12. DEEP-MARINE SEDIMENTS AND SEDIMENTARY · PDF fileBouma, the so-called Bouma turbidite...

INTRODUCTION

The deep marine is a unique sedi-

mentary environment compared to all

others because of its inaccessibility

and the enormous spatial scale of

many of its constituent depositional

systems. For, example, the modern

Bengal Fan, which has been accumu-

lating sediment for only about 55 m.y.,

is 3000 km long, over 1400 km wide,

>5000 m thick, and contains an esti-

mated sediment volume of 4x106 km3

(Table 1). (For a global inventory of

fan systems, see the wall chart of

Barnes and Normark, 1985). Accord-

ingly, the principal investigative tool

for modern deep-marine environ-

ments is seismic, including an array

of techniques that range from high-

frequency, high-resolution but shal-

low-penetrating surveys, to lower fre-

quency, more deeply penetrating but

lower resolution surveys. The 3-D

seismic with its ability to image fea-

tures in both plan and cross-sectional

views has proven to be especially

useful. Over the past decade, this

work has resulted in the publication of

many stunning seismic images that

have improved greatly our under-

standing of the deep-marine sedi-

mentary system (e.g., Weimer et al.,1991; Posamentier and Kolla, 2003;

Posamentier and Walker, 2006).

Nevertheless, the seismic method

generally suffers from limited vertical

resolution – minimum vertical resolu-

tion of industry seismic data is com-

monly of the order of 10 meters. To

fill the gap, the geological community

must also utilize outcrop studies, but

here too a number of shortcomings

are recognized (Fig. 1). Perhaps

most importantly, the outcrop record

is inherently 2-dimensional with, at

best, local 3-D perspectives (i.e., val-

ley cuts). Also, in most cases, the hor-

izontal scale of most outcrops is small

compared to the spatial scale of most

deep-marine architectural elements,

and, more profoundly, their parent

sedimentary system. Nevertheless,

over the past decade or so, much

research, including work on a number

of seismic-scale outcrops, is helping

to bridge the gap between the ancient

outcrop record and modern seismic

images, to merge these two inde-

pendent datasets into a single, coher-

ent package (see, for example, Nilsen

et al., 2007). Nevertheless, the

emphasis here, as in other chapters

in this volume, is on the lithological

characteristics of the strata that make

up the deep-marine sedimentary

record, and also on how these strata

are distributed spatially (space) and

vertically (time) in the geological

record. Excellent reviews of the seis-

mic-scale attributes can be found

elsewhere (e.g., Posamentier and

Kolla, 2003; Posamentier and Walker,

2006).

BRIEF HISTORY OF DEEP-MARINE

SEDIMENTOLOGY

Turbidites are a ubiquitous feature in

the deep-marine sedimentary record,

and are deposited from subaqueous

turbidity currents. However, because

the formative process is largely hid-

den from direct observation, the con-

nection between process and deposit

remained speculative and anecdotal

since as early as the late 1800s.

Knowledge of deep-marine process-

es took a major step forward following

the laying of the first successful

Atlantic seafloor telegraph cable in

1866. In 1899, Benest (in Heezen and

Ewing, 1952) reported that “accidents

to cables have already been valuable

... in directing attention to hitherto

unsuspected forces constantly in

action altering the features of the sea

bottom”. The most famous cable dis-

ruption occurred on November 18,

1929, on the southern margin of the

Grand Banks off east coast Canada.

On that day, an estimated 7.2 magni-

tude earthquake occurred and dis-

rupted telegraph communication

between North America and Europe.

Several submarine cables were

apparently broken at the time of the

earthquake, presumably as a result of

seabed movement. More puzzling,

however, was the fact that a number

of other cables became deactivated

successively after the earthquake,

some as much as 13 hours later

(Heezen and Ewing, 1952; Piper etal., 1999). Furthermore, the cables

broke progressively in a single (off-

shore) direction. It was hypothesized

that a turbidity current moving at con-

siderable speed was responsible

(Heezen and Ewing, 1952). At about

the same time, Kuenen and Migliorini

(1950) began experimenting with

sediment dispersions released into a

basin of still-standing water. Upon

release the dispersion formed a bot-

tom-hugging turbidity current that pro-

duced a deposit with a characteristic

upward decrease in grain size. Based

on these results, these authors sug-

gested that the common occurrence

of upward-fining beds in the deep-

marine geological record might reflect

deposition from turbidity currents.

Beds showing this upward-fining

character were observed in flysch

deposits of the European Alps by

Bouma (1962). In addition to the

upward fining, he also noted a char-

acteristic vertical succession of sedi-

mentary structures, the origin of

which was then unknown. Shortly

thereafter, experimentalists illustrated

the variety of bedforms that formed

under unidirectional currents of vari-

ous speeds (Simons et al. 1965; and

now many others). Based on those

results the characteristic suite of sed-

imentary structures observed by

Bouma, the so-called Bouma turbidite

sequence, reflects deposition from a

decelerating turbidity current. At

about the same time, Middleton

(1966a, b, 1967) published a series of

important papers based on flume

experiments, describing the various

1

DRAFT FORMAT

12. DEEP-MARINE SEDIMENTS AND

SEDIMENTARY SYSTEMS

R. William C. Arnott, Department of Earth Sciences, University ofOttawa, Ottawa, ON K1N 6N5

parts of a turbidity current and the

influence of sediment concentration

on depositional patterns and charac-

teristics.

Today, it has been well established

that sediment-gravity flows, principal-

ly turbidity currents and debris flows,

in addition to mass-movement

processes, are the major formative

agents of the deep-marine sedimen-

tary record. A superb illustration is the

1929 Grand Banks event, which is

now known to have been initiated by

widespread failure of a surficial layer

about 20–25 m thick (Piper et al.,1999). The slide, which comprised

numerous overlapping slumps formed

by progressive retrogressive failure,

moved downslope and in areas of

high slope transformed into a debris

flow and ultimately into a sustained,

fast-moving (~70 km/hr) turbidity cur-

rent. This current eroded on the conti-

nental slope and eventually deposited

a turbidite up to 1 m thick on the Lau-

rentian submarine fan. The turbidite

contains about 185 km3 of sediment –

using railway boxcars to conceptual-

ize volume, 185 km3 equates to 1.29

billion boxcars, or a train almost 22

million kilometers long, and one that

would wrap around the world over

558 times!

SEDIMENT TRANSPORT

MECHANISMS AND DEPOSITS

Unlike continental and shallow-

marine sedimentary systems, sand

and gravel transport in the deep

marine is dominated by sediment-

gravity flows and mass-movement

processes. Finer grained sediments,

principally silt and clay, although pres-

ent in sediment-gravity flows and their

related deposits, are transported

mostly in suspension.

Mass-Movement Deposits – Slides

and Slumps

The gravity-driven downslope move-

ment of coherent to semi-coherent

masses of sediment along discrete

2 ARNOTT

DRAFT FORMAT

Figure 1. Vertical and horizontal scales of modern deep-marine systems (Ben-

gal, Monterey and Navy fans), ancient turbidite systems (Butano), some deep-

marine sediment features/elements, and outcrop and core (adapted from

Barnes and Normark, 1985). Note how the scale of most outcrops is dwarfed by

the scale of seismic data sources, but more importantly, the size of deep marine

fans and their constituent depositional elements.

Table 1. Dimensional characteristics of five deep-marine turbidite systems. Data from wall chart of Barnes and Normark

(1984).

Length, Maximum

Name Location Age Width Thickness Volume Dominant Range

(km) (m) (km3) Grain Size

Amazon Brazilian Middle 700, 4200 7x105 mud pebbles

margin Miocene 700 (max) to mud

Bengal Bay of Eocene 2800 (min), >5000 4x106 mud mud to

Bengal 1400 (max) medium

sand

Laurentian coast Quaternary 1500 (max), 2000 1x105 mud to med- mud to

eastern 400 fine sand gravel

Canada

Mississippi Gulf of Pleistocene 540, 570 4000 2.9x105 silty mud mud to

Mexico gravel

Navy coast Late 900 75 sandy silt mud to

southern Pleistocene gravel

California

failure planes is termed mass move-

ment or mass wasting. Such move-

ments, which can range up to hun-

dreds of kilometers in distance, occur

when the driving force, gravity,

exceeds the tensile strength of the

parent sediment pile (which depends

upon a wide range of internal condi-

tions, including pore-fluid pressure,

sediment composition and consolida-

tion and the occurrence of mechani-

cally weak layers among others). The

mechanisms responsible for trigger-

ing the initial instability include over-

steepening, seismic loading, cyclic

storm-wave loading, rapid accumula-

tion and underconsolidation, gas

charging, gas-hydrate dissociation,

seepage, etc. (e.g., Locat and Lee,

2002). Once initiated, movement

continues until the resisting forces,

principally friction along the basal

failure plane, exceed the gravitation-

al driving force and en masse depo-

sition takes place. Following earlier

workers, two end-member kinds of

mass-movement deposits are recog-

nized: slides and slumps, where the

principal difference between the two

is the intensity and nature of internal

deformation. In the case of slides,

deformation is comparatively minor

and is dominated by brittle deforma-

tion, mostly in the form of bedding-

parallel surfaces of detachment,

although shear deformation may be

intense in a thin zone near the base

of the unit (Fig. 2A). In slumps, defor-

mation shows an element of rotation

and typically is more intense and

ductile in character (Fig. 2B). Soft-

sediment ‘slump’ folds are abundant,

and are commonly tightly folded with

the axial plane sub-parallel to planes

of internal shear.

Sediment-Gravity Flows

Gravity currents occur when a more-

dense fluid moves through and dis-

places a less-dense fluid. Sediment-

gravity currents are a type of gravity

current, wherein the density contrast

is produced by the presence of sus-

pended sediment. Early classifica-

tion schemes for sediment-gravity

flows recognized four types of end-

members: debris flows, grain flows,

fluidized/liquefied flows, and turbidity

currents (Middleton and Hampton,

1976), each differentiated by the pri-

mary mechanism for suspending

sediment and maintaining the densi-

ty contrast: matrix strength, grain col-

lision, escaping pore fluid and fluid

turbulence, respectively. More

recently, an alternative scheme by

Mulder and Alexander (2001) pro-

posed a classification based on flow

properties and sediment-support

mechanisms, and recognized two

types of end-member flows: cohe-

sive flows and frictional flows. Here,

a modified version of this classifica-

tion will be adopted.

