Middleton G.V. (Ed.) Encyclopedia of Sediments and Sedimentary Rocks

Подождите немного. Документ загружается.

SILICEOUS SEDIMENTS

665

shallow waters oxygenated by agitation or by cyanobacteria. It

then accumulated cither in pulses of deposition or else steadily

atnidst silica particles that were accumulatitig in such pulses

because the deposits are beautifully banded with alternating

layers of iron oxide (or iroti carbonate) and chert. The tiature

of the original iron atid silica precipitates is unktiowti. but both

were likely hydrous, amorphous phases which underwent early

diagenetic tratisformation to crystalline phases. The detnise of

iron formation deposition has been widely attributed to

gradual oxygenation of the ocean, but this does not explain

the disappearance of the associated silica deposition. One

possibility is that climatic temperatures also declined with titne

which greatly reduced the atiioutit of dissolved silica trans-

ported in the oceans. Numerous other types of silica-replaced

seditnetits t)ccur in the Archeati. The silicilication is often so

extensive that the nature of the replaced sediments can be

inferred only by study of petrographic fabrics and preserved

sedimentarv structures.

Evolution of siliceous sediments with time

For at leasl the past 540 million ycats, siliceous seditnents have

fortned pritnarily by the deposition of biologically secreted

opal particles that dissolve and reprecipitate as quartz beds

and nodules during the compaction history ofthe sediments.

Prior to this, there is tio current fossil evidence that the

sponges, radiolaria. and diatotns were present to supply

biogenic silica to sediments. Cherts and silicified seditnents

are nevertheless widespread in Precatnbrian strata. In the

Archean, it is possible that tnuch higher ocean tetnpcratures

allowed greatly enhanced transport of dissolved silica and that

large areas of the ocean floor were blanketed with opaline floes

derived from intense weathering and from deep-ocean hydro-

thennal vents. It is also possible that bacterial precipitation of

opal was widespread and produeed equivalctit amoutns of

silica as that produced by the Phatierozoic organistns.

However introduced into seditnents, silica diagenesis in the

.Archeati tintst have been particulary vigorous to produce such

extensive occurrences of bedded cherts and siliciTted seditnents.

In the late Archean and early Protcrozoic. ititerbedded

cherts and iron oxides became common. Inasmuch as the iron

is thought to have been exhaled frotn submarine hydrothermal

vents,

it is likely that tnuch ofthe siliea was also. However, the

role of organistns and the mechanisms that produced rhymthic

deposition of iron oxide and silica are stilt unknown. Earge-

scale siliceous scdimentatioti declined during the Proterozoic.

but replacement cherts retnaitied cotntnon in strotnatolitic and

other latninated carbonates throughout the Precatnbrian.

Carbonates and associated evaporites were appatxntly re-

placed with tnicrocrystalline quartz during early diagenetic

stabilization via the same processes that produced Phanerozoic

nodular cherts, but the silica source retnaitis problematic. The

amoimt of silica introduced into the Precambrian carbonates

appears no differetit from the younger cases.

The evolution of sponges atid radiolarians in the earliest

Paleozoic was coincident with the evolution of prodigeous

nutnbers of earbotiate-secreting organisms. Eimestones and

dolostones throughout the Paleozoic and Mesozoic are filled

with nodular and disseminated replacement chert that resulted

when opaline sponge spicules dissolved and reprecipitated

during early diagctietie stabilization. Thick radiolarian oozes

commonly accutnulated in the deep sea and these were

transformed into great bedded chert sequences now displayed

iti deformed and uplifted aecretionary wedges.

Silica secreting diatotns evolved within the last 100 million

years and completely transformed the nature of siliceoLts

seditnents. These prolific photosynthesizers now dotnitiate

silica geochemistry in the ocean, lakes, rivers, streams, and

soils.

Silica is so efficiently extracted frotn water by these

organisms that dissolved siliea concentrations in surface

seawater and virtually all fresh water environments are kept

well below saturation for quartz and all forms of opal. Silica

mobilization and re-precipitation in pore fluids has thus been

arrested iti sedimetits exeept for deep-sea diatomaeeoits oozes.

Any biologic opal ititroduced into shallow tnarinc carbonate

sediments sitnply dissolves back into the sea; nodular

rcplacetnetit cherts have essentially stopped formitig. Siliceous

sedimentation is now mainly a case of diatotii accumulation.

The geologic history of siliceous sediments has thus undergone

significant changes that provide another example of the

profound influence of biologic evolution on the evolution of

seditnentary rocks.

E. Paul Knauth

Bibliography

Belli. R..I,. and Giirrison, R.E., 1994. The origin of chert in the

Monterey I-ormation of Ciilifornia (USA). In lijima. A., Abed,

A.M., and Garrison. R.E. (cds,).

Siticeoti,s.

PhosphatieandGIaitco-

nitic Sediments ofthe Tertiary and Mesozoic.

Proceeiiings

ofthe 29th

Internatiotial Geotot^ical Coniiress. Part C. pp. 101 ]}2.

Litrcclit:VSP.

Biggs. D.E., 1957- Petrography and origin of Illinois nodular cherts.

Illinois State

Geologieat

Sttrvey i

'ircutar.

245.

Bramlett. M.N.. 1946. The Monterey Formation of Cafifornia and the

origin of its siliecuus roeks. United Slates Geologieal Survey

Frolcssional Paper 212.

Ciirozzi. .A.V.. I99.V Sedimetitary Petrography. Prentiee Hall.

Chowns, T.M.. and Eilkins. J,F_.. 1974. The origin of quart/' geodes

and eaulillower clieits through silieilieatlon of anhydrite nodules.

Journalof Sedintctnarv Petrology. 44:

885-90.1.

Folk. R.E., and McBride. E.F.. 1976. The Caballos Novaculite

revisited P:irt I: Origin of novaetilile members. Joarnat of Sedimen-

tary

Petrology.

46: fi59 (S69.

Gruniin. H.R.. 1965. Radiolarian locks and associated roeks in .spaee

and tiniL'.

Eclogae

Geological Hetvetiae. 58: 157 208.

Hesse. R.. 1990, Origin of chert and silica diagenesis. In Mellreath.

I.A.. Morrow. D.W,

(eds.},

Diagenesis. Gfologieal Assoeialion of

t anada. Geoscicnec Canada Reprint Series 4. pp. 227 275,

Knauth. L.P.. 1979. A model for the oriuin of ehert in limestone.

(.cologv. 7: 274 277,

Knauth. L.P., 1992. Origin luid diagenesis of cherts: an isotopie

perspective. In

Isotopic

Signatitri'samlSedimentary Records. Eecturc

Note?,

in Earth Sciences, 43. New York: Springer-Verlag, pp. 123-

152,

Knautli, L.P,. 1994. Pctrogenesis of chert. Reviews of

Mineraloi^v.

29:

233 25K.

Knauth, E.P.. and Lowe. D.R.,

20(13.

High Archean climatie

tempcratnre inferred from oxygen isotype geochemistry of eheru

in the 3.5Ga Swaziland Supergroup, South Africa..

Geological

So-

ciety of .Atnerica Bnlletin. 115 no. 2 (in press).

Maekenzie, F.T.. and Gees, R., 1971. Quartz: synthesis at earth-

surface eonditions. Science. 173: 533 535.

Miser, H.D.,and Purdue. A.H,.

1929,

Geoiogy of the Dequeen and Caddo

Ciap Quadrangles. Arkansas.

'SGeoiogicalSitrvevBulletin. S08.

Murray. R.W., Jones. D.L., and Bueholt/ ten Brink. M.R.. 1992.

Diagenetie formation of bedded chert: evidence trom chemistry of

Ihe chert-shale couplet. Geoh-gy. 20: 271 274.

Sehubel. K.A., and Simonson, B.M.. 1990. Petrography and diagenesis

ol'eherls from Lake Mayadi. Kenya, ./ournalof .Sedimentarv

Petrol-

ogy. 60: 761 776.

6b6

101-

AND SLDMI' STRUCTUKIiS

Sheppard, R.A,. and Gudc, A.J,.

