Gary Nichols. Sedimentology and Stratigraphy(Second Edition)

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Second-order cycles

Superimposed on the first-order cycles, which show a

duration of hundreds of millions of years, is a pattern

of rises and falls with durations of tens of millions of

years. The causes of second-order sea-level changes

are thought to be changes in the rates of spreading at

mid-ocean ridges, which may account for changes in

ocean water levels on a scale of tens of millions of

years (Hallam 1963; Pitman 1978). The shorter term

global cycles in the Neogene may be accounted for by

long-term trends in glaciation and deglaciation.

Third-order cycles

Rises and falls of sea level with a magnitude of several

tens of metres and a periodicity of one to ten million

years are recognised throughout the Phanerozoic stra-

tigraphic record (Fig. 23.19). There is no general

agreement on the mechanisms that may cause sea-

level fluctuations at this third order of cyclicity

(Plint et al. 1992) and it may be that more than one

mechanism is responsible. Mechanisms that are

thought to be less likely include changing the lengths

and volumes of material in mid-ocean spreading cen-

tres because this operates too slowly; thermal expan-

sion and contraction of the world’s seawater and

changes in the volume of water in and on the conti-

nents cannot generate the magnitude of change indi-

cated. More likely is glacio-eustasy because it can

generate the appropriate magnitude of sea-level

change in a short period of time (Vail et al. 1977;

Haq et al. 1988). However, this mechanism requires

the existence of ice caps on continents in polar regions,

and although there is abundant evidence for this dur-

ing periods of major glaciations in the Ordovic-

ian–Silurian, the Carboniferous–Permian and the

Neogene–Quaternary, there is some doubt about

other periods. In particular, the mid-Cretaceous was

one of the warmest periods of the Phanerozoic

and there is doubt over whether there were ice caps

present at this time, although there do seem to be signs

of global sea-level fluctuations. Other mechanisms that

may generate the frequency and magnitude implied

by the third-order cycle curve are applicable only to

individual basins. Changes in the regional tectonic

stresses acting on a basin may result in basin-wide

500

300

100

400

200

Ceno-

zoic

MesozoicPhanerozoic

Falling

Sea level

Rising

1st-order cycles

2nd-order cycles

3rd-order

cycles

Fig. 23.19 First-, second- and third-

order sea-level cycles are considered to be

global signatures due to tectonic and

climatic controls outlined in Fig. 23.18.

(Modified from Vail et al. 1977.)

378 Sequence Stratigraphy and Sea-level Changes

subsidence or uplift (Cloetingh 1988) although the

rates at which these changes occur are not well

known.

The third-order cycles are of particular relevance

to sequence stratigraphy because they are of the

appropriate magnitude and period to be responsible

for the cycles of sea-level rise and fall that generate

depositional sequences. It must be emphasised

though, that not all depositional sequences in all

sedimentary basins formed as a response to global

sea-level changes. Local tectonic activity can also

generate the relative sea-level fluctuations required

to form depositional sequences, and in many cases it

seems likely that a combination of global eustatic

and local tectonic mechanisms was responsible for

the sea-level changes that are recoded in depositional

sequences.

23.8.6 Short-term changes in sea level

In addition to these three orders of cyclicity, shorter

term changes in sea level have also been recognised.

These are fourth-order cycles of 200,000 to 500,000

years duration and fifth-order cycles lasting 10,000

to 200,000 years. The magnitude of sea-level change

in these cycles ranges from a few metres to 10 or

20 m, although short-term sea-level changes in the

Quaternary have much higher magnitudes. Evidence

for frequent relative changes in sea level of a few

metres has been found in successions throughout

the stratigraphic record and in sequence stratigraphy

terminology these are referred to as parasequences.