Cohesive Flows — Debris FlowsCohesive flows, which are more

commonly termed debris flows and

mud flows, are sediment-gravity

flows where the volume concentra-

tion of the solid and fluid phases are

of the same order of magnitude, and

the occurrence of a cohesive matrix

imparts a pseudoplastic rheology to

the flow (e.g., Mohrig et al., 1998;

Mulder and Alexander, 2001). Here-

after the term debris flow, which is

entrenched in the geological litera-

ture, will be used to refer to all cohe-

sive flows. Particles are principally

suspended by cohesive forces pro-

vided by a matrix of fluid and fine-

grained sediment (generally a

silt–clay mixture. Note that in both,

subaerial and subaqueous debris

flows (see also Chapter 6), the

amount of clay-size particles needed

to generate sufficient yield strength

to suspend larger particles can be

surprising low, possibly of the order

of 2–4% by volume. In addition to

matrix strength, buoyancy effects,

particle-particle interaction (collisions

and near misses), hindered settling,

elevated pore pressure, and in some

cases fluid turbulence may provide

additional support for suspended

particles. Deposition occurs when

one or more of the following exceed

the driving gravitational force: intrin-

sic shear strength of the sediment-

water mixture, grain-contact friction

and friction along the flow bound-

312. DEEP-MARINE SEDIMENTS AND SEDIMENTARY SYSTEMS

DRAFT FORMAT

Figure 2. Mass-movement deposits. A) Slide deposit (strata are vertically dipping due to later (Cordilleran) tectonic defor-

mation and stratigraphic top is toward the left). Base of slide is indicated by the solid black line. Note the extensively

deformed and brecciated strata in the lower part of the slide, including a large rotated block (Neoproterozoic Isaac Forma-

tion, British Columbia). B) Slump deposit with extensive internal ductile deformation (base of slide indicated by solid white

line; Carboniferous Gull Island Formation, Ireland).

aries, which then causes the flow to

freeze inward, either en masse, or

more gradually from areas of lower

shear near the flow’s surface toward

those of higher shear at the base.

Debris-flow deposits, or debrites,

form sheetlike to lobe-shaped mass-

es that have steep margins as a result

of the strength of the moving and

deposited sediment mass. Deposits

range widely in scale, but can be up

to tens of kilometers wide and over

100 m thick, although much thinner

beds are more common. At their

downflow terminus toe-thrusts are

common due to a rapid down-flow

reduction in flow speed. Bases of

debris-flow deposits are commonly

planar and non-erosional, because of

the strength of the moving mass and

the damping of large-scale fluid turbu-

lence. Nevertheless, deeply scoured

bases are observed. In some cases

they represent the opportunistic occu-

pation of a pre-existing seafloor chan-

nel (Fig. 3), whereas others are

thought to be formed by a rigid part of

the flow ‘ploughing’ through the

underlying seafloor sediment (see

Posamentier and Walker, 2006, their

Figures 155, 157) or being dragged

along the surface forming linear

grooves up to 40 m deep, several

hundred meters wide and extending

longitudinally for more than 20 km

(Posamentier and Kolla, 2003). The

upper surface can also be uneven

and ranges between flat to highly

rugose. Internally, debris-flow

deposits range from mud- to sand-

rich, typically with a disorganized,

poorly sorted character (e.g., Nemec

and Steel, 1984; Fig. 4A). On seismic

images debris-flow deposits, espe-

cially those interpreted to be mud rich,

exhibit a distinctive chaotic or reflec-

tion-free character. Where present,

clasts ranging from sand grains to

enormous blocks are generally dis-

persed throughout a fine-grained

matrix. In some deposits, clasts of

incorporated soft sediment are con-

torted and commonly show a subtle to

well-developed orientation with their

longest dimension oriented subparal-

lel to the base of the deposit (Fig. 4B).

Preferential particle alignment and

particle deformation is likely the result

of shearing within the sediment mass

during transport, which then may be

accentuated by post-depositional

compaction and consolidation.

Although cohesive, some debris

flows are capable of movement over

long distances. For example, Gee etal. (1999) reported a modern debris

flow off the west coast of Africa with a

run-out distance of about 700 km.

Such large distances may be attrib-

uted to hydroplaning and the atten-

dant reduction of friction between the

bed and the overlying flow. Elevated

hydrodynamic pressure exerted on

the forward part of the flow causes a

wedge of ambient fluid (i.e., seawa-

ter) to penetrate beneath the flow and

separate it from the bed (Mohrig et al.,1998). As a result, frictional resist-

ance at the base of the flow is signifi-

cantly reduced and the weight of the

overlying debris flow becomes borne

by the fluid. Importantly, the overrid-

ing debris flow must have sufficiently

low permeability (i.e., be sufficiently

muddy) so that the basal fluid layer

does not dissipate.

Debris flows are also capable of

spawning turbidity currents. Owing to

mixing along its leading edge, sedi-

ment is eroded from the debris flow

and cast into a developing turbulent

suspension above the flow. However,

owing to the low permeability of most

debris flows, water infiltration is mini-

mal and the amount of sediment erod-

ed and transferred into the turbidity

current is very small (<1%). Also, so

long as the debris flow and overlying

turbulent flow are moving at a similar

velocity, the amount of erosion along

the interface between them is negligi-

ble. But, if the debris flow stops, then

the turbulent suspension can rework

the top of the debris-flow deposit as it

becomes detached from the parent

debris flow and continues farther

downslope (Fig. 4C). Conversely, it

has been reported that, at an abrupt

reduction in slope, turbidity currents

can be partially transformed into a

debris flow. As the turbulent flow

decelerates rapidly and turbulence is

reduced, suspended particles fall rap-

idly toward the bed where the ubiqui-

tous occurrence of cohesive mud par-

ticles eventually increases yield

strength to the point where the parti-

cles become supported by matrix

strength.

Mass transport deposits (MTDs)

4 ARNOTT

DRAFT FORMAT

Figure 3. A) Ancient example of a partly debris-flow-filled channel. Base of channel is indicated by dashed black line. After

accumulating several meters of sandstone, the channel became partly plugged by a debris-flow deposit (DF; Neoprotero-

zoic Isaac Formation, British Columbia). Thereafter, channel-filling sandstones onlap and then overlap the debris-flow

deposit. B) Seismic image of the chaotic reflections of a debris-flow deposit that has exploited a pre-existing seafloor chan-

nel (photo courtesy of Henry Posamentier).

are observed often in shallow- and

deep-penetrating seismic images.

The MTDs are typically erosively

based deposits that occur on a range

of scales, including enormous fea-

tures that cover areas up to several

1000 km2 and are several 100s of

meters thick. The dimensions of

MTDs, including their length:width

ratio, are controlled by the geomor-

phology and position of the sediment

source, wherein large, expansive

deposits are related to instability initi-

ated along the upper part of the basin

margin, at or near the shelf edge,

whereas smaller deposits form in the

basin by failure along local slopes

(detached sediment sources;

Moscardelli and Wood, 2008). Inter-

nally MTDs consist predominantly of

a complex, typically disorganized

assemblage of slump, slide and

debris-flow deposits that can locally

be interstratified with (organized)

channel and overbank strata. The

occurrence of large MTDs in the

stratigraphic column represents a

major change in sedimentation

regime within the basin. In many

cases these changes have been

interpreted to be related to allocyclic

processes, and therefore to have

sequence-stratigraphic significance

(see sequence-stratigraphy section

below).

Frictional Flows

According to Mulder and Alexander

(2001), frictional flows form a contin-

uum from mass movements (slumps/

slides – see above) to a variety of dif-

ferent kinds of sediment suspensions

subdivided on the basis of the domi-

nant mechanisms of sediment sup-

port. The most common the friction-


DRAFT FORMAT

Figure 4. Debris-flow deposits (debrite). A) Schematic diagram showing typical characteristics of subaqueous debris-flow

deposits (from Nemec and Steel, 1984). B) Sharp-based debris-flow deposit overlying a succession of thin- and medium-

bedded turbidites (below the dashed line). Note the dispersed quartz pebbles and mudstone clasts, some of which have

been folded (arrow) due to shearing within the moving cohesive mass (pencil for scale). C) Pebbly debris-flow deposit

overlain by a Tbc turbidite (pencil for scale). Black layer is a large mudstone clast. The anomalous concentration of quartz

pebbles in the upper part of the debrite is attributed to reworking by a genetically related turbidity current that may have

deposited the overlying Tbc turbidite. The base of this bed (dashed line) has loaded passively into the underlying debrite,

which at the time must have been water saturated and poorly compacted (Neoproterozoic Isaac Formation, British Colum-

bia).

al flows, however, are turbidity cur-

rents.

Turbidity CurrentsBased on flow and sediment charac-

teristics, turbidity currents consist of

three distinct but not sharply bounded

parts: head, body and tail (see review

by Kneller and Buckee, 2000; Fig. 5).

The head of turbidity currents is the

sediment-rich part of the flow and the

site where most of the mixing with the

ambient fluid occurs. It is character-

ized by a sharp, overhanging nose,

above which the head slopes back in

the upstream direction due to the

resistance of the stationary overlying

fluid. This generates a strong shear

that sweeps sediment-rich fluid from

the head backward toward the body

of the current. To sustain the current,

the head must be continually provided

with new sediment supplied from the

body, which moves forward faster

than the head. Because of differ-

ences in settling velocity, coarser sed-

iment tends to accumulate in the

lower part of the head whereas finer

sediment is moved upward and back-

ward into the body of the current, the

consequence being, that with time,

the flow becomes longitudinally differ-

entiated in terms of grain size. At the

tail of the flow, sediment concentra-

tion is low and, as a consequence,

flow speed is slower and eventually

decreases to zero.

Increasingly it is being recognized

that most natural turbidity currents are

moderately to highly density stratified.

Moreover, most natural turbidity cur-

rents are typically of higher sediment

concentration and made up of sedi-

ment significantly more poorly sorted

and coarser grained than that con-

tained in laboratory currents. Unfortu-

nately, the effect of sediment concen-

tration, especially high concentra-

tions, on the nature of the current and

how it deposits sediment, remain

poorly understood. In the early

1980s, Lowe (1982) published a theo-

retical classification for turbidity-cur-

rent deposition, wherein he recog-

nized two kinds of turbidity current

and their related deposits: low-density

and high-density. Classical turbidites,

as described originally by Bouma

(1962), are interpreted to be deposit-

ed by low-density turbidity currents

(see also Mulder and Alexander,

2001). The adjective, low-density,

refers to the concentration of sus-

pended sediment in the flow, which,

based on the earlier work of Bagnold

(1954, in Mulder and Alexander,

2001), is thought to be approximately

9 % sediment volume, or less in low-

density flows (Fig. 6A). Above that

value, the closeness of adjacent

grains begins to damp fluid turbu-

lence, and hence turbulence alone is

insufficient to suspend sediment fully,

especially coarse sediment. Accord-

ingly, additional support mechanisms

like dispersive pressure, hindered

settling, and buoyancy are needed

and become increasingly more effec-

tive with high sediment concentration.

Turbidites deposited by low-density

turbidity currents consist of all, or part,

of the idealized succession described

by Bouma (1962), and reflects decel-

erating flow speed (Fig. 7). However,

if characteristics of the deposit were

formed simply by a decelerating unidi-

rectional shear flow, then for sediment

coarser than lower fine sand (>0.15

mm), upper stage plane bed (b-divi-

sion) should not be succeeded by cur-

rent ripples (c-division), but instead by

dunes (medium-scale cross-stratifica-

tion). In addition, in rare instances,

dune cross-stratified sandstone

occurs where it should – sandwiched

between upper-stage plane bed and

current-ripple cross-stratification.

