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Research author on mantle melting and recycled components in basalt.

Tests a recycled-crust reaction model for pyroxenite-bearing sources of Hawaiian basalt. Nature 434, 590–597; DOI 10.1038/nature03411.

Crust dominated by basalt and compositionally related intrusive rocks.

Rock with a composition similar to basalt; the adjective also describes related magma.

A rock formed by crystallization of magma or lava. Intrusive bodies crystallize below the surface, while extrusive rocks solidify at or near the surface. Granite is a common intrusive example and basalt a common extrusive example.

Lunar material enriched in potassium, rare-earth elements and phosphorus, together with other incompatible elements. It is associated with evolved residual liquids from lunar differentiation; accompanying uranium and thorium provide radiogenic heating. KREEP is a geochemical designation, not itself a synonym for all mare basalt or for a unique physical crustal layer.

A compositional description for magnesium- and iron-rich silicate material; basalt and gabbro are common mafic rocks.

Crust dominated by mafic rocks, commonly basalt and gabbro or their metamorphosed equivalents.

A rock relatively rich in magnesium- and iron-bearing silicate minerals. Basalt and gabbro are common examples; their magnetic susceptibility depends particularly on magnetic-mineral abundance.

The difference $\Delta \mathbf{M}={\mathbf{M}}_{\mathrm{target}}-{\mathbf{M}}_{\mathrm{reference}}$ between a target region and the reference material it replaces. A weakly magnetic crater infill inside magnetic basalt can have a negative induced contrast, even though both materials have positive magnetic susceptibility. The sign and direction of this contrast, including remanent magnetization, determine a total-field magnetic anomaly together with geometry and the ambient field. Geophysical density contrast does not determine magnetic contrast.

Basalt erupted at oceanic intraplate volcanic islands, used to investigate mantle heterogeneity and possible recycled components.

The decisive change was from moving continents within a supposedly permanent ocean floor to moving coherent tectonic plates that include both continental and oceanic lithosphere. Several independent observations made that change testable.

Continental connections supplied the initial problem. In the early twentieth century, Alfred Wegener assembled matching continental margins, correlatable rock belts and mountain chains, shared fossil assemblages, and paleoclimate indicators into the case for continental drift. Reassembling Gondwana, for example, makes geographically separated glacial deposits and striation directions compatible with one ice-sheet system. Fits at the continental shelf edge are physically more meaningful than exact fits to modern coastlines, which have been modified by erosion and sea-level change. The evidence suggested relative movement, but did not justify the proposed forces: continents cannot readily plough independently through strong oceanic rock. A convincing mechanism and independent motion measurements were missing.

Oceanography changed the boundary conditions. Echo-sounding surveys revealed a connected spreading ridge system, axial valleys, subduction trenches and long fracture zones. Seismic refraction surveys established systematic differences between continental and oceanic crust. Thin oceanic sediment cover, high ridge heat flux and the lack of very old exposed oceanic basement made permanent, ancient ocean basins increasingly implausible. Harry Hammond Hess proposed creation of oceanic lithosphere at ridges and recycling at trenches: seafloor spreading without requiring an expanding Earth. Slow mantle convection supplied a physically plausible connection between surface transport and internal heat loss. The mechanism need not have been completely quantified before its kinematic predictions could be tested.^[ref]

Magnetic records supplied a clock and a speedometer. Dated continental lavas established a sequence of geomagnetic reversals. Marine magnetometers then found elongated, approximately paired marine magnetic stripes on either side of ridges. The 1963 interpretation by F. J. Vine and D. H. Matthews combined repeated reversals with seafloor spreading: cooling basalt locks in remanent magnetization, and continuous motion converts a time sequence into a spatial sequence. If a polarity interval lasts $\Delta t$ and occupies width $w$ on one ridge flank, its half-spreading rate is $w/\Delta t$. Matching several reversals tests the predicted sequence, rather than merely fitting a single stripe.^[ref] Different apparent polar wander paths for different continents also became mutually consistent after continental reconstruction, favoring relative continental motion over independent magnetic-pole histories.

