Where Does New Lithosphere Form

Oceanic Lithosphere

The age t of the oceanic lithosphere is a proxy of its thermal state and thus of its ductile force.

From: Treatise on Geophysics , 2007

The drapery

Kent C. Condie , in Earth every bit an Evolving Planetary System (Fourth Edition), 2022

Oceanic lithosphere

Oceanic lithosphere is produced at ocean ridges and cools, thickens, and increases in age equally it moves away from ridges. The standard model involves cooling by conduction and increasing in thickness until nigh 70 Myr, reaching a maximum thickness of near 120 km. In contrast, the underlying asthenosphere is mixed by shallow convection and is thought to convert to lithosphere by cooling with little or no alter in composition. Between the oceanic lithosphere and asthenosphere is a thermal boundary layer nearly 80 km thick in which small-scale convection occurs (Fig. 4.6). As expected, hotspots (plumes), such as Hawaii and Iceland, are associated with slow velocities between 50 and 200 km deep. Thermal and geochemical modeling has shown that oceanic lithosphere can be thinned by as much as l km past extension over curtain plumes (White and McKenzie, 1995).

Seismic anisotropy in oceanic lithosphere develops in response to the alignment of olivine and pyroxene accompanying seafloor spreading with the [100] axes of olivine and [001] axes of orthopyroxene oriented normal to ridge axes (the higher velocity direction) (Fischer et al., 2010; Kodaira et al., 2014). Supporting evidence for alignment of these minerals comes from studies of ophiolites and upper-curtain xenoliths, and menses patterns in the oceanic upper mantle can be studied by structural mapping of olivine orientations (Nicolas, 1986). The mechanism of mineral alignment requires upper-mantle shear catamenia, which aligns minerals by dislocation glide. The crystallographic glide systems have a threshold temperature necessary for recrystallization of about 900°C, which yields a thermally defined lithosphere depth like to that deduced from seismic data (~ 100 km). Creep actively maintains mineral alignment below this boundary in the LVZ, and information technology is preserved in a fossil state in the overlying lithosphere. Because this layer resides in the same depth range in the Pacific, regardless of plate age, information technology suggests a rather compatible lithosphere thickness, whereas cooling models predict thickening of the lithosphere with age (Fischer et al., 2010). Although notwithstanding not understood, it is possible that the abiding depth range of the anisotropic layer reflects gradual cooling of the mantle below the aging oceanic lithosphere. Changes in hydration, fertility, and amount of melting below this lithosphere may also contribute to difference from the standard cooling model.

So where do we stand at nowadays in our agreement of the oceanic lithosphere? Anisotropy in surface-moving ridge tomography, the relatively constant fast seismic hat thickness below the Pacific, and the lack of historic period dependence for the depth of the base of the hat are all consistent with oceanic lithosphere that is a dry out, chemically depleted layer overlying a hydrated, fertile, and probably partially melted asthenosphere (Fischer et al., 2010). This interpretation also explains the constant depth to the acme of a high-electrical conductivity layer observed almost the East Pacific Ascent. Regardless of the factors that control the thickness of the oceanic lithosphere, the relatively rapid velocity gradients required at its base of operations past seismic wave velocity distributions indicate that the base of operations of the lithosphere cannot be explained by thermal gradients lonely; a contrast in hydration, fertility, and/or degree of melting is also required.

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The Mantle

Kent C. Condie , in World as an Evolving Planetary System (2nd Edition), 2011

Oceanic Lithosphere

Oceanic lithosphere is produced at ocean ridges and cools, thickens, and increases in age every bit it moves away from ridges. The standard model involves cooling past conduction and increasing in thickness until virtually 70 Ma, reaching a maximum thickness of almost 120 km. In contrast, the underlying asthenosphere is mixed past shallow convection and is thought to convert to lithosphere by cooling with little or no change in composition. Between the oceanic lithosphere and asthenosphere is a thermal purlieus layer about 80 km thick in which small-scale convection occurs (Figure four.six). Every bit expected, hotspots (plumes), such every bit Hawaii and Republic of iceland, are associated with irksome velocities between fifty and 200 km deep. Thermal and geochemical modeling has shown that oceanic lithosphere can exist thinned by every bit much as l km past extension over drapery plumes (White & McKenzie, 1995).

