Sedimentary Basins

Sedimentary Basins

kacimo samy   7:38 ص

Sedimentary Basins

The sedimentary veneer on the Earth’s surface varies greatly in thickness. If you stand in central Siberia or south-central Canada, you will find yourself on igneous and metamorphic basement rocks that are over a billion years old sedimentary rocks are nowhere in sight. Yet if you stand along the southern coast of Texas, you would have to drill through over 15 km of sedimentary beds before reaching igneous and metamorphic basement. Thick accumulations of sediment form only in special regions where the surface of the Earth’s lithosphere sinks, providing space in which sediment collects. Geologists use the term subsidence to refer to the process by which the surface of the lithosphere sinks, and the term sedimentary basin for the sediment-filled depression. In what geologic settings do sedimentary basins form? An understanding of plate tectonics theory provides the answers.

Categories of Basins in the Context  of Plate Tectonics Theory 

The geologic setting of sedimentary basins.

Geologists distinguish among different kinds of sedimentary basins in the context of plate tectonics theory. Let’s consider a few examples (figure above). 

1. Rift basins: These form in continental rifts, regions where the lithosphere is stretching horizontally, and therefore thins  vertically. As the rift grows, slip on faults drops blocks of crust down, producing low areas bordered by narrow mountain ridges. These troughs fill with sediment. 
2. Passive-margin basins: These form along the edges of continents that are not plate boundaries. They are underlain by stretched lithosphere, the remnants of a rift whose evolution successfully led to the formation of a mid-ocean ridge and subsequent growth of a new ocean basin. Passive-margin  basins form because subsidence of stretched lithosphere  continues long after rifting ceases. They fill with sediment carried to the sea by rivers and with carbonate rocks formed in coastal reefs. 
3. Intracontinental basins: These develop in the interiors of continents, initially because of subsidence over a rift. They  may continue to subside in pulses even hundreds of millions of years after they formed, for reasons that are not well  understood. 
4. Foreland basins: These form on the continent side of a mountain belt because the forces produced during convergence or collision push large slices of rock up faults and onto the surface of the continent. The weight of these slices pushes down on the surface of the lithosphere, producing a wedge-shaped depression adjacent to the mountain range that fills with sediment eroded from the range. Fluvial and deltaic strata accumulate in foreland basins.

Transgression and Regression 

Sea-level changes, relative to the land surface, control the succession of sediments that we see in a sedimentary basin. At times during Earth history, sea level has risen by as much as a couple of hundred meters, creating shallow seas that submerge the interiors of continents. At other times, sea level has fallen by a couple of hundred meters, exposing the continental shelves to air. Global sea-level changes may be due to a number of factors, including climate change, which controls the amount of ice stored in polar ice caps and causes changes in the volume of ocean basins. Sea level at a location may also be due to the local uplift or sinking of the land surface.

The concept of transgression and regression, during deposition of sedimentary sequence.

When relative sea level rises, the shoreline migrates inland. We call this process transgression. When relative sea level falls, the coast migrates seaward. We call this process regression (figure above). The process of transgression and regression leads to the formation of broad blankets of sediment.

Diagenesis 

Earlier we discussed the process of lithification, by which sediment hardens into rock. Lithification is an aspect of a broader phenomenon called diagenesis. Geologists use the term diagenesis for all the physical, chemical, and biological processes that transform sediment into sedimentary rock and that alter characteristics of sedimentary rock after the rock has formed. 

In sedimentary basins, sedimentary rocks may become very deeply buried. As a result, the rocks endure higher pressures and temperatures and come in contact with warm groundwater. Diagenesis, under such conditions, can cause chemical reactions in the rock that produce new minerals and can also cause cement to dissolve or precipitate.

As temperature and pressure increase still deeper in the subsurface, the changes that take place in rocks become more profound. At sufficiently high temperature and pressure, metamorphism begins, in that a new assemblage of minerals forms, and/or mineral grains become aligned parallel to each other. The transition between diagenesis and metamorphism in sedimentary rocks is gradational and occurs between temperatures of 150C and 300C. In the next chapter, we enter the realm of true metamorphism.

