Sedimentary basins

(Chapter 23 – Pgs 299-315)

Introduction:
The North Sea and the Gulf of California are modem-day examples of sedimentary basins.
The underlying control on the formation of sedimentary basins is plate tectonics and basins
are normally classified in terms of their position in relation to divergent, convergent and
strike-slip tectonic settings. The characteristics of sedimentation and the stratigraphic
succession which develops in a rift valley can be seen to be distinctly different from those of
an ocean trench.

1. Tectonics of sedimentary basins

The places where sediments accumulate into stratigraphic successions are collectively
referred to as sedimentary basins. At a simple level three main settings of basin formation
can be recognized:
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 context of these three settings nine main basin types cad be recognized (Fig. 1). In
addition, hybrid forms exist because of the complexities of plate tectonic processes.

1.1 Sedimentary basins and geosynclines

Large folds, hundreds of kilometres across, in the crust were envisage. The large downfolds
were termed geosynclines. A geosyncline was considered a site for sediment accumulation,
Fig. 1 Classification of sedimentary basins in
terms of tectonic setting.
which gradually sank, at least partly under the weight of the added material. Eventually,
the lower parts of the geosyncline would become deformed, and igneous intrusions and
extrusions would occur.
The forces which cause the subsidence and later deformation of the layers of strata in
geosynclinal theory are not adequately explained. On the contrary, the of plate tectonics
theory is a more comprehensive.

2. Basins related to crustal extension

Horizontal stresses associated with plate movements create extension along lines of
weakness in the lithosphere. These lines may lie within continental crust and may extend
into oceanic crust where they are zones of formation of new crust at spreading centers.
They terminate at triple junctions where three zones of extension meet. The triple junction
is normally associated with a hot-spot’, an area of increased heat flow in the crust generated
by thermal plumes in the mantle. This association suggests that there may be a relationship
between the formation of extensional plate boundaries and the location of hot mantle
plumes.
Rift basins start to form within continental crust (Fig. 1) and the earliest deposits are
typically terrestrial. Following the initiation of rifting, there are two possible courses for the
further development of the area.
1- Rifting may continue and split the continental crust completely to form a proto-oceanic
trough. Further extension leads to the formation of an ocean basin flanked by passive
margins.
2- Alternatively, the extension may cease followed by subsidence over a broad area within
the craton to form an intracratonic basin.