Both P- and Southward-wave splitting is observed in the depth range of eighty–220 km in the oceanic lithosphere. This anisotropy develops in response to the alignment of olivine and pyroxene that accompanies seafloor spreading with the [100] axes of olivine and [001] axes of orthopyroxene oriented normal to ridge axes (the higher velocity direction) (Estey & Douglas, 1986). Supporting evidence for alignment of these minerals comes from studies of ophiolites and upper mantle xenoliths, and flow patterns in the oceanic upper mantle tin be studied by structural mapping of olivine orientations (Nicholas, 1986). The mechanism of mineral alignment requires upper pall shear catamenia, which aligns minerals past dislocation glide. The crystallographic glide systems have a threshold temperature necessary for recrystallization of about 900°C, which yields a thermally defined lithosphere depth similar to that deduced from seismic data (~100 km). Creep actively maintains mineral alignment beneath this boundary in the LVZ, and it is preserved in a fossil land in the overlying lithosphere. Together with a relatively constant seismic lid thickness below the Pacific, results advise that the oceanic lithosphere is a dry, chemically depleted layer overlying hydrated, fertile, and possibly partially molten asthenosphere. Because this layer lies in the aforementioned depth range in the Pacific, regardless of plate historic period, it suggests a rather uniform lithosphere thickness, whereas cooling models predict thickening of the lithosphere with historic period (Fischer et al., 2010). Although yet not understood, it is possible that the constant depth range of the anisotropic layer reflects gradual cooling of the pall beneath the aging oceanic lithosphere. Changes in hydration, fertility, and amount of melting below this lithosphere may also contribute to deviation from the standard cooling model.

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Seafloor Processes

Carol A. Stein , in Encyclopedia of Bounding main Sciences (3rd Edition), 2019

Data for Thermal Modeling

Oceanic lithosphere forms at midocean ridges, where hot magma upwells, and then cools to class plates as the material moves away from the spreading center. As the plate cools, heat flow decreases and the seafloor deepens ( Fig. three). Nonetheless, only shallow (< 1 km) measurements of lithospheric temperatures are possible. Hence, the two principal data sets used to constrain models for the variation in lithospheric temperature with age are seafloor depths and heat flow. The depth, corrected for sediment load, depends on the temperature integrated over the lithospheric thickness. The heat menstruation is proportional to the temperature gradient. Initially seafloor depths speedily increment, with the average increase relative to the ridge crest proportional to the square root of the crustal age. Even so, for ages greater than about 50–70 My, the average increment in depth is slower and the bend is said to "flatten." Mean heat menstruation too decreases rapidly away from the ridge crest, with values approximately proportional to the inverse of the square-root of the historic period, merely afterward about 50 My this curve also "flattens."

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Geophysical Estrus Flow*

C.A. Stein , R.P. Von Herzen , in Encyclopedia of Ocean Sciences (Second Edition), 2001

Information for Thermal Modeling

Oceanic lithosphere forms at midocean ridges, where hot magma upwells, and then cools to form plates as the material moves away from the spreading center. As the plate cools, heat menses decreases and the seafloor deepens ( Effigy 3). Even so, just shallow (less than 1 km) measurements of lithospheric temperatures are possible. Hence, the two master data sets used to constrain models for the variation in lithospheric temperature with age are seafloor depths and estrus flow. The depth, corrected for sediment load, depends on the temperature integrated over the lithospheric thickness. The oestrus flow is proportional to the temperature gradient. Initially seafloor depths rapidly increase, with the average increase relative to the ridge crest proportional to the square root of the crustal age. However, for ages greater than about 50–lxx My, the boilerplate increment in depth is slower and the curve is said to 'flatten.' Hateful estrus flow besides decreases rapidly away from the ridge crest, with values approximately proportional to the changed of the square-root of the historic period, but later on nearly l My this curve as well 'flattens.'

