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 sediments. Show all posts
Showing posts with label sediments. Show all posts

Siccar Point - the world's most important geological site and the birthplace of modern geology


Siccar Point is world-famous as the most important unconformity described by James Hutton (1726-1797) in support of his world-changing ideas on the origin and age of the Earth.

James Hutton unconformity with annotations - Siccar Point 



In 1788, James Hutton first discovered Siccar Point, and understood its significance. It is by far the most spectacular of several unconformities that he discovered in Scotland, and very important in helping Hutton to explain his ideas about the processes of the Earth.At Siccar Point, gently sloping strata of 370-million-year-old Famennian Late Devonian Old Red Sandstone and a basal layer of conglomerate overlie near vertical layers of 435-million-year-old lower Silurian Llandovery Epoch greywacke, with an interval of around 65 million years.
Standing on the angular unconformity at Siccar Point (click to enlarge). Photo: Chris Rowan, 2009
As above, with annotations. Photo: Chris Rowan, 2009





Hutton used Siccar Point to demonstrate the cycle of deposition, folding, erosion and further deposition that the unconformity represents. He understood the implication of unconformities in the evidence that they provided for the enormity of geological time and the antiquity of planet Earth, in contrast to the biblical teaching of the creation of the Earth. 

   
How the unconformity at Siccar Point formed.



At this range, it is easy to spot that the contact between the two units is sharp, but it is not completely flat. Furthermore, the lowest part of the overlying Old Red Sandstone contains fragments of rock that are considerably larger than sand; some are at least as large as your fist, and many of the fragments in this basal conglomerate are bits of the underlying Silurian greywacke. These are all signs that the greywackes were exposed at the surface, being eroded, for a considerable period of time before the Old Red Sandstone was laid down on top of them.
The irregular topography and basal conglomerate show that this is an erosional contact. Photo: Chris Rowan, 2009

The Siccar Point which is a rocky promontory in the county of Berwickshire on the east coast of Scotland.

Rock layers

Rock layers


In geology and related fields, a stratum (plural: strata) is a sedimentary rock layer or soil with inside reliable qualities that recognize it from different rock layers. The "stratum" is the crucial unit in a stratigraphic section and structures the study's premise of stratigraphy.

Characteristics of rock layers

Every rock layer is for the most part one of various parallel rock layers that lie one upon another, set around characteristic procedures. They may stretch out over a huge number of square kilometres of the Earth's surface. Strata are normally seen as groups of diverse shaded or contrastingly organized material uncovered in bluffs, street cuts, quarries, and waterway banks. Individual groups may fluctuate in thickness from a couple of millimetres to a kilometre or more. Every band speaks to a particular method of affidavit: stream residue, shoreline sand, coal bog, sand ridge, magma bed, and so on.

Naming of rock layers

Geologists study rock strata and sort them by the material of beds. Each particular layer is normally doled out to the name of sheet, generally in view of a town, stream, mountain, or locale where the arrangement is uncovered and accessible for study. For instance, the Burgess Shale is a thick introduction of dim, once in a while fossiliferous, shale uncovered high in the Canadian Rockies close Burgess Pass. Slight refinements in material in an arrangement may be portrayed as "individuals" (or now and again "beds"). Arrangements are gathered into "gatherings" while gatherings may be gathered into "supergroups".

Formation

An formation or geological formation is the basic unit of lithostratigraphy. An arrangement comprises of a sure number of rock strata that have an equivalent lithology, facies or other comparable properties. Developments are not characterized on the stone's thickness strata they comprise of and the thickness of distinctive formation can thus change broadly. 
The idea of formally characterized layers or strata is key to the geologic control of stratigraphy. Arrangements can be separated into individuals and are themselves regularly divided in gatherings.

Usefulness of formation

The definition and acknowledgement of formations permit geologists to correspond geologic strata crosswise over wide separations in the middle of outcrops and exposures of rock strata. 

Developments were at first depicted to be the crucial geologic time markers in view of relative ages and the law of superposition. The divisions of the land time scale were the formations depicted and put in sequential request by the geologists and stratigraphers of the eighteenth and nineteenth hundreds of years. 

Current modification of the geologic sciences has limited formations to lithologies, in light of the fact that lithological units are shaped by depositional situations, some of which may continue for a huge number of years and will transgress chronostratigraphic interims or fossil-based routines for relating rocks. For instance, the Hamersley Basin of Western Australia is a Proterozoic sedimentary bowl where up to 1200 million years of sedimentation is saved inside of the in place sedimentary stratigraphy, with up to 300 million years spoke to by a solitary lithological unit of grouped iron arrangement and shale. 

Geologic developments are typically sedimentary rock layers, yet might likewise be transformative rocks and volcanic streams. Molten nosy rocks are for the most part not separated into formations.

Defining lithostratigraphic formations

Formations are the main formal lithostratigraphic units into which the stratigraphic section all over ought to be partitioned totally on the premise of lithology. 

