Mud Deposition at the Shoreface: Wave and Sediment

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1986

Mud Deposition at the Shoreface: Wave andSediment Dynamics on the Chenier Plain ofLouisiana.George Paul KempLouisiana State University and Agricultural & Mechanical College

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8710568

K em p, G eorge Paul

MUD DEPOSITION AT THE SHOREFACE: WAVE AND SEDIMENT DYNAMICS ON THE CHENIER PLAIN OF LOUISIANA

The Louisiana State University and A gricu ltura l and M echanical Col. Ph.D.

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MUD D E PO SIT IO N AT T H E SH O R E FA C E : W AVE AND SED IM EN T D Y N A M ICS

ONT H E C H E N IE R PLA IN O F L O U ISIA N A

A Dissertation

Submitted to the Graduate Faculty of the Louisiana State University and

Agricultural and Mechanical College in partial fulfillment of the

requirements for the degree of Doctor of Philosophy

in

The Department of Marine Sciences

byG. Paul Kemp

B.S., Cornell University, 1975 M.S., Louisiana State University, 1978

December 1986


©1987

GEORGE PAUL KEMP

All Rights Reserved


A C K N O W L E D G E M E N T S

Support for this project was provided by the Louisiana Sea Grant

University Program, a part of the National Sea Grant University Program

maintained by the National Oceanic and Atmospheric Administration of the

U.S. Department o f Commerce. Additional field support was generously

furnished by the Louisiana Universities Marine Consortium (LUMCON),

and the U.S. Army Corps of Engineers staff at Freshwater Bayou Canal

Lock. .

John Wells introduced me to the mudflat problem and supervised this

research. Dag Nummedal participated actively in the field program, and was

unfailing in his encouragement. William Wiseman, Jr., patiently schook i

me in the art of time-series analysis. Charles Adams, Jr., gave invaluable

assistance with boundary layer theory. James Coleman and James Gosselink

contributed sound advice throughout an unorthodox student career. W.

David Constant assisted ably as a member of the examining committee.

Floyd Demers, Annie McK. Prior, and Gerald McHugh at the Coastal

Studies Institute, and Mike DeTraz of LUMCON, were instrumental in

bringing this project to a successful completion. Celia Harrod and Kerry

Lyle provided much needed assistance with the graphics. Scott Dinnel, Ivor

LI. van Heerden, Lauro Calliari, and Allison Drew deserve recognition

among the many students who contributed.

A special thanks is extended to my understanding colleagues at

Groundwater Technology, Inc., to the Joseph Lipsey, Sr., Scholarship award

committee, and to Geraldine Newman at the Department of Marine Sciences.

The greatest measures of appreciation go to my parents, and to Linda

Fowler, my partner in all endeavors.

ii


TABLE OF CONTENTS

PageACKNOW LEDGEM ENTS...................................................................... ii

LIST OF TABLES.................................................................................... v

LIST OF FIGURES................................................................................... vi

A BSTRACT........................... x

INTRODUCTION...................................................................................... 1

PREVIOUS STUDIES............................................................................... 7

FIELD A REA .............................................................................................. 14

G eology.............................................................................................. 16Inner-Shelf C irculation.................................................................. 18Storms and M odem Mudflat Sedimentation.............................. 20

DATA ACQUISITION............................................................................. 23

Sediment Survey............................................................................. 23Dynam ics Experim ent ........................................................... 25

RESU LTS..................................................................................................... 29

Sediment Survey............................................................................. 29Geomorphology ................................................................... 29M udflat Sedimentation........................................................ 36Mineralogy and Rheology.................................................. 41

Dynamics Experim ent.................................................................... 50Meteorologic Influences on the Tide Record.................. 52Waves and Oscillatory Flow............................................. 52SuspendedSediments............................................................ 61Infragravity W ater Level Changes and

Non-oscillatory Flow................................................ 64


DISCUSSION 80

Formation and Transport of Fluid M ud...................................... 85Kinematics of Fluid Mud Deposition at the Shoreface............... 91Shelf Sediment Dispersal................................................................ 100Future Chenier Plain Progradation............................................... 102Implications for Other Coasts and the Geologic Record 103

CONCLUSIONS.......................................................................................... 106

BIBLIOGRAPHY......................................................................................... I l l

APPEND ICES.............................................................................................. 133

A. Beach Profiles.......................................................................... 133B. Program for Grant-Madsen Boundary Layer Closure 142

V ITA .................................................................................................................. 148

APPROVAL SHEETS............................................................................... 149

iv


LIST OF TABLES

Table . Page

1. Chenier Plain Shoreline Characteristics...................................... 31

2. Clay Mineral Abundances.............................................................. 46

3. Sea and Swell Wave Statistics....................................................... 53

4. Total Suspended Sediment............................................................. 62

5. W ater Level Variance Below 0.005 cps....................................... 67

6. Non-Oscillatory Onshore-Offshore Flow.................................... 70

7. Steady Alongshore Flow............................................................... 73

8. Characteristic Alongshore Velocity Profile Parameters 78

9. Grant-Madsen Boundary Layer Model Results........................... 88

v


LIST OF FIGURES

Figure Page

1. Location of chenier plain study area............................................ 6

2. Map of study area, showing major shoreline featuresand beach survey stations.................................................. 15

3. Cross-section of shoreface at station 6, showingdeployment of sensors during cold frontexperim ent............................................................................ 26

4. Nearshore surficial sediments and bathymetry in thestudy area.............................................................................. 30

5a-b. Muddy shoreface at station 7 (5a); reactivatedchenier beach (station 2) at Cheniere au Tigre (5b)........ 32

6a-b. Time-lapse views of the vegetated mudflat mapped by Morgan et al. (1953) at station 3; eroding marsh scarp, May, 1982 (6a), replaced by clastic washover beach, May, 1985 (6b)......................... 34

7. Bar migration at Tigre Point (station 4)..................................... 35

8. Oblique aerial view of western limb of large mud barformed in Spring, 1982, between stations 7 and 8.......... 37

9. Profile time-series of mudflat accretion and erosionat station 8, February - July, 1981................................... 37

lOa-b. Shoreface mud at station 7; Freshly-deposited fluidmud, Spring 1982 (10a); eroding surface showingdevelopment o f shore-normal flutes 5 monthslater (10b).............................................................................. 38

11. Log of vibracore from marsh now covering mudflatdescribed by Morgan et al. (1953) west of Cheniere au Tigre at station 3............................................ 40

vi


12a-c. Bedding structure in radiograph positives of mudflatsequences............................................................................... 42

13. Texture and bulk density changes with depth in uppermeter of newly deposited m udflat.................................... 43

14. X-ray diffractograms of smears from the <1 jx fractionof mudflat sediments treated as indicated........................ 45

15. Shear stress-strain rate plot from capillaryviscometer measurements on mudflat sediment 48

16. Fluid mud yield strength plotted as a function ofsediment concentration (from McCave 1984)................. 49

17. Winds and tides during cold front experiment,1 0 - 1 3 December, 1982..................................................... 51

18. Representative energy density spectra from 20 minutepressure and flowmeter time-series for pre-frontal, frontal, and post-frontal runs............................................ 55

19a-b. Concurrent spectra, and cross-spectral estimates of coherence and phase for records from run 3; between seaward and landward sensors located 10 m apart (19a), and between near-surface and near-bottom flowmeters (19b)............................. 57

20. The areas of application of the several wave theoriesas a function of H/d and d/L (from Komar 1976)........... 59

21a-c. Relationships between wave parameters determined from the spectra and cross-spectra, and those predicted by shallow water LWT; wavelength (21a), and orbital velocity (21b)................................................... 60

22. Mean suspended sediment concentrations fromrepresentative pre-frontal, frontal, and post-frontal runs............................................................................. 63

vii


23. Relationship between vertical sediment concentrationgradient and the relative heights of sea andsw ell........................................................................................ 65

24. Low frequency water level and onshore-offshoreflowmeter data from frontal run 3.................................... 68

25. Sequence of vertical velocity profiles during onshore-offshore flow event in run 3.............................................. 71

26. Low frequency water level and alongshore flowmeterdata from frontal run 5 ....................................................... 74

27. Characteristic logarithmic alongshore velocityprofiles for frontal runs 4 and 5, and post-frontalrun 8........................................................................................ 79

28. B-L-S model for sorting of cohesive and non-cohesivesediments in the boundary layer of a turbidity current

(from Stow and Bowen 1980)............................................ 82

29. Generalized shoreface geometry, showing slopes usedin determining standing wave node position, and changes in the cross-section affecting crossshore velocities...................................................................... 94

30a-b. Cross-shore trends in low-frequency flow understanding waves (30a), and bed shear stress dueto waves attenuated at a constant H/d ratio of0.2 (30b).................................................................................. 95

A -l. Station 1 profile sequence................................................................ 134

A-2. Station 2 profile.sequence.......................... 135

A-3. Station 3 profile.sequence................................................................ 136

A-4. Station 4 profile sequence................................................................ 137

A-5. Station 5 profile.sequence............................................................... 138

A-6. Station 6 profile sequence................................................................ 139viii


A-7. Station 7 profile sequence................................................................. 140

A-8. Station 8 profile sequence................................................................. 141

ix


A B STR A C T

Sedimentological data, and high-resolution current velocity time-

series were acquired nearshore along the muddy coast of Louisiana's chenier

plain. Results suggest that the 2 to 10 cm thick, laminated and graded

silt/shell + mud beds which characterize shoreface mudflat sedimentation

form through intermittent dumping of fluid mud suspensions. Cohesive and

non-cohesive sediments are first separated and concentrated by shear within a

wave boundary layer (WBL). Rapid sedimentation, or "freezing", of the

mud fraction occurs when the yield strength of the fluid mud suspension

equals that of the WBL shear stress.

Incident waves were well described by shallow water linear wave

theory, but exhibited the progressive nearshore attenuation and absence of

breaking characteristic of muddy shorelines. An onshore decrease in wave

shear stress results, which is the reverse of that on surf-dominated coasts.

Fluid mud layers, characterized by significant yield strength, are proposed to

form when high concentrations of suspended sediments are introduced

nearshore, but are prevented from depositing by WBL shear stresses close to

the bed. In cross-section, WBL sorting causes a near-bottom discontinuity in

sediment concentration, and the associated vertical yield strength gradient.

Expanded spatially, this point of weakness forms a plane along which the

fluid mud layer can move in response to shear exerted at its upper margin by

an accelerating flow.

Adjustment of the steady alongshore current velocity profile to the

fluid mud layer restricts coastwise sediment transport. Significant cross

shore transport of fluid mud may occur, however, under the influence of

unsteady flows. Such low-frequency flows were found to be associated with

x


shore-amplified water level oscillations which have characteristics of a

standing wave with an antinode at the shoreline. The onshore decreasing

wave shear stress gradient leads to preferential deposition at the shoreface.

Widespread renewal of chenier plain shoreline advance through

mudflat accretion is predicted when the Atchafalaya delta, now forming 100

km to the east in Atchafalaya Bay, emerges on the inner shelf, and the

nearshore supply of fine-grained sediment is dramatically increased.


IN T R O D U C T IO N

The past 15 years have witnessed a growth in knowledge about silt and

clay sedimentation which was recently likened to an "information explosion"

(Stow and Piper 1984). An early casualty of this "explosion" was the belief

that marine muds accumulate only in quiescent environments, isolated from

currents exceeding a few cm-s_l . Nowhere does the "quiescent

environment" paradigm contrast more with what is observed than on a

significant fraction of the world's coastlines where fine-grained sediment is

deposited at the beach face itself. There, mud sedimentation occurs in non-

estuarine settings exposed to high-velocity wave-generated flows, in addition

to significant steady currents (Dolan et al. 1972, Rine and Ginsburg 1985).

In most locations it alternates spatially and temporally with the more familiar

coarse-grained clastic deposition (Delaney 1963, Wells and Coleman 1981a,

b). Elsewhere, it appears to occur only during the high energy conditions

accompanying storms (Morgan et al. 1958, Nair 1976, Martins et al. 1979).

Modem mud coasts can provide guidance in the interpretation of

ancient sequences in which marine shales are observed to merge laterally or

vertically into non-marine deposits without the transitional beach and tidal

facies normally considered indicative of shoreline processes (Rine and

Ginsburg 1985, Walker and Harms 1971,1976; Keulegan and Krumbein

1949, Moore 1929). An understanding of the hydrodynamic and

oceanographic factors which give rise to any muddy shoreface would greatly

improve the utility of a facies model of widespread applicability. Unlike the

dynamics of littoral onshore-offshore clastic sediment transport, which have

been studied for nearly a century (Cornish 1898), those pertaining to

1


2

nearshore mud deposition have been largely overlooked. The investigation

reported here was initiated to, at least partially, close this gap. The overall

goal of this study was to identify and describe the processes by which a coast

incorporates influxes of mud introduced from offshore. More specifically, I

sought to develop a physical model for the onshore translation and deposition

of mud under waves which could explain the distinctive bedding observed in

coastal mudflat cores.

Shoreline mud accumulations are commonly linked by along-isobath

"advective mud streams" (McCave 1972) to river deltas. Such turbid

currents restrict terrigenous sediment transfer from the continental shelf to

deep water, and can deliver silt and clay to nearshore locations hundreds to

thousands of kilometers distant from the original point of introduction (Delft

Hydraulics Laboratory 1963, Drake 1976, Gibbs 1976).

Some of the best known inner-shelf mud streams play a significant role

in coastal progradation by providing a source of unconsolidated sediments

that are captured for onshore transport and deposition. The mud streams of

the Amazon, Yangtze, and Mississippi Rivers are responsible for extensive

mudflat deposition along 1400 km of the Guiana (Allersma 1971), 800 km of

the East China Sea (Meie et al. 1983), and 60 km of the Louisiana chenier

plain coasts (Morgan et al. 1953, Coleman 1966, Wells and Kemp 1981),

respectively.

Littoral mud deposits have also been described flanking the deltas of

rivers that discharge into epicontinental seas, including the Gulf of Bohai,

China (Zenkovich 1967, Meie et al. 1983), the Gulf of Carpentaria, northern

Australia (Rhodes 1982), the Firth of Thames, New Zealand (Woodroffe et

al. 1983), and the Gulf of California (Thompson 1968). Less is known about


3

apparently non-deltaic coastal mudflat deposition reported adjacent to lagoon

systems on the south-west coast of India (Nair 1976) and on the Rio Grande

do Sul coast of southern Brazil (Delaney 1963, Martins et al. 1979).

These sources document littoral mud deposition in all tide ranges and

wave energy regimes. The coastal features formed exhibit a spectrum of

morphologies. Ephemeral mud "arcs", spits and bars with a few square

kilometers of subaerial expression have been described from Louisiana

(Morgan et al. 1958, Wells and Kemp 1981). In contrast, migratory "mud

banks" a hundred times larger have been tracked for decades off the coast of

Surinam and the Guianas (NEDECO 1968, Rine and Ginsburg 1985).

Elsewhere, coastal mud deposits have been designated only with the generic

terms "flat" or "tidal flat", suggesting that features with discrete alongshore

boundaries are not observed.

While our knowledge of the processes important to fine-grained

sedimentation under waves is growing rapidly, lessons learned in the

laboratory about the complexity of this subject have led some to believe that

little might be gained from field research. The singular investigation by

Wells and Coleman (1981a, b) of the relationship between shallow-water

waves and mud suspension on the coast of Surinam suggests, however, that

such work can provide insights which might not be recognized at laboratory

scales.

The eastern margin of Louisiana's chenier plain coast today offers an

ideal location for study of the dynamics of fine-grained sediment deposition

and erosion in a wave-dominated environment (Fig. 1). This area is

presently receiving an influx of mud from the Atchafalaya River, a

Mississippi distributary now building a delta lobe in a bay 100 km to the east


/

4

(van Heerden et al. 1981). In some locations, fine-grained sediments

accumulate near shore under conditions similar to those described by Wells

and Coleman (1981a) in a process which appears to involve deposition from

highly concentrated mud-water suspensions or fluid mud. Nearby, however,

muddy shorelines exposed to the same wind and wave conditions are

erosional, while on others, normal surf zone processes and sandy beach

sedimentation prevail (Wells and Kemp 1981, Kaczorowski and Gemant

1980).

A two-staged approach was adopted to meet the objectives. First, a

survey of offshore bathymetry and sediment composition was combined with

a 2.5 year beach profiling program to supply regional coverage of the

temporal and spatial patterns of sand and mud deposition along the 30 km of

coast indicated in Figure 1. Cores from marsh, mudflat, and nearshore

environments were analyzed to provide information on depositional

processes, as well as the mineralogy and rheology of the mud itself.

Second, a dynamics experiment was conducted during the passage of a

winter cold front, when a range of energy conditions, typical to this coast,

could be examined and compared over a short time interval. Rapid-response

flowmeters, pressure sensors, and a multi-level suspended sediment sampler

were deployed in the shallow nearshore zone where mud deposition occurs.

This equipment was used to record wave characteristics, and to monitor

suspended sediment concentrations and flow velocities at various elevations

above the bed.

Evidence of major depositional events was acquired during the long

term monitoring program. Although significant mud sedimentation did not

take place during the four day period of the dynamics experiment, analysis of


/

the time-series records yielded new insight into the nature of nearshore fine

grained sediment transport under waves.


6

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PR E V IO U S STU D IES

Even a brief survey of the literature reveals that the mechanics which

govern deposition of fine-grained sediments close to shore in the presence of

waves are largely unknown (Wells and Kemp 1985). The mere presence of

mud at the shoreface is troubling, as it appears to violate the consensus of

nearly a century of work establishing the preferential onshore transport of

the coarsest sediment fraction in clastic settings (Cornish 1898). The current

lack of understanding stems, first, from the absence of a general theory

relating mud deposition to fluid and sediment parameters, and, second, from

a paucity of field data. I. N. McCave (1984) addressed both of these factors

when he stated in a recent review that:

"there are no good ways of predicting the state of aggregation of polydisperse, multimineralic suspensions . . . , thus in natural situations empiricism reigns, but based on very few measurements".

Muddy coasts pose unique difficulties for researchers not only in the

theoretical sense, but also from the standpoint of logistics and

instrumentation. The development of new sensors, continued progress in the

laboratory, and an increased awareness of the disparity between what is

known about littoral sand and mud transport, have provided new impetus for

field efforts.

Gibbs (1970,1976) inferred from work in the vicinity of the Amazon

River mouth that a nearshore concentration of fine-grained sediments

occurred down coast through an Ekman-type coastal upwelling process.

Suspended sediments introduced from the Amazon are initially carried

7


8

offshore and the sand fraction is deposited relatively close to the river mouth.

The great majority of the suspended load is, however, deflected to the west

by the Guiana Current. There, it becomes concentrated in bottom waters

with a significant shoreward velocity component which returns only the clay

fraction to the coast.

Van Straaten and Kuenen (1958) and Postma (1961) provided the first

physical model to account for the selective nearshore deposition of fine

grained sediments in a coastal marine setting, albeit one dominated by tidal

rather than wave generated flows. On the North Sea tidal flats studied,

deposits of mud were restricted to the higher elevation surfaces which

experienced only the final stages of the flood tidal current. These workers

developed the well known "scour-lag" hypothesis to explain this size

segregation. In it, they proposed that fine-grained sediments which settled

from the waning flood current and during slack tide formed deposits

sufficiently competent to resist removal by the subsequent ebb tidal scour.

These models provide an indication of the span of spatial and temporal scales

at which processes important to nearshore mud deposition can be expected to

operate, but do not specifically address the basic dynamics problem at the

shoreface, that is, how mud is deposited under waves.

