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198 DALRYMPLE

through which the flow passes. For

example, Chesapeake Bay on the east

coast of the United States is microtidal

(tidal range 1 m) but has a very small

mouth relative to its size. Therefore,

maximum, near surface speeds at the

entrance are typically greater than 0.5

mls, and locally exceed 1 ds. Within

the broader inner estuary, current

speeds are much less. Similar tidal

current speeds typify most tide-domi-

nated environments, including such

areas as the macrotidal Bay of Fundy

and the southern North Sea and

English Channel (Fig.

6).

Only very

locally in these areas do speeds

exceed 1.5

mls. Thus, the major influ-

ence of a large tidal range is not to

cause appreciably faster currents, but

to increase the area that experiences

strong tidal influence.

In general, the fastest tidal currents

occur in coastal areas because on-

shore shallowing increases the tidal

range and decreases the cross-sec-

tional area. Topographic constrictions

associated with tidal inlets and estu-

aries also accentuate the currents.

Strong tidal currents may also occur on

the outer part of the continental shelf if

the offshore increase in the volume of

water flowing through any shore-par-

allel section is not compensated by a

sufficient increase in water depth. A

particularly well-developed example is

given by Georges Bank, at the outer

edge of the shelf between Nova Scotia

and Cape Cod. Tidal current speeds

exceed 0.5

m/s in all areas shallower

than 60

m, with peak values up to

1.2

m/s (Twichell, 1983). In the deeper

water to the north, current speeds are

typically less than 0.2

ds.

Ideally, the speed and duration of the

ebb and flood currents are equal, and

no net transport of water or sediment

occurs. In many coastal areas, how-

ever, deformation of the tidal wave and

the interaction of the currents with the

bottom topography produce inequalities

between the flood and ebb currents. As

a result, many areas experience a

net

residual

transport of sediment in the

direction of the stronger

(dominant)

current. The weaker current which

flows in the opposite direction is termed

the

subordinate

current. Typically,

areas of ebb and flood dominance al-

ternate spatially in a system of mutu-

ally-evasive, flood-dominated and

ebb-dominated channels, producing

complicated patterns of sediment

transport, as in Bay of Fundy estuaries

(Dalrymple

et a/.,

1990).

TIDAL SEDIMENTARY

STRUCTURES

The

periodic

changes in current speed

and direction which characterize tidal

systems produce several sedimentary

structures which are diagnostic of tidal

deposition. A full discussion is given

by Nio and Yang (1991).

Cross bedding containing some type

of regularly spaced, internal disconti-

nuities will be formed in areas where

the maximum speed of the dominant

current is capable of producing dunes.

If the subordinate current is capable of

eroding the lee face of the dunes

formed by the preceding dominant

current,

reactivation surfaces

will be

produced (Fig. 7). Mud drapes may

also be deposited on the lee face

during one or both slack-water periods

if suspended sediment concentrations

are high enough (Figs. 7, 8). Much of

this mud may be in the form of fecal

pellets rather than individual clay

Dominant Current Stoge

DOMINANT CURRENT DIRECTION

"'\\\$Foreset Deposition

---zzzz

Slack Water Stage after the Dominant Stogg

First Mud Drape

-\-

Subordinate Current Stagg

SUBORDINATE CURRENT DIRECTION

ReactivatiOn Surface

Tidal Bundle

Second Mud Drape

Figure

Structures produced in a dune

during

Wal

cyde. In this example, the currents

exhibit a pronounced asymmetry and

suspended

sediment concentrations are moderately

high. A tidal bundle is the deposit of

the

dominant

portion

of the tidal cycle. After Visser

980).

---

11. TIDAL SYSTEMS 199

flakes. Typically the amount of sand

deposited by the subordinate current

is small, and hence the mud drapes

deposited after the dominant and sub-

ordinate tides are closely spaced (first

to second drapes in Figure 7). Mud

drapes may not be deposited if there

is wave action, or if there are rotary

tides that have no slack-water period.

The deposit of a single, dominant tide

is called a

tidal bundle,

whether

bounded by reactivation surfaces or

mud drapes (Fig. 7).