I^Sf*.

Magadi-type cherl—a

distinctive diiigenetic variety from lai-iisirinc dcposils. Uniled

States

Geologicai

Survey Bulletin. 1578: .VV^ 145.

Tiirr, W.A.. 1917. Origin of the eliert in ihc Btirlington Limestone.

Ameriedii Journal of Seieme., 44: 4()'-) 452.

Van Tuyl, F.M., 1918. The origin ofchen. AnwrieaiiJourmdoJSeienee.

45:

449 456,

Cross-references

Diagenesis

Diagenetic Struetures

Geodes

Ironstones and Iron Kormalions

Isolopic Mclliods in Sedimentology

Magadiiic

Neomorpliism and Recryslalli.^iitton

Silcrete

SLIDE AND SLUMP STRUCTURES

Slide and slump structures are common geological phenomena

on subnerial and subaqtieotis slopes (Maltman. 1994). They

range in size and volume from a few tens of sentimenters to

several thousand cubic kilometers (Jansen ci ul., 1987:

Martinsen. 1994). and encompass gliding unconsolidated

sediments as well as lilhitied blocks (olisottiths). Slides and

slumps are significant geological processes which have an

important impact on human life, known from recent catastro-

phies (Maltman. 1994. and references therein). In addition,

sliding and slumping can seriously affect the stability of

offshore installations (Prior and Coleman. 1982). and generate

tsunamis {Driscoll

(•/(//.,

2000). Products of slides and slumps

provide important information on the depositionai setting of a

sedimentary rock unit.

Basic characteristics and theory

Slumps and slides arc downslope movements of sediments

above a basal shear surface where there is, respectively,

significant and insignificant internal distortion of ihe bedding

(Stow, 1986). In slumps. Ihc bedding should be recognizable,

otherwise they classify as debris flows. There is a continuous

transition between slides, slumps and debris flows, and some

units may show characteristics of all three modes of transport

depending on the internal deformation state (Martinsen. 1994).

Slumping and sliding are common especially where there are

significant fine-grained seditnents. The structures form above a

basal shear surface (Figures S29 and S30). the depth to which

is decided mainly by the pressure gradient in the sediment.

Where the pore pressure approaches or balances the normal

stress induced by the weight of the overburden, the shear

strength is sufficiently reduced to allow slippage along the

basal shear surfaee. given a sufficiently high shear stress. These

relationships are given by the equations:

Shear strength (i)=C+(CT-/»)tan^ (Eq. 1; Hampton, 1979J

and

Shear stress {S) = pg.v Ftana (Rq. 2: Middleton and

Southard,1978)

where C is sediment cohesion, o is normal stress (or weight

of overburden). /) is pore pressure, tp is the angle of internal

friction, p is sediment density. ,sr is acceleration due to gravity.

,s is solidity (or the completiientary of porosity). Y is sediment

thickness and

is slope angle.

Tributaries

Deformed early concretion

SLUMPS

Figure S29 Idealized slump model. Modified from Martinsen (1994).

SLIDE AND SLUMP STKUCTUKFS

667

HEAD

Extensional fault family

TOE

'Pressure ridges'

SLIDES

Leading-^dge

folds/

Figure S;JO Idf.ilized slidf modci. Modified from Martinsen (1994).

Once initiated, the slide or slump basal shear surface will

propagate iipslope in a radial fashion from its nucieation point

(Williams and Chapman. 1983), leading to the tbrmation of a

scoop-shapcd. concave-downslope depression or scar, often

with an irregular outline with tributaries (Martinsen, 1994),

The shear surface is probably initiated as a slope-parallel

feature, but al some point steepens to intersect the sediment

surface. The transition may occur at lilhofacies boundaries, or

at sites of pore pressure jumps where material strength

contrasts are present (Crans ctal.. 1980),

Slumps and slides are triggered by a variety of processes.

Seismic triggering, cyclic wave-loading, high sedimentation

rates and methane generation causing overpressuring. slope

oversleepening. and tectonic movements are al! possible

initiators (Martinsen, 1994).

Slumps and slides form on low-gradient slopes, as little as

O-l (Prior and Coleman. I97S). They range in thickness from

0.5m (Martinsen. 1994). to several hundreds of meters on

eontinental slopes (Janscn cltiL, 1987).

Slump structures

M(ning slumps deform intensely internally and produce a wide

variety of Jeformatioiial structures. Folds, boudins. micro-

faults,

internal shear surfaces and faults are common

structures in slumps (Figure S29), These structures suggest

that the slumps go through a main phase of plastic/ductile

deformation whereby folds and boudins are formed. The

ductile phase is followed by a late brittle phase where the faults

form. It is common to see strain overprinting where early-

Ibrmcd folds are truncated by late faults (Figure S.^l).

Figure S31 Slump unit Irom the Upper C^irboniierous Gull IsNind

Form.ilion,

County CLire, Ireland, Note ihe inlerniil distorted bedding

ol Ihc unil in the ne.ir clit'f, which is approximalely ISm

high.

The slump folds arc mainly shealh folds formed by simple

shear (F""igure S31; Martinson. 1994), although buckle folds

also occur (Woodcock. 1976), Faults can be extensional and

contractional (Figure S29). related to local obstacles or shear

surface irregularities. Models of slumps and slides (e,g.. Allen.

1985) show a well-defined upper extensional zone and a

downslope contractional zone. Flowevcr. it is likely that a lal-

cnil compaction will oecur when fine-grained slumps come to a

halt, preventing development o\' toe contractional zones,

leaving opcii-fnch-dloiii-s (Garfunkel. 1984),

Slump scars often show evidence of retrogressive failure,

from footwall unloading (Martinsen. 1994), The slump sears

are generally filled by iinc-grained sediment, or expanded by

668

SLOPE SEDIMENTS

current erosion and filled by coiirsci'-grained turbidity current

deposits, ln this case, it may be impossible to differentiate the

slump scar from an erosional channel.

Slumps form hummocky sediment masses in plan view,

particularly at their downslope ends if toe-zones are developed.

The topography may be organized into lobate, convex-

downslope ridges (Prior el al.. 19S4) which are the surface

manifestation of toe-zone thrust faults (Martinsen and

Bakken, 1990).

Slide structures

Slides include bank collapse in river channels, subaerial

mudslides, sediments displaced by delta-front growth faults,

shclf-cdgc faults and submarine glide-blocks and olistoliths.

Slides are either rotational or translalloiial (Allen. 1985),

depending on the thickness to movemeni distance ratio.

In their most simple tbrm. slides are spoon-shaped features

with a three-partite morphology of an extensional head region,

a middle, translational, "rigid" zone and downslope, contrae-

tional toe zone (Figure S30; Brunsdcn, 1984) . In rigid zones.

there may internal slip between beds in the form of sheath

folds or microfaults (Figure S.^1). The downslope toe region

can be "open-ended'" with no contractional structures devel-

oped because of lateral compaction {Garfunkel, 1984: see also

above).

The head region usually develops a •'family"' of listric faults

(Crans et of. I9S0) which soles out at a basal decollement.

Antithetic extensional faults are common. Fault families in

slides is an essential difference from slumps, where most faults

are single faults. The margins of the slides are dominated by

strike-slip deformation, but the width variability of the slide

scar can cause both iranspressional and transtensional move-

ment if the slide scar narrows or widens.

Contractional det\irmatton is most commonly expressed as

thrust iaults. The thrusts tbrm classic duplex and imbricate

zone geometries (Figure S31; Dingle, 1977). Three ditTerent

types of eontractional zones occur in slides (Martinsen and

Bakken. 1990): (I) The thick package zones; similar to

mountain bell thrust zones, but several orders of magnitude

smaller: (2) Thin package zones developed at the frontal

margin of heterogeneous slides; (3) Basal contractional zones

related to obstacles such as ramps.

Toe zones of slides may be confused with contractional

zones formed by gravityspreuding. Gravity spreading forms in

progradational settings (e.g., deltas) due to progressive

sediment loading (Schaek Pedersen. 1987).