A very detailed record of sea-level changes has

been established for the Quaternary and related to

estimates of palaeotemperature from the oxygen iso-

tope record (21.5.3). These indicate that changes in

global climate caused periodic wasting and accretion

of ice masses with a cyclicity of tens to hundreds of

thousands of years. These global climatic variations

have been related to the behaviour of the Earth in its

orbit around the Sun and changes in the axis of

rotation. Three orbital rhythms were recognised and

their periodicity calculated by the mathematician

Milankovitch, and these cycles are commonly

known as Milankovitch cycles (Fig. 23.20). The

longest period rhythm, approximately 100,000

years, is due to changes of the eccentricity of the

Earth’s orbit around the Sun, that is, the orbit is

elliptical and changes its shape with time. The rota-

tion of the Earth on its axis shows two patterns of

variation. The axis of rotation is oblique with respect

to the plane of the Earth’s rotation about the Sun and

the angle of tilt changes over a period of 40,000 years

between 21.58 and 24.58. The shortest rhythm is

21,000 years caused by the precession or ‘wobble’

in axis of rotation analogous to the behaviour of a

spinning top. These three cycles, working indepen-

dently and in combinations, are believed to exert

fundamental controls on the global climate leading

to cycles of global warming and cooling and hence

sea-level fluctuations.

Earth

Orbit

Changes in the eccentricity of the Earths orbit around the Sun

100 ka cycle

Changes in the obliquity (tilt) of the Earths axis of rotation

41 ka cycle

Sun

Earth

Axis of rotation

(Currently 23.5°)

Precession of the axis of rotation

22 ka cycle

Sun

Tilt of the axis changes

from being inclined

towards the Sun to

being inclined away

from the Sun

Fig. 23.20 Milankovitch cycles: the eccentricity of the

Earth’s orbit of the Sun, changes on the obliquity of the axis

of rotation of the Earth and the precession of the axis of

rotation may result in global climatic cycles on the scale of

tens of thousands of years.

Causes of Sea-level Fluctuations 379

23.8.7 Global synchroneity of

sea-level fluctuations

The sea-level curves presented by Vail et al., (1977)

and Haq et al. (1987) were originally considered to be

global signatures. Dating of the strata in which they

were recognised led to assertions that the curve itself

could be used as a correlation tool: if there was evi-

dence of a sea-level fall within a succession of strata

of, say, early Oligocene, then the sequence boundary

formed by the sea-level fall could be dated by match-

ing it to one of the ‘global’ sea-level falls. There are a

number of objections to using this approach to corre-

lation. First, the sequence boundary in the section

being examined may be the result of a local relative

sea-level fall and not related to global eustasy. Second,

some authors doubt whether all of the sea-level fluc-

tuations shown on these published charts are actually

global signatures. Third, most sedimentary succes-

sions are dated biostratigraphically, and in many

cases there are known to be errors of hundreds of

thousands of years when relating them to radiometric

ages: the length of some of the cycles falls within the

error in the dating of these cycles, so confidence in the

accuracy of the dates placed on the curve is not

assured (Miall 1992, 1997). The procedure of dating

depositional sequences by comparison with a global

sea-level chart has largely fallen out of favour,

although sea-level curves for particular areas have

been published more recently (Hardenbol et al.

1998). It is reasonable to construct a sea-level curve

for a particular continental margin and use this as a

tool in correlation across that margin, but detailed

global correlation is probably not appropriate. How-

ever, there is little doubt that the signature of global

eustasy is present in many parts of the stratigraphic

record, and in the latter half of the Cenozoic there

does seem to be a global signature.

23.9 SEQUENCE STRATIGRAPHY:

SUMMARY

The sequence stratigraphy approach to the analysis

of sedimentary successions has taken some time

to become widely accepted and widely used.

There were two issues in the earlier stages of its

development which caused problems. First, the con-

cept was initially linked with the idea that a global

eustatic sea-level curve could be established and used

as a means of correlating strata: doubts about the

curve led to doubts about the methodology, but

in fact the principles underlying sequence strati-

graphy are sound and valid whether there is a global

sea-level curve or not (Posamentier & James 1993).

