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

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BEDDING AND INTERNAL STRUCTURES

(d) morphology of set boundaries (erosional vs non-erosional;

tabular vs concave or trough). Such characters can help, for

instance, to distinguish subaqueous from eolian deposits (the

latters show thick to very thick beds and lack ripple scale

lamination) or bedforms grown on fiat bottoms from those

topping larger forms, such as ripples on dunes or dunes on

sandwaves (typical of bars and sandwaves is the hierarchical

organization in cross-bedded sets).

Backset laminae (dipping opposite to the fiow) ean form on

the stoss side of so-called regressive antidunes (Allen, 1982),

but are not usually preserved; the few examples reported in the

literature for ancient deposits, especially turbidites, have been

reinterpreted as regular foreset laminae left by inverted or

reflected flows. Only in pyroclastic strata, in particular those

deposited by turbulent surge flows, are backset laminae

commonly found (Allen, 1982, figures 10-15), probably owing

to the peculiar conditions (temperature/moisture, density

contrast between solid particles and fluid). In general, tractive

lamination is beautifully preserved in these deposits, and

foreset laminae inclined at 20° or less can be observed in dune-

like or dunoid (Ricci Lucchi, 1995a) bedforms. Counter-flow

ripples are sometimes seen at the toe of large subaqueous

structures, for example, tidal sandwaves; they are the product

of recirculation in separation bubbles.

Sets and eosets of cross-laminae can show a climbing

attitude of the corresponding bedforms (more commonly

ripples, but also subaqueous dunes), with a variable angle of

climb indicated by the stoss sides. This is a clear evidence of

traction combined with fallout, and the cross-bedding style has

been called ripple-drift cross-lamination (Jopling and Walker,

1968);

three varieties (types A, B, and C) have been

distinguished on the basis of relative amounts of erosion and

deposition on the two sides of the bedforms. Climbing cross-

bedding normally derives from deeaying currents and shows

bedform profiles preserved by a mud cover.

Low angle laminae occur in sets usually separated by low

angle erosional surfaces, both planar and curved; resulting

laminasets are wedge-shaped to sigmoidal. It is not easy,

especially in limited outcrops, to decide whether they are

related to a current, an oscillatory motion or a combination of

the two (see below). Low angle lamination has been described

in "proximal" sandy turbidites (Mutti and Ricci-Lucchi, 1972,

facies B) and fluvial deposits (couplets of planar laminated

sand and gravel are a common feature of sheetfiood deposits

within alluvial plain/alluvial fan depositional systems).

Oscillatory and combined flows

Two kinds of "short term" (non-geological) periodicity are

recorded in sediments, that of wind waves and that of tidal

waves.

Wind-generated waves may form within any body of water

and exert traction below a critical depth referred to as the wave

base.

Wave orbits increase their ellipticity toward the bottom,

where they culminate as straight lines representing to-and-fro

motion. This motion results in grain movement by rolling and

saltation on an initially plane sand bottom. Wave-formed

ripples vary greatly in size, with /; (height) commonly ranging

between 0,003 m and 0,25 m, and I (wavelength) between

0,009 m and 2 m, Their ripple index varies between 4 and 13,

Wave ripples are characterized by a symmetrical profile

near wave base and tend to become more asymmetrical

approaching the shoreline (steeper side facing landward). In

2D-sections, they display diagnostic internal features, such as

inconsistency between ripple morphology and internal struc-

ture,

structural dissimilarity between adjacent sets, presence of

chevron and bundled upbuildings of foreset laminae, and

irregular and undulating lower set boundaries (wave-truncated

ripples of Campbell, 1967; de Raaf era/,, 1977), In plan view,

wave ripples display straight to slightly sinuous crests, with

characteristic bifurcations, that are uncommon in current

ripples. Wave ripples are typically found within lower shore-

face sands, but have also high preservation potential in

shallow-water environments affected by oscillatory waves,

such as lagoons and large lakes.

In intertidal environments, partial erosion of wave ripples at

low tide conditions results in characteristic truncation of their

tops.

Actually, in very shallow water depths (foreshore

deposits, swash/backwash zone) upper fiow regime conditions

prevail and planar surfaces, both erosional and depositional,

are the dominant feature. Foreshore sands are thus character-

ized by the stacking of gently seaward-dipping sets of parallel

laminated sand, formed during fair-weather conditions,

separated by erosional surfaces related to storm events

(wedge-shaped, low-angle cross-stratification).

While wave ripples are generated by fair-weather waves,

storm waves tend to produce a structure of similar or greater

size,

called hutnmoeky cross-stratification (HCS) by Harms et

al., (1975), More precisely, HCS is interpreted as the result

of combined, waning unidirectional and oscillatory flows, but

anyway a diagnostic feature of storm-dominated processes.