Ole J. Martinsen

Bibliography

AIIL'11,

J.R.L.. 19S5. Principles of Physical Sedimentology. London:

George Allen and Unwin.

Brunsden, D., 1984. Mudslides. In Brunsdcii, D.. and Prior. D.B.

(eds.).

SlopeInstithilitv. Cliichester: .lohn Wiley and Sons. pp. .163

418.

Crans.

W.. Maiidl, G.. and Harcmboure, J.. I9S0. On ihe theory of

growth-faulting a gt;omechanical delta model based on gravity

sliding. Join nal of Petroleum Geology. 2: 265-307.

Dingle. R.V.. 1977. The anatomy of a large submarine slump on a

sheared continental margin (SI: Africa). Jtntrnal of the Geological

Society Lottdim. IM:

29?.

.31(1.

Driscoll. N.W . Wcissci. J.K.. and

Goff.

J.A..

200(1.

Poicntial for large-

scale subniiirine slope lailuic and tsunami generation along the US

mid-AllantIc coasl. Geology. 28: 407 4IO.~

Garfunkel. Z.. 1984. Largc-seale submarine rotational slumps and

yrowlh faults in Ihe eastern mediterranean. Marine Geologv. 55:

.M15-324.

Hampton, M.A.. 1979. Buoyancy in debris flows,

.Iota

nal of Seditnen-

tary Petrology. 49: 753-7*58.

Jansen, E.. BetVing. S., Bugge. T,. Eidvin. T.. Holtedalil. H.. and

Sejrup, H.P.. 1987. Large submarine slides on the Norwegian

conlinental margin: sediments, transport and liming. Marine Geol-

ogy. 78: 77-107.

Maltman. A.. 1994. The

Geological Defm-mation

of Seditnents. London:

Chapman and Hall.

Martinsen. O.J.. 1994. Mass movements. In Maltman, A. (ed.). The

Geological

Deformation of Sedintents. London: Chapman and Hall,

pp.

127 165.'

Martinsen. O.J.. and Bakken, B.. 1990. E.xtensiona! and compressional

zones in slumps and slides in the Namurian of County Clare,

Ireland. Journal of

the

Geological

Society.

Lotidon. 147: 153 164.

Middleton. G.V,, and Southard. J.B., 1978. Mechanics uf Sediment

Xfoyement. Tutsa: SEPM Short

ourse, ?.

Prior. D.B.. Bornhoid, B.D.. and Johns. M.W.. 1984. Depositional

characteristics of a submarine debiis flow. Journalof Geologv. 92:

707-727.

Prior, D.B., and Coleman. J.M., I97S. Disintegrating retrogressive

landslides on very low-angle subaqueous slopes, Mississippi delta.

Marine

Gt'otcchnology.

3: 37 60.

Prior. D.B,. and Coleman. J.M., 1982. Active slides and tlows in

nndLTLonsolidated marine sediments on the slopes of the

Mississippi delta. In Sa.xov, S., and Nieuwenhiiis. J.K. (eds.).

,1/;;/-

itie Slides and Other Mass Movements. NATO Workshop. Plenum

Press,

pp. 21-49.

Schaek Pedersen. S.A.. 1987. Comparative studies of gravity tectonics

in quaternary sediments and sedimentary roeks related to fold

belts.

In Jones. M.E.. and Preston. R.M.I", (eds.). Defornuttion of

Sediment.s and Scdin]entarv Roeks. Geological Soeiety of London.

Special Publication. Volume 29. pp. 165 'l80.

Stow. D.A.V., 1986. Deep elastic seas, in Reading. H.G. (ed,). Sedi-

tnentary Enyirontnentsand Facies, 2nd edn. Blaekwell, pp. 399-444.

Williams. G.D.. and Chapman. P.. 1983. Strains developed in the

hangmgwalls of thrusts due to their slip/propagation rate: a

dislocation model. Joitrnal of Structural

Geologv.

5: 563 572.

Woodcock, N.H.. 1976. Structural style in slump sheets: Liidlow series,

Powys, Waies. Jotirnat of the

Geological

Societv, London, 132: 399-

415.

Cross-references

Bedding and Internal Structures

Debris Elou

Gravity-Driven Mass Elows

Rivers and Alluvial Fans

Slope Sediments

Storm Deposits

Submarine Eans and Channels

Synsedimentary Structures and Growth Eaults

Tectonic C~onirois of Sedimentation

Tsunami Deposits

SLOPE SEDIMENTS

The slope is a relatively steep transition zone between a low-

gradient tnarine shelf (<l gradient) and the deep seafloor.

Mostly, slopes occupy the hatfiyal zone of the ocean basins

(200 2000m). Benthic macrofauna are rare, including thin-

shelved bivalves, brachiopods, and eehjnoderms. Benthic

tbraminifera, where present, are useful in providing water-

depth esiitnates along ancient slopes. Many slopes are located

SLOPE .SEDIMENTS

669

the

edges

continents; others fringe volcanic islands

intracratonic basins

continental crust. Around

the

passive

marjzins

ol" the

Atlantic Ocean,

the

(.•onlincntal slope extends

from ihe,s7/(V/ slopehreak

{

--I20m average water depth)

to the

top ot"

the

coiuinciuiihisc (--2.400 m average depth: Emery

and

LJchupi. 1984). Along subdiiction margins,

the

base

of the

slope

much deeper

at the

landward edge

of the

trench.

The

depth

of the

shelf .shipchrcak varies widely between

20 500 m

(Fairbridge

and

Bourgeois. 1978).

and may be

within

the

photic zone along carbonate platform margins where reefs

grow into

the

surf 7one immediately landward

of a

precipitous

upper-slope escarpment.

Basin slopes

are

distinguished from adjacent shelves

by a

steeper gradient

and by a

lack

o\'

signilicant wave

and

tidal-

current action. Instead, scd'uiieni'gravity fUiws predominate,

sometimes associated with ilwimohalineiui-retus which sweep

the distal parts

some slopes. Gradients range from

3 6 or

less

mud-dominated passive-margin slopes (Heezen. 1959).

to 10-30"

erosional slopes flanking strike-slip basins like

those along

the

North Anatolian Transform Fault

north-

western Turkey (Aksu

el al..

2000),

local near-vertical

escarpments seaward

variably cemented carbonate barrier

reefs.

The

prodelta slopes

marine coarse-grained deltas

commonly include sands

and

gravels transported

gravity

currents

and

slope failures (Nemec, 1990).

Cementation, organic binding

and

reef development

carbonate platform margins

may

lead

to the

development

of a

very steep upper-slope segment (Coniglio

and Dix.

1992),

Carbonate slopes

ol"

this type

are

strongly concave

shape

because

the

steep upper slope segment gives

way in

deeper

water

to a

much gentler gradient. Some carbonate slopes

can

be steepened

erosion, which

may

cause

the

upper slope

retreat lens

kilometers (Jansa.

1981:

Mullins

el uL.

1986),

Sandy carbonate sediments form steeper slopes than carbonate

muds (30-40' versus

<5 ;

Kenter, 1990).

A fraction

of the

sediment that moves across

the

shelf-slope

break accumulates

on the

slope

itself,

leading

aggradation

and progradation.

partial filling

o\'

slope basins along

convergent continental margins.

greater portion, however.

bypasses

the

slope

deeper water areas like

the

continental

rise

and

abyssal plain (passive margins)

oceanic trenches

(convergent margins). Seismic data provide considerable

information

continental slopes. Various patterns

reflections, designated

seismic fades, have been recognized

(e,g,. Sangree

(•/(//..

1978), These include:

(1)

mounded ehaotic

facies.

the

result

slumps, slides

and

other mass movements:

(2) onlapping till facies.

due to

infilling

slope irregularities

by turbidity currents:

(3)

divergent

and

parallel-layered facies.

a rcsuk

deposition from low-energy turbidity currents

and

liemipelagic sedimentalion:

and (4)

sheet drape facies. resulting

from bUmketing

slope irregularities

heniipelagic sedi-

ments. Hemipelagic sediments accumulate sufficiently slowly

that they

are

usually thoroughly bioturbated

representa-

tives

of the

Zotiplivcos trace-fossil assemblage. Gravity-flow

deposits

and

sediment slides

are

eniplaced rapidly,

preserve

physical structures (e,g.. lamination formed

traction

transport: deformation structures t~ormed

shearing).