The limited data available suggests that the presence of gel-like fluid

mud is characteristic of most open littoral settings where significant

shoreface deposition occurs (Morgan et al. 1953, Delft Hydraulics

Laboratory 1962, NEDECO 1965,1968; Zenkovitch 1967, Nair 1976,

Martins et al. 1979, Wells and Kemp 1985). This term has been loosely

applied to unconsolidated clay and silt suspensions with sediment

concentrations ranging from 10 to 300 g-1"1 (Einstein and Krone 1962).


/

9

Chronic channel siltation associated with the occurrence of fluid mud

in San Francisco Bay led Einstein and Krone (1962) to conduct a classic

series of laboratory experiments on the modes of cohesive sediment transport

in salt water. Suspensions with salinities of l ° / o o and greater were found to

be of sufficient ionic strength to ensure that collisions between clay particles

would result in aggregates with lasting bonds. They proposed that fluid mud

forms from coagulation of dispersed fine-grained sediments in a flocculation

process involving collisions caused by both Brownian motion and fluid shear.

In the dispersed suspension, away from the bed, collisions between

particles take place primarily as a consequence of Brownian motion. Closer

to the bed, however, where both the particle population and the range of

particle sizes are greater, Einstein and Krone (1962) showed that local shear

could cause collisions at a rate far greater than Brownian motion alone.

Under these conditions large floes have a high probability of collision with

small ones and thereby continue to grow in size as they fall toward the bed.

The maximum size attainable by cohesive particles in the suspension is

thought to be determined by a complex and largely unknown interplay

between settling velocity, shear at the bed, and the strength of aggregate

bonds (McCave 1984). Einstein and Krone (1962) proposed that destruction

of floes by shear stress near the bed might return disaggregated particles to

the suspension and thereby restrict clay sedimentation.

Stow and Bowen (1980) suggest from their examination of silt

laminated turbidite muds that shear within the bottom boundary layer may

indeed act to sort silts and clays by concentrating clays in suspension while

allowing silts to deposit. In order to explain the graded and laminated silt-

mud couplets they observed, they proposed that at some critical


10

concentration, shear is "overcome" and clays deposit rapidly as a blanket

over the silt laminae. Rine and Ginsburg (1985) found silt and clay couplets

in radiographs of Surinam mudflat deposits and have proposed that a similar

sorting process may operate in this coastal setting. No mechanism was

identified in either study to explain the required rapid clay deposition. More

than 40 years ago, however, Einstein (1941) explained sudden deposition, or

"freezing", of mud-charged subaqueous underflows as a consequence of the

tendency for such suspensions to acquire internal, or yield, strength in excess

of the shear stress applied.

It has long been known from viscometer measurements, that clay

suspensions in the fluid mud range behave as plastic solids rather than true

fluids (Bingham 1922, p. 215). As such, they are characterized by a capacity

to resist deformation up to some critical level of imposed shear stess

(Einstein 1941, Einstein and Krone 1962). The magnitude of the applied

stress required to initiate flow, or conversely, the level below which flow

ceases, defines the yield strength of the suspension. Krone (1962) found that

the yield strength of fluid mud from San Francisco Bay increased in

proportion to the sediment concentration raised to the power of 2.5. The

results of subsequent work on harbor muds from other locations suggest that

this empirical relationship may have broad applicability (Owen 1975,

Hydraulic Research Station 1979).

Because of the cohesive nature of the clay particles, deposits formed

from fluid mud suspensions produce beds which also exhibit non-Newtonian

rheological characteristics (Faas 1981). Maa and Mehta (in press) have

recently prepared a useful review of the extensive engineering literature on

this topic, which is briefly summarized here.. Beds of cohesive sediments


11

respond to wave-induced shear in a complex manner which ranges from

elastic deformation to mass erosion (Alishahi and Krone 1964, Migniot

1968). Orbital motion and pressure oscillations associated with surface

waves have been observed to cause horizontal and vertical displacements at

and below the mudline both in the laboratory (Lhermitte 1960, Migniot

1968, Doyle 1973), and in the field (Tubman 1977). Furthermore, response

to one set of wave conditions may change with time. Mud beds "soften"

under cyclic loading (Thiers and Seed 1968, Schuckman and Yamamoto

1982) or suddenly liquify when pore pressures build beyond some critical

point (Turcotte et al. 1984).

The work performed by waves at and below the mudline effectively

dissipates wave energy, thereby attenuating wave heights at a far greater rate

than does normal friction over a consolidated sand bed. To explain wave

damping observed over the "mud hole", a region of fluid mud located

offshore of the central Louisiana coast, Gade (1958) provided an analytical

solution for the dissipation of wave energy by a deformable mud bed. In this

case, the bed was considered a viscous fluid. Since then, other analyses have

assumed bed response to be elastic (Gade 1959, Mallard and Dalrymple 1977,

Dawson 1980) or viscoelastic (Tubman 1977, Macpherson 1980, Hsiao and

Shemdin 1980, Maa 1986).

The study of Wells and Coleman (1981a) was the first to address the

effect of shoaling waves on mud suspension and transport nearshore.

Working along the coast of Surinam, they observed not only attenuation as

waves propagated over a gently sloping (0.0005) bottom blanketed with fluid

mud, but also a deformation of the characteristic wave profile from

sinusoidal in deep water to the steep, symmetrically crested, flat-troughed


/

12

shape best described by solitary wave theory. Plunging breakers were

observed at mud-free interbank locations, but at stations fronted by fluid

mud, the long-crested solitary-like waves disappeared completely before

reaching the shoreline. Wave height decreased with water depth at a constant

ratio (H/d) of 0.23, far below 0.78, the critical value at which breaking can

theoretically occur (McGowan 1894). By monitoring pressure variations in

the upper 0.5 m of fluid mud, Wells and Coleman were able to obtain time-

series records of sediment suspension and deposition. These data indicated

that sediment was suspended at frequencies ranging from tidal (T=12.4 hr) to

that of the incident waves (T=10 s). An intermediate, infragravity range was

also identified (T=0.5 to 5 min) in this low-slope, dissipative environment.

Suspended sediment concentrations were highest in areas where

solitary-like waves were observed. Arguing that these waves might, in fact,

be waves of translation, Wells et al. (1979) proposed that high rates of

onshore sediment transport might be possible without invoking breaking

waves or nearshore circulation cells.

Solitary wave theory predicted free-surface elevations well, but as

Wells et al. (1979) point out, several of the assumptions upon which this

theory is based (Boussinesq 1872), particularly the requirements for an

inviscid fluid, irrotational motion, and a rigid bottom, are certainly violated.

High quality velocity time-series from the muddy nearshore are clearly

needed, but, heretofore, have been unavailable.

In the past 10 years, the advent of rapid-response flowmeters has

permitted direct measurements of nearshore flows associated with shoaling

and breaking waves (Huntley and Bowen 1973, Meadows 1976). Spectral

analysis of such time-series has been used effectively to isolate the low


/

13

frequency components of the nearshore velocity field which govern the net

transport of sediments suspended by waves (Meadows 1976). When

flowmeter data is acquired simultaneously at several levels above the bed,

boundary layer models, developed by analogy from laminar flow theory, can

be used to parameterize the turbulent shear flow. Results generated by these

simple models, when applied to sediment-laden flows (Smith and McLean

1977a, b; Adams and Weatherly 1981a, b), and to flows affected by waves

(Grant and Madsen 1979), have guided recent efforts to understand factors

critical to sediment erosion, transport, and deposition.

In the present study, I extend the work begun by Wells and Coleman

(1981a, b) to include nearshore velocity data. The equatorial coast of

Surinam enjoys relatively constant wind and wave conditions year-round and

is rarely affected by storms. In contrast, the muddy coast o f Louisiana,

located 30° to the north, outside the trade wind belt, experiences a far greater

range of weather types. Data was collected on time-scales ranging from

seasonal to wave frequency in order to place a quantitative study of the

nearshore sediment dynamics in the context o f this evolving shoreline.


F IE L D A REA

The 30 km section of coast selected for study lies west of the Southwest

Pass of Vermilion Bay on the eastern margin of the Louisiana chenier plain

(Fig. 2). This area has historically been the primary locus for chenier plain

mudflat deposition (Morgan et al. 1953, Morgan and Larimore 1957,

Coleman 1966, Morgan and Morgan 1983). It also includes segments of

three other shoreline types which are more representative of the rest of the

modem chenier plain coast, namely, unprotected marsh scarps, perched

washover beaches, and reactivated strandline or chenier beaches

(Kaczorowski and Gemant 1980, Wells and Kemp 1981).

In plan view, this shoreline projects into the G ulf as a gradual headland

with an apex at Tigre Point. The curvature is more pronounced on the

eastern arm of this bulge, where the broad, shallow Trinity Shoal platform

extends 30 km offshore. The presence of this shoal results in a lessening of

the offshore slope from 0.001 (0.06°) at Dewitt Canal on the western margin

of the study area to 0.0005 (0.03°) at the eastern edge, 10 km west of

Southwest Pass.

From offshore, this low-lying shoreline appears almost featureless,

broken only by the outlet of Freshwater Bayou Canal. This artificial ship

channel is periodically dredged to maintain a 4 m depth in its offshore

approach, but discharge of freshwater and sediment is regulated by locks 2

km inland. Otherwise, the only landmarks are spoil banks associated with

pipeline canals, and the 2 to 3 m relief of Cheniere au Tigre, a truncated

beach complex which obliquely intersects the modem coast 5 km east of

Tigre Point.

14


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92-25'

DEWITT CANAL^ CHEN|EREAUT!iatti

PIPELINE CANALS

D jU .S .A .C .d .E .' TsTIDE GAGE ; :

/FRESHWATER 'BAYOU CANAL. iTIGRE:

.POINT;

□ Marsh □ Beach Ridge IB Tidal Mudflat Q Profile Station

Figure 2. Map of study area, showing major shoreline features and beach survey stations.

o

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16

GEOLOGY

Early workers recognized that the chenier plain evolved as a result of

the coastwise transport and deposition of sediments derived from Mississippi

deltaic systems to the east. Russell and Howe (1935) were the first to

hypothesize mudflat formation as the primary mechanism for Recent coastal

progradation. They interpreted the linear, generally coast-parallel sand/shell

ridges, locally called "cheniers", which now form "islands" surrounded by

marsh, as beaches built by waves during intervals of slackened fine-grained

sediment supply.

Fisk (1948) suggested on the basis of limited borings that it might be

possible to correlate major progradational episodes on the chenier plain with

the proximal location of specific Mississippi subdeltas. Price (1955)

generalized the facies relationships described by Fisk (1948) into what he

called the "chenier plain" type coast and was the first to note the analogy

between this shoreline and that of the Guianas in South America. However, it

was not until the completion of a major drilling program in 1959 that the

geochronology and stratigraphic complexity of the chenier plain really began

to be understood. Gould and McFarlan (1959) and Byrne et al. (1959) were

then able to reconstruct its genesis on the basis of radiocarbon dating and

paleofaunal analysis, respectively.

Briefly, their interpretation indicates that as post-glacial sea level rose

from -5 m to near its present level between 5600 and 3000 years B.P., the

Gulf of Mexico inundated a dissected Pleistocene prairie surface forming a

complex coastline of shallow lakes and bays. Sea-level rise ushered in a new

era of delta development to the east as lobes spread out over large areas of the

shallow shelf. Fine-grained sediments accumulated in the estuaries and filled


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17

depressions in the Pleistocene surface. The oldest, most inland chenier ridges

formed as spits or detached barriers coincident with the gradual leveling of

the rate of eustatic rise at approximately 3000 years BP.

Subsequently, two Mississippi delta complexes, the Teche, and the

Lafourche, built out from the coast in the area between the present river

mouth and the chenier plain (Kolb and Van Lopik 1958). The supply of both

fine and coarse-grained sediment to the chenier plain was increased during

these episodes, and progradation accelerated. Sand and shell chenier ridges

continued to form, but, from this point on, mudflats, which previously were

confined to back-barrier bays, began to appear at the coast itself, onlapping

beaches on the Gulf side. These mudflats make up the bulk of the late Recent

regressive sequence.

In geological terms, the stratigraphy of the chenier plain documents a

major progradational episode. Along a 150 km long front, the position of the

shoreline was extended 20 to 30 km south into the Gulf of Mexico, advancing

at a mean rate of 5 to 10 m -y r l . The presence of the abandoned coarse

grained chenier ridges indicates, however, that this has not been a continuous

process. Today, in fact, truncated and eroding chenier deposits all along the

modem shoreline signify that the long-term progradational trend is, at least

temporarily, in abeyance (Kaczorowski and Gemant 1980). Indeed,

historical records suggest, paradoxically, that coastal retreat on the chenier

plain has averaged nearly 10 m -y r l in the 200 years since the first reliable

surveys were made (Morgan and Larimore 1957, Morgan and Morgan

1983).

Wave reworking of chenier beaches provides most of the clastic

sediments currently in transport along the coast (Beall 1968). It is only


18

within the coastal bulge encompassed by the study area, and to a lesser extent,

near Rollover Bayou, 20 km farther west (Wells and Kemp 1981), that

significant fine-grained deposition and localized progradation continue

today.

Despite the regional erosional trend, these areas experienced a

dramatic influx of mud in the early 1950's. It was at this time that Morgan et

al. (1953) mapped subaerial mudflats fronting the coast throughout the study

area from Cheniere au Tigre west. These workers attributed this deposition,

which meant economic disaster to the small beach resort developed at

Cheniere au Tigre (Bailey 1934), to the growing importance of the

Atchafalaya River as a Mississippi distributary, and a source of fine-grained

sediments to the shelf.

INNER-SHELF CIRCULATION

Atchafalaya River discharge, comprising roughly a third of the

combined Mississippi and Red River flows at Simmesport, Louisiana,

averages 5126 m ^-s-l (USACOE 1974). It is dispersed onto the shelf

through a dredged navigation channel and, to a lesser extent, through other

breaks in an oyster reef at the seaward margin of Atchafalaya Bay (Roberts et

al. 1980, Wells et al. 1983). The most significant influx of low salinity,

sediment-laden waters from the Bay occurs during the spring flood which

typically builds from late December through April, and declines rapidly

thereafter (Angelovic 1976, Dinnel and Wiseman 1986, Cochrane and Kelly

1986). Flood discharge commonly exceeds the average rate by 300 percent

(USACOE 1974).

Tidal and wind-driven currents control the mixing of Atchafalaya and

Gulf waters, and the dispersal of river-borne sediments on the inner


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19

continental shelf. Mixed diumal and semi-diurnal tides in the study area have

a mean amplitude of 60 cm (Manner 1954; NOAA, NOS 1981). Currents

associated with these tides rotate in a clockwise manner with average speeds

inside the 10 m isobath of less than 10 cm-s- * (Murray 1976). Because of the

well-behaved rotation and low velocity of these flows, wind-driven currents

with speeds ranging up to 50 cm-s“l (Adams et. al. 1982) play a more

significant role in shelf sediment transport except in the immediate vicinity

of the tidal passes of Atchafalaya and Vermilion Bays (Todd 1968).

Cochrane and Kelly (1986) have recently integrated a variety of shelf

current data to provide a coherent regional perspective. Their work

indicates that flow on the Texas-Louisiana shelf is dominated for 10 months

out of the year by cyclonic circulation around a mid-shelf geopotential low

which appears off Texas in September and elongates to the east through the

winter and spring. Westerly flow in the coastal limb is balanced by an

easterly countercurrent over the outer shelf. Cochrane and Kelly (1986)

show convincingly that the nearshore portion of this gyre is a wind-driven

coastal boundary current (Csanady 1977), which is set up by the mean

alongshore component of the wind stress, directed to the west. This

interpretation is supported by the rapid change in circulation which occurs

when the regional wind stress shifts to easterly in July and August. The

cyclone which controls the inner-shelf current regime is abruptly replaced

for these two months by anti-cyclonic flow. The geopotential high around

which this circulation develops is located just off the coast of the study area,

but drives significant northerly and easterly nearshore currents only along

the Texas coast. Cyclonic circulation is re-established in September with the

onset of winds from the east.


20

High winds promote vertical mixing of the water column and build

waves which put bottom sediments into suspension. Barring hurricanes and

tropical storms, sustained high winds of regional extent are common only

between September and April when extratropical cyclones or cold fronts

cross the coast every 5 to 10 days.

Femandez-Partagas and Mooers (1975) have described the detailed

structure of surface wind systems associated with cold fronts traversing the

northern Gulf of Mexico. They found that hourly observations of wind

speed and direction made at a single station exhibit a characteristic sequence.

A pre-frontal period of steady southerlies of increasing magnitude ends with

a rapid clockwise shift in direction coincident with front passage. After a

180° rotation, a second steady wind regime, this time with a northerly

orientation, becomes established, and wind speeds gradually diminish.

The dominant cross-shelf component of the wind stress during these

fronts results in sequential water level set-up and set-down along the east-

west oriented coast of the study area (Chuang and Wiseman 1983). Wave

energy also builds during the steady southeasterlies which preceed front

passage. Wind-forced flow is to the west and onshore (Daddio et al. 1978).

The onset of northwesterly offshore winds gives rise to short-lived, but

relatively high-velocity, eastward setting currents (Crout et al. 1985, Adams

et al. 1982), and strong inertial oscillations (Daddio et al. 1978).

STORMS AND MODERN MUDFLAT SEDIMENTATION

Some of the same investigators who originally mapped the mudflats on

the eastern margin of the chenier plain (Morgan et al. 1953) revisited this

coast five years later, following the passage of Hurricane Audrey in June,

1957 (Morgan et al. 1958). This devastating storm went ashore near the


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Texas-Louisiana border and generated a 3 to 4 m storm surge in the study

area. Morgan et al. (1958) noted that although the hurricane caused an

average of nearly 100 m of coastal retreat along most of the chenier plain, the

mudflat-fronted shoreline experienced very little erosion. In two locations,

in fact, shoreline progradation was documented. There, mud was deposited

as discrete shore-welded "mud arcs" or bars 2 m thick with alongshore

dimensions of nearly 4 km. The mud in these features had abrupt, easily

defined lateral boundaries. The sharpness of the contact suggested to the

researchers "that the mass of fluid mud was transported by the storm tide and

deposited as a unit". Seeing the effects of the hurricane brought to mind an

observation made during the initial surveys in the spring of 1952. In 1958,

Morgan et al. wrote:

"There is morphological evidence of the effectiveness of suspended mud during the height of the storm. With normal storms this material is moved toward the shore where some of it may be permanently or temporarily incorporated into the beach. During field work in 1952 a beach profile was surveyed across an essentially non- mudflat area. Three days later following a m inor storm the area was revisited. The storm had blanketed the foreshore with a layer of gelatinous clay which had a maximum thickness of about 18 inches."

In the mid-1960s, Coleman (1966) found little change with respect to

the distribution of mud in the study area, but added much new information on

the physical and biological properties of the mudflat sediments. He X-rayed

mudflat cores which, to the unaided eye, appeared almost featureless.

Radiographs of "massive" sections disclosed multiple sets of parallel

laminations relatively undisturbed by biological reworking. The laminae


/

were composed of concentrated silts slightly coarser and denser than the

predominantly clay matrix.

The observations made by Morgan and his colleagues in 1952 and

1957, together with the sedimentological data provided by Coleman (1966),

suggest, somewhat surprisingly, that chenier plain mudflats are deposited

rapidly under the highest energy conditions this coast experiences.


