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49-797: The Black Warrior basin was a foreland basin during the Ouachita Orogeny during the Pennsylvanian and Permian Periods. The basin also received sediments from the Appalachian orogeny during the Pennsylvanian. The western margin of the basin lies beneath the sediments of the Mississippi embayment where it is contiguous with the Arkoma Basin of northern Arkansas and northeastern Oklahoma . The region existed as

98-407: A subaerial wedge is flanked with terrestrial or shallow marine foreland basins". The temperature underneath the orogen is much higher and weakens the lithosphere. Thus, the thrust belt is mobile and the foreland basin system becomes deformed over time. Syntectonic unconformities demonstrate simultaneous subsidence and tectonic activity. Foreland basins are filled with sediments which erode from

147-424: A certain proportion of its mass below the surface of the water. If snow falls to the top of the iceberg, the iceberg will sink lower in the water. If a layer of ice melts off the top of the iceberg, the remaining iceberg will rise. Similarly, Earth's lithosphere "floats" in the asthenosphere. When continents collide, the continental crust may thicken at their edges in the collision. It is also very common for one of

196-425: A change in crust loading) provide information on the viscosity of the upper mantle. The basis of the model is Pascal's law , and particularly its consequence that, within a fluid in static equilibrium, the hydrostatic pressure is the same on every point at the same elevation (surface of hydrostatic compensation): h 1 ⋅ρ 1 = h 2 ⋅ρ 2 = h 3 ⋅ρ 3 = ... h n ⋅ρ n For the simplified picture shown,

245-535: A characteristic wave number As the rigid layer becomes weaker, κ {\displaystyle \kappa } approaches infinity, and the behavior approaches the pure hydrostatic balance of the Airy-Heiskanen hypothesis. The depth of compensation (also known as the compensation level , compensation depth , or level of compensation ) is the depth below which the pressure is identical across any horizontal surface. In stable regions, it lies in

294-533: A local hydrostatic balance. A third hypothesis, lithospheric flexure , takes into account the rigidity of the Earth's outer shell, the lithosphere . Lithospheric flexure was first invoked in the late 19th century to explain the shorelines uplifted in Scandinavia following the melting of continental glaciers at the end of the last glaciation . It was likewise used by American geologist G. K. Gilbert to explain

343-537: A quiescent continental shelf environment through the early Paleozoic from the Cambrian through the Mississippian with the deposition of shelf sandstones, shale, limestone, dolomite and chert. 33°19′N 87°13′W  /  33.317°N 87.217°W  / 33.317; -87.217 This Alabama state location article is a stub . You can help Misplaced Pages by expanding it . This article about

392-607: A regime's tectonic origin and development as well as the lithospheric mechanics. Migrating fluids originate from the sediments of the foreland basin and migrate in response to deformation. As a result, brine can migrate over great distances. Evidence of long-range migration includes: 1) correlation of petroleum to distant source rocks , 2) ore bodies deposited from metal-bearing brines, 3) anomalous thermal histories for shallow sediments, 4) regional potassium metasomatism and 5) epigenetic dolomite cements in ore bodies and deep aquifers. Fluids carrying heat, minerals, and petroleum, have

441-490: A region, the land may rise to compensate. Therefore, as a mountain range is eroded, the (reduced) range rebounds upwards (to a certain extent) to be eroded further. Some of the rock strata now visible at the ground surface may have spent much of their history at great depths below the surface buried under other strata, to be eventually exposed as those other strata eroded away and the lower layers rebounded upwards. An analogy may be made with an iceberg , which always floats with

490-455: A specific United States geological feature is a stub . You can help Misplaced Pages by expanding it . Foreland basin A foreland basin is a structural basin that develops adjacent and parallel to a mountain belt . Foreland basins form because the immense mass created by crustal thickening associated with the evolution of a mountain belt causes the lithosphere to bend, by a process known as lithospheric flexure . The width and depth of

539-454: A vast impact on the tectonic regime within the foreland basin. Before deformation, sediment layers are porous and full of fluids, such as water and hydrated minerals. Once these sediments are buried and compacted, the pores become smaller and some of the fluids, about ⁠ 1 / 3 ⁠ , leave the pores. This fluid has to go somewhere. Within the foreland basin, these fluids potentially can heat and mineralize materials, as well as mix with

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588-528: Is defined as the Bouger anomaly minus the gravity anomaly due to the subsurface compensation, and is a measure of the local departure from isostatic equilibrium. At the center of a level plateau, it is approximately equal to the free air anomaly . Models such as deep dynamic isostasy (DDI) include such viscous forces and are applicable to a dynamic mantle and lithosphere. Measurements of the rate of isostatic rebound (the return to isostatic equilibrium following

