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Thursday, September 11, 2014

Hotspot (geology)

Hotspot (geology)

From Wikipedia, the free encyclopedia
Diagram showing a cross section though the Earth's lithosphere (in yellow) with magma rising from the mantle (in red)

The places known as hotspots or hot spots in geology are volcanic regions thought to be fed by underlying mantle that is anomalously hot compared with the mantle elsewhere. They may be unanimously hot, and provide a great deal of molten magma. They may be on, near to, or far from tectonic plate boundaries. There are two hypotheses to explain them. One suggests that they are due to hot mantle plumes that rise as thermal diapirs from the core-mantle boundary.[1] The other hypothesis postulates that it is not high temperature that causes the volcanism, but lithospheric extension that permits the passive rising of melt from shallow depths.[2][3] This hypothesis considers the term "hotspot" to be a misnomer, asserting that the mantle source beneath them is, in fact, not anomalously hot at all. Well known examples include Hawaii and Yellowstone.

Background

Schematic diagram showing the physical processes inside the Earth that lead to the generation of magma. Partial melting begins above the fusion point.

The origins of the concept of hotspots lie in the work of J. Tuzo Wilson, who postulated in 1963 that the Hawaiian Islands result from the slow movement of a tectonic plate across a hot region beneath the surface.[4] It was later postulated that hotspots are fed by narrow streams of hot mantle rising from the Earth's core-mantle boundary in a structure called a mantle plume.[5] Whether or not such mantle plumes exist is currently the subject of a major controversy in Earth science.[3][6] Estimates for the number of hotspots postulated to be fed by mantle plumes has ranged from about 20 to several thousands, over the years, with most geologists considering a few tens to exist. Hawaii, Réunion, Yellowstone, Galápagos, and Iceland are some of the currently most active volcanic regions to which the hypothesis is applied.

Most hotspot volcanoes are basaltic (e.g., Hawaii, Tahiti). As a result, they are less explosive than subduction zone volcanoes, in which water is trapped under the overriding plate. Where hotspots occur in continental regions, basaltic magma rises through the continental crust, which melts to form rhyolites. These rhyolites can form violent eruptions. For example, the Yellowstone Caldera was formed by some of the most powerful volcanic explosions in geologic history. However, when the rhyolite is completely erupted, it may be followed by eruptions of basaltic magma rising through the same lithospheric fissures (cracks in the lithosphere). An example of this activity is the Ilgachuz Range in British Columbia, which was created by an early complex series of trachyte and rhyolite eruptions, and late extrusion of a sequence of basaltic lava flows.[7]

The hotspot hypothesis is now closely linked to the mantle plume hypothesis.

Comparison with island arc volcanoes

Hotspot volcanoes are considered to have a fundamentally different origin from island arc volcanoes.
The latter form over subduction zones, at converging plate boundaries. When one oceanic plate meets another, the denser plate is forced downward into a deep ocean trench. This plate, as it is subducted, releases water into the base of the over-riding plate, and this water mixes with the rock, thus changing its composition causing some rock to melt and rise. It is this that fuels a chain of volcanoes, such as the Aleutian Islands, near Alaska.

Hotspot volcanic chains

Over millions of years, the Pacific Plate has moved over the Hawaii hotspot, creating a trail of underwater mountains that stretch across the Pacific
Kilauea is the most active shield volcano in the world. The volcano has erupted nonstop since 1983 and it is part of the Hawaiian-Emperor seamount chain.
Mauna Loa is a large shield volcano. It's last eruption was in 1984 and it's part of the Hawaiian-Emperor seamount chain.
Bowie Seamount is a dormant submarine volcano and it's part of the Kodiak-Bowie Seamount chain.
Axial Seamount is the youngest seamount of the Cobb-Eickelberg Seamount chain It's last eruption was possibly in 2011.
Mauna Kea is the tallest volcano in the Hawaiian-Emperor seamount chain. It's currently dormant and it has cinder cones growing on the volcano.
Hualalai is a massive shield volcano in the Hawaiian-Emperor seamount chain. It's last eruption was in 1801.