DATA A C Q U ISIT IO N

Data were collected over a 2.5 year hurricane-free period from

December, 1980, to May, 1983, and during a follow-up reconnaissance made

in May, 1985. The bathymetry and surficial sediment distribution was

surveyed along the 30 km length of the study area to 4 km offshore, or

roughly the 3 m isobath. Geomorphological, sedimentological, and

rheological data were acquired in beach and nearshore environments at

shoreline stations spaced 3 to 5 km apart (Fig. 2). This work, summarized in

the first subsection below, provided the information necessary to design a

dynamics experiment and interpret its results in a regional context. The

second subsection covers the experiment conducted in December, 1982.

SEDIMENT SURVEY

Depth and positioning data for the offshore survey were obtained with

a Raytheon 731 depth recorder and a Decca-DelNorte trisponder system,

respectively. One hundred-eighty bottom samples were collected at a 0.5 km

spacing along arced traverses which intersected the coast at 1 km intervals. A

long-handled scoop-type sampler was used to minimize loss of fine-grained

sediments in this shallow water setting. Samples were analyzed for percent

sand, shell, silt and clay using standard techniques (Folk 1980). Sizing of the

silt and clay fractions was accomplished by Coulter Counter (Sheldon and

Parsons 1967).

Benchmarks were established at eight field stations and beach profiles

were surveyed at these locations every 3 to 4 months. The stations were

located using the trisponder system and benchmark elevation was determined

by referencing local water levels at each station to the datum of a recording

23


24

tide gauge maintained by the U. S. Army Corps of Engineers at Freshwater

Bayou Canal. Survey lines were initiated on a stable marsh surface landward

of washover effects, carried over the berm, if present, and continued

offshore as far as visibility and water depths permitted (approximately 200

m). Where soft muds were present, the bottom was defined as the upper

surface of the mud layer. Profiles were digitized relative to mean sea level

(MSL = +0.24 m Gulf Coast Low Water Datum). The shoreline was defined

as the point where the MSL datum intersects the profile. Shoreline

progradation or retreat was determined from the translation of this point

relative to the benchmark.

Push-cores (Van Straaten 1954) up to one meter long were obtained to

characterize deposition in nearshore sub-environments. These cores were

split and described from X-ray radiographs of 1.5 cm thick slabs using the

methods of Roberts et al. (1976). They were compared with sequences

identified in a 5.5 m long vibracore (Laneski et al. 1979) recovered at profile

station 3 from a vegetated mudflat formed in the early 1950s (Morgan et al.

1953). In addition to being X-rayed, this core was sampled at 5 cm intervals

and analyzed for textural variability in the same way as the offshore

sediments.

A push-core obtained in the sub-aqueous section of a newly deposited

mudflat was analyzed to determine the mass physical properties of this

sediment. The core was transported to the lab in a vertical position, and was

subsampled with a glass tube composed of pre-weighed, detachable 5 cm

sections of known volume. The tube was vibrated slightly and allowed to

sink into the sediment as described by Sikora et al. (1981). Upon extraction,

the tube was segmented and each section weighed to determine bulk density.


25

In this way a qualitative record of the vertical variation in density and degree

of consolidation was obtained. These data were then compared with grain

size measures determined by pipetting (Folk 1980).

Surficial sediments from the same location were subjected to X-ray

diffraction to determine clay mineral composition (Schultz 1964, Griffin

1971) and analyzed to determine viscosity and Bingham yield strength in a

capillary viscometer linked to a manometer to allow application of variable

known head pressures. This equipment and the method were first described

by Einstein (1941), and more completely later by Das (1970).

DYNAMICS EXPERIMENT

Nearshore wave characteristics, flow velocities, and suspended

sediment concentrations were monitored at profile station 6 between 9 and 13

December, 1982. A cold front traversed the study area during this time, and

data was collected in three sessions corresponding to the pre-frontal, frontal,

and post-frontal periods. At the time of the experiment, nearshore soft mud

accumulations averaged 0.5 m in thickness. A tide record was obtained for

the duration of the experiment using a pressure transducer water-level gauge

fixed to a platform at the mouth o f Freshwater Bayou Canal, 2 km east of

station 6. Hourly observations of wind speed and direction were recorded at

the Freshwater Bayou Locks.

The smooth mud bottom at station 6 had a slope of 0.004 (Fig. 3).

Time-series data were acquired at two shallow water locations (d ~ 1.0 m) 10

m apart along the shore-normal transit. The landward station (B) was

approximately 90 m from shore. At each location, a 2 cm-diameter steel rod

was driven vertically into the bottom. The transducer of a pressure-type


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I N S T R U M E N T S U P P O R T

R O D S

E L E C T R O N I C R E L E A S E

A N A L O G R E C O R D E R S

S U S P E N D E DS E D I M E N TS A M P L E R S B I - D I R E C T I O N A L

DUCTE M P E L L O O W M E T E H S

■' ■ / / / / / / / / / / / / / / . ' / / . '' / / / / / / / /

/ / / / ••' / // / .. 7 7 - ‘ / / / / / / / / / / / / / / . / ’. .B O T T O M 7 / V /V * * A / L V Y / / / / / /

j- L .S E N S O R S

Figure 3. Cross-section of shoreface at station 6, showing deployment of sensors during cold front experiment.

rv>CD

27

wave gage (Fredericks and Wells 1980) was clamped to the landward rod

such that the diaphragm of the sensor was fixed at the sediment-water

interface. An identical transducer was attached to the seaward rod (A) at the

same level as the first to ensure that the frequency response factor was the

same for both sensors.

Three miniature (8 cm diameter), ducted, bidirectional impellor-type

flowmeters fabricated at Coastal Studies Institute were fixed with a common

orientation to a 2.5 cm diameter stainless steel sleeve. The sleeve was

dropped over the landward steel rod and lowered until it rested against the

housing of the pressure transducer. In this position, by rotating and

clamping the sleeve, the flowmeters could be oriented to measure either

onshore-offshore or alongshore flows at 13, 30, and 60 cm above the bed.

All data were recorded in analog form aboard a small boat moored 30 m

from the instrument array. Signals from the wave gages were logged by a 2-

channel Gould Brush strip chart recorder. The sense o f rotation and

intervals between discrete on-off signals generated by the rotating flowmeter

contacts were converted electronically to positive and negative instantaneous

velocities which were logged on heat-sensitive paper by two Astromed

recorders.

Nine 20 minute data runs, in addition to a trial run (No. 1, not used in

this analysis), were completed during the course of the experiment, including

4 paired runs in which the flowmeters were sequentially oriented onshore

and then alongshore. During the last two runs, a recorder malfunction

limited the number of time-series obtained. During these runs, pressure data

was acquired only at the shoreward station (B), and flowmeter data was

limited to the 60 and 13 cm elevations.


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28

During runs 2 through 10, 39 instantaneous suspended sediment

concentration profiles were obtained within 2 m of the instrument array

using a multi-level sampler modified from that of Kana (1979). The sampler

consisted of four 1-liter Van Dorn bottles stacked on a frame and fitted with a

common release. When triggered with the bottom of the frame at the bed,

samples were obtained simultaneously at mean elevations of 10, 35, 65, and

95 cm above the bottom.

In the laboratory, measured aliquots of the water samples were passed

through Millipore 0.45 micron filters using a pressure filtration system. The

filters were weighed to determine sediment concentration following standard

procedures (Meade et al. 1975).

Pressure and flowmeter records were digitized at a 0.5 s time step

which was determined to preclude aliasing (Nyquist frequency = 2 cps). The

power spectra were computed (Bendat and Piersol 1966) using a Fast Fourier

Transform (FFT) algorithm (Cooley et al. 1969). Coherence-squared

statistics and phase relations were calculated between similar time-series

(Jenkins and Watts 1968). Low-pass smoothing of the records was

accomplished in the frequency domain by suppressing Fourier coefficients

with a frequency greater than 0.005 cps (T = 200 s).


R E SU L T S

SEDIMENT SURVEY Geomorphologv

The nearshore bathymetry and distribution of sediments in the study

area are mapped in Figure 4. Isobaths show the effect of Trinity Shoal on the

offshore gradient and outline a 3 km^ mudflat between profile stations 7 and

8. Poorly sorted muds with a median size of 2.5 microns and a silt content

ranging from 20 to 40 percent dominate the nearshore surficial sediment

suite. Fine sand (3 phi median) fronts the coast east of Freshwater Bayou

Canal. There, it expands west to east from a ribbon a few 10's of meters wide

near station 5 to an apron nearly a kilometer wide off station 1. A second

band of sandy sediments in this area also trends to the east from Tigre Point

somewhat seaward of the first. The only sand-sized sediments on the other

side of the canal are found trailing west from the dredged spoil disposal area

on the margin of the navigation channel.

A summary of grain-size, morphological, and shoreline change

statistics at the eight survey stations is given in Table 1. These results

confirm indications in the offshore data that Freshwater Bayou Canal divides

the shoreline of the study area into two distinct sedimentary provinces.

Unvegetated mudflats 10's to 100's of meters wide characterize the coast west

of the canal (Fig. 5a). There, clastic sedimentation is limited to pockets of

small bivalve shells (Mulinea lateralis. Nuculana concentricaV and organic

"coffee grounds". In contrast, sand and shell beaches are continuous east of

the canal, except for a gap of eroding marsh between station 3 and Cheniere

au Tigre. This space of less than 2 km separates two beach systems which

differ significantly in composition. Oyster shell fragments fCrassostrea

29


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DEWITT CANALCHENIERE AUTIGR

PIPELINE CANALS

mSEHR

FRESHWATER- CANALo OINTJ

saxet

CONTOUR INTERVAL 0 .5 METERS BELOW MEAN LO W W ATER O A tU M

□ M arsh □ B sach Rldga □ Tidal Mudflat Q Profit* Station □ Mud □ Silt E3 Sandy Mud E3 Muddy Sand ♦ B ottom SamplingS tation

Figure 4. Nearshore surficial sediments and bathymetry in the study area.

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Table 1Chenier Plain Shoreline Characteristics

December, 1980 - May, 1983

Station (E to W) 1 2 3 4 5 6 7 8

Morphology!') OWB CB MS OWB OWB M M OWB/M

Foreshore (MSL to berm crest)Median size (phi) 1.5 0 .0 3.0 3 .0 -4 .0 <4.0 <4.0 <4.0Height (m) 1.9 2 .4 1.1 1.4 1.4 1.1 1.1 1.6S lope .05 .08 .08 .05 .04 .02 .02 .01

Inshore (MSL to100 m offshore)Median size (phi) 3.0 3 .0 3.0 3 .0 3 .7 5 <4.0 <4.0 <4.0

Features!2) B B B FM B/FM FM FM FMSlope .002 .002 .003 .009 .009 .005 .004 .004

Shoreline change between surveys (3 to 4 months)Max advance (+m) 8.0 6.5 2.0 5 .5 1.0 20.5 144.0 107.5Month observed Jul Jul Jul Jul Oct Mar Mar May

Max retreat (-m) 10.0 13.5 6.2 5 .8 4.5 27 .5 90 .0 19.0Month observed Dec Oct May May Mar Jul Sep Jul

Net 2 year change:May, 1981-May, 1983 (m) -15.5 -1.0 -7.8 +1.5 -9.0 +15.0 +20.0 -3 3 .0

0)OWB = overwash beach; CB = chenier beach; MS = marsh scarp; M = mudflat. <2)B = sand bars; FM = fluid mud

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A

Figures 5a-b. Muddy shoreface at station 7 (5a); reactivated chenier beach (station 2) at Cheniere au Tigre (5b).

33

virginical eroding out of the reactivated beach ridge form an important

component of the relatively steep beaches extending from Cheniere au Tigre

to the east (Fig. 5b). In contrast, low-profile overwash beaches with little

shell characterize the shoreline extending east and west of Tigre Point.

An abrupt scarp marks the seaward edge of the marsh in the area

between the two beach systems. This marsh is the remains of a large expanse

of unvegetated mudflats which Morgan et al. (1953) observed during the

1950's. Significant clastic beach sedimentation did not occur at survey

station 3, which was located in this section, throughout the duration of the

field program (Fig. 6a), although bars of fine sand resting on consolidated

marsh sediments were often present just offshore. When this location was

revisited in 1985, however, a low-profile sandy beach covered the survey

station (Fig. 6b). This deposition was clearly a consequence of westerly

longshore transport, as the boundary of the sandless shoreface could still be

found 0.5 km to the east.

Fine details of shoreline change are not resolved in beach profiles

surveyed at 3 to 4 month intervals. The results of the 2.5 year program do,

however, provide important information about seasonal and spatial patterns

of deposition and erosion. A complete set of the profiles obtained at the eight

survey stations is compiled in Appendix A.

The sandy beaches east of the canal undergo an annual cycle of winter

erosion followed by onshore translation of a bar, or multiple bars, during the

summer and fall (Fig. 7). The most extensive bar development occurs at the

easternmost stations, 1 and 2. Although the position of the shoreline was

stable at Tigre Point (Sta. 4) and Cheniere au Tigre (Sta. 2), the remainder of

the stations east of the canal experienced between 8 and 16 m of coastal


34

■ h

Figure 6a-b. Time-lapse views of the vegetated mudflat mapped by Morgan et al. (1953) at station 3; eroding marsh scarp, May, 1982 (5a), replaced by clastic washover beach, May, 1985 (5b).


Non-winter Accretion at Clastic Station 4

14 DEC'8115 MAR'82

1.0m

Sand Bar Formation250m50mMSL

200m100m 150m

13 JUL'82

15 MAR'82

20 OCT'8213 JUL‘82

Onshore Bar Translation

Figure 7. Bar migration at Tigre Point (station 4).


36

retreat over the two-year period between May, 1981, and May, 1983 (Table

1).

Mudflat Sedimentation

The beach survey program demonstrated that, non-hurricane, littoral

deposition of low-density muds takes place west of the canal only between the

months of January and May. The most extensive mudflat observed during

the study period formed between stations 7 and 8 after a survey in mid-

December, 1981, but before the next visit in late February, 1982. It was

initially arcuate in form, and closely resembled those mapped by Morgan et

al. (1958) after Hurricane Audiy. The shore-normal and alongshore

dimensions were estimated at 300 m, and 1500 m, respectively, and the

maximum thickness at 2 m. Although the eastern arm of the arc was not

visible after 6 months, the western arm survived as a subaerial spit obliquely

intersecting the coast for at least 1.5 years (Fig. 8).

An elevated bar or berm on the seaward margin of the mudflat is

typical of new deposits (Fig. 9) of highly fluid sediment (Fig. 10a).

Subsequently, the bar and other regularly exposed areas dry and the surface

cracks into ever smaller polygons (Fig. 5a). When these areas are subject to

wave action, pieces of the surface crust are ripped up and the mudflat

acquires an irregular surface. The bar is usually gone by the beginning of the

winter following deposition (Fig. 9). Mud clasts, shells, and "coffee

grounds" accumulate in offshore-oriented erosional grooves or flutes (Fig.

10b). If new mud is deposited the next spring, that part of the mudflat which

survives the winter will be protected from further erosion. Oyster grass

(Spartina altemifloral then propagates onto this newly available land.

Freshly deposited mud is easily removed, and mudflats undergo


/

Figure 8. Oblique aerial view of western limb of large mud bar formed in Spring, 1982, between stations 7 and 8.

P r o l l la S ta t io n 8

ACCRETION1.5m -

EROSION1.0m

13 FEB’81

0.5m2 9 MAY’8 1

MSL

1.0m -

0 .5 m2 9 MAY’S 1

MSL 50 m 3 0 0 m1 0 0 m 2 5 0 m2 9 J U L '8 1

1 5 0 m

-0 .5 m

Figure 9. Profile time-series of mudflat accretion and erosion at station 8, February - July, 1981.


38

r

Figure 10a-b. Shoreface mud at station 7; Freshly-deposited fluid mud, Spring 1982 (10a); eroding surface showing development of shore-normal flutes 5 months later (10b).


39

erosion in all seasons. The most rapid retreat, however, occurs during the

summer immediately following a major depositional event (Table 1). Then,

the bar loses definition as sediments are redistributed both landward and

seaward along the profile section (Fig. 9). Progradation may more than

offset shoreline retreat in any given year, but the required sedimentation

does not occur at a uniform rate, or at the same locations each year. As a

result, zones experiencing major deposition in any single spring may be

separated by long stretches of shoreline which receive little or no mud. The

data summarized in Table 1 shows that between May, 1981, and May, 1982,

shoreline translation, whether through accretion or erosion, was far more

active at mudflat stations 6,7, and 8, than at any of the stations east of the

canal. The shoreline shifted 15 and 20 m seaward in this period at stations 6

and 7, respectively, but moved 33 m landward at station 8. It is illustrative of

the dynamic nature of the mud-fronted coast that between February and May,

1981, the months just prior to the two year period considered above, the

shoreline at station 8 moved seaward 108 m (Fig. 9).

A log of the 5.5 m vibracore from station 3 is provided in Figure 11.

Coring took place 20 m from the shoreline in the marsh which now covers a

mudflat described by Morgan et al. (1953). The core penetrated a

bioturbated sandy mud at 3.2 m below the marsh surface which matches

Coleman's (1966) description of nearshore marine sedimentation.

The upper 2 m o f the core record the mudflat and marsh sequence.

The basal 0.5 m of this sequence fines up and is composed of poorly-defined

layers of silty-sand and clayey-silt with no discernible erosion surfaces.

Radiographs of this section show wavy bedding distorted by water escape

structures (Fig. 12a) which may be artifacts of the coring process. In


/

CHENIER AU TIGRE-CORE T 3

«SAND tSH E L L «ORG C LfTHOLOGV 0 SO 0 10 0 2

UNIT DESCRIPTION

wim Imbricated eJwie

2to lOcmttttoKbattocomooeedof lemmatad dbownoyrnaeaaamudiacwand M common at baaa of ootoata; btrrowa. roots, mud cracfca, and concisions tHupaauaon: •hall haen and aand present atbaaa ofutt ;iarmnationa change from oonvaua and tontiojnr naar oaaa to wavy and flat hmg farther up; Aama and toed svucaras convnon naar Oaaa

dean aand and anal; tomnaoona ndacrrt btf no burrows teoma wavy iamrae andamel- seals i-baddtoq; tnoncatedshalel baaa

Intaroaddad aand end mud;wedge-shaped poorty-aonedlaminattons; shelt-noah at baaa; no burrowa

heevfy bisrowed. faw tontteufar and wavy aand/attt laminae dUcernaols ; aftail haan secumuettone{burrowingrHenwty Ineraaaaa witti depth

DEPOSmONALENVIRONMENT

BEACH/M ARSH

MUDFLAT

BAR

TR A N SITIO N A L

(BAR TROUGH)

N EA R SH O R EMARINE

Figure 11. Log of vibracore from marsh now covering mudflat described by Morgan et al. (1953) west of Cheniere au Tigre at station 3; see Figure 6a-b.


contrast, bedding is well-preserved in the overlying sand-free mudflat

sediments and shows repeated 2 to 10 cm thick beds consisting of flat and

lenticular silt laminae overlain conformably by a layer of featureless silty-

clay. The clay layer makes up an estimated 60 to 90 percent of the thickness

of each couplet. Erosional surfaces separating the couplets are marked by

scour-and-fill structures (Fig. 12b), and, in some cases, by thin layers of

imbricated Mulinea shells (Fig. 12c). Five of these cycles were recognized

below the bioturbated, organic carbon-rich marsh horizon. A 0.2 m thick

cross-bedded sand and shell washover deposit caps the rooted zone.