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International Handbook of Earthquake and Engineering Seismology, Function A

T.A. Minshull , in International Geophysics, 2002

1 Introduction

The oceanic lithosphere covers approximately 60% of the Earth'south surface. Well-constrained measurements of the seismic velocity structure of the crust and upper pall beneath the bounding main floor are thin. However, the crustal structure appears to vary systematically with a rather minor number of parameters: the age of the lithosphere; the spreading rate at which the crust was formed; its position with respect to offsets in the ridge axis; and the proximity or otherwise of thermal or chemical perturbations such as mantle plumes at the time of crustal formation or later in its history. Pregnant portions of this parameter space remain to exist explored. Nonetheless, the available data do let various empirical correlations to exist made, and numerical modelling of the processes governing the formation and evolution of the oceanic crust allow some of these correlations to exist tied to the physics of these processes, then they can exist used every bit predictive tools in poorly explored areas. The printed version of this chapter is abstracted from a longer, fully referenced version on the attached Handbook CD, to which the reader is referred for further details.

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Ocean floor tectonics

C.Chiliad.R. Fowler , in Regional Geology and Tectonics: Principles of Geologic Assay, 2012

26.one Introduction

The oceanic lithosphere, which makes upwards almost 2-thirds of the surface of the solid Globe, has been formed along the mid-ocean ridges during the past 180 Ma or so. The surface of the solid World is bimodal with a hateful state elevation of 0.84 km, and a mean depth for the oceans of 3.8 km. Bathymetric profiles and side scan sonar data illustrate the calibration and rugged nature of the seabed topography ( Figs. 26.1 and 26.two). The mid-body of water ridge axes are at a rather standard depth of two.5 km and the bounding main basins (often called "abyssal plains") are at depths of some five–6 km. The deepest point on the surface of the Earth, the Challenger Deep in the Marianas Trench (Fig. 26.three), is over 11.5 km below sea level, and Mauna Kea on the isle of Hawaii rises to 4.ii km in a higher place sea level from the 5 km deep ocean basin. Although oceans obscure the seabed and over time sediment covers much of the most rugged topography, much of the detail of its surface is visible in both sea surface and gravity variations (Figs. 26.two and 26.3). Pelagic sedimentation rates are very low (a few metres per million years), so sediment accumulations in the deep ocean rarely exceed 1 km, in contrast to the thick passive margin deposits (Fig. 26.4).

Along passive margins, the continental shelf is typically many hundreds of kilometres wide, is covered by shallow water and is underlain by continental chaff. The seaward extent of the shelf, the shelf break, is marked by an abrupt increment in gradient from ~0.1° to the ~3° gradient of the continental slope. Along the base of operations of the continental gradient, the slope decreases once more and the h2o gradually deepens to the abyssal plain – this region is the continental rise. The transition from continental to oceanic chaff lies beneath the continental slope. Along active margins, the continental shelf is often narrow. Where the plate boundary is a transform fault, the seabed characteristically drops rapidly from the shelf to oceanic depths, simply subduction zones are usually bordered by a trench, frequently many kilometres deep (Figs. 26.1B and 26.2).