A second barrier to acceptance was the plethora

of new terms that were introduced as the concepts

were developed. To some extent these served to

make the subject appear more complicated than it

actually is, because sequence stratigraphy is really

quite a straightforward and elegant approach to the

analysis of sedimentary successions. It is now applied

very extensively, especially in the hydrocarbon

exploration industry, and several texts are now avail-

able that provide comprehensive accounts of the

principles and applications of sequence stratigraphy.

FURTHER READING

Catuneanu, O. (2006) Principles of Sequence Stratigraphy.

Elsevier, Amsterdam.

Coe, A.L. (Ed.). (2003) The Sedimentary Record of Sea-level

Change. Cambridge University Press, Cambridge.

De Boer, P.L. & Smith, D.G. (Eds) (1994) Orbital Forcing and

Cyclic Sequences. Special Publication 19, International

Association of Sedimentologists. Blackwell Scientific Pub-

lications, Oxford.

Emery, D. & Myers, K.J. (Eds) (1996) Sequence Stratigraphy.

Blackwell Science, Oxford.

Miall, A.D. (1997) The Geology of Stratigraphic Sequences.

Springer-Verlag, Berlin.

Posamentier, H.W. & Allen, G.P. (1999) Siliciclastic Sequence

Stratigraphy: Concepts and Applications. Society of Economic

Paleontologists and Mineralogists, Tulsa, OK.

Van Wagoner, J.C., Mitchum, R.M., Campion, K.M. &

Rahmanian, V.D. (1990) Siliciclastic Sequence Stratigraphy

in Well Logs, Cores and Outcrop: Concepts for High Resolution

Correlation of Time and Facies. Methods in Exploration

Series 7, American Association of Petroleum Geologists,

Tulsa, OK.

380 Sequence Stratigraphy and Sea-level Changes

Sedimentary Basins

Sedimentary basins are regions where sediment accumulates into successions hundreds

to thousands of metres in thickness over areas of thousands to millions of square kilo-

metres. The underlying control on the formation of sedimentary basins is plate tectonics

and hence basins are normally classified in terms of their position in relation to plate

tectonic setting and tectonic processes. Each basin type has distinctive features, and the

characteristics of sedimentation and the stratigraphic succession that develops in a rift

valley can be seen to be distinctly different from those of an ocean trench. A stratigraphic

succession can therefore be interpreted in terms of plate tectonics and places the study

of sedimentary rocks into a larger context. The sedimentary rocks in a basin provide a

record of the tectonic history of the area. They also provide the record of the effects of

other controls on deposition, such as climate, base level and sediment supply.

24.1 CONTROLS ON SEDIMENT

ACCUMULATION

The issues of how and where sediment is preserved

could perhaps have been considered before a discussion

of environments of deposition, because not every river,

lake, delta, estuary or so on is necessarily a place where

sediments will accumulate and form a succession of

strata. In fact, the preservation of deposits that will

eventually form part of the sedimentary record is actu-

ally the exception, rather than the rule. The transitory

nature of deposition is most obvious in upland areas.

The deposits left by glaciers retreating a few thousand

years ago may be familiar as lateral and terminal mor-

aines in some of our modern landscapes, but they occur

in areas that are undergoing erosion, and will not be

preserved as glacial features in the stratigraphic record.

Similarly, the sediment that we currently see in rivers,

estuaries, deltas and coasts is mostly only passing

through on its way to the open seas where they may

be preserved on the shelf or in the deep seas.

The concept of ‘the present being the key to the past’,

introduced in Chapter 1, can be difficult to apply,

because most of what we see happening today in mod-

ern environments of deposition is not necessarily repre-

sentative of events that will lead to the formation of

sedimentary rocks. For example, tidal currents may

form bars of cross-bedded sands in an estuary, but

those sands may be washed back and forth by the tide

for millennia, with some material added by the river,

and some moved out to sea. To create a set of strata from

these processes, something else has to happen, usually

some form of change in the environment. At a small

scale this may be the change in the position of a river

due to avulsion leading to abandonment of the old river

course, or the shift in a lake shoreline due to a change in

climate covering the old lake margin deposits with

water and more sediment. However, at a larger scale it

is tectonic subsidence, local and regional changes in

the vertical position of the crust, that ultimately

allows sediment to become preserved as strata.