637-685: Is found closer to the orogen and oil is found further away. Isostasy Isostasy (Greek ísos 'equal', stásis 'standstill') or isostatic equilibrium is the state of gravitational equilibrium between Earth 's crust (or lithosphere ) and mantle such that the crust "floats" at an elevation that depends on its thickness and density. This concept is invoked to explain how different topographic heights can exist at Earth's surface. Although originally defined in terms of continental crust and mantle, it has subsequently been interpreted in terms of lithosphere and asthenosphere , particularly with respect to oceanic island volcanoes , such as

686-408: Is more geologically accurate within a specific region. Seismicity determines where active zones of seismic activity occur as well as measure the total fault displacements and the timing of the onset of deformation. Foreland basins form because as the mountain belt grows, it exerts a significant mass on the Earth's crust, which causes it to bend, or flex, downwards. This occurs so that the weight of

735-472: Is most rapid near the moving thrust sheet. Sediment transport within the foredeep is generally parallel to the strike of the thrust fault and basin axis. The motion of the adjacent plates of the foreland basin can be determined by studying the active deformation zone with which it is connected. Today GPS measurements provide the rate at which one plate is moving relative to another. It is also important to consider that present day kinematics are unlikely to be

784-400: Is provided by loading and downflexure to form foreland basins, in contrast to rift basins, where accommodation space is generated by lithospheric extension. Foreland basins can be divided into two categories: DeCelles & Giles (1996) provide a thorough definition of the foreland basin system. Foreland basin systems comprise three characteristic properties: The wedge-top sits on top of

833-445: Is the acceleration due to gravity, and P ( x ) {\displaystyle P(x)} is the load on the ocean crust. The parameter D is the flexural rigidity , defined as where E is Young's modulus , σ {\displaystyle \sigma } is Poisson's ratio , and T c {\displaystyle T_{c}} is the thickness of the lithosphere. Solutions to this equation have

882-476: The Baltic Sea and Hudson Bay . As the ice retreats, the load on the lithosphere and asthenosphere is reduced and they rebound back towards their equilibrium levels. In this way, it is possible to find former sea cliffs and associated wave-cut platforms hundreds of metres above present-day sea level . The rebound movements are so slow that the uplift caused by the ending of the last glacial period

931-534: The Hawaiian Islands . Although Earth is a dynamic system that responds to loads in many different ways, isostasy describes the important limiting case in which crust and mantle are in static equilibrium . Certain areas (such as the Himalayas and other convergent margins) are not in isostatic equilibrium and are not well described by isostatic models. The general term isostasy was coined in 1882 by

980-520: The American geologist Clarence Dutton . In the 17th and 18th centuries, French geodesists (for example, Jean Picard ) attempted to determine the shape of the Earth (the geoid ) by measuring the length of a degree of latitude at different latitudes ( arc measurement ). A party working in Ecuador was aware that its plumb lines , used to determine the vertical direction, would be deflected by

1029-560: The Pratt hypothesis as overlying regions of unusually low density in the upper mantle. This reflects thermal expansion from the higher temperatures present under the ridges. In the Basin and Range Province of western North America, the isostatic anomaly is small except near the Pacific coast, indicating that the region is generally near isostatic equilibrium. However, the depth to the base of

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1078-404: The adjacent mountain belt. In the early stages, the foreland basin is said to be underfilled . During this stage, deep water and commonly marine sediments, known as flysch , are deposited. Eventually, the basin becomes completely filled. At this point, the basin enters the overfilled stage and deposition of terrestrial clastic sediments occurs. These are known as molasse . Sediment fill within

1127-403: The balancing of lithospheric columns gives: where ρ m {\displaystyle \rho _{m}} is the density of the mantle (ca. 3,300 kg m ), ρ c {\displaystyle \rho _{c}} is the density of the crust (ca. 2,750 kg m ) and ρ w {\displaystyle \rho _{w}} is the density of

1176-418: The crust does not strongly correlate with the height of the terrain. This provides evidence (via the Pratt hypothesis) that the upper mantle in this region is inhomogeneous, with significant lateral variations in density. The formation of ice sheets can cause Earth's surface to sink. Conversely, isostatic post-glacial rebound is observed in areas once covered by ice sheets that have now melted, such as around

