Oceanic crust is formed at anoceanic ridge, while the lithosphere is subducted back into the asthenosphere at trenches
Oceanic trenches are prominent, long, narrowtopographicdepressions of theocean floor. They are typically 50 to 100 kilometers (30 to 60 mi) wide and 3 to 4 km (1.9 to 2.5 mi) below the level of the surrounding oceanic floor, but can be thousands of kilometers in length. There are about 50,000 km (31,000 mi) of oceanic trenches worldwide, mostly around thePacific Ocean, but also in the easternIndian Ocean and a few other locations. The greatest ocean depth measured is in theChallenger Deep of theMariana Trench, at a depth of 10,994 m (36,070 ft) belowsea level.
Oceanic trenches are a feature of the Earth's distinctiveplate tectonics. They mark the locations ofconvergent plate boundaries, along whichlithospheric plates move towards each other at rates that vary from a few millimeters to over ten centimeters per year. Oceanic lithosphere moves into trenches at a global rate of about 3 km2 (1.2 sq mi) per year.[1] A trench marks the position at which the flexed,subductingslab begins to descend beneath another lithospheric slab. Trenches are generally parallel to and about 200 km (120 mi) from avolcanic arc.
Much of the fluid trapped insediments of the subducting slab returns to the surface at the oceanic trench, producingmud volcanoes andcold seeps. These support uniquebiomes based onchemotrophic microorganisms. There is concern thatplastic debris is accumulating in trenches and threatening these communities.
There are approximately 50,000 km (31,000 mi) ofconvergent plate margins worldwide. These are mostly located around the Pacific Ocean, but are also found in the easternIndian Ocean, with a few shorter convergent margin segments in other parts of the Indian Ocean, in the Atlantic Ocean, and in the Mediterranean.[2] They are found on the oceanward side ofisland arcs and Andean-typeorogens.[3] Globally, there are over 50 major ocean trenches covering an area of 1.9 million km2 or about 0.5% of the oceans.[4]
Trenches are geomorphologically distinct fromtroughs. Troughs are elongated depressions of the sea floor with steep sides and flat bottoms, while trenches are characterized by a V-shaped profile.[4] Trenches that are partially infilled are sometimes described as troughs, for example theMakran Trough.[5] Some trenches are completely buried and lack bathymetric expression as in theCascadia subduction zone,[6] which is completely filled with sediments.[7] Despite their appearance, in these instances the fundamentalplate-tectonic structure is still an oceanic trench. Some troughs look similar to oceanic trenches but possess other tectonic structures. One example is theLesser Antilles Trough, which is the forearc basin of theLesser Antilles subduction zone.[8] Also not a trench is theNew Caledonia trough, which is an extensionalsedimentary basin related to theTonga-Kermadec subduction zone.[9] Additionally, the Cayman Trough, which is apull-apart basin within atransform fault zone,[10] is not an oceanic trench.
Trenches are related to, but distinct from, continental collision zones, such as theHimalayas. Unlike in trenches, in continental collision zonescontinental crust enters a subduction zone. When buoyant continental crust enters a trench, subduction comes to a halt and the area becomes a zone of continental collision. Features analogous to trenches are associated withcollision zones. One such feature is theperipheral foreland basin, a sediment-filledforedeep. Examples of peripheral foreland basins include thefloodplains of theGanges River and theTigris-Euphrates river system.[2]
Trenches were not clearly defined until the late1940s and1950s. Thebathymetry of the ocean was poorly known prior to theChallenger expedition of 1872–1876,[12] which took 492 soundings of the deep ocean.[13] At station #225, the expedition discoveredChallenger Deep,[14] now known to be the southern end of theMariana Trench. The laying oftransatlantic telegraph cables on the seafloor between the continents during the late 19th and early 20th centuries provided further motivation for improved bathymetry.[15] The termtrench, in its modern sense of a prominent elongated depression of the sea bottom, was first used by Johnstone in his 1923 textbookAn Introduction to Oceanography.[16][2]
During the 1920s and 1930s,Felix Andries Vening Meinesz measured gravity over trenches using a newly developedgravimeter that could measuregravity from aboard a submarine.[11] He proposed thetectogene hypothesis to explain the belts of negative gravity anomalies that were found near island arcs. According to this hypothesis, the belts were zones ofdownwelling of light crustal rock arising from subcrustal convection currents. The tectogene hypothesis was further developed by Griggs in 1939, using an analogue model based on a pair of rotating drums.Harry Hammond Hess substantially revised the theory based on his geological analysis.[17]
World War II in the Pacific led to great improvements of bathymetry, particularly in the western Pacific. In light of these new measurements, the linear nature of the deeps became clear. There was a rapid growth of deep sea research efforts, especially the widespread use ofechosounders in the 1950s and 1960s. These efforts confirmed the morphological utility of the term "trench." Important trenches were identified, sampled, and mapped via sonar.