Credits: Stephen Marshak (Essentials of Geology)

Ocean basins

Ocean basins

kacimo samy   2:00 ص

Altogether 71% of the area of the globe is occupied by ocean basins that have formed by sea-floor spreading and are floored by basaltic oceanic crust. The midocean ridge spreading centres are typically at 2000 to 2500 m depth in the oceans. Along them the crust is actively forming by the injection of basic magmas from below to form dykes as the molten rock solidifies and the extrusion of basaltic lava at the surface in the form of pillows. This igneous activity within the crust makes it relatively hot. As further injection occurs and new crust is formed, previously formed material gradually moves away from the spreading centre and as it does so it cools, contracts and the density increases. The older, denser oceanic crust sinks relative to the younger, hotter crust at the spreading centre and a profile of increasing water depth away from the mid-ocean ridge results down to around 4000 to 5000 m where the crust is more than a few tens of millions of years old. The ocean basins are bordered by continental margins that are important areas of terrigenous clastic and carbonate deposition. Sediment supplied to the ocean basins may be reworked from the shallow marine shelf areas, or is supplied directly from river and delta systems and bypasses the shelf. There is also intrabasinal material available in ocean basins, comprising mainly the hard part of plants and animals that live in the open oceans, and airborne dust that is blown into the oceans. These sources of sediment all contribute to oceanic deposits. The large clastic depositional systems are mainly found near the margins of the ocean basin, although large systems may extend a thousand kilometres or more out onto the basin plain, and the ocean basin plains provide the largest depositional environments on Earth. The problem with these deep-water depositional systems, however, is the difficulty of observing and measuring processes and products in the present day. The deep seas are profoundly inaccessible places. Our knowledge is largely limited to evidence from remote sensing: detailed bathymetric surveys, side-scan sonar images of the sea floor and seismic reflection surveys of the sediments. There are also extremely localised samples from boreholes, shallow cores and dredge samples. Our database of the modern ocean floors is comparable to that of the surface of the Moon and understanding the sea floor is rather like trying to interpret all processes on land from satellite images and a limited number of hand specimens of rocks collected over a large area. However, our knowledge of deep-water systems is rapidly growing, partly through technical advances, but also because hydrocarbon exploration has been gradually moving into deeper water and looking for reserves in deep-water deposits.

Morphology of ocean basins

Continental slopes typically have slope angles of between 28 and 108 and the continental rise is even less. Nevertheless, they are physiographically significant, as they contrast with the very low gradients of continental shelves and the flat ocean floor. Continental slopes extend from the shelf edge, about 200m below sea level, to the basin floor at 4000 or 5000 m depth and may be up to a hundred kilometres across in a downslope direction. Continental slopes are commonly cut by submarine canyons, which, like their counterparts on land, are steep-sided erosional features. Submarine canyons are deeply incised, sometimes into the bedrock of the shelf, and may stretch all the way back from the shelf edge to the shoreline. They act as conduits for the transfer of water and sediment from the shelf, sometimes feeding material directly from a river mouth. The presence of canyons controls the formation and position of submarine fans. The generally flat surface of the ocean floor is interrupted in places by seamounts, underwater volcanoes located over isolated hotspots. Seamounts may be wholly submarine or may build up above water as volcanic islands, such as the Hawaiian island chain in the central Pacific. As subaerial volcanoes they can be important sources of volcaniclastic sediment to ocean basins. The flanks of the volcanoes are commonly unstable and give rise to very largescale submarine slides and slumps that can involve several cubic kilometres of material. Bathymetric mapping and sonar images of the ocean floor around volcanic islands such as Hawaii in the Pacific and the Canary Islands in the Atlantic have revealed the existence of very large-scale slump features. Mass movements on this scale would generate tsunami around the edges of the ocean, inundating coastal areas. The deepest parts of the oceans are the trenches formed in regions where subduction of an oceanic plate is occurring. Trenches can be up to 10,000 m deep. Where they occur adjacent to continental margins (e.g. the Peru–Chile Trench west of South America) they are filled with sediment supplied from the continent, but mid-ocean trenches, such as the Mariana Trench in the west Pacific, are far from any source of material and are unfilled, starved of sediment.

Depositional processes in deep seas

Deposition of most clastic material in the deep seas is by mass-flow processes. The most common are debris flows and turbidity currents, and these form part of a spectrum within which there can be flows with intermediate characteristics.