Because of the variations in tidal

current speed associated with the

neap-spring cycle, sequences of tidal

bundles commonly show cyclic varia-

tions in thickness, with thicker bundles

forming during spring tides when cur-

rents are stronger and the dune mi-

grates further. Thinner bundles form

during neap tides (Fig.

9).

If the tides

are semidiurnal, there should be 28

tidal bundles within a complete neap-

spring cycle of 14.77 days (Fig. 1).

There will be fewer than 28 if the

maximum current speeds drop below

the threshold for sand movement

during neap tides (Visser, 1980).

There will be 14 or fewer bundles if the

tides are diurnal. Excellent ancient ex-

amples of tidal bundles have been de-

scribed by Allen (1982) and Allen and

Homewood

984).

The above structures form only with-

in dunes which do not bear smaller,

superimposed dunes. The internal

structure of large compound dunes

(which are commonly called

sand

waves)

is much more complicated, as

discussed below.

Herringbone cross stratification,

which is marked by opposed

(bipolar)

dip directions in adjacent sets of cross

bedding, is widely used as an indication

of tidal deposition. However, bipolar

cross bedding is not universally devel-

oped in tidal sands because the devel-

opment of either flood or ebb domi-

nance commonly produces a unidirec-

tional paleocurrent pattern. Herring-

bone cross bedding is most easily

recognized in planar tabular cross

bedding. In limited outcrops of trough

cross bedding, care must be taken to

avoid misinterpreting a partial end-on

view of troughs.

In many areas, current ripples are the

stable

bedform instead of dunes. The

structures formed in these situations

generally belong to the continuum from

flaser bedding

lenticular bedding

(Fig.

10;

Reineck and Wunderlich,

1968). In the ideal case, a unit of

ripples will be formed by the maximum

currents of the dominant tide, followed

by deposition of a mud layer during the

ensuing slack-water period (Fig. 11).

If the subordinate current is strong

enough, a second rippled unit with an

opposite paleocurrent direction will be

deposited on top of the first mud drape,

followed by another mud drape during

the second slack-water interval. More

commonly, however, the subordinate

current is too weak to deposit sand,

and the two mud drapes are amalga-

mated. Tidal currents which are too

slow to produce ripples may still

deposit thin sand layers from suspen-

sion that alternate with mud laminae.

Beds which show cyclic changes in

layer thickness due to neap-spring

variations in tidal current speed are

termed

tidal rhythmites

(Kvale and

Archer, 1990; Dalrymple et al., 1991;

Williams, 1991).

Figure

Bundled cross bedding in subtidal sands of the Oosterschelde estuary, the Netherlands.

Outcrop photo showing repetitive mud

drapes in the

foresets and bottomsets of two large-scale cross beds. The dark flecks are either organic debris or mud pebbles. Trowel

long. B) Close-up of

foresets showing a complete tidal bundle bounded by mud couplets that enclose subordinate-current ripples. The tidal

bundle is approximately

cm thick.

200

DALRYMPLE

TIDE-DOMINATED ENVIRONMENTS

Two environments are common to

many tide-dominated depositional

systems,

sandy to muddy tidal flats

which fringe the shoreline, and

elongate

tidal sand ridges

or bars, with

their superimposed sand waves, that

occur offshore from the flats, in estu-

arine, deltaic and open shallow marine

settings.

Tidal flats

In areas with a large tidal range, tidal

flats typically rim the margins of back-

barrier lagoons, estuaries and the

open coast. They may be up to sev-

eral km wide, and form flat to gently

sloping areas (typically

<lo)

that

extend from the

supratidal zone,

through the

intertidal zone,

and into

the shallow portion of the

subtidal

zone

(Fig.

12;

see also Fig.

22).

The

outer edge of the flats in back-barrier

and estuarine settings is commonly

bordered by a major tidal channel in

which flow is subparallel to the local

shoreline. The speeds of both the

flood and ebb tidal currents are great-

est in these channels, and decrease

landward. Consequently, the sedi-

Figure

Flaser bedding (A), wavy

("tidal") bedding

(B),

and lenticular bedding

(C). The progression from A to C results

from a net decrease in current speed and

increased deposition and preservation of

mud drapes. After Reineck and Singh

980).

ments adjacent to the channel are typ-

monly contain dune cross bedding in

ically sandy, and pass gradually into

areas with high current speeds, and

muds near the high-tide line. In some

ripple cross lamination in places where

lagoons, the tidal flats pass outward

the current speed is lower (Fig.