The difference in lithology between arrangements required to legitimize their foundation shifts with the multifaceted nature of the geography of an area and the point of interest required for geologic mapping and to work out its geologic history. 

Formations must have the capacity to be depicted at the size of geologic mapping honed in the area. The thickness of developments may run from not as much as a meter to a few thousand meters. 

Geologic arrangements are normally named for the geographic territory in which they were initially portrayed. 

Entirely, developments can't be characterized on whatever other criteria aside from essential lithology. Nonetheless, it is frequently helpful to characterize biostratigraphic units in light of paleontological criteria, chronostratigraphic units taking into account the stones' age, and chemostratigraphic units in view of geochemical criteria. 

Succession stratigraphy is an idea which challenges the thought of strict lithostratigraphic units by characterizing units in light of occasions in sedimentary bowls, for example, maritime relapses and transgressions. These groupings are a mix of chronostratigraphic units, connected by time, and depositional environment connected by the geologic occasions which happened around then, paying little respect to the grain size of the silt. 

The expression "formation" is regularly utilized casually to allude to a particular gathering of rocks, for example, those experienced inside of a sure profundity range in an oil well.

Flows, sediments and bedforms

Bedform


A bedform is a morphological feature formed by the interaction between a flow and cohesion less sediment on a bed. Ripples in sand in a flowing stream and sand dunes in deserts are both examples of bedforms, the former resulting from flow in water, the latter by airflow. The patterns of ripples and dunes are products of the action of the flow and the formation of bedforms creates distinctive layering and structures within the sediment that can be preserved in strata. Recognition of sedimentary structures generated by bedforms provides information about the strength of the current, the flow depth and the direction of sediment transport. A fluid flowing over a surface can be divided into a free stream, which is the portion of the flow unaffected by boundary effects, a boundary layer, within which the velocity starts to decrease due to friction with the bed, and a viscous sublayer,a region of reduced turbulence that is typically less than a millimetre thick. The thickness of the viscous sublayer decreases with increasing flow velocity but is independent of the flow depth. The relationship between the thickness of the viscous sublayer and the size of grains on the bed of flow defines an important property of the flow. If all the particles are contained within the viscous sublayer the surface is considered to be hydraulically smooth, and if there are particles that project up through this layer then the flow surface is hydraulically rough. As will be seen in the following sections, processes within the viscous sublayer and the effects of rough and smooth surfaces are fundamental to the formation of different bedforms. The following sections are concerned mainly with the formation of bedforms in flowing water in rivers and seas, but many of the fluid dynamic principles also apply to aeolian (wind-blown) deposits.

Current ripples



Flow within the viscous sublayer is subject to irregularities known as turbulent sweeps, which move grains by rolling or saltation and create local clusters of grains. These clusters are only a few grains high but once they have formed they create steps or defects that influence the flow close to the bed surface. Flow can be visualised in terms of streamlines in the fluid, imaginary lines that indicate the direction of flow. Streamlines lie parallel to a flat bed or the sides of a cylindrical pipe, but where there is an irregularity such as a step in the bed caused by an accumulation of grains, the streamlines converge and there is an increased transport rate. At the top of the step, a streamline separates from the bed surface and a region of boundary layer separation forms between the flow separation point and the flow attachment point downstream. Beneath this streamline lies a region called the separation bubble or separation zone. Expansion of flow over the step results in an increase in pressure and the sediment transport rate is reduced, resulting in deposition on the lee side of the step. 
Current ripples are small bedforms formed by the effects of boundary layer separation on a bed of sand. The small cluster of grains grows to form the crest of a ripple and separation occurs near this point. Sand grains roll or saltate up to the crest on the upstream stoss side of the ripple. Avalanching of grains occurs down the downstream or lee side of the ripple as accumulated grains become unstable at the crest. Grains that avalanche on the lee slope tend to come to rest at an angle close to the maximum critical slope angle for sand at around 308. At the flow attachment point there are increased stresses on the bed, which result in erosion and the formation of a small scour, the trough of the ripple.

Current ripples and cross-lamination


A ripple migrates downstream as sand is added to the crest and accretes on the lee slope. This moves the crest and hence the separation point downstream, which in turn moves the attachment point and trough downstream as well. Scour in the trough and on the base of the stoss side supplies the sand, which moves up the gentle slope of the stoss side of the next ripple and so a whole train of ripple troughs and crests advance downstream. The sand that avalanches on the lee slope during this migration forms a series of layers at the angle of the slope. These thin, inclined layers of sand are called cross-laminae, which build up to form the sedimentary structure referred to as cross-lamination. 
Linguoid ripple
When viewed from above current ripples show a variety of forms. They may have relatively continuous straight to sinuous crests (straight ripples or sinuous ripples) or form a pattern of unconnected arcuate forms called linguoid ripples. The relationship between the two forms appears to be related to both the duration of the flow and its velocity, with straight ripples tending to evolve into linguoid forms through time and at higher velocities. Straight and linguoid ripple crests create different patterns of cross-lamination in three dimensions. A perfectly straight ripple would generate cross-laminae that all dipped in the same direction and lay in the same plane: this is planar cross-lamination. Sinuous and linguoid ripples have lee slope surfaces that are curved, generating laminae that dip at an angle to the flow as well as downstream. As linguoid ripples migrate, curved cross-laminae are formed mainly in the trough-shaped low areas between adjacent ripple forms resulting in a pattern of trough cross-lamination.

