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

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TAFHONOMY: SEDIMENTOIOGICAL IMPLICATIONS OF FOSSIL PRESLRVATION

725

SETTLING

WAVE (OSCILL,)

Mytilus

Turfitella

RIVERS, TIDAL STREAMS

High drag (unstable)

Low drag (stable)

Concavo - Convex Shells B

CURRENT

PREFERRED Concordant

ORIEN-

TATION ,.„ _ ^, ,

Perpendicular/

Nesting

Figure T2 Orientations and fabrics tjl shells under varying tonditions,

(A) sfltling ot concavo-convex shells under free f.ill conditions;

(B) re-orientdti(in of concavo-convex shells hy currents; (C) himndal

orientalions of shells in oscillaling

urrents, (D) alignment of shells in

unidirectional currents; rose diagrams show directions of apices of

shells; arrows indicale current directions; lE) various types of shell

fabrics in pl.in and cross sectionol view. Modified from Allen (1492)

and Kidwell and FHolland (1991),

CoiiCL-ntriitioiis of shells in localized pockets commonK

display coiivex-downwiird or Uiteritl-obliqiic positions iilong

lincLir leatLires th^il may be interpreted as subtle scours or

giUter casts (Fulterer. 1978), 'Ihcir presence is a significant

indicator of siorm-getierated erosion and redeposition.

Concavo-convex shells mosl commonly display prelerred

convex-upward orientations (Figure T2(B)). Even gentle

currents will cause shells to flip into a hydrodynamically stable

convex upward orientation (Middleton. 1967; Brenchley and

Newall.

1970). Thus, the occurrence oi abundant shells in

convex-upward positions on bedding planes provides e\idence

for reuorking oi shells by one or more storni-generaled

gradient currents.

Beds oi preferentially concave-upward shells are not as

common.

When the shells are allowed to resettle from

suspension they will generally settle eoneave-upward

(Figure T2(A)). Such reorientation might occur in areas close

to storm wave-base, where gentle storm-generated waves

would lift shells olT the bottom and allow free-fall back to

the substrate. Conca\e-upward accumulations in rippled, cross

bedded sands also nia> arise from free-fall of shells down the

lee sides of migrating bedforms (Brenchley and Newall. 1970),

Preferred orientation of elongate skeletons has been the

subject ol' numerous experimental studies (see Nagle, 19fi7;

Brenchley and Newall, 1970; Reyment. 1980; Kidwell and

Bosenee, 1991). Elongate particles typically orient themselves

selectively in a current. Cylindrical objects may roll perpendi-

ctilar to the current direction. Hlongaic objects that do not

roll,

e.g.,

rhabdosomes of graptolites, will become oriented parallel

to the current. Similarly, conical shells, including many

nautiloid cephalopods, tentaculitids. and gastropods, will

become aligned parallel lo the current with the apex directed

up-current (Figure T2(D)). Hence, depending upon the type of

fossil,

it may be possible to determine not only Ihe line of

ctirrent action, but also the direction of propagation.

Bimodal orientations are eommon and may result from

more than one cause. In some instanees. the bimodality is size

dependent; for example, objects, such as fragments of large

nautiloids. may approach a cylindrical shape and roll with the

current, whereas smaller cone-shaped shells will oiient in a

preferred up-current direction. Conversely, conical shells of

similar size and shape may show nearly perfect

180 bimodality. aligned parallel to subtle wave ripples

(Figure T2(C)), Such fossil evidence indicates oscillatory

currents or waves.

Stacks of edgewise or shingled shells (Figure T2(E))

apparently occur where densely packed shells were affected

by oscillatory, storm-generated waves and indicate deposition

within storm wave-base (Seilacher and Meischner, 1965:

Grinnell.

1974), A similar mode of preservation in shell

aeeumulations involves eup and saucer stacking in which

concavo-convex shells are bundled together in nested group-

ings.

These, too. seem to reflect the effects of settling and

concentration of shells during turbulence events and they are

commonly associated with minor sediment grading.

Another, perhaps related, phenomenon involves the cone-

in-cone interpcnctration of nauiiloids observed in some

densely packed cephalopod limestones (Figtire T2(E)), These

appear to represent turbulence effects that lirst align and then

telescope cephalopods: perhaps septa had been weakened

somewhat by dissoltition or corrosional effects.

Upright amtnonoids and nautiloids have been observed in

a number oi sections in the geologic record. Cephalopod

phragmocones will initially settle in a \ei tical attitude during

the Hooding oi the chambers. There is considerable variation

based on shell shape and type, but most shells will settle

vertically oniy up to about lOm (Raup. 1973); beds with

abundant upright ammonoids may thus provide a useful depth

indicator. However, in orthoconic nautiloids, ballasting of the

apical end of the phragmocone by cameral deposits may also

cause shells to sink vertically and embed with apex downward

following deeay and removal of the sof( body parts. In either

case,

preservation of this attitude demonstrates that the

sediments were relatively soft to allow shells to penetrate

some distance into the seafloor.

Sorting and fragmentation

Skeletal parts may be acted upon by various hydrodynamic

processes, which serve to sort and fragment them. Sorting of

skeletal elements by size, density, and even shape, is eommon

in environments above wave base. Differential transport of left

versus right valves of equivalved pelecypods

(e,g,.

Figure

r2(C)) may occur in environments where waves approach a

coast obhquely (Muller, 1979). Hence, left-right sorting may be

an excellent indicator of wave swash environments.