Recently, however, sedimentary structures very similar to HCS

have also been observed to characterize flood-dominated fan

delta and river delta systems (Mutti etal., 1996),

HCS consists of convex-up, large amplitude (1-5 m) and

low-relief (10-50

cm),

irregular domal features (hummocks),

separated by broad troughs (swales). Gently undulating

(5-15°) lamination is the most common internal feature,

displaying a general concordance with the basal erosion

surfaces and systematic lateral variations of laminae thickness

(pinching and swelling of individual laminae). Where concave-

up sets are well preserved, swaley cross-stratification (SCS)

is developed (Leekie and Walker, 1982), HCS is generally

restricted to coarse silt to fine sand, and is particularly

abundant within lower shorefaee deposits, in association with

symmetrical wave ripples, SCS may be abundant in shallower

(upper shorefaee) and coarser grained deposits. Storm events

in the offshore-transition area, above storm wave base,

characteristically develop idealized sequences (Dott and

Bourgeois, 1982) that rest on scoured surfaces, passing upward

into HCS, and capped by fiat lamination and symmetrical

ripples at top of the sandy layers. The sequences are capped by

bioturbated mudstone intervals.

In flood-generated delta-front sandstone lobes, large-scale

HCS is developed within 3-15 m thick sandstone packets,

separated by highly fossiliferous and bioturbated muddier

deposits.

Tidal structures and bedding

Tidal eurrents regularly change direction from the flood tide

current, which flows landward between the low and high tide,

and the ebb tide current, which flows in the opposite direction

as the water level returns to low tide. Bipolar cross-bedding,

forming the characteristic herringbone cross-stratification,

is the typical, expression of alternating tidal currents in

56 BEDDING AND INTERNAL STRUCTURES

high-energy, subtidal settings. Herringbone cross-stratifica-

tion, however, is not very common in the rock record, because

of the fact that one half of the current generally is much

stronger than the other, and ebb and flood currents may follow

different paths, Bedforms migrating in one direction, under the

influence of the dominant current, can be only slightly

modified by the subordinate current. Diurnal clianges in

bedform migration controlled by asymmetrical tidal currents

may thus result in discrete sets of unidirectional cross-strata

(sand foresets form on the lee side of large bedforms),

separated by gently sloping erosion surfaces, which are termed

reactivation surfaces. This pattern of alternating undirectional

cross beds and reactivation surfaces is characteristic of large,

linear bedforms (sand waves), with wavelengths of tens to

hundreds of meters and heights exceeding

m, which are

commonly observed in open-shelf environments.

Regular changes in magnitude and direction of tidal

currents over the tidal cycle allow transport and deposition

of sand at some stages and fallout of mud from suspension

during periods of slack waters. Tidal structures resulting from

this rhythmic alternation of traction and fallout processes are

termed tidal rhythmites (Tessier, 1993) and generally char-

acterize intertidal fiats. Tidal rhythmites include sand layers,

showing small-scale current ripples, alternating with mud

drapes, refiecting deposition during stillstand phases.

Distinctive sedimentary structures of intertidal fiat environ-

ments include fiaser bedding, wavy bedding and lenticular

bedding (Reineck and Wunderlieh, 1968), Flaser bedding (see

Flaser) is characterized by isolated thin drapes of mud

amongst a cross-laminated sandy unit. In contrast, lenticular

bedding is made up of isolated sand ripples within a mud-

dominated deposit. Wavy bedding is intermediate between the

two,

consisting of approximately equal proportion of sand

and mud.

In higher-energy, subtidal settings, a characteristic feature of

tidal sedimentation at the scale of diurnal, ebb-flood tidal

cycles is double mud drapes, formed during the two slack

water stages of the tidal cycle. Double mud drapes, commonly

a few mm thick, typically separate unidirectional, bipolar cross

beds showing current reversals. These diagnostic tidal struc-

tures,

forming a characteristic sigmoidal stratification, are

referred to as tidal bundles (Visser, 1980), and are commonly

developed in subtidal (estuarine) channels. In intertidal

settings, the tidal bundles include just one mud drape. Cyclic

variations in strength of tidal currents at the scale of neap-

spring cycles result in cyclic variations in thickness of tidal

bundles, forming characteristic tidal bundle sequences (Yang

and Nio, 1985),

Internal structures in episodic (catastrophic) deposits

Episodic events include tsunami waves, seismic shocks, violent

hurricanes, catastrophic fiash floods, and sediment gravity-

flows. Resulting beds are lenticular to tabular and internally

disorganized (random arrangement of particles) to organized.

Organization means size grading (graded bedding), grading in

the lower part plus lamination in the upper portion of

sandstone, or thoroughly laminated sandstone, Beds are

commonly capped by mud, but can be separated by erosion

surfaces ("amalgamated"). The different parts of organized

beds are called intervals or divisions.

Gravity flows are moving dispersions (mixtures of solid

particles and water) classified on the basis of the different

proportions of the two phases and the grain-support mechan-

isms.

Following Middleton, Lowe, Postma and Mutti (see

Mutti, 1992, for references), different flow types can be

identified. In cohesive debris flows, the fluid has pseudoplastie

characteristics and grain transport is a result of matrix

strength, which is in many instances adequate to support large

boulders. Coarser grains move to regions of lower shear, and

are rafted on the top or edge of flow, resulting in inverse

grading. In grain flows, shear stress is transmitted through the

flow by a dispersive pressure resulting from intergranular

collisions, Frictional freezing, which occurs when the applied

shear can no longer overcome the flow strength, is the

dominant depositional mechanism from both cohesive debris

flows and grain flows.