Creating and preserving cross-lamination

Starved ripples
Current ripples migrate by the removal of sand from the stoss (upstream) side of the ripple and deposition on the lee side (downstream). If there is a fixed amount of sand available the ripple will migrate over the surface as a simple ripple form, with erosion in the troughs matching addition to the crests. These starved ripple forms are preserved if blanketed by mud. If the current is adding more sand particles than it is carrying away, the amount of sand deposited on the lee slope will be greater than that removed from the stoss side. There will be a net addition of sand to the ripple and it will grow as it migrates, but most importantly, the depth of scour in the trough is reduced leaving cross-laminae created by earlier migrating ripples preserved. In this way a layer of cross-laminated sand is generated. 
When the rate of addition of sand is high there will be no net removal of sand from the stoss side and each ripple will migrate up the stoss side of the ripple form in front. These are climbing ripples. When the addition of sediment from the current exceeds the forward movement of the ripple, deposition will occur on the stoss side as well as on the lee side. Climbing ripples are therefore indicators of rapid sedimentation as their formation depends upon the addition of sand to the flow at a rate equal to or greater than the rate of downstream migration of the ripples.

Constraints on current ripple formation

The formation of current ripples requires moderate flow velocities over a hydrodynamically smooth bed. They only form in sands in which the dominant grain size is less than 0.6 mm (coarse sand grade) because bed roughness created by coarser sand creates turbulent mixing, which inhibits the small scale flow separation required for ripple formation. Because ripple formation is controlled by processes within the viscous sublayer their formation is independent of water depth and current ripples may form in waters ranging from a few centimetres to kilometres deep. This is in contrast to most other subaqueous bedforms (subaqueous dunes, wave ripples), which are water-depth dependent. Current ripples can be up to 40 mm high and the wavelengths (crest to crest or trough to trough distances) range up to 500 mm. The ratio of the wavelength to the height is typically between 10 and 40. There is some evidence of a relationship between the ripple wavelength and the grain size, approximately 1000 to 1. It is important to note the upper limit to the dimensions of current ripples and to emphasise that ripples do not ‘grow’ into larger bedforms.

Dunes

Sand dunes
Beds of sand in rivers, estuaries, beaches and marine environments also have bedforms that are distinctly larger than ripples. These large bedforms are called dunes: the term ‘mega-ripples’ is also sometimes used, although this term fails to emphasise the fundamental hydrodynamic distinctions between ripple and dune bedforms. Evidence that these larger bedforms are not simply large ripples comes from measurement of the heights and wavelengths of all bedforms. The data fall into clusters which do not overlap, indicating that they form by distinct processes which are not part of a continuum. The formation of dunes can be related to large-scale turbulence within the whole flow; once again flow separation is important, occurring at the dune crest, and scouring occurs at the reattachment point in the trough. The water depth controls the scale of the turbulent eddies in the flow and this in turn controls the height and wavelength of the dunes: there is a considerable amount of scatter in the data, but generally dunes are tens of centimetres high in water depths of a few metres, but are typically metres high in the water depths measured in tens of metres.

Dunes and cross-bedding

Trough cross bedding.
Tabular cross bedding
The morphology of a subaqueous dune is similar to a ripple: there is a stoss side leading up to a crest and sand avalanches down the lee slope towards a trough. Migration of a subaqueous dune results in the construction of a succession of sloping layers formed by the avalanching on the lee slope and these are referred to as cross-beds. Flow separation creates a zone in front of the lee slope in which a roller vortex with reverse flow can form. At low flow velocities these roller vortices are weakly developed and they do not rework the sand on the lee slope. The cross-beds formed simply lie at the angle of rest of the sand and as they build out into the trough the basal contact is angular. Bedforms that develop at these velocities usually have low sinuosity crests, so the three-dimensional form of the structure is similar to planar cross-lamination. This is planar cross-bedding and the surface at the bottom of the cross-beds is flat and close to horizontal because of the absence of scouring in the trough. Cross-beds bound by horizontal surfaces are sometimes referred to as tabular cross-bedding. Cross-beds may form a sharp angle at the base of the avalanche slope or may be asymptotic (tangential) to the horizontal. At high flow velocities the roller vortex is well developed creating a counter-current at the base of the slip face that may be strong enough to generate ripples (counter-flow ripples), which migrate a short distance up the toe of the lee slope. 
A further effect of the stronger flow is the creation of a marked scour pit at the reattachment point. The avalanche lee slope advances into this scoured trough so the bases of the cross-beds are marked by an undulating erosion surface. The crest of a subaqueous dune formed under these conditions will be highly sinuous or will have broken up into a series of linguoid dune forms. Trough cross-bedding formed by the migration of sinuous subaqueous dunes typically has asymptotic bottom contacts and an undulating lower boundary.