Biased ratios of skeletal parts may provide an indieator of

time-averaging in shell-rieh deposits. F\ir example, brachiopod

assemblages commonly show several times as many whole

pedicle as brachial valves (Holland. I9S8; Brett and Bordeaux.

1990). Ihis indicates that (he pedicle valves, which possess

hea\v hinue dentition and a thickened interarea. were betler

726

TAPHONOMY: SEDIMENTOLOCICM IMi'LICATIONS OF FOSSIL PRESERVATION

able to withstand destrtietive processes. Offshore, low energy

facies thai show strotigly biased valve ratios represent time-

averaged accumulations. During prolonged intervals of

exposure, occasional storm disturbances fragmenied shells,

preferentially removing the more fragile skeletal parts.

Fragmentation of shells may oceur v/(/ biolie or abiotic

means. Much fragmentation is thought to reflect physical

impact of shells with one another. Generally, high proportions

of fragmented skeletal material indicate repeated processing by

high energy events.

Shattering of septa in nautiloid and amtiionoid cephalopod

shells is eommon in some offshore faeies and has been used as

a paleobathyinetric indicator (Westermann, 1985)- The shat-

tered phragmocones may refleet hydrostatic pressures on

an empty body chamber that shattered septa. Given this

assumption, and septal strength indices for internally

shattered cephalopods. approximate depths of implosion can

be inferred- However, more reeent work by Hewitt and

Westermann. (1996) suggests that some of the shattering

may aetually take place due to explosion: i.e,. septa blown

forward owing to back pressures. Punctures in the shells may

allow water to penetrate into back chambers of the shell to

cause shattering oi weaker early septa at much shallower

depths than those predicted by the itnplosion model. If these

two aspects of shell internal fragtnentation can be sorted out,

then shattered phragmocones may indeed provide a useful

paleobathynietric indieator,

Corrasion and encrustation

It is often difficult to distinguish between shells that have been

physically abraded and those corroded by biogeochemical

processes. Brett and Baird (1986) suggested the term corrasion

for cases of unspccilied corrosion and abrasion. Even such

unspecified shell damage may indicate exposure in the

depositional environment.

Physically abraded fossils are common in many sandstones

and siltstones. but rare in mudroeks. as clay sized particles are

not affective abrasive agents. Abraded fossils in shales might

indicate a complex burial history in which shells had been, at

some point and time, affected by sediments coarser than the

matrix in which they reside.

Corrosion and/or bioerosion of shells predominates over

physical abrasion in offshore, low-energy environments

(Kidwell and Bosence. 1991), Many apparently abraded shells

in tnudrocks have probably been chemically etched and/or

acted upon by microboring organisms.

Microborings of endolithic organisms, identified by SEM.

may be useful indicators of exposure and relative water depth,

Endolithic algae of various sorts, for example, are confined to

various portions of the euphotic to upper dysphotic zone

(Vogel

(•/(//..

1987). Absence of these sorts of endoliths on

shells from sediments, in which independent evidence indicates

shallow water deposition, mighl be used as an indieator of

turbidity or elevated seditnentation rates.

Even in life, skeletons may become encrusted with epibiontic

organisms, such as bryozoans. After death of their occupants,

shells may be encrusted both externally and internally. Internal

encrustation provides an excellent indication that shells

remained unburied after death. Skeletons that have been

corroded,

or abraded may be subseqtiently encrusted and

provide evidence for more prolonged exposure. Fossil corals

that show one side with the epitheca perfectly intact and the

other side corroded and subsequently encrusted (Brett and

Baird,

1986) must have lain partially buried in the mud; the

upper surface was exposed to destructive and encrusting

agents.

Fossil diagenesis: geochemical processing

Fossil diagenetic features also prove valuable in the inter-

pretation of sedimentary environments. Early diagenetic

phenomena may provide information regarding the geochem-

istry of bottom waters and the upper, "taphonomically

active zone" (TAZ) of the sediment column (Martin. 1999;

Figure T3). Diagenetic features of note inelude evidence for

early dissolution, compaction, and mineralization of Ibssils.

Skeletal minerals and dissolution

Organisms secrete a variety of minerals, including calcite.

aragonite. silica, and phosphate (apatite). Different skeletal

mineralogies have varying stabilities.

Carbonate skeletons are commonly preserved in normal

marine environments, although those composed of aragonite

are more prone to dissoltition. The large biocrystals of

echinoderm plates are exceptionally resistant to dissolution

and may be the only skeletal earbonates remaining after

prolonged exposure.

Living organisms

(Epifaunal

Products

of organic

maner decay

Geochem.

Zones

Figure T3 The realms ofbiostriitinomy and ejrly diagenesis in marine

sediments overlain by oxic water. Lett columns show transition uf

organism remains from epifaunal and infaunal community into the

physically/biologically mixed or taphonomitally active zone (TAZ)

and ils transition to the permanent fossil record. Note some upward

reworking of skeletal remains in the TAZ; also note intact skeletons in

abru|)lly buried obrution deposil, which is also the site of pyrite

nucleation.

Middle columns highlight bacterially-mediated diagenetic

processes and produt

fs;

right columns ifientify geochemical zones.

Abbreviations: RPD: redox potential discontinuity (oxic-anoxic

boundary); Trans: transitional redox conditions. Adapted from

diagrams in Martin (1999) and Wignall

(19941.

