Gary Nichols. Sedimentology and Stratigraphy(Second Edition)

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Lift component

(Bernoulli effect)

Fluid force

Drag component

(Frictional drag)

Flow

Gravity

Fig. 4.4 The lift force resulting from the

Bernoulli effect causes grains to be

moved up from the base of the flow.

0.1

1000

100

Flow velocity (cm s

1

)

0.001 0.01 0.1 1 10 100

Grain size (mm)

Clay Silt VF S F S M S C S VC S Gran. Pebbles Cobbles Boulders

Curves are approximate

for flow depth of 1 metre.

The positions of the

curves vary for different

flow depths and different

sediment characteristics

Erosion of unconsolidated mud

rosion of sand

sitio

1000

EROSION AND TRANSPORT

TRANSPORT IN SUSPENSION

DEPOSITION

TRANSPORT AS BEDLOAD

Fig. 4.5 The Hju

lstrom diagram shows the relationship between the velocity of a water flow and the transport of loose grains.

Once a grain has settled it requires more energy to start it moving than a grain that is already in motion. The cohesive

properties of clay particles mean that fine-grained sediments require relatively high velocities to re-erode them once they are

deposited, especially once they are compacted. (From Press & Siever 1986.)

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velocity and particles that are already in motion.

This shows that a pebble will come to rest at around

20 to 30 cm s

1

, a medium sand grain at 2 to 3 cm s

1

and a clay particle when the flow velocity is effect-

ively zero. The grain size of the particles in a flow

therefore can be used as an indicator of the velocity

at the time of deposition of the sediment if depos-

ited as isolated particles. The upper, curved line

shows the flow velocity required to move a particle

from rest. On the right half of the graph this line

parallels the first but at any given grain size the

velocity required to initiate motion is higher than

that to keep a particle moving. On the left side of the

diagram, there is a sharp divergence of the lines:

counter-intuitively, the smaller particles require a

higher velocity to move them below coarse silt size.

This is due to the properties of clay minerals that

will dominate the fine fraction in a sediment. Clay

minerals are cohesive (2.4.5) and once they are

deposited they tend to stick together making it diffi-

cult to entrain them in a flow. Note that there are

two lines for cohesive material. ‘Unconsolidated’ mud

has settled but remains a sticky, plastic material.

‘Consolidated’ mud has had much more water

expelled from it and is rigid.

The behaviour of fine particles in a flow as indicated

by the Hju

lstrom diagram has important conse-

quences for deposition in natural depositional envi-

ronments. Were it not for this behaviour, clay would

be eroded in all conditions except standing water, but

mud can accumulate in any setting where the flow

stops for long enough for the clay particles to be

deposited: resumption of flow does not re-entrain the

deposited clay unless the velocity is relatively high.

Alternations of mud and sand deposition are seen in

environments where flow is intermittent, such as tidal

settings (11.2).

4.2.5 Clast-size variations: graded bedding

The grain size in a bed is usually variable (2.5) and

may show a pattern of an overall decrease in grain

size from base to top, known as normal grading,ora

pattern of increase in average size from base to top,

called reverse grading (Fig. 4.6). Normal grading is

the more commonly observed pattern and can result

from the settling of particles out of suspension or as a

consequence of a decrease in flow strength through

time.

The settling velocity of particles in a fluid is deter-

mined by the size of the particle, the difference in the

density between the particle and the fluid, and the

fluid viscosity. The relationship, known as Stokes

Law, can be expressed in an equation:

V ¼ g  D

 (r

 r

)=18m

where V is the terminal settling velocity, D is the grain

diameter, (r

 r

) is the difference between the den-

sity of the particle (r

) and the density of the fluid (r

)

and m is the fluid viscosity; g is the acceleration due to

gravity. One of the implications of this for sedimentary

processes is that larger diameter clasts reach higher

velocities and therefore grading of particles results

from sediment falling out of suspension in standing

water. Stokes Law only accurately predicts the set-

tling velocity of small grains (fine sand or less)

because turbulence created by the drag of larger

grains falling through the fluid reduces the velocity.

The shape of the particle is also a factor because the

drag effect is greater for plate-like clasts and they

therefore fall more slowly. It is for this reason that

mica grains are commonly found concentrated at the

tops of bed because they settle more slowly than

quartz and other grains of equivalent mass.

A flow decreasing in velocity from 20 cm s

1

1cm s

1

will initially deposit coarse sand but will

progressively deposit medium and fine sand as the

velocity drops. The sand bed formed from this decel-

erating flow will be normally graded, showing a

reduction in grain size from coarse at the bottom to

fine at the top. Conversely, an increase in flow veloc-

ity through time may result in an increase in grain

size up through a bed, reverse grading, but flows that

gradually increase in strength through time to pro-

duce reverse grading are less frequent. Grading can

occur in a wide variety of depositional settings: nor-

mal grading is an important characteristic of many

turbidity current deposits (4.5.2 ), but may also result

from storms on continental shelves (14.2.1), over-

bank flooding in fluvial environments (9.3) and in

delta-top settings (12.3.1).