Earthquakes identified creation, translation and consumption. Improved global seismology located shallow earthquakes along ridges and transforms, and dipping Wadati-Benioff zones beneath trenches. Their progressive focal depths identified cold, coherent slabs entering the mantle. Focal mechanisms showed extension at spreading boundaries, tangential slip at transforms and convergence at trenches. In 1965, John Tuzo Wilson's transform fault interpretation made a particularly discriminating prediction: active slip occurs between offset ridge segments, with a sense that can be opposite to that inferred by treating the entire lineament as an ordinary offset fault. The outer fracture zones are inherited traces, not continuing active plate boundaries. Agreement of focal mechanisms with the transform prediction linked magnetic and seismic evidence.

The late-1960s synthesis made these observations one quantitative theory. Relative motion of rigid tectonic plates on a sphere can be described by an Euler angular velocity vector, with $\mathbf{v}=\mathbf{\omega }\times \mathbf{r}$. Spreading rates, transform azimuths and slip directions must fit the same rotation, and relative angular velocities must close around plate circuits. The resulting framework accounted for global boundary geometry and earthquake mechanisms without moving continents separately from the seafloor. Deep-ocean drilling beginning in 1968 subsequently confirmed that basement and overlying sediments generally become older away from ridges; this was a powerful confirmation of an already emerging synthesis, not a prerequisite discovered decades earlier.^[ref]

Thus plate tectonics succeeded through agreement between geological reconstruction, magnetic chronology, ocean-floor structure, seismology and rigid-plate kinematics. Modern satellite geodesy verifies motions directly, but belongs to the later confirmation of the theory, not its original 1960s evidence.

The geological record is a collection of thermometers with different sampling biases, not a direct time series of the whole mantle.

Primary melt compositions constrain melting temperatures. Reconstruct primary magma compositions from dated, suitably unaltered basalts by correcting for fractional crystallization and checking equilibrium with olivine. Experimental phase equilibrium constraints and mineral–melt thermometers then constrain the pressure and temperature of melt generation. In laboratory calibrations, a candidate melt and mineral assemblage are equilibrated at controlled pressure and temperature, quenched, and their crystals and glass analyzed. Repetition across adjacent conditions brackets the relevant phase equilibrium boundaries. Correcting for decompression and latent heat then allows comparison as mantle potential temperature. Age-controlled, non-arc basalt suites are especially valuable for separating changes in ambient mantle from exceptional mantle plumes. The 2010 reconstruction by Claude Herzberg and colleagues combines this approach with the temporal distribution of high-Mg parental melts.^[ref]

Komatiites provide evidence for unusually hot melting, with qualifications. Many ancient high-Mg lavas, their olivine compositions and spinifex textures indicate large degrees of melting and high parental temperatures. Their abundance in Archean terrains and rarity in the recent record are consistent with a hotter past. However, a hot mantle plume, dissolved water, unusual source composition or crystal accumulation can change the interpretation. A komatiite temperature is not automatically the ambient mantle temperature, and spinifex texture alone is not a calibrated thermometer.

Melting residues and crustal products provide independent checks. Depleted peridotite mantle xenoliths, preserved mineral chemistry and compatible-element depletion constrain former melt fractions. More extensive melting can produce thicker mafic crust and more refractory residual mantle; these are predictions to compare with preserved rock packages and seismic or petrological constraints on ancient lithosphere. Metamorphic facies and pressure–temperature paths provide contextual evidence about geothermal regimes, but tectonic burial, decompression and preservation can mimic temporal changes. They do not directly measure $T_p$.

The most persuasive inference is therefore agreement between dated melt thermometry, melting residues and their predicted crustal consequences, while separating source fertility, volatile content, mantle plume anomalies, alteration and selective preservation. Those controls explain why hotter ancient melting is better established than one exact global cooling curve.