24.1.1 Tectonics of sedimentary basins

The importance of tectonic subsidence as a mechan-

ism for creating accommodation space has already

been considered in the context of sequence stratigra-

phy (23.1.4), but there is a broader implication of this

process. Put simply, without tectonics creating areas

that are ‘lows’ on the Earth’s surface, there would be

no long-term accumulation of sediment, no sedimen-

tary rocks and no stratigraphy as we know it. Places

where sediment accumulates are known as sedimen-

tary basins and they can range in size from a few

kilometres across to ocean basins covering half the

planet. A ‘basin’ can also be a geomorphological

feature, a bowl-shaped depression on the land surface

that may or may not be a place where sediment is

accumulating – in geology we are really only con-

cerned with basins that preserve strata, and provide

us with our record of depositional environments

through Earth history.

Distinct areas of sediment accumulation were

recognised by geologists in the late 19th and early

20th centuries. They were then referred to as ‘geo-

synclines’ and defined as broad down-folds in the

crust where successions of strata were first preserved

and then subsequently deformed. With the advent of

plate tectonic theory the geosynclinal concept became

redundant and it is now conventional to categorise

sedimentary basins in terms of their plate tectonic

setting (Ingersoll 1988; Busby & Ingersoll 1995).

24.1.2 Climate, sediment supply and

base-level controls

The role of tectonics in creating the accommodation

for sediment to accumulate is fundamental to sedi-

mentology and stratigraphy, but there are a number

of other factors that control the volume, type and

distribution of sediment. These are summarised in

Fig. 24.1. Climate, tectonics, bedrock geology and

Tectonics

Climate

Connection

to ocean

tidal and wave

processes

relief in

hinterland

accommodation

Facies and

facies distributions

sediment

pathways

Hinterland

bedrock

nature/volume

of detritus

weathering

processes

basin

dimensions

subsidence

regional/global

sea-level change

rates of

erosion

Fig. 24.1 The facies of

deposits in sedimentary

basins and their distribu-

tions in three dimensions

are controlled by climatic

and tectonic factors, the

nature of the hinterland

bedrock and the connection

with the oceans.

382 Sedimentary Basins

ocean connection/base level all interact within and

around all types of sedimentary basin to govern the

character of the basin-fill succession.

Connection to oceans and sea-level changes

The role of relative changes in sea level has been

considered during the discussion of sequence strati-

graphy in Chapter 23. In shallow marine environ-

ments the sea level directly determines the amount

of accommodation available for sediment to accumu-

late, but it also influences fluvial deposition and deep-

sea sedimentation. Sea-level changes do not necessa-

rily affect all basins because some are wholly within

continental landmasses and have no link or direct

exchange of water with the oceans. These basins of

internal drainage (or ‘endorheic’ basins) can form in

a variety of tectonic settings, principally as rifts, fore-

land basins and strike-slip basins (see below). They

may be dominated by lacustrine conditions, but in

more arid climates fluvial and aeolian processes dom-

inate (Nichols 2005).

Climatic effects of weathering, transport

and deposition

The significance of climate as a control on processes

has been considered in the context of a number of

different surface processes and depositional environ-

ments. To start with, weathering processes are deter-

mined by the availability of water and the

temperature (6.4): under warm, humid conditions

more clay minerals and ions in suspension are gener-

ated, whereas colder environments form more coarse

clastic material. The transport of sediment by water,

ice or wind is also climatically controlled, both in

terms of the amount of water available and the tem-

perature. Depositional processes in all continental

environments and many coastal settings are sensitive

to the climate: a comparison of clastic lagoons formed

in a temperate or tropical setting (13.3.2) and an

evaporite lagoon formed in an arid environment

(15.2.2) makes clear the importance of climate in

determining depositional facies.

