While temperature drop across the mantle’s basal thermal boundary layer (TBL) is likely $>$1000 K, the temperature anomaly of plumes believed to rise from that TBL is only up to a few hundred Kelvins. Reasons for that discrepancy are still poorly understood and a number of causes have been proposed. Here we use the ASPECT software to model plumes from the lowermost mantle and study their excess temperatures. We use a mantle viscosity that depends on temperature and depth with a strong viscosity increase from below the lithosphere towards the lower mantle, reaching about $10^{23}$ Pas above the basal TBL, consistent with geoid modelling and slow motion of mantle plumes. With a mineral physics-derived pyrolite material model, the difference between a plume adiabat and an ambient mantle adiabat just below the lithosphere is about two thirds of that at the base of the mantle, e.g. 1280 K vs.\ 835 K. 3-D models of isolated plumes become nearly steady-state >10-20 Myr after the plume head has reached the surface, with excess temperature drop compared to an adiabat for material directly from the CMB usually less than 100 K. In the Earth, plumes are likely triggered by slabs and probably rise preferrably above the margins of chemically distinct piles. This could lead to reduced excess temperatures, if plumes are more sheet-like, similar to 2-D models, or temperature at their source depth is less than at the CMB. Excess temperatures are further reduced when averaged over the plume conduit or melting region.

While temperature drop across the mantle’s basal thermal boundary layer (TBL) is likely $>$1000 K, the temperature anomaly of plumes believed to rise from that TBL is only up to a few hundred Kelvins. Reasons for that discrepancy are still poorly understood and a number of causes have been proposed. Here we use the ASPECT software to model plumes from the lowermost mantle and study their excess temperatures. We use a mantle viscosity that depends on temperature and depth with a strong viscosity increase from below the lithosphere towards the lower mantle, reaching about $10^{23}$ Pas above the basal TBL, consistent with geoid modelling and slow motion of mantle plumes. With a mineral physics-derived pyrolite material model, the difference between a plume adiabat and an ambient mantle adiabat just below the lithosphere is about two thirds of that at the base of the mantle, e.g. 1280 K vs.\ 835 K. 3-D models of isolated plumes become nearly steady-state $>$ 10-20 Myr after the plume head has reached the surface, with excess temperature drop compared to an adiabat for material directly from the CMB usually less than 100 K. In the Earth, plumes are likely triggered by slabs and probably rise preferrably above the margins of chemically distinct piles. This could lead to reduced excess temperatures, if plumes are more sheet-like, similar to 2-D models, or temperature at their source depth is less than at the CMB. Excess temperatures are further reduced when averaged over the plume conduit or melting region.

Controls on the deformation pattern (shortening mode and tectonic style) of orogenic forelands during lithospheric shortening remain poorly understood. Here, we use high-resolution 2D thermomechanical models to demonstrate that orogenic crustal thickness and foreland lithospheric thickness significantly control the shortening mode in the foreland. Pure-shear shortening occurs when the orogenic crust is not thicker than the foreland crust or thick, but the foreland lithosphere is thin (< 70-80 km, as in the Puna foreland case). Conversely, simple-shear shortening, characterized by foreland underthrusting beneath the orogen, arises when the orogenic crust is much thicker. This thickened crust results in high gravitational potential energy in the orogen, which triggers the migration of deformation to the foreland under further shortening. Our models present fully thick-skinned, fully thin-skinned, and intermediate tectonic styles in the foreland. The first tectonics forms in a pure-shear shortening mode whereas the others require a simple-shear mode and the presence of thick (> ~4 km) sediments that are mechanically weak (friction coefficient < ~0.05) or weakened rapidly during deformation. The formation of fully thin-skinned tectonics in thick and weak foreland sediments, as in the Subandean Ranges, requires the strength of the orogenic upper lithosphere to be less than one-third as strong as that of the foreland upper lithosphere. Our models successfully reproduce foreland deformation patterns in the Central and Southern Andes and the Laramide province.