The positive centers in the supplied maps are examples of lunar mascons, especially associated with large impact basins. An empty excavation by itself removes mass and would not explain a positive central anomaly. Basin formation can thin the low-density lunar crust and raise denser lunar mantle beneath the floor; impact melting, collapse and later cooling can also create dense material and lock in a mass excess. Where a basin was later flooded, dense mare basalt adds another contribution. Mare infill is therefore one possible contributor, not a necessary explanation of every mascon.^[ref]

The difference between the maps is the modeled topographic contribution. On a common reference surface, write schematicallyThe free-air gravity anomaly retains relief-related attraction. Rims, excavated floors and central peaks therefore contribute fine structure. The Bouguer gravity anomaly removes the attraction predicted from the measured topography at an assumed crustal density, making internal density and interface structure more apparent. A low basin floor has a negative topographic contribution; subtracting it can enhance an underlying positive central anomaly. Removing a positive rim contribution can weaken a rim high. The supplied lower map is correspondingly smoother at short wavelengths and shows prominent residual basin centers and surrounding rings. This is a change of gravity reduction, not a physical change of the craters.^[ref]

For scale, a Bouguer slab approximation gives $\Delta g\simeq 2\pi G\Delta \rho d$. A $5$ km layer with geophysical density contrast $300\mathrm{kg}{\mathrm{m}}^{-3}$ produces about $63$ mGal above an ideal infinite slab. Actual basin geometry, burial, porosity and mantle uplift require a finite three-dimensional model. Smaller craters can have central lows, damaged porous interiors or more complex rings; the maps do not imply that every crater has a positive anomaly or that positive anomalies measure basalt thickness uniquely.

Choose an ocean–continent convergent margin. The essential relationship is recycling of oceanic lithosphere by subduction, together with production and modification of crust in a volcanic arc.

The incoming tectonic plate consists of thin oceanic crust, chiefly basalt and gabbro, above colder lithospheric mantle. As that oceanic lithosphere approaches the subduction zone, bending produces the subduction trench. Its cold interior can become denser than the surrounding mantle; the resulting slab pull helps sustain subduction. The descending slab carries both oceanic crust and lithospheric mantle into the interior. Thus the recycled unit is an entire portion of lithosphere, rather than just its crust.

At the subduction trench, some incoming sediment is scraped off and stacked by thrust faults into an accretionary prism. Material can also be added beneath the overriding plate by tectonic underplating. A forearc basin collects sediment between the accretionary prism and the volcanic arc. This is an accreting example: other convergent margins lose upper-plate material through tectonic erosion, so an accretionary prism is not universal.

The cold slab undergoes increasing pressure and metamorphism. Conversion of basaltic rocks and gabbro towards eclogite increases their density, while slab dehydration releases water from hydrated minerals. The rising fluid lowers the solidus of hot peridotite in the mantle wedge, causing flux melting. The usual source of arc magma is therefore the fluid-modified mantle wedge; extensive slab melting is not required. Thermal-petrological calculations show that the depth and amount of slab dehydration depend on the incoming slab's thermal structure.^[ref]

Buoyant magma rises through the overriding plate, freezes in igneous intrusions or erupts in the volcanic arc. Fractional crystallization, interaction with existing continental crust and repeated intrusion can produce more evolved arc magmas and add new crust. Some material merely reworks old crust, and simultaneous tectonic erosion can remove it: volcanism alone does not measure net continental growth. Tectonic shortening, folding and thrust faults can thicken and raise the upper plate.

The Wadati-Benioff zone marks earthquakes along and within the cold descending slab. Shallow interface earthquakes reflect unstable slip where the plates are coupled; deeper earthquakes require mechanisms such as dehydration embrittlement and are not simply dry shallow friction operating at every depth. In contrast, the hot mantle wedge deforms mainly by material creep. Together, the subduction trench, accretionary prism, forearc basin, volcanic arc and inclined seismicity connect surface features to the processes recycling and creating crust.

The volume alone does not determine a unique change of elevation. A local isostatic response to an added surface layer depends on its thickness and density, so the Deccan Traps' footprint must be specified or retained as a parameter.