Therefore, what intuitively should be

the norm is in fact the exception – but

why? One explanation that has been

advanced is that most of the deceler-

ating turbidity currents passed too

quickly through the dune stability

field. Although appealing, it has been

argued that dunes can be formed

from a flat bed in a matter of a few

tens of minutes, and that many natu-

ral turbidity currents persisted at flow

speeds in the dune stability field for

much longer periods (Arnott and

Hand, 1989). Another suggestion has

been that, under high rates of sedi-

ment fallout from suspension, upper-

stage plane bed remains stable

because the formation of dunes and

in some cases, also ripples, is inhibit-

ed (Lowe, 1988). However, in many

turbidites the ripples that formed the

c-division show a negligible angle of

climb, indicating that fallout rates are,

in fact, commonly low. An alternative

explanation for the absence of dunes

might be the effect of high near-bed

sediment concentration on the incep-

tion of dunes. Under such conditions,

bed-surface defects are prevented

from being amplified into dunes and,

as consequence, plane bed persists,

even at flow speeds that in a clear-

water flow would form dunes. Finally,

when near-bed sediment concentra-

tions have been sufficiently reduced

that bed defects can grow into bed-

forms, the flow is moving too slowly

and/or the sediment is too fine to form

dunes, and ripples form instead. In

the case of a Tbd turbidite, where nei-

ther dune nor ripple cross-stratifica-

6 ARNOTT

DRAFT FORMAT

Figure 5. A) Well-developed head of an experimental turbidity current. B) Line diagram illustrating the typical shape and

velocity profile of a turbidity current (modified after Kneller and Buckee, 2000). Note that, unlike an open-channel flow such

as a river, the velocity maximum occurs in the lower part of the flow. Note also the extensive mixing (due to interfacial

instability) that occurs along the upper part of the current.

tion is present, it is argued that sedi-

ment concentration remains suffi-

ciently high for long enough that nei-

ther dunes or current ripples form

and plane bed remained stable until

the end of traction transport.

A significant part of the sand- and

gravel-rich deep-marine sedimentary

record consists not of classical

Bouma turbidites, but instead of

structureless, normally graded and,

less commonly, ungraded or inverse-

ly graded sandstone and conglomer-

ate, which equate to the Ta division of

Bouma (1962), or the S3/R3 division

of Lowe (1982). In general, such

beds are several decimeters to a few

meters thick (Fig. 6C) and are either

amalgamated or separated by a thin

finer grained interval. Also, beds are

characteristically poorly sorted and

coarse-tail graded, wherein only the

coarsest part of the grain-size distri-

bution fines upward, typically with an

upward decrease in the abundance

of the coarsest grains (e.g., Sylvester

and Lowe, 2004; Fig. 6D). Mudstone

intraclasts are common and, in many

cases, are concentrated near the top

of the bed. In addition, water-escape

features, including pillar and dish

structures, are common in some

beds. However, in spite of their abun-

dance, the origin of structureless

beds remains a major source of

debate in the geological literature

(see, e.g.,, Stow and Johansson,

2000), although there is growing con-

sensus that they are deposited rapid-

ly from high-concentration (high-den-

sity) suspensions – sedimentation

being so rapid that any lamination

that might normally be produced by

bed-load transport is not visible.

Under these conditions sediment

‘raining’ from suspension (i.e.,

capacity-driven sedimentation; seeChapter 4) entraps fine-grained sedi-

ment that otherwise could be main-

tained in suspension, but neverthe-

less gets incorporated into the accu-


DRAFT FORMAT

Figure 7. Line diagram illustrating a complete idealized turbidite (after Blatt et al., 1984). The lowermost a-division con-

sists of normally graded sandstone or conglomerate deposited from suspension, which then is overlain by the planar strat-

ified b-division (upper plane bed). This, in turn, is overlain by small-scale cross-stratified sandstone (formed by current

ripples) of the c-division, overlain by the subtly to well-interlaminated sandstone/siltstone and mudstone d-division capped

finally by mudstone of the e-division. Photo on right is an outcrop example of a Tbcde turbidite that ranges from medium-

grained sandstone at its base to silty mudstone at the top. The a-division is missing because deposition started with bed-

load transport.

Figure 6. Idealized sediment-concentration profile in a low concentration A) and high concentration B) turbidity current.

C) Succession of stacked decimeter- to meter-thick Ta/S3 beds. Inset depicts a typical structureless, coarse-tail graded

Ta/S3 bed. D) Coarse-tail, graded, structureless Ta/S3 bed. Arrows in C, D indicate stratigraphic tops.

mulating bed, producing a relatively

poorly sorted deposit (e.g., Sylvester

and Lowe, 2004). With time and a

reduction in sedimentation rate, beds

become better sorted and, in some

cases, are overlain by tractional sedi-

mentary structures, especially planar-

lamination, producing Tab-type suc-

cessions; but what about the origin of

beds that do not grade upward into

planar lamination or other tractional

structures? The most obvious expla-

nation would be that the finer grained

upper part of the bed was eroded by a

later transport event. Although

appealing, it cannot explain beds

overlain by a fine-grained, typically

mud-rich, layer. Here, the lack of trac-

tional structures at the top of the bed

must be related to highly efficient sed-

iment bypass following the earlier

episode of rapid sediment fallout.

During bypass, sediment fallout all

but ceases and dunes and ripples are

prevented from forming by the main-

tenance of high sediment concentra-

tion, which lasts until low-energy con-

ditions and fine-grain sediment fallout

takes place from the tail of the flow.

DEEP-MARINE ARCHITECTURAL

ELEMENTS

The origin and characteristics of

deep-marine clastic systems depend

on a complex assemblage of auto-

genic and allogenic processes,

including: changes of global sea level,

tectonics, sediment flux and composi-

tion, and the nature of the sediment

supply system. For example, sedi-

ment supply controls the volume and

internal stratigraphy of the system,

whereas the number and nature of

sediment entry points controls its mor-

phology and sediment distribution.

Also, grain size, which is a function of

climate and provenance, controls

sedimentation patterns. Based on

these controls, Reading and Richards

(1994) classified turbidite systems

based on morphology, recognizing

point-source fans, multiple-source

ramps, and line-source aprons. Fur-

ther, based on the dominant grain

size, they also distinguished between

gravel-rich, sand-rich, mixed sand-

mud, and mud-rich systems (Fig. 8).

Gravel- and sand-rich systems tend

to be small (radius of a few to a few

tens of kilometers) and grade rapidly

to fine-grained basin-floor deposits.

Sand-mud systems, however, are

much larger (radius up to a few hun-

dreds of kilometers) and exhibit a sys-

tematic change in depositional ele-

ments and their internal stratigraphy

down the transport pathway. It is

these systems that make up much of

the sandstone-rich part of the ancient

deep-marine sedimentary record, and

accordingly form much of the subse-

quent discussion. Mud-rich systems

are the largest, and range in radius

from several tens to a few thousand

kilometers, and today represent the

most voluminuous deep-marine sedi-

mentary systems (see Table 1). Col-

lectively, mud-, sand- and gravel-rich

deep-marine sedimentary systems

have generally been termed ‘deep-

sea fans’ because of their common

semi-conical shape. However many

modern deep-sea ‘fans’, and by

extension ancient fans, are in fact

elongate or irregularly shaped, and

therefore the more generic term ‘tur-

bidite system’ is more appropriate

(Bouma et al., 1985). In addition, con-

fusion exists because the various

parts of a turbidite system, which in a

downflow direction consist of the

upper, middle and lower fan, have

been variously defined by different

authors. The recent classification by

Pirmez et al. (2000), which is based

on the modern Amazon Channel, sub-

divides the system based on the spa-

tial patterns of sediment erosion and

deposition (Fig. 9). From proximal to

distal, the fan subdivisions are: sub-

marine canyon – zone of net erosion;

upper fan – a zone of net sediment

bypass, wherein the channel thalweg

(the deepest part of the channel),

which is bounded on both sides by

constructional levees, lies at about

the same elevation as the surround-

ing seafloor (i.e., the area external to

the channel); middle fan – zone of net

sediment deposition caused by flow

expansion related to loss of flow con-

finement, and accordingly, where the

thalweg lies generally above the sur-

rounding seafloor; and lower fan – the

area lying downflow from the middle

fan where the rate of net deposition

decreases and the thalweg elevation

more closely approximates the eleva-

tion of the surrounding seafloor.

Internally turbidite systems are

made up of an assemblage of deposi-

tional elements, which according to

Mutti and Normark (1991), “are the

basic mappable components of both

modern and ancient turbidite systems

and are characterized by a distinctive

assemblage of facies and facies

associations”. In this chapter, four

depositional elements are discussed

in the context of a point-sourced fan

system: channels, levees, overbank/

crevasse splays, and depositional

lobes. These elements, and this kind

of turbidite system, appear to make

up much of the sandstone-dominated

part of the deep-marine record (seealso Mutti et al., 2003; Wynn et al.,2007).

Channels

As in many sedimentary systems,

channels are a common element in

deep-marine settings. A channel is a

negative topographical element pro-

duced mostly by confined turbidity

currents that transport sediment along

a major, long-term pathway. Howev-

er, channels can also be sites of sed-

iment deposition or erosion. Like

channels in the continental realm, the

condition of erosion, bypass or depo-

sition is controlled by the sediment

characteristics and boundary condi-

tions of the system – changes in one

or both of these parameters will effect

a change in the channel system.

Deep-marine channels, like fluvial

channels, continually seek a longitu-

dinal profile graded to a base level,

which in the deep marine is generally

taken to be gravity base, but, more

practically, is defined as the position

where the flow becomes unconfined

at the upcurrent end of the terminal

lobe (e.g., Pirmez et al., 2000).

Where the channel gradient and sed-

iment-transporting flows are in equi-

librium, channels migrate laterally

along a plane parallel to the equilibri-

um profile and most of the sediment is

bypassed to areas farther downflow.

Erosion and deposition, on the other

hand, represent conditions where the

channel profile lies above and below

the graded profile, respectively (seeChapters 2 and 6).

Based on observations from mod-

ern and ancient systems, three kinds

of deep-marine channels are recog-

nized: erosional, depositional and

mixed erosional-depositional (Fig.

10). Erosional channels are bounded

by a scour surface that clearly trun-

8 ARNOTT

DRAFT FORMAT

cates older strata. In contrast, depo-

sitional channels are bounded by

well-developed channel-margin lev-

ees that, with time, become progres-

sively elevated above the surround-

ing seafloor. Mixed channels show a

combination of levee deposition and

channel-axis erosion, and, in which,

the channel floor may lie above, or

below, the level of the adjacent

seafloor.

Channels can also be classified on

the degree of confinement of the

channel, which typically changes

systematically downflow (e.g., Posa-

mentier and Kolla, 2003). Highly

confined channels are those in which

flow is contained mostly within the


DRAFT FORMAT

Figure 8. Idealized models for deep-marine sedimentary systems based on volume, caliber and nature of sediment input

(from Stow and Mayall, 2000, based on Reading and Richards, 1994).

Figure 9. A) Plan view of the Amazon Fan (modified after Flood et al., 1991). Note the basinward change from a single-

thread channel to a distributive network, the longest system being the modern Amazon Channel, which extends 900 km

beyond the present shelf-slope break). B) Basinward transect along the modern Amazon Channel (from Pirmez et al.,2000). Subdivision of the system is based on the relationship between the elevation of the channel thalweg and levee crest

relative to the adjacent seafloor surface. Datum (zero line) is the general level of the seafloor outside of the channel

channel. Generally, these channels

occur in the upflow parts of a turbidite

system and include submarine

canyons and erosional channels, the

dimensions of which are commonly of

the order of kilometers to several kilo-

meters wide and >100 m deep. Far-

ther downflow, these channels

change into leveed-channel systems,

where channel dimensions range up

to several to a few kilometers wide

and generally up to about 100 m deep

(i.e., the relief between the thalweg

and the adjacent levee crests). Here,

flows and, in particular, the upper

parts of flows, can escape the con-

fines of the channel to build channel-

bounding levees. These systems, in

turn, are succeeded downflow by a

complex of poorly confined channels

associated with a lobate sedimentary

body, commonly termed a deposition-

al lobe or fan, at the downstream end

of the channel.

