BASIN ANALYSIS PREFACE

Transcription

1 PREFACE 1 BASIN ANALYSIS Basin analysis is a tool and attempts to answer these questions: Why should sedimentation occur in one place at a particular time? What is the spatial organization of large volumes of sediment? What are the factors that control their facies? And how do petroleum and mineralizing fluids move within basins? Geologists study sedimentary rocks to develop a critical understanding of their geologic history or to evaluate their economic potential. Effective study requires utilization of all the sedimentological and stratigraphic principles. Because most sedimentary rocks were deposited in basins, Basin analysis (Bogges, 2006) is an integrated program of study that involves application of sedimentologic, stratigraphic, and tectonic principles to develop a full understanding of the rocks that fill sedimentary basins for the purpose of interpreting their geologic history and evaluating their economic importance. The spatial distribution of depositional facies and variations in the environment of deposition through time will depend upon the tectonic setting, so a comprehensive analysis of the sedimentology and stratigraphy of an area must take place in the context of the basin setting. Sedimentary basin analysis (Nichols, 2009) is the aspect of geology that considers all the controls on the accumulation of a succession of sedimentary rocks to develop a model for the evolution of the sedimentary basin as a whole. Basin analysis studies aimed to understanding and predicting basin formation within the framework of plate tectonics and mantle convection; hydrocarbon generation and migration during basin evolution; present and historic ground-water flow and chemical transport; changes in basin fill and thermal evolution with tectonic environment; spatial and temporal variations of subsurface porosity and permeability; and the record of tectonics, climate, and sea-level change preserved in sedimentary basins.

2 SEDIMENTARY BASIN CONCEPT Sedimentary basins depressions on Earth's surface over geologic time are filled with sediments and organic materials that have been transported by wind, rivers, and ocean currents, they are come in many shapes and sizes, pervasive on Earth and form in response to complex geologic processes. Sedimentary basins can be hundreds to thousands of kilometers in horizontal dimensions and contain more than m 3 of buried materials. A sedimentary basin (Bogges, 2006) is a depression of some kind capable of trapping sediment. Sedimentary basins (Nichols, 2009) are regions where sediment accumulates into huge successions thickness over giant areas. The basin-filled materials is important in two respects. First, it preserves unique information regarding the history of tectonic, biologic, oceanographic, and atmospheric events during Earth's evolution. Second, basin fill contains most of the fuel and water, and many of the mineral resources, that are critical for society and industrial civilization. Some Basins are filled with strata deposited entirely in terrestrial environments, others with strata deposited below sea level in marine environments; many basins include both kinds of sediment. The formation of sedimentary basins is ultimately controlled by three elements: topography that defines the surface depressions that receive the sediments, the elevated regions that provide sediment sources, and the topographic and bathymetric gradients that transport sediments from source to basin. Understanding the evolution of sedimentary basins, and the reasons for their existence in particular places at specific times, can provide fundamental insights into a wide range of Earth processes. The imprint of geologic events left on the materials of sedimentary basins is the most detailed record of the history of Earth. Herein the research on basins overlaps almost the entire spectrum of earth sciences and thereby provides a unifying focus for research efforts in a wide range of sub disciplines. MECHANISMS OF BASIN FORMATION (SUBSIDENCE) Subsidence of the upper surface of the crust must take place to form a depression. Mechanisms that can generate sufficient subsidence to create basins are summarized in Table 1. Note in Table1 that isostatic compensation is an important aspect of loading. This concept assumes that: local compensation of the crust occurs as if Earth consists of a series of free-floating blocks. Adjacent blocks of crust of different 2

3 thickness and / or density structure will have different relative relief. Thus, adding a load to the crust (e.g., filling a basin with sediment) causes subsidence; removing a load (e.g., erosion of the crust) causes uplift. Then a basin originally filled with water will be deepened by the sediment load as the basin gradually accumulates sediment. In addition to the effects of loading, flexing of the crust also occurs, to various degrees depending upon the rigidity of the underlying lithosphere, because of tectonic forces: over thrusting, underpulling, underthrusting of dense lithosphere. Finally, thermal effects (e.g., cooling of lithosphere, increase in crustal density caused by changing temperature/pressure conditions) may also be important in basin formation. Table 1: MECHANISMS OF BASIN FORMATION (SUBSIDENCE) Crustal thinning: Mantle-lithospheric thickening: Sedimentary volcanic loading: Tectonic loading: Subcrustal loading: and Asthenospheric flow: Crustal densification: Extensional stretching, erosion during uplift, and magmatic withdrawal Cooling of lithosphere following either cessation of stretching or heating due to adiabatic melting or rise of asthenospheric melts Local isostatic compensation of crust and regional lithospheric flexure, dependent on flexural rigidity of lithosphere, during sedimentation and volcanism Local isostatic compensation of crust and regional lithospheric flexure, dependent on flexural rigidity of underlying lithosphere, during tectonic forces(overthrusting and/ or underpulling) Lithospheric flexure during underthrusting of dense lithosphere Dynamic effects of asthenospheric flow, commonly due to descent or delamination of subducted lithosphere Increased of crust density due to changing pressure/ temperature conditions and/ or emplacement of higherdensity melts into lower-density crust. Marine Environments Subdivisions Marine environments are classified into the benthic, for the sea bottom, and the pelagic, for the water mass. (Fig. 1) summarizes categories that are frequently used. (Rich, 1951) depended on effective wave base level to divide marine environments into shelf, slope, and basin floor Basin plains Several boundary plains at the sea bottom and within the water column are commonly used in a vertical subdivision of marginal-marine and marine environments. Essential critical interfaces that control sedimentary patterns and facies and the distribution of organisms are: 3