Let the deposited basalt have uniform thickness $H=V/A$ and density ${\rho }_b$, where $V=10^{15}{\mathrm{m}}^3$ and $A$ is the loaded area. Let $w$ be downward displacement of the original land surface, and let ${\rho }_m$ be the density of displaced mantle. Assume a subaerial, laterally extensive deposit, negligible air density, unchanged thickness and density of the pre-existing crust, and complete local isostatic adjustment. Take the magma supply to be external to the local compensation column, and neglect cooling, erosion, magmatic underplating and any deep thermal or source-depletion anomaly.

Choose a compensation depth beneath all the displaced interfaces. The added column mass per unit area is ${\rho }_{b}H$. Downward translation of the pre-existing crust replaces a thickness $w$ of mantle within the column, reducing its mass by ${\rho }_{m}w$. Equality of hydrostatic pressure at the compensation depth therefore givesThe newly formed top of the basalt stands $H$ above the depressed original surface. Consequently the change of the region's top-surface elevation isThus the original substrate subsides, but the top of the deposit rises if ${\rho }_b<{\rho }_m$. Distinguishing these two surfaces prevents reporting the large downward displacement as the change in plateau elevation.

For an explicitly illustrative choice $A=10^{12}{\mathrm{m}}^2$, ${\rho }_b=2900\mathrm{kg}{\mathrm{m}}^{-3}$ and ${\rho }_m=3300\mathrm{kg}{\mathrm{m}}^{-3}$, the deposit is 1000 m thick, the substrate subsides about 879 m, and the final surface rises about 121 m. This is a conditional estimate, not a number fixed by the paper. Half that footprint would double all three heights under the same assumptions. For a nonuniform deposit, local isostasy gives the same relation pointwise with $H(x,y)$; the total volume still cannot specify a unique local elevation.

The expected broad contrast is a less mature, commonly thinner early juvenile continental crust, with more mafic material and abundant sodic tonalite-trondhjemite-granodiorite, rather than the more differentiated, potassium-rich granitoid component characteristic of much Phanerozoic continental crust. This is a secular tendency, not a single stratified column applicable throughout the Archean.

Archean terrains commonly expose tonalite-trondhjemite-granodiorite gneiss associated with greenstone belts containing basalt, komatiite and metasedimentary rocks. Many later granitoids are richer in potassium, including granite and granodiorite. The proportions changed during the Archean: a reconstruction from trace element ratios in sedimentary rocks infers a markedly mafic upper continental crust before about 3.0 Ga and a more felsic upper continental crust by 2.5 Ga. Thus “all Archean continental crust was mafic” would be wrong. A preserved tonalite-trondhjemite-granodiorite-dominated nucleus can be strongly felsic, including at depth; its composition is not identical to that of its original juvenile precursor.^[ref]

Thickness needs a stated comparison. An influential seismic compilation found preserved Archean crust about 27–40 km thick, commonly without the thick high-velocity basal layer of its 40–55 km thick Proterozoic comparison provinces. That supports a relatively thin endmember, but is not a measurement of every original Archean continental crust, nor a demonstration that every Archean region was thinner than today's 35–40 km reference column. Tectonic thickening, magmatic underplating, erosion, delamination and preservation bias all alter what survives. The very thick lithospheric mantle beneath a craton must also not be mistaken for thick continental crust.^[ref]

The transferable answer is therefore more mafic early juvenile precursors, widespread sodic tonalite-trondhjemite-granodiorite production, and variable thickness evolving through time; late Archean felsic nuclei can already resemble substantial modern continental crust.

Hotter Archean mantle would generally produce thicker, more magnesium-rich oceanic crust, together with a more strongly depleted residual mantle. Greater melt production, not merely thermal expansion of an unchanged modern column, is the principal reason for the increased crustal thickness.

At a spreading ridge, upwelling peridotite undergoes adiabatic decompression melting. A higher mantle potential temperature makes the ascending mantle intersect the solidus at greater pressure and depth. It therefore melts over a longer decompression interval and generally reaches a larger integrated melt fraction. With comparable melt focusing, more melt is supplied per unit area of newly formed oceanic crust.