Constraints on the formation of dunes

Dunes range in size from having wavelengths of about 600 mm and heights of a few tens of millimetres to wavelengths of hundreds of metres and heights of over ten metres. The smallest are larger than the biggest ripples. Dunes can form in a range of grain sizes from fine gravels to fine sands, but they are less well developed in finer deposits and do not occur in very fine sands or silts. This grain size limitation is thought to be related to the increased suspended load in the flow if the finer grain sizes are dominant: the suspended load suppresses turbulence in the flow and flow separation does not occur. The formation of dunes also requires flow to be sustained for long enough for the structure to build up, and to form cross-bedding the dune must migrate. Dunescale cross-bedding therefore cannot be generated by short-lived flow events. Dunes are most commonly encountered in river channels, deltas, estuaries and shallow marine environments where there are relatively strong, sustained flows.

Bar forms

Bars are bedforms occurring within channels that are of a larger scale than dunes: they have width and height dimensions of the same order of magnitude as the channel within which they are formed. Bars can be made up of sandy sediment, gravelly material or mixtures of coarse grain sizes. In a sandy channel the surfaces of bar forms are covered with subaqueous dune bedforms, which migrate over the bar surface and result in the formation of units of cross-bedded sands. A bar form deposit is therefore typically a cross-bedded sandstone as a lens-shaped body. The downstream edge of a bar can be steep and develop its own slip-face, resulting in large-scale cross-stratification in both sandstones and conglomerates. Bars in channels are classified in terms of their position within the channel (side and alternate bars at the margins, mid-channel bars in the centre and point bars on bends: and their shape.

Plane bedding and planar lamination

Horizontal layering in sands deposited from a flow is referred to as plane bedding in sediments and produces a sedimentary structure called planar lamination in sedimentary rocks. Current ripples only form if the grains are smaller than the thickness of the viscous sublayer: if the bed is rough, the small-scale flow separation required for ripple formation does not occur and the grains simply roll and saltate along the surface. Plane beds form in coarser sands at relatively low flow velocities (close to the threshold for movement), but as the flow speed increases dune bedforms start to be generated. The horizontal planar lamination produced under these circumstances tends to be rather poorly defined. 
Plane bedding is also observed at higher flow velocities in very fine- to coarse-grained sands: ripple and dune bedforms become washed out with an increase in flow speed as the formation of flow separation is suppressed at higher velocities. These plane beds produce well-defined planar lamination with laminae that are typically 5–20 grains thick. The bed surface is also marked by elongate ridges a few grain diameters high separated by furrows oriented parallel to the flow direction. This feature is referred to as primary current lineation (often abbreviated to ‘pcl’) and it is formed by sweeps within the viscous sublayer that push grains aside to form ridges a few grains high which lie parallel to the flow direction. The formation of sweeps is subdued when the bed surface is rough and primary current lineation is therefore less well defined in coarser sands. Primary current lineation is seen on the surfaces of planar beds as parallel lines of main grains which form very slight ridges, and may often be rather indistinct.

Supercritical flow

Flow may be considered to be subcritical, often with a smooth water surface, or supercritical, with an uneven surface of wave crests and troughs. These flow states relate to a parameter, the Froude number (Fr), which is a relationship between the flow velocity (y) and the flow depth (h), with ‘g’ the acceleration due to gravity. The Froude number can be considered to be a ratio of the flow velocity to the velocity of a wave in the flow. When the value is less than one, the flow is subcritical and a wave can propagate upstream because it is travelling faster than the flow. If the Froude number is greater than one this indicates that the flow is too fast for a wave to propagate upstream and the flow is supercritical. In natural flows a sudden change in the height of the surface of the flow, a hydraulic jump, is seen at the transition from thin, supercritical flow to thicker, subcritical flow. 
Where the Froude number of a flow is close to one, standing waves may temporarily form on the surface of the water before steepening and breaking in an upstream direction. Sand on the bed develops a bedform surface parallel to the standing wave, and as the flow steepens sediment accumulates on the upstream side of the bedform. These bedforms are called antidunes, and, if preserved, antidune cross-bedding would be stratification dipping upstream. However, such preservation is rarely seen because as the wave breaks, the antidune bedform is often reworked, and as the flow velocity subsequently drops the sediment is reworked into upper stage plane beds by subcritical flow. Well-documented occurrences of antidune cross-stratification are known from pyroclastic surge deposits, where high velocity flow is accompanied by very high rates of sedimentation.