The joint mantle plume/hotspot hypothesis envisages the feeder structures to be fixed relative to one another, with the continents and seafloor drifting overhead. The hypothesis thus predicts that time-progressive chains of volcanoes are developed on the surface. Examples are Yellowstone, which lies at the end of a chain of extinct calderas, which become progressively older to the west. Another example is the Hawaiian archipelago, where islands become progressively older and more deeply eroded to the northwest.

Geologists have tried to use hotspot volcanic chains to track the movement of the Earth's tectonic plates. This effort has been vexed by the lack of very long chains, by the fact that many are not time-progressive (e.g. the Galápagos) and by the fact that hotspots do not appear to be fixed relative to one another (e.g., Hawaii and Iceland.[8])

 Postulated hotspot volcano chains

An example of mantle plume locations suggested by one recent group.[9] Figure from Foulger (2010).[3]

List of volcanic regions postulated to be hotspots

Distribution of selected hotspots. The numbers in the figure are related to the listed hotspots on the left.

Eurasian Plate

African Plate

Antarctic Plate

South American Plate

North American Plate

Indo-Australian Plate

Nazca Plate

Pacific Plate

Over millions of years, the Pacific Plate has moved over the Bowie hotspot, creating the Kodiak-Bowie Seamount chain in the Gulf of Alaska

Former hotspots

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Mantle (geology)

Mantle (geology)

From Wikipedia, the free encyclopedia
 
The internal structure of Earth
 
The mantle is a part of a terrestrial planet or other rocky body large enough to have differentiation by density. The interior of Earth, similar to the other terrestrial planets, is chemically divided into layers. The mantle is a layer between the crust and the outer core. Earth's mantle is a silicate rocky shell about 2,900 kilometres (1,800 mi) thick[1] that constitutes about 84% of Earth's volume.[2] It is predominantly solid but in geological time it behaves like very viscous liquid. The mantle encloses the hot core rich in iron and nickel, which occupies about 15% of Earth's volume.[2][3] Past episodes of melting and volcanism at the shallower levels of the mantle have produced a thin crust of crystallized melt products near the surface, upon which we live.[4] Information about structure and composition of the mantle either result from geophysical investigation or from direct geoscientific analyses on Earth mantle derived xenoliths and on mantle exposed by mid-oceanic ridge spreading.

Two main zones are distinguished in the upper mantle: the inner asthenosphere composed of plastic flowing rock about 200 km (120 mi) thick,[5] and the lowermost part of the lithosphere composed of rigid rock about 50 to 120 km (31 to 75 mi) thick.[6] A thin crust, the upper part of the lithosphere, surrounds the mantle and is about 5 to 75 km (3.1 to 46.6 mi) thick.[7] Recent analysis of hydrous ringwoodite from the mantle suggests that there is between one[8] and three[9] times as much water in the transition zone between the lower and upper mantle than in all the world's oceans combined.

In some places under the ocean the mantle is actually exposed on the surface of Earth.[10] There are also a few places on land where mantle rock has been pushed to the surface by tectonic activity, most notably the Tablelands region of Gros Morne National Park in the Canadian province of Newfoundland and Labrador.

Structure

The mantle is divided into sections which are based upon results from seismology. These layers (and their thicknesses/depths) are the following: the upper mantle (starting at the Moho, or base of the crust around 7 to 35 km (4.3 to 21.7 mi) downward to 410 km (250 mi)),[11] the transition zone (410–660 km or 250–410 mi), the lower mantle (660–2,891 km or 410–1,796 mi), and anomalous core–mantle boundary with a variable thickness (on average ~200 km (120 mi) thick).[4][12][13][14]

