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Chemical modification of residual rock by extraction of melt. In mantle peridotite, depletion commonly changes mineral proportions and trace-element contents, and can accompany volatile loss. Compositional buoyancy and water content must be assessed separately from thermal density changes when explaining a long-lived craton.

Repeated assembly and dispersal are supported by mutually consistent paleogeography, paleomagnetism, dated collision belts and subsequent rifted margins. A useful reconstruction must explain both former connections and later separation, at the same geological ages.

For the youngest well-constrained example, Pangaea, matching continental outlines become more compelling when the fit uses the continental edges rather than mobile modern shorelines. More discriminating evidence includes the continuation of mountain belts, sedimentary successions and distinctive basement rocks across now-separated continents. Correlated terrestrial fossil assemblages and the distribution of glacial deposits become geographically coherent when the continents are restored. Each observation alone can have alternatives, but a single reconstruction that satisfies several independent tests is much harder to explain by coincidence.

For older examples, the most useful surviving record lies within the cratons and their margins. Dated suture zones, ophiolites, metamorphic belts and pressure-temperature-time paths document former ocean closure and continental collision. Restoring now-separated belts to continuity tests an assembly. Subsequent continental rifting, matching dike swarms, large igneous provinces and the transition to passive margins test its breakup. Uranium-lead dating of zircon, including detrital zircon populations and sedimentary provenance, helps establish whether the proposed connections existed at the same time.

Paleomagnetism supplies an independent quantitative test. A primary remanent magnetization constrains ancient latitude and orientation. After a proposed relative rotation brings two blocks together, compatible apparent polar wander paths over an interval support their travelling together; divergence of those paths can document separation. Radiometric dating, a paleomagnetic fold test and a paleomagnetic reversal test help identify primary records and distinguish them from remagnetization. There remain important ambiguities: conventional paleomagnetism does not fix absolute longitude, and sparse ages, later deformation and true polar wander complicate a reconstruction.

The broad succession is Columbia or Nuna, assembled mainly around 1.9–1.6 billion years ago; Rodinia, assembled mainly around 1.1–0.9 billion years ago and dispersed in stages during the later Proterozoic; and Pangaea, assembled by roughly 300 million years ago and breaking apart from roughly 200 million years ago. These are broad intervals rather than synchronous global events. A quantitative Nuna reconstruction combines paleomagnetism with the histories of continental margins rather than relying only on matching rock ages.^[ref]

Repeated global peaks in dated magmatism and metamorphism offer supporting evidence, but cannot prove a supercontinent cycle by themselves: erosion, burial and preferential preservation also change the observed age distribution. The persuasive case is the recurrence of dated assembly–breakup sequences in different continental configurations, with stronger geometric constraints for Pangaea than for the older supercontinents. A Wilson cycle describes an individual ocean's opening and closure; a supercontinent cycle concerns the much larger-scale continental assembly and dispersal.

Use a rock strength envelope: at each depth plot the smaller differential stress needed for frictional failure and for sustained ductile material creep at a specified strain rate. Shallow rock generally strengthens with depth because effective confining pressure increases; sufficiently deep hot rock weakens because thermally activated material creep becomes efficient.

At shallow depth, Byerlee's law or a Mohr-Coulomb failure criterion gives a relation of the formwhere $C$ is cohesion, $\mu$ a coefficient of friction, ${\sigma }_n$ the compressive normal stress and $P_f$ the pore pressure. Increasing overburden generally increases effective stress, hence the frictional strength. Conversion from fault shear stress to differential stress depends on the stress regime and fault orientation; a sketch should not label a universal numerical slope. High pore pressure or weak fault gouge can lower the frictional strength substantially.

A simplified dislocation creep relation isHere $Q$ is activation energy, $V^{\ast }$ activation volume, $R$ the molar gas constant and $T$ absolute temperature; $A$ and $n$ depend on composition and deformation mechanism. At fixed strain rate, increasing temperature usually decreases the required stress strongly. The competing effects of pressure, water weakening of rocks, grain size and composition also matter. The intersection with frictional failure gives a brittle-ductile transition, not a change from solid to liquid.

For the Indian continental interior, a useful first sketch has an upper-crustal strength peak, a weaker lower continental crust, and renewed strength in lithospheric mantle beneath the Moho. Quartz-bearing rocks can undergo ductile material creep at conditions where colder, drier olivine-rich mantle retains appreciable strength. This creates a possible two-peak, or “jelly sandwich”, rock strength envelope. With further depth the mantle also weakens. A cold craton can retain strong lithospheric mantle to much greater depths than a tectonically heated continental region.

For the central Pacific, show thin mafic oceanic crust above a relatively thick layer of strong lithospheric mantle. There need not be a conspicuous weak lower-crustal layer. Much of the integrated strength can reside below the Moho, with a broad strength peak followed by ductile weakening towards the asthenosphere. Cooling and age control the thickness: young hot oceanic lithosphere would give a much shallower and weaker profile than an old central-Pacific plate.