Submarine CanyonsSubmarine canyons are the primary

conduits for sediment transport into

the deep sea. In the modern oceans,

submarine canyons range up to more

than 2.5 km deep and 100 km wide

(see Normark and Carlson, 2003),

whereas the examples of ancient

canyons described so far are typically

much smaller, reaching only slightly

more than 1 km deep and 10 km

wide. The fill of a submarine canyon is

typically stratigraphically complex and

lithologically variable. In part, this

relates to temporal and spatial differ-

ences in sediment source, which

varies from local canyon-wall col-

lapse, to an up-dip feeder system

sampling a shelf and/or continental

sediment source. Fills dominated by

local wall collapse are typically fine-

grained with common mud-rich mass-

movement (slump and slide) and

debris-flow deposits (Fig. 11A, B); an

example is provided by the subsur-

face late Paleocene Yoakum and

Lavaca canyon fills (Galloway et al.,1991). Fine-grained sediment

deposited from suspension may also

occur as a drape infilling part or all of

the canyon relief. Coarse sandstone

and conglomerate occur as isolated

elements in mud-rich fills, but also

dominate some fills, especially those

formed in tectonically active areas

(Fig. 11C). In most cases, coarse sed-

iment occurs as thick-bedded, struc-

tureless beds deposited by high-con-

centration turbidity currents (Fig.

11D).

Erosional ChannelsAlthough the smaller scale erosional

channels down-system from subma-

rine canyons also owe their existence

to erosion, their fill is different,, and

typically consists of a significant pro-

portion of sand- and gravel-rich strata

deposited by turbidity currents and

other frictional flows. Although the

geological literature is replete with

erosional channel-fill models (e.g.,

Beaubouef et al., 1999; Mayall and

Stewart, 2000; Samuel et al., 2003),

the stages common to most models

include: channel inception, sediment

bypass, channel filling and channel

abandonment. Channel inception is

marked by a period of successive

flows with high transport efficiency

that scour-out a throughgoing topo-

graphical feature that serves as the

conduit for later flows. Sediment

transported during this stage is car-

ried further basinward and deposited

in more distal areas. This stage is

then succeeded by the channel-bypass stage wherein flows range

between complete bypass (with no

further erosion) to incomplete bypass.

Incomplete bypass is commonly indi-

cated by the deposition of laterally

discontinuous beds (due to erosion by

subsequent flows), intercalation of

coarse- and fine-grained deposits, the

common patchy occurrence of trac-

tional sedimentary structures, espe-

cially dune cross-stratification and

coarse-grained lags, and thin drapes

of fine-grained sediment, a heteroge-

neous assemblage of lithofacies

herein termed the ‘bypass facies’. The

next stage, channel fill, is character-

ized by a change toward flows that

have a lower transport efficiency that

initiate sediment deposition within the

channel, eventually filling part, or all,

of the channel. Later, as a result of a

diversion of flow at a point upstream

(avulsion), or diminution of flow for

other reasons (e.g., sea-level rise),

the channel system is abandoned and

becomes a site of mostly fine-grained

deposition that drapes any residual

topography. Superimposed on this

idealized succession of events are

episodes of reactivation, particularly

during the channel-filling stage, which

10 ARNOTT

DRAFT FORMAT

Figure 10. Nomenclature for deep-marine channels (after Pickering et al.,1995).

serve to temporarily rejuvenate the

system but not reverse its long-term

filling trend (e.g., Samuel et al.,2003).

Erosional channels are commonly

reported from seismic images, and

much less commonly from the

ancient outcrop record. This disparity

may be the result of two factors:

1. many seismically resolved exam-

ples of erosional channels are

very large, commonly measuring

more than 100 m deep (thick) and

a kilometer to many kilometers

wide (e.g., Mayall and Stewart,

2000; Deptuck et al., 2003; Abreu

et al., 2003) and therefore are on a

scale significantly larger than most

outcrops; and

2. if strata inside and outside the

channel margin are of similar

lithology, or are poorly exposed,

recognizing the channel-bounding

surface in outcrop may be difficult,

even though the channel fill and

the surrounding deposits may

have distinctly different acoustical

properties and therefore are easily

differentiated on seismic.

Seismic images also show that ero-

sional channels are commonly bor-

dered by well developed levees.

Deptuck et al. (2003) recognized two

end-member kinds of erosional

channels and their related levee

deposits. Large-scale channels and

their related ‘outer’ levees represent

the master erosional channel that is

typically several kilometers wide and

>100 m deep (Fig. 12). These fea-

tures were infilled, partly or com-

pletely, by smaller scale channels

with their associated ‘inner’ levees,

and are described in the next sec-

tion.

Heterogeneous strata deposited

by incompletely bypassed flows form

the basal unit of the channel fill (Fig.

13). This is commonly succeeded by

contorted seismic reflectors interpret-

ed to represent slump/slide (mass-

movement) and debris-flow deposits

produced by collapse of the channel

margins. These strata are overlain by

a thick succession of sandstone or,

less commonly, conglomerate that

forms a tabular or sheetlike unit com-

prising amalgamated small-scale

channel fills. During this early stage

of channel fill, these smaller scale

channels tend to be poorly confined

with low to moderate sinuosity and a

high width-to-depth ratio (see leveed

channels below). Upward, these

channels are succeeded by channels

with higher sinuosity and a lower


DRAFT FORMAT

Figure 11. A, B) Submarine-canyon strata consisting of contorted mudstone-rich slump deposits most probably sourced

from local canyon-wall collapse (Pennsylvanian Jackfork Formation, Arkansas). C) 135-m-deep canyon incised into thin-

bedded and very thin-bedded continental-slope turbidites. Canyon fill consists mostly of graded, structureless coarse-

grained sandstone and conglomerate. D) (Hector Formation, Lake Louise, Alberta).

width-to-depth ratio (see later: con-

fined sinuous channels). Further-

more, these channels commonly

show an upward increase in angle of

climb, reflecting an increased rate of

channel aggradation relative to the

rate of lateral migration (Peakall et al.,2000). Eventually, the entire channel

system, which at this point has been

either partly or completely filled, is

abandoned and blanketed by a layer

of thin-bedded turbidites and

hemipelagic suspension deposits!

Leveed ChannelsLeveed channels are more commonly

reported from outcrops because of

their smaller size compared to large

erosional channels,. Although smaller,

leveed-channel fills still range up to a

few kilometers wide and 100 m thick

(Fig. 14). In addition to occurring as

an independent channel element, lev-

eed channels are also important

stratal components in larger scale

erosional channel fills as described

above.

Leveed channels typically have a

sinuous planform (Fig. 14A, B) with

levees of varying development along

their margins. Based on the degree of

confinement, two end-member types

of leveed channels are recognized:

poorly confined and highly confined.

In both types, channel-fill strata termi-

nate abruptly along an erosion sur-

face defining the outer-bend margin

of the channel. The strata, cut by the

erosion surface, are either genetically

related levee deposits or strata relat-

ed to an older channel. Along the

inner-bend side, however, channel-fill

strata either grade continuously into

levee deposits in poorly confined

channels, or onlap them in highly con-

fined channels. Also, owing to the

effects of (fluid) inertia and the result-

ing tendency of a current to continue

along a straight line while the channel

floor bends beneath it, levees and

levee deposits are always best devel-

oped along the outer bend of all chan-

nels.

Poorly Confined Leveed ChannelsChannel deposits

The base of poorly confined leveed

channels is commonly asymmetric,

with a steeper margin along one side

(analogous to the cut bank of a sinu-

ous fluvial channel; Fig. 15). In addi-

tion, the channel base is often charac-

terized by a step-flat morphology,

indicating the episodic but systematic

step-like lateral migration of the entire

channel system (e.g., Eschard et al.,2003; Navarro et al., 2007). This

asymmetry is present also in the

nature of the relationship between the

channel-fill and adjacent levee strata.

Along the steep margin, levee strata

are either in erosional contact with

channel strata (see ‘outer-bend’ levee

deposits below) or are separated by a

thin, fine-grained bypass unit

(Beaubouef et al., 1999). In contrast,

on the opposite side of the channel

(analogous to the point bar of a sinu-

ous fluvial channel), channel-fill strata

either onlap or grade laterally into

levee deposits (see ‘inner-bend’ levee

deposits below). The fill of leveed

channels is, in fact, composed of the

fill of myriad smaller channels, which,

because of extensive amalgamation,

are difficult to trace individually in out-

crop. Nevertheless, the fill of an indi-

vidual channel is of the order of a few

to several meters thick and tens to, at

most, a few hundred meters wide.

The 3-D amalgamation of these chan-

nels forms a channel unit, which is

probably the most readily identified

channel succession in outcrop. Chan-

nel units range from several meters to

a few tens of meters thick and com-

monly show a well-developed fining-

and thinning-upward trend. As noted

above, their base is typically marked

by the presence of coarse sediment

and abundant mudstone intraclasts.

Also, in many cases, coarse, amalga-

mated strata grade abruptly laterally

into finer, more stratified deposits.

Strata in the axial part are generally

thick- to very thick-bedded, coarse-tail

normally graded or structureless con-

glomerate or sandstone. Beds are

typically amalgamated with variable

lateral continuity (ranging from tens to

hundreds of meters laterally). These

coarse-grained deposits are Ta tur-

bidites deposited by gravel- and

sand-rich, high-concentration turbu-

lent flows. Toward the margin of chan-

nel units, especially those higher in

the leveed-channel fill, coarse-

grained strata tend to thin rapidly

(generally over <100 m), and grade

from amalgamated sandstone to less

amalgamated, more thin- to thick-

12 ARNOTT

DRAFT FORMAT

Figure 12. Uninterpreted and interpreted seismic profile of a channel-levee sys-

tem in the Indus Fan, Bay of Bengal (after Deptuck et al., 2003). Outer levees

bound master channels that are several kilometers wide, which are filled with

the deposits of smaller channels (HARs) and their associated inner levees.

bedded sandstone with intervening

mudstone. Sandstone beds consist

of Ta and Tab turbidites and are inter-

calated with thin-bedded Tcd, Tcde

Bouma turbidites. The thinning and

fining of strata toward the margin of

channel units suggest that the high-

er-energy, axial part of the channel is

flanked by lower-energy conditions.

Levee deposits

Subaqueous levees build upward by

the addition of sediment when flows

overtop the margins of the channel.

Flow overspill and related levee

aggradation occurs in three ways:

flow stripping, inertial overspill, and

continuous overspill. Flow stripping

and inertial overspill occur at channel

bends, and preferentially along the

outer bank. In the case of flow strip-

ping, the upper fine-grained part of

the flow becomes separated from the

lower, coarse-grained part that

remains confined to the channel.