The top of the mantle is defined by a sudden increase in seismic velocity, which was first noted by Andrija Mohorovičić in 1909; this boundary is now referred to as the Mohorovičić discontinuity or "Moho".[12][15] The uppermost mantle plus overlying crust are relatively rigid and form the lithosphere, an irregular layer with a maximum thickness of perhaps 200 km (120 mi). Below the lithosphere the upper mantle becomes notably more plastic. In some regions below the lithosphere, the seismic velocity is reduced; this so-called low-velocity zone (LVZ) extends down to a depth of several hundred km. Inge Lehmann discovered a seismic discontinuity at about 220 km (140 mi) depth;[16] although this discontinuity has been found in other studies, it is not known whether the discontinuity is ubiquitous. The transition zone is an area of great complexity; it physically separates the upper and lower mantle.[14] Very little is known about the lower mantle apart from that it appears to be relatively seismically homogeneous. The D" layer at the core–mantle boundary separates the mantle from the core.[4][12]

Characteristics

The mantle differs substantially from the crust in its mechanical properties which is the direct consequence of chemical composition change (expressed as different mineralogy). The distinction between crust and mantle is based on chemistry, rock types, rheology and seismic characteristics. The crust is a solidification product of mantle derived melts, expressed as various degrees of partial melting products during geologic time. Partial melting of mantle material is believed to cause incompatible elements to separate from the mantle, with less dense material floating upward through pore spaces, cracks, or fissures, that would subsequently cool and solidify at the surface. Typical mantle rocks have a higher magnesium to iron ratio and a smaller proportion of silicon and aluminium than the crust. This behavior is also predicted by experiments that partly melt rocks thought to be representative of Earth's mantle.
Mapping the interior of the Earth with earthquake waves.

Mantle rocks shallower than about 410 km (250 mi) depth consist mostly of olivine, pyroxenes, spinel-structure minerals, and garnet;[14] typical rock types are thought to be peridotite,[14] dunite (olivine-rich peridotite), and eclogite. Between about 400 km (250 mi) and 650 km (400 mi) depth, olivine is not stable and is replaced by high pressure polymorphs with approximately the same composition: one polymorph is wadsleyite (also called beta-spinel type), and the other is ringwoodite (a mineral with the gamma-spinel structure). Below about 650 km (400 mi), all of the minerals of the upper mantle begin to become unstable. The most abundant minerals present, the silicate perovskites, have structures (but not compositions) like that of the mineral perovskite followed by the magnesium/iron oxide ferropericlase.[17] The changes in mineralogy at about 400 and 650 km (250 and 400 mi) yield distinctive signatures in seismic records of the Earth's interior, and like the moho, are readily detected using seismic waves. These changes in mineralogy may influence mantle convection, as they result in density changes and they may absorb or release latent heat as well as depress or elevate the depth of the polymorphic phase transitions for regions of different temperatures. The changes in mineralogy with depth have been investigated by laboratory experiments that duplicate high mantle pressures, such as those using the diamond anvil.[18]

Composition of Earth's mantle in weight percent[19] [20]
Element Amount   Compound Amount
O 44.8    
Mg 22.8 SiO2 46
Si 21.5 MgO 37.8
Fe 5.8 FeO 7.5
Ca 2.3 Al2O3 4.2
Al 2.2 CaO 3.2
Na 0.3 Na2O 0.4
K 0.03 K2O 0.04
Sum 99.7 Sum 99.1

The inner core is solid, the outer core is liquid, and the mantle solid/plastic. This is because of the relative melting points of the different layers (nickel–iron core, silicate crust and mantle) and the increase in temperature and pressure as depth increases. At the surface both nickel–iron alloys and silicates are sufficiently cool to be solid. In the upper mantle, the silicates are generally solid (localised regions with small amounts of melt exist); however, as the upper mantle is both hot and under relatively little pressure, the rock in the upper mantle has a relatively low viscosity. In contrast, the lower mantle is under tremendous pressure and therefore has a higher viscosity than the upper mantle. The metallic nickel–iron outer core is liquid because of the high temperature, despite the high pressure. As the pressure increases, the nickel–iron inner core becomes solid because the melting point of iron increases dramatically at these high pressures.[21]