2.1 Rift basins

Rifts are structural valleys bound by extensional (normal) faults extending deeply into
continental crust. In extensional regimes the downfaulted blocks are referred to as graben
and the upfaulted areas as horsts. The bounding faults may be planar or listric, and they
form asymmetric valleys referred to as half-graben which are basins deeper on one side
than on the other. The axis of the rift lies more or less perpendicular to the direction of the
stress. Due to structural weakness in the crust volcanic activity is commonly associated with
rifting of continental crust, Thinning of the continental crust raises the geothermal gradient
in the area by bringing hot mantle closer to the surface.
Sediment in terrestrial rift valleys develop by aeolian, fluvial, alluvial and lacustrine
processes. Alluvial fans form along the rift margins. Sediment bodies locations are
controlled by the structure of rift, particularly the patterns of the extensional faults at the
edges and the tilting of the surface over large areas. The centers of rift basins may be
regions of fluvial sedimentation and/or regions of lake formation.
In maritime rifts, coarse sediment is deposited as
fan deltas at the basin margins and distributed in
the basin by wave, tide, storm and gravity
currents. In the absence of terrigenous clastic
material, carbonate sedimentation may be
dominant.
EXAMPLE OF A RIFT BASIN
The Gulf of Suez in Egypt is the north-western arm of
the Red Sea (Fig. 2). It is a rift basin which started to
form in mid-Tertiary times due to extension between
the Arabian plate and north Africa. The oldest rift
deposits are red beds of fluvial sandstone and
mudstone winch lie unconformably on pre-rift Eocene
limestone strata. Basaltic dykes, sills and lavas
amongst these red beds indicate that a small amount
of igneous activity occurred during the initial sifting
phase.
Widespread late Miocene evaporite beds indicate that
restriction of the connection to the marine waters of
the Mediterranean occurred at this time. A marine
connection was subsequently established with the Red
Sea and Gulf of Aden. The Gulf of Suez is no longer Fig. 2 Example of s basin formed by crustal
actively extending because the movement between the extension: Gulf of Suez, Egypt.
African and Arabian plates is now taken up along the
Gulf of Aqaba - eastern side of the Sinai Peninsula.
2.2 Intracratonic basins
¾ 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. When rifling stops the crust in the
region of the rift starts to cool and thicken. Cold rock is
denser, so as the continental crust cools it contracts and
sinks, resulting in thermal subsidence. This develops into
a broad area of subsidence within the continental block
(craton) and becomes an intracratonic basin.
¾ Intracratonic basins are typically broad but not very
deep and the rate of subsidence due to the cooling of the
lithosphere is slow. Fluvial and lacustrine sediments are
commonly encountered although flooding from an
adjacent ocean may result in a broad epicontinental sea.
Intracratonic basins are very sensitive to climate
fluctuations such as increased temperature.
EXAMPLE OF AN INTRACRATONIC BASIN
The Chad Basin in West central Africa is a basin of
internal drainage with a central lake, Lake Chad (Fig. 3). Fig. 3 An intracratonic basin: Chad
Basin map and cross-section.
Sediments in the Chad Basin are mostly fine-grained
elastic sediments formed on the river floodplains and
within the lake. The basin lies adjacent to and partly over
the Benue Trough (rift formed inMesozoic and Tertiary).
2.3 Proto-oceanic troughs: the transition from rift to ocean
¾ Continued extension within continental crust leads to thinning and eventual complete
rupture. Basaltic magmas form new oceanic crust along the axis of the rift. Where there is a
thin strip of basaltic crust in between two halves of a rift system the basin is called a proto-
oceanic trough. The basin will be wholly or partly flooded by sea water and the trough has
the form of a narrow seaway between continental blocks.
¾ Sediment supply comes from the relatively uplifted flanks of the trough. Rivers will feed
sediment to shelf areas and into deeper water. Connection to the open ocean may be
intermittent during the early stage of basin and in areas of high evaporation the basin may
periodically desiccate. Evaporites may form part of the succession in these circumstances.
EXAMPLE OF A PROTO-OCEANSC TROUGH