Debris flow deposits are characteristically very poorly

sorted and lack almost any internal structure (the preponder-

ance of plug flow hampers the development of shear fabric or

sorting). When large clasts are present, they are randomly

distributed (the bed has a chaotic aspect) or concentrated near

the top, locally protruding, A rough development of inverse

grading can thus be observed throughout the bed or at the

base,

combined with lack of significant basal scours.

Debris flows show marked downslope transformations, as

they decelerate and mix with ambient waters; such transforma-

tions are reflected in internal fabric and structures, A high-

density, inertia-driven mixture of sediment and water that

moves downslope under conditions of high pore-fluid pressure

is termed hypereoneentrated flow. The typical structureless

debris flow deposits may thus be replaced in a downcurrent

direction by deposits from hyperconcentrated flows; these are

characterized by deep basal scours, large rip-up mudstone

clasts,

a coarser grained matrix, and tendency of larger clasts

to concentrate in the lower part of the bed. Deposits from

hyperconcentrated flows may also include evidence for

traetional bedforms.

Traction carpets, resulting from a combination of fluid

turbulence, hindered settling, and dispersive pressure produced

by grain collision, may develop in the basal portion of high-

density flows, and are a relatively common feature of grain

flows. Traction carpets, which are commonly followed by

en-masse deposition, display characteristic, inversely graded

stratification bands and show upward transition to ungraded,

crudely graded or well-graded beds.

At modest levels of sediment concentration, the flow is

eohesionless and may be internally turbulent. In fluidized and

liquefied flows, escaping pore fluid provides a full or partial

supporting mechanism, respectively. In low-density turbidity

currents, particles are supported by the flow turbulence, and

sediment deposition occurs through grain settling and traction-

plus-fallout processes. The typical pattern of grain size

change within a bed which is formed from a decelerating flow

is referred to as normal grading. Larger particles achieve a

higher terminal velocity and settle out of suspension faster

than smaller grains. Two types of grading have been

distinguished in turbidites (Middleton, 1967): distribution

and coarse-tail. The first type is shown by all size classes, the

second one by coarser particles only. Grading may pass

upward into, and/or be superimposed by, tractive lamination,

as observed since the thirties of last century by geologists in the

Apennines looking for way-up criteria in zone of vertical and

overturned beds. Later Bouma (1962) formalized a vertical

sequence of internal structures, the now familiar "Bouma

sequence".

BEDDING AND INTERNAL STRUCTURES

The Bouma sequence includes, from base to top: normally

graded, structureless medium to fine sand, with frequent

dewatering pipes and dish-and-pillar structures (division a);

planar laminated to undulating fine sand, with parting

lineations and inversely graded stratification (division b);

small-scale cross-laminated very fine sand, with abundant

climbing ripples (division c); current-laminated coarse silts and

interlarninated silts and mud (division d). The succession is

capped by a homogeneous and massive mudstone (division e),

actually a pelitic bed or parting. The transition from

thoroughly current-laminated sand to structureless mud is

more or less gradual and marks the passage from traction-

plus-fallout processes to pure grain settling from suspension.

In more proximal turbidites, richer in sand, Lowe (1982)

defined different varieties of structure sequences, but they seem

to have a less universal application.

Sharp-based layers with distinctive internal fining-upward

tendencies, being the expression of decelerating, waning

turbulent flows, are not exclusive of submarine gravity fiows,

but may be commonly encountered in different depositional

environments, such as lakes, river deltas (hyperpycnal plumes

related to river fioods) and shallow-marine realms (storm-

generated currents).

Intraformational mud clasts, referred to usually as clay chips

and a variety of other terms, and deriving from intrabasinal

sources, are a common feature of debris flow, hyperconcen-

trated fiow and structureless to graded portions of turbidity

flow deposits. They are typically embedded in sand and occur:

a) concentrated at or near the base of the bed; b) scattered

throughout; c) aligned (and relatively sorted) in the uppermost

structureless or graded sandstone, often at the transition with

tractive laminae. These situations reflect different distances of

travel within the flow and a marked tendence to buoyancy.

This should be considered a fabric more than a structure; the

same applies to clast imbrication (see Imbrication), showing in

both massive and selective (e.g., beach) deposits. Flattened

clasts are inclined to the horizontal, so that the plane formed

by the long and intermediate axis of the clasts dips up-current,

according to the most stable position. Imbrication represents a

very useful paleocurrent indicator and is a characteristic

feature of gravel-bed rivers and gravel beaches (where dip is

seaward). It indicates a transport of pebbles in temporary, or

intermittent, suspension.

Internal scours may be found in gravity flow and other

episodic deposits, owing to fiow pulsations or bottom

irregularities during deposition; care should be taken, how-

ever, not to mistake such local features with parts of larger

erosion or amalgamation surfaces separating distinct beds

(events).

Internal structures and facies sequences

The "Bouma sequence" just described represents an example

of a genetic package, or association, of structures, that is, a

systematic vertical succession of structures consistent with a

specific mechanism evolving in time (decaying suspension

fiow). In this case, the time is short (hours, days) being the

duration of a single depositional event of episodic nature.

