Plate tectonics and sedimentary basins

The lithosphere is the crust together with the cold upper part of the mantle (more detail). It is not one continuous shell. It is a mosaic of lithospheric plates. Inside, the plates hardly deform, like rigid sheets. All the main events — earthquakes, eruptions, the growth of mountains — happen at their boundaries. A plate can carry a continent and an ocean floor together. The Eurasian Plate includes almost all of Europe and most of Asia, with the adjoining parts of the oceans. Kazakhstan lies in its interior, far from the present-day boundaries.

There are seven or eight large plates: the Pacific, North American, South American, Eurasian, African, Antarctic and Indo-Australian (often divided into the Indian and Australian plates). There are about a dozen medium-sized ones — for example, the Arabian, Philippine, Caribbean, Nazca and Cocos plates — and several dozen small microplates. The number depends on which moving blocks are counted as separate plates.

Why do plates move? The lithosphere is the upper, cold layer of the mantle's convection, the slow circulation by which the mantle loses its heat (more detail). As oceanic lithosphere cools, it becomes heavier than the mantle beneath it. In subduction zones (where one plate dives beneath another) it sinks into the mantle. By current understanding, the main force is the pull of the heavy, cold plate sinking into the mantle (slab pull). Another force adds to it: the push away from the raised, hot mid-ocean ridges (ridge push). That is why plates with long subduction zones, for example the Pacific Plate, move fastest.

We can measure the speed in two independent ways. The first uses the magnetic anomalies of the sea floor. As basalt cools at the ridge axis, it becomes magnetised in the direction of the Earth's magnetic field. From time to time the field reverses its polarity (north and south swap). So the sea floor on both sides of a ridge is made of stripes of normal and reversed magnetisation, parallel to the axis. We know the age of each polarity reversal from the reversal time scale. The distance from the axis to a stripe boundary, divided by its age, gives the mean speed over millions of years. The second way is satellite geodesy (GPS, the Global Positioning System, and other systems). Stations on different plates move closer together or farther apart by millimetres year after year. Where both measurements exist, they usually agree. Plates move almost uniformly over times from years to millions of years.

At a divergent boundary the plates move apart. Under the axis of a mid-ocean ridge, hot mantle material rises. The pressure on it falls, and part of it melts. The melt solidifies and adds to the oceanic crust. This is sea-floor spreading. When a continent splits, a continental rift forms first: a long depression between normal faults (faults along which the upper block slides down), like the East African Rift. If the stretching continues, the rift turns into a narrow young ocean like the Red Sea, and later into a wide ocean.

At a convergent boundary the plates move together. If at least one of them carries ocean floor, the denser plate sinks beneath the other. This is subduction. A deep-sea trench forms at the surface above it. Deeper down, water is released from the sinking plate and melts the mantle above it, and a chain of volcanoes forms. It is an island arc when ocean goes under ocean (the Mariana Islands), or a volcanic belt on the edge of a continent (the Andes). Earthquakes can be traced along the sinking plate to a depth of about 700 km. Continental crust is too light to sink deeply. So when the ocean between two continents closes, they collide. This is collision: the crust is crumpled and thickened, and high mountains grow (the Himalayas, the Alps). The Urals are the ancient seam (suture) of such a collision: in the Late Carboniferous and Permian the East European and Kazakhstan continents came together here.

At a transform boundary the plates slide past each other along a vertical strike-slip fault (a fault where the two sides move sideways). Crust is neither created nor destroyed. Short transform faults cut the mid-ocean ridges into segments. Large ones come onshore: the San Andreas Fault in California and the North Anatolian Fault in Turkey. Where a strike-slip fault bends, its sides in some places move apart and form depressions, and in other places are squeezed together and push up uplifts (see the section on strike-slip basins).

The type of boundary also decides the stress field in the lithosphere. Where the lithosphere is stretched, the smallest stress is horizontal and the faults are normal faults. Where it is compressed, the vertical stress becomes the smallest, and reverse faults and thrusts appear (faults where the upper block is pushed up and over the lower one). At strike-slip boundaries both the greatest and the smallest stresses are horizontal (tectonic stress regimes). The rigid lithosphere carries stresses from the boundaries over thousands of kilometres. So even deep inside a plate they set the directions of fractures and the stability of wellbores (Rock mechanics for hydraulic fracturing).