The Red Sea is the only modem example of a basin in the stage of transition from a rift
basin to an ocean basin (Fig. 4).
Extension in the Red Sea started in the mid-Tertiary.
The southern end of the sea lies at a triple junction with
the Gulf of Aden, and the East African rift system. The
floor of the Red Sea is thinned continental crust in the
north and newly formed oceanic crust in the south. This
reflects greater extension across the south the basin
than the north. Rivers from the African and Arabian
margins fed the Red Sea with elastic detritus, and
carbonate sedimentation occurs in warm waters. In
early stages of basin evaporites was important due to
intermittent connection with the open ocean. Fig. 4 The Red Sea is a protooceanic trough.
2.4 Passive margins
¾ Passive margins are the regions of continental crust along the
edges of ocean basins. The continental crust is commonly thinned in
this region and there may be a zone of transitional crust before
becoming a fully oceanic crust. Subsidence of the passive margin is
due mainly to continued cooling of the lithosphere as the heat
source, spreading centre, becomes further away, augmented by the
load on the crust due to the pile of sediment accumulates.
¾ Morphologically the passive margin is the continental shelf and
slope. Rates of erosion and transport affect the thickness and
distribution of elastic deposits on the passive margin. Adjacent to
large rivers very thick piles of sediment build out and vice versa
applies to adjoining desert area. In favorable climatic conditions,
high biogenic productivity and in the absence of terrigenous
detritus, carbonate occurs.
EXAMPLE OF A PASSIVE MARGIN
The eastern seaboard of North America is the passive margin of the
continental crust (Fig. 5). Rifting started to separate the continent
from Europe and Africa in the Triassic but sea floor spreading did
not commence until the Jurassic. The North Atlantic has continued
to open since that time. Along the northern part of the margin
shallow marine elastic sediments dominated the Mesozoic-Tertiary
succession, but further south carbonate sedimentation dominated
through most of the margin history. Fig. 5 The North America eastern seaboard is a
passive margin of the Atlantic Ocean.
2.5 Ocean basins
¾ As the basin grows in size by new magmas created along the spreading ridges, older
crust moves away from the hot mid-ocean ridge. Cooling of the crust increases its density
and decreases relative buoyancy, so it sinks. Mid-ocean ridges are typically at depths of
around 2500m. The depth of the ocean basin increases away from the ridges to 4000-5000m
where the basaltic crust is old and cool.
¾ The ocean floor is not a flat surface. Spreading ridges tend to be irregular, offset by
transform faults which create some areas of local topography. Isolated volcanoes and linear
chains of volcanic activity such as the Hawaiian islands form submerged seamounts or
exposed islands. In the deeper parts of the ocean basins sedimentation is mainly pelagic,
consisting of fine-grained biogenic detritus and clays. Nearer to the edges of the basins
terrigenous elastic material may be deposited as turbidities.
EXAMPLE OF AN OCEAN BASIN
The stratigraphy of an ocean basin will not be preserved as a succession of rocks. Oceanic
crust is denser than continental crust and at most convergent plate boundaries the ocean
basin sediments are either subducted or are deformed as they are incorporated into an
accretionary prism. It is only in situations where obduction occurs that relatively
undeformed parts of the ocean basin succession are found exposed on the continents. The
Pacific Ocean is the largest modern ocean basin. Away from the margins pelagic
sedimentation dominates in the basin. In equatorial areas biogenic productivity is high and
calcareous deposits occur. In more temperate regions the sea floor is below the calcite
compensation depth (CCD) and deposits are siliceous muds. Reefs form around volcanic
seamounts.
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 oceanic crust. Outcrops of oceanic crust preserved in these situations
are known as ophiolites. 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 which formed on the ocean floor (Fig. 6).
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. In absence of carbonates,
red clays and siliceous oozes are
lithified to form red mudstones and
cherts. Metalliferous ores formed by
hydrothermal processes close to the
volcanic vents are common.