4 (1) The sea water level (horizontal plain identify general level of sea surface). (2) lower and the upper boundaries of the tides (control the distribution of organisms), (3) The base of the photic zone (controls the distribution of lightdependent phototrophic organisms), (4) The effective wave base level (the plain where wave effect becoming zero, above where bottom currents and wave action may lead to erosion and cementation or, the plain, which is separate high-energy traction deposits from low energy suspended deposits) (5) The storms wave base level (base of storms action on the sea bottom) (6) The O2 minimum zone (strongly limiting life on and in the sea bottom), (7) The thermocline (the layer of water that is too cold for most carbonateproducing organisms) (8) The pycnocline (the layer of water where salinity is too high for most organisms). (9) Aragonite compensation depth (ACD): Level in the oceans where aragonite is dissolved, about 3 Km. (10) Calcite compensation depth (CCD): The level in the deep oceans where the rate of dissolution of calcium carbonate (calcite) balances the rate of deposition and below which carbonate-free sediments accumulate. The level is characterized by a transition from carbonate ooze to deepmarine clay or siliceous ooze, about 4-5 Km. The CCD varies between ocean basins. The bathymetrical position of the ACD and the CCD depends on the fertility of the surface waters and the degree of undersaturation of deep water. (11) Silica compensation depth (SCD): Level in the oceans where silica (SiO2) is dissolved, about 6 Km. (12) Depositional base level (the interface between the sediments and liquids (water, air), it may be horizontal or inclined at 30 then called depositional dip) (13) Tectonic base level (the interface between basement of basin and sediment or it s a structural plain of faults that forming basin).tectonic level is the main plain affecting directly on sedimentation processes, there are two types: 4

5 a- Level without faults or smooth basements as platforms or low subsiding areas. The sediments here are undeformed with clear structures and features. b- Level with faults interface that exist on deeply faulted basement as Horst and Graben,the sediments here are deformed by faults without tectonic effect, these faults called growth faults(non-tectonic causes, compaction cause, recognized in quick sedimentation basin as deltas ex: Mississippi,Niger) the good example is Sirit basin in libya. The discrimination between two types are important in basin analysis. Vertical Zonations Benthic depth zones: The depth of the sea bottom and critical levels controlling the sedimentation subdivide the benthic environments into six zones: (1) Coastal supralittoral, supratidal zone (above high tide). (2) Littoral, intertidal zone, foreshore zone (between high and low tide), (3) Sublittoral, subtidal zone, shoreface zone (between low tide and effective/storme wave base, corresponding to the major part of the continental shelf), (4) Bathyal, offshore zone (approximately equal to the continental slope), (5) Abyssal, offshore zone (corresponding to the abyssal plains) (6) The hadal, offshore zone (deep-sea trenches). Pelagic depth zones: Five zones are defined by the vertical distribution of floating and swimming life. These are: the epipelagic zone (the upper of the oceans, extending to a depth of about 200 m), the mesopelagic, bathypelagic, abyssopelagic and hadopelagic zone (corresponding to oceanic zones below about m). Note that the boundaries of the benthic and pelagic depth zones are not fixed accurately. These boundaries reflect the situation in modern oceans that are not necessarily equivalent to depth zones visualized for ancient oceans. Horizontal (surface) Zonation The lateral distribution of pelagic organisms with respect to their distance from the coast characterizes two major zones of the ocean: 5