Oceanic sediments

Pelagic sediments

The term pelagic refers to the open ocean, and in the context of sedimentology, pelagic sediments are made up of suspended material that was floating in the ocean, away from shorelines, and has settled on the sea floor. This sediment comprises terrigenous dust, mainly clay and some silt-sized particles blown from land areas by winds, very fine volcanic ash, particularly from major eruptions that send fine ejecta high into the atmosphere, and airborne particulates from fires, mainly black carbon. It also includes bioclastic material that may be the remains of calcareous organisms, such as foraminifers and coccoliths, and the siliceous skeletons of Radiolaria and diatoms. All of these particles reside in the ocean water in suspension, moved around by currents near to the surface, but when they reach quieter, deeper water they gradually fall down through the water column to settle on the seabed. The origin of the terrigenous clastic material is airborne dust, and much of this is likely to have come from desert areas. The particles are therefore oxidised and the resulting sediments are usually a dark red-brown colour. These ‘red clays’ are made up of 75% to 90% clay minerals and they are relatively rich in iron and manganese. They lithify to form red or red-brown mudstones. These pelagic red mudrocks are a good example of how the colour of a sedimentary rock should be interpreted with caution: it is tempting to think of all red beds as continental deposits, but these deep-sea facies are red too. The accumulation rate of pelagic clays is very slow, typically only 1 to 5 mm/yr , which means it could take up to a million years of continuous sedimentation to form just a metre of sediment. Pelagic sediments with a biogenic origin are the most abundant type in modern oceans, and two groups of organisms are particularly common in modern seas and are very commonly found in strata of Mesozoic and Cenozoic age as well. Foraminifera are single-celled animals that include a planktonic form with a calcareous shell about a millimetre or a fraction of a millimetre across. Algae belonging to the group chrysophyta include coccoliths that have spherical bodies of calcium carbonate a few tens of microns across; organisms this size are commonly referred to as nanoplankton. The hard parts of these organisms are the main contributors to finegrained deposits that form calcareous ooze on the sea bed: where one group is dominant the deposits may be called a nanoplankton ooze or foraminiferal ooze. Calcareous oozes accumulate at rates ten times that of pelagic clays, around 3 to 50 mm/yr . This sediment consolidates to form a fine-grained limestone, which is a lime mudstone using the Dunham Classification, although these deposits are often called pelagic limestones. The foraminifers are normally too small to be seen with the naked eye, and the coccoliths are only recognisable using an electron microscope.

An electron microscope is also required to see any details of the siliceous biogenic material: diatoms are only 5 to 50 mm across while Radiolaria are 50 to 500 mm, so the larger ones can be seen with the naked eye. They are made of opal, a hydrated amorphous form of silica that is relatively soluble, and diatoms in particular are often dissolved. Accumulations of this material on the sea floor are known as siliceous ooze and they form more slowly than calcareous oozes, at between 2 and 10 mm/yr . Upon lithification siliceous oozes form chert beds. The opal is not stable and readily alters to another form of silica such as chalcedony, which makes up the chert rock. Deep sea cherts are distinctive, thinly bedded hard rocks that may be black due to the presence of fine organic carbon, or red if there are terrigenous clays present. The Radiolaria can often be seen as very fine white spots within the rock and where this is the case they are referred to as radiolarian chert. These beds formed from the lithification of a siliceous ooze deposited in deep water (primary chert) should be distinguished from chert formed as nodules due to a diagenetic silicification of a rock (secondary chert). Secondary cherts are developed in a host sediment (usually limestone) and have an irregular nodular shape: they do not provide information about the depositional environment but may be important indicators of the diagenetic history.

Distribution of pelagic deposits

Pelagic sediments form a significant proportion of the succession only in places that do not receive sediment from other sources, so any ocean areas close to margins tend to be dominated by sediment derived from the land areas, swamping out the pelagic deposits. The distribution of terrigenous and bioclastic material on the ocean floors away from the margins is determined by the supply of the airborne dust, the biogenic productivity of carbonate-forming organisms, the productivity of siliceous organisms, the water depth and the ocean water circulation. The highest productivity of the biogenic material is in the warmer waters near the Equator and also in areas where there is a good supply of nutrients provided by ocean currents. In these regions there is a continuous ‘rain’ of calcareous and, to a much lesser extent, siliceous biogenic material down towards the sea floor: this ‘rain’ is less intense in cooler regions or areas with less nutrient supply. The solubility of calcium carbonate is partly dependent on pressure as well as temperature. At higher pressures and lower temperatures the amount of calcium carbonate that can be dissolved in a given mass of water increases. In oceans the pressure becomes greater with depth of water and the temperature drops so the solubility of calcium carbonate also increases. Near the surface most ocean waters are near to saturation with respect to calcium carbonate: animals and plants are able to extract it from seawater and precipitate either aragonite or calcite in shells and skeletons. As biogenic calcium carbonate in the form of calcite falls through the water column it starts to dissolve at depths of around 3000 m and in most modern oceans will have been completely dissolved once depths of around 4000 m are reached. This is the calcite compensation depth (CCD). Aragonite is more soluble than calcite and an aragonite compensation depth can be defined at a higher level in the water column than a calcite compensation depth. The calcite compensation depth is not a constant level throughout the world’s oceans today. The capacity for seawater to dissolve calcium carbonate depends on the amount that is already in solution, so in areas of high biogenic productivity the water becomes saturated with calcium carbonate to greater depths and higher pressures are required to put the excess of ions into solution. The depth of the CCD is also known to vary with the temperature of the water and the degree of deep water circulation that is present.