Temperature

In the mantle, temperatures range between 500 to 900 °C (932 to 1,652 °F) at the upper boundary with the crust; to over 4,000 °C (7,230 °F) at the boundary with the core.[21] Although the higher temperatures far exceed the melting points of the mantle rocks at the surface (about 1200 °C for representative peridotite), the mantle is almost exclusively solid.[21] The enormous lithostatic pressure exerted on the mantle prevents melting, because the temperature at which melting begins (the solidus) increases with pressure.

Movement

This figure is a snapshot of one time-step in a model of mantle convection. Colors closer to red are hot areas and colors closer to blue are cold areas. In this figure, heat received at the core-mantle boundary results in thermal expansion of the material at the bottom of the model, reducing its density and causing it to send plumes of hot material upwards. Likewise, cooling of material at the surface results in its sinking.

Because of the temperature difference between the Earth's surface and outer core and the ability of the crystalline rocks at high pressure and temperature to undergo slow, creeping, viscous-like deformation over millions of years, there is a convective material circulation in the mantle.[12] Hot material upwells, while cooler (and heavier) material sinks downward. Downward motion of material occurs at convergent plate boundaries called subduction zones. Locations on the surface that lie over plumes are predicted to have high elevation (because of the buoyancy of the hotter, less-dense plume beneath) and to exhibit hot spot volcanism. The volcanism often attributed to deep mantle plumes is alternatively explained by passive extension of the crust, permitting magma to leak to the surface (the "Plate" hypothesis).[22]

The convection of the Earth's mantle is a chaotic process (in the sense of fluid dynamics), which is thought to be an integral part of the motion of plates. Plate motion should not be confused with continental drift which applies purely to the movement of the crustal components of the continents. The movements of the lithosphere and the underlying mantle are coupled since descending lithosphere is an essential component of convection in the mantle. The observed continental drift is a complicated relationship between the forces causing oceanic lithosphere to sink and the movements within Earth's mantle.

Although there is a tendency to larger viscosity at greater depth, this relation is far from linear and shows layers with dramatically decreased viscosity, in particular in the upper mantle and at the boundary with the core.[23] The mantle within about 200 km (120 mi) above the core-mantle boundary appears to have distinctly different seismic properties than the mantle at slightly shallower depths; this unusual mantle region just above the core is called D″ ("D double-prime"), a nomenclature introduced over 50 years ago by the geophysicist Keith Bullen.[24] D″ may consist of material from subducted slabs that descended and came to rest at the core-mantle boundary and/or from a new mineral polymorph discovered in perovskite called post-perovskite.

Earthquakes at shallow depths are a result of stick-slip faulting; however, below about 50 km (31 mi) the hot, high pressure conditions ought to inhibit further seismicity. The mantle is considered to be viscous and incapable of brittle faulting. However, in subduction zones, earthquakes are observed down to 670 km (420 mi). A number of mechanisms have been proposed to explain this phenomenon, including dehydration, thermal runaway, and phase change. The geothermal gradient can be lowered where cool material from the surface sinks downward, increasing the strength of the surrounding mantle, and allowing earthquakes to occur down to a depth of 400 km (250 mi) and 670 km (420 mi).

The pressure at the bottom of the mantle is ~136 GPa (1.4 million atm).[14] Pressure increases as depth increases, since the material beneath has to support the weight of all the material above it. The entire mantle, however, is thought to deform like a fluid on long timescales, with permanent plastic deformation accommodated by the movement of point, line, and/or planar defects through the solid crystals comprising the mantle. Estimates for the viscosity of the upper mantle range between 1019 and 1024 Pa·s, depending on depth,[23] temperature, composition, state of stress, and numerous other factors. Thus, the upper mantle can only flow very slowly. However, when large forces are applied to the uppermost mantle it can become weaker, and this effect is thought to be important in allowing the formation of tectonic plate boundaries.