Inertial overspill occurs when an

energetic flow is unable to follow the

sinuous thalweg and ‘runs-up’ the

channel margin, allowing even the

lower parts of the flow to escape the

channel. Continuous overspill takes

place where the thickness of the flow

exceeds the depth of the channel,

leading to loss of the flow above the

height of the levee along both sides

of the channel. Levee growth along

the inner-bend side of the channel

and the straight segments between

channel bends comes about by con-

tinuous overspill. Upon escaping the

channel, the flow expands rapidly

and collapses, resulting in elevated

rates of sedimentation immediately

adjacent to the channel, with rapidly

decreasing rates of deposition far-

ther from the channel. This lateral

variation in average sedimentation

rate gives levee deposits their dis-

tinctive ‘gull wing’ geometry on seis-


DRAFT FORMAT

Figure 13. Idealized model for the fill of an erosional channel (after Mayall and

Stewart, 2000).

Figure 14. Sinuous leveed channels. A) Low sinuosity, large-scale (width >1 km) (Pleistocene) Einstein Channel in the

Gulf of Mexico (modified after Posamentier and Walker, 2006). B) High sinuosity, small-scale (100s m wide) modern Ama-

zon channel-levee system in about 3500 m water depth (from Pirmez et al., 2000). Seismic profiles across multiple (C)

and single (D) channel-levee complexes in the Amazon Fan (from Hiscott et al. (1998) and Piper and Normark (2001)

respectively). In C, note the organized offset stacking, termed compensational stacking, of successive complexes (i.e.,

younger channels are located preferentially in the topographic lows between two older systems). Note also the extensive

blanket of mass-transport deposits that underlies the uppermost stack of channel-levee complexes.

mic images (Fig. 14C). As a levee

aggrades, the relief between the

channel floor and levee crest increas-

es, eventually allowing only the upper,

more dilute portion of the through-

going flows to overspill. This process

is responsible for the upward thinning

and fining trends associated with

many levee deposits.

Owing to their fine-grained nature,

levee deposits are typically not well

exposed, although notable exceptions

exist. Along the outer-bend side of

the channel, levee deposits tend to be

thick and sand-rich, whereas those on

the inner-bend side are significantly

finer and thinner. Paleocurrents, typi-

cally measured from the c-division of

turbidites, are generally oriented very

oblique to paleoflow in the main chan-

nel. Closest to the channel and along

the outer bend, (the proximal levee

facies) strata consist mostly of thin- to

medium-bedded, fine to medium-

grained sandstone Tbc turbidites inter-

stratified with thin-bedded, very fine-

to fine-grained sandstone/siltstone

Tcde turbidites and common thick Ta

sandstones (Fig. 16A). In many mod-

ern and interpreted ancient proximal-

levee deposits, small-scale (ripple)

cross-lamination commonly shows

evidence of vertical aggradation (i.e.,

climbing) because of sediment fallout

from suspension as the flow expands,

with the angle of climb generally

decreasing away from the channel

margin. However, the presence of

climbing ripples is not a universal fea-

ture of levee deposits as shown by

the Windermere Supergroup (Navarro

et al., 2007). This may be due to a

somewhat coarser grain size and/or

poorer sorting, but the cause remains

uncertain.

With increasing distance from the

channel, the strata become thinner.

Typically, thicker beds thin rapidly

over 100s of meters whereas thinner

beds show little lateral change. As a

consequence, strata occurring a few

to several 100s of meters from the

channel in the distal levee are com-

posed predominantly of thin-bedded,

very fine to fine-grained sandstone/

siltstone Tcde turbidites (Fig. 16B). The

c-division consists of one to at most a

few sets of non-climbing ripple cross-

stratification. Locally, distal levee stra-

ta are intercalated with overbank or

crevasse-splay deposits (see below).

Levee strata on the inner side of

channel bends are distinctly thinner

and generally finer grained than those

on the outer side of the bend.

Deposits consist predominantly of fine

sandstone and siltstone Tcde tur-

bidites, which, locally, are interbedded

with thicker sandstone beds. These

latter beds consist typically of medi-

um- to thick-bedded, fine to medium

sandstone Tc turbidites composed of

multiple (3 or 4) ripple cross-stratified

sets. Like on the opposite side of the

channel, strata thin and fine away

from the channel, but the rate of

change is much greater and reflects

the lower-energy nature of over-

spilling flows on the inner-bend side

of the channel. An important differ-

ence between inner- and outer-bend

levee strata is that, in many places,

inner-bend strata are continuous with

and grade laterally into channel-fill

strata, indicating a continuum

between channel and levee deposi-

tion.

Highly Confined Leveed ChannelsChannels of this variety are signifi-

cantly smaller than those described

above –channel width and depth are

of the order of tens to a few hundreds

of meters and a few tens of meters,

respectively; sinuosity is generally

higher too (Fig. 17). In outcrop, highly

14 ARNOTT

DRAFT FORMAT

Figure 15. Poorly confined leveed-channel deposits.

A) Channel fill 1 is a sharp-based, up to about 80 m-

thick sandstone/conglomerate unit sharply overlain by

a second channel system (Channel fill 2), the base of

which is indicated by the dashed orange line. Note the

sharp, terraced channel base (solid orange line) along

the base of Channel fill 1 that ascends obliquely

upward toward the right, and is the result of combined

vertical and lateral channel migration. Lateral channel

migration and aggradation also causes deposits of

Channel 1 to terminate abruptly as they overstep stra-

ta of their genetically related outer-bend levee B).

Along the opposite (left) side, channel strata fine and

thin continuously into strata of the inner-bend levee

(Neoproterozoic Isaac Formation, western Canada).

(C, D) Sharp, terraced margin along the outer bend of

laterally migrating, vertically aggrading channels

(Upper Cretaceous Tres Pasos, southern Chile).

confined channels occur mostly as

disconnected channels fills that

locally are clustered laterally and/or

vertically. Clustering is attributed to

younger channels exploiting remnant

seafloor topography formed by

incompletely filled older channels

(Fig. 18). In addition, lateral-accre-

tion deposits are well developed,

indicating systematic deposition on

the inner bend of laterally migrating

sinuous channels (Arnott, 2007a;

Fig. 19A, B).

Channel bases are generally pla-

nar and horizontal with only local

small-scale (few cm) scours. In

places, however, the surface shows

a step-like geometry, rising abruptly

upward by a few to several meters.

The top of the sandy channel fill

interfingers with thinly bedded tur-

bidites (Fig. 19A, B). The strata with-

in the channel fill show a distinctive

and consistent dip of up to about 7 to

12º, and, like similar features

observed in meandering fluvial sys-

tems, are interpreted to be lateral-

accretion deposits (LADs) formed on

the inner bend of a laterally migrating

sinuous channel. The fill of the chan-

nel consists of lower and upper parts.

Beds in the lower part are typically

thick- and very thick-bedded, and

generally amalgamated (Fig. 20).

They consist mostly of sharp-based,

graded sandstone and less common

granule or fine pebble conglomerate

(Fig. 19C). Bases of beds are com-

monly scoured and, in places, com-

pletely erode underlying beds. Mud-

stone generally occurs as localized

patches of intraclasts. Stratigraphi-

cally upward, sand-/gravel-rich strata

change little in grain size but thin,

typically becoming medium bedded;

mudstone intraclasts are absent. In

the upper part of the channel fill, fine-

grained strata (mudstone and thinly

bedded turbidites) become interstrat-

ified with the coarse-grained deposits

(Figs. 19A, B). It is noteworthy that

coarse-grained beds terminate

abruptly up-slope (Figs. 19D, E, 20),

which contrasts markedly with the

gradual lateral trend observed in

poorly confined leveed channels.

Also, very near the terminus of each

bed, coarse strata consist of a small

number of graded, poorly sorted

beds capped by planar laminated or

dune cross-stratified sandstone.

Down-dip from the terminus of each

coarse bed, fine-grained strata are

truncated as the coarse beds amal-

gamate (Figs. 19E, 20). Near their

termination, fine-grained strata con-

sist of almost complete Bouma-divi-

sion turbidites (Fig. 19E) that up dip

thin, fine, and become dominated by

upper-division turbidites (Tcde).

The coarse-grained beds in the

LADs represent coarse sediment

deposition on the lower part of the

channel margin. In the LADs fine-

grained beds fine and thin abruptly


DRAFT FORMAT

Figure 16. A) Proximal outer-bend levee deposits composed of medium- and

thin-bedded turbidites. Thicker beds typically consist of Tbcde turbidites that thin

rapidly away from the channel margin, whereas thinner beds consist of Tcde tur-

bidites that change little in thickness laterally. B) Thin to very thin-bedded Tcde

and Tde turbidites of the distal levee (field notebook for scale). As in the proximal

levee, thin beds in the distal levee show little thickness change laterally. Photos

A) and B) are from Neoproterozoic Isaac Formation, western Canada.

Figure 17. Highly sinuous small-scale (<1 km wide) leveed channel system, the

Joshua Channel (Pleistocene), Gulf of Mexico (Posamentier and Kolla, 2003;

Posamentier and Walker, 2006). A) Seismic horizon slice illustrating the highly

sinuous nature of the channel. Note also the meander cut-offs (oxbows) indicat-

ed by arrows. B) Seismic curvature map showing the well-developed levees that

bound the channel. Note the common slump scars along the levees on both

sides of the channel.

Figure 18. Organized channel pattern created by the lateral offset pattern of

successive channel-fill elements (images courtesy of Henry Posamentier). Inset

sketch from Posamentier and Kolla, (2003).

beyond the termini of the coarse

LADs, and represent the fine-grained

part of the inner-bend levee onto

which the coarse-grained LADs

onlap. On the opposite, or outer-bend

margin of the channel, coarse-

grained LADs typically terminate

abruptly against fine, thin-bedded tur-

bidites of a slightly older levee sys-

tem.

The rhythmic interstratification of

coarse- and fine-grained beds, which

is especially well developed and pre-

served in the upper part of channel

fills, reflects recurring changes in sed-

iment transport through the channel

system. Deposition of the fine beds

most likely represents periods of high-

ly efficient turbidity currents that

bypass the channel bend and trans-

port much of their coarse sediment

load farther downdip. Periodically,

these conditions are interrupted by

episodes of less efficient turbidity cur-

rents that result in deposition of a

small number of beds that make up

each coarse interval. Currently, the

cause for the rhythmic alternation of

fine and coarse beds in the LADs is

poorly understood, but most likely

relates to repetitive changes in local

or regional flow and/or channel condi-

tions (Arnott, 2007a).

Overbank and Crevasse Splays

Overbank splays are lobate and

sheetlike features formed by ener-

getic flows that overtopped and

escaped the channel in an unconfined

manner. Crevasse splays, on the

other hand, are larger scale lobe-

16 ARNOTT

DRAFT FORMAT

Figure 19. A, B) Lateral-accretion deposits (LADs) in a highly confined leveed channel from the Neoproterozoic Isaac For-

mation, western Canada. Note the dipping LADs, which are inclined at an angle of 7-12º, and are interpreted to have accu-

mulated on the inner-bend margin (point bar) of a laterally accreting sinuous channel (from Arnott, 2007a). Also, note the

sharp, planar basal contact, but the interfingering of the sandy beds with thin, mud-rich deposits at the top. C) Structure-

less, graded sandstone near the base of LADs. At their upper end, such coarse beds pinch-out abruptly upward (open

arrow in E) into thin-bedded Tcde and Tde turbidites D), which in turn become truncated downward by the coarse beds (solid

arrows in E).