6 The neritic zone is the water that overlies the continental shelf, today generally with water depth less than 200 m and covering about 8% of the ocean floor. The oceanic zone refers to the water column beyond the shelf break, overlying the slope and the deep-sea bottoms, generally with water depths greater than 200 m and down to more than m. The term neritic is often used to describe sea bottom environments below the neritic water column, or shallow-marine environments characterized by significant terrigenous influx. Again, these water depths are not compatible with the situation in many ancient oceans. Fig. 1: summarizes categories and terms that are frequently used in marine environments Basin axis: (Fig 2) 1-Basin axis: a line connecting the lowest structural points of the basin, as in a synclinal axis, similarly the axes of troughs may be plunge. 2-topographic axis: a line connecting the lowest topographical points of the basin. 3-Depocenter axis: the part of the basin with the thickest sedimentary fill, this axis may migrate along basin with time. It is very important to note that the depocenter and basin axis need not coincide with one another, nor indeed with the topographic axis. This is particularly true of asymmetric basins with large amounts of terrigenous sedimentation on the limb of maximum uplift. In gently subsiding basins, with pelagic fine- 6

7 grained and turbidite fill, depocenter, axis, and topographic nadir may coincide. It is a common feature of many basins that the depocenter moves across the basin in time. This may reflect a migration of the topographic axis of the basin, or merely a lateral progradation of the main depositional site across an essentially stable basin floor. A B Fig 2:(a) basin axis, (b) lateral migration of depocenter with time of two basins 7

8 CONTROLS ON SEDIMENT ACCUMULATION 1-Tectonics of sedimentary basins: Plate tectonics provides a first-order control on sedimentation through its influence on the sediment source area. Tectonic forces control the type, size, shape, and location of the basins. Tectonic processes, together with sediment loading, further determine the rate of basin subsidence and thus the space available (accommodation) for sediment accumulation. The Tectonic processes controlling on creation of places where sediment accumulates are known as sedimentary basins, range in size from a few kilometres across to ocean basins covering half the planet, with specific geomorphological feature, and may or may not be a place where sediment is accumulating and/or preserving. Without tectonics creating areas that are lows on the Earth s surface, there would be no long-term accumulation of sediment, no sedimentary rocks and no stratigraphy as we know it. 2-Connection to oceans and sea-level changes: In shallow marine environments, the sea level directly determines the amount of accommodation available for sediment to accumulate, but it also influences fluvial deposition and deep-sea sedimentation. Sea-level changes do not necessarily affect all basins because some are wholly within continental landmasses and have no link or direct exchange of water with the oceans, as lacustrine conditions, fluvial / aeolian processes in more arid climates. 3-Climatic effects of weathering, transport and deposition: Weathering processes are determined by the availability of water and the temperature: under warm, humid conditions, more clay minerals and ions in suspension are generated, whereas colder environments form more coarse clastic material. The transport of sediment by water, ice or wind is also climatically controlled, in both, the amount of water available and the temperature. Depositional processes in all continental environments and many coastal settings are sensitive to the climate: a comparison of clastic lagoons formed in a temperate or tropical setting and an evaporite lagoon formed in an arid environment gives a clear importance of climate in determining depositional facies. 8

9 4-Bedrock and topography controls on sediment supply: The availability of sediment is principally determined by tectonic controls on uplift in the hinterland, but climate and bedrock character also play a role. Sediment supply is an important factor, in both, character and volume of material. The character: It is obvious that a delta cannot be a site of deposition of sand if no sand is supplied by the river, and similarly, a deposit derived from the weathering /erosion of basaltic rock will have a very different character to one derived from a limestone terrain. The volume of sediment supply: has an impact on the nature of the whole basin fill. -If the rate of sediment supply exceeds the rate of tectonic subsidence, the basin fills up (is overfilled) and the facies will be shallow marine or continental. -A low supply compared with subsidence rate results in a basin that is underfilled or starved: In a marine setting, these basins will accumulate mainly deep-water facies. Continental basins that are underfilled may end up below sea level (e.g. the Dead Sea, Jordan, and Death Valley, USA). Principles of plate tectonics: Earth surface includes many plates floating over asthenosphere (melted part of upper mantel zone with high density, ductile, and high viscosity). These plats moved because convection currents through asthenosphere. Plate-tectonic processes bring about major changes in continental masses and ocean basins through time. Continents break up and drift apart to create ocean basins as much as 500 km wide, which can subsequently close again as ocean-floor crust is subducted in trenches. The opening and closing of an ocean basin is referred to as a Wilson cycle (after Wilson, 1966). Wilson cycles begin with the formation of rift basins (floored by continental crust), which subsequently evolve into proto-oceanic troughs (partially floored by oceanic crust), and eventually into ocean basins, floored by oceanic crust with mid oceanic ridges and bordered by passive continental margins. After tens of millions of years or more, subduction zones develop around the ocean margins (active margins) and the ocean begins to close. Closure culminates with continental collision and the formation of an orogenic belt. 9