Oceans open and close. About 200 million years ago almost all the continents were joined in the supercontinent Pangaea. Its break-up opened the Atlantic. The oceans that existed before closed in collisions and left sutures like the Urals. The sequence “rift — young ocean — wide ocean with passive margins — subduction — collision” is called the Wilson cycle, after the Canadian geophysicist John Tuzo Wilson. Each stage makes its own basins, and in one area their sequences often lie one on top of another.

A sedimentary basin is an area hundreds of kilometres across where the surface of the lithosphere sank for a long time and collected a sequence of sediments from a few kilometres up to 20 km thick. For sediments to pile up, they need room: a depression below the level to which water can build them up. This room is called accommodation. Subsidence creates it, and so does a rise in sea level. Without subsidence the depression would fill up quickly, and after that the sediments would be carried past it into other, deeper depressions.

Subsidence caused by the lithosphere itself is called tectonic subsidence. It has three main causes:

1. Stretching. The light crust gets thinner, and denser mantle takes its place. The rock column becomes heavier and sinks. This is the rapid subsidence of rifts while the faults are active.
2. Cooling. The stretched lithosphere is hot. As it cools, it shrinks and becomes denser, and the surface keeps sinking slowly for tens of millions of years after the stretching. Passive margins and basins above ancient rifts sink in this way.
3. Flexure (bending) under a load. The thrusts of growing mountains, volcanic islands and thick sequences of sediments bend the elastic lithosphere, as a weight bends a plank floating on water. Foreland basins form this way.

Other causes are also discussed: currents in the mantle that drag the surface down or up over large areas, and the rocks of the lower crust becoming denser. The weight of the sediments does not create a basin, but it deepens it. Sediments are heavier than the water they push out, and with local isostasy a depression filled with sediments ends up 2–3 times deeper than one filled with water (isostasy and the load of sediments).

Stretching acts on the lithosphere in two ways. First, it thins the crust by a factor of $\beta$. This factor is the stretching factor: at $\beta = 2$ a crust 35 km thick becomes 17.5 km thick. Dense mantle takes the place of the light crust, so the column becomes heavier and the surface sinks. Second, the mantle part of the lithosphere thins together with the crust. Hot asthenosphere, which is slightly less dense, takes its place, and this raises the surface. If the crust is thick enough, the first effect wins and a depression forms. Normal faults break the brittle upper crust into tilted blocks, and half-grabens (depressions with a fault on one side) form between them. This is the syn-rift stage, the stage while the faults are active.

When the stretching stops, excess heat remains. The isotherms (lines of equal temperature) in the stretched lithosphere lie $\beta$ times closer together, and the heat flow is higher. The heat escapes to the surface. The lithosphere cools, thickens again and shrinks, so the surface sinks slowly, now without any help from faults. This is the post-rift (thermal) stage. Its sediments lie undisturbed and cover both the rift and its shoulders. So in cross-section a rift basin often looks like a narrow wedge beneath a broad bowl. Stretching can also be asymmetric, along a gently dipping fault that cuts the whole lithosphere. Then the subsidence is shared unevenly between the two sides of the rift.

The lithospheric stretching model describes both stages in numbers. The British geophysicist Dan McKenzie proposed it in 1978. In the model, the lithosphere is stretched instantly and evenly by a factor of $\beta$ and then left to cool. The model calculates the subsidence from isostasy for a depression filled with water. It then allows for filling with sediments by the same multiplier as in the lesson on the structure of the Earth.

A passive margin is the edge of a continent after a rift has turned into an ocean. Its lithosphere is stretched most strongly next to the ocean, so it also sinks most there. Sediments are carried onto it from the continent, and over tens of millions of years a wedge of sediment 10 km or more thick grows along the margin. It contains deltas, submarine fans (fan-shaped heaps of sediment on the sea floor) and carbonate platforms. If the young ocean was narrow and poorly connected to the open sea, salt collects in it at an early stage. Later the salt flows under the load of the sediments and forms domes and walls that create traps. Margins that split apart from one rift, for example the coasts of Brazil and West Africa, have a similar structure. So discoveries on one side of the ocean suggest where to look on the other.

