Carbonate Petrography

Carbonate petrography is the study of limestones, dolomites and associated deposits under optical or electron microscopes greatly enhances field studies or core observations and can provide a frame of reference for geochemical studies.

25 strangest Geologic Formations on Earth

The strangest formations on Earth.

What causes Earthquake?

Of these various reasons, faulting related to plate movements is by far the most significant. In other words, most earthquakes are due to slip on faults.

The Geologic Column

As stated earlier, no one locality on Earth provides a complete record of our planet’s history, because stratigraphic columns can contain unconformities. But by correlating rocks from locality to locality at millions of places around the world, geologists have pieced together a composite stratigraphic column, called the geologic column, that represents the entirety of Earth history.

Folds and Foliations

Geometry of Folds Imagine a carpet lying flat on the floor. Push on one end of the carpet, and it will wrinkle or contort into a series of wavelike curves. Stresses developed during mountain building can similarly warp or bend bedding and foliation (or other planar features) in rock. The result a curve in the shape of a rock layer is called a fold.

Showing posts with label basin. Show all posts
Showing posts with label basin. Show all posts

Basics of Basin Analysis



·         A sedimentary basin is an area in which sediments have accumulated during a particular time period at a significantly greater rate and to a significantly greater thickness than surrounding areas.


·         A low area on the Earth’s surface relative to surroundings e.g. deep ocean basin (5-10 km deep), intramontane basin (2-3 km a.s.l.)


·         Basins may be small (kms2) or large (106+ km2)


·         Basins may be simple or composite (sub-basins)


·         Basins may change in size & shape due to:

1.      erosion


2.      sedimentation

3.      tectonic activity

4.      eustatic sea-level changes

·         Basins may overlap each other in time


·         Controls on Basin Formation


1.      Accommodation Space,


a.       Space available for the accumulation of sediment

b.      T + E = S + W T=tectonic subsidence E= Eustatic sea level rise S=Rate of sedimentation W=increase in water depth

2.      Source of Sediment

a.       Topographic Controls

b.      Climate/Vegetation Controls

c.       Oceanographic Controls (Chemical/Biochemical Conditions)

·         The evolution of sedimentary basins may include:


1.      tectonic activity (initiation, termination)


2.      magmatic activity

3.      metamorphism

4.      as well as sedimentation

·         Axial elements of sedimentary basins:


1.      Basin axis is the lowest point on the basement surface


2.      Topographic axis is the lowest point on the depositional surface

3.      Depocentre is the point of thickest sediment accumulation

·         The driving mechanisms of subsidence are ultimately related to processes within the relatively rigid, cooled thermal boundary layer of the Earth known as the lithosphere. The lithosphere is composed of a number of tectonic plates that are in relative motion with one another. The relative motion produces deformation concentrated along plate boundaries which are of three basic types:


1.      Divergent boundaries form where new oceanic lithosphere is formed and plates diverge. These occur at the mid-ocean ridges.

2.      Convergent boundaries form where plates converge. One plate is usually subducted beneath the other at a convergent plate boundary. Convergent boundaries may be of different types, depending on the types of lithosphere involved. This result in a wide diversity of basin types formed at convergent boundaries.

3.      Transform boundaries form where plates move laterally past one another. These can be complex and are associated with a variety of basin types.

·         Many basins form at continental margins.

Using the plate tectonics paradigm, sedimentary basins have been classified principally in terms of the type of lithospheric substratum (continental, oceanic, transitional), the position with respect to a plate boundary (interplate, intraplate) and the type of plate margin (divergent, convergent, transform) closest to the basin.