The gradients

and

morphologies

siliciclastie slopes

are

strongly controlled

tectonics, sediment supply,

and

salt

mud diapirism.

passive inargins. gradients

are low and the

seabed

may be

quite featureless. Exceptions

are

erosional

gullies

and

submarine canyons

in the

upper slope that

fiuinel shelf-derived sediment seaward,

and

slide scars

and

mound-shaped slide

and

slump deposits that characterize some

slopes.

high latitudes, where continental

ice

sheets recently

delivered sediment directly

to the

upper slope, gullies

are

replaced

widely separated, glacially

cut

transverse troughs,

and

the

slope tends

to be

smooth, with

without down-to-

the-hasin listric faults (Aksu

and

Hiscott.

1989.

1992). During

interglacial times, these glacially

fed

slopes

may

become

extensively dissected

headward erosion

gullies

and

canyons (Hesse

and

Klaucke. 1995),

The slopes along convergent plate boundaries lead from

narrow

shelf,

across

accretionary prism,

to the

floor

of a

deep-sea trench. Thrust f"aults

and

local

mud

diapirs

in the

accretionary prism produce

complex slope morphology

(Figure

S32)

with intraslope basins that

are

elongated parallel

the

plate boundary (Pickering

ei al.. 1995:

Moore

el a!.,

2001), Turbidity currents travelling down

the

slope

are

diverted into these intraslope basins where they

can

become

ponded.

On basin slopes seaward

major river deltas (e.g.. Gulf

Mexico. Niger Delta area), loading

water-rich muds

deeply buried salt

by a

thick delta

and

slope succession leads

complex diapiric intrusions, thrust faulting,

and

down-to-the-

basin growth faulting which

can

create

rugged scalloor

characterized

intraslope basins

and

hills (e.g,. Sigsbee

Knolls

on the

Texas -Louisiana slope; Sattertield

and

Behrens,

1990),

Sedimentary processes and typical deposits

ParticLilaie sediments are transported beyond the shelf-slope

break by three main mechanisms: (1) bottom-hugging sedi-

ment gravity flows; (2) thermohaline bottom currents: and

(3) surface wind-driven currents or river plumes that carry

suspended sediment of"f the

shelf",

lidal currents and surface

waves are only locally important as transport agents on the

upper parts of slopes and in the heads of some submarine

canyons (Shepard etal.. 1979). Sediment gravity flows that

develop on basin-margin slopes include turhidiiy currents, li-

quefied flows,

and

debris flows.

Where shelves are narrow (e.g.. offshore California), plumes

of suspended sediment from river deltas can extend beyond the

Figure S32 Multichannel seismic clip [jroiiie midway down the slope

oi the

N.inkai

accretionary prism near longitude t34,5E, south of

japan (water dt^pih ^3,000

m),

A turbidite-tilled slope hjsin has been

created alop the stac

keH

thrust sheets of the [irism. The white line

marks a boftoni-simulating reflector, altribuled lo the presence of f^as

hydrates. Seismic profile discussed in Moore efj/, I200J), Reproduced

by permission of C,F, Moore,

670

SLOPE SEDIMFNTS

shelf slope break (Emery and Miiliman, 1978: Thornton.

1981.

1984). directly contributing fine-grained sediments to the

slope. In polar areas, sediment-laden spring meltwater may

actually flow from the land across the surface of floating sea

ice.

to deposit its load direetly onto the continental slope

(Reimnit/and Bruder. 1972).

Mud that leaves the shelf either by high-level escape in river

plumes, by dilute turbidity-current i\ow. or by movement

along density interfaces in the water column over the slope

eventually settles to the seafloor to form hemipelagic deposits.

Deposition rates are on the order of 10-60cm/l000 years

(Krissek. 1984: Nelson and Stanley. 1984; Coniglio and Dix.

1992).

Hemipelagic sediments contain a pelagic component

(biogenic skeletons, volcanic and windblown dust, ice-rafted

detritus), generally 5 50 percent but locally as much as 75

percent, with the remainder consisting of terrigenous mud

supplied t"rom the iidjaeent land mass. The finest partieles are

probably carried to the bottom as aggregates in the form of

faecal pellets (Calvert. 1966; Schrader. 1971: Dunbar and

Berger. 1981), Unlike turbidites, which fill low areas on the

slope (Figure S32). hemipelagic deposits form even-thickness

drapes (Figure S33).

Dilute low-concentration turbidity currents tend to travel at

relatively low velocities, are potentially many tens of meters

thick, and may be turned by the Coriolis Ef"f"ect to flow oblique

or parallel with the bottom contours before reaching the base

of the slope (Hill. 1984). On carbonate slopes, the fine-grained

material exported from the platform is ref"erred to as periplai-

form sediment (Sehlager and James. 1978: Coniglio and Dix.

1992),

Suspended sediment concentrations in shelf areas may

be quite high due to input of mud-laden river water, or stirring

of the bottom by strong waves, tidal currents, or internal

waves at density intert^aees (Caeehione and Southard. 1974).

This suspended sediment may be adveeted ol"f the shelf by

ambient eurrents. possibly wind-driven, or by transport in

cascading cold water that tnay How off the shelf in the winter

months (McCave. 1972; McGrail and Cames. 1983; Wilson

and Roberts. 1995), Suspensions of fme-grained sediment may

also leave the shelf as dilute turbidity currents (lutite flows).

Isvee

Figure S33 Single-channel seismic strike profile across the middle

slope offshore Brunei, Btjrneo (Hiscott. in press). Waler depth is

-1,000

The slope gradient, toward the reader, is -2 , A number of

leveed channels have fjrown vertically during the Quaternary,

principally through fallout of suspended load from the Baram Delta to

form

even-thickness hemipelagic drape. The channel to the

SW has been most recently active—levees have enhanced sediment

thicknesses and the floor of the channel appears lo have

central

thalv^/eg channel flanked by low, possihiy sandy inner levees (cf. Piper

et al., 1999). The patterned lines mark reflections that can he widely

traced across the Brunei shelf and upper slope.

moving along the bottom onto the lower slope and rise, or

along density interfaces in the ocean water. These dilute

suspensions may move down a smooth upper slope as

unconflned sheet flows, or may be confmed by gullies or

canyons in which suspension is assisted by weak tidally-forced

flows (Shepard el al.. 1979: Gorsline ei al.. 1984), Most

mud transported off the shelf by dilute turbidity currents

(lutite flows) bypasses the slope, leading to maximum rates

of accumulation on the continental rise (Nelson and Stanley,

1984).

In carbonate settings, platform-derived sediment can

be traced >100km from the shelf edac (Heath and Mullins.

1984).

High-concentration turbidity currents carrying significant

sand load arc confined to gullies and larger canyons on the

slope, permitting the sediment to bypass this area to

accumulate in deeper water (base-of"-s!ope fans, or on the

abyssal plain). Komar( 1971) argued that turbidity currents are

likely to be siipercririeal (Froude number >1.0) on slopes

greater than ^O.i'^. Such turbidity currents can effectively

carry sand-sized detritus across basin-margin slopes without

leaving a deposit: there may even be net erosion in these areas.

For example. Weaver and Thomson (1993) demonstrated

slope bypassing by large mud-rich turbidity currents entering

the Madeira Abyssal Plain by examining their reworked

coccolith assemblages. Such bypassing requires that the

turbidity current be at the least self-sustaining on the slope,

or while flowing through slope channels. This condition of

.self

nuiinlcnancc (Southard and Mackintosh. 1981) was named

aiilosiispeiision by Bagnold (1962),

Gully margins and adjacent areas of the upper slope may be

sites of sediment failures as coherent slides and intertially

folded and deformed slumps. The downslope component of

gravity can cause sediment masses previously deposited on a

slope to move by increments or during a single episode into

deeper water. Slow downslope movement without slip along a

single detachment surface, i.e.. without failure, is referred to as

creep. No structures in ancient successions have been

unambiguously attributed to creep, although features in

seismic profiles have been interpreted to have formed by this

mechanism (Hill ei al.. 1982: Mulder and Cochonat. 1996),

More rapid downslope movements generate sediment slides

and debris Hows. Sediment slides can result in little-deformed

to intensely folded, faulted and brecciated masses (Barnes and

Lewis. 1991) that have translated downslope from the original

site of deposition, leaving a headwall scarp and slide sear

where the failure occurred. If the primary bedding is entirely

destroyed by internal mixing, with soft muds and/or water

mixed into the slide, then it may transform into a debris flow.