Above the CCD the remains of the less abundant siliceous organisms are swamped out by the carbonate material; below the CCD the skeletons of Radiolaria can form the main biogenic component of a pelagic sediment. High concentrations of siliceous organisms need not always indicate deep waters. The cold waters of polar regions favour diatoms over calcareous plankton and in pre-Mesozoic strata calcareous foraminifers and nanoplankton are not present. At water depths of around 6000 m the opaline silica that makes up radiolarians and diatoms is subject to dissolution because of the pressure and an opal compensation depth (or silica compensation depth) can be recognised. In the deepest ocean waters it may be expected that only pelagic clays would be deposited. In some parts of the world’s oceans this is the case, and there are successions of red-brown mudrocks in the stratigraphic record that are interpreted as hadal (very deep water) deposits. In some instances, these deepwater mudstones include thin beds of limestone and chert: radiolarian chert beds also sometimes include thin limestone beds. The occurrence of these beds might be explained in terms of fluctuations in the compensation depths, but a simpler explanation is that these deposits are actually turbidites and this can be verified by the presence of a very subtle normal grading within the beds. Carbonate, for example, can be deposited at depths below the CCD if it is introduced by a mechanism other than settling through the water column. If the material is brought into deep water by turbidity currents it will pass through the CCD quickly and will be deposited rapidly. The top of a calcareous turbidite may subsequently start to dissolve at the sea floor, but the waters close to the sea floor will soon become saturated with the mineral and little dissolution of a calcareous turbidite deposit occurs.

Hemipelagic deposits

Fine-grained sediment in the ocean water that has been directly derived from a nearby continent is referred to as hemipelagic. It consists of at least 25% non-biogenic material. Hemipelagic deposits are classified as calcareous if more than 30% of the material is carbonate, terrigenous if more than half is detritus weathered from the land and there is less than 30% carbonate, or volcanigenic if more than half is of volcanic origin, with less than 30% of the material carbonate. Most of the material is brought into the oceans by currents from the adjacent landmass and is deposited at much higher rates than pelagic deposits (between 10 and over 100mm/yr ). Storm events may cause a lot of shelf sediment to be reworked and redistributed by both geostrophic currents and sediment gravity underflows. A lot of hemipelagic material is also associated with turbidity currents: mixing of the density current with the ocean water results in the temporary suspension of fine material and this remains in suspension for long after the turbidite has been deposited. The provenance and hence the general composition of the hemipelagic deposit will be the same as that of the turbidite. Consolidated hemipelagic sediments are mudrocks that may be shaly and can have a varying proportion of fine silt along with dominantly clay-sized material. The provenance of the material forms a basis for distinguishing hemipelagic and pelagic deposits: the former will be compositionally similar to other material derived from the adjacent continent, whereas pelagic sediments will have a different composition. Clay mineral and geochemical analyses can be used to establish the composition in these cases. Mudrocks interbedded with turbidites are commonly of hemipelagic origin, representing a long period of settling from suspension after the short event of deposition directly from the turbidity current.

Chemogenic sediments

A variety of minerals precipitate directly on the sea floor. These chemogenic oceanic deposits include zeolites (silicates), sulphates, sulphides and metal oxides. The oxides are mainly of iron and manganese, and manganese nodules can be common amongst deep-sea deposits. The manganese ions are derived from hydrothermal sources or the weathering of continental rocks, including volcanic material, and become concentrated into nodules a few millimetres to 10 or 20 cm across by chemical and biochemical reactions that involve bacteria. This process is believed to be very slow, and manganese nodules may grow at a rate of only a millimetre every million years. They occur in modern sediments and in sedimentary rocks as rounded, hard, black nodules. At volcanic vents on the sea floor, especially in the region of ocean spreading centres, there are specialised microenvironments where chemical and biological activity result in distinctive deposits. The volcanic activity is responsible for hydrothermal deposits precipitated from water heated by the magmas close to the surface. Seawater circulates through the upper layers of the crust and at elevated temperatures it dissolves ions from the igneous rocks. Upon reaching the sea floor, the water cools and precipitates minerals to form deposits localised around the hydrothermal vents: these are black smokers rich in iron sulphide and white smokers composed of silicates of calcium and barium that form chimneys above the vent several metres high. The communities of organism that live around the vents are unusual and highly specialised: they include bacteria, tubeworms, giant clams and blind shrimps. Ancient examples of mid-ocean hydrothermal deposits have been found in ophiolite suites but fossil fauna are sparse.