10 The entire process of basin formation and destruction requires perhaps 50 to 150 million years. The geologic record suggests that there have been many Wilson cycles in the history of each continent. Thus, few sedimentary basins remain unchanged with time, or in fixed positions, except perhaps some basins located on cratons well within continental margins. During the opening phases of a Wilson cycle, tectonic plates are moving apart (by rifting) to form divergent (passive) continental margins. The closing stages of a Wilson cycle are characterized by, plates moving toward each other, as oceanic crust is subducted (consumed) in trenches. Continental margins formed during ocean closing stage are called convergent (active) margins. During opening or closing of an ocean basin, some parts of plates may slide past each other without either diverging or converging. Such a setting is referred to as a transform margin. During a Wilson cycle, various kinds of sedimentary basins form in divergent, convergent, and transform settings, as well as in intraplate settings. Fig: 3: The Wilson cycle of ocean formation and closure.continental extension (a) oceanic spreading center (b) ocean enlargement(c) subduction ocean floor (d) closure ocean basin. Oceanic ridge subduction (e) continental collision (f). 10

11 Main earth phenomena resulted from plate movement 1- ocean and sea spreading (sea floor spreading) 2- The basaltic magma that extruded from mid oceanic ridges and faults rejuvenated. Thus, as the rocks are near of this positions, they are younger than the other rocks. 3- Creation of mid-oceanic ridges along shear faults, which are forming volcanic islands over or below seawater. 4- High rate of volcanic, seismic activities and thermal leakage in midoceanic ridges. 5- Homogeneity in topographical, lithological, and magnetic characteristics along both sides of fault in mid-oceanic ridges. 6- Presence of Ophiolite complex rocks (rhythemic succession of pelagic and basaltic rocks) away of mid-oceanic ridges. 7- Passive margins are developed (shelf, slope, and rise) in divergent continental margins by sea floor spreading. 8- Subduction occurs due to subsiding of, denser oceanic crust (SiMa) below continental crust (SiAl) as a result of convergent state. 9- Active margins are developed (Trenches) in convergent continental margins by subduction. 10- Subduction may occurs in two or one side according to active margins development. 11- Consuming of ocean floor accompanied with faults, earthquakes, and volcanic activities. 12- After consumption all of oceanic floor, continent-continent collision occurs, finally thrust faults and high mountains are formed. Tectonic setting classification of sedimentary basins All different tectonic settings are also areas where sediment can accumulate, and at a simple level, three main settings of basin formation can be recognized: A. basins associated with regional extension (within and between plates). B. basins associated with convergent plate boundaries (subduction). C. basins associated with strike-slip plate boundaries. D. basins associated with crustal loading. E. complex and hybrid basins 11

12 A- BASINS RELATED TO LITHOSPHERIC EXTENSION (DIVERGENT) The motion of tectonic plates results in some areas where lithosphere is under extension and other places where it is under compression. Horizontal stress within continental crust causes brittle fracture in the surface strata while the stretching is accommodated by ductile flow in the lower part of the lithosphere. In the early stages of this extension (stretching), rifts form and are typically sites of continental sedimentation. If the stretching continues, the continental lithosphere may rupture completely (brittle) and the injection of basaltic magmas results in the formation of new oceanic crust within the zone of extension, known as a proto-oceanic trough and it is the first stage in the initiation of an ocean basin: the remnant flanks of the rift become the passive margins of the ocean basin as it develops. However, not all-crustal extension follows the same path: continental rift basins may exist for long periods without making the rift to drift transition of forming an ocean basin, especially if the driving force for the extension fades. One tectonic setting where lithospheric extension occurs is associated with a hot spot, an area of increased heat flow in the crust generated by thermal plumes in the mantle. Rupture of the continental lithosphere over a plume creates three branches along which extension occurs, a triple junction of plates that can be seen today centered on the Afarica Triangle. These three extensional regimes are in different stages of development continental rift (East African Rift valley), proto-oceanic trough (the Red Sea) and young ocean basin (the Gulf of Aden). On the other side of Africa, an older triple junction now centered on the Niger Delta had two arms forming the South Atlantic, while the third arm, the Benue Trough, was a failed rift that subsequently became an area of intracratonic subsidence. Not all-lithospheric extension is related to hot spots and the formation of new ocean basins. Areas of thickened crust and high heat flow due to asthenospheric upwelling, such as the Basin and Range Province in western USA, are also regions of widespread rift basin development as the upper layer of the crust responds to the doming. Furthermore, in arc trench systems, local tectonic forces lead to the rifting of the crust and the formation of intra-arc and subsequently back arc basins due to extension. 12