The base of the slide mass is marked by an iniraformaiio/ial

iriiiicalion surface and shear zones (Coniglio. 1986).

Initiation of slides in muddy terrigenous sediments depends

on: (1) bottom slope (Moore, 1961); (2) sedimentation rates

(Hein and Gorsline. 1981): and (3) the response of the

sediment to cyclic stress produced by earthquake shaking

(Morgenstern. 1967). Sedimentation rates on basin-margin

slopes vary widely, but Hein and Gorstine (1981) eonclude

that a rate of 30 mg cm - per year must be attained before

slope failures become common. For example, in the Santa

Barbara Basin. Calit"ornia Borderland, sedimentation rates

exceed 50 mg cm ' per year, and debris flows are widespread

on slopes of <r (Hein and Gorsline, 1981), In many areas,

sediment slides occur on very low slopes, suggesting the

involvement of water-rich, rapidly deposited sediments.

SLOPE SEDIMENTS

671

The sedimentation rate effectively determines the water

content and shear strength of the sediment, although shear

strength is also a function of other variables such as content of

organic matter (Keller. 1982) and generation of gas in the

sediment by decay of organics or by decomposition of gas

hydrates (Carpenter. 198U, According to Keller (1982. p,197).

"cohesive sediments with greater than about 4 5 percent

organic carbon [have]... (1) unusually high water content:

(2) very high liquid and plastic limits: (3) unusually low wet

bulk density: (4) high undisturbed shear strength; (5) high

sensitivity: (6) high degrees of apparent overeonsolidation: and

(7) high potential for failure [by liquef"action] in situations of

excess pore pressure'".

Slides and debris-flow deposits {debriics] are common slope

sediments. In carbonate settings, the individual blocks in

dcbritcs may be larger in diameter than 100m (Hiscott and

James. 1985: Hine cl al.. 1991), Single flow deposits can be

many tens of meters thick (Figure S34). and taper to convex-

upward snotils at their downslope and lateral margins (Hiseott

and James. 1985: Hiscott and Aksu, 1994), Shingled, down-

slope-elongate debrites are common in high-latitude slopes

(Aksu and Hiscott. 1992). where glaeiers delivered large

quantities of poorly sorted sediment directly to the upper

slope during Quaternary glacial advances.

The deeper segments of slopes on the western sides of ocean

basins are affected by geostrophic. thermohaline currents

moving at mean speeds of IO-3Ocm per second (McCave ela!..

1980:

Hollister and McCave. 1984). with short "gusts"

reaching about 7()cm per seeond (Richardson ci al., 1981),

These currents follow bathymetric contours, and are known as

conlourttirrcnls. They carry a dilute suspended load, generally

< 0.1 0,2gm '. that I'orms the thick bottom nepheloidlayer

(Fv\ing and Thorndike. 1965: Biscaye and Eittreim. 1977),

Concentrations may briefly exceed 12gm" (Biscaye cl al..

1980.

Gardner eial.. 1985). Most of the f"ine-grained suspended

material is derived by winnowing of the seafloor: the rest is

probably added to the current by cascades of cold shelf water

or lutite flows originating at the shelf slope break. The water

depths to which thermohaline eurrents influence the seabed

have changed from glacial to interglacia! times in high-latitude

oceans (Hiliaire-Marcel el ai. 1994). in some cases shifting

from the continental rise to the lower eontinental slope. It is

reasonable to assume that similar excursions of the axes of

contour eurrents occurred along ancient slopes and rises.

Contour current deposits, or coniouriies (Hollister and

Heezen. 1972). can be treated as two end members: (I) muddy

eontourites. and (2) sandy eontourites. Muddy eontourites are

flne grained, mainly homogeneous and structureless, thor-

oughly bioturbated. and only rarely show irregular layering,

lamination and lensing. They are poorly sorted silt- and clay-

size sediments with up to 15 percent sand. They range from

flner grained homogeneous mud to coarser grained mottled silt

and mud. and their composition is most commonly mixed

biogenic and terrigenous grains. According to Hollister and

McCave (1984). short-term depositional rates can be extremely

high, about

cm per year, followed by rapid biogenic

reworking by burrowers. It may be difficult to distinguish

between mud turbidites and muddy eontourites. although

some criteria ha\e been suggested (Bouma. 1972: Stow; 1979).

Sandy eontourites cotiiprise irregular layers <5cm thick

that arc either structureless and thoroughly bioturbated. or

that possess some primary parallel or cross lamination which

may be accentuated by hca\y-mineral or fbraminiferal

concentrations (Bouma and Hollister. 1973). Grading may be

normal or inverse, and bed contacts may be sharp or

gradational. Grain size ranges from coarse silt to. rarely,

medium sand, with poor to moderate sorting. The sandy

f"acies. which may be rich in biogenic sand grains, is produced

Debrite

snout

TWT(s)

,-2 7

,-2050 m

.2,8

-2,9

TRANSVERSE

TROUGH FLOOR

Figure S34 Single-chdnnel seismic dip profile from the lower slope offshore Baffin Island, Canadian Arctic (modified from Hiscott and Aksu,

19941,

The line drawing shows the rflatiunship of the lower slope to the entire margin. Stratified hemipelagic deposits (including ice-rafled

sedimentsi are punttualed by large-volume, mounded debris-flow tongues (DF1-DF3) derived from the upper slope when glaciers were

tle|)ositing till <ind outwash at the shelf edge. The debriles have clist.il snouts that testify to the intrinsic strength of !he viscous muddy debris.

672

SLOPt SEf^lMENTS

by winnowing of fines by stronger flows (Driscoll eial.. 1985),

Where currents are focused by seabed topography, as in the

Florida Strait, coarse grained eontourites may include lag

gravels and cemented hardgrounds (Coniglio and Dix. 1992),

Reworking of sandy turbidites by contour currents can result

in bottom-current-modilicd turbidite sands, believed to be

common on some continental slopes.

Muddy and sandy contourites commonly occur together in

vertical upward coarsening to upward fining sequences

(Faugeres el al.. 1984; Stow and Piper. 1984). A complete

seqtience shows inverse grading from a fine homogeneous

mud. through a mottled silt and mud. to a fine-grained sandy

contouritc facies, and then back to a muddy contourite. The

changes in grain size, sedimentary structures and eomposition

probably are related to long-term (1 kyr to 30kyr) lluctuations

in the mean current velocity (Stow and Piper, 1984),

Siliciclastie slopes typically have gradients of

6', whereas the

upper segments of" carbonate slopes can be much steeper

because of biological reef development and early eementation.

The upper slope is generally an area of bypass, cut by gullies

and slide scars. Any draping sediments are hemipelagic or

periplatform deposits (the latter for carbonate settings).

Farther downslope. turbidites and debrites punctuate the

slope succession, particularly in bathymetric lows like those

atop accretionary prisms at convergent margins. The distal

slope may be influenced by thermohaline bottom currents

producing muddy and sandy contourites. In areas where

oceanic upwelling enhances biological productivity, organic-

rich deposits, phosphates and diatomitcs may accumulate

under dysaerobic to anaerobic conditions.