Mineralogy and Rheologv

The cyclic bedding observed in the upper mudflat section of the

vibracore produces variations in bulk density with depth. Overall, the trend

in a push-core obtained through newly deposited mud is from 1.2 g-cm-3 at

the surface to 1.4 g-cm-3 a meter below (Fig. 13). The density increase

occurs as a series of steps, however, increasing downward within each silt-

mud or shell-mud couplet to the point where grain-size data indicates the base

of the cycle. Bulk density drops off abruptly below the erosional interfaces

so that the upper part of each cycle is underconsolidated. The percent of

sand-sized particles (mainly shell) in the mudflat samples from this core

ranged from near zero to 14 percent. Silt content varied between 20 and 50

percent, while clay made up the remainder, between 50 and 80 percent of the

total. The within-bed pattern of consolidation, apart from the trend which

characterizes the deposit as a whole, suggests that each of these couplets

represents sedimentation occuring during a single depositional event.

X-ray diffraction traces from the <1 micron fraction of mudflat

sediments subjected to the standard suite of laboratory treatments are shown


Reproduced

with perm

ission of the

copyright ow

ner. Further


without

permission.

Figures 12a-c. Bedding structure in radiograph positives of mudflat sequences

(light tones indicate higher density silt and shell); flat and lenticular

laminae distorted by loading and water escape (12a), flat-lying silt

laminae filling scour depressions in underlying erosional surface (12b),

lamina of disarticulated, imbricated M ulinea valves at erosional surface (12c).

43

C O A R S EF R A C T I O N

(% > 63 M) D E P T H

D E N S I T Y(g -c m " 3)

30

50

70

90

Figure 13. Texture and bulk density changes with depth in upper meter of newly deposited mudflat.

Reproduced with permission of the copyright owner. Further reproduction prohibited without permission./

44

in Figure 14. The major clay mineral groups identified were illite, smectite,

and kaolinite. As is shown in Table 2, the two analytical methods used to

obtain semi-quantitative estimates of the abundances of these minerals give

somewhat different results for the dominant expanding-lattice illite and

smectite groups.

The method of Griffin (1971) is similar to techniques which have been

used in previous studies of sediments in Atchafalaya Bay and the inner shelf

(Powers 1957, Johns and Grim 1958, Whitehurst et al. 1960, Ho and

Coleman 1969, Brooks 1970, Mobbs 1981, Van Heerden 1983). Mobbs

(1981) worked both in the Bay and just offshore, and found insignificant

differences in composition between the two locations. Comparing his results

with earlier work, he did, however, suggest a long-term trend toward an

increased illite fraction in the most recent deltaic sediments. Van Heerden

(1983) noted a similar trend in cores containing both bay bottom sediments

and those associated with the newly emergent subaerial delta. He further

observed some indication of an increased illite content in flood deposition

relative to low-flow sedimentation. Mean percentages for the < 2 micron

fraction from Mobbs (1981) are presented in Table 2, and show a somewhat

higher percentage of illite relative to smectite than do the comparable

mudflat results.

A mixed-layer illite-smectite group was identified in the mudflat

sediment using the more detailed method of Schultz (1964). This mixed-

layer clay makes up almost half of the smectite identified following Griffin

(1971). This mineral group can form when degraded illite enters seawater

and regains much of its original potassium content (Johns and Grim 1958).

The dominance of the transitional mixed-layer illite-smectite clay in


/

45

A IR -D RI ED

GLYCOLATED

3 3 0

5 5 0

10 15DIFFRACTION ANGLE (29 )

2520 3 0ANGLE

Figure 14. X-ray diffractograms of smears from the < 1ji fraction of mudflat sediments treated as indicated: S=Smectite, l=lllite,K=Kaolinite, and Q=Quartz.


/

46

Table 2 Clay Mineral Abundances

Chenier Plain Mudflat and Atchafalaya Bay

Clay Mineral Atchafalaya Bay <2 jim

Kao Unite

Il l i te

Sm ectite

M ixed-LayerIllite -S m e c tite

%Mobbs (1981)

16d)

53

31

Chenier Plain <1jim

Methods--------Griffin(1971) Schultz(1964)

— % ..............17 19

%

43

39

31

20

30

(^Includes Chlorite


47

sediments from the study area suggests a close connection between mudflat

deposition and the illite-rich Atchafalaya suspended load.

Results of viscometer runs on newly deposited mudflat sediment with a

bulk density of 1.26 g-cm-3 (sediment concentration = 416 g-l_l) and a

salinity of 22 °/OQ are summarized in the shear stress - strain rate plot in

Figure 15. The data describe an approximately linear relationship (r2 =

0.93) between the applied shear stress, T, and the rate o f deformation, du/dy.

The equation for this line describes the behavior of a Bingham plastic

T = T p + |i (du/dy) (1)

where Tp is the shear stress corresponding to the yield strength, and |i is the

sediment viscosity (Bingham 1922). Deformation ceases when the applied

shear stress falls below the yield strength. The yield strength, 171

dynes-cm"2, and viscosity of the mudflat sediment, 52 dyne-s-cm_2, are

determined by the intercept and slope, respectively, in Figure 15.

Krone's (1962) relationship between yield strength and sediment

concentration has the form

Tp = FC2.5 (2)

where C is the sediment concentration in g-1‘1, if Tp is measured in dynes-

cm-2. The experimental results indicate a value of 4.9 * 10"5 for the

proportionality factor, F. If Krone’s empirical relation holds, determination

of this factor allows estimation of yield strength for a range of sediment

concentrations. The line described by this equation is plotted on Figure 16

(adapted from McCave 1984) along with results for other marine muds

(Krone 1962,1963, Owen 1975, Hydraulic Research Station 1979).


48

5 0 0 i

4 0 0

/* = 51.7dyne-s-cm-2

S t 3 0 0OC o < >

Z 2 0) >9 2 0 0

100 Tp = 1 7 0 .9 dynes-cm -2

0 1 32 4s-1

DEFORMATION RATE(du/dy)

Figure 15. Shear stress-strain rate plot from capillary viscometer measurements made at 22° C on mudflat sediment with a density of 1.26 g -cnr3 (sediment conc. = 416 g-M ), and a pore water salinity of 22 °/00. Bingham plastic behavior is indicated by the presence of

a Bingham yield strength, T p , of 170.9 dynes-crrr2


49

1 0 0 0 .0

C9 1 0 0 .0

1.0

"S .

0.440 60 1002 0 2 0 0 40010

SEDIMENT CONCENTRATION (g-l-1 )

Figure 16. Fluid mud yield strength plotted as a function of sediment concentration (from McCave 1984). Data from Krone (1963), Owen (1970), and Hydraulics Research Station (1979) for a variety of marine muds plot with approximately the same trend (T p = F C 2-5). This trend is extrapolated from the mudflat data along

the dashed line shown.


/

50

DYNAMICS EXPERIMENT

The wind record for 10 through 13, December, 1982, in Figure 17,

shows the 180° shift in direction which marks the passage of a cold front in

the early morning hours of the 1 l^h . Wave characteristics, flow velocities,

and suspended sediment concentrations were monitored in the three sessions

indicated.

The pre-frontal regime began at midday on the lO^1 with the onset of

easterlies ranging between 10 and 15 knots. The steady pre-frontal regime

lasted 15 hours and includes the period of the first two instrument runs.

Starting at midnight, wind direction rotated in a clockwise manner from

easterly to northwesterly during the next 8 hours. Speeds averaged less than

10 knots during this shift, which we identify as the frontal period. A rapid

increase in velocity after runs 3 through 6, on the afternoon on the 1l^ 1,

signaled the onset of the post-frontal regime. Northwesterly winds

strengthened over the next 8 hours to sustained speeds between 25 and 35

knots. Eight hours later, velocity dropped to between 15 and 25 knots, but

continued steady from the north throughout the 12^ . Wind azimuth then

shifted clockwise a second time in the early hours of the 1 3 ^ and speeds

dropped off to 10 to 15 knots. By midday, during the period of runs 7

through 10, winds were from the east once again.

The effects of front passage were manifested in departures of the

observed tide record from that predicted, in changes in the gravity wave

regime, and in the distribution of suspended sediments above the bed. These

effects are described in the first three sections below. The final results

section is focused on the relationship between infragravity water level

oscillations and low pass flow velocities.


/

RUNS1 - 2

RUNS7-10100 PREDICTED

E3 50zO< MGL >UJmi

OBS ERVEDUJHI “50 O

\ - /

HZsN<

(0g 30Z*o 201UUJa(0a 10 z5

10 13

Figure 17. Winds and tides during cold front experiment, 1 0 - 1 3 December, 1982. Tide level recorded at the experiment site are compared with those predicted at the mouth of the Mermentau River (NOAA, NOS 1981). Wind data are hourly observations at Freshwater Bayou Canal Lock. The periods of the pre-frontal, frontal, and post- frontal instrument runs are indicated by stippling.


52

Meteorological Influences on the Tide Record

W ater level at Freshwater Bayou Canal is plotted along with that

predicted at the mouth of the Mermantau River in Figure 17 (NOAA, NOS

1981). This gauging station, located 70 km west of the Canal, is the closest

for which predictions are made. Although the mean tide range there exceeds

that at Freshwater Bayou by 0.1 m, a general comparison can be made.

W ater level is an average of 0.2 m above predicted for the pre-frontal period

on the 10th Predicted and observed elevations are in agreement on the 11 &

coincident with front passage. Following the onset of the strong

northwesterlies later that day and continuing through the 12th, Freshwater

Bayou levels average 0.3 m below predicted. On the 13^ , the water level

depression was reversed as easterlies again set in.

Waves and Oscillatory Flow

Four kilometers offshore of station 6, significant wave height during

the first 6 data runs (Pre-frontal and frontal periods) was estimated at 1 m.

The wave field at this time included a long-crested swell which approached

the shoreline from the south-southeast at a 5 to 10 degree angle of incidence.

Following front passage, wave height offshore was less than 0.5 m, but a

short-crested, confused sea made the dominant direction of approach

impossible to assess.

Wave data acquired at the experiment site are summarized in Table 3.

W ater depth ranged from 85 to 105 cm, and wave heights averaged less than

20 cm. These low-amplitude waves formed spilling breakers at a point 30 m

landward of the instrument array during the first two runs (pre-frontal

period). In all subsequent runs, waves were completely attenuated before

reaching the shoreline, and breaking waves were not observed.


Reproduced

with perm

ission of the

copyright ow

ner. Further


without

permission.

Table 3Sea and Swell Wave Statistics

Cold Front Experiment 1 0 - 1 3 December, 1982

Run No. Time FlowmeterOrientation

W aterDepth

d(cm)

Period T (s)

Height ^rms (cm)

Wavelength L (cm)

Orbital Vel. Urms (cm-s-1)

Swell/Sea Swell/Sea Swell/Sea Swell/Sea

Pre-Frontal 12/10 2 14:20 On/Off 85 6.8 3.0 10.4 10.2 2097 965 17 18

Frontal 12/113 13:07 On/Off 100 6.9 2.8 11.3 4.4 2174 829 22 84 13:38 Along 100 7.3 3.5 11.0 4.6 2242 1107 — —

5 14:23 Along 100 5.8 3.5 13.0 4.3 1806 1107 — —

6 15:09 On/Off 90 6.1 3.2 8.9 3.6 2000 976 20 9

Post-Frontal 12/137 10:51 On/Off 85 6.5 2.8 5.6 6.3 2067 829 9 118 11:30 Along 90 8.0 3.2 5.3 5.7 2397 976 — —

9 12:34 Along 100 7.3 3.4 5.1 4.8 2290 1025 — —

10 13:01 On/Off 105 6.8 2.5 5.0 4.9 2097 761 8 8

Mean 96 6.8 3.1 8.4 5.4 2130 953 15 11Std. Dev. 8 0.7 0.4 3.1 1.9 173 123 6 4

cnco

54

The effects of front passage on wave height and oscillatory flow can be seen

in the sequence of pressure and flowmeter spectra shown in Figure 18.

Spectra computed from representative 20 minute pressure and flowmeter

records acquired before, during, and after front passage all show energy

density peaks at two frequencies in the gravity wave range. These peaks

define distinct wave trains; swell with an average period of 6.8 s (0.1471 cps)

and locally generated seas of 3.1 s period (0.3226 cps). The energy

contributed by linear trends and infragravity oscillations of period greater

than 200 s (0.005 cps), combine to form a third peak which appears close to

the origin in all the spectra. Root-mean-square wave heights (Hrmg) and

oscillatory flow velocities (Um s ) calculated for the principal gravity wave

peaks (Thompson 1981) are also included in Table 3.

Two changes in the wave field associated with the transition from pre-

to post-frontal conditions are apparent. First, the total gravity wave energy,

consisting of the combined contributions of both sea and swell, decreases

through the sequence. Second, the distribution of energy between swell and

sea frequencies shifts significantly. Hnns and UnxiS in both bands are of

comparable amplitude in the pre-frontal runs. Energy associated with the

swell is retained during the second sampling period coincident with the shift

to northwesterly winds, while that of the seas is greatly diminished. Finally,

a reduced, residual swell and building seas characterize the final, post-frontal

records.

The differences in inferred wave height between seaward (A) and

landward (B) pressure sensors 10 m apart were less than 5 percent in all runs

where simultaneous data were obtained. Although it was noted qualitatively

that, except during the pre-frontal period, waves were completely attenuated


PERIOD (•)10 5

300 -IPRESSURE

SWELLSIa

FLOWMETER 6 0 CMRUN 2 RUN 2

■ SWELL200-I 400- SEA

100-1 200-1

3001 600

RUN 3 RUN 3

200 ■! 400-

200 -

600300-;

RUN 7 RUN 7

200 - 400 -j

1 0 0 - 200

0.25 1.00FREQUENCY (cps)

0.50 0.75 0.25 0.50 0.75 1.00

Figure 18. Representative energy density spectra from 20 minute pressure and flowmeter time-series for pre-, frontal, and post- frontal runs. Energy concentrations at the characteristic sea and swell frequencies are indicated.


56

before reaching the shoreline, comparison of water level spectra at the two

stations indicates that in four runs, wave height actually increased slightly in

the landward direction. Such an increase can be seen in spectra from run 3

plotted in Figure 19a.

Coherence and phase estimates computed from the cross-spectra of

simultaneous water level records from run 3 are also shown in Figure 19a.

The relationships between coherence, phase, and frequency shown in this

figure are representative of all the runs and provide information about the

incident waves which cannot be deduced from individual spectra. As is

expected, coherence between the records is highly significant through all

frequencies with appreciable energy. The strength of the correlation lends

credibility to the phase trend shown.

Below 0.005 cps, in the infragravity peak, the two records are in

phase. The gradual increase in lag with increasing frequency suggests that

the spectral peaks observed in the gravity wave range are associated with

progressive waves rather than standing waves or spectral harmonics

(Suhayda 1974). Given that the spacing between the sensors is known

exactly, the magnitude of the phase lag at the principal gravity wave

frequencies provides a direct measure of swell and sea wavelengths. Mean

wavelengths of the swell and sea calculated in this way were 21.3 m and 9.5

m, respectively (Table 3).

Spectra of flowmeter records obtained at 60 and 13 cm above the bed

during the same run are shown in Figure 19b. Energy densities at swell

frequency during this run were somewhat higher at 13 cm than at 60 cm, but

in most onshore-offshore runs, gravity wave energy was almost identical at

all levels. Coherence and phase estimates from the cross-spectra indicate that


/

10 sPERIOD (a )1 10 5

3 0 0 -l

O 2 0 0 -MlU

100

PRESSURE: RUN 3 ••a w a rd sanaor (A)

•horaw ard sanaor (B)

600

AOf

• VV

400

200i

i ife c J

FLOWMETERS: RUN 3* . . . . . 6 0 cm

— 13 cm

■ <»Jo-f v

1-0-1 1.0-1

ttluziusUlzoo

0.5-0 .5 -

i 0 . 0 -

(B) lags (A)

0.250.75 0.50 0.750.25 0.50 1.00FREQUENCY ( c p a )

Figures 19a-b. Concurrent spectra, and cross-spectral estimates of coherence and phase for records from run 3; between seaward and landward pressure sensors located 10 m apart (19a), and between near-surface and near-bottom flowmeters (19b).


58

oscillatory flow at the two elevations is highly coherent and in phase. These

observations suggest that oscillatory wave motion is reasonably

approximated by potential flow, at least down to 13 cm above the bed

(Komar 1976, p. 39).

A knowledge of wave height, water depth, and wavelength is sufficient

to classify surface gravity waves within the bounds o f specific wave theories.

As is shown in Figure 20, adapted from Komar (1976, p. 61), both sea and

swell fall within the range of Airy or Linear Wave Theory (LWT), well

below the line beyond which wave steepness is limited by breaking. Because

of their limited wavelength, the seas are considered "intermediate water"

waves even in 1 m water depth and lie outside the region where LWT shallow

water approximations are most valid.

Wavelengths determined from the cross-spectra average 6 percent

longer than those calculated using shallow water LWT, as is indicated by a

slope slightly less than unity in Figure 21a, but otherwise are well predicted

by this theory (r2 = 0.92). The oscillatory velocity data plotted in Figure 21b

also shows a good fit (r2=0.99) between Urms determined from spectra, and

the theoretical maxima predicted. Umax exceeds Unns, as is expected. The

62 percent difference is, however, an indication that orbital velocities

calculated from the spectra provide a somewhat inadequate representation of

oscillatory flow for the shoaling waves considered here (Cornish 1898). In

all runs, at all depths, an assymetry was observed between maximum onshore

and offshore velocities, both before and after correction for the low-

frequency mean flow component. Corrected peak onshore orbital velocities

exceeded offshore maxima by an average of 28 percent.


Lina of l imiting w i v « s t a a p n a s a

£ ' / S W E L L > v EA

H /d 0.1

AIRY WAVE THEORY

Airyd e e pw a t e r

Airy Airy’■N

sha l l ow 'o i n t e r m e d i a t e w a t e r w a t e r

0 .010 .01 0.1 1 10

Figure 20. The areas of application of the several wave theories as a function of H/d and d/L (from Komar 1976, p. 61). The regions occupied, within the span of Airy or Linear W ave Theory (LWT), by seas and swell monitored during the cold front experiment are shown in stippling.


60

(A)2 2

*y = 0 . 9 1 x ♦ 0 . 5 7

140 . 9 2

10

6 1 0 1 4 18 2 2 2 6WAVELENGTH FROM CROSS-SPECTRA (m)

Q

(B)

« 3 0

y = 1 . 6 1x ♦ 0 . 1 7

2 2 0 . 9 9

1 4 2 2 3 0

Urms FROM SPECTRA (cm -*’ 1 )

3 8 4 6

x<2

Figures 21a-b. Relationships between wave parameters determined from the spectra and cross-spectra, and those predicted by shallow water LWT; wavelength (21a), and orbital velocity (21b).


61

Suspended Sediments

Total suspended sediment concentration results from the multi-level

sampler are summarized for the pre-frontal, frontal, and post-frontal

periods in Table 4. Concentrations ranged from 160 to 9400 mg-l“l. Both

concentration and between-sample variability decreased with elevation above

the bed. The highest concentrations, which approached those of fluid mud,

were monitored at 10 cm above the bed during the frontal period (runs 3 - 6 ) .

Concentrations at this elevation, during these runs, averaged more than 10

times greater than during the pre-frontal period, and 3 times higher than the

post-frontal mean.

Ratios calculated between concentrations at the 65 and 10 cm levels are

shown in Table 4. Near-bottom concentrations (10 cm elevation) during the

frontal period were 550 percent higher than at 65 cm, but only 2 and 54

percent greater during the pre- and post-frontal periods, respectively. These

results indicate that a significant vertical gradient in sediment concentration

was developed in the frontal runs which was not present during the other two

periods. Representative concentration profiles from runs 2, 3, 5, and 7,

plotted in Figure 22, show the progression of gradient development through

the frontal sequence.