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.

Mixed carbonate-clastic environments


The depositional environments described are made up of ‘pure’ carbonate and evaporite deposits that do not contain terrigenous clastic or volcaniclastic material. There are, however, modern environments where the sediments are mixtures of carbonate and other clastic materials, and in the stratigraphic record many successions consist of mixtures of limestones, sandstones and mudstones. These typically occur in shallow-marine settings. The changes from carbonate to non-carbonate deposition and vice versa are the result of variations in the supply of terrigenous clastic material and this is in turn determined by tectonic or climatic factors, or fluctuations in sea level. Climate plays an important role in determining the supply of sand and mud to shallow marine environments. Under more humid conditions, the increased run-off on the land surface results in more sediment being carried by rivers, which are themselves more vigorous and hence deliver more sediment to the adjacent seas. A change to a wetter climate on an adjacent landmass will therefore result in increased deposition of sand and mud, which will suppress carbonate production on a shelf. Alternation of beds of limestone with beds of mudstone or sandstone may therefore be due to periodic climatic fluctuations of alternating drier and wetter conditions. However, other mechanisms can also cause fluctuations in the supply of detritus from the continent to parts of the shelf. Tectonic uplift of the landmass can also increase the sediment supply by increasing relief and hence the rate of erosion. Tectonic activity can also result in subsidence of the shelf, which will make the water deeper across the shelf area: a relative sea-level rise will have the same effect. With increased water depth, more of the shelf area will be ‘starved’ of mud and sand, allowing carbonate sedimentation to occur in place of clastic deposition. Fluctuations in sea level may therefore result in alternations between limestone and mudstone/sandstone deposition.


Carbonate deposits can co-exist with terrigenous clastic and volcaniclastic sediments under certain conditions. Deltas built by ephemeral rivers in arid environments may experience long periods without supply of debris and during these intervals carbonates may develop on the delta front, for example, in the form of small reefs that build up in the shallow marine parts of ephemeral fan-deltas. Time intervals between eruption episodes in island arc volcanoes may be long enough for small carbonate platforms to develop in the shallow water around an island volcano, giving rise to an association between volcanic and carbonate deposition.

Characteristics of shallow marine carbonates 

  • lithology – limestone. 
  • mineralogy – calcite and aragonite. 
  • texture – variable, biogenic structures in reefs, well sorted in shallow water. 
  • bed geometry – massive reef build-ups on rimmed shelves and extensive sheet units on ramps. 
  • sedimentary structures–cross-bedding in oolite shoals. 
  • palaeocurrents – not usually diagnostic, with tide, wave and storm driven currents. 
  • fossils – usually abundant, shallow marine fauna most common. 
  • colour – usually pale white, cream or grey. 
  • facies associations – may occur with evaporites, associations with terrigenous clastic material may occur.

Characteristics of marine evaporites 

  • lithology – gypsum, anhydrite and halite. 
  • mineralogy – evaporite minerals. 
  • texture – crystalline or amorphous. 
  • bed geometry – sheets in lagoons and barred basins, nodular in sabkhas. 
  • sedimentary structures – intrastratal solution breccias and deformation. 
  • palaeocurrents – rare. 
  • fossils – rare. 
  • colour – typically white, but may be coloured by impurities. 
  • facies associations – often with shallow marine carbonates.

Marine evaporites


Evaporite deposits in modern marine environments are largely restricted to coastal regions, such as evaporite lagoons and sabkha mudflats. However, evaporite successions in the stratigraphic record indicate that precipitation of evaporite minerals has at times occurred in more extensive marine settings.

Platform evaporites

In arid regions the restriction of the circulation on the inner ramp/shelf can lead to the formation of extensive platform evaporites. On a gently sloping ramp a sand shoal can partially isolate a zone of very shallow water that may be an area of evaporite precipitation; the subtidal zone here often merges into a low-energy mudflat coastline. Shelf lagoons behind rims formed by reefs or sand shoals can create similar areas of evaporite deposition, although the barrier formed by a reef usually allows too much water circulation. Evaporite units deposited on these platforms can be tens of kilometres across.