Three of the best known ophiolites

are in Newfoundland, Cyprus and
Oman. The Cyprus Ophiolite and
the Semail Ophiolite in Oman are
both related to the closure of the
Tethys Ocean in the Mesozoic.

Fig. 6 Ophiolite Suites, show a layering which can be related to the

structure of oceanic crust.
3 Basins related to subduction
¾ At convergent plate margins the downgoing ocean plate descends into the mantle
beneath the over-riding plate. 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 and partially melts at 90-150 Km depth. The magmas generated
rise to the surface through the over-riding plate to create a line of volcanoes, a volcanic arc.
¾ The arc-trench gap (distance between the axis of the ocean trench and the line of the
volcanic arc) depends on the angle of subduction: steep angles the distance will be as llttle
as 50 Km, and shallow angle it may be over 200km (Fig. 7).

Fig. 7 Arc-trench systems at destructive plate margins are sites of sediment accumulation in the
trench, forearc and backarc areas. (After Dickinson & Seely 1979)
¾ Arc-trench systems are regions of plate convergence, yet there may be local extension as
well as compression in the upper plate. If convergence is faster than subduction some of the
shortening is taken up in the over-riding plate to form a retroarc foreland basin. If
convergence is slower than subduction at the trench the upper plate is in net extention and
an extensional backarc basin forms. A balance between the two results in a ‘neutral’ arc—
trench system.
3.1 Trenches
Ocean trenches are elongate, gently curving troughs which form where an oceanic plate
bends as it enters a subduction zone. The inner margin of the trench is the leading edge of
the over-riding plate of the arc-trench system. The bottoms of modern trenches are up to
10,000 m below sea level, twice as deep as the ocean floors. They are also narrow, as little as
5 km across, although they may be thousands of kilometers long. Trenches flanked by
continental crust tend to be filled with sediment derived from the adjacent land. Intra-
oceanic trenches are often starved of sediment. Transport of coarse material into trenches is
by mass flows, especially turbidity currents.
EXAMPLE OF A TRENCH BASIN
The Chile Trench off the western coast of South America developed where Pacific Oceanic
lithosphere is being subducted beneath the continental crust of South America. The Chile
Trench is over 2500km long and 30km wide. It varies in depth from 7000—8000m in the
northern part to around 5000 m in the south. The amount of sediment in the trench is
extremely variable.
3.2 Accretionary complexes
The sedimentary pile on the ocean crust and in the trench is not necessarily subducted
along with the crust. The pile of sediments may be wholly or partly scraped off the
downgoing plate and accrete on the leading edge
of the over-riding plate to form an accretionary
complex or accretionary prism. These prisms or
wedges of oceanic and trench sediments are best
developed where there are thick successions of
sediment in the trench. A subducting plate can be
thought of as a conveyor belt bringing ocean basin
deposits to the edge of the over-riding plate (Fig.
Fig. 8 Accretionary prisms form on the inner sides of ocean
8). Along the Sunda Trench, where the Indian trenches by the accretion of material from the downgoing
Ocean plate is sub- dueling beneath the plate on to the leading edge of the over-riding plate.

continental island of Sumatra, the accretionary

prism has grown up to sea level to form a chain of
islands between the trench and the coast of
Sumatra (Fig. 9). If ocean closure is complete, the
accretionary prism can become incorporated into
an orogenic belt e.g. Southern Uplands of
Scotland “deep water sediments of Ordovician
and Silurian arranged in a structural pattern
which is consistent with the formation of an
accretionary prism.

Fig. 9
3.3 Forearc basins
The width of a forearc basin depend on the arc- trench gap determined by the angle of
subduction. Its inner margin is the edge of the volcanic arc and the outer limit the
accretionary complex. The basin may be underlain by either oceanic crust or a continental
margin. The thickness of sediments will be partly controlled by the height of the
accretionary complex. Subsidence in the forearc region is due only to the load from the
sediment pile.
The main source of sediment is the volcanic arc Given sufficient supply of detritus a forearc
basin succession will consist of deep water deposits at the base shallowing upward to
shallow marine, deltaic and fluvial sediments at the top. Volcaniclastic debris present in
almost all cases.
EXAMPLE OF A FOREARC BASIN
The Indian Ocean subducts along a trench which lies
offshore of the island of Sumatra (Fig. 9). The Sumatra
forearc basin contains several thousand meters of strata
which date back to the early Miocene. Rivers from the
Sumatran highlands are still actively supplying sediment to
the basin, building up a coastal plain which passes into
deltas which prograde seawards.