It is clear that suspended sediment concentrations are not directly

related to wave energy as the highest values near the water surface were

found during the post-frontal period when wave heights were lowest. One

interpretation of the distinctive character of the frontal profiles, relative to

those from the pre- and post-frontal periods, is that gradient development

may be linked to the frequency composition of the gravity wave energy,

rather than to wave amplitude. Ratios of sea to swell wave


/

Reproduced

with perm

ission of the

copyright ow

ner. Further


without

permission.

Table 4 Total Suspended Sediment


Run No. Time Water No. Height Ratio Mean Cone. (Std. Dev.) Cone. Ratiodepth Samples Sea/Swell sampled at 4 elev. 65cm/10cm

d (cm) 95 cm 65 cm 35 cm 10 cm....................... - m g - h 1 - - ..................................................Pre-Frontal 12/10

2 14:20 85 4 0.98 — 318( 53) 320( 45) 325( 56) 0.98(.03)Pre-Frontal Mean 4 0.98 — 318( 53) 320( 45) 325( 56) 0.98(.03)

Frontal 12/113 13:07 100 3 0.39 205(28) 256( 92) 350( 61) 3213(3531) 0.13(.07)4 13:38 100 2 0.42 304(21) 355( 4) 634(232) 4650(2792) 0.09(.06)5 14:23 100 3 0.33 360(54) 341 ( 25) 362( 11) 5666(3597) 0.09(.06)6 15:09 90 3 0.40 337( 95) 515(129) 1155( 358) 0.30(.09)

Frontal Mean 11 0.39(.04) 290(99) 317( 76) 443(152) 3582(3037) 0.16(.11)

Post-Frontal 12/137 10:51 85 6 1.12 820( 61) 901 ( 82) 1355( 398) 0.66(.21)8 11:30 90 6 1.08 582( 46) 618( 86) 1131(455) 0.61 (.28)9 12:34 100 6 0.94 413(12) 426( 16) 439( 29) 732( 250) 0.65(.23)

10 13:01 105 6 0.98 431(20) 477( 20) 466( 18) 773( 391) 0.69(.27)Post-Frontal Mean 24 1.03(.08) 425(19) 568(168) 606(196) 998( 443) 0.65(.23)

Grand Mean 96 39 0.74(.34) 364(93) 473(185) 533(200) 1658(2020) 0.54(.33)

CDro

100

5 0

Eo

aUi“ 10 iu > o m <

tTl fcml

RUN 2

— ,— imu

RUN 3

* Vinmi~lQmi|lmjni|i

RUN 5-i 50RUN 7

J ,li minmfan'lilfl.nnmri101000 10000100

SUSPENDED SEDIMENT ( m g - f1)

Figure 22. Mean suspended sediment concentrations from representative pre-frontal (2), frontal (3 and 5), and post-frontal (7) runs. Error bars show two standard deviations about the mean.


64

heights, given in Table 4, provide one measure of the relative importance of

these components of the wave field. When the near-surface/near-bottom

concentration ratios (65cm/10cm) from each run are plotted against the

respective wave height ratios, the data from the pre- and post-frontal periods

form one cluster distinctly separate from a second frontal cluster, as is shown

in Figure 23. The concentration ratios exhibit large within-run variability,

as is indicated by the error bars in the figure. Within-run variability in the

wave height ratios cannot be estimated. Even so, when all the data are

considered together, a weak linear relationship emerges (r2 = 0.69). This

result suggests that swell waves may contribute to the formation of a

suspended sediment gradient, which does not develop if higher frequency

seas make up an important part of the wave field.

Infragravitv W ater Level Changes and Non-Oscillatorv Flow

Energy contained in the low-frequency peak below 0.005 cps

constitutes between 30 and 60 percent of the total variance in the water level

data, and from 7 to 22 percent of that in onshore-offshore flowmeter

records. An average of 54 percent of the low-frequency variance in the

pressure records is contributed by uni-directional tidal trends in water level

which are long with respect to the 20 minute sample length. These can be

approximated by a straight line and are readily removed using a linear

detrending procedure. Non-tidal water level fluctuations contribute the

remaining low-frequency variance. As was noted earlier, concurrent water

level spectra obtained 10 m apart differ by less than 5 percent in the gravity

wave frequencies, and, overall, indicate no consistent cross-shore pattern

with respect to wave height. In contrast, a systematic amplitude increase in

the direction of the shoreline is observed below the gravity wave range. The


/

o5 1 - °oczoj= 0.8<OCz Eg S 0 .6 -

siO o

IA <0

O*£<<>0.4UJ2Ew 0.2</>Q UJa zUJ CL CO

1.00 .2 0.4 0.6Hsea/Hswei!

1.20.8

WAVE TYPE RATIO

Figure 23. Relationship between vertical sediment concentration gradient, indicated by the ratio between near-surface and nearbottom values, and the relative significance of sea and swell, indicated by the ratio of their respective heights determined from the spectra. Circled numerals correspond to run number, and error bars indicate two standard deviations about the mean concentration ratio for each run.


6 6

sum of variances in spectral frequency bands below 0.005 cps from the

detrended water level records are listed by station in Table 5.

Between-station differences provide a relative measure of the

landward amplification of the non-tidal, low-frequency water level changes

occurring during each run. This amplification was less than 5 percent in 2 of

the 7 runs from which data at both (A) and (B) are available, but ranged from

11 to 33 percent in the remainder, and averaged 12 percent overall.

Concurrent water level records are in phase at frequencies below 0.005 cps,

as the cross-spectral results in Figure 19a show for run 3. An amplitude

increase in the direction of the shoreline, without a progressive phase shift

over the same distance, is characteristic of a standing wave with an offshore

node and an antinode at the shoreline (Suhayda 1972, Waddell 1973).

Attention is drawn to the variance values for run 3 in Table 5 which

exceed the means for all runs by nearly 300 percent. Low-pass wave and

flowmeter data (T > 200 s) from run 3 are shown in Figure 24 and clearly

show the source of this energy. Mean water level drops gradually in the first

half of this segment, then abruptly rises 10 cm in 340 s. During this period of

rapid water level rise, flowmeters at all depths register significant onshore

flow. Most of the low-frequency adjustment in flow velocities occurs away

from the bed at the upper flowmeter elevations. The records shown in

Figure 24 clearly establish that coherent onshore-offshore flows lasting for

minutes are associated with non-tidal fluctuations in nearshore water

elevation. Local and regional wind data of comparable resolution are

lacking. It appears, however, that the wind-forced sequential water level set

up and set-down which characterizes the frontal tide record is reflected in

higher-frequency changes observed in 20 minute records.


Table 5Water Level Variance Below 0.005 cps

in Concurrent Detrended Records from Stations A and B10 m Apart

Run Variance (cm)2 D ifference %No. --------------------

Station A Station B

P re -fro n ta l2 1.59 1.78 +0.19 + 12

Frontal3 7.74 8.63 +0.86 + 11

4 1.18 1.21 +0.03 + 3

5 1.54 1.71 +0.17 + 11

6 1.20 1.21 +0.01 + 1

P o st-fro n ta l7 0.87 1.16 +0.29 + 33

8 1.02 1.14 +0.12 + 11

Mean Std. Dev.

2.162.47

2.412.76

+0.240.29

+ 12 10


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ner. Further


without

permission.

RUN 3 LOW PASS

60cmT "O N S H O R E

IEo sr- 30cma111IIIQ.0)

1 3cmO F F S H O R E

300 1050750 900600 TIME (s)

450150

F igure 2 4 . Low freq u en cy w a ter le v e l and o n sh o r e -o ffsh o r e flow m eter d a ta from frontal run 3 . Low p a s s w ater le v e l is su p e r p o s e d on d e m e a n e d p r e ssu r e record in upper fram e. Low p a s s flow m eter v e lo c it ie s ab o u t a true z e r o point a re sh o w n in th e low er fram e. Logarithm ic flow p er io d s stip p led .

O)00

69

The existence of a connection between infragravity water level

changes and cross-shore flow is supported by analysis of the flowmeter

variance statistics presented in Table 6. Flowmeter variance less than 0.005

cps increases with elevation above the bed. Covariances were computed

between the low-frequency velocity time-series and the low-pass water level

records. Then the same comparison was made between the velocity data and

the time-series which results from evaluating the change in water level

between each time step. This simple derivative of the low-pass water level

time-series exhibits positive values when water level rises, and negative

values when it falls.

Change in the rate of water level rise and fall, represented by the

derivative time-series, is positively correlated with the low-frequency

flowmeter records and explains, on average, 60 percent of the variance in

velocity below 0.005 cps. The undifferentiated water level record explains

only 22 percent of the low-frequency onshore-offshore flow. The poor

correlation between untreated water level and velocity is not unexpected

given that maxima and minima in water level represent end points which are

out of phase with cross-shore flow. Physically, as the surface elevation

fluctuates about a mean position, high and low water stands are both periods

of diminished shore-normal transport.

The covariance data in Table 6 indicates that the correlation between

the rate of water level change (derivative) and shore-normal velocity is far

stronger at the 60 and 30 cm levels than it is closer to the bed, at the 13 cm

level. The reason for the decreased correlation there can be seen in Figure

25. This figure shows a progression of velocity profiles extracted at 50 s

intervals during the low-frequency flow event shown in Figure 24 (run 3).


Table 6Non-Oscillatory Onshore-Offshore Flow:

Covariance Between Water Level and Flowmeter RecordsBelow 0.005 cps

Run Flowmeter Flowmeter % Flowmeter Variance No. Elevation Variance Explained by:

(cm) (cm-s-1 )2 Untreated Water Level

Derivative Water Level

2 60 4.5 - 16 + 7030 3.4 - 15 + 7013 2.4 - 7 + 62

3 60 3.9 - 4 + 9030 2.6 - 6 + 8813 1.2 - 60 + 49

6 60 3.0 + 10 + 5830 2.2 + 15 + 6313 0.9 - 1 + 49

7 60 3.9 - 29 + 7630 2.9 - 22 + 6613 1.2 - 20 + 43

10 60 2.9 - 42 + 4530 N/A N/A N/A13 0.7 - 41 + 19

Mean 2.6 - 22 + 60Std. Dev. 1.2 17 19


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ner. Further


without

permission.

80

708080

40

30

20ouoUl>oD<Zo

70I-<>111_]ui

60

6040

3020

T=675»10

-3 -2 1 0 21 43

O F F 8 H O R E O N S H O R E

T=575 8 T=626 s

1 30 1 42

V E L O C I T Y C C M - S - 1 )

!

Figure 25. Sequence of vertical velocity profiles from 525 to 775 s during onshore-offshore flow event in run 3 (see Fig. 24).

72

The sequence demonstrates that both onshore and offshore transport begin as

flow reversals close to the bed. This observation lends credibility to the

standing wave concept introduced earlier, as near-bottom flow reversals are

characteristic of transport predicted under both standing (Longuet-Higgins

1953) and partially reflected progressive waves (Carter et al. 1973). The 20

minute length of the records obtained during this study does not permit

resolution of the periodicity of the low-frequency water level fluctuations

observed, if, indeed, any such periodicity exists. Establishment of periodic

forcing is not required with respect to the kinematics, however, as any

perturbation of the free surface will involve a redistribution of mass

requiring shore-normal currents. These currents emulate those developed

under a true standing wave characterized by a particular node position and

frequency.

Spectra from flowmeters oriented alongshore have a significant peak

only at frequencies below 0.005 cps as the ducts eliminated an estimated 95

percent of the gravity wave signal. Alongshore flowmeter statistics compiled

in Table 7 lead to three important observations. First, alongshore flow was

to the west throughout the course of the experiment. Second, mean velocities

remained relatively constant, ranging from 7.6 to 8.1 cm-s"l at the 60 cm

elevation, through the frontal and post-frontal sequence when wave heights

dropped significantly. Finally, mean flow velocities in all runs increase with

elevation above the bed.

Low-pass wave and flowmeter data from run 5 are shown in Figure

26. The dependence of velocity on flowmeter elevation is more pronounced

in this alongshore data than in the onshore-offshore record from the same

period (Fig. 24). W ater level drops through most of this segment but rises


/

73

Table 7 Steady Alongshore Flow:

Covariance Between Water Level and Flowmeter RecordsBelow 0.005 cps

Run Flowmeter Mean % Flowmeter VarianceNo. Elevation Westward Explained by:

Velocity ......................................(cm) (cm-s-1) - Untreated

Water LevelDerivative

Water Level

4 60 7.8 + 52 + 1130 4.5 + 49 + 1013 1.7 + 56 + 11

5 60 7.9 + 57 + 3530 5.6 + 66 + 4013 1.3 + 81 + 24

8 60 8.1 + 54 + 2030 6.2 + 48 + 1013 3.0 + 40 + 3

9 60 6.0 + 86 - 1630 N/A N/A N/A13 1.7 + 82 + 24

Mean 4.9 + 61 + 19Std. Dev. 2.6 16 11


Reproduced

with perm

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without

permission.

RUN 5 LOW PASS

30cm

n. ■-w o -----------------------------1-----------------------------1-----------------------------1-----------------------------1-----------------------------1---------------------------- 1-----------------------------1-----------------------------1-----------------------------< ■ ■■ i wISO 300 450 600 750 900 1050

TIME (s)Figure 26. Low frequency water level and alongshore flowmeter data from frontal run 5. Low pass water level is superposed on demeaned pressure record in upper frame. Westerly low pass flowmeter velocities about a true zero point are shown in the lower frame. Logarithmic flow periods indicated by stippling.

■vl

75

toward the end. Flowmeter velocities generally mirror the water level

curve, decreasing when water level falls, and increasing when water level

rises. As before, covariances were computed first between the low-

frequency flowmeter time-series and the untreated water elevation record,

and then between the velocity data and the derivative water level time-series.

For the alongshore runs, the untreated water elevation data is positively

correlated with velocity (positive to the west) and explains an average of 61

percent of the low-frequency flow variance. The derivative of the water

level time-series explains, on average, only 19 percent of the flowmeter

variance below 0.005 cps. Physically, the positive correlation between

alongshore flow velocity and water level means that maximum and minimum

velocities occur at high and low water stands, respectively. A rise in water

level, then, results in an increase in westerly alongshore flow velocities at the

instrument array.

The depth dependence of the mean alongshore velocities in Table 7

suggests that the alongshore current might be well characterized as a

boundary shear flow. The physical constraint that velocities must be zero at

the bed gives rise in a neutrally stable flow to a boundary modified layer in

which mean steady horizontal velocities typically approach free-stream

values as a logarithmic function of elevation above the bed (Adams and

Weatherly 1981a). The equation which expresses the relationship of

horizontal velocity to depth in the bottom boundary layer has the form,

u* zU = — In — (3)

K zo


/

76

where U is the velocity measured at a reference height z above the bed and K

is von Karman's constant (0.4). This profile is fitted to the data by

minimizing the square of the velocity deviations with respect to a characteristic shear velocity, u*, and an apparent bottom roughness length,

z0 (Wilkinson 1984).

In the present investigation, low-pass velocities averaged over 5 s

intervals at each of the three flowmeter elevations were fit to equation (3) to obtain 240 profiles, with associated estimates of u* and z0, for each 20

minute record. A profile was defined as logarithmic when A , the coefficient

of determination, exceeded 0.9755, the t-test criterion for a 90 percent

confidence level in a three point fit (Adams et al. 1982). Low-frequency

onshore-offshore flows met this test between 5 and 25 percent of the time.

Alongshore flow, in contrast, was logarithmic for 65 to 85 percent of the

runs where three point velocity data was available (runs 4, 5, and 8), and was

close to logarithmic most of the rest of the time. Periods of significant

logarithmic flow are shown in the onshore-offshore record in Figure 24 and

in the alongshore data in Figure 26.

A subsample of velocity profiles meeting the logarithmicity test was

selected from periods of steady flow when the mean current was undergoing

neither acceleration or deceleration. It is important to note here, for

consideration later, that no periods of steady flow were found in the onshore-

offshore data. Ten steady flow profiles with r^ values above 0.99 were

identified, however, in each of alongshore runs 4, 5, and 8. Mean values of

u*, z0, and the velocities at each reference elevation derived from these

subsamples are given in Table 8 and were used to generate characteristic

steady flow profiles for each of these runs.


/

77

The characteristic velocity profiles are shown in Figure 27. In these semi-log plots the y-intercept corresponds to z0, the apparent bed roughness,

while the slope is K/u*. The profiles for runs 4 and 5, which were obtained

less than 1 hour apart, are similar, while that from run 8 in the post-frontal

period differs significantly. At 60 cm, velocities in runs 4 and 5 exceed that

for run 8, while at 30 and 13 cm the reverse is true. Values for z0 in all of

these profiles suggest the presence of bottom roughness elements 2 to 5 times

the amplitude of any actually found on the smooth mud bed at the experiment

site (2 cm maximum).

In briefly reviewing the results of the cold front experiment, we note

that both sea and swell waves at the study site were well described by shallow

water LWT. The major effect of front passage on the wave field appears to

be a modulation of the relative importance of the high-frequency sea

component. A significant vertical gradient in suspended sediment

concentration was developed during the frontal period under essentially

monochromatic swell waves (runs 3 - 6). This gradient was not present

during the pre- or post-frontal runs, when seas contributed more

significantly to the spectral energy distribution. W ater level oscillations in

the infragravity region were observed to be correlated with unsteady cross

shore flows. The low pass alongshore current, in contrast, was found to be a

fully developed, and relatively steady, boundary shear flow.


78

Table 8Characteristic Alongshore Velocity Profile Parameters

Run Tim e Water u* z0 Velocity %— at— Log

No. Depth 60cm 30cm 13cm.(cm) (cm-s-1) (cm) (cm-s-1)

Frontal Period 12/114 13:38 100 1.9 10.1 8.3 5.0 1.2 85

5 14:23 100 1.8 10.1 8.0 4.8 1.1 65

Post-Frontal Period 12/138 12:34 85 1.2 5.1 7.4 5.3 2.8 80


79

100

— RUN 4in O m RUN 5

r u n 3

aUJaUJ>Oa<zo<>UJmlUJ

0 2 4 0 8 10 12 14.1

VELOCITY

Figure 27. Characteristic logarithmic alongshore velocity profiles for frontal runs 4 and 5, and post-frontal run 8.


D ISC U SSIO N

The results o f this study show that mudflat deposition has ceased east of

Freshwater Bayou Canal, but remains active in the western half of the study

area. Mudflats form there only in late winter and early spring, and are

characterized in cross-section by distinctive 2 to 10 cm thick beds displaying

both grading and flat-lying basal silt or shell laminae. The goal of

monitoring nearshore dynamics over a broad range of wave energy

conditions was achieved during the December cold front experiment. Fine

grained sediments were in motion during this event, but sequential bed

profiles conducted in the course of the experiment showed no significant

changes indicative of accretion or erosion. A conspicuous lack of

sedimentation is not surprising given the intermittency which characterizes

mud deposition in this area. It is, however, an important indication that all

factors necessary for nearshore mud deposition were not present at the time

of the experiment.

Significant water level set-up and set-down were associated with the

characteristic clockwise rotation in wind direction which accompanies front

passage. These effects were clearly manifested in departures from the

predicted tide, but also showed up as infragravity oscillations in 20 minute

nearshore water level records. These oscillations exhibited characteristics of

a standing wave with an antinode at the shoreline. Significant low-frequency,

onshore-offshore flows were correlated with these shore-amplified water

level changes and support an assumption of standing wave motion.