12).

into subtidal muds.

The thickness of the cross bed sets

The

sand flats

which occupy the

generally decreases toward the shore.

lower portion of most tidal flats

com-

Parallel lamination may also occur in

Figure

Variation of bundle thickness within a cross bed from the Oosterschelde estuary,

the Netherlands.

neap tides; S

spring tides. The average number of bundles between

successive_neap tides is

27.

The tendency for the thickness of successive bundles to be al-

ternately larger and smaller is due to the diurnal inequality. After Visser (1980).

FLOOD

EBB

Figure

Variation in tidal current speed over a single tidal cycle and the origin of tidal

rhythmites. The tidal currents in this example are of equal speed in both directions. As the

difference in speed between the dominant and subordinate currents increases, the thickness

of the two sand layers will become more unequal. After

Dalrymple

etal.

(1991).

_._-_____-I

11.

TIDAL SYSTEMS

201

the sand flats located in the head-

ward portions of macrotidal estuaries

(Dalrymple et al.,

1990).

Mixed flats,

which mud layers become more abun-

dant as the distance from the channel

increases, lie shoreward of the sand

flats.

Mud flats

consisting of laminated

muds with relatively little sand lie still

further landward (Fig.

12).

Flaser

bedding passing landward into lentic-

ular bedding is typical of the transition

from mixed flats to mud flats. Burrows

belonging to the

Skolithos

ichnofacies

(Chapter

may partially to completely

obliterate the physical structures on

any part of the tidal flat where

sedi-

ment movement is not too intense.

The highest parts of the tidal flats may

become vegetated to produce

salt

marshes

where the stratification is

largely destroyed by rootlets. Salt-

water and freshwater peats can accu-

mulate here. Desiccation cracks are

most abundant in the upper intertidal

and supratidal zones.

Tidal flats along exposed, open

coasts exhibit the landward-fining

trend described above, but because of

greater wave action they may be

coarser grained, and contain more

wave-generated structures, than tidal

flats in protected areas (Thompson,

1968;

Larsonneur,

1975).

During

periods of reduced sediment supply,

wave erosion concentrates the sand

and shells into beach ridges or swash

bars that pass seaward into a thin ero-

sional lag.

In tropical climates, sea grasses and

mangroves commonly colonize large

parts of the tidal flats

(Gostin

al.,

1984).

The carbonate content is gen-

erally higher than in temperate tidal

flats, and evaporites may form in the

upper intertidal and supratidal zones if

the climate is arid. By contrast, tidal

flats in mid to high latitudes commonly

contain ice-rafted debris and

soft-sedi-

ment deformation produced by the

grounding of ice blocks.

The muddy parts of the intertidal flats

tide level-

Figure

Block diagram of a typical siliciclastic tidal flat. The tidal flats fine toward the high-tide level, passing gradationally from sand flats,

through mixed flats, to mud flats and salt marshes. An example of the upward-fining succession produced by tidal flat progradation is shown

in the upper left corner. The sedimentary structures reflect a landward decrease in tidal-current speeds. The cross bedding deposited on the

lower portion of the sand flats, and in the adjacent tidal channel, is commonly oriented parallel to the local coastline (which parallels the back

edge of the diagram).

202

DALRYMPLE

are dissected by a network of small- to

medium-sized meandering

tidal gullies

that increase in width and depth as they

coalesce seaward (Fig.

12).

The floors

of the gullies commonly contain a lag

of shells and mud

clasts. Sand is only

abundant in the more seaward parts.

The main site of deposition is on the

point bars where lateral-accretion

bedding (also called

inclined heterolithic

stratification)

is developed (Fig.

13;

Bridges and Leeder,

1976;

Thomas

i987).