13 1- Rift basins (Fig.4) In regions of extension continental crust fractures to produce rifts, which are structural valleys bound by extensional (normal) faults. The axis of the rift lies more-or-less perpendicular to the direction of the stress. The downfaulted blocks are referred to as graben and the up-faulted areas as horsts. The bounding faults may be planar or listric, and if the displacement is greater on one side they form asymmetric valleys referred to as half-graben. The structural weakness in the crust and high heat flow associated with rifting may result in volcanic activity. Uplift on the flanks of rifts due to regional high heat flow and the effect of relative movements on the rift-bounding faults creates local sediment sources for rift valleys. The controls on sedimentation in rift valleys are a combination of tectonic factors that determine the rift flank relief and hence availability of material, as well as the pathways of sediment into the basin, and climate, which influences weathering, water availability for transport and facies in the rift basin. Connection to oceans is also important. Death Valley, California, is a terrestrial rift valley, isolated from the sea and has an arid climate, such that alluvial fan, desert dune and evaporative lake environments are dominant. In contrast, the Gulf of Corinth, Greece, is a maritime rift and is the site of fan-delta and deeper marine clastic deposits. Extensional basins with low clastic supply may be sites of carbonate deposition. The patterns of sedimentation in rifts evolve as the basins deepen, separate basins combine and links to the marine realm become established. The best examples are: East African Rift and Gulf Of Suez Basins (fig:4), the anticlockwise divergent movement of Arabian pensula plate away of east Africa plate occurred in Miocene. The Oligocene rock may be eroded because doming uplift, fluvial silicaclastic rocks deposited on unconformable surface consequently and associated with volcanic rocks. Shallow marine carbonate facies during early Miocene indicating that the water covered the basin. The later silicaclastic formations referred to fandelta along faulted basin margins. The late Miocene evaporates rocks referred to restricted basin without connection to Mediterranean sea, this enhanced later after red sea opened due to continued movement along alaqba gulf. 13

14 Fig:4: rifting stages, gulf of Suez modern example. 14

15 2-Intracratonic sag basins (Fig:5) Areas of broad subsidence within a continental block (craton) away from plate margins or regions of orogeny are known as intracratonic basins. The cratonic crust is typically ancient, and with low relief: the area may be very large, but the amount of subsidence is low and the rate is very slow. The mechanism of subsidence varies; some are apparently related to antecedent rifting episodes, others are not. When continental crust is extended (early rifting phase), it is thinned and this brings hotter mantle material closer to the surface (high geothermal gradient). When rifting stops the geothermal gradient is reduced and the crust in the region of the rift starts to cool, contract and sink by thermal subsidence. Intracratonic basins that apparently have no precursor rift history may also be a product of thermal subsidence. Mantel Irregularities in the temperature distribution associated with cold crustal slabs relict from longextinct subduction zones create areas with downward movement. Cratonic areas above these zones may be subject to subsidence and the formation of a broad, shallow basin. Long-wavelength lithospheric buckling has also been suggested as a mechanism for forming intracratonic basins. Fluvial and lacustrine sediments are commonly encountered in intracratonic basins, although flooding from an adjacent ocean may result in a broad epicontinental sea. Intracratonic basins in wholly continental settings are very sensitive to climate fluctuations as increased temperature may raise rates of evaporation in lakes and reduce the water level over a wide area. Recent example of intracratonic basin is Chad lake basin western Africa. This basin related to antecedent arm of triple rifting episodes that are formed southern Atlantic ocean later. 15

16 A B Fig:5: A- Chad lake basin western Africa. This basin related to antecedent arm of triple rifting episodes. B- Broad shallow intracratonic sag basin. 16

17 3- Proto-oceanic troughs (fig: 6) The transition from rift to ocean, continued extension within continental crust leads to thinning and eventual complete rupture. Basaltic magmas rise to the surface in the axis of the rift and start to form new oceanic crust. Where there is a thin strip of basaltic crust in between two halves of a rift system the basin is called a proto-oceanic trough The basin will be wholly or partly flooded by seawater with the time, the trough has the form of a narrow seaway between continental blocks. Sediment supply to this seaway comes from the flanks of the trough, which will still be relatively uplifted. Rivers will feed sediment to shelf areas and out into deeper water in the axis of the trough as turbidity currents. Connection to the open ocean may be intermittent during the early stage of basin formation and in arid areas with high evaporation rates, the basin may periodically desiccate. Evaporates may form part of the succession in these circumstances and this phase of basin development may be recognized by beds of gypsum or halite in the lower part of a passive margin succession. Red sea is the one exclusive recent example of proto-oceanic troughs, where extension begin in middle-tertiary with triple junction included east African rift, Aden gulf, and red sea. The southern parts of red sea are more rifted (proto-ocean) than northern ones (early rifting). Sediments are clastics (from c-margins), carbonates, and evaporates (if restricted). Fig :6: a-east Africa rift, b- Gulf of Suez, c-red sea proto-oceanic basin 17