Debris-flow deposits

Remobilisation of a mass of poorly sorted, sedimentrich mixture from the edge of the shelf or the top of the slope results in a debris flow, which travels down the slope and out onto the basin plain. Unlike a debris flow on land an underwater flow has the opportunity to mix with water and in doing so it becomes more dilute and this can lead to a change in the flow mechanism and a transition to a turbidity current. The top surface of a submarine debris flow deposit will typically grade up into finer deposits due to dilution of the upper part of the flow. Large debris flows of material are known from the Atlantic off northwest Africa and examples of thick, extensive debris-flow deposits are also known from the stratigraphic record. Debris-flow deposits tens of metres thick and extending for tens of kilometres are often referred to as megabeds.

Turbidites

Dilute mixtures of sediment and water moving as mass flows under gravity are the most important mechanism for moving coarse clastic material in deep marine environments. These turbidity currents carry variable amounts of mud, sand and gravel tens, hundreds and even over a thousand kilometres out onto the basin plain. The turbidites deposited can range in thickness from a few millimetres to tens of metres and are carried by flows with sediment concentrations of a few parts per thousand to 10%. Denser mixtures result in high-density turbidites that have different characteristics to the ‘Bouma Sequences’ seen in low- and medium-density turbidites. Direct observation of turbidity currents on the ocean floor is very difficult but their effects have been monitored on a small number of occasions. In November 1929 an earthquake in the Grand Banks area off the coast of Newfoundland initiated a turbidity current. The passage of the current was recorded by the severing of telegraph cables on the sea floor, which were cut at different times as the flow advanced. Interpretation of the data indicates that the turbidity current travelled at speeds of between 60 and 100 km. Also, the deposits of recent turbidity flows have been mapped out, for example, in the east Atlantic off the Canary Islands a single turbidite deposit has been shown to have a volume of 125 km cube.

High- and low-efficiency systems

A deep marine depositional system is considered to be a low-efficiency system if sandy sediment is carried only short distances (tens of kilometres) out onto the basin plain and a high-efficiency system if the transport distances for sandy material are hundreds of kilometres. High-volume flows are more efficient than small-volume flows and the efficiency is also increased by the presence of fines that tend to increase the density of the flow and hence the density contrast with the seawater. The deposits of low-efficiency systems are therefore concentrated near the edge of the basin, whereas muddier, more efficient flows carry sediment out on to the basin plain. The high-efficiency systems will tend to have an area near the basin margin called a bypass zone where sediment is not deposited, and there may be scouring of the underlying surface, with all the deposition concentrated further out in the basin.

Initiation of mass flows

Turbidity currents and mass flows require some form of trigger to start the mixture of sediment and water moving under gravity. This may be provided by an earthquake as the shaking generated by a seismic shock can temporarily liquefy sediment and cause it to move. The impact of large storm waves on shelf sediments may also act as a trigger. Accumulation of sediment on the edge of the shelf may reach the point where it becomes unstable, for example where a delta front approaches the edge of a continental shelf. High river discharge that results in increased sediment supply can result in prolonged turbidity current flow as sediment-laden water from the river mouth flows as a hyperpycnal flow across the shelf and down onto the basin plain. Such quasi-steady flows may last for much longer periods than the instantaneous triggers that result in flows lasting just a few hours. A fall in sea level exposes shelf sediments to erosion, more storm effects and sediment instability that result in increased frequency of turbidity currents.

Composition of deep marine deposits

The detrital material in deep-water deposits is highly variable and directly reflects the sediment source area. Sand, mud and gravel from a terrigenous source are most common, occurring offshore continental margins that have a high supply from fluvial sources. Material that has had a short residence time on the shelf will be similar to the composition of the river but extensive reworking by wave and tide processes can modify both the texture and the composition of the sediment before it is redeposited as a turbidite. A sandstone deposited by a turbidity current can therefore be anything from a very immature, lithic wacke to a very mature quartz arenite. Turbidites composed wholly or partly of volcaniclastic material occur in seas offshore of volcanic provinces. The deep seas near to carbonate shelves may receive large amounts of reworked shallow-marine carbonate sediment, redeposited by turbidity currents and debris flows into deeper water: recognition of the redeposition process is particularly important in these cases because the sediment will contain bioclastic material that is characteristic of shallow water environments. Because there is this broad spectrum of sandstone compositions in deep-water sediments, the use of the term ‘greywacke’ to describe the character of a deposit is best avoided: it has been used historically as a description of lithic wackes that were deposited as turbidites and the distinction between composition and process became confused as the terms turbidite and greywacke came to be used almost as synonyms. ‘Greywacke’ is not part of the Pettijohn classification of sandstones and it no longer has any widely accepted meaning in sedimentology.