Similar genetic packages can be found in multilayer sedimen-

tary bodies, built up by several events, and in normal or

selective deposits. The best known case is that of the fining-

upward sequence displayed by fluvial and tidal point bars

formed by lateral accretion (Allen, 1982, his figures 2-23).

Notice that Allen uses the term structure not only for the

cross-bedding, parallel lamination, etc., inside the body, but

also for its overall geometry ("lateral accretion structure").

Alternatively the same object can be described as a facies

sequence, a facies association or an architectural element.

The following description summarizes that of Allen (1982):

in tidal fiats, mud-sand interbedded points bars form the inner

bank of small meandering gullies. They are rarely thicker than

1.5 m and consist of curved, sigmoidal layers, from parallel-

laminated to current rippled, both ebb and flood directions

often being preserved. The base of the bar is an irregular

erosional surface, generally overlain by gravel of bivalve and

other shells mixed with pebbles of muddy sediment derived

from cut banks.

River point bars, in intermediate to large sand-bedded

rivers,

where the secondary [transversal] fiow is fully devel-

oped, have also a basal, diachronous erosion surface, with

flutings, pot-holes and other scours. Sediments in the lateral

accretion deposit fine upward in general but interfinger in

detail, and Internal sedimentary structures appear upward in

order of declining flow power, an arrangement which refiects

the distribution of bedforms over the bar surface, in turn

revealing the inward decline of current strength. Curved

bedding surfaces, generally sigmoidal in radial profile, define

the lateral accretion structure

[

= master bedding].

Structures related to deformational processes

Primary sedimentary structures are subjected to disturbances

by fluid movement, compaction and gravitational instability if

the sediment remains soft and deformable. Soft sediment defor-

mation is the general term for changes to the fabric and

layering of a bed of unconsolidated sediment. It affects most

commonly multilayer deposits, especially sand-mud alterna-

tions,

and can be diffused in the whole packet or concentrated

along bedding planes or within individual beds. No systematic

classification exists, apart from a broad subdivision into

structures characterized by vertical movement and structures

showing a lateral component of motion (gravitational or

current shearing). For nomenclature, see Table B2.

Concerning the first group, trapping of water induced by

rapidly deposited sediment may lead to disruption of previous

structures within a bed (more commonly sand) or a bedset.

This may occur due to shocks by earthquakes or oscillations in

water pressure, which may temporarily liquefy the sediment.

Beside liquefaction, deformation can result from sediment

contraction (e.g., desiccation) or dilatation, unequal loading,

upward injection of fluid or paste-like material, downward

injection in fractures, compaction, downslope movements,

current or ice drag or combined effects.

Sedimentary structures formed by liquefaction, injection,

etc.

are often associated with deposits from high-density

turbidity currents and subaqueous slumps; those that can be

related to earthquakes may be called seismites, but this

attribution is often conjectural.

Sand volcanoes are formed by the violent expulsion from

the subsurface of sand and water mixtures. Dewatering (fluid-

escape) structures result from forceful expulsion of water from

a consolidating fiuidized bed, due to loading or a shock. The

water moves upward, disrupting overlying layers and forming

dish-and-pillar structures. Dishes are subhorizontal, gently

concave-upward clay or organic-rich laminae, deformed at

their margins by upward fiow. Pillars are near-vertical

BEDDING AND INTERNAL STRUCTURES

Table B2 Soft-sediment deformation structures reported in modern and ancient sediments. For references, see Ricci Lucchi (1995b)

Modern

Ancient

Liquefaction structures

Sand boil

Sand volcano, sand-vented volcano (Obermeier etal., 1986)

Sand blow (Sieh, 1978; Doig, 1990)

Sand slough (Fuller, 1914; Rascoe, 1975)

Patterned ground (Washburn, 1950)

Mima mounds (Fuller,1914; Arkley and Brown, 1954; Berg,l990)*

Patterned ground (Washburn, 1950)

Ball-and-pillow structure (Smith,1961; Pettijohn and Potter, 1964)

Load structures

Flame structure (Cooper, 1943)

Injection structures, diapiric structures

Flow rolls (Pepper etal., 1954)

Pseudonodule (Macar, 1948)

Convolute lamination (Kuenen, 1953)

Cycloid (Hempton and Dewey, 1983)

Dish structure (Wentworth, 1967; Lowe and Lo Piccolo, 1974)

Pillar structure, or dish-and-pillar structure (Lowe and

Lo Piccolo, 1974)

Water escape structures

Sand volcano

Sandstone dike

Sandstone pipe

Clastic (sedimentary) sill

Clastic (sedimentary) dike

Slump ball (Kuenen, 1948)

Slump structures

Boudinage

Pull-apart (Natland and Kuenen, 1950)

tee-shear (ice-push) structures

Glaciotectonic structures

Drag fold

Hydroplastic deformation

Wrinkle marks (Hantzschel and Reineck, 1968; Allen, 1985)

* Fuller (1912) regarded this structure as aseismic, Berg (1990) reinterpreted it as seismic

dewatering pipes, interrupting the dishes and made up of

structureless sand. Dishes are found in both massive and

crudely laminated sandstones and are related to horizontal

movements of water or liquidized sand. In the dark laminae,

preexisting (semipermeable membranes?) or newly formed,

fine particles are concentrated by elutriation.