Exploration

Exploration of the mantle is generally conducted at the seabed rather than on land because of the relative thinness of the oceanic crust as compared to the significantly thicker continental crust.

The first attempt at mantle exploration, known as Project Mohole, was abandoned in 1966 after repeated failures and cost over-runs. The deepest penetration was approximately 180 m (590 ft). In 2005 an oceanic borehole reached 1,416 metres (4,646 ft) below the sea floor from the ocean drilling vessel JOIDES Resolution.

On 5 March 2007, a team of scientists on board the RRS James Cook embarked on a voyage to an area of the Atlantic seafloor where the mantle lies exposed without any crust covering, mid-way between the Cape Verde Islands and the Caribbean Sea. The exposed site lies approximately three kilometres beneath the ocean surface and covers thousands of square kilometres.[25][26] A relatively difficult attempt to retrieve samples from the Earth's mantle was scheduled for later in 2007.[27] The Chikyu Hakken mission attempted to use the Japanese vessel 'Chikyu' to drill up to 7,000 m (23,000 ft) below the seabed. This is nearly three times as deep as preceding oceanic drillings.

A novel method of exploring the uppermost few hundred kilometres of the Earth was recently proposed, consisting of a small, dense, heat-generating probe which melts its way down through the crust and mantle while its position and progress are tracked by acoustic signals generated in the rocks.[28] The probe consists of an outer sphere of tungsten about one metre in diameter with a cobalt-60 interior acting as a radioactive heat source. It was calculated that such a probe will reach the oceanic Moho in less than 6 months and attain minimum depths of well over 100 km (62 mi) in a few decades beneath both oceanic and continental lithosphere.[29]

Exploration can also be aided through computer simulations of the evolution of the mantle. In 2009, a supercomputer application provided new insight into the distribution of mineral deposits, especially isotopes of iron, from when the mantle developed 4.5 billion years ago.[30]
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Lithosphere

Lithosphere

From Wikipedia, the free encyclopedia
 
The tectonic plates of the lithosphere on Earth
Earth cutaway from core to crust, the lithosphere comprising the crust and lithospheric mantle (detail not to scale)

A lithosphere (Ancient Greek: λίθος [lithos] for "rocky", and σφαῖρα [sphaira] for "sphere") is the rigid,[1] outermost shell of a rocky planet, and can be identified on the basis of its mechanical properties. On Earth, it comprises the crust and the portion of the upper mantle that behaves elastically on time scales of thousands of years or greater. The outermost shell of a rocky planet, the crust, is defined on the basis of its chemistry and mineralogy.

Earth's lithosphere

Earth's lithosphere includes the crust and the uppermost mantle, which constitute the hard and rigid outer layer of the Earth. The lithosphere is broken into tectonic plates. The uppermost part of the lithosphere that chemically reacts to the atmosphere, hydrosphere and biosphere through the soil forming process is called the pedosphere. The lithosphere is underlain by the asthenosphere, the weaker, hotter, and deeper part of the upper mantle. The boundary between the lithosphere and the underlying asthenosphere is defined by a difference in response to stress: the lithosphere remains rigid for very long periods of geologic time in which it deforms elastically and through brittle failure, while the asthenosphere deforms viscously and accommodates strain through plastic deformation. The study of past and current formations of landscapes is called geomorphology.

History

The concept of the lithosphere as Earth’s strong outer layer was developed by Joseph Barrell, who wrote a series of papers introducing the concept.[2][3][4][5] The concept was based on the presence of significant gravity anomalies over continental crust, from which he inferred that there must exist a strong upper layer (which he called the lithosphere) above a weaker layer which could flow (which he called the asthenosphere). These ideas were expanded by Harvard geologist Reginald Aldworth Daly in 1940 with his seminal work "Strength and Structure of the Earth"[6] and have been broadly accepted by geologists and geophysicists. Although these ideas about lithosphere and asthenosphere were developed long before plate tectonic theory was articulated in the 1960s, the concepts that a strong lithosphere exists and that this rests on a weak asthenosphere are essential to that theory.