18 4- Passive margins (Fig. 7) The regions of continental crust and the transition to oceanic crust along the edges of spreading oceans basins are known as passive margins. The continental crust is commonly thinned in this region and becoming transitional zone crust to fully oceanic crust of the ocean basin. Transitional crust forms by basaltic magmas injecting into continental crust in a diffuse zone as a proto-oceanic trough develops. Subsidence of the passive margin is due mainly to continue cooling of the lithosphere as the heat source of the spreading center becomes further away, and/ or increased by loading of sediments accumulation. Morphologically the passive margin is the continental shelf and slope and the clastic sediment supply is largely from the adjacent continental land area. The climate, topography and drainage pattern on the continent therefore determines the nature and volume of material supplied to the shelf. Passive margins are important areas of accumulation of both carbonate and clastic sediment: they may extend over tens to hundreds of thousands of square kilometres and develop thicknesses of many thousands of metres. In the absence of terrigenous detrital supply, the shelf may be the site of accumulation of large amounts of biogenic carbonate sediment, although the volume and character of the material will be determined by the local climate. Adjacent to desert areas the clastic supply is low, and the margin will be a starved margin, experiencing a low clastic sedimentation rate. In contrast, a large river system may carry large amounts of detritus and build out a large deltaic wedge of sediment onto the margin. The shelf are also sensitive areas to the effects of eustatic changes in sea level. because most of the deposition occurs in water depths of up to 100 m, Sea-level fluctuations of tens of meters result in significant shifts in the patterns of sedimentation on passive margins and the effects of a sealevel rise or fall can be correlated over large distances in a passive margin setting. Eastern coast of North America (western of Atlantic Ocean) represents the recent example of passive margin basin. Atlantic rifting expected in Triassic, ocean floor spreading took place in Jurassic. Clastics shallow marine deposites are dominated during Mesozoic and Tertiary in the northern parts more than the southern parts (proto). Now, in Florida carbonates deposited in passive margin shelf. 18

19 A B Fig : 7: A-passive margin basin subdivisions and characteristic features. B- North America passive margin 19

20 5- Ocean basins (Fig. 8) As the basin grows in size by new magmas created along the spreading ridges, basaltic crust formed at mid-oceanic ridges is hot and relatively buoyant and older crust moves away from the hot mid-ocean ridge. Cooling of the crust increases its density and decreases relative buoyancy, so as crust moves away from the ridges, it sinks. Mid-ocean ridges are typically at depths of around 2500 m. The depth of the ocean basin increases away from the ridges to between 4000 and 5000m where the basaltic crust is old and cool. The ocean floor is not a flat surface. Spreading ridges tend to be irregular, offset by transform faults that create some areas of local topography. Isolated volcanoes and linear chains of volcanic activity related to hotspots (mantle plumes) such as the Hawaiian Islands form submerged seamounts or exposed islands. In addition to the formation of volcanic rocks in these areas, the shallow water environment may be a site of carbonate production and the formation of reefs. In the deeper parts of the ocean basins sedimentation is mainly pelagic, consisting of fine-grained biogenic detritus and clays. Nearer to the edges of the basins, terrigenous clastic material may be deposited as turbidites. Since the oceanic crust are denser than continental, it is subducted or destroyed as prisms. Oceanic basins stratigraphic units are not well preserved as continental basins, so, presence of non-destroyed oceanic successions are only found in obducted plates (Ophiolite complexes). Deep sea drilling gives a good record of oceanic succession. Pacific Ocean is the biggest modern oceanic basin. Pelagic sedimentation dominates in central deeper parts, whereas organic sediments dominates at tropical parts. Clastics dominates in slope (as turbidites), and silicates sediments (chert beds) in oceanic floor under CCD, and coral reef dominates around of the volcanic mountains. 20

21 Fig: 8: ocean floor basin 21

22 B- BASINS RELATED TO SUBDUCTION (CONVERGENT) (Fig: 9). At convergent margins, the oceanic lithosphere plate descends into the mantle beneath the overriding plate (either piece of oceanic lithosphere or a continental margin). As the downgoing plate bends to enter the subduction zone, an ocean trench trough is created between the two plates. The descending slab is heated as it goes down and partially melts. The magmas generated rise to the surface through the overriding plate to create a line of volcanoes, a volcanic arc. The magmas start to form when the downgoing slab reaches 90 to 150km depth. The arc trench gap (distance between the axis of the ocean trench and the line of the volcanic arc) will depend on the angle of subduction: at steep angles, the distance will be as little as 50 km and where subduction is at a shallow angle it may be over 200 km. Arc trench systems (forearc) are regions of plate convergence. the plate of an active arc must be in extension in order for magmas to reach the surface. The amount of extension is governed by: the relative rates of plate convergence / subduction and this is in turn influenced by the angle of subduction. If the angle of subduction is steep (if the downgoing plate consists of old, cold crust), then convergence is slower than subduction at the trench, the upper plate is in net extension and an extensional backarc basin forms. However, not all backarc areas are under extension: some are sites of the formation of a flexural basin due to thrust movements at the margins of the arc massif (retroarc basins). Fig : 9: basins related to subduction (convergent) 22