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Mantle Dynamics

A. Davaille , A. Limare , in Treatise on Geophysics, 2007

seven.03.6.three Stagnant-Lid Authorities and Lithosphere Cooling

The oceanic lithosphere thickens at it moves away from a mid-ocean ridge ( see Capacity 7.07 and 6.05). It has been proposed that common cold thermal instabilities develop in its lower part. Such 'modest-scale convection' (SSC) was starting time invoked to explain the flattening of heat flux and bathymetry of the oceans at old ages (Parsons and Sclater, 1977; Parsons and McKenzie, 1978) and later several other phenomena occurring on different time and length scales, such as minor (150–500 km) geoid anomalies in the central Pacific (due east.thousand., Haxby and Weissel, 1986) and central Indian (Cazenave et al., 1987) oceans, differences in subsidence rates (Fleitout and Yuen, 1984; Cadet, 1987; Eberle and Forsyth, 1995), ridge segmentation (e.one thousand., Sparks and Parmentier, 1993; Rouzo et al., 1995), delamination of the lithosphere under hot spots (Sleep, 1994; Moore et al., 1998; Dubuffet et al., 2000), and patterns of anisotropy beneath the Pacific bounding main (Nishimura and Forsyth, 1988; Montagner and Tanimoto, 1990; Davaille and Jaupart, 1994; Ekström and Dziewonski, 1998).

Davaille and Jaupart (1994) applied the scaling laws derived from their experimental results, using for the curtain a Newtonian creep law with activation enthalpy H:

[13] $ν (T) = μ_{r} exp (\frac{H}{RT})$

For H varying between 150 and 500 kJ mol⁻¹ (the value depends on the creep machinery and the water content; run across Chapter two.14), ΔT _eff ranges between 310 °C and 90 °C according to [eleven] and [13]. Since pocket-sized-scale instabilities develop when the local Rayleigh number based on the characteristics of the unstable part of the lithosphere (temperature, viscosity, thickness) exceeds a critical value, the onset of convection depends merely on the thermal structure of the lithosphere, and on the rheology of curtain textile. Following Davaille and Jaupart (1994), the onset time can exist written every bit

[14] $τ_{c} = \frac{a}{κ} \times {(\frac{α g Δ T_{eff}}{κ ν_{m}})}^{- 2 / 3} \times {[1 + \frac{b}{f (Δ T / Δ T_{eff})} (\frac{Δ T}{Δ T_{eff}} - 1)]}^{2}$

where ν_m is the viscosity of the asthenosphere at the mantle temperature T _m, a = 51.84 and b = 0.3013 are ii constants determined from the laboratory experiments, and ΔT is the temperature drop across the lithosphere. The part f depends on the cooling model and for a half-space conductive cooling is given in Figure 36(a) . For a thin stagnant lid (or ΔT/ΔT _eff < 3), f ∼ 1. This approximation applies well to laboratory experiments and was therefore adopted by Davaille and Jaupart (1994; thin line labeled DJ94 on Figure 36(b) ). However, as pointed out by subsequent studies (Dumoulin et al., 2001; Zaranek and Parmentier, 2004), ΔT/ΔT _eff > iv for the curtain lithosphere and the approximation f ∼ 1 leads to overestimate the SSC onset fourth dimension ( Effigy 36(b) ) compared to the prediction of [14] without the approximation (thick black line on Figure 36(b) ). In the last years, several numerical studies, using either Arrhenius or exponential viscosity laws, have allowed to extend the range of ΔT/ΔT _eff studied (Choblet and Sotin, 2000; Dumoulin et al., 2001; Korenaga and Jordan, 2003; Huang et al., 2003; Zaranek and Parmentier, 2004), and to develop new scaling laws. The discrepancies betwixt the different data sets reveal the relative importance of the onset fourth dimension measurement, the type of viscosity law (10–20% of change simply), and the blazon of perturbation (initial-time/at-all-times, numerical-noise/finite-size-perturbation) introduced numerically which volition grow toward convective instabilities. The latter has the strongest influence (Zaranek and Parmentier, 2004): the longest onset times are obtained when the numerical perturbations initially introduced are smallest (Choblet and Sotin, 2000). The tendency of the laboratory experiments seems to compare best with numerical experiments which introduce thermal noise around 10^−three × ΔT at all times ( Effigy 36(b) ). Then, in the parameter range relevant for the Earth's mantle (ΔT/ΔT _eff ≥ four), numerical studies agree with [14] inside 25%. So, if the magnitude of the temperature perturbations in the lithospheric mantle is like to the laboratory one, an SSC onset time between 10 and 100 My is predicted, depending on the asthenospheric drape temperature.