Bedrock and topography controls

on sediment supply

Sediment supply is a further important factor, both in

terms of character and volume of material. It is

obvious that a delta cannot be a site of deposition of

sand if no sand is supplied by the river, and similarly a

beach cannot form a foreshore body of rounded peb-

bles if there is no gravel available to be deposited

there. A deposit derived from the weathering and

erosion of basaltic rock will have a very different

character to one derived from a limestone terrain.

The nature of all facies in all depositional environ-

ments is ultimately determined by the grain size of the

sediment available, the mineralogical and petrological

character of the detritus and the chemistry of the

water. The relationship between sediment supply

and accommodation was considered in the context

of relative sea-level changes in Chapter 23, but the

volume of sediment supply has an impact on the

nature of the whole basin fill. The availability of sedi-

ment is principally determined by tectonic controls on

uplift in the hinterland, but climate and bedrock char-

acter also play a role. If the rate of sediment supply

exceeds the rate of tectonic subsidence, the basin fills

up (is overfilled) and the facies will be shallow mar-

ine or continental. A low supply compared with sub-

sidence rate results in a basin that is underfilled or

starved: in a marine setting these basins will accumu-

late mainly deep-water facies. Continental basins that

are underfilled may end up below sea level (e.g. the

Dead Sea, Jordan, and Death Valley, USA).

24.1.3 Tectonic setting classification

of sedimentary basins

The movement of tectonic plates results in mountain

belts where two areas of continental crust collide,

subduction zones with associated volcanic arcs

where oceanic crust is consumed at plate margins,

oceans form at places where plates are moving apart

and major fault zones where plates move past each

other. All these different tectonic settings are also

areas where sediment can accumulate, and at a sim-

ple level three main settings of basin formation can be

recognised:

1 basins associated with regional extension within

and between plates;

2 basins related to convergent plate boundaries;

3 basins associated with strike-slip plate boundaries.

In the following discussion the main basin types

and the transitions between them are considered in

terms of the plate tectonic setting. An elementary

knowledge of plate tectonic processes and the nature

Controls on Sediment Accumulation 383

of continental and oceanic lithosphere is assumed, as

is an understanding of the basic terminology of struc-

tural geology. A more detailed consideration of the

tectonic setting of both modern and ancient basins

(Ingersoll 1988; Busby & Ingersoll 1995) indicates

that at least 20 types can be recognised. In addition

hybrid forms exist because of the complexities of plate

tectonic processes, for example where crustal exten-

sion is oblique and resulting basins have characteris-

tics of both a rift and a strike-slip setting.

24.2 BASINS RELATED TO

LITHOSPHERIC EXTENSION

The motion of tectonic plates results in some areas

where lithosphere is under extension and other places

where it is under compression. Horizontal stress

within continental crust causes brittle fracture in the

surface layers while the stretching is accommodated

by ductile flow in the lower part of the lithosphere. In

the early stages of this extension, rifts form and are

typically sites of continental sedimentation. If the

stretching continues, the continental lithosphere

may rupture completely and the injection of basaltic

magmas results in the formation of new oceanic crust

within the zone of extension. This stage is known as a

‘proto-oceanic trough’ and is the first stage in the

initiation of an ocean basin: the remnant flanks of

the rift become the passive margins of the ocean

basin as it develops. However, not all crustal exten-

sion follows the same path: continental rift basins

may exist for long periods without making the rift

to drift transition of forming an ocean basin, espe-

cially if the driving force for the extension fades.