It is useful to draw a distinction between grading

that is a trend in grain size within a single bed and

trends in grain size that occur through a number of

beds. A pattern of several beds that start with a coarse

clast size in the lowest bed and finer material in the

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highest is considered to be fining-upward. The

reverse pattern with the coarsest bed at the top is a

coarsening-upward succession (Fig. 4.6). Note that

there can be circumstances where individual beds are

normally graded but are in a coarsening-up succes-

sion of beds.

4.2.6 Fluid density and particle size

A second important implication of Stokes Law is that

the forces acting on a grain are a function of the

viscosity and density of the fluid medium as well as

the mass of the particle. A clast falling through air will

travel faster than if it was falling through water

because the density contrast between particle and

fluid is greater and the fluid viscosity is lower.

Furthermore, higher viscosity fluids exert greater

drag and lift forces for a given flow velocity. Water

flows are able to transport clasts as large as boulders

at the velocities recorded in rivers, but even at the

very high wind strengths of storms the largest rock

and mineral particles carried are likely to be around a

millimetre. This limitation to the particle size carried

by air is one of the criteria that may be used to

distinguish material deposited by water from that

transported and deposited by wind. Higher viscosity

fluids such as ice and debris flows (dense slurries of

sediment and water) can transport boulders metres or

tens of metres across.

4.3 FLOWS, SEDIMENT AND

BEDFORMS

A bedform is a morphological feature formed by the

interaction between a flow and cohesionless 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 air-

flow. 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 pro-

vides information about the strength of the current,

the flow depth and the direction of sediment transport.

To explain how bedforms are generated some

further consideration of fluid dynamics is required (a

comprehensive account can be found in Leeder

1999). 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 (Fig. 4.7). 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 sub-

layer 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.

Normal grading

in a bed

Reverse grading

in a bed

Fining-up of a series

of beds

Coarsening-up of a

series of beds

Fig. 4.6 Normal and reverse grading within indivi-

dual beds and fining-up and coarsening-up patterns in a

series of beds.

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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: these are

considered in more detail in Chapter 8.

4.3.1 Current ripples

Flow within the viscous sublayer is subject to irregu-

larities 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

(Fig. 4.8). 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 (Fig. 4.8). 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 (the Ber-

noulli effect, 4.2.3) and the sediment transport rate

is reduced, resulting in deposition on the lee side of

the step.

Current ripples (Figs 4.9 & 4.10) are small bed-

forms formed by the effects of boundary layer separa-

tion on a bed of sand (Baas 1999). The small cluster of

grains grows to form the crest of a ripple and separa-

tion 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 down-

stream 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 (Fig. 4.9).

When viewed from above current ripples show a

variety of forms (Fig. 4.11). They may have relatively

continuous straight to sinuous crests (straight rip-

ples or sinuous ripples) or form a pattern of uncon-

nected arcuate forms called linguoid ripples. The

Viscous

sublayer

Boundary

layer

Smooth boundary:

Thick viscous sublayer (low velocities)

and/or small grain diameters

Viscous

sublayer

Boundary

layer

Rough boundary:

Thin viscous sublayer (high velocities)

and/or large grain diameters

Flow depth

Fig. 4.7 Layers within a flow and flow surface roughness:

the viscous sublayer, the boundary layer within the flow and

the flow depth.

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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 (Baas

1994). Straight and linguoid ripple crests create differ-

ent 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. Sinu-

ous 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-lamina-

tion (Fig. 4.9).

Creating and preserving cross-lamination

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 depos-

ited 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

Attachment point

(increased turbulence and erosion)

Separation point

Flow lines

1. Erosion in the trough of a bedform

Flow lines

2. Development of counter-currents in lee of bedform

Roller vortex

(separation eddy with

counter-current flow)

Fig. 4.8 Flow over a bedform: imaginary

streamlines within the flow illustrate the

separation of the flow at the brink of the

bedform and the attachment point where

the streamline meets the bed surface,

where there is increased turbulence and

erosion. A separation eddy may form in

the lee of the bedform and produce a

minor counter-current (reverse) flow.

Fig. 4.9 Current ripple cross-lamination in fine sandstone:

the ripples migrated from right to left. The coin is 20 mm

in diameter.

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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 (Allen 1972)

(Fig. 4.12). 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

(see above). 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 inde-

pendent of water depth and current ripples may form

in waters ranging from a few centimetres to kilo-

metres deep. This is in contrast to most other subaqu-

eous 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

Current

direction

Sinuous crested and isolated ripples

Trough cross-lamination

Current

direction

Straight crested ripples

Planar cross-lamination

Fig. 4.10 Migrating straight crested ripples form planar

cross-lamination. Sinuous or isolated (linguoid or lunate)

ripples produce trough cross-lamination. (From Tucker

1991.)

Flow direction

Straight Sinuous Isolated (linguoid)

Fig. 4.11 In plan view current ripples may have straight,

sinuous or isolated crests.

low angle of climb

moderate rate of

deposition

high angle of climb

high rate of

deposition

Fig. 4.12 Climbing ripples: in the lower part of the figure,

more of the stoss side of the ripple is preserved, resulting

in a steeper ‘angle of climb’.