·         Plate Tectonic Setting for Basin Formation


1.      Size and Shape of basin deposits, including the nature of the floor and flanks of the basin


2.      Type of Sedimentary infill

·         Rate of Subsidence/Infill


·         Depositional Systems

·         Provenance


·         Texture/Mineralogy maturity of strata


3.      Contemporaneous Structure and Syndepositional deformation


4.      Heat Flow, Subsidence History and Diagenesis

·         Interrelationship Between Tectonics - Paleoclimates - and Eustacy


1.      Anorogenic Areas------>


·         Climate and Eustacy Dominate


2.      Orogenic Areas--------->


Sedimentation responds to TectonismPlate Tectonics and Sedimentary Basin


   Types


SB = Suture Belt

RMP = Rifted margin prism

S C = Subduction complex

FTB = Fold and thrust belt

RA = Remnant arc
Wilson Cycle
 
about opening and closing of ocean basins and creation of continental crust.




Structural Controls on Sedimentary Systems in Basins Forming:


Stage 1: Capacity < Sediment


Fluvial sedimentation


Stage 2: Capacity = Sediment


Fluvial lacustrine Transition



Stage 3: Capacity > Sediment


Water Volume > excess capacity

Shallow-water lacustrine sedimentation



Stage 4: Capacity >> Sediment


Water volume = excess capacity

Deep-water lacustrine sedimentation



Stage 5: Capacity > Sediment


Water volume < excess capacity

Shallow-water lacustrine sedimentation     
Contributed by:


Rehan.A Farooqui

M.Sc Geology,,

University of Karachi.

Banded-iron formations (BIFs) - Evidence of Oxygen in Early Atmosphere

Our knowledge about the rise of oxygen gas in Earth’s atmosphere comes from multiple lines of evidence in the rock record, including the age and distribution of banded iron formations, the presence of microfossils in oceanic rocks, and the isotopes of sulfur.
However, this article is just focus on Banded Iron Formation.

BIF (polished) from Hamersley Iron Formation, West Australia, Australia

Summary: Banded-iron formations (BIFs) are sedimentary mineral deposits consisting of alternating beds of iron-rich minerals (mostly hematite) and silica-rich layers (chert or quartz) formed about 3.0 to 1.8 billion years ago. Theory suggests BIFs are associated with the capture of oxygen released by photosynthetic processes by iron dissolved in ancient ocean water. Once nearly all the free iron was consumed in seawater, oxygen could gradually accumulate in the atmosphere, allowing an ozone layer to form. BIF deposits are extensive in many locations, occurring as deposits, hundreds to thousands of feet thick. During Precambrian time, BIF deposits probably extensively covered large parts of the global ocean basins. The BIFs we see today are only remnants of what were probably every extensive deposits. BIFs are the major source of the world's iron ore and are found preserved on all major continental shield regions. 

Banded-iron formation (BIF)
is 
consists of layers of iron oxides (typically either magnetite or hematite) separated by layers of chert (silica-rich sedimentary rock). Each layer is usually narrow (millimeters to few centimeters). The rock has a distinctively banded appearance because of differently colored lighter silica- and darker iron-rich layers. In some cases BIFs may contain siderite (carbonate iron-bearing mineral) or pyrite (sulfide) in place of iron oxides and instead of chert the rock may contain carbonaceous (rich in organic matter) shale.

It is a chemogenic sedimentary rock (material is believed to be chemically precipitated on the seafloor). Because of old age BIFs generally have been metamorphosed to a various degrees (especially older types), but the rock has largely retained its original appearance because its constituent minerals are fairly stable at higher temperatures and pressures. These rocks can be described as metasedimentary chemogenic rocks.



                     Jaspilite banded iron formation (Soudan Iron-Formation, Soudan, Minnesota, USA
Image Credits: James St. John



In the 1960s, Preston Cloud, a geology professor at the University of California, Santa Barbara, became interested in a particular kind of rock known as a Banded Iron Formation (or BIF). They provide an important source of iron for making automobiles, and provide evidence for the lack of oxygen gas on the early Earth.