The early phase of trench exploration reached its peak with the 1960 descent of theBathyscapheTrieste to the bottom of the Challenger Deep. FollowingRobert S. Dietz' andHarry Hess' promulgation of the seafloor spreading hypothesis in the early 1960s and the plate tectonic revolution in the late 1960s, the oceanic trench became an important concept inplate tectonic theory.[11]
Cross section of an oceanic trench formed along an oceanic-oceanicconvergent boundary ThePeru–Chile Trench is located just left of the sharp line between the blue deep ocean (on the left) and the light blue continental shelf, along the west coast of South America. It runs along an oceanic-continental boundary, where the oceanicNazca plate subducts beneath the continentalSouth American plate
Oceanic trenches are 50 to 100 kilometers (30 to 60 mi) wide and have an asymmetric V-shape, with the steeperslope (8 to 20 degrees) on the inner (overriding) side of the trench and the gentler slope (around 5 degrees) on the outer (subducting) side of the trench.[18][19] The bottom of the trench marks the boundary between the subducting and overriding plates, known as the basal plate boundary shear[20] or the subductiondécollement.[2] The depth of the trench depends on the starting depth of the oceanic lithosphere as it begins its plunge into the trench, theangle at which the slab plunges, and the amount of sedimentation in the trench. Both starting depth and subduction angle are greater for older oceanic lithosphere, which is reflected in the deep trenches of the western Pacific. Here the bottoms of the Marianas and the Tonga–Kermadec trenches are up to 10–11 kilometers (6.2–6.8 mi) below sea level. In the eastern Pacific, where the subducting oceanic lithosphere is much younger, the depth of the Peru-Chile trench is around 7 to 8 kilometers (4.3 to 5.0 mi).[18]
Though narrow, oceanic trenches are remarkably long and continuous, forming the largestlinear depressions on earth. An individual trench can be thousands of kilometers long.[3] Most trenches are convex towards the subducting slab, which is attributed to the spherical geometry of the Earth.[21]
The trench asymmetry reflects the different physical mechanisms that determine the inner and outer slope angle. The outer slope angle of the trench is determined by the bending radius of the subducting slab, as determined by its elastic thickness. Since oceanic lithosphere thickens with age, the outer slope angle is ultimately determined by the age of the subducting slab.[22][20] The inner slope angle is determined by theangle of repose of the overriding plate edge.[20] This reflects frequent earthquakes along the trench that prevent oversteepening of the inner slope.[2]
As the subducting plate approaches the trench, it bends slightly upwards before beginning its plunge into the depths. As a result, the outer trench slope is bounded by anouter trench high. This is subtle, often only tens of meters high, and is typically located a few tens of kilometers from the trench axis. On the outer slope itself, where the plate begins to bend downwards into the trench, the upper part of the subducting slab is broken by bending faults that give the outer trench slope ahorst and graben topography. The formation of these bending faults is suppressed where oceanic ridges or large seamounts are subducting into the trench, but the bending faults cut right across smaller seamounts. Where the subducting slab is only thinly veneered with sediments, the outer slope will often showseafloor spreading ridges oblique to the horst and graben ridges.[20]
Trench morphology is strongly modified by the amount of sedimentation in the trench. This varies from practically no sedimentation, as in the Tonga-Kermadec trench, to completely filled with sediments, as with theCascadia subduction zone. Sedimentation is largely controlled by whether the trench is near a continental sediment source.[21] The range of sedimentation is well illustrated by the Chilean trench. The north Chile portion of the trench, which lies along theAtacama Desert with its very slow rate of weathering, is sediment-starved, with from 20 to a few hundred meters of sediments on the trench floor. The tectonic morphology of this trench segment is fully exposed on the ocean bottom. The central Chile segment of the trench is moderately sedimented, with sediments onlapping ontopelagic sediments or ocean basement of the subducting slab, but the trench morphology is still clearly discernible. The southern Chile segment of the trench is fully sedimented, to the point where the outer rise and slope are no longer discernible. Other fully sedimented trenches include the Makran Trough, where sediments are up to 7.5 kilometers (4.7 mi) thick; the Cascadia subduction zone, which is completed buried by 3 to 4 kilometers (1.9 to 2.5 mi) of sediments; and the northernmost Sumatra subduction zone, which is buried under 6 kilometers (3.7 mi) of sediments.[23]