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distances) range up to 500 mm (Leeder 1999). 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 (Leeder 1999). 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.

4.3.2 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 (Fig. 4.13): the term ‘megaripples’ is also some-

times used, although this term fails to emphasise the

fundamental hydrodynamic distinctions between rip-

ple and dune bedforms. Evidence that these larger

bedforms are not simply large ripples comes from

measurement of the heights and wavelengths of all

bedforms (Fig. 4.14). 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 tur-

bulence 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 gen-

erally 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 (Allen

1982; Leeder 1999).

Dunes and 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

(Figs 4.15 & 4.16). 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 (Fig. 4.17).

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 (Fig. 4.17). 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 (Fig. 4.18). Cross-beds

may form a sharp angle at the base of the avalanche

slope or may be asymptotic (tangential) to the hori-

zontal (Fig. 4.17). At high flow velocities the roller

Fig. 4.13 Dune bedforms in an estuary:

the most recent flow was from left to right

and the upstream side of the dunes is

covered with current ripples.

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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

(Fig. 4.17).

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 undu-

lating 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 (Fig. 4.15) 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

Wavelength

20%

15%

1cm 1m

10%

Ripples

100m10m10cm 1000m

Subaqueous dunes

Bedform height

0.1mm

20%

15%

1mm 10m

10%

Ripples

10cm1cm 1m

Subaqueous dunes

Fig. 4.14 Graphs of subaqueous ripple and subaqueous

dune bedform wavelengths and heights showing the absence

of overlap between ripple and dune-scale bedforms. (From

Collinson et al. 2006.)

Current

direction

Sinuous crested and isolated dunes

Trough cross-bedding

Current

direction

Straight crested dunes

Planar cross-bedding

Fig. 4.15 Migrating straight crested dune bedforms form

planar cross-bedding. Sinuous or isolated (linguoid or lunate)

dune bedforms produce trough cross-bedding. (From Tucker

1991.)

Fig. 4.16 Subaqueous dune bedforms in a braided river.

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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 (Leeder 1999). 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. Dune-

scale 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 rela-

tively strong, sustained flows.

4.3.3 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 (Bridge

2003). Bars can be made up of sandy sediment, grav-

elly 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 conglom-

erates. 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: Collinson et al. 2006) and their

shape (9.2).

4.3.4 Plane bedding and planar lamination

Horizontal layering in sands deposited from a flow is

referred to as plane bedding in sediments and pro-

duces a sedimentary structure called planar lamina-

tion in sedimentary rocks. As noted above, 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 for-

mation 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 – 4.2.4), 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 veloc-

ities in very fine- to coarse-grained sands: ripple and

dune bedforms become washed out with an increase

Flow

counter-current ripples

sharp, angular contact

curved, asymptotic contact

Fig. 4.17 The patterns of cross-beds are determined by

the shape of the bedforms resulting from different flow

conditions.

Fig. 4.18 Planar tabular cross-stratification with tangential

bases to the cross-beds (the scale bar is in inches and is

100 mm long).

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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 (Bridge 1978)

(Fig. 4.19). The bed surface is also marked by elongate

ridges a few grain diameters high separated by fur-

rows oriented parallel to the flow direction. This fea-

ture is referred to as primary current lineation

(often abbreviated to ‘pcl’) and it is formed by sweeps

within the viscous sublayer (Fig. 4.7) that push grains

aside to form ridges a few grains high which lie par-

allel 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 coar-

ser 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.

4.3.5 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:

Fr ¼ y=

g  h

The Froude number can be considered to be a ratio of

the flow velocity to the velocity of a wave in the flow

(Leeder 1999). 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 bed-

form surface parallel to the standing wave, and as the

flow steepens sediment accumulates on the upstream

side of the bedform. These bedforms are called anti-

dunes, 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 (17.2.3), where high velocity flow is accom-

panied by very high rates of sedimentation (Schminke

et al. 1973).

4.3.6 Bedform stability diagram

The relationship between the grain size of the sedi-

ment and the flow velocity is summarised on

Fig. 4.20. This bedform stability diagram indic-

ates the bedform that will occur for a given grain

size and velocity and has been constructed from

experimental data (modified from Southard 1991,

and Allen 1997). It should be noted that the upper

boundary of the ripple field is sharp, but the other

boundaries between the fields are gradational and

there is an overlap where either of two bedforms

may be stable. Note also that the scales are logarith-

mic on both axes. Two general flow regimes are

recognised: a lower flow regime in which ripples,

dunes and lower plane beds are stable and an upper

flow regime where plane beds and antidunes form.

Flow in the lower flow regime is always subcritical

and the change to supercritical flow lies within the

antidune field.

The fields in the bedform stability diagram in

Fig. 4.20 are for a certain water depth (25 to 40 cm)

Fig. 4.19 Horizontal lamination in sandstone beds.

Flows, Sediment and Bedforms 57

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