The plotted Moho depths of about 38 km and 7 km, and all plotted stresses, are illustrative assumptions. “India” alone does not specify a unique geothermal gradient, water content or strain rate, and the Pacific includes plates of different ages. The robust comparison is the possible crust–mantle layering in a continent versus the thin oceanic crust and mantle-dominated strength of an old oceanic plate. The Moho is a compositional boundary; neither it nor the deepest small plotted strength is automatically the seismic lithosphere-asthenosphere boundary.

Elasticity lets the lithosphere share a load with neighbouring regions through lithospheric flexure. Beneath the centre of a localized load, subsidence is generally smaller than the local isostasy estimate, so the deposited surface stands higher. The surrounding plate also bends down, and a more distant flexural forebulge may rise. The extra support is regional bending resistance, rather than removal of the load's mass.

The controlling comparison is load width against the lithosphere's flexural length, which increases with flexural rigidity and decreases with the relevant buoyant geophysical density contrast. Equivalently, the flexural support ratio becomes large for short-wavelength loading. A very broad nearly uniform load produces little curvature in its interior, so even a strong elastic lithosphere has a nearly local isostatic adjustment there. Bending matters most near its edges. A narrow load requires more curvature and can be supported appreciably by the surrounding elastic lithosphere.

The Deccan Traps form a broad continental large igneous province. The volcanic edifices of the Hawaiian Islands are more localized loads on an old oceanic lithosphere, for which regional bending and a flexural moat are prominent. Under the usual broad-Deccan versus narrow-Hawaiian-load comparison, elastic strength matters more for the Hawaiian Islands. The more coherent mantle-dominated strength of an old oceanic plate can reinforce this contrast. It cannot be deduced simply from “oceanic” versus “continental”: a cold Indian craton can be strong, and a site's effective elastic thickness depends on its thermal and loading history.

For an initially submerged Hawaiian load, replacing seawater changes both the load and buoyant restoring contrast. At fixed deposited thickness, the added load relative to the initial water column uses the rock-minus-water density, and the restoring term uses the mantle-minus-water density. An emerged island requires separate treatment of its subaerial and submerged portions. These corrections change the local isostasy reference but do not remove the importance of load width for lithospheric flexure. The broad Hawaiian swell should also be distinguished from the narrower volcanic load: deeper thermal or dynamic support of the Hawaiian swell is not the same mechanism as elastic support of the volcanic ridge.

Finally, over long periods material creep, yielding and thermal modification can reduce elastic support. Without measured load geometries, ages and effective elastic thicknesses, the ranking is a justified qualitative expectation rather than a theorem valid for every continental and oceanic load.

The asthenosphere exists because the mantle's ability to deform is strongly dependent on temperature, pressure, composition, water and deformation mechanism. Cooling from the surface creates a strong lithosphere. Beneath it, hotter mantle deforms by material creep at much lower stress on geological timescales. The asthenosphere is predominantly solid weak rock, not a global ocean of magma. On the much shorter timescale of a seismic wave, it still supports S waves.

The hot shallow mantle can be close to its solidus. Water weakening of rocks and, in some settings, a small connected melt fraction can further lower effective viscosity. At greater depths, changes in pressure, homologous temperature, mineral structure and water storage modify the deformation laws; increasing depth does not imply indefinitely decreasing viscosity. Explaining a weak upper-mantle interval does not require every location to contain the same amount of melt.

There are several physically distinct boundaries to test. The thermal lithosphere-asthenosphere boundary describes a thermal transition, the mechanical boundary a contrast in long-term strength or viscosity, and the seismic lithosphere-asthenosphere boundary a change in seismic velocity or other wave properties. They need not have identical depths. A smooth geothermal gradient can explain broad thermal weakening, but a very sharp decrease of seismic velocity may require changes in melt fraction, hydration, composition or seismic anisotropy. A sharp shallow reflector within a craton might instead be an internal lithospheric boundary.^[ref]

Begin with a thermal null model. A half-space cooling model givesIf a thermal boundary corresponds to a fixed temperature $T_{\ast }$, its depth isso it deepens as the square root of plate age. Test that prediction against independently determined seafloor ages, bathymetry and heat flux. An oceanic plate cooling model adds finite thickness and basal heat supply, allowing the old-age depth to approach a finite value. For continents, include radiogenic heating, crustal thickness, the history of continental rifting or tectonic thickening, and advective heating by magma or mantle flow. A cold, dry craton and a recently extended volcanic region should not share the same thermal profile.

Then calculate the observable consequences of competing mechanisms rather than merely assigning a boundary to a map:

Compare models against the observed boundary depth and its sharpness, including uncertainty and resolution. An age–depth trend with matching heat flux favours thermal control; a sharp discontinuity that fails the thermal prediction but coincides with an electrical anomaly can support a melt- or hydration-related transition. Neither observation alone proves a unique cause. The spatially varying boundary is best understood as the result of thermal evolution interacting with composition, fluids and deformation history.

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.

A seismically inferred transition associated with the base of the lithosphere, often identified by a decrease in shear-wave velocity. Its depth and sharpness constrain combinations of temperature, composition, hydration, melt fraction and seismic anisotropy. A sharp discontinuity inside a craton can instead lie within its thick lithospheric mantle.^[ref]

A continental assembly containing most of the Earth's continental area or major cratons. Its reconstruction requires age-compatible geological connections and paleomagnetism; similarity of modern continental outlines alone is insufficient.