Classification of sedimentary basins

Classification of sedimentary basins

kacimo samy   1:14 ص

Sedimentary basins are classified by tectonic (and, specifically, plate-tectonic) setting. That's fairly easy to do for modern basins, but it's rather difficult to do for ancient basins. List of some of the important criteria that could be used, ranging from more descriptive at the top of the list to more genetic at the bottom of the list:

• nature of fill
• geometry paleogeography
• tectonic setting 

Intracratonic Basins 

Location and tectonic setting: in anorogenic areas on cratons.

Tectonic and sedimentary processes: These basins have no apparent connection with plate tectonics. They are thought to reflect very slow thermal subsidence (for times of the order of a hundred million years) after a heating event under the continental lithosphere. But the reasons for depression below the original crustal level are not understood. Erosion during the thermal uplift seems untenable, as does lithospheric stretching. Subsidence is so slow that there seems to have been no depression of the upper surface of the lithosphere, so depositional environments are mostly the same as those in surrounding areas; the succession is just thicker. These successions are also more complete, however, there are fewer and smaller diastems (A diastem is a brief interruption in sedimentation, with little or no erosion before sedimentation resumes), so at times the basin must have remained under water while surrounding areas were emergent.

Size, shape: rounded, equidimensional, hundreds of kilometres across.

Sediment fill: shallow-water cratonal sediments (carbonates, shales, sandstones), thicker and more complete than in adjacent areas of the craton but still relatively thin, hundreds of meters.

Aulacogens 

Location and tectonic setting: extending from the margins toward the interiors of cratons.

Tectonic and sedimentary processes: Aulacogens are thought to represent the third, failed arm of a three-armed rift, two of whose arms continued to open to form an ocean basin. In modern settings, aulacogens end at the passive continental margin. An example is the Benue Trough, underlying the Congo River Basin in West Africa. The ocean eventually closes to form an orogenic belt, so in ancient settings aulacogens end at an orogenic belt; an example is the basin filled by the Pahrump Group (Proterozoic) in Nevada and California.

Size, shape: long, narrow, linear; tens of kilometres wide, many hundreds of kilometres long

Sediment fill: very thick (up to several thousand meters); coarse to fine siliciclastics, mostly coarse, minor carbonates; mostly non-marine, some marine; contemporaneous folding and faulting; the succession often passes upward, with or without major unconformity, into thinner and more widespread shallow-marine cratonal siliciclastics and carbonates.

Rift Basins

Location and tectonic setting: Within continental lithosphere on cratons.

Tectonic and sedimentary processes: Lithospheric extension on a craton, presumably by regional sub-lithospheric heating, causes major rifts. Some such rifts continue to open and eventually become ocean basins floored by oceanic rather than continental crust; the basin description here then applies to this earliest stage of the rifting. In other cases, the rifts fail to open fully into ocean basins (perhaps some adjacent and parallel rift becomes the master rift), so they remain floored by thinned continental crust rather than new oceanic crust. A modern example: the East African rift valleys. An ancient example: the Triassic–Jurassic Connecticut and Newark basins in eastern North America. Sediment supply from the adjacent highlands of the uplifted fault blocks is usually abundant, although in the East African rifts the land slope is away from the rim of the highlands, and surprisingly little sediment reaches the rift basins.

Size, shape: long, narrow, linear; tens of kilometres wide, up to a few thousand kilometres long

Sediment fill: Coarse to fine siliciclastics, usually non-marine; often lacustrine sediments; inter-bedded basalts.

Oceanic Rift Basins 

Location and tectonic setting: In a narrow and newly opening ocean.

Processes: This category of basins is transitional between intracontinental rift basins and passive-margin basins. Basins have opened wide enough to begin to be floored with oceanic crust but are still so narrow that the environment is either still non-marine or, if marine, has restricted circulation. Modern examples are the Red Sea and the Gulf of Aden. In the ancient, the sediment fill of such a basin is likely to underlie passive-margin sediments deposited later in the history of ocean opening.

