Tectonic geomorphology · deformation
Collisions thicken the crust, faults break it, earthquakes jolt it, and the whole plate bends under the weight. Every one of those leaves a mark in the landscape you can measure.
Relief asks how fast landscapes wear down. Uplift asks how they are pushed up and bent — how thick crust holds a plateau aloft, why carving valleys makes peaks rise, how faults grow by linking, how drowned forests and lifted beaches record earthquakes nobody wrote down, and why a mountain belt digs a basin beside itself.
Wide mountain belts float. Thicken the crust and the surface rises by only about a fifth of what you added — the rest becomes root. Erode it evenly and the surface drops by only a fifth of what you removed. But erode it unevenly, cutting valleys, and the peaks between them climb.
For the basic calculation — how deep a mountain’s root must be — see Relief → Isostasy. This tab picks up where that leaves off.
India has been pushing its crust under Asia for tens of millions of years. If every bit of that crust stayed in the plateau at double thickness, a cross-section balance says the plateau’s width is simply W = u·t: convergence rate times duration.
Rivers and glaciers cut valleys but leave the ridge crests alone. The crust responds to the average mass removed, so the whole block rises — including peaks that lost nothing. Drag the sliders and watch the three quantities England & Molnar (1990) insisted we keep separate.
The assumption to know: this works because valleys are spaced much closer together (kilometers) than the distance over which the lithosphere spreads a load (a hundred kilometers or more — see the Flexure tab). A single isolated canyon would not lift its rims this way.
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Textbook faults are infinite planes between rigid blocks. Real faults end, and the crust around them bends. That is why a normal fault lifts its footwall into a range, why a strike-slip fault builds hills at its bends, and why faults grow by joining up.
Normal faults follow a scaling rule: maximum throw is proportional to length, S = γL. Dawers, Anders & Scholz (1993) found this linear relation holding over three orders of magnitude of fault length on the Volcanic Tableland near Bishop, California. Link two segments and L jumps — so the throw the joined fault can carry jumps too. Tap the gaps to link or unlink.
A fault slips when shear stress reaches its strength: τ ≥ S + tanφ·(σ − P). Water pressure in the pores, P, holds the fault faces apart and cuts the effective normal stress σ′ = σ − P. Raise P and watch the point slide left toward failure.
Rock barely compresses, so conservation of mass says ∂u/∂x + ∂v/∂y + ∂w/∂z = 0: squeeze material horizontally and it must go up; stretch it and the surface drops. A strike-slip fault that ends pushes rock into the quadrants ahead of each moving block and pulls it away from the quadrants behind — four lobes of uplift and sag around the tips. Bends do the same on a larger scale: a restraining bend squeezes up hills, a releasing bend opens a basin or sag pond.
Written earthquake records go back a few hundred years at best. The landscape goes back thousands. Offset streams give slip rates, trenches give event histories, drowned forests give subsidence, and lichens on fallen boulders give dates.
Divide an offset by the age of the feature it cuts and you have a long-term slip rate. Divide the slip in one typical earthquake by that rate and you have how often such earthquakes must happen. At Wallace Creek on the San Andreas, offset and dated stream channels give 33.9 ± 2.9 mm/yr (Sieh & Jahns, 1984).
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Where one plate dives under another and the contact is locked, the overriding plate is dragged down near the trench and bows up further inland. When the lock breaks, both motions reverse in seconds. The line separating uplift from subsidence is the hinge line. Schematic profile, not to scale.
In an estuary inland of the hinge, slow interseismic uplift lets a salt marsh build up into freshwater marsh and forest. Then the earthquake drops it below the tide in an instant: trees drown standing (a ghost forest), a tsunami may lay a sheet of sand, and tidal mud buries it all. Tap any layer. A fictional estuary.
Tap a layer
Each layer has a story. Start at the bottom.
How many earthquakes does this bank record?
Strong shaking shakes boulders loose all at once, and lichens start growing on the fresh faces. Calibrate the growth rate on gravestones of known date, measure the largest lichen on each fallen boulder, and the boulders pile up into peaks — one per earthquake. A fictional valley and a fictional cemetery; the method is Bull & Brandon’s (New Zealand).
A wave-cut platform is a level line drawn by the sea. Lift the coast and the line is carried up out of reach; do that through several ice-age sea-level cycles and you get a flight of terraces. Date one and you have the uplift rate — if you remember where sea level was when it was cut.
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Each high stand of the sea cuts a platform. Uplift carries it up. It survives only if it has risen above every later high stand — otherwise the sea comes back and planes it off. Slow uplift erases most of the record; fast uplift keeps it. Sea-level high stands are rounded, simplified values; published estimates for each differ by several meters.
The lithosphere is not a raft of separate blocks; it is an elastic plate. Load it and it bends: a moat under the load, a low bulge beyond it. Whether a load is held up by the plate’s stiffness or by buoyancy depends on one thing above all — how wide it is.
A note on D. The flexural rigidity is D = E·Te³ / 12(1 − ν²) (Turcotte & Schubert, 2002). If you see it written with (1 − ν), that is a slip: with ν = 0.25 it makes D 25% too large and α about 6% too large. This page uses E = 70 GPa and ν = 0.25.
The degree of compensation C is the share of a load’s weight carried by buoyancy. C near 1: isostatic, and the simple balance from tab 01 works. C near 0: the plate’s stiffness holds the load up and there is almost no root.
The equations and the vocabulary in one place.