The sea and swell waves monitored during the cold front experiment

were found to be well described by shallow water linear wave theory, but

80


81

also experienced the nearshore attenuation documented by Wells and

Coleman (1981a, b) on the muddy coast of Surinam. A significant near

bottom gradient in suspended sediment concentration was observed under

monochromatic swell waves, at the time of front passage, which was absent

when higher-frequency seas contributed significantly to the total wave

energy. Mean alongshore velocities near the surface responded to low-

frequency changes in the depth of flow at the flowmeter array, rather than to

changes observed in the incident wave field. This current consistently set to

the west in spite of the sequential shifts in wind direction which characterized

the cold front progression.

The vertical velocity distribution of the alongshore current, unlike that

of the the low-frequency shore-normal transport, was indicative of a steady,

logarithmic boundary shear flow. Apparent bottom roughness estimates (z0)

were, however, 2 to 5 times greater than the amplitude of any physical

roughness elements present at the experiment site. A marked decrease in both apparent bottom roughness, and the characteristic shear velocity (u*),

was noted in the transition between frontal and post-frontal profiles.

Now that we have some information about the nearshore dynamics, it

is appropriate to re-examine in more detail the boundary layer sorting model

(B-L-S model) proposed by Stow and Bowen (1980). Their representation

of the near-bottom dynamical environment is reproduced in Figure 28. Stow

and Bowen (1980) inferred, on the basis of sedimentological data alone, that

the origin of graded and laminated silt + mud beds must depend on a

fundamental difference in the way suspended cohesive and non-cohesive

sediments respond to boundary layer shear. Differential settlement of these

two types of particles in a waning flow could not explain the within-bed

Reproduced with permission of the copyright owner. Further reproduction prohibited without permission

SEA SURFACE

— z = h

TURBIDITYCURRENT

SUSPENSION

Nl

SHEAR

BOUNDARYLAYER10-

>VISCOUS SUB- { \ LAYERlam inarSUB-LAYER

SETTLING

Figure 28. B-L-S model for sorting of cohesive and non-cohesive sediments in the boundary layer of a turbidity current (from Stow and Bowen 1980). Elevation, z, measured from the bed in a millimeter logarithmic scale.


83

sorting they found in fine-grained, shelf margin turbidite sedimentation, as

clay floes and silt grains with equivalent settling velocities would be expected

to deposit together. According to the B-L-S model, initially only silt grains

can settle through the high-shear zone to form the characteristic basal

laminae, while clay floes approaching the bottom are disaggregated and

returned to the flow interior. Stow and Bowen (1980) suggest that clay

concentrations build up above the boundary shear layer to the point that

aggregates form which are sufficiently large to "overcome" shear. At some

critical concentration, they penetrate the high shear zone, "break-up", and

deposit rapidly to form the characteristic mud "blanket" overlying the silt

laminae. Following deposition of a silt + mud couplet, the process is repeated

at the new surface, such that a sequence of stacked laminated and graded beds

is developed.

Observations made during the course of the present study indicate that

the first part of this process, the selective passing o f non-cohesive particles

through the shear zone, could well explain the flat-lying basal laminations

found in mudflat sedimentation. It is not clear, however, how large clay

floes, which initially could not survive passage to the bed, could later resist

disaggregation by the same shear stress. An alternative mechanism for

deposition of the mud part of the silt + mud couplet is proposed here which

does not require any diminution of boundary layer shear. This approach is

founded on the assumption that the mud suspension supported above the high

shear zone acquires yield strength as its sediment concentration rises.

Sudden deposition might then occur, as suggested by Einstein (1941), without

break-up o f the floes, through a "freezing" process, similar to that observed

in mudflows on land (Johnson 1970). Freezing would be initiated when the


yield strength of the suspension'reached equivalence with the boundary layer

shear stress. As yield strength is a property which characterizes fine-grained

sediment suspensions only at fluid mud concentrations, deposition by the

mechanism proposed requires the presence of a fluid mud layer above the

high shear zone.

The dynamics data collected in the present study permit a more

detailed investigation of parameters relevant to boundary layer sorting than

has been possible in the past. In the first section below, we address the

problem, also recognized by Einstein (1941), of how turbulence, generated

in this case primarily by wave shear at the bed, is prevented from

propagating into the flow interior, where it would prevent development of

the significant near-bottom sediment concentrations required. This analysis

establishes a potential for the intact transport of a fluid mud layer by a uni

directional current. In the second section below, we consider the potential

for nearshore deposition of fluid mud transported along the unusual cross

shore shear stress gradient developed under progressively attenuated waves.

A scaling argument is advanced to explain why deposition did not occur

during the cold front investigated, and, more importantly, when it could.

This leads, in the third section, to a discussion of the temporal and spatial

connections between coastal mud deposition, the shelf sediment dispersal

system, and the Atchafalaya River. Predictions for the future of the chenier

plain are made in a fourth part, based on projections of Atchafalya delta

development. Finally, generalizable features of the chenier plain model are

extrapolated to other muddy coasts, both modem and ancient.

permission of the copyright owner. Further reproduction prohibited without permission.

85.

FORMATION AND TRANSPORT OF FLUID MUD

Implicit in the simple logarithmic boundary layer model defined by

Equation (3) is the assumption that a single boundary layer length scale is

sufficient to fully characterize interaction between the flow and the bottom

(Tennekes and Lumley 1972, p. 146). Elevated estimates of the bottom

roughness, such as those derived from the alongshore velocity profiles in this

study, are common in shear flows where additional length scales play a role

(Adams and Weatherly 1981a). Multiple scales arise when the flow is

influenced by turbulence from sources other than that due to the steady

current boundary shear, or when turbulence production is restricted by a

density gradient.

Turbulence generated by unsteady oscillatory wave flows at the bed

gives rise to a second, shorter length scale, that of the wave boundary layer

(WBL), which must be considered in addition to the scale associated with the

steady current boundary layer. Grant and Madsen (1979) have developed a

widely applied method for parameterizing the effects on the steady current

which result from combined wave and current interactions in the WBL.

Their analysis, which will not be repeated here, indicates that this interaction

is non-linear and results in estimates of bottom shear stress which are greater

than the sum of wave or current shear stresses treated separately. The steady

current in the region above the WBL, in which wave motion is characterized

by non-turbulent potential flow, experiences a resistance which depends not

only on the physical bottom roughness, but also on characteristics of the

WBL. The velocity profile above the WBL has a form similar to that given

in Equation (3):


86

u* 30z

U = — In —

K ZcW

Z > 8yy (4)

where the only new term is ZcW, an "apparent roughness length" containing

the combined effects of the physical roughness (z0), and that due to the

presence of a WBL of thickness Sw. Inside the WBL, the appropriate

roughness is once again z0, the amplitude of the bottom relief, and the

velocity profile is determined by:

u* u* 30z

U = — ( — ) In — z < 8W (5)

K u*cw zo

where u*cw is the characteristic turbulence intensity inside the W BL due to

the nonlinear wave and current interaction. Maximum shear stress at the bed

due to both waves and the current is given by:

where p is the density of the fluid. The thickness of the WBL is estimated as:

2K(U*CW)

(0

WBL thickness is thus proportional to u*cw and inversely proportional to co,

the radian wave frequency.


Tew,max = P(u*cw)^ (6)

(7)

87

Values for u*, u*Cw> zcw> zo> Tcw,max> and §w are determined from

the velocity at a reference elevation in an iterative closure dependent on a

small number of wave and current parameters (See FORTRAN listing in

Appendix B). Results of this analysis, based on current velocities at 13 cm

and swell parameters for the three alongshore, cross-wave runs in Figure 27

are given in Table 9.

Calculated WBL thicknesses (8W) for the runs range from 1.5 to 2.0

cm, roughly 2 orders of magnitude less than that of the current boundary

layer, which can be considered to cover the full depth of the flow (100 cm). The characteristic turbulence intensity in this limited region, u*cw, is nearly

an order of magnitude greater than that determined in the model for the flow

above the WBL, from 1.5 cm-s-! in run 8 to 2.2 cm-s-1 in run 4. The

estimated bed shear stress due to the combined wave and current motion

ranges from 2.1 to 4.6 dynes-cm-2 for runs 8 and 4, respectively.

Calculated WBL shear velocities are competent to suspend fine to

medium sand (McCave 1984). More significantly, bed shear stresses equal or

exceed Bingham yield strengths experimentally determined by Krone (1962,

1963) and Owen (1975) for fluid mud suspensions in the 20 to 100 g-l“l

range (Fig. 16). These shear stresses are, however, far below those required

to maintain in suspension a slurry with the concentration of the surficial

mudflat sediment, which, the results of this study suggest, has a yield strength

greater than 100 dynes-cm-2 (Fig. 15).

The large magnitude of the wave-induced flow, relative to that of the

steady current, leads to estimates of an apparent roughness length, zcw,

which range up to 36 cm (Table 9). This is the amplitude of physical

roughness elements which the model predicts would produce the same drag


/

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ner. Further


without

permission

Table 9Grant-Madsen Boundary Layer ModelO)

Wave Effects on Steady Alongshore Flow


-cwRun Depth Velocity Maximum Oscillatory Wave u* u*cw ;No. at 13 cm Orbital Excursion Freq.

ref. elev. Vel. Amp.(2) 05 ----------------(cm) (cm-s-1) (cm-s-1) (cm) (rad-s-1) (cm-s-1).......................Model Input........................ ........

Frontal Period 12/114 100 1.2 18 20 0.87 0.23 2.20 0.02 36

cw.max 'w

(cm) (dynes-cnr2) (cm) — Model Output................

100 1.1 18 20 1.08 0.23 2.12 0.02 34

4.6

4.3

2.02

1.58

Post-Frontal Period 12/13 8 90 2.8 10 11 0.79 0.40 1.47 0.02 11 2.1 1.48

OIGrant and Madsen (1979) (2)From LWT

OOoo

89

on the steady current as the combination of physical roughness, and the non

linear interaction of oscillatory and non-oscillatory flows in the WBL. This

resistance is transmitted to the flow interior in calculated values for u* (Table

9) which are nearly an order of magnitude lower than estimates derived from

the three-point velocity measurements given in Table 8. As a result, the

Grant-Madsen model predicts, for all the alongshore runs, velocity profiles

far steeper than the characteristic profiles shown in Figure 27. This

observation suggests that the current does not "feel" the drag generated in the

WBL, at least to the degree predicted by the model. One explanation for this

result may be that W BL turbulence, while important for maintaining

sediment in suspension close to the bed, is largely destroyed before it reaches

the interior of the flow.

An assumption of neutral stability, or the lack of significant turbulence

damping features in the flow, is implicit in both boundary layer models introduced thusfar. The increased slope (decreased u*) and lower z0 which

characterize the post-frontal velocity profile (run 8) in Figure 27, relative to

those from the two earlier runs is strongly suggestive of a change from a

stably stratified flow to one that is more neutral. Gradients in suspended

sediment concentration, like those due to salinity and temperature, can

produce a stable density field capable of inhibiting the vertical eddy diffusive

flux of turbulence (Smith and McLean 1977a). The suspended sediment

profiles shown in Figure 22, and the more complete data in Table 4, clearly

indicate the presence of a significant near-bottom concentration gradient

during runs 4 and 5, which was developed to a far lesser degree during run 8.

Gravity acts on the suspended sediment particles to oppose the upward

flux of turbulence produced by shear at the bed. The flux Richardson


90

number provides a dimensionless measure of the relative magnitudes of these

forces. Smith and McLean (1977a) give the appropriate form of this ratio as:

(ps-p) dCs duR f = - g r~ ~ ■ —1(~)-2 (8)

p dz dz

where g is the force of gravity, dCs/dz is the volumetric sediment

concentration gradient (z positive up), du/dz is the horizontal velocity

gradient, and ps and p are the respective densities of the sediment and the

fluid. It can be seen that the gravity force, which acts on the excess mass of

sediment, represented by the bracketed term, is balanced against turbulence

produced by the velocity gradient. A flow characterized by a positive Rf

number is stably stratified, and acts as a sink for turbulent energy. Rf values

calculated for the velocity and suspended sediment gradients between 13 and

30 cm elevations during the alongshore runs range from +0.01 for run 8 to

+0.10 in run 5. At values of Rf greater than approximately +0.20,

laboratory experiments show that turbulent flux is completely suppressed

(Tennekes and Lumley 1972, p. 99).

Einstein and Krone (1962) found that differential consolidation of

fluid mud, a process which reduces water content in the lower part of a layer

at a more rapid rate than in the upper part, results in a characteristic increase

in sediment concentration with decreasing elevation above the bed. If fluid

mud was present below the 10 cm elevation o f the bottom-most sampler

during runs 4 and 5, and turbulence associated with the wave motion is generated within a W BL 1 to 2 cm thick, the magnitude of the R f number

might be expected to increase at least to the top of the WBL. The effect of


91

such a stable density gradient will be to limit the vertical flux of turbulence

generated at the bed and thereby restrict sediment dispersal into the flow

interior.

One consequence of a restriction on the vertical flux of turbulence is

the emergence of a mechanism whereby a coherent fluid mud layer might be

transported in the direction of the mean flow. As Einstein and Krone (1962)

argued cogently 24 years ago, the vertical density and yield strength

gradients which characterize such suspensions render their transport en

masse by a uni-directional current alone improbable. Shear stress applied

only at the upper surface would initiate movement there first, with the result

that sediment would be stripped off and dispersed, while the denser

suspension closer to the bed remained stationary. Returning once more to the

B-L-S model (Fig. 28), however, we note that a sorting mechanism which

acts to segregate silt particles from a much larger population of clay floes in

the mixed suspension above the high shear zone must cause a break in the

sediment concentration gradient close to the bed. Given the relationship

between sediment concentration and yield strength, a concentration drop

imposed by shear at the bottom of a fluid mud layer would also be a point of

reduced yield strength. Transport of an intact fluid mud layer might then

occur, under the influence of shear exerted at its upper margin, along a shear

plane defined in 2 dimensions by the drop in yield strength at the top of the

high shear zone, in this case the WBL.

KINEMATICS OF FLUID MUD DEPOSITION AT THE SHOREFACE

Having developed the potential for transport of sediments contained

within a fluid mud layer under the influence of a unidirectional flow, the task

is now to identify such flows nearshore, particularly those with an onshore


92

orientation. A re-examination o f the characteristic alongshore velocity

profiles (Fig. 27) suggests, however, that near-bottom current velocities in

these steady flows are so low that significant transport of sediment in a fluid

mud layer would be unlikely. All nearshore flows monitored during the cold

front experiment were not steady, however, and would not be expected to

reflect this equilibrium condition. This is particularly true of the shore-

normal flows observed, which, though sometimes logarithmic, continually

undergo acceleration and deceleration. This evolution is apparent in the

sequence o f velocity profiles from the low-frequency flow event captured

during run 3 (Fig. 25). Unlike the characteristic alongshore profiles,

velocity profiles associated with shore-normal flow suggest a potential for

non-equilibrium sediment transport close to the bed.

As noted previously, the low-frequency fluctuations in water level,

which are correlated with significant unsteady onshore-offshore flow, have

characteristics of a standing wave with an antinode at the shoreline and an

offshore node. Suhayda (1972) has shown that, for a linear wave field on a

sloping bottom, the frequency of a standing wave which has a node X meters

offshore is given by:

2.4 (g tan <X>)l/2f ------------------------------------------------------------ ( 9 )

4 tcX1/2

where tan O is the bottom slope. The period of the low frequency oscillation

in run 3 (Fig. 24) is estimated to be 600 s (0.0017 cps). The bottom slope is

0.004. Equation (9) can be rearranged to solve for the distance of the node

from shore,

with permission of the copyright owner. Further reproduction prohibited without permission.

93

5.76 g tan OX = -------------- (10)

(4 7 i f ) 2

which, in this case, is found to be 400 m.

It was observed that the amplitude of the low frequency oscillations

increases in the direction of the shoreline or antinode. In the present

example, the amplitude (H/2) of the low-frequency oscillation at sensor B

was 4.7 cm. The amplitude at sensor A, 10 m seaward, was 0.17 cm less, a

value which yields a maximum between-sensor slope of the still water

surface, tan N, of 0.00017. When this slope is linearly extrapolated offshore,

the node position is predicted to lie 369 m from the shoreline, a value in

reasonable agreement with the result from Equation (10).

The shoreface geometry at station 6 is shown schematically in Figure

29. A control volume approach was adopted to estimate the depth and time-

averaged velocities required to accomodate standing wave-induced flux

through water column cross-sections spaced at 25 m intervals along the

shore-normal transit shown in this figure. Onshore/offshore velocities

between the shoreline and 800 m offshore were calculated for 3 standing

waves, each with a 4.7 cm amplitude at sensor B, but with node positions

ranging from 200 to 800 m offshore, and the corresponding periods

determined from Equation (9). Results are plotted in Figure 30a for

conditions of rising water at the shoreline antinode. Except for a change in

sign, the results are identical for falling water conditions.


1.0

ANTINODEANTINODE TAN N NODE

0 .5

800I6 0 04 0 02 0 0

P ressu reS e n so rs

5.76g tancp- ° - 5 r ( b ) x^ V ( a ) X=( 4 n f )

/ /S . TAN0

- 1 . 5

- 2.0

- 2 . 5

Figure 29. Generalized shoreface geometry, showing slopes used in determining standing wave node position, and changes in the cross- section affecting cross-shore velocities.


95

(A)

200m NODE

400m NODE800m NODE

i"'*"‘*......

>HOo- IUl>

ONSHORE FLOWOFFSHORE FLOW

- 2

-4

l i t * '

.1*'*“ — SEAS (T: 3.1s)

0 "

— SWELUT: 6.8s)

_i_____ .— .— i100 200 300 400

X (m)DISTANCE FROM SHORE

500 600 700 800

Figure 30a-b. Cross-shore trends in low-frequency flow under standing waves (30a), and bed shear stress due to waves attenuated at a constant H/d ratio of 0.2 (30b).


96

The model predicts that a standing wave with the amplitude observed

during run 3, and a node located 400 m offshore, would generate a mean

depth-averaged onshore velocity during the rising water phase of 3.8 cm-s-1

at sensor B. This is twice the actual depth- and time-averaged low frequency

onshore velocity monitored at this point between 600 and 900 s in run 3 (Fig.

24), but, given the many simplifications, the result is encouraging.

It can be seen in Figure 30a that, although the higher frequency,

shorter wavelength standing waves generate greater shore-normal velocities

for the same amplitude, cross-shore current speeds associated with a 750 s

oscillation (node 800 m from shore) would still be significant. The shape of

these curves is more important than the velocities calculated in this example.

For all three standing waves, the onshore velocity peak is shifted significantly

shoreward of the node position it would occupy if the bottom did not slope.

As a result, deceleration of onshore, and acceleration of offshore flows, due

to standing waves of any period, will be most pronounced within 100 m of

the shoreface. The existence of waxing and waning flows at a position fixed

by geometry is of great interest with respect to nearshore sediment transport

and deposition. It does not completely explain preferential deposition at the

shoreface, however, as both onshore and offshore velocities under a standing

wave are of the same magnitude.

Waves observed during the frontal and post-frontal periods were

attenuated completely before reaching the shoreface. Wells (1978) found

that attenuation across a Surinam mud bank was a linear function of depth.