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Mineral Physics

T. Irifune , T. Tsuchiya , in Treatise on Geophysics (Second Edition), 2015

ii.03.4.i Chemical Compositions and Density Calculations

Subducting oceanic lithosphere is modeled by layers of basaltic oceanic crust of ~ 6 km thickness, underlain by thicker layers (~ 50–100 km) of residual harzburgite and fertile lherzolite, which are covered with sparse (~ one km) terrigeneous and/or pelagic sediments. Typical chemical compositions of these lithologies are listed in Tabular array ane . Most parts of the sedimentary materials are believed to be trapped to form accretion terrains underneath island arcs upon subduction of slabs at ocean trenches, although geochemical evidence suggests certain parts of such materials may be subducted deeper into the curtain (e.g., Loubet et al., 1988). At to the lowest degree part of the bottom warmer lherzolite layer of a slab may too be assimilated to the surrounding mantle during subduction in the upper pall and mantle transition region, and thus the slab approaching the 660 km seismic discontinuity can reasonably be modeled by a layered structure of basaltic and harzburgitic rocks (Ringwood and Irifune, 1988).

Table one. Representative chemical compositions of lower curtain and those related to subducting slabs

	Lower mantle
	Chondrite	Pyrolite	Harzburgite	MORB	Continental crust
SiO₂	53.8	44.5	43.six	50.4	66.0
TiO_two	0.two	0.ii	–	0.6	0.5
Al₂O₃	3.8	four.three	0.7	xvi.1	15.2
Cr₂O₃	0.four	0.4	0.five	–	–
FeO	3.5	8.6	vii.8	7.7	4.5
MgO	35.1	38.0	46.iv	10.five	two.ii
CaO	2.8	3.5	0.5	13.1	4.2
Na_twoO	0.three	0.4	–	1.9	3.ix
K₂O	–	0.one	–	0.one	3.4

Chondrite, Liu (1982); pyrolite, Sun (1982); harzburgite, Michael and Bonatti (1985); MORB, Green et al. (1979); continental crust, Taylor and McLennan (1985).

The chemic composition of the lower mantle has been a major controversial issue in the mineralogy of the World'southward interior. Some (e.g., Ringwood, 1962) believe peridotitic or pyrolitic materials are ascendant in the whole mantle, while others (e.g., Anderson, 1989; Hart and Zindler, 1986; Liu, 1982) claim more Si-rich chondritic materials should be representative for the composition of the lower pall ( Table i ). The difference is based on rather philosophical arguments on the origin and subsequent differentiation processes of the Globe, which are critically dependent on the models of condensation/evaporation processes of elements and compounds in the primordial solar system and the possible formation of a deep magma body of water in the early phase of the formation of the Earth. As the elastic properties, specially those related to shear moduli, of loftier-force per unit area phases take non been well documented nether the pressure level and temperature atmospheric condition of the lower drape, it is hard to unambiguously evaluate the feasibility of these two alternative limerick models in the lite of mineral physics and seismological observations (due east.g., Bina, 2003; Mattern et al., 2005). Some efforts accept been made to directly measure the rubberband backdrop under the P, T conditions of the lower mantle (Murakami et al., 2012), but we need further data on the elastic backdrop of major minerals in the lower mantle and their pressure, temperature, and compositional dependencies to reach a decisive conclusion on the feasible chemical composition. Moreover, the knowledge of variation of temperature with depth is vital to address this issue, but this as well has non been well constrained in the lower mantle.