Modern normal oceanic crust is about 6–7 km thick and contains basaltic lavas, a sheeted dike complex, and gabbro with mafic cumulates at depth. An illustrative reconstruction for hot ambient Archean mantle gives about 25–35 km of oceanic crust, rather than a universal fixed thickness. Its primary liquids are more magnesium-rich than typical modern primary ridge liquids; their crystallization produces more olivine- and pyroxene-rich cumulates. The extracted liquid and its crystal products must be distinguished: the entire final oceanic crust need not be a homogeneous komatiite layer.^[ref]

The residual peridotite loses more melt components, becoming more refractory and commonly harzburgite-rich. Thick buoyant basaltic crust, its later eclogitization, and the depleted underlying mantle can change the mechanics of subduction. A hotter mantle does not by itself establish how fast plates moved or whether a globally connected modern plate tectonics regime operated.

Komatiites, with very magnesium-rich primary liquids and high inferred eruption temperatures, provide evidence for hot melting, but may sample especially hot mantle plumes. Estimates of ambient mantle conditions also use less exceptional non-arc basalts. Variable source composition, water content, cooling, melt extraction and later alteration all limit the precision of a single inferred Archean thickness. The robust prediction is a larger melting column and melt yield, thicker igneous oceanic crust, and a more depleted residue.

I would not assign the lunar mare an automatic terrestrial 100–125 km compensation level. For present-day mare basins, a local equal-column-mass model may have no physically valid compensation depth; first test whether lithospheric flexure supports the load. For ancient thermally driven subsidence that did relax, choose a common datum beneath the cooling layer in deformable lunar mantle appropriate to that epoch.

The Moon is smaller than Earth and has lost a larger fraction of its original heat. Its cold, mechanically strong outer shell can be much thicker relative to its radius than the terrestrial oceanic lithosphere. Lunar mare basalts filled pre-existing impact basins: their loading, basin excavation, mantle uplift, cooling and the evolving strength of the lithosphere all matter. A thick elastic lithosphere can support lateral stress differences, so equal column mass need not hold even after a long time. The strong gravity anomalies of lunar mascons, measured by GRAIL, are evidence that the present mare-basin system is not everywhere locally compensated.^[ref]

For an explicitly assumed ancient relaxed thermal model, estimate its thermal penetration depth from thermal diffusivity and the elapsed cooling time, infer the contemporaneous mechanical lid from gravity-topography admittance or basin deformation, and place $z_c$ below that layer where lateral pressure anomalies relax. For example, a hypothetical several-hundred-kilometre-thick lid requires a deeper datum than that lid, not the shallow crust-mantle boundary. The paper supplies neither an epoch-specific thermal model nor a load wavelength, so a unique kilometre value is not determined. If the chosen datum is a substantial fraction of the lunar radius, spherical geometry and radial variation in gravity must replace the flat, constant-gravity column approximation.

Does the Moon have an asthenosphere? Not a demonstrated shallow, globally connected Earth-like layer underlying mobile plates; but “no weak or partially molten mantle anywhere, ever” is too strong. Apollo lunar seismology indicates a cold outer interior and more attenuating deep regions. Reanalyses support a deep low-velocity or partially molten region near the lunar core, while lunar laser ranging and tidal Love numbers constrain deep deformability and tidal dissipation. Such a deep weak region is sometimes called a lunar asthenosphere; it is not automatically the correct local compensation surface for a mare load.^[ref]

The inferred melt fraction and distribution remain model dependent: Dissipation at tidal and seismic frequencies in a melt-free Moon by Francis Nimmo and others demonstrates that much of the dissipation can also be explained by hot solid-state anelasticity, although its model does not reproduce every frequency-dependent constraint. The Moon was hotter during mare emplacement and could relax loads differently then. Choose the datum from the relevant thermal and mechanical history, or use a flexural model when local isostasy fails.