Evaporitic basins (saline giants)

Evaporite sedimentation occurs only in situations where a body of water becomes partly isolated from the ocean realm and salinity increases to supersaturation point and there is chemical precipitation of minerals. This can occur in epicontinental seas or small ocean basins that are connected to the open ocean by a strait that may become blocked by a fall in sea level or by tectonic uplift of a barrier such as a fault block. These are called barred basins and they are distinguished from lagoons in that they are basins capable of accumulating hundreds of metres of evaporite sediment. To produce just a metre bed of halite a column of seawater over 75 m deep must be evaporated, and to generate thick succession of evaporite minerals the seawater must be repeatedly replenished. Deposition of the thick succession can be produced in three ways each of which will produce characteristic patterns of deposits. 
  1. A shallow-water to deep-basin setting exists where a basin is well below sea level but is only partly filled with evaporating seawater, which is periodically replenished. The deep-water setting will be evident if the basin subsequently fills with seawater and the deposits overlying the evaporites show deep marine characteristics such as turbidites. 
  2. A shallow-water to shallow-basin setting is one in which evaporites are deposited in salterns but continued subsidence of the basin allows a thick succession to be built up. The deposits will show the characteristics of shallow-water deposition throughout. 
  3. A deep-water to deep-basin setting is a basin filled with hypersaline water in which evaporite sediments are formed at the shallow margins and are redeposited by gravity flows into deeper parts of the basin. Normally graded beds generated by turbidites and poorly sorted deposits resulting from debris flows are evidence of redeposition. Other deep-water facies are laminated deposits produced by settling of crystals of evaporite minerals out of the water body. As a basin fills up, the lower part of the succession will be deeper water facies and the overlying succession will show characteristics of shallow-water deposition.
Deep-basin succession can show two different patterns of deposition. If the barred basin is completely enclosed the water body will gradually shrink in volume and area and the deposits that result will show a bulls-eye pattern with the most soluble salts in the basin centre. In circumstances where there is a more permanent connection a gradient of increasing salinity from the connection with the ocean to the furthest point into the basin will exist. The minerals precipitated at any point across the basin will depend on the salinity at the point and may range from highly soluble sylvite (potassium chloride) at one extreme to carbonates deposited in normal salinities at the other. If equilibrium is reached between the inflow and the evaporative loss then stable conditions will exist across the basin and tens to hundreds of metres of a single mineral can be deposited in one place. This produces a teardrop pattern of evaporite basin facies. Changes in the salinity and amount of seawater in the basins result in variations in the types of evaporite minerals deposited. For example, a global sea-level rise will reduce the salinity in the basin and may lead to widespread carbonate deposition. Cycles in the deposits of barred basins may be related to global sea-level fluctuations or possibly due to local tectonics affecting the width and depth of the seaway connection to the open ocean. Organic material brought into the basin during periods of lower salinity can accumulate within the basin deposits and be preserved when the salinity increases because hypersaline basins are anoxic. There are no modern examples of very large, barred evaporitic basins but evidence for seas precipitating evaporite minerals over hundreds of thousands of square kilometres exist in the geological record. These saline giants have over 1000 m thickness of evaporite sediments in them and represent the products of the evaporation of vast quantities of seawater. Evaporite deposits of latest Miocene (Messinian) age in the Mediterranean Sea are evidence of evaporative conditions produced by partial closure of the connection to the Atlantic Ocean. This period of hypersaline conditions in the Mediterranean is sometimes referred to as the Messinian salinity crisis.

Sedimentology

Sedimentology is the study of processes which are formation, transportation and deposition of sediments of continents as well the marine environments which forms the sedimentary rocks.

Sedimentary processes

Sedimentary rock natures varies with difference of the grains origin, size, shape and composition. Grains and pebbles particles may be derived from the erosion of existing rocks or from volcanic eruption. Organism also contributes to sediments which varies from micro organism to reefs. These have calcium carbonate shells which rest at floors to make sedimentary rock, limestone. Sedimentary rock formation involves transportation which can be aided by any of the medium air, water, ice, gravity or chemical, biological growth of sediments at depositional site. Few factor influence the accumulation of sediments ultimate results on the formation of rock which are chemistry, temperature and biological characteristics of desired location. Transportation and depositional processes can be interpretted by looking at the sediments layers, their size, shape and distribution pattern gives knowledge of the processes of transportation. Sedimentary structures also proves the type of transportation medium example ripple marks. As present is the key to past so keeping this in mind physical, chemical and biological characters of the sedimentary rock can be examined to interpret the sedimentation process which resulted the formation of rock. 

Sedimentary environment and facies

Every environment have its own characteristics either chemical, biological or physical. Studying these properties of the sedimentary rocks deposited and exposed on the surface reveals the secrets of environments at which they deposited. As an example of the fluvial environment, river base deposits will have gravely sandy material deposited while in floods on the banks of the river fine sediments will deposits providing habitat to vegetation. The channel will have lens of sand or conglomerate while the floodplain will have clay producing mud rock and sand at the base.
As every environment has own characteristic rocks so these are termed facies. Facies is a rock body with specific characteristics that reflects the condition under which it formed. For describing a facie its all characteristics as lithology, texture, sedimentary structures and fossil content are documented in order to fully interpret its environment of deposition.