Richard N. Hiscott

Oxygenation and organic-matter deposition

In some slopes, particularly along the eastern sides of ocean

basins (e.g., offshore west Africa. California, and Peru), the

wind-driven surface circulation causes upwelling of nutrient-

rich deep water, and enhaneed biological productivity above

the slope. The high biological productivity leads to organic

carbon and phosphate deposition and. locally, accumulation

of diatomaceous deposits. Oxygen is consumed to oxidize this

material, and eventually sediment pore water and possibly

even the marine bottom water become anoxic, creating an

o.xygen minimum along the slope. For example, along the

eastern Pacific continental margin, an oxygen minimum layer

occurs in intermediate water masses, commonly at depths of

250 I5()()m (Ingle. 19S1).

Two levels of oxygen depletion of bottom waters are

recognized by Byers (1977); disaerobk conditions (U.I l.OmI

O2/L HiO) under which only resistant infauna survive to

bioturbate sediment: and anaerobic conditions (<0,l ml OVL

H:.O) where no benthic organisms live and the sediment

remains undisturbed by burrowers. commonly being finely

laminated. Whore the sediment pore water is anaerobic but the

overlying bottom water is oxygenated, some filter-feeding

organisms may populate the seabed, but there will be no

infauna and therefore little bioturbation.

Facies successions and architecture of slope deposits

There is consensus that slope successions do not consist of

predictable, depositional cycles (Pickering cial.. 19H9; Coniglio

and Dix. 1992), Instead, the facies architecture is disordered.

Stow (1986) identified differences in facies associations

between muddy facies of the upper versus lower slope. The

upper slope is marked by more abundant coarse grained

gravity-flow deposits, slide scars, erosional gullies and

channels; in contrast, the lower slope is characterized by more

fine-grained turbidites. relatively rare submarine channel fills,

debrites. slide and slump scars, and local interfmgering with

contourite drifts and/or submarine-fan deposits.

Summary

Basin slopes extend from the shelf-slope break at --lOOm

depth, 10 the top of the eontinental rise on passive margins

(•^25()()m depth), or to the landward edge of the treneh along

subduction margins. This includes the entire bathyal zone.

Bibliography

AkMi. A,L,. iiiid Hiscott. R.N,. 1489, Slides und debris llovvs on tlie

higli-latitudc continental slopes of BalTm Bay. Geologv. 17: -885-

8S8.

Aksu, A,F.. iiiul Hiscou. R,N.. 1492, Sliiriglcd Upper Quulcrnary

debris flow lenses on the NL NevvfoiiiKllarid slope. Seilimeiuniotiv.

39:

193 206,

Aksu. A,E.. Calon. T,J,. Miseotl. R.N,. and Ya^ir. D.. 201)0. Anatomy

of tfie North Anatoloian Fault Zone in the Marmara Seii. western

Turkey: extensional basins above a eouiincntal transform. GSA

Tihiciy.

10 Id): ? 7.

Bagnold. R,A,. 1962, Aulo-suspeiision of transported seduiieul:

turbidity euircnis, Pifniviliii^si-fihcRoyiilSaticlvojLi>iiiitiii. Series

A,265:

315 -^14,

Uarues. P,M,, and Lewis. K.B,. 1441. Sheet slides and rotational

taiUircs on a eonvergcnt margin: the Kidnappers Slide. New

Zealand.

.Si-dimcnlolay'v.

38: 205-222.

Biscaye. P.K,. and Eitireim. S,L.. 1977. Suspended particuliite loads

and transports in the nepheloid layer of tfie abyssal Atlantic Ocean.

Marine

Geol(\^'y.

23: 155 172,

Biseaye, P.E.. Giirdnei'. W,[),. Zancveld. J.R,V,. Pak. H.. and

Tueholke. B,,

198(1,

Nepliels! Have we got nephels! EOS. Traimu-

liotisoflhv AnieiUaii Cicujihysiiul I iiiaiu 61: 10(4,

Bouma. A.H,. 1472, Reeent and aneient turbidites and contouriies.

Tninsaiiidiisiif liwCulf

(.'oasi

As.sdcialhinofGfdlo'iim!Soiiclii-s. 22:

205-221.

Bouma. A.H,. and Holli.ster. CD.. 1973. Deep oecan basin sedimenta-

lion. In Middleton. G,V.. iind Bouma. A,H, (eds.). TurhUlik:\anil

Deep Wilier

SeiiJDU'iilalion.

Analicim: Pacifie Seetion. Society of

Econoniie Paleotologists and Mineralogists. Short Course Notes.

pp.

74-118,

Byers.

C.W,, 1977, Biofaeies patterns in euxinic basins: a general

model. In Cook, H,E,. and Euos. P, (eds,). Dvcp-Wiuer Carhdiuitc

EiivirDiiiufiiis. Society of Economic Paleontologists and Mineral-

ogists. Special Publication. 25. pp, 5-17,

Cacehlone. D,A,. and Southard. J,B.. 1974. Ineipient sediment

movement by shoaling internal gravity waves. JiiuriuiIofGeophvsi-

iuiResvcirch. 70: 2237-2242,

Calvert. S,R.. 1466, Aecumulation of diatomaceous silica in the

sL'cJhiients of the Gulf of California, Gcclo^iculSocieivof.Anieriai.

Biil/eliii. 77: 569 596,

Carpenter. G,. 14S1. Coincident sediment slump/ekithiate comple,\cs

Oil the LIS Atlantic eontinental slope. Gt'o-Mnriiw

Lellcr.s^

1: 29 32,

Coniglio. M,. 1486, Synsedimentaiy submarine slope failure and

teetonic deformation in deep-waler earhonates. Cow Head group.

western Newfoundland. Cuiuiiiiun

.linii-iittl

of Earth Seicmcs. 23:

476 490.

C\)nigiio. M,. and Dix. G,R,. 1442, Carbonate slopes. In Walker.

R.G,. and James. N.P. (eds.). Eiuii-sMmlels: Responseii>Sea Eevcl

Change. Geological Association of Canada, pp. 349-374.

Driseoll. M.L,. Tueholke. B,E.. and McCave. 1,N,, I9S5, Seafloor

zonation in sediment texture on the Nova Scotian lower

continental rise, MtiiincGenloi^v. 66: 25 41,

SLOPE SEDIMENTS f.73

Dunbiir. R.I) . and Ucrgcr. W.M., 1981.

l\.'cal

pellet llux to modern

hoitom sediinciU ol'Santa Barbani

B;isin

(California) based on

sedi-

ment irapping.

Geological Soeiety

Atnerica.

Bulletin 92; 212 218.

Emery. K.O,. iind Milliman. J.D.. 1978, Suspended matter in surfaee

waters: influence t>f river discharge and of upwellint;.

Sedimentol-

ogy. 25: 125-140.

Fmery, K.O., and Lichupi. E,, 1984-

The Geology of the .4ttatitic

Ocean.

New York: Springer-Verlag.

Hwinj!,

M.. and Thorndike. E.M..

\^)b'>.

Siispeiuled matter in deep

ocean water.

Science.

Ul. 1291 1244.

Fairbridge, R.W\. and Bourgeois, J., iy78. Submarine (balliyal) slope

scdiincntalion. In I

airbridge.

R.W.. and Bourgeois. J. (eds.).

Encyclopedia

Sedimentologv.

Stroudsbiirg: Dowden. Ikitehiiison

and'

Ross, pp.' 774 777.

F;Higeres, J.-C.. Stow,

D.A.V,,

and Gontliier. F_., 11^X4. Contourlte

drift moulded by deep Mediterninean outflow.

Geologv.

12: 296

M)().

Gardner. W.D.. Biseaye, P.E., Zancvcld, J.R.V., and Richardson. M.J,.

1985.

Calibration and eomparison of

the

LDGO nephclometer and

the OSLf iraiismissometer on ihe Nova Scotian rise. Marine

Geologv.

66: .^23 344.

Gorsline. D.S.. Kolpack. R,l...

Karl.

H.A.. Drake, D.F.., Fleischer, P..

Thonilon.

S.E.. Sehwalbaeh. J.R.. and Svarda. C.E., 19H4. Studies

of fine-grained sediment transport processes and products in (he

California Condnemal Borderland. In Slow.

D.A.V.,

and Piper.

D.J.W.

(cds.),

Tine-Grained

Sediments:

Deep-Hater Proeesses

and

Facie-..