1225-448: The deep crust, but in active regions, it may lie below the base of the lithosphere. In the Pratt model, it is the depth below which all rock has the same density; above this depth, density is lower where topographic elevation is greater. When large amounts of sediment are deposited on a particular region, the immense weight of the new sediment may cause the crust below to sink. Similarly, when large amounts of material are eroded away from

1274-406: The deformation of the rigid crust. These elastic forces can transmit buoyant forces across a large region of deformation to a more concentrated load. Perfect isostatic equilibrium is possible only if mantle material is in rest. However, thermal convection is present in the mantle. This introduces viscous forces that are not accounted for the static theory of isostacy. The isostatic anomaly or IA

1323-411: The depth of the mountain belt roots (b 1 ) is calculated as follows: where ρ m {\displaystyle \rho _{m}} is the density of the mantle (ca. 3,300 kg m ) and ρ c {\displaystyle \rho _{c}} is the density of the crust (ca. 2,750 kg m ). Thus, generally: In the case of negative topography (a marine basin),

1372-403: The flexural rigidity of the lithosphere approaches zero. For example, the vertical displacement z of a region of ocean crust would be described by the differential equation where ρ m {\displaystyle \rho _{m}} and ρ w {\displaystyle \rho _{w}} are the densities of the aesthenosphere and ocean water, g

1421-421: The foredeep acts as an additional load on the continental lithosphere. Although the degree to which the lithosphere relaxes over time is still controversial, most workers accept an elastic or visco-elastic rheology to describe the lithospheric deformation of the foreland basin. Allen & Allen (2005) describe a moving load system, one in which the deflection moves as a wave through the foreland plate before

1470-430: The foreland basin is determined by the flexural rigidity of the underlying lithosphere, and the characteristics of the mountain belt. The foreland basin receives sediment that is eroded off the adjacent mountain belt, filling with thick sedimentary successions that thin away from the mountain belt. Foreland basins represent an endmember basin type, the other being rift basins . Space for sediments (accommodation space)

1519-500: The gravitational attraction of the nearby Andes Mountains . However, the deflection was less than expected, which was attributed to the mountains having low-density roots that compensated for the mass of the mountains. In other words, the low-density mountain roots provided the buoyancy to support the weight of the mountains above the surrounding terrain. Similar observations in the 19th century by British surveyors in India showed that this

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1568-679: The lithosphere beneath the mountain range becomes ductile almost entirely, except a thin (about 6 km in the center) brittle layer near the surface and perhaps a thin brittle layer in the uppermost mantle." This lithospheric weakening underneath the orogenic belt may in part cause the regional lithospheric flexure behavior. Foreland basins are considered to be hypothermal basins (cooler than normal), with low geothermal gradient and heat flow . Heat flow values average between 1 and 2 HFU (40–90 mWm . Rapid subsidence may be responsible for these low values. Over time sedimentary layers become buried and lose porosity. This can be due to sediment compaction or

1617-424: The load system. The deflection shape is commonly described as an asymmetrical low close to the load along the foreland and a broader uplifted deflection along the forebulge. The transport rate or flux of erosion, as well as sedimentation, is a function of topographic relief. For the loading model, the lithosphere is initially stiff, with the basin broad and shallow. Relaxation of the lithosphere allows subsidence near

1666-457: The local hydrostatic head. Orogen topography is the major driving force of fluid migration. The heat from the lower crust moves via conduction and groundwater advection . Local hydrothermal areas occur when deep fluid flow moves very quickly. This can also explain very high temperatures at shallow depths. Other minor constraints include tectonic compression, thrusting, and sediment compaction. These are considered minor because they are limited by

1715-436: The mountain belt can be compensated by isostasy at the upflex of the forebulge. The plate tectonic evolution of a peripheral foreland basin involves three general stages. First, the passive margin stage with orogenic loading of previously stretched continental margin during the early stages of convergence. Second, the "early convergence stage defined by deep water conditions", and lastly a "later convergent stage during which

1764-545: The moving thrust sheets and contains all the sediments charging from the active tectonic thrust wedge. This is where piggyback basins form. The foredeep is the thickest sedimentary zone and thickens toward the orogen. Sediments are deposited via distal fluvial, lacustrine, deltaic, and marine depositional systems. The forebulge and backbulge are the thinnest and most distal zones and are not always present. When present, they are defined by regional unconformities as well as aeolian and shallow-marine deposits. Sedimentation

1813-511: The physical or chemical changes, such as pressure or cementation . Thermal maturation of sediments is a factor of temperature and time and occurs at shallower depths due to past heat redistribution of migrating brines. Vitrinite reflectance, which typically demonstrates an exponential evolution of organic matter as a function of time, is the best organic indicator for thermal maturation. Studies have shown that present day thermal measurements of heat flow and geothermal gradients closely correspond to