Convolute laminations are deformational structures devel-

oped in rapidly deposited fine-grained sands and silts, and

commonly associated to water-escape structures and climbing

ripple cross-lamination in turbidites and river flood deposits.

They are either symmetrical about a vertical plane or leaning

and asymmetrical, and usually exhibit narrow vertical

upturned laminae, often truncated at the top, separated by

broader synclinal downfolds, with wavelengths of a few

centimeters or decimeters. Some convolutions are associated

with water-escape dishes, pillars, etc.

Many explanations have been offered for the origin of this

structure; the asymmetrical varieties involve shearing, prob-

ably caused by an overflowing dense current. Downslope

movement (slumping) seems to be excluded by the

common occurrence in basinai sediments, especially turbidites,

and the dying out of the folds both upward and downward

into flat surfaces, without any evidence of decollement.

Liquefaction or hydroplasticity of some or all laminae

involved (which contain variable amounts of organic material

or clay) and density instability are the most invoked

mechanisms. Timing of formation can be during bed deposi-

tion (syndepositional type of Allen, 1982), just before or

immediately after deposition ceases (metadepositional), or

definitely after (postdepositional).

Load casts are the commonest type of deformation of sand-

mud interfaces and "represent the instability of discontinuous

layered systems" (Allen, 1982) due to inverted density

gradients (Dzulynski, 1966). The structure consists of lobes

of sagging sand alternating laterally with sharp-crested diapers

("flames"), projecting from a mud that was quite fluid when it

was loaded. Load casts can be found at various scales ranging

from millimeters to decameters (in outcrops).

Similar to load casts, but more deformed and complex in

shape and internal structure are pseudonodules or ball-and-

pillow structures. In this case, sand is pierced or completely

disrupted and detached from the "mother bed" and embedded

in the underlying mud (something analogous to stoping in

magmas). Internal laminae follow the curvature of the outer

surfaces, indicating that they are formed by viscous shearing

within the ball or pillow during its descent. The most plausible

explanation is again the physical state of the materials

(Rayleigh-Taylor instability of sand-mud interbeds) combined

with an external perturbation, that is, seismic shocks. Down-

slope movements are not necessarily implied, even though the

pillows may end up in slumped masses.

Deformations in slump sheets is best observed where sliding

materials are heterolithic and/or in different state of compac-

tion and cohesion. A variety of geometries is thus displayed,

from disharmonic folds, cascade folds, box folds, concentric

folds,

isolated hinges, shear planes, slabs and blocks, breccia-

tion and even, in most conspicuous slump sheets (tens of

meters thick), cleavage in mudstones (Ricci Lucchi, 1995b).

Sand volcanoes, other fiuid-escape structures and burrows may

occur at the top. Axial surfaces vary from steeply inclined to

flat-lying, and axes tend to be parallel to depositional strike,

but with a wide dispersion of orientation.

Structures related

advancing

ice

or sand bodies are shown by

layered muds in the form of folding, crumpling and contortion.

BEDSET

AND

LAMINASET 59

at places extremely "baroque" (complicated). Antiform folds

can be arranged en

echelon,

more or less parallel to the edge of

the encroaching body, and may be associated with shear

planes. These structures are the combined effect of differential

ioading and drag.

Sedimentary dykes can be produced at various scales by

both forceful injection of fluid or hydroplastic sediment from

below and infilling of open cracks from above. Ptygmatic folds

can develop in dykes owing to subsequent compaction or to

increasing resistance encountered by the injection.

Smaller injected dykes are indistinguishable from the pillars

associated to dishes in sandstones; larger structures are called

pseudodiapirs, mudlumps, etc., and originate in multilayered

successions rich in mud. Infilled dykes range from desiccation

cracks to large fissures and fractures formed in lithified

sediments, especially limestone, subjected to extensional

stresses ("neptunian dikes").

Deformed eross-hedding has various aspects. It can affect

either the topmost part of foreset laminae, overturning or

disharmonically folding them, or the mid-lower portions,

where small folds, discordant slabs, small normal (gravity)

faults or patches of vaguely laminated or structureless

sediment appear (Allen, 1982). Fluid drag on subaqueous or

subaerial, water-saturated sediment is generally held respon-

sible for the first case, whereas in the other, some form of

gravity induced downslope slip is more plausible. The lee side

of eolian dunes, for example, is prone to sliding when wet.

Detachment and sliding of slightly cohesive sand laminasets

can occur without internal deformation.

Franco Ricci-Lucchi and Alessandro Amorosi

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Mutti,

E., and

Ricci-Lucchi,

F., 1972. Le

torbiditi dellAppennino

Settentrionale: introduzione allanalisi

facies. Memorie Societh

Geologica

Italiana, 11(2): 161-199.

Mutti,

E., 1992.

Turhidite Sandstones.

San

Donato Milanese: AGIP,

275

pp.

Mutti,

E.,

Davoli,

G.,

Tinterri,

R., and

Zavala,

C, 1996. The

importance

ancient fluvio-deltaic systems dominated

catastrophic flooding

tectonically active basins. Memorie

Scienze

Geologiche

(Padova, Italy), 48:

233-291.