Hither, we assume the whole mantle is of a pyrolitic composition to address the phase transitions and associated density changes in the lower mantle, as there are significant chemical variations in the chondritic models and as well because high-pressure experimental data on these compositions take been deficient to date. We also assume the major lithologies transported into the lower drapery via subduction of slabs are of MORB and harzburgite compositions. Some recent studies advise that the continental crust may also exist delivered to the mantle transition region (Ichikawa et al., 2013; Kawai et al., 2013), only it should be difficult to subduct further into the lower mantle, considering the density relations between the continental crust material and surrounding mantle at around the 660 km discontinuity (Irifune et al., 1994; Ishii et al., 2012; Kawai et al., 2013). Phase transitions in MORB have been extensively studied down to the depths near the drapery–core purlieus. In contrast, although virtually no experimental information be on harzburgite compositions in this depth region, they tin be reasonably estimated from those available on the loftier-pressure phases with uncomplicated chemical compositions, as harzburgitic compositions have a minor corporeality of Atomic number 26 and but very minor amounts of Ca, and Al, and are well approximated past the MgO (FeO)–SiO_ii system.

We calculated the density changes in pyrolite, harzburgite, and MORB compositions using available experimental data every bit follows: The densities of the individual loftier-pressure phases appeared in these lithologies were calculated at given pressures using the thermal EoS combining third-order Birch–Murnaghan EoS and Debye theory forth an appropriate geotherm, using the PVT-EoS parameters given in Table two . The resultant density changes of individual phases forth the geotherm are depicted in Figure 9 . The density changes in the bulk rocks were then calculated using the proportions of the individual phases with force per unit area forth the geotherm.

Table 2. PVT-EoS parameters of lower curtain phases determined from diverse experimental and theoretical information and their systematics

	Mg-Pv	Fe-Pv	Al_twoO_three-Pv	Mg-PPv	Ca-Pv	MgO	FeO	SiO_ii-St	SiO₂-α-PbO_ii
Five ₀ (cm³ mol^{− 1})	24.45	25.48	24.77	24.6	27.45	eleven.36	12.06	14.02	13.81
B ₀ (GPa)	257	281	232	226	236	158	152	314	325
B′	4.02	4.02	4.iii	4.41	3.9	4.four	4.nine	four.iv	iv.2
Θ _D (K)	1054	854	1020	1040	984	725	455	1044	1044
γ	1.48	i.48	ane.48	1.55	1.53	1.five	1.28	1.34	1.34
q	1.ii	1.2	1.2	1.2	1.six	i.five	1.5	ii.4	ii.4

Mg-Pv (Fiquet et al., 2000; Shim and Duffy, 2000; Sinogeikin et al., 2004; Tsuchiya et al., 2004a, 2005b), Atomic number 26-Pv (Jeanloz and Thompson, 1983; Kiefer et al., 2002; Mao et al., 1991; Parise et al., 1990), Al₂O₃-Pv (Thomson et al., 1996; Tsuchiya et al., 2005c), Mg-PPv (Tsuchiya et al., 2004a, 2005b), Ca-Pv (Karki and Crain, 1998; Shim et al., 2000b; Wang et al., 1996), MgO (Fiquet et al., 1999; Sinogeikin and Bass, 2000), FeO (Jackson et al., 1990; Jacobsen et al., 2002; Systematics), SiO_ii (Andrault et al., 2003; Karki et al., 1997a; Ross et al., 1990; Tsuchiya et al., 2004c).

The high-T Birch–Murnaghan equation was applied just for the hexagonal aluminous phase with parameters: V ₀ = 110.07 cm³ mol^{− one}, B ₀ = 185.five GPa, B' = four (fix), dB/dT =− 0.016 GPa M^{− one}, and α ₀ = 3.44 × ten^{− 5} One thousand^{− ane} (Shinmei et al., 2005; Sanehira et al., 2005).