23 1- Trench basin (Fig. 10) Ocean trenches are elongate, narrow, very deep, gently curving, and starved troughs that form where an oceanic plate bends as it enters a subduction zone. The inner margin of the trench is formed by the overriding plate of the arc trench system. The bottoms of modern trenches are up to 10 km below sea level, twice as deep as the average bathymetry of the ocean floors. They are also narrow, sometimes as little as 5 km across, although they may be thousands of kilometers long. Trenches formed along margins flanked by continental crust tend to be filled with sediment derived from the adjacent land areas. Intra-oceanic trenches are often starved of sediment because the only sources of material (apart from pelagic deposits) are the islands of the volcanic arc. Transport of coarse material into trenches is by mass flows, especially turbidity currents that may flow for long distances along the axis of the trench. Chile trench in western coast of South America is good recent example of trench basin. When the Pacific Ocean crust plate subducted downward beneath of South America overriding continent plate. Chile trench dimension is 2500 km in long, 30 km in width, and 8 km in depth. The accumulated sediment thickness is various along this trench basin depending on source area relief. The sediment transported by turbidity currents by rivers across submarine canyons to forming submarine fans environments. Fig. 10: Trench basin and features 23

24 Obducted slabs (ophiolites complexes) Most oceanic crust is subducted, but in some cases, it is obducted up onto the overriding continental or oceanic crust plate. Ophiolites may represent the stratigraphic succession formed in an ocean basin. An ophiolite suite consists of: the ultrabasic and basic intrusive rocks of the lower oceanic crust (peridotites and gabbros), a dolerite dyke swarm, which represents the feeders to the basaltic pillow lavas that formed on the ocean floor. The lavas are overlain by deep-ocean deposites (micrite, mudstones, or cherts deposited at or close to the spreading center) depending on relation of CCD with spreading location. Concentrations of metalliferous ores are common, formed as hydrothermal deposits close to the volcanic vents. Accretionary complexes: The strata accumulated on the ocean crust and in a trench are not necessarily subducted with the crust at a destructive plate boundary. The mainly pelagic and turbidites sediments may be wholly or partly scraped off the downgoing plate and accrete on the leading edge of the overriding plate to form an accretionary complex or accretionary prism. These prisms or wedges of oceanic and trench sediments are best developed where there are thick successions of sediment in the trench. 2- Forearc basins (Fig.11) The area between the volcanic arc- and -the edge of accretionary complex formed on trench. The width of a forearc basin will therefore be determined by the dimensions of the arc trench gap, which is in turn determined by the angle of subduction. The basin floor either oceanic crust or a continental margin (subduction type). The sediments thickness in a forearc setting is partly controlled by the height of the accretionary complex. Subsidence here is due only to sedimentary loading. The main source of sediment to the basin is the volcanic arc. Forearc basin succession will consist of deep-water deposits at the base, shallowing up to shallow marine, deltaic and fluvial sediments at the top. Volcaniclastic debris is likely to be present in almost all cases. The good example is forearc basin between Sumatra island arc in Indonesia and Australian subducted plate. 24

25 Fig.:11: Forearc basin. 3- Backarc basins (Fig.12) backarc basins form where the rate of subduction is greater than the rate of plate convergence and the angle of subduction of the slab is steep. With further extension, the backarc basin may be developed to grow new ocean by spreading. Extensional backarc basins can form in either oceanic or continental plates (subduction type). The principal source of sediment in a backarc basin formed in an oceanic plate will be the active volcanic arc. More supplies are available if there is continental crust around the basin. Backarc basins are typically starved basin, containing mainly deep-water sediment of volcaniclastic and pelagic origin. Sea of Japan is the example. Fig.:11: Backarc basin. 25

26 C- BASINS RELATED TO STRIKESLIP TECTONICS (Fig:12) Since plate boundaries are not straight, they are consisting of a network of branching and overlapping individual faults, and the motion is not purely parallel. Therefore, areas of localized subsidence and uplift create topographic depressions for sediment accumulation and the uplifted source areas to supply them. Most basins in strike-slip belts are generally termed transtensional basins and are formed by three main mechanisms. First, the overlap of two separate faults can create regions of extension between them known as pull-apart basins. Such basins are typically rectangular or rhombic in plan with widths and lengths of only a few or tens of kilometers. They are unusually deep, especially compared with rift basins. Second, where there is a branching of faults, a zone of extension exists between the two branches forming a basin. Third, the curvature of a single fault strand leading to releasing bends (locally extensional and form elliptical zones of subsidence). Most strike-slip basins are bounded by deep faults are relatively small (hundred- thousand km 2 ), rapid subsidence, and often contain thicker successions (high rate) than basins of similar size formed by other mechanisms.. Typically, the margins are sites of deposition of coarse facies (alluvial fans and fan deltas) and these pass laterally over very short distances to lacustrine sediments in continental settings or marine deposits. In the stratigraphic record, facies are very varied and show lateral facies changes over short distances. Fig: 12: Strike slip basins 26