Reineck,

H.E., and

Wunderlich,

F., 1968. Zur

Unterscheidung

von

asymmetrischen Oszillationsrippeln

und

Schomungsripplen.

Senckenbergiana Lethaea, 49: 321-345.

Ricci Lucchi,

F.,

1995a. Sedimentographica.

Photographic Atlas

Sedimentary Structures. Columbia University Press.

Ricci Lucchi,

F.,

1995b. Sedimentological Indicators of Paleoseismicity.

Association

Engineering Geolgists, Special Publication,

7-18.

Tessier,

B., 1993.

Upper intertidal rhythmites

in the

Mont-Saint-

Michel

Bay (NW

France); perspectives

for

paleoreconstruction.

Marine Geology, 110: 355-367.

Visser,

M.J.,

1980. Neap-spring cycles reflected

Holocene subtidal

large-scale bedform deposits:

preliminary note. Geology,

543-546.

Yang, C.-S.,

and Nio,

S.-D.,

1985.

The

estimation

paleohydrody-

namic processes from subtidal deposits using time series analysis

methods. Sedimentology, 32: 41-58.

Cross-references

Accretion, Lateral

and

Vertical

Bedset

and

Laminaset

Clastic (Neptunian) Dykes

and

Sills

Convolute Lamination

Cross-Stratification

Deformation

Sediments

Dish Structure

Fiaser

Fluid Escape Structures

Grading, Graded Bedding

Grain FJow

Grain Size

and

Shape

Gravity-Driven Mass Flows

Hummocky

and

Swaley Cross-Stratification

Imbrication

and

Flow-Oriented Clasts

Liquefaction

and

Fluidization

Pillar Structure

Planar

and

Parallel Lamination

Sediment Transport

Tides

Sediment Transport

Unidirectional Water Flows

Sediment Transport

Waves

Storm Deposits

Surface Forms

Tides

and

Tidal Rhythmites

Turbidites

BEDSET

AND

LAMINASET

According to McKee and Weir (1953), a bed is a sedimentary

stratum greater than

mm thick, whereas a lamina is less

than

mm thick. The divide at 10 mm is arbitrary, and not

60 FiFDSETAND LAMINASET

sedimentary log

McKee & Weir

(1953)

Campbell (1967).

Reineck & Singh

(1973)

Blatt, Middleton

& Murray (1980)

set set

cosel

set set

coset

composite set

set set

coset

composite set

bed bed

bedset

bed bed

bedset

bed

bedset

division

layer

bed

division

layer

sets of strata

or cross strata

(beds or laminae)

beds may have

internal laminae,

Laminaset=bed

Figure B12 Comparison of terminology for describing stratification (from

Bridgf,

199.'i).

point bai

channel bell

groups of large-scale inclined strata sets (sandstone

body):

channel belt

sei of large-scale inclined strata (storey); channel bar and fill

large-scale inclined stratum:

accretion on channel bar

simple conapound

medium-sciile cross strata: dunes

medium-scale cross strata

superimposed on simple

large-scale inclined strata:

dunes on unit bar

Figure B13 Proposed terminology for the case of superimposed scales of strata, using river channel deposits as an example (modified from

Bridge, 1993).

BENTONiTES AND TONSTEINS

based on objective statistical analysis ofstratal thickness, nor

on any genetic implication. Beds and laminae are recognized

on the basis of changes in sediment texture and/or composition

within or between them. Beds and laminae generally occur in

sets

{hccl.sci.s.

laminasci.s) that arc distinctive in terms of the

orientation, texture and composition oi beds or laminae

within the set (Figure BI2). If the beds or laminae within a

set are inclined at more than a few degrees relative to the set

boundaries, the beds or laminae are given the prefix cross or

inclined {c.^., set of cross laminae, or cross-lamina set). Bedsets

or laminascts also occur in sets. For example, a set of similar

cross-lamina sets or cross-bed sets is referred to as a simple

set or cosvi (Figure B12). according to McKce and Weii"

(1953).

A set or different types ol" bedsets or laminasets is

called a ciimpcsih-scl (Figure B12).

This terminology for defining beds, laminae, bedsets. and

laminasets is widely used but is not without its problems

and detractors. Modifications to the terminology shown in

Figure B12 suggest that there was a desire to use the term bed

for a larger scale of stratum than that defined by McK.ee and

Weir (1953). The term bed may therefore refer to a set of strata

rather than a single stratum with no internal subdivision.

Another problem with McKec and Weirs terminology is that

botli kiininae and beds may occur in the same set of strata.

This problem arises because of the arbitrary division of strata

into beds and laminae. A single strataset containing beds and

laminae could logically be called a composite set. but this term

has already been used for a set of different sets rather than a

set of differcnl strata. Perhaps the greatest difficulty with all of

the different methods for defining strata and sets of strata is

that there is no consistent use of terms for referring to the

many different superimposed scales of strata and sets of

strata in sedimentary basins. The terms used (or the smaller

stratasels (bedset, lamiiiaset) have not been applied to the

larger-scale stratasets. Instead, poorly defined, inexplicit terms

such a

.siorvy.

archilecluralelement and

para.secjitencc

have been

used.