Sediments are sometimes transported along the axis of an oceanic trench. The central Chile trench experiences transport of sediments from source fans along an axial channel.[24] Similar transport of sediments has been documented in the Aleutian trench.[2]
In addition to sedimentation from rivers draining into a trench, sedimentation also takes place from landslides on the tectonically steepened inner slope, often driven bymegathrust earthquakes. The Reloca Slide of the central Chile trench is an example of this process.[25]
Convergent margins are classified as erosive or accretionary, and this has a strong influence on the morphology of the inner slope of the trench. Erosive margins, such as the northern Peru-Chile, Tonga-Kermadec, and Mariana trenches, correspond to sediment-starved trenches.[3] The subducting slab erodes material from the lower part of the overriding slab, reducing its volume. The edge of the slab experiences subsidence and steepening, with normal faulting. The slope is underlain by relative strong igneous and metamorphic rock, which maintains a high angle of repose.[26] Over half of all convergent margins are erosive margins.[2]
Accretionary margins, such as the southern Peru-Chile, Cascadia, and Aleutians, are associated with moderately to heavily sedimented trenches. As the slab subducts, sediments are "bulldozed" onto the edge of the overriding plate, producing anaccretionary wedge oraccretionary prism. This builds the overriding plate outwards. Because the sediments lack strength, their angle of repose is gentler than the rock making up the inner slope of erosive margin trenches. The inner slope is underlain byimbricatedthrust sheets of sediments. The inner slope topography is roughened by localizedmass wasting.[26] Cascadia has practically no bathymetric expression of the outer rise and trench, due to complete sediment filling, but the inner trench slope is complex, with many thrust ridges. These compete with canyon formation by rivers draining into the trench. Inner trench slopes of erosive margins rarely show thrust ridges.[19]
Accretionary prisms grow in two ways. The first is by frontal accretion, in which sediments are scraped off the downgoing plate and emplaced at the front of the accretionary prism.[2] As the accretionary wedge grows, older sediments further from the trench become increasinglylithified, and faults and other structural features are steepened by rotation towards the trench.[27] The other mechanism for accretionary prism growth is underplating[2] (also known as basal accretion[28]) of subducted sediments, together with someoceanic crust, along the shallow parts of the subduction decollement. TheFranciscan Group ofCalifornia is interpreted as an ancient accretionary prism in which underplating is recorded as tectonic mélanges and duplex structures.[2]
Oceanic trench formed along an oceanic-oceanicconvergent boundary TheMariana Trench contains the deepest part of the world's oceans, and runs along an oceanic-oceanic convergent boundary. It is the result of the oceanicPacific platesubducting beneath the oceanicMariana plate.
Frequentmegathrust earthquakes modify the inner slope of the trench by triggering massive landslides. These leave semicircular landslide scarps with slopes of up to 20 degrees on the headwalls and sidewalls.[29]
Subduction of seamounts andaseismic ridges into the trench may increaseaseismic creep and reduce the severity of earthquakes. Contrariwise, subduction of large amounts of sediments may allow ruptures along the subduction décollement to propagate for great distances to produce megathrust earthquakes.[30]
Trenches seem positionally stable over time, but scientists believe that some trenches—particularly those associated with subduction zones where two oceanic plates converge—move backward into the subducting plate.[31][32] This is calledtrench rollback orretreat,hinge rollback orretreat,slab rollback orretreat and is one explanation for the existence ofback-arc basins.