Size, shape: long, narrow; straight or piecewise straight; tens to a few hundreds of kilometres wide, up to a few thousand kilometres long.

Sediment fill: mafic, volcanics and coarse to fine non-marine siliciclastics, as in intracratonic rift basins described above, passing upward and laterally into evaporites, lacustrine deposits, and fine marine sediments, often metal-rich from hydrothermal activity at the spreading ridge.

Passive Margin Basins 

Location and tectonic setting: Along passive continental margins, approximately over the transition from continental to oceanic crust formed by rifting and opening of a full-scale ocean basin

Processes: As an ocean basin opens by spreading, the zone of heating and extensional thinning of continental crust on either side of the ocean basin subsides slowly by cooling. Sediments, either siliciclastics derived from land or carbonates generated in place, cover this subsiding transition from continental crust to oceanic crust with a wedge of sediment to build what we see today as the continental shelf and slope. In the context of the ancient, this represents the miogeocline. The subsidence is accentuated by loading of the deposited sediments, resulting in a prominent down bowing of the continental margin. Deposition itself therefore does not take place in a basinal geometry, but the base of the deposit is distinctively concave upward.

Size, shape: Straight to piecewise straight, often with considerable irregularity in detail; a few hundreds of kilometres wide, thousands of kilometres long.

Sediment fill: Overlying and overlapping the earlier deposits laid down earlier during rifting and initial opening are extensive shallow-marine siliciclastics and carbonates of the continental shelf, thickening seaward. These sediments pass gradually or abruptly into deeper marine fine sediments of the continental slope and rise, often grading or inter-fingering seaward into deep marine coarse and fine siliciclastics or re-sedimented carbonates in the form of turbidites building submarine fans at the base of the slope and filling the deepest parts of the ocean basin to form abyssal plains.

 

Trenches 

Location and tectonic setting: In the abyssal ocean, at the line of initial down bending of the subducted oceanic-crust plate in a subduction zone.

Tectonic and sedimentary processes: (1) Open-ocean pelagic sediments (mainly abyssal brown clay and organic oozes) are conveyor-belted to the trench, and underlie the sediments, thin to thick, deposited in the trench itself. (2) What happens to the sediments delivered to or deposited in the trench? While still within the trench they are little deformed, but they don't stay that way long. They are either scraped off the descending plate to form an accretionary wedge, whose structure ranges from chaotically mixed material in a subduction mélange to a fairly regular imbricated succession of under-thrusted sheets dipping toward the arc (the thrust sheets themselves get younger downward, but within a given plate the sediments get younger upward), or they are dragged down the subduction zone. 

Size, shape: long and narrow (tens of kilometres wide, thousands of kilometres long), arcuate, with convex side toward the oncoming subducted plate.

Sediment fill: varies from thin (hundreds of meters) pelagic sediments (fine abyssal muds, volcanic ash) to thick (thousands of meters) arc-derived coarse siliciclastics and volcaniclastics, as local fans built perpendicular to the trench axis or oblong fan like bodies built parallel to trench axis.

 

Trench-Slope Basins 

Location and tectonic setting: On the inner (arcward) wall of subduction-zone trenches.

Tectonic and sedimentary processes: The basins are formed as low areas, with closed contours, between adjacent thrust sheets in the growing accretionary wedge. Near-surface folding of sediment in the accretionary wedge may also be a factor in the development of the basins. These basins intercept some of the sediment carried as turbidity currents from upraised older parts of the accretionary complex, or from the more distant arc itself.

Size, shape: Small (no larger than kilometres wide, tens of kilometres long, often smaller); linear, and elongated parallel to the trench.

Sediment fill: deep-marine silts and muds sedimented directly into the basins or slumped into the basins from higher on the slope; also coarser siliciclastics supplied from farther upslope by turbidity currents.

 

Fore-Arc Basins 

Location and tectonic setting: In subduction zones; between the upraised subduction complex just inboard of the trench and the volcanic arc (in the case of ocean-ocean subduction) or the overriding continent (in the case of ocean–continent subduction).