At a convergent margin, the incoming oceanic lithosphere is a chemically and mechanically layered package: sediments, oceanic crust and lithospheric mantle. Subduction does not deliver this whole package unchanged to one destination. Solids can be accreted, underplated, exhumed or recycled deeply, while released fluids and melts carry selected components to the overriding plate. The outcome depends on incoming composition and hydration, slab age, convergence rate, thermal structure, interface coupling and subsequent tectonic history.

The incoming material already contains a history. Sediments include terrigenous debris, biogenic carbonates, siliceous deposits, clay minerals, pore water and organic matter. Oceanic crust consists mainly of basalt, sheeted dikes and gabbro; hydrothermal alteration can add water, carbonates and mobile elements before subduction. Beneath it, residual peridotite forms the lithospheric mantle. Faulting at a spreading environment, a transform boundary or an outer-rise bending region can let seawater enter and cause serpentinization. The lithospheric mantle is not uniformly hydrated to its full thickness. Consequently, sediment thickness alone does not measure the incoming volatile inventory.

Sediments: separation starts near the trench. Compaction and increasing effective pressure expel pore fluids, which can escape through the forearc, transport dissolved constituents and influence pore pressure on the plate interface. Some sediments are scraped off into an accretionary prism, shortened by thrusts and retained in the upper plate. Others pass beneath the leading edge, then detach and undergo tectonic underplating beneath the forearc. Deeper underplated material may return to the surface during tectonic exhumation, preserving a record of burial and deformation. At a sediment-starved erosive margin, little prism may form, and tectonic erosion can entrain material from the overriding plate as well. Thus net continental gain and net continental loss are both possible; they can even alternate at one margin.

Sediments that remain attached to the slab undergo metamorphism, dehydration, and potentially decarbonation or carbonate dissolution. The released fluids can transport boron, barium, strontium, lead and other soluble constituents. Carbon can leave as dissolved species, remain in carbonate, or occur as reduced carbon depending on redox conditions and reaction history. Along sufficiently hot paths, sediment-derived melts or mixtures of sediment and altered rock can transfer material into the mantle wedge; buoyant mixed material may rise as a mélange diapir. Cold paths can instead retain more solid material and some volatile-bearing phases to greater depths. These are alternative reaction and transport pathways, not a single compulsory sequence. A carbon-budget study explicitly includes fluid dissolution, sediment diapirism and storage in the upper plate, showing why volcanic carbon dioxide alone cannot close the whole subduction carbon budget.^[ref]

Oceanic crust: metamorphic transformation and chemical filtration. Altered basalt and gabbro are progressively buried. Depending on their pressure-temperature paths and reaction kinetics, they can pass through blueschist facies or warmer assemblages and eventually form eclogite, dominated by garnet and omphacite. Eclogitization increases density and can strengthen negative slab buoyancy, but a kinetic delay can leave metastable rock. It is incorrect to assign every slab the same sequence or the same fixed depths for these changes.

During burial, breakdown of hydrous phases releases water-rich fluids; different mineral assemblages retain different quantities of water. The fluids migrate through or along the interface and into the mantle wedge, where they cause metasomatism and lower the solidus of peridotite. Flux melting of the mantle wedge generates much of the primary magma feeding a volcanic arc. This is different from saying that the entire descending oceanic crust melts beneath the arc. Most of its refractory mass can remain solid. Partial melting of oceanic crust is more plausible for unusually hot or young slabs, or other favorable thermal conditions, and produces a distinct crustal contribution.

The depth of dehydration is controlled by thermal structure and composition. Global thermal-petrological calculations, for example, predict more efficient deep water transport in old, fast, cold slabs than in hot ones, with substantial differences between volcanic crust, gabbro and hydrated peridotite. Such a calculation supports conditional deep retention; its modeled proportions should not be imposed on every margin.^[ref]

Mechanical detachment adds more branches. Upper-crustal slices, seamounts or oceanic plateaus may be accreted or underplated instead of continuing down the slab. A complete or partial ophiolite can be emplaced onto another plate by obduction in suitable convergence histories. Most intact oceanic crust is not saved this way. High-pressure fragments may be entrained in an interface channel or mélange and exhumed; their small preserved volume should not be mistaken for the dominant global mass flux. Magma generated after slab-derived transfer may also pond and crystallize beneath the upper crust by magmatic underplating, which is a different process from mechanical tectonic underplating of solid slices.

