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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.

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.

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]