Figure 20. Idealized model of lateral-accretion deposits (LADs) formed by a lat-

erally migrating deep-marine sinuous channel. Each coarse and fine LAD con-

sists of several beds and indicates that there were longer term repetitive alter-

nations in the nature of the flows.

shaped features formed immediately

downflow of a crevasse channel that

is incised into the channel margin

and proximal levee. To date, few

unequivocal examples of crevasse-

or overbank-splay deposits have

been reported from the geological

record. Work on the Amazon Fan

(e.g., Flood et al., 1991) identified a

distinctive seismic facies that com-

monly underlies channels (Fig. 21A).

Termed a HARP (high-amplitude

reflection package) because of its

high acoustic impedance, these stra-

ta have a sheet-like geometry and

are assumed to be sand-rich. These

packages are interpreted to repre-

sent crevasse splays formed down-

flow of a breach in the levee of an

adjacent active channel (Fig. 21B,

C). In the Windermere Supergroup,

sand-rich strata interpreted to be cre-

vasse-splay deposits are interstrati-

fied with fine-grained, thin-bedded

basin-floor and distal-levee deposits

(Fig. 22; Arnott 2007b). These cre-

vasse-splay successions, which are

up to several meters to a few tens of

meters thick, consist of decimeter to

several meter-thick units of medium-

to thick-bedded structureless sand-

stone containing common mudstone

intraclasts. In places, these strata

are interbedded with units composed

of upper division (Tcde) sandstone tur-

bidites. Structureless sandstones

are poorly sorted and coarse-tail

graded with a matrix of fine sand-

stone, siltstone and mudstone, and

are interpreted to have been deposit-

ed rapidly by capacity-driven deposi-

tion immediately downflow of an area

of rapid flow expansion. The thinner

graded beds are deposited on the

periphery of the collapsing sediment

cloud. The intercalation of the two

types of beds is related to the lateral

wandering of the zone of flow expan-

sion.

At its headward end, a crevasse

splay is joined to the breach in the

adjacent parent channel by a cre-

vasse channel, which forms a sharp-

based, distinctly coarser grained unit

within a background of fine-grained

levee deposits. Crevasse-channel

deposits tend to be thin (up to a few

meters thick) and consist of amalga-

mated thick-bedded, massive or nor-

mally graded sandstone or (less

commonly) conglomerate. Discon-

tinuous beds of single-set-thick dune

cross-stratified sandstone are com-

mon also.

Overbank splays, on the other

hand, form from large magnitude,

high concentration, coarse-grained

turbidity currents that overspill the

adjacent channel without confine-

ment (Fig. 22). Their deposits con-

sist of single to multiple beds forming

units up to 2–5 m thick. Internally,

units consist of amalgamated thick-

bedded, medium-grained sandstone

turbidites that commonly comprise

complete Bouma turbidites.

Basin-floor deposits

In the proximal part of the basin floor,

leveed channels terminate downflow

in a thick, laterally extensive sedi-

ment body variously termed a depo-

sitional lobe, distributary-channel

complex, sand-sheet deposits, and

frontal-splay complex, for it is here

that highly confined flows emanating

from the leveed channels become

unconfined and depositional (Fig.

23). The depositional lobes range up

to several tens of meters thick and

100 km wide. Farther basinward,

they become finer and thinner, and

eventually are replaced by

hemipelagic and pelagic deposits.

Loss of confinement and the lateral

spreading of the flow can be the

result of a reduction in slope and/or

loss of sufficient fine-grained sedi-

ment to build channel-margin levees

(Posamentier and Kolla, 2003).

Characteristics of the transition from

channels to depositional lobes are

principally controlled by grain size. In

coarse-grained systems, the deposi-

tional lobe connects directly with the

leveed channel, but in mud-rich sys-

tems, the lobe is separated by a tran-

sitional zone marked by large scours,

sediment waves and sediment

mounds. Interpreted transition-zone

deposits in the ancient stratigraphic

record consist also of an array of

lithofacies suggesting both sediment

bypass and sediment deposition.

Bypass features include shallow

scour surfaces draped by fine-

and/or coarse-grained (lag) sediment

that commonly is planar laminated or

dune cross-stratified, and units up to

about 5 m thick composed of com-

pensationally stacked scour-based

lenticular sandstone. Depositional

features include sheetlike sand-

stones that are similar to sheetlike

splay deposits described below.

Depositional Lobes The planform geometry and size of a

depositional lobe depends principally

on the sand-mud ratio of the sedi-

ment supply. In sand- and gravel-

rich systems, these features tend to

be areally restricted because of rapid

deposition, and form discrete ele-

ments on the order of a few kilome-


DRAFT FORMAT

Figure 21. A) Seismic profile and interpretive sketch of channel-levee complex-

es in the Amazon Fan (Flood et al., 1991). High amplitude, sheet-like packages,

termed HARPs, commonly underlie the high-amplitude reflectors (HARs) that

are interpreted to be channel deposits. HARPs are interpreted to be crevasse-

splay deposits (B) formed during the initial stages of an avulsion, which in many

cases are overlain by their genetically related channel (C).

ters wide and a few meters to several

tens of meters thick that decrease in

thickness and grain size rapidly

downflow.

Recent work by Deptuck et al.(2007) showed that such late Pleis-

tocene features off the coast of Corsi-

ca are of the order of 2–19 km2 in

area, 9–20 m thick, have a length-to-

width ratio of <1–2, and show no evi-

dence of progradation but instead

backstep upslope. (A similar back-

stepping pattern is noted in proximal

mouth-bar deposits; see Chapter 10).

Based on seismic expression and

shallow piston cores, strata are domi-

nated by amalgamated sand that rep-

resents 50% of the planform area and

75% of the total volume of the deposi-

tional lobe. In contrast, more mud-

rich systems form broader, more

sheetlike features that are up to about

10 km wide and become finer and

thinner more gradually downflow.

Also, their stratigraphic composition is

significantly more complex than those

in sand-rich systems.

Based on seismic interpretation

and observations in the rock record,

such lobe systems consist commonly

of three recurring architectural ele-

ments: deep channels, shallow chan-

nels, and sheetlike splay deposits

(Figs. 23 B, C, D; 24). Deep channels

show up to a few tens of meters of

incision, with channel margins that

are locally steep (few to several

meters of relief over a lateral distance

of only a few tens of meters; e.g.,

Meyer and Ross, 2007). These chan-

nels are filled with a variety of lithofa-

cies, including heterogeneous assem-

blages of coarse- and fine-grained

strata related to incomplete bypass,

and also thick-bedded, amalgamated

sandstone (e.g., Johnson et al., 2001;

Fig. 24C). Deep channels are inter-

preted to represent the principal con-

duit that supplies sediment to the

depositional-lobe complex. Shallow

channels occur downflow of the deep

channels; these channels show

minor relief along their base (scour

only a few meters deep over a few

hundred meters laterally), and exhibit

a consistent internal stratigraphy,

which mostly is aggradational. In the

channel axis, strata range up to about

5–10 m thick and consist of thick- to

very thick-bedded amalgamated

sandstone (Ta beds), which, when

traced laterally over a few hundred

meters, show a gradual but systemat-

ic change to thick-/medium-bedded

complete turbidites that become pro-

gressively thinner, finer and less com-

plete turbidites, and eventually pass

into single-set thick-, thin-bedded,

fine-sandstone Tcde turbidites (Fig.

24B). Infilling of a shallow channel is

probably because of reduced efficien-

cy caused by deposition farther down-

flow in the sheetlike splay element,

which then is followed by, or is coeval

with, the initiation of a new channel

and lobe element elsewhere. Out-

board of the terminus of the shallow

channels, flows become unconfined

and highly depositional, forming the

sand-rich sheetlike splay element that

farther downflow forms a distal apron

of thin- and very thin-bedded fine-

grained turbidites intercalated with

hemipelagic and pelagic mudstone.

Currently, details of this transition are

poorly understood. Nevertheless, in

the more proximal sand-rich part of

the depositional lobe, strata consist of

laterally extensive, tabular units that

range from a few meters to a few tens

of meters thick, but typically are of the

order of 15 m thick (e.g., Meyer and

Ross, 2007; Fig. 24D). Typically, the

base of each unit is marked by a

sharp increase in grain size com-

pared to the underlying strata, which,

in many cases, consists of a few

meter-thick succession of structure-

less, mudstone-intraclast-rich,

coarse-tail-graded sandstone beds

that resemble overbank and cre-

vasse-splay deposits. The sheetlike

splay element consists almost entire-

18 ARNOTT

DRAFT FORMAT

Figure 22. Aerial photo of strata from the Neoproterozoic Isaac Formation,

western Canada showing overbank-splay deposits (OB1.1-OB1.4), and cre-

vasse-splay deposits (CS2.1-CS2-3) and their genetically related crevasse-

channel fills (C1, C2) – note, strata are vertically dipping (see Arnott, 2007b).

Overbank-splay deposits occur as units, one to a few beds thick, comprising

medium- to thick-bedded, coarse-grained, more complete turbidites (Tbcde)

interbedded with thin-bedded, upper-division turbidites (Tcde, Tde). This succes-

sion is then overlain abruptly by a thick crevasse-splay deposit consisting of

three several-meter-thick packages of matrix-rich structureless sandstone

(CS2.1–CS2-3) intercalated with few-meter-thick packages composed of classi-

cal turbidites. Crevasse-splay deposits are then sharply overlain by a lateral-

accreting channel deposit that comprises two separate channel fills (C1, C2).

ly of amalgamated sandstone that

can be mapped over several kilome-

ters. Although difficult to discern

because of amalgamation, beds are

generally thick to very thick bedded

and seemingly form a random pat-

tern of bed-scale cut-and-fill with no

organized internal architecture such

as compensational stacking of beds.

CONTOURITES

Although first theorized in the mid-

1930s to exist, the occurrence of

contour-following deep-sea currents,

or contour currents, was first demon-

strated some 30 years later on the

continental rise off eastern North

America. Deposits of these currents,

which are known as contourites,

were shown to be characterized by

features that distinguish them from

better known turbidites, and were

later discovered to be the principal

constituent in large (tens to hundreds

of km long, few tens of kilometers

wide, and up to over 1 km high),

elongate, slope-parallel sediment

bodies termed sediment drifts.

These deposits owe their existence

to deep-water bottom currents that

form part of the global thermohaline

or wind-driven circulation system.

These semi-permanent currents gen-

erally flow parallel to the slope, but

locally, especially because of topo-

graphical effects, can be diverted

obliquely up or down the slope.

In polar regions, cold surface

water and also more saline water

formed by surface-water freezing

descends to the basin floor initiating

a large-scale (global) flow system. As


DRAFT FORMAT

Figure 23. A) Seismic image of a Pleistocene depositional-lobe complex developed at the downflow terminus of a leveed

channel, Gulf of Mexico (Posamentier and Kolla, 2003). B) Isopach map of a Pleistocene depositional-lobe complex,

Indonesia (Saller et al., 2008). C) Map illustrating the internal stratigraphic complexity of the lobe complex in (B), which

consists of 18 (A-R) discrete lobe deposits, which formed during the lowstand systems tract. During the late lowstand to

early transgressive systems tract, however, a change to more mud-rich flows caused the leveed channel at the headward

end of the lobe complex to extend basinward (‘final upper fan channel’), which caused lobe deposition to shift basinward

too. D) Detail of a single depositional lobe (Lobe D – outlined in red in C) showing that it too is composed of several even

smaller, discrete splay elements (labeled 1-6). It is these smaller splay elements that are generally observed in the ancient

record and here are termed sheet-like splay elements.

a consequence of the rotation of the

Earth and the Coriolis effect, these

currents, which have speeds of about

1–2 cm/s, become deflected toward

the western side of ocean basins.