The characteristic ratio of H/d identified in Surinam (0.2) also appears to act

as a limit to wave heights monitored in this study (Fig. 20). If this ratio is

assumed to be constant across the nearshore profile shown in Figure 29, peak


97

bed shear stresses generated by waves alone may be estimated at any point

using an equation derived by Krone (1976) from LWT:

pjc flHrms/T

T m ax = 1-7 p (— -------------------------------- (11)T \x sinh (2jcd/L)

where ji is the coefficient of viscosity. The empirical coefficient of 1.7 is

included to account for laboratory results which indicate that, without it, this

equation underestimates shear stress based on significant wave height (H j/3 =

1.4 Hrms) by 20 percent (Krone 1976). It should be noted that, for shallow

water waves, both the period, T, and the depth, d, enter Equation (11)

through the wavelenth, L = T (gd)l /2} as well as where explicitly stated.

Results of shear stress determinations at 25 m intervals are plotted in Figure

30b for the characteristic sea and swell wave periods, 3.1 and 6.8 s,

respectively.

Two points arise from this analysis. First, it is clear that shorter

wavelength seas generate greater peak bed shear stresses than do swells of

equal amplitude, attenuated at the same rate. Second, shear stresses due to

waves of both periods increase with distance from the shoreline.

The presence of a shore-normal wave shear stress gradient which is the

reverse of that found on surf-dominated coasts has clear implications for

sediment deposition, particularly when considered together with the

possibility for cross-shore transport of fluid mud. If a fluid mud layer

formed and maintained at one level of WBL shear is carried shoreward, the

potential exists for deposition when its yield strength exceeds the lower


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98

nearshore WBL shear. Mud transported offshore remains supported and is

less likely to deposit.The effect of wave period is less apparent. Though hardly conclusive

on this point, the results suggest that high-frequency seas may inhibit

development of a fluid mud layer. Wells (1978) has indicated that the

converse also appears to be true; that is, that fluid mud filters high-frequency

waves from the spectrum. Data from the chenier plain do not permit

resolution of this question. The thickness of the W BL is, however, known to

be a function of frequency (Eq. 7). It is probable that other factors related to

the state of development of the WBL may also influence the ability of a flow

to support a coherent fluid mud layer, particularly, if, as in most real

systems, several wave trains are superposed. The theoretical difficulties in

characterizing a simple system, with only two appropriate boundary layer

length scales, are significant, but they are magnified dramatically if multiple

WBL length scales must be considered. A range of frequencies then become

important, and, particularly for locally generated seas, a variety of

directions.Mud deposition through boundary layer dumping is triggered in the

modified B-L-S model whenever the yield strength of the fluid mud

suspension attains equivalence with wave generated shear stress. An abrupt

reduction in near-bottom suspended sediment concentration would occur

immediately after each dumping event. This reduction would permit

reformation of the WBL above the new deposit. The cycle of sediment

sorting, concentrating, and dumping might then be repeated, and give rise to

the stacked sequences noted in mudflat cores. It is important to note that

accretion can occur, however, only if sediment continues to be added


/

99

primarily from above the WBL, rather than from resuspension of sediment

already deposited.

High rates of mudflat sedimentation through WBL dumping are

possible if suspended sediment is continuously advected into the vicinity of

the shoreface. No high-resolution measurements of the rate of mudflat

sedimentation have been made. The best available observation is that of

Morgan et al. (1958) who noted that a layer of mud 18 inches thick (46 cm)

was deposited over a span of 3 days. Based on this, a lower bound may be

placed on the sedimentation rate if it is assumed.that the deposit has a density

of 1.3 g-cm-3, and, less realistically, that sedimentation was continuous

throughout the 3 day period. A deposition rate is calculated on the order of 1

g-m -2-s-l, 2000 times greater than the average rate for continental shelf

sedimentation estimated by McCave (1984).

Instantaneous rates several orders of magnitude higher could apply

during the boundary layer dumping events envisioned, such that a single bed

10 cm thick might be deposited in a matter of seconds. Sedimentation rates

ranging between 2 and 50 kg-m-2-s-l are calculated if deposition is assumed

to occur over a period of 1 to 30 s, and the density of the deposit formed is

again estimated at 1.3 g-cm-3. Deposition at this rate from a i m thick water

column implies a depth-averaged sediment concentration of nearly 50 g -H .

If the nearshore introduction of sediments at these concentrations is required

for mudflat sedimentation, it is clear why mudflat formation is such an

intermittent process, and, more specifically, why deposition did not occur

during the cold front monitored. The highest depth-averaged concentrations

found during this experiment, estimated by logarithmically extrapolating the

mean concentration data from the frontal period to the bed, were less than


/

100

2.5 g -l" l. One interpretation of this scaling is that, during the periods when

mudflat sedimentation is active, the flux of mud advected to the coast must be

an order of magnitude higher than was observed during the cold front

experiment.

The stated purpose of this investigation is to examine the processes by

which a coast incorporates influxes of mud intercepted from a shelf dispersal

system. A purely kinematic model has been proposed to explain preferential

deposition of mud at the shoreface. This is an important first step. Results of

the long-term shoreline monitoring program, historical data, and the

sedimentation rate scaling just concluded, indicate, however, that factors

affecting sediment supply operate at far greater temporal and spatial scales

than have been considered thusfar. Attention is directed now to these

regional influences.

SHELF SEDIMENT DISPERSAL

The high abundances of illite and illite-smectite mixed-layer clays in

mudflat sediment suggest an origin in modem Atchafalaya flood discharge,

rather than reworked older shelf deposits (Johns and Grim 1958, Brooks

1970), or low-flow Atchafalaya input (Van Heerden 1983), both of which

tend to be richer in smectite. A close temporal connection between mudflat

sedimentation on the chenier plain and the flux of Atchafalaya sediments to

the shelf seems likely. Littoral mud deposition in the study area occurred

only during late winter and in the spring when the River was in flood. The

high winds, waves, and water level fluctuations accompanying front passages

are also important, but frontal weather systems traverse this region from

September through April, a far longer period, which includes, but is not

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101

limited to, the 5 month window identified for coastal deposition. The fact

that mud sedimentation did not occur during the December experiment, and

does not occur generally, prior to peak flow on the Atchafalaya in late winter

and early spring, suggests the need for a volume or availability of advected

sediments nearshore which is attained only at that time.

The dynamical environment on the shelf is quite different from that

identified nearshore along the muddy coast. Turbulent shelf currents have

the potential to transport high concentrations of fine-grained sediments

dispersed through the full depth of the flow, rather than confined close to the

bed. Sediment concentrations at the surface within the shelf "mud stream"

offshore of Atchafalaya Bay range up to 800 m g -H during river flood

(Wells and Kemp 1981), and are higher than any measured near-surface on

the muddy coast (Table 4). This sediment appears to become concentrated in

near-bottom fluid mud layers where the characteristic turbulent eddy length

scale within the shelf flow is restricted, perhaps by a decrease in depth, and

WBL shear stress prevents deposition. Unsteady onshore-offshore flows,

associated with the wind-forced, shore-amplified water level oscillations

which precede front passage, then capture fluid mud for transport across the

shear stress gradient developed under progressively attenuated waves.

Stacked graded and laminated mud beds form accretionary deposits at the

shoreface through repeated dumping of fluid mud suspensions by the wave

boundary layer. Coastal progradation occurs when these mudflat deposits

are stabilized by marsh vegetation.

While the kinematics remain the same, a direct temporal link with the

period of river flood is not required in the special case of mudflat deposition

during hurricanes. Under normal cold front conditions, deep-water waves


102

offshore are too small to exert significant shear stress at the bed (Adams et al.

1982). Hurricane-generated waves, however, would be capable of

reworking and suspending large volumes of mud from deep-water shelf

storage sites. Fluid mud might form farther offshore, and be transported a

greater distance cross-shelf, by the "standing wave" flow associated with a

storm surge several meters in magnitude.

FUTURE CHENIER PLAIN PROGRADATION

The spread of clastic beaches in the area east of Freshwater Bayou

Canal provides an indication that littoral mud deposition is modulated by

changes in sediment flux over periods far longer than a single year. Mudflat

deposition and shoreline progradation west of the Canal continues at

approximately the same rate today as during the 1950’s and 1960's (Morgan

and Larimore 1957, Coleman 1966, Morgan and Morgan 1983). The lateral

extent of the muddy coast has decreased, however, as the eastern margin of

the zone of active mudflat formation has shifted 10 km west without any

corresponding extension on its western boundary. It is clear that less mud is

being deposited in the study area today than 20 or 30 years ago.

The subaerial delta in Atchafalaya Bay is now 13 years old, and covers

nearly 30 km^ (Wells et al. 1983). Sediments entering the bay in river flow

are dispersed across the delta through an intricate system of bifurcating

channels which did not exist 20 years ago. Prior to the emergence of sandy

islands, scour by waves maintained a constant depth and limited

sedimentation in the shallow, open bay (Thompson 1955). Now, fine

grained sediments which formerly reached the shelf, particularly during low

flow periods, are deposited in channel fills and on extensive algal flats


103

formed between natural levee ridges (Van Heerden et al. 1981, Van Heerden

1983). The amount of mud trapped relative to total fine-grained sediment

flux to the shelf is unknown, but evidence from the chenier plain suggests that

the presence of the new delta may be exerting an influence on downdrift

deposition.

Wells et al. (1983) have projected that the subaerial delta will continue

to grow within the confines of Atchafalaya Bay through the end of this

century, but will emerge on the inner shelf by 2030. Mudflat deposition on

the eastern chenier plain may, therefore, be expected to continue at

approximately the rate observed in this study for the next 50 years. When

delta development begins on the open coast outside the shell reef, it will be

exposed to greater erosion by waves, and can be expected to expand at a

slower rate. More fine-grained sediments will be delivered to the coastal

mud stream, particularly under low flow conditions, and mudflat

sedimentation downdrift should increase dramatically.

IMPLICATIONS FOR OTHER COASTS AND THE GEOLOGIC RECORD

Observations made during this study indicate that the only strict

requirement for mud sedimentation at the shoreface is that a large flux of

fine-grained sediments must be delivered close to shore. A high wave energy

regime, for example, is no barrier, as waves, and die boundary layers they

generate, play an important role in the formation of fluid mud layers

nearshore. A large tidal excursion, likewise, poses no restriction on fine

grained sedimentation at the coast. Rapid changes in water level are critical

to the onshore transport of sediment contained in fluid mud layers, therefore

semi-diurnal, rather than diurnal tides should be more effective for the same

range. A 3.2 m semi-diurnal tide almost certainly provides the same forcing


104

for cross-shore flow along the Surinam coast (Wells and Coleman 1981b,

Rine and Ginsburg 1985) as the wind-generated water level oscillations

observed on Louisiana's chenier plain. Where a turbid coastal current is not

developed, large storms can cause intermittent mud deposition at the beach.

Storm waves suspend large volumes of fine-grained sediments stored

offshore, and storm surges provide the cross-shore flows necessary for

transport to the shoreface. Such deposition has been observed outside of the

period of river flood along the Louisiana coast (Morgan et al. 1958) and on

the Rio Grande do Sul coast of southern Brazil (Delaney 1963, Martins et al.

1979).

Large-scale coastal progradation associated with mud deposition at the

shoreface requires a continuous, rather than intermittent, supply o f large

volumes of mud, as sediment-starved littoral deposits are quickly destroyed

by normal wave action. Continuity of fine-grained sediment provenance is

dependent on the steadiness of sediment introduction to the shelf, and on the

spatial stability of shelf dispersal. Seasonal and longer-term changes in these

factors characterize all real systems to a greater or lesser degree. Thus, all

muddy coasts can be expected to exhibit a mix,proportionate to this

variability, of coarse-grained and muddy facies at the receiving shoreface.

Onshore transport of fluid mud by unsteady flow under standing

waves is developed only on the foreshore, where the significant change in

bottom slope plays an important role. This restricts mud deposition to

shallow water in the immediate vicinity of the shoreface, and thereby limits

the maximum thickness of the deposit.

Intermittent littoral mud deposits associated with storms will have

poor preservation potential, but mudflats that form part of a major

B-«rr*iWW

Reproduced with permission of the copyright owner. Further reproduction prohibited without permission/

105

progradational sequence may survive if regional subsidence or uplift

produce gradual changes in the relative position of sea level. In cross-

section, all shoreface mud deposits should exhibit the characteristic laminated

and graded bedding which Stow and Bowen (1980) ascribe to boundary layer

sorting. Lower beds will be little disturbed by marine organisms, because of

the high rate of deposition associated with boundary layer dumping. Layers

close to the upper, subaerially exposed surface will, however, display cracks,

burrows, trails, and evidence of rooting.

W alker and Harms (1971,1976) have described in detail what they

interpret as coastal mudflat deposits from an ancient chenier plain setting.

This facies is found in more than 20 predominantly muddy, marine-

nonmarine sequences in the Upper Devonian Catskill Formation of central

Pennsylvania. The sequences represent episodes of coastal progradation on a

subsiding coast, which, in some cases, can be traced for 30 km down basin.

Transgressive sands define the boundaries of each cycle.

Indeed, the shales which lie at the top of the regressive sequences

display precisely the structures which are predicted for an accreting littoral

mudflat deposit. Rootlets and dessication cracks in some cases occur less than

2 m above marine faunas.


/

CONCLUSIONS

Fieldwork conducted between December, 1981, and May, 1983, along

a 30 km section of Louisiana's chenier plain coast disclosed widespread

shoreline retreat, averaging 5 to 10 m-yr-1. Intermittent sedimentation of

shore-welded mudflats was observed, however, to cause localized reversal of

this erosional trend on a 15 km long muddy coast extending west from

Freshwater Bayou Canal. This deposition takes place only in late winter and

early spring, synchronous with peak flood on the Atchafalaya River, which

discharges into Atchafalaya Bay 100 km to the east o f the study area.

A four day investigation of nearshore dynamics, conducted over the

broad range of wave and current conditions which accompanied passage of a

winter cold front showed that waves incident on the muddy shoreface were

well described by shallow water linear wave theory, but exhibited the

progressive nearshore attenuation and absence of breaking described by

Coleman and Wells (1981a) from Surinam. This attenuation gives rise to an

onshore decrease in wave shear stress at the bed which is the reverse of that

noted on clastic coasts dominated by surf zone processes.

A significant vertical gradient in suspended sediment concentration

was observed under low-frequency swell which was not developed when

higher-frequency seas formed an important component of the wave energy

spectrum. Concentrations 10 cm above the bed ranged up to the lower limit

for fluid mud (Einstein and Krone 1962) when this gradient was present.

The existence of a concentrated mud suspension near bottom at this time was

also apparent in elevated estimates of bed roughness identified in logarithmic

velocity profiles of the steady alongshore flow.

106


107

The formation of fluid mud layers appears to occur when sediments

suspended through the full depth of the coastal boundary current are

introduced into a less turbulent environment nearshore, but are prevented

from depositing by high shear stresses close to the bed within a wave

boundary layer (WBL). An important property of fine-grained sediment

suspensions at fluid mud concentrations, which was demonstrated for the

mudflat sediments in this study, is the development o f yield strength, which is

the ability to resist deformation up to some critical level of shear stress

(Einstein 1941). Concentration of the sediment load near-bottom in a gel

like fluid mud layer decouples the steady alongshore flow from turbulence

generated within the 1 to 2 cm thick WBL. Adjustment of the steady current

velocity profile to the presence of a fluid mud layer restricts further

transport of this sediment alongshore. Cross-shore transport of sediment

contained in this layer is proposed, however, under the influence of unsteady

onshore-offshore flows which were observed to be associated with low-

frequency, shore-amplified water level oscillations. These oscillations,

apparently forced by pre-frontal winds, have characteristics of a standing

wave with a shoreline antinode. Because of the slope of the foreshore,

onshore and offshore flows associated with standing wave-like motion

experience maximum acceleration and deceleration within 100 m of the

shore.

Intact transport of the fluid mud layer in the direction of an

accelerating flow is proposed as a consequence of the wave-induced

oscillatory shear stress exerted at its lower margin. Boundary-layer sorting

of cohesive and non-cohesive sediments at this point, of the type described by

Stow and Bowen (1980), implies the existence of a break in the vertical


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108

sediment concentration gradient at the top of the WBL. Because of the

relationship between sediment concentration and yield strength in fine

grained sediment suspensions in the fluid mud range, yield strength should

also exhibit a significant reduction at this point. Expanded to two

dimensions, this point of weakness forms a plane along which the fluid mud

layer can move in response to shear exerted by an accelerating flow at its

upper margin.

Rapid deposition of the fluid mud layer is proposed through a

boundary layer dumping process when the yield strength of the fluid mud

layer supported above the WBL reaches equivalence with WBL shear stress.

The onshore decreasing shear stress gradient developed under attenuating

waves gives rise to preferential deposition close to shore. Repeated

sequences of boundary layer sorting, followed by "freezing" of the fluid mud

layer, give rise to the vertically stacked 2 to 10 cm thick silt/shell + mud beds

which characterize mudflat cores. Sedimentation rates associated with this

process are estimated to be as high as 50 kg-m-2-s-l.

Shoreface accretion by boundary layer dumping gives rise to a

characteristic internal mudflat structure which exhibits little disturbance by

burrowing marine organisms. Littoral mudflats deposited as part of a major

coastal progradational episode have good preservation potential on a

subsiding coast, and are well displayed in the regressive sequences described

by Walker and Harms (1971,1976) from the Middle Devonian.

The spread of clastic beaches into an area of former mudflat deposition

in the eastern half of the chenier plain study area indicates that the flux of

mud to this coast has decreased over the past 20 years. Trapping of fine

grained sediment in the subaerial delta which has formed in Atchafalaya Bay


109

since 1973 is proposed as one possible cause for this reduction. Because of

the close temporal relationship between mudflat formation and fine-grained

sediment flux from Atchafalaya Bay to the shelf, no increase in littoral mud

sedimentation on the chenier plain can be expected until 2030, when the

subaerial delta is projected to expand onto the inner shelf (Wells et al. 1983).

Regional coastal retreat along the chenier plain is, therefore, likely to

continue for the next 50 years. Renewed large-scale advance of the marsh

shoreline across mudflats into the Gulf of Mexico, the progradational process

which has characterized this coast generally over the past 3000 years, is

predicted when the new delta emerges from Atchafalaya Bay, and is exposed

to increased wave action on the shelf.

Some variation of the dynamics proposed for fluid mud formation and

transport in the littoral setting investigated are likely to be manifested in the

many shelf and estuarine environments where highly concentrated near

bottom suspensions of fine-grained sediment are observed. Evidence

acquired nearshore, in this study, of the boundary layer sorting originally

proposed by Stow and Bowen (1980) to explain shelf margin turbidite

sedimentation suggests that this mechanism may be important wherever

multiple boundary layer length scales are involved.

The dynamical inferences drawn by Stow and Bowen (1980) from

sedimentological evidence, coupled with analysis of time-series

measurements made on muddy coasts, begun by Wells and Coleman (1981a,

b), and continued in this study, have led to proposal o f a new mode for fine

grained sediment transport in the marine environment. At the heart of this

progress is a recognition that properties of fine-grained sediment

suspensions, yield strength, for example, which do not emerge from


/

110

consideration of the behavior of individual particles, play an important role

in mud sedimentation.

The importance of continued field research is clear, as some processes,

like wave attenuation, which are critical to fluid mud formation, transport,

and deposition, are poorly developed at laboratory scales. The results of this

work suggest that the world's many muddy shorelines, which remain largely

untouched by the most basic research efforts, are appropriate "laboratories"

for the investigation of a broad spectrum of fine-grained sediment problems

which also arise in other environments less accessible for study.