Oxford: Geological Soeie(y of London, Speeial Publiealion

15.

pp. .ly? 415.

Heath,

K.C., and Mullins, H.T., 1984. Open-ocean. nlT-bank transpori

of One-grained carbonate sediment in the northern Bahamas. In

Stow,

D.A.V..

and Piper.

D.J.W.

(eds.), Fitie-GrainedSedintetits:

Deep-Hater Proces,ses

ami

Facies.

Geological Soeiety ol' London.

Special Publication, 15, pp. 199-208.

Hee/en.

B.C., I9.'i9. Dynamic processes of abyssal sedimentation:

erosion,

transportation, and redcposition on the deep-sea lloor.

Geophysicafhntrtial.

2: 142 163.

Hein.

F.J.. and (jorsline, D.S., 1981. Geotechnical aspects of fine

grained mass llow deposits: California Continental Borderland.

Geo-Marine

Letters.

1: 1-5.

Hesse. R.. and Klaucke, I., 1995. A eontiiuious along-slope seismic

prolile from the upper Labrador slope. In Pickering. KT.. Hiscott.

R.N., Kenyon, N.H.. Rieei Lueehi. F.. and Smith, R. (eds.).

Atla.s

Architectural Stvles

Turhidite

Systems.

London: Cliupman and

Hall.

pp. IS-22. '

Hill.

P.R.. 1984. Faeies and sequence analysis of nova scotian slope

muds:

lurbidite versus "hemipelagic"' deposition. In Stow.

D.A.V.,

and Piper.

D.J.W.

(eds.).

Fine-GraitieilSediments:

Deep-Hater

Pro-

ees.\es

and

Faeies.

Geoloeieal Society oC London. Special Publica-

tion.

15. pp. .MI-3I8-

Hill.

P.R.. Moran. K,M.. and Blaseo. S.M.. 1982, Creep deformation

of slope sediments in the Canadian Beaufort Sea. Geo-Murine

Letters. 2: !63 !7().

Millaire-Mareel. C. de Vernal. A.. Lucotte. M.. Mucci. A.. Bilodeau.

G.. Roehoii. A. Vallieres. S.. and Wu. G.. 1994. Proiluelivite et llu\

de earbone dans la mer du Labrador an cours des derniers 40 000

ans.

Catiadian

Journal of

Earth

Sciences.

31: 139 158.

Hiiie.

A.C.. Locker. S.D.. Hallock. P.. and Mullins. HT.. 1991,

Strongly contrasting modes ol' slope failure and erosion along the

carbonate margins of ihrce open seaways, northwest Nicaraguan

rise.

American

A.s.weiation

Petroleum Geologists

Bulletin.

75: 595

596.

Hiscott, R.N.. in press. Latesi Quaternary Barani prodetta, north-

western Borneo. In

Sidi.

IML. Nummedal. D,. Imbert. P.. Darman.

H., and Posamentier. H.W. (cds.).

Iropieal Deltas ot Southeast .4sia

Seditnentology

Stratigraphy,

anil

Petroleum

Geology.

Tulsa: Society

of Economic Paleontologists und Mineralogists. Special Publica-

tion 76.

Hiscott. R.N.. and Aksu, A.E.. 1994. Submarine debris Hows and

eoniinental slope evolution in front of Oualernary lee Sheets.

Baffin Bay. Canadian Aretic. Anieriam

.4ssoeiatt'on

Petroleum

Geologists

Bulletin.

78: 445 460.

Hiscott. R.N.. and James. N.P.. 1985. Carbonate debris Hows. Cow

Head Group, western Newfoundland. Jout-nal of

Sedimentarv

Petrologv.

55: 7.15-745.

Ilollistcr. CD., and Heezen. B.C.. 1972. Geologie effects of ocean

bottom currents: western North Atlantic. In Gordon, A.L. (cd.).

.Siudies

in Physical

Oceanography.

2. New York: Gordon and

Breach,

pp. 37-66.

Hollister. CD., and McCave. I.N.. 1984. Sedimentation under deep-

sea storms.

.\'ature.

309: 220-225.

Ingle,

J.C Jr.. 1981. Origin (if Neogene diatomites around the North

Pacific rim. In Garrison. R.F.. and Douglas, R.G. (eds.). The

.Monterey

Formittioti atid Related Siliceous Rocks of California.

Anaheim:

Paeitie Section Soeiety of Economic Paleontologists

and Mineralogists. Special Publication, pp. 159 179.

Jansa.

L.I'., 1981. Meso/oic carbonate platforms and banks of the

eastern North American margin. Marine

Geology.

44: 97-117.

Keller, G.H.. 1982. Organie matter and tlie geoteehnieal properties of

submarine sediments.

(ieti-,Marine

Letters.

2: 191 198.

Kenter. J.A.. 1990. Carbonate platform tlanks: slope angle and

sediment fabric.

Seilimentology.

37: 777 794.

Komar. P.D,. 1971. Hydraulic jiuiips in turbidity currents.

Geological

Society

America

Bulletin. 82: 1477 1488.

Krissek. L.A.. 1984. Continental source area contributions to tine-

grained sediments on the Oregon and Washington continental

slope.

In Stow.

D.A.V..

and Piper.

D.J.W.

(eds.). Eine-Grained

Seilimctits:

Deep-Wttter Proees.sesand

Eaeies.

Geologieal Soeiety of

London Special Publication 15, pp. 363-375.

McCave. I.N., 1972. Transport and escape of fine-grained sediment

from shelf areas. In Swift. D.J.P., Duane. D.B.. and Pilkey. O.H.

jcds.).

Shelf Sedimettt

Tritnsport:

Process

and

Pattertt.

Stroudsburg:

Hiitchinson and Rcss, pp. 225-248.

MeCave, LN., Lonsdalc, P.F.. Hnllister. CD., and Gardner. W.D.,

1980.

Sediment transport over the Hatton and Gardar contourite

drifts,

Jourttal

Sedimentary

Petrology.

50: 1(149 1062.

McGraii,

D.W.. and Carnes, M.. 1983. Shclfcdgc dynatnies and the

nepheloid layer in the northwestern Gulf of Mexico. In Stanley,

D.J.. and Moore. G.f. (eds.).

The

Shelfhreak.

Ctitical

Ititerfaceon

(ontinciital Margins. Society of Econotiiic Paleontologists and

Mineralogists, Special Publication, 33, pp. 251-264.

Mooie.

D,G..

1961,

Submarine slumps. Journatof

Sedimentarv

Petrol-

ogy. 31: .M3-357.

Moore, G.F.. Taira, A.. Bangs. N.L.. Kuramoto, S., Shipley, T.H.,

Alex. CM.. Gulick. S.S.. Hills, D.J., Ike, T., lto. S.. Leslie. S.C,

McCutcheon. A.J., Mochi/uki. K., Morita, S., Nakamura, Y.,

Park. J.O.. Taylor. B.L.. Toyama, G.. Yagi. H., and Zhao. Z..

2001.

Data report: structural setting of the Leg 190 Muroto

Transect. In Moore, G.F.. Taira. A,, and Klaus. A. (eds.). Pro-

ceedings

the Ocean

Driltitig

Program.

Initial Reports, Leg 190.

College Station. Texas: Ocean Drilling Program.

Morgenstern. N.R.. 1967. Submarine slumping und the initiation of

turbidity currents. In Richards, A.F. (cd.). Marine

(ieoteehnic/ue.

tirbana: Illinois University Press, pp. 189-220.

Mulder. \.. and Cochonat. P.. 1996. Classification of offshore mass

movements,

./ournalof.SedintetitaryRe.seaich.

A66: 43-57.

Mullins. H.T.. Gardulski. A.F.. andHine. A,C. 1986. Catastrophic

collapse of the wesl Florida carbonate plalform mamin. Geology.

14:

167 170.

Nelson.

T,A., and Stanley, D.J., 1984. Variable depositional rales on

the slope and rise off the Mid-Atlantic states.

Geti-Maime

Letters.

3: 37-42.

Nemec. W., 1990. Aspectsof sediment movement on steep delta slopes.