These

deposits

'Onsist

in-

Figure

Cut bank of tidal creek in Jade Bay, Germany, showing lateral-accretion

terbedded

sand

and

mud

which

may

bedding of an earlier tidal channel. The exposure is approximately

m high.

contain tidal

rhythmites. Burrowing is

generally less intense than on

thead-

jacent flats because deposition is more

rapid. The vertical thickness of the

lateral-accretion sets approximates the

depth of the channel, and ranges from

a few tens of centimetres to many

metres in the main tidal channels. The

dips are typically

5-15",

but may

exceed

25".

The proportion of a tidal

flat which is underlain by point-bar de-

posits varies widely. In some North Sea

tidal flats, lateral-accretion bedding

comprises a high proportion of the tidal

flat deposits, whereas in other areas

(e.g., the Bay of Fundy and Korean

west coast)

gulleys, if present, have

stable positions, and most of the de-

posits are horizontally bedded

(Alexander

et al., 1991;

Dalrymple

a/., 1991).

Progradation of a tidal flat generates

an upward-fining succession (Weirner

Figure

Extensive field of three-dimensional megaripples with well-developed scour pits,

eta/., 1982;

Tewindt,

1988).

The

sue-

on a tidal sand bar in the outer part of the Cobequid Bay-Salmon River estuary. These

cessions typically begin with an ero- dunes have a wavelength of approximately

m. Note the late-ebb stage current ripples that

sional base that is scoured by tidal

cover the dunes. Similar dunes occur on tidal sand ridges in deltaic and shelf environments.

Figure

High-resolution seismic-reflection profiles and interpreted sections through flood-asymmetric sand waves from the northern

French coast. The vertical exaggeration is

x4, and maximum lee-face slopes are approximately

20°.

All of the stratification surfaces shown are

master bedding planes, and the structures belong to classes

IV or V of Allen

(1980;

Fig.

16).

From Bern6

a/.

988).

11. TIDAL SYSTEMS 203

channels during a local transgression.

Above this there is a gradual upward

decrease in the grain size and thick-

ness of sand beds, and an increase in

the proportion of mud (Fig. 12). Com-

monly encountered structures in the

sands of the

subtidal and lower inter-

tidal zone include mud-draped foresets,

reactivation surfaces and local herring-

bone cross bedding. Variable amounts

of wave-generated structures, perhaps

including hummocky cross stratification,

are present in open coast settings. The

intertidal mud flats contain abundant

flaser and lenticular bedding, and ero-

sionally based tidal

gulley sediments

with lateral-accretion bedding. Rooted

horizons, coals

and/or paleosols occur

in the capping salt marsh. Bioturbation

may range from minimal to intense in

any facies. The thickness of such suc-

cessions ranges from about 1 m to

more than 10 m. Some workers have

attempted to estimate the tidal range by

determining the sediment thickness be-

tween the interpreted levels of high and

low tide (Fig. 12; Terwindt,

1988), but

such attempts have met with limited

success.

Tidal sand waves and sand ridges

Sand waves and sand ridges are

common in shallow

subtidal areas such

1) the axial portions of tide-domi-

nated estuaries, 2) distributary mouth

and subaqueous delta plain regions of

tide-dominated deltas, and

large

parts of tide-dominated shelves. These

environments commonly experience

tidal currents with speeds of 0.5-1.5

mls, and the sediments consist pre-

dominantly of sand. As a result, dunes

of various types (megaripples and sand

waves of previous authors) are widely

developed (Fig. 14). These

bedforms

are commonly superimposed on still

larger features called tidal sand ridges

or bars that are oriented nearly parallel

to the main tidal flow.

Tidal sand waves

Large-scale, compound dunes (i.e.,

sand waves) are particularly abundant

in sandy

subtidal settings. These bed-

forms range in height from about 1 m to

more than 20 m, and have wavelengths

of 10 m to several hundred metres

(McCave, 1971

;

Belderson

et al.,

1982).

In general, the larger

bedforms occur in

deeper water. Because of the

wide-

IIA

spread dominance of flow in one direc-

tion, sand waves are typically asym-

metrical (Fig.