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A PPE N D IX A

BEACH PROFILES

The eight Chenier Plain beach stations shown in Figure 2 were

reoccupied 9 times after establishment in the winter o f 1980/1981. The last

surveys were completed in May, 1983, thus the morphodynamic record

extends over 2.5 years. The results of these surveys are compiled in Figures

A -l through A-8. All profiles are tied to mean sea level (MSL) which lies

0.24 m above Gulf Coast Low Water, the standard datum referenced in local

charts and tide tables. Vertical exaggeration is 50 times.

133


F IG U R E A -l17 OEC'eO

LA. CHENIER PLAIN

PROFILE STA . 1

1 7 D E C '8 0 -2 5 M AY'83

1.0m

M8L'

28 JUL'S 1

A

20 MAY 81

VmffHPJQIT. t l - - *8 JUL Q1• 14 DEC'81 V

6 OCT 61

DEC'81?

_ 6 MAR 82

\ 12 JUL'82

100m250m200m

rTTT»U*l)‘u|l|iMMmiiff

18 DEC 82

20 OCT’82

SrwyMiyuiw,^.^.... w1 --------- ~

16 DEC 82

26 MAY'83 — s

........... nil.............


FIGURE A-2

29 MAY'S 1LA. CHENIER PLAIN

PRO FILE ST A . 2

1 7 D E C '8 0 - 2 5 MAY’8 320 JUL‘81

I.IIIM^i^lMAWAWIAHIAWAIIIAIBiW

, 8 OCT'81■ VMlVMilUWUlWr14 DEC'S 1

1 .0 m

100mMSI* 200m 250m150m

14 DEC'81

50mDEC'82 RW.lMl.i1 ■ iw ra w im m ^ —

12 JUL'62

8 DEC'82

^ I H I I I i m n i i i i H H i m i n , . , .


136

1.0m

MtL

FIGURE A-3

.A . CH EN IER PLAIN

PR O FIL E S T A . 3

2110m

|IH IIH H H lH H H in in ilH |,|


137

FIGURE A-4LA. CHENIER PLAIN

PROFILE 6T A . 420 MAY'S 1

1 7 D E C '6 0 -2 5 M AY'8317 DEC'80

20 JUL'S 1

10 ocrei

1.0m

SOmMSL 300m200m150m100m

13 JUL'82

10 DEC'8213 JUL'82 15 MAR'82

14 DEC'81

26 MAY'8310 DEC'82


FIGURE A-5UL CMENIER PLAIN

PROFILE ETA. 6

1 4 J A N '0 1 -2 6 M AY'83

1.0m

M8L

t t MAY 81

JUL 6 1

14 DEC'S 1

16 JUL'S 1

/ 10 OCT'S 1

_____

200mIS MAR 62

13 JUL’82

14 DEC 01

SOOm


139

1 •On*

USL.

F IG U R E A-6LA. CHENIER PLAIN

PROFILE STA. 6

1 4 JAN '81 “ 2 5 MAY '8 3

20 MAY'81

14 JAN '81

15 DEC'81

.......... ....

15 DEC'81

300m250m150m 200m50nf 100m

i OOCT'8

10 DEC '82

10 DEC’82

lC s a s B ==,


/

140

FIGURE A-7

LA. C H EN IER PLAIN

PROFILE ST A. 7

13 FEB’e i “26MAY’83

.'8115DEC81

13 OCT‘81

^ ....1.Qm

24 JUN '82

16 MAR *82

MSL 300r250m200m150m50nv 100m15 DEC *81

'" " " ••I I IM h i i IHIII

150EC *82

'{ I m m y ^ i y jM I M H I I H H l y 20 SEP '82H >H lll l lH H l ||n t „ )|

15 DEC'82

28 MAY'83

... .....


141

1.0m

MSL

FIGURE A-8

LA. CHENIER PLAIN

PROFILE ST A. 8

13 FE B ’8 1 - 2 6 MAY'83

20 MAY '81

20 JUL '61

j? " " 11.....

\ 15DEC81130CT*81

23JUN '82

8 APfl *82

— i s o m — '« « .7 > 2 0 0 m^ -iLt.3 0 0 m250m

*""1111111 """•••liUlllllMMIliiimmnTTii

100m

23 JUN'82

15DEC82

• ’ M ill,26 MAY *63

15 DEC '82

Utu


A PPE N D IX B

PROGRAM FOR GRANT-MADSEN BOUNDARY LAYER CLOSURE

The FORTRAN program listed on the next five pages was used to

perform the boundary layer closure referenced in the text. Documentation is

provided in the comments on the first 3 pages. References of the form "GM

EQ X" are to equations which appear in the journal article from which this

program was developed (Grant and Madsen 1979).

142


C * PROGRAM (GMMODEL) *C * THIS PROGRAM PERFORMS GRANT-MADSEN(1979) CLOSURE FOR *C * WAVE-CURRENT BOUNDARY LAYER PROBLEM. ANGULAR DEVIATION *C * OF MEAN CURRENT FROM DIRECTION OF WAVE PROPAGATION PHICB)*C * IS ASSUMED FIXED AND MUST BE POSITIVE AND LESS THAN 90 *C * DEGREES. A NIKURADSE BOTTOM ROUGHNESS(XKB) MUST ALSO BE *C * SPECIFIED. INITIALLY, GUESS A VALUE FOR UA/UB(RATIO) AND*C * WAVE-CURRENT FRICTION FACTOR(FCW). OBTAIN A NEW VALUE *C * FOR FCW THROUGH ITERATION IN NESTED LOOP. RETURN TO MAIN*C * LOOP TO CALCULATE AN APPARENT BOTTOM ROUGHNESS(XKBC1) *C * WHICH IS COMPARED TO VALUE CALCULATED FROM FIELD DATA *C * (XKBC2) FOR MEAN CURRENT VELOCITY(UCR) MEASURED AT *C * REFERENCE ELEVATION(ZR) ABOVE BED. IF A SPECIFIED *C * TOLERANCE(TOL2) IS EXCEEDED, MAIN LOOP IS RE-EXECUTED *C * WITH AN ADJUSTED VALUE FOR UA/UB(RATIO). THIS PROCESS IS*C * REPEATED UNTIL CONVERGENCE CRITERION IS SATISFIED. *C ************************************************************C INPUTSC FIELD VALUES FOR WAVE CHARACTERISTICS(LINEAR WAVE THEORY)C TITLE RUN NO.C FREQ WAVE FREQUENCY IN RADSC UB MAX NEAR-BOTTOM ORBITAL VELOCITY(CM/S)C AB NEAR-BOTTOM OSCILLATORY EXCURSION AMP(CM)CCCCCCC FIELD VALUES FOR CURRENT CHARACTERISTICSC PHICB DEVIATION OF MEAN STRESS VECTOR FROM DIRECTION OFC WAVE PROPAGATION IN RADS(LESS THAN 90 DEG)C PHIC DEVIATION OF MEAN CURRENT IN WBL FROM DIRECTION OFC WAVE PROPAGATION IN RADS. GM EQ 18.C UCR MEAN CURRENT VELOCITY AT REFERENCE ELEVATION IN CM/SC ZR REFERENCE ELEVATION-POSITIVE UP IN CM.C XKB NIKURADSE ROUGHNESS IN CMCCCCCC TOLERANCES FOR ITERATIONSC TOL1 PERMISSABLE ERROR FOR FCW ITERATIONC TOL2 PERMISSABLE ERROR FOR KBC ITERATIONCCCCC VARIABLES FOR WHICH STARTING ESTIMATES ARE NEEDED C RATIO RATIO OF UNKNOWN CURRENT REFERENCE VELOCITY TO MAXC NEAR-BOTTOM ORBITAL VELOCITY (UA/AB). USE (UCR/AB)


ccccccccccccccccccccccccccccccccc

TO START.FCW WAVE-CURRENT FRICTION FACTOR.

GM FIG.4.CHOOSE START VALUE FROM

CCCCCCCCCCCCC

VARIABLES FOR WHICH NO STARTING ESTIMATES ARE NEEDED VARIABLES DEPENDENT ON RATIO AND PHIC ONLY V2 GM EQ 16ALPHA GM EQ 20

TERMS DEPENDENT ON V2, ALPHA, AND PHICB ONLY Q1 LAST TERM LFT SIDE GM EQ 54Q2 RT SIDE GM EQ 54

TERMS CALCULATED IN FCW ITERATIONUCW CHARACTERISTIC TURBULENCE INTENSITY IN WBL(U*CW)

GM EQ 21XL PRANDTL MIXING LENGTH(L)

GM EQ 29ZETAO NO-SLIP WBL BOTTOM BOUNDARY PARAMETER

GM EQ 31X ARGUMENT OF KELVIN FUNCTION (2*SQRT(ZETAO))

TERMS CALCULATED IN IMSL SUBROUTINE MMKELOc BER BER X (ORDER 0)c BEI BEI X (ORDER 0)c XKER KER X (ORDER 0)c XKEI KEI X (ORDER 0)c IER ERROR PARAMETER

XK BRACED TERMS IN GM EQ 53GFCW NEW ESTIMATE OF FCW FROM TRANSPOSED GM EQ 54Q3 SECOND TERM LFT SIDE TRANSPOSED GM EQ 54Q4 DUMMY VARIABLE USED IN CALCULATING GFCWXDIF1 (GFCW - FCW)


oo

oo

o

on

o

oo

oo

oo

oo TERMS DEPENDENT ON RESULTS OF FCW ITERATION

UC CHARACTERISTIC SHEAR VELOCITY ABOVE WBL GM EQ 15BETA EXPONENT IN GM EQ A9XKBC1 ESTIMATE OF XKBC FROM TRANSPOSED GM EQ A9XKBC2 FIELD BASED ESTIMATE OF XKBC FROM GM EQ 49XDIF2 XKBC2-XKBC1

DIMENSION TITLE(2)READ(5,10) TITLE,FREQ,UB,AB READ(5,20) PHICB,PHIC,UCR,ZR.XKB READ(5,20) T0L1.T0L2 READ(5,20) RATIO,FCW

10 F0RMAT(2A4,2X,3F10.5) _ .... - - -'20 F0RMAT(7F10.5)

WRITE(6,100) TITLE WRITE(6,200) FREQ,UB,AB WRITE(6,300) PHICB,PHIC,UCR,ZR,XKB WRITE(6,400) T0L1.T0L2 WRITE(6,500) RATIO,FCW

100 F0RMAT('1',2X,2A4,10X,’INITIAL VALUES')200 FORMATC'O',2X,'WAVE PARAMETERS’/2X,'FREQ =',F10.5,2X,'UB =',

1 F10.5,2X,'AB =',F10.5)300 F0RMAT('0',2X,'CURRENT PARAMETERS’/2X, 'PHICB =',F10.5,15X,

1 'PHIC ='.F10.5/2X,'UCR =',F10.5,2X,'ZR =',F10.5,2X,'XKB =',1 F10.5)

400 F0RMAT('0',2X,'CONVERGENCE TOLERANCES'/2X, 'FCW ITERATION*,F11.5/ 1 2X,'XKBC ITERATION',F10.5)

500 FORMATC'O',2X,'STARTING ESTIMATES'/2X,'RATIO'.F10.5/2X,'FCW',1 F12.5)

XKAPPA =0.4 PI =3.14159

30 DO 1 J = l , 21 C EXIT LOOP IF NO CONVERGENCE AFTER 20 ITERATIONS

I F ( J .EQ. 2 1 ) GO TO 210V2 = ( ( 2 . * R A T I 0 ) / P I ) * S Q R T ( 4 . - ( 3 . * ( S I N ( P H I C ) ) * * 2 . ) )ALPHA = 1 .+ (R A T I0**2 .)+2 .*R A T IO *C O S (P H IC )Q3 = ( ( ( SQRT(SQRT(ALPHA)) ) * * 3 . ) / 4 . ) - ( ( V 2**2 . ) / ( 4 . *SQRT(ALPHA)) )Q2 =(V2/(2 .*SQR T(SQ RT(A LPH A )) ) )*COS(PHICB)

W R IT E (6 ,6 0 0 ) J ,V 2,A LPH A ,Q 2,Q 3 600 F O R M A T ( ' l ' , 2 X , ’ TERMS DEPENDENT ON RATIO, PHIC, AND PHICB ONLY'/

1 2X,'ITERATION = ’ , I 3 / 2 X , 'V 2 = ' ,F 1 0 .5 /2 X , 'A L P H A = ' , F 1 0 . 5 / 2 X , ' Q2 = ' 1 F 1 0 . 5 , 2 X , ' Q3 = ' , F 1 0 . 5 )


40 DO 2 K=1,21C EXIT LOOP IF NO CONVERGENCE AFTER 20 ITERATIONS.

IF(K .EQ. 21) GO TO 220UCW =SQRT(.5*FCW*ALPHA*UB**2.)

UC =SQRT(0.5*FCW*V2*(UB**2.))XL =(XKAPPA*UCW)/FREQ ZETAO =XKB/(30.*XL)X =2.*SQRT(ZETA0)WRITE(6,700) K,UCW,UC,XL,ZETAO,X CALL MMKELO CX,BER,BEI,XKER,XKEI,IER)WRITE(6,800) BER,BEI,XKER,XKEI,IER XKEI =ABS(XKEI)XK =1./(X*(SQRT((XKER**2.)+(XKEI**2. ))))Q1 =.097*SQRT(XKB/AB)*(XK/FCW**(3./4.))Q4 =XK* v Q.9J*SQRTCXKB./AB.)#( ( Ql+( 2 .*Q2 -) ) /Q3 )......-.......GFCW =(Q4**(4./3.))WRITEC6,900) FCW,GFCW,XK,Q1,Q2,Q3,Q4 XDIF1 =(GFCW-FCW)IF(ABS(XDIF1) .LE. T0L1) GO TO 50 FCW =GFCW

2 CONTINUE700 FORMATCl','TERMS USED IN MMKELO. ITER = ',I3/2X,'UCW =',F10.5,

1 2X,'UC =' ,F10.5/2X,1 XL =',F10.5,2X,'ZETAO =',F10.5,2X,'X 1 F10.5)

800 FORMATC'O',2X,'MMKELO OUTPUT1/2X,'BERX =',F10.5,2X,'BEIX =',1 F10.5/2X,'XKER =',F10.5,2X,'XKEI =',F10.5,2X,1IER =',F10.5)

900 FORMATC'O',2X,'FACTORS USED TO FIND GFCW'/2X,'FCW =',F10.5,2X,1 'GFCW =' ,F10.5,2X,'XK = ',F10.5/2X,'Q1 =',F10.5,2X,'Q2 1 F10.5,2X,'Q3 =',F10.5,2X,'Q4 =’,F10.5)

50 FCW =GFCWUC =SQRT(0.5*FCW*V2*CUB**2.))UCW =SQRTC.5*FCW*ALPHA*UB**2. )BETA =1.-(UC/UCW)WRITEC6,910) K,UC,UCW,BETA,FCW

910 FORMATC'1',2X,'CONVERGENCE REACHED ON FCW AFTER',13,1 ' ITERATIONS'/2X,'TERMS NEEDED TO FIND XKBC’/2X,'UC =',F10.5,1 'UCW =' ,F10.5,2X,'BETA =',F10.5,2X,'FCW =',F10.5)XKBC1 =XKB*CC24.*(UCW/UB)*(AB/XKB))**BETA)XKBC2 =30.*ZR/EXP(UCR*XKAPPA/UC)XDIF2 =XKBC2-XKBC1 WRITEC6,920) J,RATIO,XKBC1,XKBC2,XDIF2

920 FORMATC'O',2X,'RESULTS OF ITERATION',13,' ON KBC'/2X,'RATIO =', 1 F10.5,2X,'XKBC1 =',F10.5,2X,'XKBC2 =',F10.5,2X,'XDIF2 1 F10.5)IFCABSCXDIF2) .LE. T0L2) GO TO 60 IFCABSCXDIF2) .LE. 5.) GO TO 55 RATIO =RATI0*SQRTCXKBC1/XKBC2)GO TO 1

55 RATIO =RATI0*SQRT(SQRTCXKBC1/XKBC2))1 CONTINUE

210 WRITEC6,310)GO TO 999

220 WRITEC6,320)GO TO 999


310 FORMATC' 1 ' , ' NO CONVERGENCE ON RATIO AFTER 20 ITERATIONS. ABORT') 320 FORMATC’ 1 NO CONVERGENCE ON FCW AFTER 20 ITERATIONS. ABORT')

60 XKBC =XKBC2W R ITE(6 ,9 3 0 ) J ,T IT L E

930 FORMATC'l',2X,'CONVERGENCE REACHED ON RATIO A F T E R ',13 ,1 1 IT E R A T IO N S '/2 X .2 A 4 ,1 0 X , ' INITIAL VALUES')WRITEC6.200) FREQ, UB, AB WRITEC6,3 0 0 ) PHICB,PHIC,UCR,ZR,XKB WRITEC6 , 4 0 0 ) T 0L 1 ,T 0L 2 WRITEC6,9 4 0 )WRITE C 6 , 9 5 0 ) RATI0 , FCW, XKBC, UCW, UC, XL, ZETAO, V 2, ALPHA

940 FORMATC'O1,2 2 X , 'F IN A L RESULTS').................950 FORMATC'O', ' RATIO = ' ,F 1 0 .5 ,2 X , 'F C W = ' ,F 1 0 .5 ,2 X , 'X K B C = ’ , F 1 0 . 5 /

1 2 X , ' UCW = ' , F 1 0 . 5 , 2 X , 'U C = ’ , F 1 0 . 5 / 2 X , ' XL = ' , F 1 0 . 5 ,1 2X,'ZETAO = ' . F 1 0 . 5 / 2 X , 'V 2 = ' ,F 1 0 . 5 , 2 X . ' ^ f H A = ' , F I 0 . 5 ) ............. .......

: • w. r_ : g gj g 'CONTI NT)E r - - ; ...... -...r - - - - -.....STOPEND


V IT A

George Paul Kemp was bom on May 2,1954, in Port Jefferson, New

York. He graduated cum laude from Saint Paul's School in Garden City,

New York, in May, 1971. He enrolled in the Fisheries Science program at

Cornell University in Ithaca, New York, and received the B.S. degree in May

1975. Mr. Kemp began graduate studies later that year in the Department of

Marine Sciences at Louisiana State University. Following award of the M.S.

degree in May, 1978, he pursued his interest in estuarine geochemistry as a

Research Associate in the Department of Engineering Research, also at

L.S.U., through the fall o f 1979. He then re-entered the Department of

Marine Sciences to begin a graduate program in coastal sedimentology

through the Coastal Studies Institute. Mr. Kemp was awarded the Joseph

Lipsey, Sr., Scholarship in 1981, and spent 1984 in Washington, D.C., where

he served on the legislative staff of Senator Edward M. Kennedy as a Sea

Grant Congressional Fellow. Since returning to Louisiana, Mr. Kemp has

been employed by Groundwater Technology, Inc., in Mandeville. Mr. Kemp

is currently a candidate for the Ph.D. degree in the Department of Marine

Sciences.

148


DOCTORAL EXAMINATION AND DISSERTATION REPORT

Candidate: G. Paul Kemp

Major Field: Marine S c ien ces

Title of Dissertation: MUD DEPOSITION AT THE SHOREFAGE: WAVE AND SEDIMENTDYNAMICS ON THE CHENIER PLAIN OF LOUISIANA

Approved:

Major Professor and .Chairman

Dean of the Graduatee School

EXAMINING COMMITTEE:

Date of Examination:

November 26, 1986

n"7—7*


Documents

Mud Deposition at the Shoreface: Wave and Sediment