There, they are constrained by the

continental slope and their speed

increases to about 10–20 cm/s and,

where locally constricted, can reach

speeds of over 2 m/s. However, even

at their generally lower speeds, con-

tour currents are approximately at the

threshold velocity for very fine and

fine-grained sand, and hence are an

important sediment-transporting

agent in the deep sea.

The characteristics of contourites

are controlled largely by sediment

supply (Stow et al., 2002). They can

be composed of terrigeneous clastic,

volcaniclastic, or carbonate sediment,

and grain size can range from mud to

sand and admixtures of both,

although mud and silt are most com-

mon. In addition, gravel particles

occur, but are restricted to local areas

of high energy and attendant seafloor

reworking and winnowing. (Note that

in most cases gravel is brought to the

area by glacial ice rafting). Sorting is

generally moderate to good, except in

areas with low current speeds and

slow sedimentation where bioturba-

tion has mixed the deposit. Moreover,

bioturbation, which generally is domi-

nated by forms of the Nereites inch-

nofacies (see Chapter 3), can be

moderate to intense, and can destroy

primary sediment layering. Traction

structures are common and, depend-

ing on flow speed and grain size,

include small-scale (current ripple)

and large-scale (dune) cross-stratifi-

cation, and also scour marks. Pale-

ocurrents are commonly, but not

strictly, parallel to the slope. Con-

tourites often occur as composite

units 20-30 cm thick (Fig. 25A). Most

successions consist of a basal

upward-coarsening interval consisting

of muddy to silty to sandy contourites

overlain by a unit that fines upward to

muddy contourites (Gao et al., 1998;

Stow et al., 2002). This upward

change suggests a systematic tempo-

ral change in flow speed and/or sedi-

ment supply, and which recent evi-

dence suggests occurs on time

scales that closely parallel

Milankovitch periodicities, suggesting

a relationship between orbital forcing

of climate and changes in bottom-cur-

rent velocity (Stow et al., 2002).

Where turbidites and contourites co-

exist, they may be difficult to differen-

tiate. However in an interpreted Plio-

Pleistocene turbidite–contourite suc-

cession, Ito (1996) suggested that

contourites can be identified based on

minor inverse grading within the ripple

cross-stratified unit, in addition to

intercalated layers or drapes of mud

within ripple cross-stratified sand-

stone/siltstone units (Fig. 25B). More-

over, contourites commonly contain

internal erosion surfaces, typically

lack an ordered vertical succession of

features like a classical turbidite, and,

where interstratified with turbidites,

are bounded sharply on their base

and top. Collectively these features

indicate fluctuating bottom-current

speed, and, in many cases, the oscil-

lation between traction and suspen-

sion sedimentation.

SEQUENCE STRATIGRAPHY OF

DEEP-WATER DEPOSITS

Early sequence-stratigraphic models

for the deep-marine siliciclastic

deposits were based on wide passive

continental margins with a well-devel-

oped shelf–slope break. Along such

margins, the supply of continent-

derived sediment into deep water

depends on the state of the continen-

tal shelf, which, in large part, is con-

trolled by the position of relative sea

level. During highstand when the

shelf is wide, clastic sediment is

sequestered in marginal-marine and

continental settings and sediment flux

into the deep sea is much reduced.

Lowstand, on the other hand, and

especially when rivers reach the

shelf–slope break, is a time of volumi-

nous sediment supply and active dep-

osition in deep water. It has been

pointed out, however, that in situa-

tions where the continental shelf is

narrow, for example along the coast

20 ARNOTT

DRAFT FORMAT

Figure 24. A) Ancient basin-floor deposits (Neoproterozic upper Kaza Group,

western Canada) comprising 3 principal depositional elements: shallow chan-

nels, deep channels, and sheetlike depositional lobes — note strata are vertical-

ly dipping. Yellow arrows indicate location of figures B, C, D. In all photos strati-

graphic top is to the left. B) Shallow, erosionally based channel with only subtle

relief along its base (black dashed line). From the margin toward the axis (direc-

tion indicated by red arrow), strata show a rapid increase in sandstone:mud-

stone ratio as thin-bedded turbidites pass laterally into amalgamated sandstone

in the channel axis. C) Deep channel with prominently scoured base (solid black

line), overlain by a heterogeneous bypass facies. D) Sheetlike depositional

lobes of the order of 15 m thick that form laterally extensive bodies composed

of amalgamated sandstone (hammer (circled) for scale). Base and top of lobe

indicated by double-headed arrow.

of California, or where submarine

canyons have incised into the

coastal rivers, for example in the

modern Congo Delta, sediment,

especially sand, is almost continu-

ously supplied to deep water irre-

spective of the position of relative

sea level. Notwithstanding this

important variant, much of the follow-

ing discussion will follow the lead of

the early models and consider a

point-sourced sediment system

where characteristics of the sedi-

ment supply, namely the sediment

flux, sediment caliber and sand:mud

ratio, change in a systematic manner

over a relative sea-level cycle (Fig.

26).

Falling-Stage and Early Lowstand

Systems Tracts

During falling relative sea level (RSL)

as the influence of the river-supplied

sediment begins to be felt at the shelf

edge, the deep-marine system

becomes reactivated with thick, mud-

rich (but also sand transporting),

high-density flows (i.e., high trans-

port-efficiency flows). Accordingly,

the character of the sediment-gravity

flows change and causes them to be

out of grade with the existing surface

gradient, which, in this case, triggers

erosion, with the depth of incision

decreasing downflow. On the upper

and mid slope, submarine canyons

and large erosional channels

become excavated and the eroded

sediment is moved downflow. At the

base of slope, the depth of incision is

minor, but nevertheless this area too

is generally an area of bypass. In

contrast, the basin floor becomes an

active depocenter, the deposits con-

sisting mostly of fine sediment erod-

ed from the slope. Any coarse sed-

iment is sourced from slope erosion,

but can also be derived from relict

and palimpsest sediment on the shelf

and, as time passes, increasingly

from the advancing shelf deltas.

The falling stage is characterized

also by widespread gravitational

instability along the upper slope and

outer shelf, caused by factors such

as hydro-eustatic uplift (due to the

reduced thickness of the overlying

water column), stratal overpressures,

seafloor fluid seepage, gas-hydrate

decomposition and sublimation,

oversteepened slope gradient (as


DRAFT FORMAT

Figure 25. A) Characteristics of an idealized coutourite deposit (after Stow etal., 2002). B) Small-scale (core and outcrop) characteristics of contourite

deposits (Ito, 1996). Temporal variations in the speed and/or direction of the

contour current cause many of the contourite structures (e.g., mud flasers, reac-

tivation surfaces) to resemble features that occur in tidal deposits (see Chapter

9), although complete reversals are generally absent in contourite deposits.

Figure 26. Idealized model relating characteristics of sediment transport and

deposition to positions of relative sea level (RSL) on a passive continental mar-

gin with a single point-source sediment supply (modified after Posamentier and

Walker, 2006). With falling RSL widespread gravitational instability on the outer

shelf/upper slope results in extensive mass wasting and incision of canyons and

slope channels, and bypass of sediment to the base of slope and basin floor. As

the shoreline approaches the shelf-slope break and flows become more

enriched in sand, a period of deposition occurs in the slope channels to bring

them into grade, whereafter flows bypass the slope and feed basin-floor depo-

sitional lobes. Later, the sand:mud ratio begins to decrease, which enhances

levee growth and the development of leveed channel systems. Eventually as

sediment flux decreases and the deep-marine part of the system becomes pro-

gressively more starved of sediment, ongoing slope instability may cap the suc-

cession with mass-transport deposits. Eventually a condensed interval forms at

the next highstand.

the system progrades into rapidly

deepening water near the relict shelf

edge), and erosion along margins of

canyons and channels. The conse-

quence is extensive erosive mass

wasting in the form of slumps, slides

and cohesive debris flows, which

singly, or in combination with turbidity

current erosion, excavate major

canyon systems. The eroded sedi-

ment is moved downslope producing

mass-transport deposits (MTDs) on

the slope and basin floor. Observa-

tions from the modern indicate that

MTDs occur on a range of scales, and

include those of enormous scale such

as the late Pleistocene MTDs on the

Amazon Fan, which are of the order

of 50–100 m thick and cover an area

of 104 km2 (Piper et al., 1997). Howev-

er it is important to note that MTDs,

which are common during the falling-

stage systems tract, can occur at any

position in a depositional sequence.

Lowstand to Late Lowstand

Systems Tracts

Shelf deltas have now migrated far

out onto the shelf, and in some cases

may form shelf-edge deltas perched

at the top of the slope. As a conse-

quence, sediment supply into the

basin changes significantly: sediment

flux most probably increases, but

more importantly the sand:mud ratio

increases causing the transport effi-

ciency of flows to be reduced. In the

submarine canyons flows still gener-

ally bypass and canyon-wall collapse

continues. Nevertheless, coarse-

grained canyon fills indicate out-of

grade systems, which upon re-estab-

lishing equilibrium return to being a

sediment transfer conduit. On the

middle and lower part of the slope,

channels are mostly low to moderate

sinuosity and poorly confined. They

are initially at grade and neither erode

nor aggrade. With time and increas-

ing sand:mud ratio, channel gradients

become less than that of the (equilib-

rium) graded profile promoting exten-

sive in-channel deposition and aggra-

dation. The net result of these condi-

tions is the formation of a sheetlike

sandstone body consisting of amalga-

mated channel sandstones that

upward becomes increasingly more

interstratified with fine-grained strata

(i.e., channel fills become separated

by out-of-channel deposits). At the

same time, but downflow in the depo-

sitional lobe, sediment supply has

been significantly reduced because of

deposition upslope channels, and

because the increased sand:mud

ratio causes this part of the system to

retrograde (i.e., backstep upslope).

As the lowstand progresses, slope

channel-levee systems eventually

reach an equilibrium profile, and as a

consequence, sediment flux into

more distal parts of the turbidite sys-

tem increases. Accordingly, deposi-

tional lobe complexes become the

loci of sediment deposition and pro-

grade basinward, although, because

of extensive lateral compensational

stacking, the rate of progradation may

be low (see, e.g.,, Saller et al., 2008).

Transgressive Systems Tract

During the early transgressive sys-

tems tract, sediment supply into the

deep part of the basin begins to

diminish as sediment becomes

retained in the now aggrading coastal

plains and deltas (Posamentier and

Walker, 2006). Moreover, the

sand:mud ratio decreases. Within the

canyon instability along canyon walls,

in addition to parts of the shelf edge

and slope that are re-equilibrating to

the newly developing basin conditions

(Posamentier and Kolla, 2003),

results in the common emplacement

of MTDs. Flows that bypass the

canyon, however, form highly sinu-

ous, confined channel systems in the

middle and lower parts of the slope.