15), with their steeper

---

(lee1 face inclined in the direction of net

IIIA

Figure

Theoretical models of the structures within tidal sand waves as proposed by

Allen (1980).

mud drape;

bioturbation. The letters

refer to the hierarchy of erosion

surfaces;

surfaces are the master bedding planes. The hydrodynamic interpretation of

the various structures provided by Allen (1980) has been questioned (Dalrymple,

1984), but

the structures depicted are matched qualitatively by many ancient examples.

sediment transport and

bedform migra-

tion. Lee face inclinations vary from the

angle of repose

(32-35") to nearly zero,

but are typically about

(Belderson

al.,

1982). Allen (1980) has produced a

set of theoretical models of the internal

structure of large sand waves (Fig. 16).

Two examples are shown in Figures 17

and 18. At one extreme, the deposits

consist of high-angle cross stratification

(dips

>25-30"; Fig. 16

and II) in which

there are few erosional breaks (Fig. 17).

This structure most likely forms where

the subordinate current is below the

threshold of sediment movement (Allen,

1980), and where any superimposed

dunes are small (Dalrymple, 1984). As

the power of the subordinate and domi-

nant currents becomes more equal, or

as the superimposed dunes become

larger, the number of erosional surfaces

(the

master bedding planes

of Allen,

1980) increases and their angle of dip

decreases, until the deposit consists

entirely of

compound cross bedding

(smaller-scale cross beds separated by

inclined set boundaries; Figs. 16

and

VI, 18) that is produced as smaller

dunes migrate down (or up) the lee

face of the sand wave. Note that bipolar

cross stratification becomes more

pre-

DALRYMPLE

valent as the inclination of the master

bedding planes decreases. The abun-

dance of mud drapes, mud clasts and

bioturbation varies markedly between

examples. At one extreme, mud and

bioturbation may be entirely absent,

but in other cases, the bottomsets and

lower part of the lee face contain signif-

icant amounts of mud. Such deposits

commonly show a concave-upward

geometry and have been termed

sig-

moidal cross stratification

(Mutti

et a/.,

1 985).

High-resolution, seismic-reflection

profiles through modern sand waves

confirm the general features of Allen's

models (Fig. 15; Berne

et al.,

1988).

The reflectors in these profiles have

dips of less than

15" (do not be fooled

by the vertical exaggeration!) and rep-

resent master bedding planes. Cores

show that the structures between

these planes consist of smaller cross

beds.

Tidal sand ridges

Sand ridges cover thousands of

square kilometres in tide-dominated

shallow-marine settings (Fig. 19) and

in the delta front region of tide-domi-

nated deltas. Morphologically similar

features also occur in estuaries where

they are called tidal sand bars. These

features have spacings of 1-30

km,

lengths of 10-120 km, and heights of

7.5-40 m (Houbolt, 1968; Swift, 1975;

Stride

et al.,

1982; Twichell, 1983).

They are nearly flow parallel, but are

always oriented at a small angle

(<20°)

to the maximum tidal currents. The

local, net sediment movement is in op-

posite directions on either side of the

crest (Fig. 20). The ridges are gener-

ally asymmetrical, with

steeper (lee)

side facing in the direction of regional

sediment transport. Shelf ridges which

have been examined seismically

contain gently-inclined reflectors which

lie parallel to this "steepn face (Figs.

20, 21; Houbolt, 1968; Johnson

eta/.,

1982; Yang and Sun, 1988). Dips of

these lateral-accretion surfaces range

from 5-10" in ridges which are pres-

ently active, to less than

2" in moribund

(inactive) ridges. No published informa-

tion exists on the internal architecture

of modern estuarine and delta mouth

ridges.

Active ridges in all settings are

covered by sand waves and

megarip-

ples, and the limited core information in-

dicates that these ridges contain cross

bedding of various types (Houbolt,

1968;

Dalrymple

et al.,

1990). Shelf

ridges which have become moribund

due to the postglacial sea level rise lack

large-scale bedforms. They may be

draped by strata containing ripple cross

lamination

andlor storm deposits with

greater degrees of burrowing. These

sediments may also contain larger

amounts of mud than the deposits of

active ridges. The vertical succession of

grain sizes in ridges may either coarsen

or fine upward, depending on the depo-

sitional setting and sea level history.