TAPHONOMY: SfDIMFNTOLOGICAL IMPLK ATIONS OF FOSSIL I'KESFRVATION

727

C;ii-bonate skeletons thai dissolved prior to compaction may

be preserved as piastically deformed molds. Such presei'vaiion

would indicate undersaluraiion with respect (o aragonite and/

or calcite in the upper sediments. Conversely, many fossil

molds display mosaic fracture patterns on their surfaces (Brett

and Baird, 1986). Such shells survived without dissolution until

after early phases of compaction and brittle fracture of the

shells.

Even more stable than echinoderm plates, are skeletal

elements composed of dense calcium tluophosphate (apatite).

In certain environments, particularly dysoxic. organic-rich

sediments, bone apatite (organic fiuophosphate) may be

replaced by the far more stable form calcium carbonate

liuorapatite, Fluorapatitc is stable over a wider pH/Eh range

than calcite. and is one ol' the most stable biogenieally

associated components (Lucas and Prevot, 1991). Once bones.

leeth. conodonts. or fossil molds have been replaced by

calcium fluorapatite. they are resistant to further dissolution.

Such materials may form bone beds, which remain stably on

the sea bottom for periods a\' htindreds of thousands to

perhaps millions of years (Kidwell and Bosence, I99i; Martill,

1991).

Early diagenetic mineralization

Early diagenetic minerals in fossils, form as a result oi' the

action of anaerobic bacteria (Allison and Briggs. 1991; Briggs,

1995).

They may provide valuable information on sediment

geochemistry and rates of burial,

Pyrite is common in marine mudrocks because of the

availability of iron in terrigenous muds and the ubiquitous

abundanee of dissolved sulfate in marine (but not in fresh)

water. In anoxie sediments generally low in organic matter

early diagenetic pyritization tends to be localized in the vicinity

o\'

anaerobically decaying organisms (Hudson. 1982; Brett

and Baird. !9S6). As a result of local suifate reduction.

H:S is liberated around decomposing organic material

(higure T3), Dissolved iron will react at these sites of sulfide

production

Well preserved, pyritized fossils tend to occur in bioturbated

mudstones. indicating environments with limited bottom water

oxygenation. and ano,\ic. sultidic microenvironments within

the sediment (Hudson. 1982; Brett and Baird. 19S6; Canlield

and Raiswell. 1991a),

I\>ssils enclosed within carbonate concretions ;ire typically

three-dimensional; hence, the concretions formed before shells

could be crushed by overburden pressure.

Most anaerobic decay processes generate bicarbonate

(HCO,) which may initiate the precipitation of ealeite or

siderite concretions around decaying organic matter in the

/ones of nitrate or sulfate reduction (Berner. 1981; Canfield

and Raiswell. 1991b; figure 3).

However, mass balance calculations have shown that the

bicarbonate generated by decaying carcasses is insufficient lo

!orm a concretion. Carbonate concretions, including those

whieh enclose well preserved fossils, were largely generated by

the decay of disseminated organic matter within the sediment.

The development of large concretions requires that the

sediment surrounding a local nucieation center remains within

the /one of sulfate reduction or methanogcnic zone (Kigure T3)

in the upper lew decimeters of the sediment column for

prolonged periods, probably at least a few thousand years

(Raiswell. 1976; Canfield and Raiswell. 1991b; Wignall. 1994),

Hence, horizons of fbssil-bearing concretions may reflect

episodes of sedimentation that were then followed by sediment

starved intervals. Not surprisingly, many concretion beds

underlie recognized diastems and concretions may also become

reworked into erosion lag beds (Brett and Baird. 1986).

The decay of organic matter and reduction of phosphatie

iron hydroxides liberates phosphate-bearing compounds

(Lucas and Prevot. 1991; Wignall. 1994; Figure T3), Dissolved

phosphate may be released to the water column, if anoxia

persists to the sediment-water interface. However, if a thin

oxic zone exists in the upper sediment then the phosphates

may be reprecipitated (Berner. 19S1; Allison. 1988). especially

around phosphatie skeletal nuclei, such as bones or lingulid

brachiopods.

If the sedimentation rate is very low and geochemical

conditions remain relatively constant, then phosphorus com-

pounds will be sufTicientiy concentrated that phosphate

minerals ean precipitate. These minerals may replace organie

remains or can form concretions. Thus, the occurrence of

phosphatie fossil molds or concretions is nearly always an

indicator of sediment starvation, commonly associated with

deepening episodes.

Under less strongly reducing conditions the precursors of

ironstones, such as berthierine or chamosite may precipitate in

the upper part of the sediment, commonly as oolitic coatings

or impregnations of skeletal grains. Later diagenesis may

further chemically alter these precursors to hematite or

goethite. "Fossil ores"'commonly occur near the interfaces of

reduced iron rich muds and carbonate shoals, suggesting

critical redox boundary conditions that existed in relatively

low sedimentation areas of mixed siliciclastic-carbonate

deposition. Widespread beds of "fossil ore"", i.e,, hematite or

chamosite coated fossils are also excellent indicators of

sediment starved discontinuities; they commonly intergrade

with phosphatie nodule beds (Brett. 1995).

Prefossili/ed material may be identified by breakage along

formerly continuous structures, rotated geopetal structures,

and shell infillings that do nol match surrounding matrix

sediment. Perhaps the most diagnostic feature is abrasion.

bioerosion. or encrustation of diagenetically altered surfaces

(e,g.. abraded or bored phosphatie steinkerns) and preferential

alignment (Kidwell and Bosence. 1991), In rare cases, even

trace fossils may be diagenetically cemented and then

reworked.