Fig. 9 Subduction of Indian Ocean crust along the Sunda Trench has led to the formation of a trench, accretionary
prism and forearc basin on the south-western side of the island of Sumatra.
3.4 Backarc basins
When the upper plate in an arc—trench system is under extension it rifts in the region of the
volcanic arc where the crust is hotter and weaker. Initially the arc itself rifts and splits into two
parts, an active arc with continued volcanism closer to the subduction zone and a remnant arc.
As extension between the remnant and active arcs continues, a new spreading centre is formed to
generate basaltic crust. This region of extension and crustal formation is the backarc basin. The
basin continues to grow by spreading until renewed rifting in the active arc. Once a new backarc
basin is formed the older one is abandoned. Extensional backarc basins can form in either
oceanic or continental plates.
The principal source of sediment is the active volcanic arc. More abundant supplies are available
if there is continental crust. Backarc basins are typically under- filled, containing mainly deep
water sediment of volcaniclastic and pelagic origin.
EXAMPLE OF A BACICARC BASIN
Most of the active intra-oceanic backarc basins lie in
the western Pacific. The Sumisu rift is a backarc basin
lies south of Japan (Fig. 10). The backarc basin is over
100 km long and 40km wide. It is around 2000m deep
and has formed during the past 2 million years. A
thousand meters of sediment is developed: the lower
third consists of volcaniclastic breccia and flow
deposits, whilst the remainder is composed of
calcareous mudrock. These two sediment types reflect
the two main sources, the active volcanic arc and
pelagic sedimentation.
Fig. 10 The Sumisu rift is an example of a backarc basin
south of Japan.
3.5 Retroarc foreland basins
In compressional convergence regimes the over-riding continental plate shortens by the
development of a mountain belt. Magmas from the subduction zone also add material to the
upper plate in an arc along the mountain belt. Thickening of the crust results in the upward
and outward movement of masses of rock along thrusts and as nappes. As these thrust slices
move on to the continent on the opposite side of the arc to the trench, they add a load on to
it. Loading of the lithosphere causes it to bend and it is as a result of this flexure that a
basin forms.
The continental crust will be close to sea level so most of the sedimentation occurs in fluvial,
coastal and shallow marine environments. The main source of detritus is the mountain belt
and volcanic arc.
EXAMPLE OF A RETROAEC FORELAND BASIN
The eastern Andean basins are present-day
examples of basins formed by the convergence
between the eastern Pacific oceanic plates and the
South American continent (Fig. 11). It is up to
200km wide and contains up to 8000 m of
sediment. Throughout the basin evolution
sedimentation has been in a continental
environment but the facies show lateral variations
due to differences in climate. In arid regions
aeolian and playa environments are important,
whilst in the wetter regions fluvial facies dominate.
Fig. 11 In the eastern part of the Andes in South America lie retroarc
basins formed by the loading of the crust by the Andean mountain belt.
4 Basins related to continental collision
When an ocean completely closes, the two continental margins eventually converge. Where
two continental plates converge subduction does not occur because the thick, low-density
continental lithosphere is too buoyant to be subducted. Collision of plates involves a
thickening of the lithosphere and the creation of an orogenic belt, a mountain belt formed
by collision of plates.
The two continental margins which collide are likely to be thinned, passive margins.
Shortening initially increases the lithosphere thickness up to ‘normal’ values before it over-
thickens. As the crust thickens deformation with metamorphism occur in the lower parts of
the crust and faulting and folding at shallower levels in the mountain belt. Thrust faults
form a thrust belt along the edges of the mountain chain within which material is moved
outwards, away from the centre of the orogenic belt.
4.1 Peripheral foreland basins
The thrust belt moves material out on the foreland crust either side of the orogenic belt.
Under this load the crust flexes to form a peripheral foreland basin. The basin width
depends on sediments load and the flexural rigidity of the foreland lithosphere.
Early foreland basin sediments occur in a deep water basin with the rate of subsidence
exceeding the rate of sediments supply. Turbidites are typical of this stage. When the
orogenic belt is more mature there is an increase in sediment supply to the foreland basin.
Foreland basin stratigraphy typically shallows up from deep water to shallow marine
followed by continental sedimentation which dominates the later stages of foreland basin
sedimentation.
Foreland basin stratigraphy is often complicated by later thrusting deformation. These
thrusts may subdivide the basin into piggy-back basins which lie on top of the thrust sheets
and foredeep basins which lie in front of all the thrusts.
EXAMPLE OF A PERIPHERAL FORELAND BASIN
The Apennines are a mountain chain along the axis bf Italy (Fig. 12). The Adriatic Sea is
the foreland area to the north-east thrust movement of the Apennines belt. The northern
extension of the Adriatic foreland basin is the Po Basin in northern Italy. The Adriatic
foredeep has been an area of turbidite deposition since the Oligocene. The Po Basin has
received sediment from the Alps to the north; fluvial and shallow marine deposition is
currently occurring in this area.

