.
  9. ^ a b c Johnston, T. M. S.; P. E. Holloway & M. A. Merrifield (2003). "Internal tide scattering at the Line Islands Ridge". J. Geophys. Res. 108. (C11) 3365 (C11): 3365. Bibcode:2003JGRC..108.3365J. doi:10.1029/2003JC001844.
  10. ^ St. Laurent; L. C.; C. Garrett (2002). "The Role of Internal Tides in Mixing the Deep Ocean". J. Phys. Oceanogr. 32 (10): 2882–2899. Bibcode:2002JPO....32.2882S. doi:10.1175/1520-0485(2002)032<2882:TROITI>2.0.CO;2. ISSN 1520-0485.
  11. ^ MacKinnon, J. A.; K. B. Winters (2005). "Subtropical catastrophe: Significant loss of low-mode tidal energy at 28.9 degrees". Geophys. Res. Lett. 32 (15): L15605. Bibcode:2005GeoRL..3215605M. doi:10.1029/2005GL023376.
  12. ^ a b Nash, J. D.; R.W. Schmitt; E. Kunze & J.M. Toole (2004). "Internal tide reflection and turbulent mixing on the continental slope". J. Phys. Oceanogr. 34 (5): 1117–1134. Bibcode:2004JPO....34.1117N. doi:10.1175/1520-0485(2004)034<1117:ITRATM>2.0.CO;2. ISSN 1520-0485.
  13. ^ Garrett, C.; E. Kunze (2007). "Internal tide generation in the deep ocean". Annu. Rev. Fluid Mech. 39 (1): 57–87. Bibcode:2007AnRFM..39...57G. doi:10.1146/annurev.fluid.39.050905.110227.
  14. ^ a b c Wunsch, C.; R. Ferrari (2004). "Vertical mixing, energy, and the general circulation of the ocean". Annu. Rev. Fluid Mech. 36 (1): 281–314. Bibcode:2004AnRFM..36..281W. CiteSeerX doi:10.1146/annurev.fluid.36.050802.122121.
  15. ^ Munk, W.; Wunsch, C. (1998). "Abyssal recipes II: Energetics of tidal and wind mixing". Deep-Sea Research. 45 (12): 1977–2010. Bibcode:1998DSRI...45.1977M. doi:10.1016/S0967-0637(98)00070-3.
  16. ^ Munk, W. (1966). "Abyssal recipes". Deep-Sea Research. 13 (4): 707–730. Bibcode:1966DSROA..13..707M. doi:10.1016/0011-7471(66)90602-4.
  17. ^ Sandstrom, J. W. (1908). "Dynamische Versuche mit Meerwasser". Ann. Hydrodyn. Marine Meteorology: 6.
  18. ^ Alfonso-Sosa, E. (2002). Variabilidad temporal de la producción primaria fitoplanctonica en la estación CaTS (Caribbean Time-Series Station): Con énfasis en el impacto de la marea interna semidiurna sobre la producción (PDF). Ph. D. Dissertation. Department of Marine Sciences, University of Puerto Rico, Mayagüez, Puerto Rico. UMI publication AAT 3042382. p. 407. Retrieved 2014-08-25.
  19. ^ Alfonso-Sosa, E.; J. Morell; J. M. Lopez; J. E. Capella & A. Dieppa (2002). "Internal Tide-induced Variations in Primary Productivity and Optical Properties in the Mona Passage, Puerto Rico" (PDF). Retrieved 2015-01-01.
  20. ^ Sharples, J.; V. Krivtsov; J. F. Tweddle; J. A. M. Green; M. R. Palmer; Y. Kim; A. E. Hickman; P. M. Holligan; C. M. Moore; T. P. Rippeth & J. H. Simpson (2007). "Spring–neap modulation of internal tide mixing and vertical nitrate fluxes at a shelf edge in summer" (PDF). Limnol. Oceanogr. 52 (5): 1735–1747. Bibcode:2007LimOc..52.1735S. doi:10.4319/lo.2007.52.5.1735. Retrieved 2014-08-25.