Summary discussion: preservational

processes and taphofacies

Various aspects of fbssif preservation may be combined to

recogni/e laphnfucics characteristic of various sedimentary

environments (Speyer and Brett. 1986. 1988), Together with

lit

ho-,

bio-, and ichnofacies. taphofacies tend to vary

predictably with sedimentary environments as shown by

studies in modern marine settings (e.g.. Parsons and Brett.

1991).

Fossils and their modes of preservation may provide

important insights into a number of features of sediment

deposition. These include: (i) sedimentary environment (depth,

temperature, salinity, oxygen level, substrate consistency);

(ii) dynamics of sediment accumulation, average rates, as well

as evidence for episodicity of sedimentation and erosion;

(iii) temporal scope of individual mudrock units; and

(iv) sediment geochemistry and early diagenetic environments.

728

TAPHflNOMY: SEDIMENTOLOCICAL IMPLICATIONS OE EOSSIL PRESERVATION

Taphonomic features. (e,g,. articulation of delicate skeletal

elements) may provide unambiguous evidence for episodic

sedimentation; conversely, highly corroded fossil material

provides a distinctive signature of time-averaging. The overall

condition of skeletons may be assessed qualitatively or

quantitatively, but certainly should be noted in the field.

Semi-quantitative indices (e.g.. of corrosion or fragmentation)

may include the proportion of skeletal parts in different,

arbitrarily defined preservational states, ln such cases, it is

commonly useful to develop a set of standards with whieh

particular shells may be compared and assigned to a category

in much the same way ihat grain shapes and roundness indices

have long been assessed on the basis of standardized profiles

by sedimentoiogists. Skeletal condition is particularly valuable

for recognizing qualitatively the differing relative extents of

sedimentary time-averaging that may be related to sequence

stratigraphy (Brett. 1995).

Overall, condensed, highly time averaged, shelly accumula-

tions will display high frequencies of disarticulation. fragmen-

tation, valve ratio biasing, corrosion, and/or abrasion

(FUrsich. 1978; Kidwell. 1991; Brett. 1995). Condensed

intervals, associated with marine transgressions, thus may be

readily recognized.

Conversely, rapidly accumulated sediments, typical of later

highstands. will display higher proportions of articulated

material with little or no shell breakage, including intact,

delicate f\irms. such as branching bryozoans. little or no

biasing of valve ratios, and few. if any. fossils that are heavily

corroded, abraded, or otherwise altered in their surf\ices

(Kidwell. 1991; Brett, 1995),

Taphonomic observations should accompany any interpre-

tation of fossiliferous sediments. By making eveti a qualitative

assessment of the preservation states of fossils, a sedimen-

tologist may be able to determine a great deal abotit the

dynamics of sediment accumulation and many other paleoen-

vironnienta! parameters,

Carlton E. Brett

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verte-

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The Processes

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Middfeton. Ci.V,.

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Nagle.

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Tupfionomic processes

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Endofith associaliotis

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263-290,

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septal implosion

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W..

1994,

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Monographs on

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Geophysics. Volume

30.

f27pp.

depositional system, which

act

independently

of" the

extrinsic

vitriables.

primary goal

in the

study

sediment

and

sedimentary rocks

is to

deeiplier

the

relative effects

of the

e,xtrinsic

und

intrinsic processes

on the

depositional history

a sedimentitry basin.

Tectonics afTeets sediinentation

in two

different, btit often

related ways: subsidence

of the

crust

and

uplift ol"

the

land.

Althotigh sediment

may be

deposited ternporarily

tectoni-

citlly inactive, topographically

lou-

areas

of the

crust,

it is

subsidence over geologically significant intervals

time that

continually provides

the

accommodation space

for the

aceiimulation

and

burial

sediment

sedimcnlary basins,

Stibsidence

may be

driven

vertical ofTsct

faults, lateral

flou'

the mantle, cooling

and

concomitant increase

density

of the crust and/or mantle,

and

isostatic response

the crust, including

the

load provided

sediment within

the

basin (Ingersoll

and

Busby. 1995). Often working

concert

with subsidence, uplift

the land provides detrital sediment

depositional basins through weathering

and

erosion. Vertical

tectonic tnovement

of the

crust

required

maintain high

regions

they will ultimately

reduced

elevation

near

sea level.

Early attempts

classify sedimentary basins were

ham-

pered

by the

lack

ofa

comprehensive tnodel

explain crustal

deformation (Schuchert.

1923; Kay.

1951).

It is now

widely

accepted that

the

dominant tectonic process affecting

the

crust

and upper mantle

of the

earth during most

of its

history

plate tectonics, with mantle plumes

secondary importanee.

Consequently,

the

following discussion

sedimentary basins

is organized according

their plate-tectonic settings. Other

references

this topie include Allen

and

Allen (199(1)

and

Busby

and

Ingersoll (1995).

Cross-references

Algal

and

Bacterial

(

arbonate Sediments

Bacteria

Sediments

Bioeiasts

Bioerosion

Biogenic Sedimentary Structures

Black Shales

t oal BalK

Diagenetif Structures

Micritization

Mudrocks

Piileocurrent .-Analysis

Phosphorites

Reefs

Sediment Fluxes

and

Rates

Sedimentation

Sedimeniologists

Storm Deposits

Substrate-C\>ntrolled Ichnof'aeies

Sullide Minerals

Sediments

TECTONIC CONTROL OF SEDIMENTATION

Introduction

Sedimentation

influenced

three e,\triiisic variables:

tectonics, climate,

and sea

level, with

the

latter

two

potentially

dependent upon each other

and

upon tectonics.

addition,

sedimetitation

affected

processes inherent

to the

Sedimentary basins associated with diverging plates

Initial breakup

continental crust produces

long, linear

extensional terrane referred

to as a

continental rift.

some

cases rifting initially takes place simultaneously along three

arms (triple junction),

one of

which eventually fails while

the

other

two

evolve into

the

rifted continental margin.