Bridge (1993) argued that, in order to describe hierarchies of

different-scale strata, ii is desirable to use a reasonably stnall

nninber of explicit lerms consistently irrespective of scale, and

to use qualifying terms to describe relative scale of strata.

Figure B13 shows an example of this approach applied lo river

deposits, and includes interpretation as well as descriptive

terminology. An alternative method of describing hierarchies

of strata ol" different scale is by numerically ordering their

bounding surt"accs (e.g.. Brookfield. 1977: Allen, l983rMiall,

1996).

As discussed by Bridge (1993), this approach is difficult

to use in practice.

John S. Brid"c

Bibliography

Alloii, .I.K.t... 19S3. Studies in Iliivialilc scdiiiicntiilioii: bars, bar-

compfc\es ;iiid sandstone slu-cl^ (low-sinuosity braided streams) in

the Brounsloncs (Lower Devonian). Wclsli Borders. Scdiniciuiirv

(k-ologv. 33:

^.•^T-:^,!.

Blall. l-l,. Middleton. G,V.. and Murray. R.. I'JKO. OrigmofScdiim-ii-

hiry Rocks. 2nd cdn.. Prcniicc-Hall.

Bridge. J.S.. 1993. Descriplion and iniLTpreliilion ol'lluvial deposils: a

ciiticiil perspective. Sc(liiiicnloloi;y. 40: 801 81U.

Brookfield. ME.. 1977. The origin ol' bounding surfaces in ancient

eolian sandstones.

Scdimciildlot;]-.

24: 303 332.

Campbell, C,V,, 1967. Laminae, lamina sel, bed and bedsei. Scdinien-

loh'f-y. 8: 7-26.

McKee. E.D.. and Weir. G.W.. 1953. Terminology for

rai ilka lion

iind cross-siratificaiion in sedimentary roeks. Bullcuii of ihc Gcolo-

^Uul Socictx of America. 64: 381 -390.

Miall. A.D.. 1996. The Gcohs^y of Fliniul

Dcpo.Kiis.

New York:

Springer-Veriag.

Reineck. H,E.. and Singh. I,B.. 1973. DvposiiioiuilSctliiiwniaryEinir-

oiiiiu-ius. Berlin; Springer,

Cross-references

Bedding and Imernal Slrticlures

C.'ross-siralilicalion

Grading. Graded Bedding

BENTONITES AND TONSTEINS

Definition

Bentoniies are thin, widespread elay-altercd ash layers that are

commonly dominated by smectite (Fisher and Schminke, 19f^4).

This isthedefinition used by sedimenUilogists. Knight (1898) first

used ihe name benionite for an unusual clay

the Ben ton Shale of

Cretaceous age in Wyoming. The absorbent properties and

swelling nature of thisand similar clays were later shown lo be due

to the presenee of the smectite group of minerals ((/,v,) of which

montmorillonite is the most common. In the definition of

bentonite the emphasis on the origin is generally attributed to

Rossand Shannon

(1926).

They noted that the clay was

product

of the "alteration of a glassy igneous material, tisiially a tuff or

volcanic ash" and the resultant clay "is usually the mineral

moniniorillonite. hut less often beidcllite" (a member of the

smeclite group). Smectite-rich bentonites are valuable raw

materials and some authors would restrict the bentonite

definition to smectite-rich deposits irrespective of origin (Grim

and Guvcn, 1978; Roen and Hosterman. 1982). This definition

cannot be reconciled with that used by sedimentologists and it has

to be accepted that both definitions will continue to be used, albeit

f"or different purposes.

Mesozoic bentonites usually consist of

smectite.

As tempera-

ture rises diu'ing burial diagenesis smectite layers are progres-

sively converted into itlite. via a mi.\ed-layer illitc-smectite.

These beds are known as K-bentonites because of the

introduction ol" potassium in the illitization reaction. Most

Paleozoic clay-altered volcanic ashes are of this type. In coal-

bearing strata o\' all ages the volcanic ash usually alters to

kaolinite forming thin beds known as lonstcins. Only since the

1980s has the volcanic origin been generally accepted.

Fisher and Schminckc (1984) suggest thai the nomenclature

of bentonites should reflect the dominant clay mineral species

with smeclite (s)-bentonite. kaolinite (k)-bentonitc and illite

(i)-bcntonite. K-bentonite is a well-established term and Huff

(1983) has made a good case for its retention. Likewise the

term tonstein is also well-established and will continue to be

used by coal geologists. Possibly smeetite-bentonite should be

adopted as this would eliminate potential confusion between

62 BENTONITES AND TONSTEINS

bentonite as a general term for the whole group and a specitic

member within the group.

Recognition of bentonites

Field characteristics

Bentonites typically occur as thin beds with a wide lateral

extent of up to hundreds of kilometers. The contacts with the

enclosing sediment are nearly always sharp. If the material is

unweathered internal lamination and graded bedding may be

visible. The expansion of smectite during weathering causes

preferential loss of bentonites in surface exposures. Tonsteins

on the other hand may be more durable than the host

sediments because kaoiinite does not expand and there is an

interlocking texture. Many tonsteins have the appearance of

flint clays, but others are more granular. Bentonites are mainly

associated with shales and coals because a low energy

depositional environment is required for their preservation.