TIDE-DOMINATED

DEPOSITIONAL SYSTEMS

It is generally not appreciated that the

new concepts of allostratigraphy and

sequence stratigraphy discussed in

Chapter 1 were primarily developed in

wave-dominated

settings. Few studies

have examined how

tidal

systems res-

pond to sea level change. As a result,

the material in the following sections on

the stratigraphy of tide-dominated

systems contains a certain amount of

"educated speculation" that is based to

varying degrees on documented ex-

amples.

Figure

Large-scale cross bedding in the Lower Greensand (L. Cretaceous) near

Leighton

Buzzard, England. The large set at the base is about

thick. It is overlain by

smaller sets which interfinger with the major foresets, indicating that the smaller sets were

produced by superimposed bedforms. Many of the

foreset beds are bioturbated. This

example would correspond to structure class

II or Ill of Figure 16.

Figure

Low-angle

(5"),

compound cross bedding in the Lower Greensand (L.

Cretaceous) near Leighton Buzzard, England. Regional bedding is horizontal. This example

corresponds with class

IV or

of Figure 16.

wet and cold Roger Walker (1.6 m) for scale.

TIDAL SYSTEMS 205

Tidal systems are potentially more

sensitive to sea level change than

wave-dominated systems, because

tidal resonance, which favours tidal

sedimentation, is a sensitive function

of basin geometry. Basin geometry, in

turn, may be dramatically altered by

small changes in sea level. Thus, tide-

dominated conditions may be turned

on and off in a "geological instant"

because of sea level rise or fall.

Transgressive systems tracts

During a lowstand of sea level, river

valleys commonly become incised. With

the succeeding transgression, the lower

reaches of these valleys are flooded by

saltwater and experience tidal action.

These areas are termed estuaries. They

trap the sediment supplied by the rivets,

and consequently little sediment

reaches the shelf. As a result,

pre-ex-

isting material, including estuarine

facies which are abandoned on the

shelf as transgression continues, may

be eroded from the

seafloor by strong

tidal currents. The mud component

commonly moves offshore into deeper

water, while the sands are either trans-

ported onshore or left in place on the

Figure

Distribution of tidal sand ridges off Norfolk, England, in the southern North Sea.

Depths in fathoms (1 fathom

6 feet, or roughly

m). The ridges are 35-40 m high, with

their crests in 10-15 m of water. The steepest ridge flank

(5")

is on the northeastern side of

each ridge because the strongest tidal currents flow to the north. From Swift (1975) after

Houbolt (1968).

shelf as the water deepens.

Tide-dominated estuaries

The Cobequid Bay-Salmon River

system in the Bay of Fundy (Dalrymple

a/., 1990, 1991) shows the typical

facies of a tide-dominated estuary

(Figs. 22, 23). A similar situation is pre-

sented by the Ord River in Australia

(Wright et

a/., 1975) and the Severn

River, England (Allen, 1990). The pat-

terns of sedimentation in these estu-

aries, and differences between tide-

dominated and wave-dominated estu-

aries are discussed by Dalrymple et

a/.

(1 992). Wave-dominated estuaries are

described in Chapter 10. It is worth

noting that the morphological distinc-

tion between tide-dominated estuaries

and deltas is not well documented,

and that many so-called

tide-domi-

nated deltas (e.g., the Ord River) may

be better classified as estuaries (see

discussions in Chapters 1 and 9).

Tide-dominated estuaries receive

sediment both from the river at the

head of the estuary, and from the ad-

jacent shelf by tidal currents. As a

result, grain sizes are coarsest at the

mouth and head of the estuary. Unlike

wave-dominated estuaries, however,

no fine-grained lagoonal facies are

present, because tidal currents pene-

trate into the estuary much more

easily than waves do (Dalrymple

etal.,

1992). Consequently, the distinct tri-

partite facies distribution (sandy barrier

muddy lagoon

sandy bay-head

delta; see Chapter 10) of

wave-domi-

nated estuaries does not occur in tide-

dominated systems. Instead, tide-

dominated estuaries are bordered

by muddy tidal flats and marshes

because tidal currents are strongest in

the sandy channels that run along the

entire length of the estuary axis (Fig.

22). The axial sands are finest grained

and contain the greatest number of

mud drapes in the vicinity of the

tidal-

fluvial transition, where the suspended

sediment turbidity maximum is situated.