The

failed

arm (aulacogen) represents

region

thiek sediment that

extends

far

into

the

continental interior (Hoffman ciaL, 1974).

Continental rift basins

are

either bordered

both sides

(full graben)

or one

side (half grabcn)

normal faults that

extend

least into

the

middle crust.

It is

movement

on the

normal faults that provides

the

primary mechanism

basin

subsidence

and

adjacent uplift.

The

margins

of a

continental

rift basin commonly contain coarse detritus deposited

alluvial fans.

In a

half graben. small catchments

on the

footwail scarp result

small fans, while larger catchments

the hanging-wall mountains produce broad fans that extend

far into

the

basin (Figure

T4;

Leeder

and

Gawthorpe. 1987).

Basin-center facies

rift basins

may

include lakes,

axial

river, and/or eolian sand field. Large lakes

are

particularly

common

in the

East African rift (Soreghati

and

C\ihen, 1996).

while

the Rio

Grande rift

of the

southwestern United States

has axial rivers (Mack

etaf,

1997). Continental rift basins also

may

inundated

shallow seas, particularly

if the

rift

near

sea

level

at the

onset

extension (Gawthorpe

cf af.

1997).

Normal faufts

continental rifts have finite lengths usuafly

measured

in a few

tens

to a

hundred kilometers. Strain

transferred f"rom

the tip of one

active fault

another

via an

730

TECTONIC CONTROL OF SEDIMENTATION

Figure T4 Schcnicitic map and cross-section ul ii Lunlincnl<il

riti,

showing fiasin-houndin^5 norm.il faults (ball ancf stick on down side

of tault). an atcummodalion zone in which the fiordt'r faults overlap

and dip in the same direction (upper part of diagram), an

acjcommodalion zone in which ihe border faults overlap and dip

toward each other, and ihe asymmetrical distributi<jn of alluvial-fiin

and axial-fluvial environments in half grabens.

accommodation zone, whose structural style cicpcnds on the

degree of overlap and dip direction of the faults (Morley

('/(//,.

1990),

Because o\' large catchments, itcconimodation zoties

may be the locus of sedimentation by large alluvial fans or

deltas (Figure T4).

Continental rifting may eventtially split the continental land

masses apart, creating between them a narrow, incipient ocean

basin (loored by nascent oceanic crust, such as exists today in

the Red Sea (Purser and Hot/l. 1988), The narrow seaway may

be initially flanked by high escarpments, narrow coastal plains,

and marine shelves, whose depositional environments include

alluvial fan. fluvial, siliciclastie shoreline, including small

deltas,

and marine carbonates. With time the marginal rift

blocks are reduced by erosion and onlapped by sediment.

Because of a narrow connection to the open ocean and poor

circulation, the incipient ocean basin may accumulate line,

organic-rich sediment and/or evaporites.

With increasing distance from the oceanic spreading ridge,

the previously rifted continental tnargin subsides slowly and

asytnmeti'ically about a hinge on the land\vai"d side. Sub-

sidence, which decreases exponentially with time, is pritnarity

related to the atnount of crustal thinning, cooling of the

passive continental margin sedimenCary prism

coastal plain shelf slope

J rise

Figure T5 Cross-section of a passive continental margin showing Ihe

major depositional settings.

lithosphere. and sediment loading (Bond cl uf. 1995). The

resultant passive continental-margin prism of sediment is

deposited in coastal plain, shallow-marine

shelf,

and deep-

water rise environments, and consists predominantly of tine

siliciclastie and/or carbonate sediment (Figure T5), Sedimenta-

tion may be locally influenced by and strata disrupted by

contemporaneous rise of diapirs of salt deposited in the

incipient ocean basin phase, Along-strike variation in passive

margin sedimentation is related to continental transfonn faults

that produce an orthogonal continental margin (Thomas.

1991),

The thickness of sediment deposited at active oceanic

spreading ridges is initially influenced by hall' grabens in the

oceanic crust. Although the earliest seditncnl may be umber,

the iron- and manganese-rich fallout from black smokers

(Robertson and Hudson. 1973). sedimentation on oceanic

crust is dotniiiatcd by pelagic sediment composed of carbonate

and siliceous shells and clay that settle from suspension. The

global distribution of pelagic sediment is primarily related to

proximity to continental land masses and to the productivity

of planktonic organisms (Davies and Gorsline. 1976), Tccton-

ism comes into play in pelagic sedimentation through

subsidence of the oceanic crust, a litnc-dcpendent process

driven by cooling of the oceanic plate as it moves away from

the spreading ridge. The subsiding plate may eventually fall

below the carbonate and silica compensation depths, inhibiting

the deposition of carbonate and siliceous oozes and promoting

deposition of clay (Heezen ci af. 1973). Voluminous out-

pouring of basaltic lava on the ocean floor, sueh as occurs

above mantle pltimes. may build intraoceanic islands. Once the

volcanism has slowed or ceased, however, the oceanic plateau

will subside to abyssal depths, producing a stratigraphy that

may include carbonate reef and shoal deposits overlain by

pelagic sediment (Clague. 19S1).