Types

There are two types of lithosphere:
  • Oceanic lithosphere, which is associated with oceanic crust and exists in the ocean basins (mean density of about 2.9 grams per cubic centimeter)
  • Continental lithosphere, which is associated with continental crust (mean density of about 2.7 grams per cubic centimeter)
The thickness of the lithosphere is considered to be the depth to the isotherm associated with the transition between brittle and viscous behavior.[7] The temperature at which olivine begins to deform viscously (~1000 °C) is often used to set this isotherm because olivine is generally the weakest mineral in the upper mantle. Oceanic lithosphere is typically about 50–140 km thick [8](but beneath the mid-ocean ridges is no thicker than the crust), while continental lithosphere has a range in thickness from about 40 km to perhaps 280 km;[8] the upper ~30 to ~50 km of typical continental lithosphere is crust. The mantle part of the lithosphere consists largely of peridotite. The crust is distinguished from the upper mantle by the change in chemical composition that takes place at the Moho discontinuity.

Oceanic lithosphere

Oceanic lithosphere consists mainly of mafic crust and ultramafic mantle (peridotite) and is denser than continental lithosphere, for which the mantle is associated with crust made of felsic rocks.
Oceanic lithosphere thickens as it ages and moves away from the mid-ocean ridge. This thickening occurs by conductive cooling, which converts hot asthenosphere into lithospheric mantle and causes the oceanic lithosphere to become increasingly thick and dense with age. The thickness of the mantle part of the oceanic lithosphere can be approximated as a thermal boundary layer that thickens as the square root of time.
\, h \, \sim \, 2\, \sqrt{ \kappa t } \,
Here, h is the thickness of the oceanic mantle lithosphere, \kappa is the thermal diffusivity (approximately 10−6 m2/s) for silicate rocks, and t is the age of the given part of the lithosphere. The age is often equal to L/V, where L is the distance from the spreading centre of mid-oceanic ridge, and V is velocity of the lithospheric plate.

Oceanic lithosphere is less dense than asthenosphere for a few tens of millions of years but after this becomes increasingly denser than asthenosphere. This is because the chemically differentiated oceanic crust is lighter than asthenosphere, but thermal contraction of the mantle lithosphere makes it more dense than the asthenosphere. The gravitational instability of mature oceanic lithosphere has the effect that at subduction zones, oceanic lithosphere invariably sinks underneath the overriding lithosphere, which can be oceanic or continental. New oceanic lithosphere is constantly being produced at mid-ocean ridges and is recycled back to the mantle at subduction zones. As a result, oceanic lithosphere is much younger than continental lithosphere: the oldest oceanic lithosphere is about 170 million years old, while parts of the continental lithosphere are billions of years old. The oldest parts of continental lithosphere underlie cratons, and the mantle lithosphere there is thicker and less dense than typical; the relatively low density of such mantle "roots of cratons" helps to stabilize these regions.[9][10]

Subducted lithosphere

Geophysical studies in the early 21st century posit that large pieces of the lithosphere have been subducted into the mantle as deep as 2900 km to near the core-mantle boundary,[11] while others "float" in the upper mantle,[12][13] while some stick down into the mantle as far as 400 km but remain "attached" to the continental plate above,[10] similar to the extent of the "tectosphere" proposed by Jordan in 1988.[14]

Mantle xenoliths

Geoscientists can directly study the nature of the subcontinental mantle by examining mantle xenoliths[15] brought up in kimberlite, lamproite, and other volcanic pipes. The histories of these xenoliths have been investigated by many methods, including analyses of abundances of isotopes of osmium and rhenium. Such studies have confirmed that mantle lithospheres below some cratons have persisted for periods in excess of 3 billion years, despite the mantle flow that accompanies plate tectonics.[16]