Cloud realized that the widespread occurrence of BIFs meant that
the conditions needed to form them must have been common on the ancient Earth, and not common after 1.8 billion years ago. Shale and chert often form in ocean environments today, where sediments and silica-shelled microorganisms accumulate gradually on the seafloor and eventually turn into rock. But iron is less common in younger oceanic sedimentary rocks. This is partly because there are only a few sources of iron available to the ocean: isolated volcanic vents in the deep ocean and material weathered from continental rocks and carried to sea by rivers.


Banded iron-formation (10 cm), Northern Cape, South Africa.
Specimen and photograph: A. Fraser
Most importantly, it is difficult to transport iron very far from these sources today because when iron reacts with oxygen gas, it becomes insoluble (it cannot be dissolved in water) and forms a solidparticle. Cloud understood that for large deposits of iron to exist all over the world’s oceans, the iron must have existed in a dissolved form. This way, it could be transported long distances in seawater from its sources to the locations where BIFs formed. This would be possible only if there were little or no oxygen gas in the atmosphere and ocean at the time the BIFs were being deposited. Cloud recognized that since BIFs could not form in the presence of oxygen, the end of BIF deposition probably marked the first occurrence of abundant oxygen gas on Earth (Cloud, 1968).
Cloud further reasoned that, for dissolved iron to finally precipitate and be deposited, the iron would have had to react with small amounts of oxygen near the deposits. Small amounts of oxygen could have been produced by the first photosynthetic bacteria living in the open ocean. When the dissolved iron encountered the oxygen produced by the photosynthesizing bacteria, the iron would have precipitated out of seawater in the form of minerals that make up the iron-rich layers of BIFs: hematite (Fe2O3) and magnetite (Fe3O4), according to the following reactions:
4Fe3 + 2O2 → 2Fe2O3
6Fe2 + 4O2 → 2Fe3O4
The picture that emerged from Cloud’s studies of BIFs was that small amounts of oxygen gas, produced by photosynthesis, allowed BIFs to begin forming more than 3 billion years ago. The abrupt disappearance of BIFs around 1.8 billion years ago probably marked the time when oxygen gas became too abundant to allow dissolved iron to be transported in the oceans.
Banded Iron Formation
Source is unknown

It is interesting to note that BIFs reappeared briefly in a few places around 700 millionyears ago,during a period of extreme glaciation when evidence suggests that Earth’s oceans were entirely covered with sea ice. This would have essentially prevented the oceans from interacting with the atmosphere, limiting the supply of oxygen gas in the water and again allowing dissolved iron to be transported throughout the oceans. When the sea ice melted, the presence of oxygen would have again allowed the iron to precipitate.

References:

1. Misra, K. (1999). Understanding Mineral Deposits Springer.
2. 
Cloud, P. E. (1968). Atmospheric and hydrospheric evolution on the primitive Earth both secular accretion and biological and geochemical processes have affected Earth’s volatile envelope. Science, 160(3829), 729–736.
3. 
James,H.L. (1983). Distribution of banded iron-formation in space and time. Developments in Precambrian Geology, 6, 471–490.

Sedimentary Basins

Sedimentary Basins

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

Ocean basins


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

Morphology of ocean basins


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

Depositional processes in deep seas

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

Debris-flow deposits

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

Turbidites

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

High- and low-efficiency systems

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

Initiation of mass flows

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

Composition of deep marine deposits

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

Classification of sedimentary basins

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

Intracratonic Basins 

Location and tectonic setting: in anorogenic areas on cratons.

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


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

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




Aulacogens 

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

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

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

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




Rift Basins

Location and tectonic setting: Within continental lithosphere on cratons.

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

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

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



Oceanic Rift Basins 

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

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

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

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






Passive Margin Basins 

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

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

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

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


 

Trenches 

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

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

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

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




 

Trench-Slope Basins 

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

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

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

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


 

Fore-Arc Basins 

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

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

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

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

Foreland Basins 

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

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

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


 

Remnant Basins 

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

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

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

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


 

Pull-Apart Basins 

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

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

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

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