Sedimentary basins associated with subduction

The process of subduction of an oceanic plate beneath another

oceanic plate or a plate with continental crust on its leading

edge results in a series of depositionat basins that are bordered

by tectonic elements related to the subduetion proeess

(Figure T6), The topographic low at the Jtinction of the two

plates is the trench, and oceanic sediment and basaltic crust

scraped off of the subducting slab onto the overriding plate

creates the topographically high subduction complex. Partial

melting of the astbenosphere above the subducting slab

generates tnagma that rises to create a volcanic arc and its

underlying plutons.

Mostly pelagic sediment, locally alTected by thertnohalinc

currents, is deposited on the outer trench slope, where Hexure

of the subdticting slab may result in extensionai basins

TE(

TONIC: CONTROL

Of SEOIMbNIAIiON

731

outer

trench trench-slopc

slope basins

Figure

ross-set

!ion

converj^ing

showing

the

subduction

complex

and arc, and

iheir

associated sedimentary

basins.

ll-'igure

T6;

Underwood

and

Moore. 1995),

addition

pelagic sediment,

the

trench

and

trench-slope basins

may

recei\e transverse-

longittidinally dispersed gravity flows

and stibmarine slides. The amount

sediment

the trench

highly variable

and

depends

on the

ratio

rate

scditnent

inllux

the rate

subduction. Developing

the subduction

complex, trench-slope basins

may

consist

grabens. linear

valleys separated

fault-

and

fold-generated ridges, and/or

thin apron

sediment.

The forearc basin, whieh

situated between

the

magmatic

arc and

tbe

subduction complex,

generally

the

largest

the

subduction-related basins (Figure

T6;

Dickinson. 1995).

Usually less than

150

km wide, forearc basins

may

extend

laterally

for

thousands

kilometers

or be

segmented itito

series

smaller basins. Sediment thickness

forearc basins

may

be up to

10 km. and the sediment commonly progressively

onlaps both

the arc and the

subduction complex. Subsidence

tnechanisms

forearc basins are poorly understood,

but

may

include downward pull

of the

subducting plate, teetonic

by the

subduction complex, cooling

of a

relict

arc.

and sediment loading.

The

principle source

siliciclastie

sediment

is the

volcanic arc. although

the

plutonic core

the

arc

may be

periodically exposed. Less commonly,

the

subduction cotiiplex

may

provide sediment

to the

forearc

basin

via

submarine slides

or by

erosion

of a

cotnplex that

builds above sea level. Many foreare basins are pritnarily deep

marine,

whicb hemipelagic muds

and

ttirbidite siinds

are

deposited.

It is

also possible

for

forearc basins

infill with

shallow-marine

and

nonmarine siliciclastie sediment,

and

marine forearc basins bordered

only

a few

volcanic islands

may experience significant carbonate deposition.

Many active volcanic arcs contain intra-arc basins, which

may

fatill-botind

simply occupy topographically

low

regions between

the

volcanoes (Figure T(i: Smith

and

Landis,

1995).

Accumulation rates

sediment

and

volcanic rocks

are

high, because

active eruptions

and

erosion

high-relief

\olcanoes. Proximal

to the

volcanoes, lava flows

and

pyroclastic flows

are

interbedded wilb coarse gravity flow

deposits, while

tnore distal settings volcanie

ash is

interbedded with tine-grained alluvial-fan, liuvial,

and

lacus-

trine sediment.

oceanic areas,

it is

common

for

part

the

intra-arc basin

be marine and contain fan-delta, delta, deep-

sea

fan.

hemipelagic,

and

carbonate deposits, Intra-arc basins

provide important infonnation about the e\olution

the are.

but

may be

diflicult

distinguish from other arc-related

basins

ancient rocks.

In some cases, lithospheric extension occurs behind

tbe

arc.

creating

back-arc basin (Figure

T6;

Marsaglia. 1995),

Back-arc basins may develop in intraoceanic arc (e,g,. Mariana

arc)

continental-margin arc (e.g.. Sea

Japan) settings and

are especially common

in the

western Pacific associated with

rapid subduction

old. steeply dipping oceanic lithosphere,

F.xtension

driven

mantle flow above the subducting plate

and/or

relative movetnent

of the

trench tov\ard

the

subdticting plate (roll back) related

steepening

the angle

of subduction. Initial breakup commonl> takes place

in the

thermally weakened

arc,

splitting

remnant

arc

from

the

active

arc.

Tectonics

and

sedimentation

back-arc basins

resemble that

rifts

and

acting spreading ridges, with

the

exception that the active arc

is an

important source

detrital

sediment

and

volcanic rocks.

Sedimentary basins associated with compressional

mountain building

Contraction

continental crust results

long (l.OOOs

km),

narrow (few lOOkm) mountains characterized

thrust faults

and associated folds. These fold-thrust belts

may

develop

continental collision,

which

the

edge

ofa

continental plate

collides uith ancuher continetit. microconiinent.

oceanic

terrane. stich

as an

arc, subduction complex,

aseismic ridge.

Fold-thrust belts comparable

style and scale

those formed

by collision may also develop

in a

back-arc setting, such as the

subandes

South America.

The

sedimentary basins thai

are

complementary

fold-tlirust belts are termed foreland basins

and

can be

separated into peripheral (associated with

collisional fold-thrust belts)

and

retroare (associated with

back-arc fold-tbrust bells) (Dickinson. 1974). Tectonic

732 TECTONIC CONTKtJL CJF SEDIMFiNTATlON

subsidence of foreland basins is primarily driven by loading of

the crust by thrust sheets (Jordan. 1981), but also may be

influenced by downward pull oi' the erust by the subducting

sliib (Royden. 1993) or mantle flow above the subducting slab

(Gumis,'l992).