Although some recent studies suggest that sub-adiabatic temperature gradients are required to friction match the observed and calculated density and bulk sound velocity for pyrolitic compositions (Bina, 2003; Mattern et al., 2005), nosotros but assumed adiabatic temperature changes throughout the lower drapery (i.e., 1900 K at 660 km and 2450 Thou at 2890 km with an averaged gradient of dT/dz = ~ 0.3 Chiliad km^{− i}; e.k., Brown and Shankland, 1981) as such conclusions are non robust, given the uncertainties in both mineral physics measurements and seismological observations. In add-on, significantly sharp temperature increases are expected to occur virtually the mantle–core boundary and presumably almost the 660 km aperture as these regions are accompanied by chemical changes and form thermal boundary layers, which are also not taken into account in the present calculations.

The mineral proportion changes in pyrolite, harzburgite, and MORB compositions are shown in Figure ten , while the calculated density changes of these lithologies are depicted in Effigy 11 . The density changes at pressures lower than 30 GPa are based on an earlier estimate of Irifune (1993), using a similar method and mineral physics parameters. Although the density change in pyrolite seems to agree well with that of PREM (Preliminary Reference World Model; Dziewonski and Anderson, 1981), the latter gauge inevitably has significant uncertainties, equally this density profile is rather indirectly determined from seismic velocities with some assumptions. The calculated density values may too take significant errors, mainly due to the uncertainty in the geotherm stated in the above. Withal, mineral physics parameters to constrain the density accept been reasonably well determined, more often than not on the basis of in situ 10-ray diffraction measurements, and the differences amid these calculated density profiles are regarded as robust results.

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Tectonophysics

Donald L. Turcotte , in Encyclopedia of Physical Scientific discipline and Technology (Third Edition), 2003

III.C Subduction

Equally the oceanic lithosphere moves away from an ocean ridge, information technology cools, thickens, and becomes more than dumbo considering of thermal contraction. Even though the basaltic rocks of the oceanic crust are lighter than the underlying mantle rocks, the colder mantle rocks in the lithosphere become sufficiently dense to brand old oceanic lithosphere heavy plenty to be gravitationally unstable with respect to the hot drape rocks beneath the lithosphere. As a consequence of this gravitational instability the oceanic lithosphere founders and sinks into the interior of the world, creating the ocean trenches. This process is known as subduction and is illustrated schematically in Fig. six.

The excess density of the rocks of the descending lithosphere results in a downwardly buoyancy force. Because the lithosphere behaves elastically, information technology can transmit stresses, i.e., it can deed every bit a stress guide. A portion of the negative buoyancy forcefulness acting on the descending plate is transmitted to the surface plate, which is pulled toward the ocean trench. This is slab pull, i of the important forces driving plate tectonics.

Body of water trenches are the sites of nigh of the largest earthquakes. Earthquakes occur on the dipping error plane that separates the descending lithosphere from the overlying lithosphere. Earthquakes at ocean trenches can occur to depths of 660 km. This seismogenic region, known as the Wadati–Benioff zone, delineates the approximate structure of the descending plate.

Volcanism is likewise associated with subduction. A line of regularly spaced volcanoes closely parallels the trend of almost all the ocean trenches. These volcanoes may result in an island arc or they may occur within continental crust. The volcanoes generally lie in a higher place where the descending plate is 125 km deep, as illustrated in Fig. six. Information technology is far from obvious why volcanism is associated with subduction. The descending lithosphere is common cold compared with the surrounding pall, and thus it acts every bit a estrus sink rather than equally a heat source. The down flow of the descending slab is expected to entrain flow in the overlying mantle wedge. Even so, this flow will be primarily downward; thus, magma cannot be produced past force per unit area-release melting. One possible source of heat is frictional heating on the fault plane betwixt the descending lithosphere and the overlying mantle.

When a subduction zone is adjacent to a continent, every bit in the instance of South America, subduction zone volcanism can course great mount belts, for example, the Andes. In some subduction zones tensional stresses tin upshot in dorsum-arc, seafloor spreading and the formation of a marginal basin as illustrated in Fig. 6. An example is the Sea of Japan.

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