The most seaward, estuarine facies

consists of elongate sand bars

(Fig.

22) which are separated by ebb-domi-

nant and flood-dominant channels

(Dalrymple et

a/.,

1990). The bars are

covered by dunes of various types and

sizes, and the bar sediments are com-

posed primarily of cross bedded,

medium to coarse sand. Lateral shift-

ing of the bars produces an

upward-

206 DALRYMPLE

fining trend (Fig. 23) because current

speeds are greatest in the channels

and decrease toward the bar crests.

On average, the tidal current speeds

increase into the estuary, so that the

area

headward of the bars is charac-

terized by extensive sand flats with

parallel lamination. Still farther

head-

ward, these flats grade into the tidally

influenced portion of the river channel

(Fig. 22). Parallel lamination is present

here, but cross bedding containing

mud drapes may become abundant

headward. Tidal rhythmites occur in

the bordering mud flats (Dalrymple et

a/., 1991), and inclined heterolithic

stratification may be present in the

meanders that occur in this region. At

the tidal limit, which can be situated

several hundred kilometres inland, the

muddy channel sands of the inner

estuary pass gradationally into coarser

fluvial sediments.

The Gironde estuary (France) is

commonly cited as a typical tkk-domi-

Figure

Model of sediment transport around, and internal structure within, a tidal sand

nated estuary, but it has many of the

ridge. The peak tidal currents flow at a slightly oblique angle to the ridge crest so that local,

attributes of a wave-dominated estu-

net transport is in opposite directions on either side of the crest. In most cases, net accretion

ary, including a tripartite

facies distri-

occurs on the steeper (regionally downstream) side of the ridge. Thus most of the small-

bution (Allen, 1991). This estuary con-

scale cross bedding in the ridge should be oriented in the opposite direction to the regional

tains tidal sand ridges, but these rep-

transport. After Stride

eta/.

982).

resent a tide-dominated bay-head

delta within the inner estuary. As a

285" Sea level

105"

result, these ridges coarsen upward

Well Bank

because they prograde seaward over

central estuary muds (Allen, 1991).

---KT\

\\\

During active transgression, the

,,/--

----

facies within a tide-dominated estuary

shift headward. This appears to be the

case in the Severn estuary today

(Allen, 1990). In the process, migrating

tidal channels coupled with wave

action erode all or part of the more

headward facies, leaving an incom-

plete record of the estuary in shallow

marine areas that lie seaward of the

final shoreline. In the ideal case of es-

sentially complete preservation, the

transgressive valley-fill sediments will

fine upwards from

fluvial sands andlor

gravels to interbedded sands and

muds deposited in the tidal-fluvial tran-

sition of the inner estuary (Fig. 24).

This is followed by an overall upward

coarsening into fine-grained,

parallel-

Figure

Seismic-reflection profiles

through tidal sand ridges from the North

Sea; see Figure

for ridge locations. The

vertical exaggeration is approximately

times; none of the inclined reflectors is

steeper than

5".

From Houbolt (1968).

215" Sea level 35"

Well Bank

160"

280"

340' (Sea level

Well Bank

235'

5or

lovet

55"

Sm~th

Knoll

A~--

---

...........

.......

11.

TIDAL SYSTEMS 207

current speeds are less than about 50

d, medium The shallow marine area surrounding

cm/s (Fig. 26), rippled sand sheets

bars. These the British

Isles is the classic example and isolated sand patches are devel-

Transgressive tidal sand sheets are

sand

of the sand flats, and by mixed flat,

m/s (Stride

a/.,

1982). Even farther fine sand and mud. The relative influ-

mud

flat and marsh deposits (Fig. 23). down the transport path, where tidal ence of storms increases as the tidal

FACIES

ZONES

1-3-

-24

Figure

Facies distribution in the Cobequid Bay

Salmon River macrotidal estuary, Bay of Fundy. Facies zone

elongate sand bars

(medium to coarse sand); zone

upper-flow-regime sand flats (fine sand); and zone

tidal-fluvial transition. The erosional foreshores bor-

dering zone

are unique to Cobequid Bay. Sand bars in estuaries such as the Severn River are bordered

mud flats and salt marshes.