One tectonic setting where lithospheric extension

occurs is associated with a ‘hot spot’, an area of

increased heat flow in the crust generated by thermal

plumes in the mantle. Rupture of the continental

lithosphere over a plume creates three branches

along which extension occurs, a triple junction of

plates that can be seen today centred on the Afar

Triangle where the East African Rift valley, the Red

Sea and the Gulf of Aden meet. These three exten-

sional regimes are in different stages of development –

continental rift, proto-oceanic trough and young

ocean basin respectively. On the other side of Africa,

an older triple junction now centred on the Niger

Delta had two arms forming the South Atlantic,

while the third arm, the Benue Trough, was a ‘failed

rift’ that subsequently became an area of intracra-

tonic subsidence.

Not all lithospheric extension is related to hot spots

and the formation of new ocean basins. Areas of

thickened crust and high heat flow due to astheno-

spheric upwelling, such as the Basin and Range Pro-

vince in western USA, are also regions of widespread

rift basin development as the upper layer of the crust

responds to the doming. Furthermore, in arc–trench

systems (24.3) local tectonic forces lead to the rifting

of the crust and the formation of intra-arc and subse-

quently backarc basins due to extension.

24.2.1 Rift basins (Fig. 24.2)

In regions of extension continental crust fractures to

produce rifts, which are structural valleys (Fig. 24.3)

bound by extensional (normal) faults (Leeder 1995).

The axis of the rift lies more-or-less perpendicular to

the direction of the stress. The down-faulted blocks

are referred to as graben and the up-faulted areas as

horsts. The bounding faults may be planar or listric,

and if the displacement is greater on one side they

form asymmetric valleys referred to as half-graben.

The structural weakness in the crust and high heat

flow associated with rifting may result in volcanic

activity. Uplift on the flanks of rifts due to regional

high heat flow and the effect of relative movements on

the rift-bounding faults creates local sediment sources

for rift valleys.

The controls on sedimentation in rift valleys are

a combination of tectonic factors that determine the

rift flank relief and hence availability of material, as

well as the pathways of sediment into the basin, and

climate, which influences weathering, water avail-

ability for transport and facies in the rift basin

Terrestrial

rift basin

continental crust

Fig. 24.2 Rift basins form by extension in continental crust:

sediment is supplied from the rift flanks or may also be

brought in by rivers flowing along the axis of the rift.

384 Sedimentary Basins

(Nichols & Uttamo 2005). Connection to oceans is

also important. Death Valley, California, is a terres-

trial rift valley, isolated from the sea and has an

arid climate, such that alluvial fan, desert dune and

evaporative lake environments are dominant. In con-

trast, the Gulf of Corinth, Greece, is a maritime rift

and is the site of fan-delta and deeper marine clastic

deposits (Leeder & Gawthorpe 1987). Extensional

basins with low clastic supply may be sites of carbo-

nate deposition. The patterns of sedimentation in rifts

evolve as the basins deepen, separate basins combine

and links to the marine realm become established

(Gawthorpe & Leeder 2000).

24.2.2 Intracratonic basins (Fig. 24.4)

Areas of broad subsidence within a continental block

(craton) away from plate margins or regions of oro-

geny are known as intracratonic basins (Klein

1995). The cratonic crust is typically ancient, and

with low relief: the area may be very large, but the

amount of subsidence is low and the rate is very slow.

The mechanism of subsidence varies between different

basins, as some are apparently related to antecedent

rifting episodes, whereas others are not. After the

cessation of rifting within continental crust there is

a change in the thermal regime of the area. When

continental crust is extended it is thinned and this

brings hotter mantle material closer to the surface.

Rifts are therefore areas of high heat flow, a high

geothermal gradient (rate of change of temperature

with depth). When rifting stops the geothermal gra-

dient is reduced and the crust in the region of the rift

starts to cool, contract and sink resulting in thermal

subsidence. Intracratonic basins that apparently

have no precursor rift history may also be a product

of thermal subsidence. Irregularities in the tempera-

ture distribution within the mantle associated with

cold crustal slabs relict from long-extinct subduction

zones create areas where there is downward move-

ment. Cratonic areas above these zones may be sub-

ject to subsidence and the formation of a broad,

shallow basin. Long-wavelength lithospheric buckling

has also been suggested as a mechanism for forming

intracratonic basins.