DeCclles and Giles (1996) divided foreland basins into four

potential depozones that may migrate laterally through time

(Figure T7). The wedge-top depozone represents locally

derived sediment deposited on top of the active, frontal part

of the fold-thrust belt, including piggyback basins (Ori and

Friend. 1984). A common feature of wedge-top strata are

progressive uneonformities. which develop by rotation of

synteetonie sedimentary beds on fault-bend folds (Riba, 1976;

Burbank cf al., 1996). The greatest volume o\' sediment in

foreland basins is deposited in the foredecp depozone. which is

typically 100 300 km wide and receives (he bulk of its sediment

from the fold-thriist belt. Foredeeps of peripheral foreland

basins often begin with deep-water sedimentation {Jiy.sch).

followed by shallow-marine, shoreline, and nonmarine sedi-

mentation

(mohi.Ksi').

Fluvial systems within the Ibredeep may

be transverse or axial, and large drainages may emerge onto

the foredeep at the tips of major thrust faults. Fle.xural bending

of the floor of ihe Ibreland basin may result in a tbrebulge

located on the craton side of the foredeep. Generally hundreds

of kilometers wide and a maximum of a few hundred meters

high, the forebulge is the site of erosion or deposition of thin

sediment derived from the fold-thrust belt or the craton. The

slowly subsiding back-bulge depozone lies eratonward of the

forebulge and contains relatively ihin, fine siliciclastie sediment

derived from the fold-thrust belt and/or the eraton. or is a site

of marine carbonate deposition.

Another type of back-are contractional deformation of

continental erust oeeiirs today in the Sierras Pampeanas of

Argentina and occurred in latest Cretaceous Eocene time in

the Western Interior of the United States (Jordan and

Allmendinger. 1986). This type of deformation is located over

500 1000 km from the active trench and is associated with

subhorizontal subduction. Deformation produces discrete,

basement-co red uplifts tens of kilometers wide and hundreds

of kilometers long bordered by thrust and reverse faults. The

uplifts are separated by intermontane basins the same seale or

larger than the uplifts. In the modern and ancient examples

cited above, the basins are entirely continental, include

alluvial-fan, fluvial, and lacustrine environments, and may

display subsidence histories and sediment distribution similar

depozones of foreland basin

wedge-top

foredeep forebulge

to that of the wedge-top or tbredeep depozones of foreland

basins (Seager ciuL, 1997).

Several other sedmientary basins may be assoeiated with

continental collisions. Deep-sea sediment derived from colli-

sional mountains may be deposited in remnant ocean basins,

which exist between and may eventually be deformed by the

eolliding continents (Graham ci al.. 1975). The best modern

example of a remnant ocean basin is the Bay of Bengal and

assoeiated Bengal deep-sea fan. Continental collisions may

also generate deformational structures hundreds to thousands

of kilometers inboard of the orogen. Such structures may

include extensional basins (impactogens) and strike-slip faults

(Sengorc/a/.. 1978: Tapponnier tvJ/.. 1982).

Sedimentary basins associated with strike-slip faults

Strike-slip faults exist in a variety of difterent tectonic settings

and can be divided into transform and transcurrent types

(Sylvester, 1988). Transform faults define plate boundaries and

penetrate the lithosphere. Fxamples of transform strike-slip

faults inciude those at oeeanic spreading ridges and the San

Andreas fault, which tbrms the boundary between the North

Ameriean and Paeilie plates. In contrast, a transeurrent strike-

slip fault, such as the North Anatolian fault in Turkey, is

confined to the erust and is located within a plate. Most strike-

slip basins are related to bends in the trace of strike-slip faults

or are associated with en echelon strike-slip t^aults (Figure T8).

Extensional fault-bound basins form at releasing bends in

strike-slip faults and tend to be elongate or lens-shaped

(Crowell. 1974). Restraining bends tend to produce thrust

faults and assoeiated transpressional basins, which subside in

response to thrust loading (Nilsen and Sylvester. 1995). In

addition, parallel but offset strike-slip faults may form

extensional stepover basins, such as in the Dead Sea rift

(Aydin and Nur. 1985). and transrotational basins form

among erustal blocks rotated about a vertieal axis between

strike-slip faults (Ingersoll, 1988).

Strike-slip basins may contain a variety of marine and

nonmarine depositionai systems arranged similar to that in rif~t

or foreland basins. Most diagnostic of strike-slip basins are

high sediment accumulation rates, abrupt lateral and vertical

facies changes, lateral migration of the depocenter produeing

stratal shingling, and stratigriiphic variations in provenanee

related to lateral displacement of source terranes (Sylvester.

t9S8;

Nilsen and Sylvester, 1995). Although generally not

restraining

bend

extensional

stepover basin

progressive

unconformilies

Figure T7 Sthemjitic m.ip

,m<.\

cross-section, nol to sctile, ot" the

depozones of a foreland basin, based on DeCelles and Gile^ 11996).

releasing bend

sag ponds

transpressional

basin

Figure T8 Schematic map of three types of sedimentary' basins

associated with en echelon, rij^ht-lateral strike-slip faults. Bali and slick

on down side of normal

faull;

iidld triangles on up side of thrust

f.iull.

TFCTONK" CONTKOI, Or SEI5IMENTATION

733

associated with distinct sedimentary basins, transform faults at

oceanic sprcuding ridges may develop vertical scarps that

supply cravity Hows to the adjacent ocean lloor (Simoiiian and

Gass.