External links

  • [1] Scripps Institution of Oceanography
  • [2] Southern California Coastal Ocean Observing System
  • [3] Internal Tides of the Oceans, Harper Simmons, by Jenn Wagaman of Arctic Region Supercomputing Center
  • [4] Principal tidal constituents in Physical oceanography textbook, Bob Stewart of Texas A&M University
  • [5] Eric Kunze's work on internal waves, internal tides, mixing, and more
Bahama Banks

The Bahama Banks are the submerged carbonate platforms that make up much of the Bahama Archipelago. The term is usually applied in referring to either the Great Bahama Bank around Andros Island, or the Little Bahama Bank of Grand Bahama Island and Great Abaco, which are the largest of the platforms, and the Cay Sal Bank north of Cuba. The islands of these banks are politically part of the Bahamas. Other banks are the three banks of the Turks and Caicos Islands, namely the Caicos Bank of the Caicos Islands, the bank of the Turks Islands, and wholly submerged Mouchoir Bank. Further southeast are the equally wholly submerged Silver Bank and Navidad Bank north of the Dominican Republic.

Carbonate platform

A carbonate platform is a sedimentary body which possesses topographic relief, and is composed of autochthonic calcareous deposits. Platform growth is mediated by sessile organisms whose skeletons build up the reef or by organisms (usually microbes) which induce carbonate precipitation through their metabolism. Therefore, carbonate platforms can not grow up everywhere: they are not present in places where limiting factors to the life of reef-building organisms exist. Such limiting factors are, among others: light, water temperature, transparency and pH-Value. For example, carbonate sedimentation along the Atlantic South American coasts takes place everywhere but at the mouth of the Amazon River, because of the intense turbidity of the water there. Spectacular examples of present-day carbonate platforms are the Bahama Banks under which the platform is roughly 8 km thick, the Yucatan Peninsula which is up to 2 km thick, the Florida platform, the platform on which the Great Barrier Reef is growing, and the Maldive atolls. All these carbonate platforms and their associated reefs are confined to tropical latitudes. Today's reefs are built mainly by scleractinian corals, but in the distant past other organisms, like archaeocyatha (during the Cambrian) or extinct cnidaria (tabulata and rugosa) were important reef builders.

Coral reef

A coral reef is an underwater ecosystem characterized by reef-building corals. Reefs are formed of colonies of coral polyps held together by calcium carbonate. Most coral reefs are built from stony corals, whose polyps cluster in groups.

Coral belongs to the class Anthozoa in the animal phylum Cnidaria, which includes sea anemones and jellyfish. Unlike sea anemones, corals secrete hard carbonate exoskeletons that support and protect the coral. Most reefs grow best in warm, shallow, clear, sunny and agitated water.

Often called "rainforests of the sea", shallow coral reefs form some of Earth's most diverse ecosystems. They occupy less than 0.1% of the world's ocean area, about half the area of France, yet they provide a home for at least 25% of all marine species, including fish, mollusks, worms, crustaceans, echinoderms, sponges, tunicates and other cnidarians. Coral reefs flourish in ocean waters that provide few nutrients. They are most commonly found at shallow depths in tropical waters, but deep water and cold water coral reefs exist on smaller scales in other areas.

Coral reefs deliver ecosystem services for tourism, fisheries and shoreline protection. The annual global economic value of coral reefs is estimated between US$30–375 billion and 9.9 trillion USD. Coral reefs are fragile, partly because they are sensitive to water conditions. They are under threat from excess nutrients (nitrogen and phosphorus), rising temperatures, oceanic acidification, overfishing (e.g., from blast fishing, cyanide fishing, spearfishing on scuba), sunscreen use, and harmful land-use practices, including runoff and seeps (e.g., from injection wells and cesspools).

List of submarine volcanoes

A list of active and extinct submarine volcanoes and seamounts located under the world's oceans. There are estimated to be 40,000 to 55,000 seamounts in the global oceans. Almost all are not well-mapped and many may not have been identified at all. Most are unnamed and unexplored. This list is therefore confined to seamounts that are notable enough to have been named and/or explored.

Mona Passage

The Mona Passage (Spanish: Canal de la Mona) is a strait that separates the islands of Hispaniola and Puerto Rico. The Mona Passage connects the Atlantic Ocean to the Caribbean Sea, and is an important shipping route between the Atlantic and the Panama Canal.

The 80 mi (130 km) stretch of sea between the four islands is one of the most difficult passages in the Caribbean. It is fraught with variable tidal currents created by the large islands on either side of it, and by sand banks that extend out for many miles from both coasts.