Fluvial and lacustrine sediments are commonly

encountered in intracratonic basins, although flooding

from an adjacent ocean may result in a broad epiconti-

nental sea. Intracratonic basins in wholly continental

settings are very sensitive to climate fluctuations as

increased temperature may raise rates of evaporation

in lakes and reduce the water level over a wide area.

24.2.3 Proto-oceanic troughs: the transition

from rift to ocean (Fig. 24.5)

Continued extension within continental crust leads to

thinning and eventual complete rupture. Basaltic

magmas rise to the surface in the axis of the rift and

start to form new oceanic crust. Where there is a thin

strip of basaltic crust in between two halves of a rift

system the basin is called a proto-oceanic trough

(Leeder 1995). The basin will be wholly or partly

flooded by seawater by the time this amount of exten-

sion has occurred and the trough has the form of a

narrow seaway between continental blocks. Sediment

supply to this seaway comes from the flanks of the

Fig. 24.3 Rift valleys are characterised by steep sides

formed by the extensional faults that form the basin (East

African Rift Valley).

Intracratonic

sag basin

continental crust

Fig. 24.4 Intracratonic basins are broad regions of subsi-

dence within continental crust: they are typically broad and

shallow basins.

Basins Related to Lithospheric Extension 385

trough, which will still be relatively uplifted. Rivers

will feed sediment to shelf areas and out into deeper

water in the axis of the trough as turbidity currents.

Connection to the open ocean may be intermittent

during the early stage of basin formation and in arid

areas with high evaporation rates the basin may peri-

odically desiccate. Evaporites may form part of the

succession in these circumstances and this phase of

basin development may be recognised by beds of gyp-

sum or halite in the lower part of a passive margin

succession.

24.2.4 Passive margins (Fig. 24.6)

The regions of continental crust and the transition to

oceanic crust along the edges of spreading oceans

basins are known as passive margins. The term

‘passive’ is used in this sense as the opposite to the

‘active’ margins between oceans and continents

where subduction is occurring. The continental

crust is commonly thinned in this region and there

may be a zone of transitional crust before fully oceanic

crust of the ocean basin is encountered. Transitional

crust forms by basaltic magmas injecting into con-

tinental crust in a diffuse zone as a proto-oceanic

trough develops. Subsidence of the passive margin is

due mainly to continued cooling of the lithosphere as

the heat source of the spreading centre becomes

further away, augmented by the load on the crust

due to the pile of sediment that accumulates (Einsele

2000).

Morphologically the passive margin is the conti-

nental shelf and slope (Fig. 11.1) and the clastic sedi-

ment supply is largely from the adjacent continental

land area. The climate, topography and drainage

pattern on the continent therefore determines the

nature and volume of material supplied to the shelf.

Adjacent to desert areas the clastic supply is low, and

the margin will be a starved margin, experiencing a

low clastic sedimentation rate. In contrast, a large

river system may carry large amounts of detritus

and build out a large deltaic wedge of sediment onto

the margin. In the absence of terrigenous detrital

supply, the shelf may be the site of accumulation of

large amounts of biogenic carbonate sediment,

although the volume and character of the material

will be determined by the local climate.

Passive margins are important areas of accumula-

tion of both carbonate and clastic sediment: they may

extend over tens to hundreds of thousands of square

kilometres and develop thicknesses of many thou-

sands of metres. They are also areas that are sensitive

to the effects of eustatic changes in sea level because

most of the deposition occurs in water depths of up to

100 m. Sea-level fluctuations of tens of metres result

in significant shifts in the patterns of sedimentation

on passive margins and the effects of a sea-level rise or

fall can be correlated over large distances in a passive

margin setting.