Forces perpendicular to the slab (the portion of the subducting plate within the mantle) are responsible for steepening of the slab and, ultimately, the movement of the hinge and trench at the surface.[33] These forces arise from the negative buoyancy of the slab with respect to the mantle[34] modified by the geometry of the slab itself.[35] The extension in the overriding plate, in response to the subsequent subhorizontal mantle flow from the displacement of the slab, can result in formation of a back-arc basin.[36]
Several forces are involved in the process of slab rollback. Two forces acting against each other at the interface of the two subducting plates exert forces against one another. The subducting plate exerts a bending force (FPB) that supplies pressure during subduction, while the overriding plate exerts a force against the subducting plate (FTS). The slab pull force (FSP) is caused by the negative buoyancy of the plate driving the plate to greater depths. The resisting force from the surrounding mantle opposes the slab pull forces. Interactions with the 660-km discontinuity cause a deflection due to the buoyancy at thephase transition (F660).[35] The unique interplay of these forces is what generates slab rollback. When the deep slab section obstructs the down-going motion of the shallow slab section, slab rollback occurs. The subducting slab undergoes backward sinking due to the negative buoyancy forces causing a retrogradation of the trench hinge along the surface. Upwelling of the mantle around the slab can create favorable conditions for the formation of a back-arc basin.[36]
Seismic tomography provides evidence for slab rollback. Results demonstrate high temperature anomalies within the mantle suggesting subducted material is present in the mantle.[37] Ophiolites are viewed as evidence for such mechanisms as high pressure and temperature rocks are rapidly brought to the surface through the processes of slab rollback, which provides space for the exhumation ofophiolites.
Slab rollback is not always a continuous process suggesting an episodic nature.[34] The episodic nature of the rollback is explained by a change in the density of the subducting plate, such as the arrival of buoyant lithosphere (a continent, arc, ridge, or plateau), a change in the subduction dynamics, or a change in the plate kinematics. The age of the subducting plates does not have any effect on slab rollback.[35] Nearby continental collisions have an effect on slab rollback. Continental collisions induce mantle flow and extrusion of mantle material, which causes stretching and arc-trench rollback.[36] In the area of the Southeast Pacific, there have been several rollback events resulting in the formation of numerous back-arc basins.[34]
Interactions with themantle discontinuities play a significant role in slab rollback. Stagnation at the 660-km discontinuity causes retrograde slab motion due to the suction forces acting at the surface.[35] Slab rollback induces mantle return flow, which causes extension from theshear stresses at the base of the overriding plate. As slab rollback velocities increase, circular mantle flow velocities also increase, accelerating extension rates.[33] Extension rates are altered when the slab interacts with the discontinuities within the mantle at 410 km and 660 km depth. Slabs can either penetrate directly into thelower mantle, or can be retarded due to the phase transition at 660 km depth creating a difference in buoyancy. An increase in retrograde trench migration (slab rollback) (2–4 cm/yr) is a result of flattened slabs at the 660-km discontinuity where the slab does not penetrate into the lower mantle.[38] This is the case for the Japan, Java and Izu–Bonin trenches. These flattened slabs are only temporarily arrested in the transition zone. The subsequent displacement into the lower mantle is caused by slab pull forces, or the destabilization of the slab from warming and broadening due to thermal diffusion. Slabs that penetrate directly into the lower mantle result in slower slab rollback rates (~1–3 cm/yr) such as the Mariana arc, Tonga arcs.[38]
As sediments are subducted at the bottom of trenches, much of their fluid content is expelled and moves back along the subduction décollement to emerge on the inner slope asmud volcanoes andcold seeps.Methane clathrates andgas hydrates also accumulate in the inner slope, and there is concern that their breakdown could contribute toglobal warming.[2]
The fluids released at mud volcanoes and cold seeps are rich inmethane andhydrogen sulfide, providing chemical energy forchemotrophic microorganisms that form the base of a unique trenchbiome. Cold seep communities have been identified in the inner trench slopes of the western Pacific (especially Japan[39]), South America, Barbados, the Mediterranean, Makran, and the Sunda trench. These are found at depths as great as 6,000 meters (20,000 ft).[2] The genome of theextremophileDeinococcus from Challenger Deep has sequenced for its ecological insights and potential industrial uses.[40]
Because trenches are the lowest points in the ocean floor, there is concern thatplastic debris may accumulate in trenches and endanger the fragile trench biomes.[41]
Recent measurements, where the salinity and temperature of the water was measured throughout the dive, have uncertainties of about 15 m (49 ft).[42] Older measurements may be off by hundreds of meters.
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