1862-478: The plates to be underthrust beneath the other plate. The result is that the crust in the collision zone becomes as much as 80 kilometers (50 mi) thick, versus 40 kilometers (25 mi) for average continental crust. As noted above , the Airy hypothesis predicts that the resulting mountain roots will be about five times deeper than the height of the mountains, or 32 km versus 8 km. In other words, most of

1911-424: The rheological structure of the lithosphere underneath the foreland and the orogen are very different. The foreland basin typically shows a thermal and rheological structure similar to a rifted continental margin with three brittle layers above three ductile layers. The temperature underneath the orogen is much higher and thus greatly weakens the lithosphere. According to Zhou et al. (2003), "under compressional stress

1960-430: The same as when deformation began. Thus, it is crucial to consider non-GPS models to determine the long-term evolution of continental collisions and in how it helped develop the adjacent foreland basins. Comparing both modern GPS (Sella et al. 2002) and non-GPS models allows deformation rates to be calculated. Comparing these numbers to the geologic regime helps constrain the number of probable models as well as which model

2009-401: The slow rates of tectonic deformation, lithology and depositional rates, on the order of 0–10 cm yr , but more likely closer to 1 or less than 1 cm yr . Overpressured zones might allow for faster migration, when 1 kilometer or more of shaley sediments accumulate per 1 million years. Bethke & Marshak (1990) state that "groundwater that recharges at high elevation migrates through

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2058-441: The subsurface in response to its high potential energy toward areas where the water table is lower." Bethke & Marshak (1990) explain that petroleum migrates not only in response to the hydrodynamic forces that drive groundwater flow, but to the buoyancy and capillary effects of the petroleum moving through microscopic pores. Migration patterns flow away from the orogenic belt and into the cratonic interior. Frequently, natural gas

2107-434: The thickened crust moves downwards rather than up, just as most of an iceberg is below the surface of the water. However, convergent plate margins are tectonically highly active, and their surface features are partially supported by dynamic horizontal stresses, so that they are not in complete isostatic equilibrium. These regions show the highest isostatic anomalies on the Earth's surface. Mid-ocean ridges are explained by

2156-408: The thickness of the crust. This hypothesis was suggested to explain how large topographic loads such as seamounts (e.g. Hawaiian Islands ) could be compensated by regional rather than local displacement of the lithosphere. This is the more general solution for lithospheric flexure , as it approaches the locally compensated models above as the load becomes much larger than a flexural wavelength or

2205-444: The thrust, narrowing of basin, forebulge toward thrust. During times of thrusting, the lithosphere is stiff and the forebulge broadens. The timing of the thrust deformation is opposite that of the relaxing of the lithosphere. The bending of the lithosphere under the orogenic load controls the drainage pattern of the foreland basin. The flexural tilting of the basin and the sediment supply from the orogen. Strength envelopes indicate that

2254-560: The uplifted shorelines of Lake Bonneville . The concept was further developed in the 1950s by the Dutch geodesist Vening Meinesz . Three principal models of isostasy are used: Airy and Pratt isostasy are statements of buoyancy, but flexural isostasy is a statement of buoyancy when deflecting a sheet of finite elastic strength. In other words, the Airy and Pratt models are purely hydrostatic, taking no account of material strength, while flexural isostacy takes into account elastic forces from

2303-401: The water (ca. 1,000 kg m ). Thus, generally: For the simplified model shown the new density is given by: ρ 1 = ρ c c h 1 + c {\displaystyle \rho _{1}=\rho _{c}{\frac {c}{h_{1}+c}}} , where h 1 {\displaystyle h_{1}} is the height of the mountain and c

2352-531: The word 'isostasy' in 1889 to describe this general phenomenon. However, two hypotheses to explain the phenomenon had by then already been proposed, in 1855, one by George Airy and the other by John Henry Pratt . The Airy hypothesis was later refined by the Finnish geodesist Veikko Aleksanteri Heiskanen and the Pratt hypothesis by the American geodesist John Fillmore Hayford . Both the Airy-Heiskanen and Pratt-Hayford hypotheses assume that isostacy reflects

2401-503: Was a widespread phenomenon in mountainous areas. It was later found that the difference between the measured local gravitational field and what was expected for the altitude and local terrain (the Bouguer anomaly ) is positive over ocean basins and negative over high continental areas. This shows that the low elevation of ocean basins and high elevation of continents is also compensated at depth. The American geologist Clarence Dutton use

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