24.2.5 Ocean basins (Fig. 24.6)

Basaltic crust formed at mid-oceanic ridges is hot and

relatively buoyant. As the basin grows in size by new

magmas created along the spreading ridges, older

crust moves away from the hot mid-ocean ridge. Cool-

ing of the crust increases its density and decreases

relative buoyancy, so as crust moves away from the

ridges, it sinks. Mid-ocean ridges are typically at depths

of around 2500 m. The depth of the ocean basin

increases away from the ridges to between 4000 and

5000 m where the basaltic crust is old and cool.

Proto-oceanic trough

continental crust

oceanic crust

Fig. 24.5 With continued extension in a rift, the

lithosphere thins and oceanic crust starts to form in a proto-

oceanic trough where sedimentation occurs in a marine

setting.

Passive continental

margin

Oceanic basin

continental crust

oceanic crust

Fig. 24.6 An ocean basin is flanked by thinned continental

crust, which subsides to form passive margins to the ocean

basin.

386 Sedimentary Basins

The ocean floor is not a flat surface. Spreading

ridges tend to be irregular, offset by transform

faults that cre ate some areas of local topography.

Isolated volcanoes and linear chains of volcanic

activity related to hotspots (mantle plumes) such

as the Hawaiian Islands form submerged sea-

mounts or exposed islands. In addition to the for-

mation of volcanic roc ks in these areas, the

shallow water environment may be a site o f carbo-

nate production and the formation of reefs. In the

deeper parts of the ocean basins sedimentation is

mainly pelagic, consisting of fine-grained biogenic

detritus and clays. Nearer to the edges of t he

basins terrigenous clastic material may be depos-

ited as turbidites .

24.2.6 Obducted slabs

Most oceanic crust is subducted at destructive plate

margins, but there are circumstances under which

slabs of ocean crust are obducted up onto the over-

riding plate to lie on top of continental or other ocean-

ic crust. Outcrops of oceanic crust preserved in these

situations are known as ophiolites (Gass 1982).

Ophiolites may represent the stratigraphic succession

formed in an ocean basin or the fill of a backarc basin.

Until drilling in the deep oceans became possible

ophiolite complexes provided the only tangible evi-

dence of oceanic crust and deep-sea sediments. An

ophiolite suite consists of the ultrabasic and basic

intrusive rocks of the lower oceanic crust (peridotites

and gabbros), a dolerite dyke swarm which represents

the feeders to the basaltic pillow lavas that formed on

the ocean floor. The lavas are overlain by deep-ocean

sediments deposited at or close to the spreading

centre. If sea-floor spreading occurred above the

CCD these sediments would have been calcareous

oozes, preserved as fine-grained pelagic limestones.

Red clays and siliceous oozes deposited below the

CCD are lithified to form red mudstones and cherts.

Concentrations of metalliferous ores are common,

formed as hydrothermal deposits close to the

volcanic vents.

24.3 BASINS RELATED TO

SUBDUCTION

At convergent plate margins involving oceanic litho-

sphere subduction occurs (Fig. 24.7). The downgoing

ocean plate descends into the mantle beneath the

overriding plate, which may be either another piece

of oceanic lithosphere or a continental margin. As

the downgoing plate bends to enter the subduction

zone a trough is created at the contact between the

two plates: this is the ocean trench. The descending

slab is heated as it goes down and partially melts.

The magmas generated rise to the surface through

the overriding plate to create a line of volcanoes,

Hinge point

100

Continental crust

Oceanic crust

Backarc

(extensional)

Volcanic arc

Forearc

Accretionary

prism

Trench

Abyssal

plain

Fig. 24.7 An arc–trench system forms where oceanic crust is subducted at an ocean trench and the downgoing plate releases

magma at depth, which rises to form a volcanic arc. Sediment may accumulate in the trench, in the forearc basin between the

trench and the arc and in the region behind the arc called a backarc basin if there is subsidence due to extension (see also

retroarc basins, Fig. 24.13).

Basins Related to Subduction 387