When plates collide, sheets of rock are thrust over one another, and the growing mountain belt presses on the edge of the lower plate. The lithosphere behaves like an elastic plate lying on the soft asthenosphere. Under a load it bends, and far from the load the bending dies away. In front of the thrusts a foreland basin forms (in Russian-language literature it is also called a marginal trough). It is an asymmetric depression: deepest next to the mountains and shallower towards the interior of the plate. Farther out, the plate bends the other way and arches upwards, forming a gentle forebulge. Beyond the forebulge there is a weak back-bulge depression.

The flexural rigidity of the plate (how stiff it is against bending) sets the width of the basin. We express it through the effective elastic thickness $T_e$: the thickness of a uniform elastic plate that would bend in the same way. This is not the thickness of the lithosphere. It is a measure of its strength. For young, hot lithosphere $T_e$ is from a few kilometres to a few tens of kilometres. For ancient, cold lithosphere it is 50–100 km by many estimates, although different methods give very different values for ancient continents.

Foreland basins are among the richest petroleum provinces in the world. Their early, deep-water sequences and the deposits of the former passive margin below them often contain source rocks. The forebulge and the basin itself provide reservoirs: sandstones washed down from the mountains, and leached carbonates. Thrusts crumple the layers into anticlines (arch-shaped folds), which are traps. An example is the Alberta Basin in Canada, in front of the Rocky Mountains. In front of the Urals, the Pre-Ural foredeep formed at the end of the Palaeozoic. Fragments washed down from the growing mountains are also found on the eastern edge of the Pre-Caspian Basin.

Intracratonic basins (in Russian-language literature also called syneclises) are broad, gentle bowls in the interior of plates, on ancient crust, far from the plate boundaries. They sink very slowly, by tens of metres per million years, but they keep sinking for hundreds of millions of years. They collect sequences of shallow-marine rocks — carbonates, shales, evaporites — that keep the same character over huge areas. Scientists do not agree on one cause of their long subsidence. The suggestions include cooling of the lithosphere after heating or weak stretching long ago, the rocks of the lower crust becoming denser, and slow flow of the mantle beneath the plate.

Strike-slip basins form where a large strike-slip fault bends, or steps from one line to another. If the bend lets the movement pull the sides of the fault apart, a diamond-shaped (rhomb-shaped) depression appears between its segments. This is a pull-apart basin. On a right-lateral fault, where the opposite side moves to the right, the pull-apart forms at a right step. With the opposite bend, the sides are compressed and push up an uplift. Such basins are small, from a few kilometres to 150–200 km long, but they sink quickly, are often hot, and are filled with coarse debris from the steep margins and with lake or marine shales in the centre.

The setting can change, and stretching can give way to compression. Old normal faults then start to work as reverse faults. The sediments that filled the rift are squeezed upwards and turn into anticlines. This is basin inversion. Inversion folds often become traps, but inversion can also destroy an existing accumulation if the faults break through the seal.

A basin is not just a pit filled with sediments. It is a “factory” where all the elements of the petroleum system (source, reservoir, seal and trap) come together (Petroleum system and the play concept).

• Source rocks. Organic matter is preserved where the water at the bottom is stagnant and oxygen-free: in deep rift lakes, in isolated marine depressions, in the deep-water centres of basins (Source rocks, maturation and migration). Subsidence creates such depressions.
• Burial and heating. For a source rock to yield oil, it has to be buried to a depth where the temperature reaches the oil window. The subsidence and the heat flow of the basin set this depth (example below).
• Reservoirs and seals. Sands of rivers, deltas and submarine fans, and carbonate platforms and reefs on the margins, are reservoirs. Shales and salts that built up during subsidence are seals (Reservoir and seal).
• Traps. The basin's own processes create them: tilted blocks and horsts (raised fault blocks) of rifts, salt domes of passive margins, anticlines of thrusts and of inversion, reefs on platform edges, and pinch-outs of layers on highs (where a layer thins out to nothing) (Traps and timing). Timing matters too: a trap must form before the oil passes through the place.

The type of basin also sets its temperature conditions. Young rifts and many strike-slip basins are hot: the lithosphere beneath them is thin, the heat flow is high, and source rocks mature (turn their organic matter into oil) at shallow depth. Foreland basins and ancient platform basins lie on thick, cold lithosphere, so a source rock has to be buried much deeper to mature.

Kazakhstan lies inside the Eurasian Plate, but its basins keep the history of ancient oceans, collisions and rifts. Oil and gas are produced in several of them. Below is what is well enough established about the main basins. Geologists keep refining the details of their history with new wells and seismic data.