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Basin analysis GEOLOGY Fourth stage ==================================================================================

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BASIN ANALYSIS

PREFACE

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.


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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 1015 m3of 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


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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: Extensional stretching, erosion during uplift, and magmatic

withdrawal

Mantle-lithospheric thickening:

Cooling of lithosphere following either cessation of stretching or heating due to adiabatic melting or rise of asthenospheric melts

Sedimentary and volcanic loading:

Local isostatic compensation of crust and regional lithospheric flexure, dependent on flexural rigidity of lithosphere, during sedimentation and volcanism

Tectonic loading: Local isostatic compensation of crust and regional lithospheric flexure, dependent on flexural rigidity of underlying lithosphere, during tectonic forces(overthrusting and/ or underpulling)

Subcrustal loading: Lithospheric flexure during underthrusting of dense lithosphere

Asthenospheric flow: Dynamic effects of asthenospheric flow, commonly due to descent or delamination of subducted lithosphere

Crustal densification: Increased of crust density due to changing pressure/ temperature conditions and/ or emplacement of higher-density 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:


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(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 light-

dependent 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 carbonate-

producing 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 deep-

marine 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:


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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 4 000 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:


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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 10 000 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-


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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


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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.


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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.


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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).


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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 mid-

oceanic 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


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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.


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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 down-

faulted 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.


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


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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 long-

extinct 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.


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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.


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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


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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 sea-

level 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.


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A

B

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

North America passive margin


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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.


21

Fig: 8: ocean floor basin


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)


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


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.


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.


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 km2), 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


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.


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 .


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, arc-

related 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.

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