Processes: As subduction proceeds, a relatively low area, usually below sea level, is formed between the relatively high outer arc upraised by subduction and the inner volcanic arc built by subduction magmatism. Ancient examples of such fore-arc basins are likely to be tectonically isolated from the originally adjacent areas. After an arc-continent collision, another variant of fore-arc basin can be formed between the outer arc and the overriding continent. These basins are likely to be filled mainly from the high land of the tectonically active continent. Later continent-continent collision would make direct reconstruction of their tectonic setting difficult.

Size, shape: tens of kilometres to over one hundred kilometres wide, up to thousands of kilometres long; commonly arcuate.

Sediment fill: non-marine, siliciclastic, fluvial to deltaic deposits at the arcward margin pass seaward into deep marine siliciclastics, mainly sediment-gravity flow deposits, all inter-bedded with arc derived volcanics flows and pyroclastics. Section thickness can be many thousands of meters.

Foreland Basins 

Location, tectonic setting, processes: There are two kinds of foreland basins: retro-arc foreland basins, which are formed on stable continental crust by loading by thrust sheets moving toward the continental interior as a result of compression and crustal shortening in an ocean-continent subduction zone, and peripheral foreland basins, formed after continent-continent collisions by loading of the continental crust of the subducted plate by development of thrust sheets in the continental crust of the subducted plate directed back away from the subduction zone. Both kinds tend to be asymmetrical, with their deepest parts nearest the emplaced thrust sheets. They tend to migrate away from the arc or suture zone with time. They are filled by sediments derived from the mountainous terrain associated with the compression and thrusting.

Size, shape: tens to a few hundreds of kilometres wide, hundreds to thousands of kilometres long; often with varying development along their length; commonly arcuate or piecewise arcuate, reflecting the geometry of subduction.

Sediment fill: Coarse, fluvial, siliciclastics, mainly as alluvial fans, thinning and fining away from the arc or suture, often passing into shallow-marine sandstone-shale successions if sea level is high enough to flood the basin. Thickness are up to many thousands of meters. The classic molasse facies, thick non-marine conglomerates, is deposited in foreland basins.

 

Remnant Basins 

Location and tectonic setting: within suture zones formed by continent-continent collision.

Processes: Continental margins and subduction zones are (for different reasons, connected with geometry of rifting and geometry of subduction) commonly irregular in plan rather than straight, so when continent-continent collision eventually comes to pass, certain salient of continental crust encounter the subduction zone before re-entrants. With further subduction and suturing, this creates isolated basins still floored by residual oceanic crust, which receive abundant sediment from adjacent strongly uplifted crust.

Size, shape: many tens to hundreds of kilometres across; irregular in shape.

Sediment fill: very thick and highly varied, with strong lateral facies changes; usually fluvial at the margins, commonly passing into deep-marine sediment gravity-flow deposits; sometimes the basin becomes sealed off from the ocean, so that facies include lacustrine sediments.

 

Pull-Apart Basins 

Location and tectonic setting: Locally along major strike-slip or transform faults, either in continental crust or in oceanic crust.

Processes: If a strike-slip fault is stepped or curved rather than straight, movement along it tends to produce tension, where the sense of the curvature and movement are such that the walls of the fault are pulled apart from one another (this kind of regime is described as transtensile), or compression, where the sense of the curvature and movement are such that the walls are pushed against one another (this kind of regime is described as transpressive). In the tensional segments, gaps or basins are produced which are filled with sediment from adjacent high crust.

Size, shape: There is a strong tendency for pull-apart basins to be rhomboidal. They range from approximately equidimensional early in their history to elongated later. Widths are kilometres to a few tens of kilometres, and lengths are kilometres to many tens of kilometres. Some basins are even smaller than this.

Sediment fill: The continental crust basins, which are the most significant sedimentologically, are filled by thick non-marine to marine coarse to fine clastics, often as alluvial fans passing into lake deposits or into deposits of restricted marine environments. In some cases thick marine turbidites fill the distal parts of the basin. There is usually sharp variation in facies laterally, and the thickness of the lithologic units may be not much greater than the lateral extent, or even less. Deposition is concurrent with elongation of the basin, so be wary of total section thickness computed by bed-by-bed measurements of the section.

How are sedimentary basins formed?