Fig. 12 The Po Basin in northern Italy is an example of a

peripheral foreland basin.
5 Basins related to strike-slip plate boundaries
If the relative plate motion purely parallel to that line there would be neither uplift nor
subsidence along strike-slip plate boundaries. In fact, such plate boundaries are not
straight, the motion is not purely parallel, and they consist not of a single fault strand but of
a network of branching and overlapping individual faults. Zones of localized subsidence
and uplift create topographic depressions for sediment to accumulate and the source areas
to supply them.
5.1 Strike-slip basins
These are generally termed transtensional
basins and are formed by a number of
mechanisms (Fig. 13). The overlap of faults
can create regions of extension between
them known as pull- apart basins. Such
basins are typically rectangular or rhombic
in plan with widths and lengths of only a few
kilometers or tens of kilometers. They are
unusually deep, especially compared to rift
basins. Where there is a branching of faults
a zone of extension exists between the two
branches forming a basin. The curvature of
a single fault results in bends which are Fig. 13 Basins formed at strike- slip plate boundaries
and in regions of intra-continental strike-slip tectonics.
either restraining (locally compressive) or
releasing (locally extensional form elliptical
small basins).
The mechanism of basins development in strike- slip belts is variable but there are a
number of common characteristics. They are relatively small, usually in the range of a
hundred to a thousand square kilometers, and often contain thick successions. Subsidence is
usually rapid and several kilometers of strata can accumulate in a few million years.
Typically the margins are sites of deposition of coarse facies (alluvial fans and fan deltas)
and these pass laterally over very short distances to lacustrine sediments in continental
settings or marine deposits.
EXAMPLE OF A STRIKE-SLIP BASIN
The Dead-Sea is a basin bound on both sides by
steep strike-slip faults (Fig. 14). Sediment is supplied
to the Dead Sea Basin by the Jordan River, by
aeolian input and by alluvial fans which have
formed on both flanks of the transtensional basin
from material shed from the high ground either side.
The Dead Sea is highly saline and 400 m below the
level of the Mediterranean Sea.

Fig. 14 The region extending from the Dead Sea to the Gulf of Aqaba
is a zone of strike-slip tectonics and the formation of sedimentary
basins.
6. Complex and hybrid basins
These are basins developed due to the interaction of more than one tectonic regime. This
most commonly occurs where there is a strike-slip component to the motion at a
convergent or divergent plate boundary. Such situations exist because plate motions are
commonly not simply orthogonal or parallel and there is oblique convergence or extension
between plates in many parts of the world.
EXAMPLE OF A HYBRID BASIN
A modern example of a basin which is controlled by
both extensional and strike tectonics is the Gulf of
California (Fig. 15). This basin lies at the southern
end of the San Andreas strike-slip fault zone. Thus,
it falls in a zone of transition from a strike-slip plate
boundary to an extensional, oceanic basin. The
overall basin characteristics are a mixture of a pull-
apart basin and proto-oceanic trough.

Fig. 15 The Gulf of California is considered to be

a hybrid basin, controlled by a combination of
extensional and strike-slip tectonics.
Comprehensive Illustration of Plate Tectonic
Behind Development of Sedimentary Basins
7. Changes in tectonic setting with time
Rift basins form and evolve into proto-oceanic
troughs and eventually into ocean basins
bordered by passive margins. After a period of
tens to hundreds of millions of years the ocean
basin starts to close, with subduction zones
around the margins consuming oceanic crust.
Final closure of the ocean results in continental
collision and the formation of an orogenic belt.
These patterns of plate movement through time
are known as the Wilson cycle (Fig. 16). The
whole cycle starts again as the continent breaks
up by renewed rifling. The Wilson cycle predicts
that at any place tectonic setting changes through
time.

Fig. 16 The Wilson cycle of extension to form s rift bssin and

ocean basin followed by basin closure and formation of an
orogenic belt.
¾ The stratigraphy of an area must be considered at different scales;
9 At the scale of mapping formations, a succession maybe considered in terms of the plate
tectonic controls and interpreted as the product of a certain sedimentary basin type.
9 At a larger scale of regional mapping of groups, a stratigraphic succession should be
considered as being formed in an evolving tectonic setting. The stratigraphy of an area may
be divided into units formed under different tectonic regimes using the characteristics of
the various basin types.

¾ It is only in the centers of stable continental areas that basins are unchanging over long
periods of geological time. The central part of the Australian continent has not experienced
the tectonic forces of plate maigins for 400 million years.
¾ In regions closer to plate margins basins typically have a life-span of a few tens of
millions of years. The backarc basins in the west Pacific appear to be active for 20 million
years or so. In contrast, passive margins of the Atlantic, have been sites of sedimentation at
the edges of the continents for over 200 million years.