In International Association of Sedimentologists. Speeial Publica-

tion.

10. pp. 29 73.

Pickering.

K.T.. Hiscoit. R.N.. and Hein, FJ.. 1989. Deep-MarineEn-

\ironinents: Clastic

Seditnentation

and

Tectonics.

London: Unwin

Hyman.

Pickering.

K.T., Underwood, M,B.. Taira, A., and Ashi. J., 1995.

lZANAGI sidescan sonar and high-resolution multichannel

seismic rejection line interpretation of Nankai aecretionary pnsm

and treneh. offshore Japan, In Pickering. K.T.. Hiscott. R.N..

Kenyon.

N.H.. Ricci Lucchi. F.. and Smith, R. (eds.). .AtlastifAr-

chitectural Styles

iti litrhtdite

Systems.

London; Chapman and

Hall,

pp.

34 49.

Piper.

D.J.W..

Hiscott. R-N,. and Normark. W.R.. 1999. Onterop-

scalc acoustic facies analysis and latest O^aternary development of

Hueneme and Dume fans, offshore California.

.Sedimentology.

46:

47 78.

674

SLURRY

Reimnitz, E., and Bruder, K.F., 1972. River discharge into an ice

covered ocean and related sediment dispersal, Beaufort Sea, coast

of Alaska.

Geological

Society of America, Bulletin, 83: 861-866.

Richardson, M.J., Wimbush, M., and Mayer, L., 1981. Exceptionally

strong near-bottom flows on the continental rise of Nova Scotia.

Science, 213: 887-888.

Sangree, J.B., Waylett, D.C, Frazier, D.E., Amery, G.B., and

Fennessy, W.J., 1978. Recognition of continental-slope seismic

facies,

offshore Texas-Louisiana. In Bouma, A.H., Moore, G.T.,

and Coleman, J.M. (eds.). Framework, Facies and Oil-Trapping

Characteristics of the

Upper

Continental Margin. American Associa-

tion of Petroleum, Geologists Studies in Geology, 7, pp. 87-116.

Satterfield, W.M., and Behrens, E.W., 1990. A Late Quaternary

canyon/channel system, northwest Gulf of Mexico continental

slope. Marine Geology, 92: 51-67.

Schlager, W., and James, N.P., 1978. Low-magnesian calcite lime-

stones forming at the deep-sea floor. Tongue of the Ocean,

Bahamas. Sedimentology, 25: 675-702.

Schrader, H.J., 1971. Fecal pellets: role in sedimentation of pelagic

diatoms. Science, 174: 55-57.

Shepard, F.P., Marshall, N.F., McLoughlin, P.A., and Sullivan, G.G.,

1979.

Currents in Submarine Canyons and Other Sea Valleys.

American Association of Petroleum Geologists, Studies in Geol-

ogy, 8.

Southard, J.B., and Mackintosh, M.E., 1981. Experimental test of

autosuspension. Earth Surface Processes and Landforms, 6:

103-111.

Stow, D.A.V., 1979. Distinguishing between fine-grained turbidites

and contourites on the Nova Scotian deep water margin. Sedi-

mentology, 26: 371-387.

Stow, D.A.V., 1986. Deep clastic seas. In Reading, H.G. (ed.). Sedi-

mentary Enviromnents and Facies, 2nd edn, Blackwell Scientific,

pp.

399-444.

Stow, D.A.V., and Piper, D.J.W., 1984. Deep-water fine-grained

sediments: facies models. In Stow, D.A.V., and Piper, D.J.W.

(eds.),

Fine-Grained Sediments: Deep-Water Processes and Facies.

Geological Society of London, Special Publication, 15, pp. 611-646.

Thornton, S.E., 1981. Suspended sediment transport in surface waters

of the California Current off southern Cahfornia: 1977-78 floods.

Geo-Marine Letters, 1: 23-28.

Thornton, S.E., 1984. Basin model for hemipelagie sedimentation in a

tectonically active continental margin: Santa Barbara basin, Cali-

fornia Continental Borderland. In Stow, D.A.V., and Piper, D.J.W.

(eds.),

Fine-Grained Sediments: Deep-Water Processes and Facies.

Geological Society of London, Special Publication, 15, pp. 377-394.

Weaver, P.P.E., and Thomson, J., 1993. Calculating erosion by deep-

sea turbidity currents during initiation and flow. Nature, 364:

136-138.

Wilson, P.A., and Roberts, H.H., 1995. Density cascading: off-shelf

sediment transport, evidence and implications, Bahama Banks.

Journal of Sedimentary Research, A65: 45-56.

Cross-references

Debris Flow

Gravity-Driven Mass Flows

Nepheloid Layer, Sediment

Oceanic Sediments

Slide and Slump Struetures

Substrate-Controlled Ichnofacies

Synsedimentary Structures and Growth Faults

Turbidites

Upwelling

SLURRY

Introduction

A slurry is a mixture of solids and fluid that is highly mobile

under gravitational and low magnitude stresses, and remains

coherent (no bulk phase separation) during transport. In

geophysical settings, a slurry is part ofa spectrum of mass-flow

phenomena known as sediment gravity flows. In agricultural,

industrial, and engineering settings, a slurry is a coherent two-

phase mixture that commonly is pumped into place rather than

driven by gravity. The criterion of high-mobility is included

here to distinguish a slurry from concentrated, low-mobility

industrial suspensions of solids and fluid such as gels, greases,

and pastes. The fluid medium of both geophysical and

industrial slurries commonly consists of water but in some

circumstances may consist of a gas or oil. The solids

component of a slurry can have widely varying composition,

size distribution, and concentration. Common geophysical and

industrial slurries can consist of cohesive fine sediment,

inorganic granular debris, rock fragments, coal, animal waste,

sewage sludge, snow, or ice. Some familiar examples of slurries

include freshly mixed concrete, drilling mud, and debris flows.

The volume concentration of solids in a slurry can vary widely.

In general, the finer grained the solids, the lower the

concentration needed to obtain a coherent solids-fluid mixture.

For example, a slurry composed of smectite and water can

become coherent at a solids concentration of as low as 3

percent by volume (Hampton, 1972), whereas a granular

debris flow typically does not become coherent until the solids

concentration reaches about 50 percent or more (e.g., Iverson

and Vallance, 2001).

Distinguishing characteristics of a slurry

A slurry differs from a dilute suspension or solids-laden fluid in

terms of momentum transfer, the pressure in the suspending

fluid, and phase separation of the solid and fluid constituents.

In dilute suspensions or solids-laden fluids, momentum is

transferred by fluid forces, and the solids are a passive

component transported by the fluid (see Gravity-Driven Mass

Flows).

In a slurry, however, significant momentum transfer

occurs through solid-phase interaction, owing to the greater

concentration of particles. Solids are no longer transported

passively, but instead move the suspending fluid in response to

a pressure gradient or gravitational body force.

Nonequilibrium fluid pressure, the total fluid pressure minus

the local hydrostatic fluid pressure, further distinguishes a

slurry from a solids-laden fluid. In a concentrated slurry,

nonequilibrium fluid pressure arises from gravitational settling

of the particles, which transfers stress to the fluid phase. If the

fluid can migrate easily through pore spaces around the solid

particles, the nonequilibrium fluid pressure dissipates readily,

otherwise it remains elevated and helps balance the weight of

the overlying mixture. As long as pore-fluid pressure remains

elevated, particle settling is hindered, and separation of the

solid and fluid phases is limited.

A lack of phase separation over timescales relevant to the

transport process can be used to further distinguish a slurry

from a solids-laden fluid. In a solids-laden fluid, the phases are

not intimately bound, and gravitational settling of particles is

controlled by the size, shape, and density of the particles, by

the density and viscosity of the fluid, and by fluid dynamics

(such as turbulence). Gravitational phase separation in

slurries, however, is related to dissipation of nonequilibrium

fluid pressure (e.g.. Major, 2000; also see Gravity-Driven Mass

Flows).

An approximate estimate of the time necessary for

phase separation to occur in most geophysical slurries is

obtained by examining the ratio h^lD, where /; is the thickness