Oceanic plateau

An oceanic or submarine plateau is a large, relatively flat elevation that is higher than the surrounding relief with one or more relatively steep sides.There are 184 oceanic plateaus covering an area of 18,486,600 km2 (7,137,700 sq mi), or about 5.11% of the oceans. The South Pacific region around Australia and New Zealand contains the greatest number of oceanic plateaus (see map).

Oceanic plateaus produced by large igneous provinces are often associated with hotspots, mantle plumes, and volcanic islands — such as Iceland, Hawaii, Cape Verde, and Kerguelen. The three largest plateaus, the Caribbean, Ontong Java, and Mid-Pacific Mountains, are located on thermal swells. Other oceanic plateaus, however, are made of rifted continental crust, for example Falkland Plateau, Lord Howe Rise, and parts of Kerguelen, Seychelles, and Arctic ridges.

Plateaus formed by large igneous provinces were formed by the equivalent of continental flood basalts such as the Deccan Traps in India and the Snake River Plain in the United States.

In contrast to continental flood basalts, most igneous oceanic plateaus erupt through young and thin (6–7 km (3.7–4.3 mi)) mafic or ultra-mafic crust and are therefore uncontaminated by felsic crust and representative for their mantle sources.

These plateaus often rise 2–3 km (1.2–1.9 mi) above the surrounding ocean floor and are more buoyant than oceanic crust. They therefore tend to withstand subduction, more-so when thick and when reaching subduction zones shortly after their formations. As a consequence, they tend to "dock" to continental margins and be preserved as accreted terranes. Such terranes are often better preserved than the exposed parts of continental flood basalts and are therefore a better record of large-scale volcanic eruptions throughout Earth's history. This "docking" also means that oceanic plateaus are important contributors to the growth of continental crust. Their formations often had a dramatic impact on global climate, such as the most recent plateaus formed, the three, large, Cretaceous oceanic plateaus in the Pacific and Indian Ocean: Ontong Java, Kerguelen, and Caribbean.

Outline of oceanography

The following outline is provided as an overview of and introduction to Oceanography.

Physical oceanography

Physical oceanography is the study of physical conditions and physical processes within the ocean, especially the motions and physical properties of ocean waters.

Physical oceanography is one of several sub-domains into which oceanography is divided. Others include biological, chemical and geological oceanography.

Physical oceanography may be subdivided into descriptive and dynamical physical oceanography.Descriptive physical oceanography seeks to research the ocean through observations and complex numerical models, which describe the fluid motions as precisely as possible.

Dynamical physical oceanography focuses primarily upon the processes that govern the motion of fluids with emphasis upon theoretical research and numerical models. These are part of the large field of Geophysical Fluid Dynamics (GFD) that is shared together with meteorology. GFD is a sub field of Fluid dynamics describing flows occurring on spatial and temporal scales that are greatly influenced by the Coriolis force.

Poor Knights Islands

The Poor Knights Islands are a group of islands off the east coast of the Northland Region of the North Island of New Zealand. They lie 50 kilometres (31 mi) to the north-east of Whangarei, and 22 kilometres (14 mi) offshore halfway between Bream Head and Cape Brett. Uninhabited since the 1820s, they are a nature reserve and popular underwater diving spot, with boat tours typically departing from Tutukaka. The Poor Knights Islands Marine Reserve surrounds the island. Beaglehole (1955) comments that the origin of the island name is not clear, and speculates that the name could be related to the Poor Knights of Windsor, or, that the islands were named for their resemblance to Poor Knight's Pudding, a bread-based dish topped with egg and fried, popular at the time of discovery by Europeans.

Undersea mountain range

Undersea mountain ranges are mountain ranges that are mostly or entirely underwater, and specifically under the surface of an ocean. If originated from current tectonic forces, they are often referred to as a mid-ocean ridge. In contrast, if formed by past above-water volcanism, they are known as a seamount chain. The largest and best known undersea mountain range is a mid-ocean ridge, the Mid-Atlantic Ridge. It has been observed that, "similar to those on land, the undersea mountain ranges are the loci of frequent volcanic and earthquake activity".

Wave base

The wave base, in physical oceanography, is the maximum depth at which a water wave's passage causes significant water motion. For water depths deeper than the wave base, bottom sediments and the seafloor are no longer stirred by the wave motion above.

Ocean zones
Sea level

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