1978)."

Cratonic sedimentary basins

Cratonic basins arc large (,>il)^ knr). generally elliptical basins

thai develop within Uic interior of continents far from active

plate boundaries. Despite relatively slow and erratic sub-

sidence, they may accumulate thousands of meters of sediment

because of histories measured in hundreds of millions of years.

The basin fill consists predoininantly of nonmarine and

shallow-marine line-grained siliciclastics, carbonates, and

evaporites. and unconformities are common. Sloss (1963)

defined si,\ Phanerozoic depositional sequences separated by

continental-scale uneonformities in the cratonic basins of

North America, some of which could be correlated with

cratonic basins in Asia (Sloss. 1972). Sotith America (Soares et

al.. 1978). and Africa (Peters. 1979), Stratigraphic similarities

of cratonic basins in Europe and India were also noted by

Klein and Hsui (1987), Throughout their histories, cratonic

basins may be affected by a variety of teetonic processes,

including various types of thermal and isostatic subsidence.

lower crustal phase changes or melting, and stress related to

plate interactions (Klein. 1995). A comprehensive model to

explain the origin and apparent synehroneity of cratonic

basins remains elusive, although Klein (1995) proposed that

they evolve in response to breakup of supercontinenis.

Summary

(1) Plate tectonics is the primary driving force for crustal

uplift and basin subsidence through the processes oi

faulting, isostatic response to loading and unloading,

lateral flow of the mantle, and ehange in density of the

crust and manlle due to heating and cooling.

(2) Sedimentary basins associated with diverging plates

include continental rifts (e.g.. East African rift), incipient

ocean basins (e.g.. Red Sea), passive continental margins

(e,g.. Atlantic coast of North America), and oceanic basins

created at active spreading ridges (e,g.. Mid Atlantic

Ridge).

(3) Snbduction-related basins may develop on the outer-

trench slope of the subducting slab, in the trench, and on

the overriding plate, including trench-slope basins, forearc

basins, and intra-arc basins. Extensional back-arc basins

may develop in areas of trench roll back or asthenospheric

upwelling .

(4) Continental collision and back-arc compression produce

fold-thrust belts and their complementary foreland basins,

whieh can be divided into wedge-top, foredeep. forebulge,

and back-bulge depozones. An usual type oi back-arc

compression associated with subhori/^ontal subduction

results in discrete uplifts separated by intermontane basins.

Continental collision may produce extensional basins and

strike-slip faults far inboard of the collisional orogen.

(5) Strike-slip faults may define plate boundaries (transform

faults) or reside within a plate (transcurrent fault). Basins

are created at releasing and restraining bends in fault

traces,

between en echelon faults, and by rotation of fault-

bound blocks about a near \ertical axis.

(6) Cratonie basins develop far from active plate margins and

may have long histories of intermittent subsidence and

erosion. Cratonic basins on different continents may

have similar histories, but their tirigin remains poorly

understood.

Greg H, Mack

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Cross-references

Calcite Compensation Depth

Coastal Sedimentary Facies

Deltas and Estuaries

Fan Delta

Cjravity-Driven Mass Flows

Laeiistrine Sedimentation

Oceanic Sediments

Rivers and Alluvial Fans

Slope Sediments

Submarine Fans and Channels

Turbidite

TIDAL FLATS

Definition

Tidal Hats arc sandy-muddy dcpositiona! systems along marine

and estiiarine shores periodically subtnergcd and exposed in the

eourse ofthe

rise

and

ftill ofthe tide ((/.Bates

and

Jackson.

1987).

They differ from tidal beaches in sp;itial extent, topographic

cotnplexity. and faeies differentiation. Perhaps the oldest

reference to tidal flats dales back to the 1st Century AD when

the Rotnan historian Pliny the Klder, after visiting the Wadden

Sea eoast in the year 47 AD. deseribed it in his epochal work

"Naturaiis historia""

as an

area which is

tiitdated by thesea with

forceful eiirrents twice a day and of whieh it was doubtful

whether it formed part ofthe land or the sea.

Physical setting, dimensions, and hydrologic

classification

On a global scale, tidal flat environments are exposed to two

dominant tidal types, a semi-diurnal and a diurnal one which

are separated by two mixed types consisting of a predomi-

nantly setni-diuriial and a predominantly diurnal one. The

loeal tidal type controls the daily frequency and rate of

inundation of a tidal tiat (Allen. 1997). The most prominent

tidal period reeorded in rhythmic tidal deposits is the

tbrtnightly spring-neap cycle

(e,g,.

Visser. 1980). Another

short-term astronotnieal feature influeneing depositional pro-

cesses on tidal Hats is the daily inequality of the tide whieh

results from the inclination of the earth's axis relative to the

plane ofthe ecliptic, and hence the tidal bulge

(e.g.,

De Boer

i't ai. 1989), The daily inequality inereases from the equator

toward higher latitudes. The largest astronomical tides in the

eourse of

year, i,e,. the highest spring and lowest neap tides,

occur at the vernal and atUumnal equinoxes. These are the

periods of maximum tidal energy iliix and henee of greatest

sediment transport potential, ln addition to these short-period

cycles, the tide is also tnodulaied over longer astronomical

periods such as the 18.6 year nodal eyele (Oost

etal.,

1993).

Good examples of rhythmie tidal deposits have been deseribed

from ancient tida! deposits

(e.g..

Archer

eiul..

1991; Williatns,

1991),

A second important t"'aetor affecting depositional processes

on tidal flats is the (idal ranee. In combination with the size of