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Mid-ocean ridge

Mid-ocean ridge

From Wikipedia, the free encyclopedia

Oceanic ridge, including a black smoker

A mid-ocean ridge is a general term for an underwater mountain system that consists of various mountain ranges (chains), typically having a valley known as a rift running along its spine, formed by plate tectonics. This type of oceanic ridge is characteristic of what is known as an oceanic spreading center, which is responsible for seafloor spreading. The uplifted seafloor results from convection currents which rise in the mantle as magma at a linear weakness in the oceanic crust, and emerge as lava, creating new crust upon cooling. A mid-ocean ridge demarcates the boundary between two tectonic plates, and consequently is termed a divergent plate boundary.

The mid-ocean ridges of the world are connected and form a single global mid-oceanic ridge system that is part of every ocean, making the mid-oceanic ridge system the longest mountain range in the world. The continuous mountain range is 65,000 km (40,400 mi) long (several times longer than the Andes, the longest continental mountain range), and the total length of the oceanic ridge system is 80,000 km (49,700 mi) long.[1]

Description

A close-up showing a mid-ocean ridge topography with magma rising from a chamber below, forming new ocean plate which spreads away from ridge

Mid-ocean ridges are geologically active, with new magma constantly emerging onto the ocean floor and into the crust at and near rifts along the ridge axes. The crystallized magma forms new crust of basalt (known as MORB for mid-ocean ridge basalt) and gabbro.

The rocks making up the crust below the sea floor are youngest at the axis of the ridge and age with increasing distance from that axis. New magma of basalt composition emerges at and near the axis because of decompression melting in the underlying Earth's mantle.[2]

The oceanic crust is made up of rocks much younger than the Earth itself. Most oceanic crust in the ocean basins is less than 200 million years old. The crust is in a constant state of "renewal" at the ocean ridges. Moving away from the mid-ocean ridge, ocean depth progressively increases; the greatest depths are in ocean trenches. As the oceanic crust moves away from the ridge axis, the peridotite in the underlying mantle cools and becomes more rigid. The crust and the relatively rigid peridotite below it make up the oceanic lithosphere.

Slow spreading ridges like the Mid-Atlantic Ridge generally have large, wide rift valleys, sometimes as wide as 10–20 km (6.2–12.4 mi), and very rugged terrain at the ridge crest that can have relief of up to a 1,000 m (3,300 ft). By contrast, fast spreading ridges like the East Pacific Rise are narrow, sharp incisions surrounded by generally flat topography that slopes away from the ridge over many hundreds of miles.

Formation processes

Oceanic crust is formed at an oceanic ridge, while the lithosphere is subducted back into the asthenosphere at trenches.

There are two processes, ridge-push and slab pull, thought to be responsible for the spreading seen at mid-ocean ridges, and there is some uncertainty as to which is dominant. Ridge-push occurs when the growing bulk of the ridge pushes the rest of the tectonic plate away from the ridge, often towards a subduction zone. At the subduction zone, "slab-pull" comes into effect. This is simply the weight of the tectonic plate being subducted (pulled) below the overlying plate dragging the rest of the plate along behind it.

The other process proposed to contribute to the formation of new oceanic crust at mid-ocean ridges is the "mantle conveyor" (see image). However, there have been some studies which have shown that the upper mantle (asthenosphere) is too plastic (flexible) to generate enough friction to pull the tectonic plate along. Moreover, unlike in the image above, mantle upwelling that causes magma to form beneath the ocean ridges appears to involve only its upper 400 km (250 mi), as deduced from seismic tomography and from studies of the seismic discontinuity at about 400 km (250 mi). The relatively shallow depths from which the upwelling mantle rises below ridges are more consistent with the "slab-pull" process. On the other hand, some of the world's largest tectonic plates such as the North American Plate are in motion, yet are nowhere being subducted.