Compared to the lowstand, flows now

are highly stratified with a coarse

basal layer supporting a thick fine-

grained upper part. In addition, chan-

nel sinuosity is typically higher, which,

in turn, favors the development of

channels having a lower width:depth

ratio and greater channel relief

(Pirmez et al., 2000). Collectively

these conditions result in the forma-

tion of well-developed fine-grained

levees bordering erosively based, lat-

erally-migrating confined channels.

Stratigraphically upward, these chan-

nels become significantly more

aggradational, suggesting that rates

of deposition have increased both in

the channel and on the levees. Far-

ther downflow, sediment supply has

diminshed, and so too the sand:mud

ratio. Enhanced levee development

causes the leveed-channel to deposi-

tional lobe transition to prograde bas-

inward, but, because of comparative-

ly low sediment flux, at a much

reduced rate (e.g., Saller et al., 2008).

Eventually sediment flux to the basin

floor becomes sufficiently reduced

that depositional lobe complexes

begin to retrograde.

Highstand Systems Tract

Rising relative sea level eventually

floods the ever-widening continental

shelf. As a consequence, sediment

supply (especially of sand and gravel)

into the deep part of the basin is effec-

tively shut off as continental sediment

becomes trapped in up-dip fluvial and

coastal estuarine systems. Farther

basinward, sedimentation rates are

low and dominated by deposition of

fine-grained siliciclastic and biogenic

material from suspension, forming a

condensed horizon rich in pelagic

sediment. Where shelf conditions are

appropriate (water depth, salinity,

temperature, nutrients, etc.), howev-

er, carbonate production may be initi-

ated, interrupting for the first time the

succession of purely siliciclastic

deposits, but then only to become

deactivated during the subsequent

fall of relative sea level.

SUMMARY AND CONCLUSIONS

The inaccessibility and immense

scale of modern deep-marine sedi-

mentary environments makes it espe-

cially challenging for sedimentological

research. Our current state of knowl-

edge has benefited significantly from

recent improvements made in seismic

acquisition and resolution. These data

suggest that much of the deep-marine

sedimentary record can be subdivid-

ed into 2 end-member categories: an

expansive unit consisting of simple,

concordant reflectors indicating most-

ly fine-grained suspension fallout

deposition that blankets the seafloor,

and a significantly more complex unit,

both locally and regionally, which

locally interrupts the fine-grained

blanket. It consists of sand and, less

commonly, gravel-size sediment that

was transported and deposited by a

variety of sediment gravity flow and

mass-movement processes. It is

these latter kinds of deposits that are

most commonly described in the sed-

imentological literature, and, based

mostly on seismic data, include:

22 ARNOTT

DRAFT FORMAT

channels, levees, overbank/crevasse

splays, and depositional lobes.

Channels are diverse in kind, size

and origin. Submarine canyons and

other kinds of erosional channels

owe their existence to erosion by

moving currents, mass wasting, or a

combination of both. Fills of these

channels are similarly diverse and

range between locally derived, typi-

cally fine-grained sediment, to sand-

stone and conglomerate composed

of particles derived from an upslope

sediment source. Ideally, these ero-

sive channels pass downflow, into

leveed channels and then farther

basinward and with an associated

loss of flow confinement, into deposi-

tional lobes. Depending on the

degree of lateral confinement, leveed

channels can be subdivided into two

kinds: poorly and highly confined

channels. Poorly confined channels

tend to have lower sinuosity, be larg-

er in width, and form much thicker

channel deposits. Along the outer

margin of channel bends, channel-fill

strata terminate abruptly. Along the

opposite side, or inner bend, chan-

nel-fill strata show two kinds of termi-

nation. In poorly confined channels,

coarse channel strata grade laterally

into fine, thin-bedded turbidites of the

inner-bend levee, suggesting signifi-

cant flow overspill along that margin.

In contrast, (coarse) channel strata

of highly confined channels onlap

abruptly and become interfingered

with fine, thin-bedded inner-bend

levee strata, suggesting containment

of at least the basal, coarse-grained

part of the flow below the top of the

inner-bend levee. In addition, chan-

nel strata on the inner side of bends

show negligible upward or lateral

change in grain size, and lateral-

accretion deposits are well devel-

oped

Leveed channels are bound on

their sides by levees, which are best

developed in channel bends. Outer-

bend levees of poorly confined chan-

nels show a systematic but rapid lat-

eral thinning and fining of moderate-

ly thick beds away from the channel

margin. This trend suggests a genet-

ic relationship between channel and

levee, and relates to lateral loss of

flow competence and/or capacity as

flows, or the upper part of flows,

escape the channel and travel

across the adjacent levee. Intercalat-

ed with the medium beds are thinner,

finer beds that show little lateral

change, and after a few hundred

meters away from the channel mar-

gin become indistinguishable from

thicker beds that had previously

thinned and fined laterally. In highly

confined channels, outer-bend lev-

ees consist typically of fine, thin-bed-

ded turbidites that change little later-

ally. In this case, the channel margin

represents an erosion surface that

has incised vertically and laterally

into levee strata deposited by an

older highly confined channel. Asso-

ciated also with outer bends, are cre-

vasse splay and overbank splay

deposits. These features, which

appear to be best developed in poor-

ly confined channels, are the result of

flow overspill, possibly aided by iner-

tial run-up. Overbank splays tend to

be smaller and thinner, and common-

ly consist of a single or small number

of often complete Bouma turbidites.

Crevasse splays are larger, longer-

lived features that comprise an array

of lithofacies, including strata indicat-

ing rapid flow expansion and high

rates of deposition (structureless,

matrix-rich sandstones).

Depositional lobes are major sedi-

ment bodies that occur downflow of

leveed channels. These units com-

prise three subunits: deep channels,

shallow channels, and sheetlike

splay deposits. Deep channels,

which are the principal sediment-

input conduit for the lobe, are deeply

incised and filled, at least in part, with

heterogeneous strata related to

incomplete bypass, and in some

cases also by thick amalgamated

sandstone. Shallow channels form a

distributive network immediately

downflow of the deep channel.

Although channel bases show

meters of erosion, it is distributed

over hundreds of meters laterally.

Typically, these channels are filled

with sand-rich strata that show a sys-

tematic axis-to-margin fining and

thinning trend. Unconfined flow at the

downflow terminus of the shallow

channels builds up the sheetlike

splay unit that consists almost exclu-

sively of amalgamated sandstone.

Eventually these strata will pass

gradually into fine-grained (suspen-

sion) deposits of the seafloor (possi-

bly interrupted locally by contourite

deposits), however at present the

details of that change are poorly

understood.

Like in most other sedimentary

systems, sedimentation in the deep-

marine environment is influenced

significantly by changes of relative

sea level, especially where the conti-

nental shelf is wide and terrestrial

sediment is point sourced. During

stages of falling relative sea level,

widespread gravitational instability

causes extensive mass wasting on

the slope and the formation of exten-

sive mass-transport deposits on the

lower slope and basin floor. During

the lowstand, slope canyons and

erosional channels pass downflow

into leveed channels, depositiional

lobes and eventually suspension

deposits on the distal basin floor.

Spatial and temporal characteristics

of deposition along this transport/

deposition pathway depend on the

complex interaction of myriad sedi-

mentological variables, including

sediment supply, sediment caliber,

sediment mineralogy, channel gradi-

ent, flow density and stratification, in

addition to possible high-frequency

changes of relative sea level. Never-

theless, as relative sea level rises,

sediment supply becomes progres-

sively reduced as river-borne sedi-

ment becomes increasingly

sequestered in shelf and coastal-

plain environments. Eventually a

condensed horizon consisting of

mostly suspension deposits is

formed, and marks the minimum

level of sediment supply into the

deeper parts of the basin.

In the last version of this book,

Walker (1992) noted that “Submarine

fan studies are presently in a state of

flux. The seismically derived infor-

mation from modern fans ... cannot

be applied very easily in the geologi-

cal record where the scale of obser-

vation is commonly very much small-

er. Conversely, features in the geo-

logical record cannot be related to

parts of modern fans”. Over the past

several years, however, the minimum

resolution of seismic images has

improved enormously and the num-

ber of seismic-scale outcrops exam-

ined has increased significantly.

Together these developments have

narrowed the gap considerably, but


DRAFT FORMAT

still more needs to be done. Of partic-

ular note is the need to ground-truth

the origin of a number of seismic-

reflection patterns. Do they faithfully

mimic the character of the geology, or

might some part of the pattern be an

artifact of geology and wave interfer-

ence or diffraction characteristics, for

example, in places were bed thick-

ness and facies change rapidly? Such

details might have dramatic impact on

how our models predict stratal archi-

tecture and connectivity. Future areas

of research must also stress the fine-

grained part of the deep-marine sedi-

mentary record. Although typically

poorly exposed, it is these strata that

make up most of the deep-marine

sedimentary record. Also, because of

their abundance and ubiquity, these

strata probably add significantly to

hydrocarbon reservoirs. Lastly, and

maybe most profoundly, a concerted

effort is needed to understand better

the mechanisms that transport and

deposit sediment in the deep marine

(e.g., McCaffrey et al., 2001). For

example, the effect of sediment con-

centration on turbulence structure

and intensity in turbidity currents, or

how different levels of density stratifi-

cation in the flow manifest itself in the

geological record are not understood

well. These and many other important

topics are now ripe for investigation

as the seemingly endless advances in

analytical instrumentation make what

was formerly impossible, possible.

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N.E., eds., 1985, Submarine Fans and

Related Turbidite Systems. Frontiers in

Sedimentary Geology: New York,

Springer-Verlag, 351 p.

This volume provides an excellent sum-mary of the seismic attributes of manymodern submarine-fan systems. Asmaller number of ancient examples arealso presented.

Kuenen, P.H. and Migliorini, C.I., 1950, Tur-

bidity currents as a cause of graded

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Seminal paper that established the mod-ern turbidite concept.

McCaffrey, W.D., Kneller, B.C. and Peakall,

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302 p.

A mostly theoretical and experimentalapproach to the study of particulategravity currents. Papers provide current

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Mulder, T. and Alexander, J., 2001, The

physical character of subaqueous sedi-

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This paper provides a thorough reviewof sediment gravity flows and theirdeposits, and proposes a classificationscheme based on sediment concentra-tion.

Mutti, E., Steffens, G.S., Pirmez, C. and

Orlando, M., eds., 2003, Turbidite: mod-

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A collection of 24 papers discussingmodern and ancient slope-to-basin-floorsedimentary systems as well as trans-port mechanisms and depositionalprocesses.

Nilsen, T., Shew, R., Steffens, G. and

Studlick, J., eds., 2007, Atlas of Deep-

Water Outcrops: American Association

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Geology 56, 504 p. + CD-ROM papers.

An extensive, well-illustrated compila-tion of case examples of deep-marinesedimentary rocks from around theworld, including important statisticaldata.

Posamentier, H.W. and Walker, R.G., 2006,

Deep Water Turbidites and Submarine

Fans, in Posamentier, H.W. and Walker,

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Recent review of deep-marine sedimen-tation with a comprehensive compilationof seismic images that illustrate a widearray of depositional elements.

Piper, D. J. W. and W. R. Normark, 2001,

Sandy fans; from Amazon to Hueneme

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26 ARNOTT

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