27 D- BASINS RELATED TO CRUSTAL LOADING When an ocean basin completely closes with the total elimination of oceanic crust by subduction the two continental margins eventually converge. Where two continental plates converge, subduction does not occur because the thick, low-density continental lithosphere is too buoyant to be subducted. Collision of plates involves a thickening of the lithosphere and the creation of an orogenic belt, a mountain belt formed by collision of plates. The Alps have formed by the closure of the Tethys Ocean as Africa has moved northwards relative to Europe, and the Himalayas are the result of a series of collisions related to the northward movement of India. The edges of the two continental margins that collide are likely to be thinned, passive margins. 1- 'peripheral foreland basin. (Fig: 13) As the crust thickens, it undergoes deformation, with metamorphism occurring in the lower parts of the crust and movement of material outwards from the orogenic belt along major deep or shallow faults. The combination of movement by thick-skinned tectonics (faults extended deeply into the crust) and thin-skinned tectonics (superficial thrust faults) transfers mass laterally and results in a loading of the crust adjacent to the mountain belt, because the mantle/asthenosphere below the lithosphere is mobile, they allow a flexural deformation of the loaded crust and formation of a peripheral foreland basin. The width of the basin will depend on the amount of load and the flexural rigidity of the foreland lithosphere. Rigid (typically older, thicker) lithosphere will respond to form a wide, shallow basin, whereas younger, thinner lithosphere flexes more easily to create a narrower, deeper trough. In the initial stages of foreland basin formation, the orogenic belt itself will not be high above sea level, and then a little detritus will be supplied by erosion of the orogenic belt. Early foreland basin sediments will therefore occur in a deep-water basin, with the rate of subsidence exceeding the rate of supply (starved). Turbidities are typical of this stage. When the orogenic belt is more mature and has built up a mountain chain there is an increase in the rate of sediment supply to the foreland basin. The Arabian gulf is a recent example. Sometimes thrusting may subdivide the basin to form piggy-back basins that lie on top of the thrust sheets and which are separate from the foredeep, the basin in front of all the thrusts. 27

28 Fig: 13: peripheral foreland basin. 2- retroarc foreland basin (Fig:14) At ocean continent convergence settings, the thick overriding continental plate and subduction related magmatism could also create a mountain belt. The loading of the crust on the opposite side of the arc to the trench leading to flexure, and the formation of a basin: these basins are called retroarc foreland basins because of their position behind the arc. The continental crust will be close to sea level at the time the loading commences, so most of the sedimentation occurs in fluvial, coastal and shallow marine environments. Continued subsidence occurs due to further loading of the basin margin by thrusted masses from the mountain belt, augmented by the sedimentary load. The main source of detritus is the mountain belt and volcanic arc. The Andes, along the western margin of South America, have been uplifted by crustal thickening and the intrusion of magma associated with subduction to the west. Thrust belts on the landward side of mountain chains in these settings result in loading and the formation of a retroarc foreland basin. Fig: 14: Retroarc foreland basin. 28

29 E- COMPLEX AND HYBRID BASINS Not all basins fall into the simple categories outlined above because they are the product of the interaction of more than one tectonic regime. This most commonly occurs where there is a strike-slip component to the motion at a convergent or divergent plate boundary. A basin may therefore partly show the characteristics of a peripheral foreland but also contain strong indicators of strike-slip movement. Such situations exist because plate motions are commonly not simply orthogonal or parallel and examples of both oblique convergence and oblique extension between plates are common. Modern example: Mississippi Embayment. THE RECORD OF TECTONICS IN STRATIGRAPHY within the Wilson Cycle, the rift basin deposits may be recognised by river and lake deposits overlying the basement, evaporates may mark the proto-oceanic trough stage, and a thick succession of shallow-marine carbonate and clastic deposits will record passive margin deposition. If this passive margin subsequently becomes a site of subduction, arcrelated volcanics will occur as the margin is transformed into a forearc region of shallow-marine, arc-derived sedimentation. Upon complete closure of the ocean basin, loading by the orogenic belt may then result in foreland flexure of this same area of the crust, and the environment of deposition will become one of deeper water facies. As the mountain belt rises, more sediment will be shed into the foreland basin and the stratigraphy will show a shallowing-up pattern. The same principles of using the character of the association of sediments to determine the tectonic setting of deposition can be applied to any strata of any age. An objective of sedimentary and stratigraphic analysis of a succession of rocks is therefore to determine the type of basin that they were deposited in, and then use changes in the sedimentary character as an indicator of changing tectonic setting. 29