The rate at which the mid-ocean ridge creates new material is known as the spreading rate, and is generally measured in mm/yr. The common subdivisions of spreading rate are fast, medium, and slow with values generally being >100 mm/yr, 100–55 mm/yr, and 55–20 mm/yr, respectively. The spreading rate of the north Atlantic Ocean is ~ 25 mm/yr, while in the Pacific region, it is 80–120 mm/yr. Ridges that spread at rates <20 a="" and="" arctic="" are="" as="" brethren.="" crustal="" different="" e.g.="" faster="" formation="" gakkel="" in="" indian="" mm="" much="" nbsp="" ocean="" on="" p="" perspective="" provide="" referred="" ridge="" ridges="" southwest="" spreading="" than="" the="" their="" they="" to="" ultraslow="" yr="">
The mid-ocean ridge systems form new oceanic crust. As crystallized basalt extruded at a ridge axis cools below Curie points of appropriate iron-titanium oxides, magnetic field directions parallel to the Earth's magnetic field are recorded in those oxides. The orientations of the field in the oceanic crust record preserve a record of directions of the Earth's magnetic field with time. Because the field has reversed directions at irregular intervals throughout its history, the pattern of reversals in the ocean crust can be used as an indicator of age. Likewise, the pattern of reversals together with age measurements of the crust is used to help establish the history of the Earth's magnetic field.

History

World Distribution of Mid-Oceanic Ridges; USGS

Discovery

Mid-ocean ridges are generally submerged deep in the ocean. It was not until the 1950s, when the ocean floor was surveyed in detail, that their full extent became known.
The Vema, a ship of the Lamont-Doherty Earth Observatory of Columbia University, traversed the Atlantic Ocean, recording data about the ocean floor from the ocean surface. A team led by Marie Tharp and Bruce Heezen analyzed the data and concluded that there was an enormous mountain chain running along the middle of floor of the Atlantic. Scientists gave the name "Mid-Atlantic Ridge" to the submarine mountain range.

At first, the ridge was thought to be a phenomenon specific to the Atlantic Ocean. However, as surveys of the ocean floor continued around the world, it was discovered that every ocean contains parts of the mid-ocean ridge system. Although the ridge system runs down the middle of the Atlantic Ocean, the ridge is located away from the center of other oceans.

Impact

Alfred Wegener proposed the theory of continental drift in 1912. He stated: the Mid-Atlantic Ridge ... zone in which the floor of the Atlantic, as it keeps spreading, is continuously tearing open and making space for fresh, relatively fluid and hot sima [rising] from depth.[3] However, Wegener did not pursue this observation in his later works and his theory was dismissed by geologists because there was no mechanism to explain how continents could plow through ocean crust, and the theory became largely forgotten.

Following the discovery of the world-wide extent of the mid-ocean ridge in the 1950s, geologists faced a new task: explaining how such an enormous geological structure could have formed. In the 1960s, geologists discovered and began to propose mechanisms for sea floor spreading. Plate tectonics was a suitable explanation for sea floor spreading, and the acceptance of plate tectonics by the majority of geologists resulted in a major paradigm shift in geological thinking.

It is estimated that 20 volcanic eruptions occur each year along earth's mid-ocean ridges and that every year 2.5 km2 (0.97 sq mi) of new sea floor is formed by this process. With a crustal thickness of 1 to 2 km (0.62 to 1.24 mi), this amounts to about 4 km3 (0.96 cu mi) of new ocean crust formed every year.[citation needed]
  • Oceanic ridge and deep sea vent chemistry
  • Age of oceanic crust. The red is most recent, and blue is the oldest.
  • Plates in the crust of the earth, according to the plate tectonics theory
  • A demonstration of magnetic striping

List of oceanic ridges

List of ancient oceanic ridges
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