How are sedimentary basins formed?

kacimo samy   4:00 ص

Sedimentary basin: A low area in the Earth’s crust, of tectonic origin, in which sediments accumulate. Sedimentary basins range in size from as small as hundreds of meters to large parts of ocean basins. The essential element of the concept is tectonic creation of relief, to provide both a source of sediment and a relatively low place for the deposition of that sediment. Tectonics is the most important control on sedimentation; climate is a rather distant second. The important effects of tectonics on sedimentation, direct or indirect, include the following:

•  nature of sediment
• rate of sediment supply 
• rate of deposition 
• depositional environment 
• nature of source rocks 
• nature of vertical succession

The only basins that are preserved in their entirety are those that lie entirely in the subsurface! Basins exposed at the surface are undergoing destruction and loss of record by erosion. So there’s an ironic trade-off between having more complete preservation in the subsurface but less satisfactory observations. How do you gather data on sedimentary basins?.

Master cross sections: With the present land surface as the most natural datum, construct several detailed physical cross sections through the basin to show its geometry and sediment fill.

Stratigraphic sections: Construct a graph, with time along the vertical axis, showing the time correlations of all the major rock units along some generalized traverse across the basin. Such a section includes hiatuses, during which there was non deposition or erosion.

Isopach maps: With some distinctive stratigraphic horizon near the top of the section as datum, draw a contour map showing isopachs (isopachs are loci of equal total sediment thickness) in the basin.

Lithofacies maps: For one or a series of times, draw a map showing distribution of sediment types being deposited at that time.

Ratio maps: Compute things like sand/shale ratio, integrated over the entire section or restricted to some time interval, and plot a contour map of the values.

Paleocurrent maps: For one or a series of times, draw a map showing the direction of paleo-currents in the basin at that time.

Grain-size maps: For the entire basin fill, averaged vertically, or for some stratigraphic interval or time interval, draw a map that shows the areal distribution of sediment grain size. This is especially useful for conglomeratic basins.

In one sense, the origin of sedimentary basins boils down to the question of how relief on the Earth is created. Basically, there are only a few ways, described in the following sections.

Local

On a small scale, hundreds to thousands of meters laterally, fault movements can create relief of hundreds to thousands of meters, resulting in small but often deep basins (some of these are called intermontane basins). Along strike-slip faults. can produce small pull-apart basins; more on them later. Relief of this kind is on such a small scale that it tends not to be isostatically compensated. It’s like setting a block of granite out on your driveway; the flexural rigidity of your driveway is great enough compared with the imposed load that the granite block is prevented from finding its buoyant equilibrium position.

Regional

Basin relief can be created mechanically on a regional scale in two very important ways: thermally or flexurally, or by a combination of those two effects). Each of these is discussed briefly below. Keep in mind that basins can also be made just by making mountain ranges, on land or in the ocean, by volcanism. 

Thermal

If the lithosphere is heated from below, it expands slightly and thus becomes less dense. This less dense lithosphere adjusts isostatically to float higher in the asthenosphere, producing what we see at the Earth’s surface as crustal uplift. If the lithosphere cools back to its original temperature, there’s isostatic subsidence back to the original level.

 But suppose that some erosion took place while the crust was elevated. The crust is thinned where the erosion took place (and thickened somewhere else, where there was deposition; that might be far away, at the mouth of some long river system), so when the crust cools again it subsides to a position lower than where it started, thus creating a basin available for filling by sediments.

 But the magnitude of crustal lowering by this mechanism is less than is often observed in basins thought to be created thermally. It has therefore been proposed, and widely accepted, that in many cases extensional thinning of the lithosphere accompanies the heating. Then, upon re-cooling, the elevation of the top of the lithosphere is less than before the heating and extension. This kind of subsidence has been invoked to explain many sedimentary basins.        

                   

Flexural

Another important way to make basins is to park a large load on some area of the lithosphere. The new load causes that lithosphere to subside by isostatic adjustment. But because the lithosphere has considerable flexural rigidity, adjacent lithosphere is bowed down also. The region between the high-standing load and the lithosphere in the far field (in the parlance of geophysics, that just means far away!) is thus depressed to form a basin. This model has been very successful in accounting for the features of foreland basins, which are formed ahead of large thrust sheets that move out from orogenic areas onto previously undeformed cratonal lithosphere.