Lithospheric mantle: residual rock and a variable water carrier. Most incoming lithospheric mantle is depleted peridotite left after extraction of ridge melts. Its cold, dense, mechanically coherent portion contributes to slab pull and generally sinks with the plate. Hydrated portions contain serpentine and possibly other hydrous minerals. Breakdown of these phases releases water at conditions that depend on the thermal path, feeding further metasomatism of the mantle wedge and sometimes dehydration embrittlement. The latter is a proposed contributor to intermediate-depth earthquakes, not an exclusive explanation for all such events. Very cold paths may transfer some water onward in high-pressure hydrous phases or as hydrogen incorporated in nominally anhydrous minerals.

Do not confuse the incoming hydrated lithospheric mantle with newly serpentinized forearc mantle in the overriding plate. The latter receives slab-derived fluids and can temporarily store their water. Hydrated, relatively buoyant interface material can facilitate return flow and tectonic exhumation, while the colder deeper bulk of the lithospheric mantle continues down. Some peridotite may also be accreted or obducted with an ophiolite. These exchanges mean that even an exposed serpentinite requires structural and chemical context before it can be assigned to its original plate.

Deep recycling and eventual return. The residual eclogite, remaining sediments and lithospheric mantle can enter the deeper mantle. A slab can temporarily stagnate around the mantle transition zone or penetrate into the lower mantle, depending on buoyancy, phase transformations, geometry and mantle viscosity. Crustal material can remain as compositionally distinct domains, mix, or react with surrounding peridotite; it need not retain the near-surface eclogite assemblage indefinitely as pressure rises. Subsequent mantle convection can return recycled components toward spreading regions or mantle plumes, where partial melting incorporates them into new basalt. A primary study of Hawaiian basalt illustrates a specific model in which melts derived from recycled crust react with peridotite to form a pyroxenite source. It is an interpretation of combined chemical evidence, not a unique label attached to every enriched basalt.^[ref]

How the pathways can be observed. Seismic reflection profiles, drilling, wedge geometry and balanced structural sections constrain accretion, fluid escape and tectonic erosion. High-pressure metamorphism, mineral zoning, radiometric dating and pressure-temperature-time paths in exhumed rocks reveal burial, reaction and return. Their reconstructable histories are especially useful when deformation and fluid inclusions are linked to the dated mineral growth. Earthquake distributions, seismic tomography, seismic attenuation and electrical conductivity constrain the geometry and physical state of the slab and mantle wedge, although fluids, melts and temperature can produce nonunique signals.

Arc magmas carry complementary chemical evidence: relatively mobile barium, boron, lead or strontium can trace fluid-mediated transfer, whereas sediment-sensitive thorium and isotope ratios can help identify other contributions. Element mobility changes with fluid composition and pressure-temperature conditions, so no one element is an infallible tracer. Boron isotope ratios, radiogenic strontium isotope ratios and lead isotope ratios should be interpreted with measured incoming sediments and altered oceanic crust. Arc gases and forearc fluid discharge constrain volatile return; ocean-island basalts, mantle xenoliths and diamond inclusions provide possible evidence of deeper recycling. Each observation samples only part of the system.

For any element $X$, a useful control-volume statement isHere permanent accretion is counted outside the chosen reactive control volume, return includes fluid, melt and exhumed-solid exports, and transient underplating or reservoir growth contributes to storage until exported. For a different boundary choice, the terms must be regrouped consistently. The three incoming layers can have very different transfer efficiencies for bulk rock, water, carbon and a particular trace element. A small mass of escaping fluid can dominate an element's return while most of the solid slab continues into the deep mantle.