Ashfalls may incorporate minor amounts of organic matter,

but the content is significantly less than in the enclosing

sediments and hence there is a usually a color contrast.

Mineralogy

The clay mineralogy of bentonites differs significantly from

that of the associated sediments. The normal detrital clay

assemblage of kaoiinite, illite and chlorite is absent and in

tonsteins the kaoiinite is well-ordered in contrast to most

detrital kaolinites. Bentonites are dominated by one clay

mineral to the extent that many are monomineralic and in this

they are exceptional.

The clay minerals form from unstable volcanic glass and to

a lesser extent from some of the pyrogenic minerals. Many of

the latter are relatively stable and survive the ash alteration.

These minerals are diagnostic ofthe volcanic origin and also of

the original ash composition. Quartz and biotite are common,

sanidine and high temperature plagioclase feldspars less so and

olivine, pyroxene and amphibole are rare. The relative

abundances largely reflect the dominance of acid volcanics,

but with weathering stability superimposed. In bentonites

there is a restricted heavy mineral suite consisting mainly of

zircon, apatite and ilmenite/magnetite. Weaver (1963) empha-

sized the interpretative value of this restricted suite, which is

unlike that normally encountered in detrital sediments.

As the thickness of the volcanic ash decreases away from the

eruptive source so the grain size of the particles decreases and

their proportions change. Minerals are generally most

abundant between 2 mm and 1/16 mm and are practically

absent below 10 microns, whereas glass shards can be much

smaller (Fisher and Schmincke, 1984). These features are

inherited by bentonites. There is also mineral fractionation in

air-fall ashes and in bentonites a noteworthy result of this

eolian fractionation is the enrichment of biotite at the base of

some units (Diessel, 1985; Huff and Morgan, 1990).

Textures

In the magma the pyrogenic minerals develop relatively freely

resulting in euhedral crystals. These are found in bentonites

and also high temperature forms such as beta quartz crystals.

During explosive volcanism crystals are shattered and quartz

splinters are very common. As would be anticipated minerals

with well-developed cleavages also fragment. There is therefore

a contrast with the shape of grains in normal detrital

sediments.

Vitric particles are the other main component of volcanic

ashes.

Cuspate and platy shards result from the shattering of

bubbles in silicic magmas and the characteristic shape may

survive the alteration to clay. Clay-altered pumice shards may

retain the shape but lose the internal structure and conse-

quently are more difficult to identify positively.

Composition

The major element composition of bentonites essentially

reflects the diagenetic conditions and the resulting clay mineral

proportions rather than the magmatic origin. There are rare

instances where vitric particles have been preserved, usually by

an early diagenetic cement. Glass inclusions in pyrogenic

minerals are more common and are sufficiently large to be

analyzed by an electron microprobe, as in the work of Lyons

etal. (1993).

Not all elements are mobile in the alteration process and the

immobile elements are concentrated. The radioactivity of

many bentonites in borehole logs drew attention many years

ago to the high concentrations of Th and U. More recently it

has been appreciated that classification schemes for altered

volcanic ashes using immobile elements (Winchester and

Floyd, 1977) also apply to bentonites (Lyons et al., 1993).

The use of these elements confirms the volcanic origin, but

more importantly provides information on the original ash

composition and the tectonic setting of the vulcanicity.

Classification of tuffs and bentonites

Tuffs are classified on the glass-mineral proportions and range

from vitric tuffs to crystal tuffs (Pettijohn, 1975). As some

minerals and textures survive the alteration process such a

classification should be applicable to bentonites. The petro-

graphy of tonsteins has been studied in the greatest detail

because of their essential role in coalfield correlations and also

tonsteins represent the worst case scenario in terms of the

extent of alteration. Most authors use SchuUers classification

(see Bohor and Triplehorn, 1993), which is kristal tonstein

(crystals of kaoiinite), graupen tonstein (grains of kaoiinite),

psudomorphosen tonstein (kaoiinite pseudomorphs after

silicates) and dichte tonstein (fine-grained kaoiinite). Bouroz

etal. (1983) and Bohor and Triplehorn (1993) have made some

progress in relating the Schuller classification to the tuff

nomenclature. Dichte tonstein is interpreted as a clay-altered

vitric

tuff.

A graupen tonstein could be equivalent to a vitric

tuff containing pumice shards. There is also the possibility

raised by Diessel (1985) that highly altered biotites which lose

the original lath-shape of the grains and become barrel-shaped

could be responsible for a texture characteristic of graupen

tonsteins. The pseudomorphosen tonsteins are clearly altered

crystal tuffs and so are some of the kristal tonsteins. The

difficulty with other kristal tonsteins is that the crystals of

kaoiinite are vermicular and may have formed from the matrix

BENTONITES AND TONSTEINS

rather than being a pyrogenic mineral. If this was resolved then

bentonites could be related to the equivalent tuff category with

greater certainty.

Rate of ash alteration and clay stability

The alteration of volcanic glass is reviewed in detail by Fisher

and Schmincke (1984). The rate of alteration and formation of

intermediate products are further discussed in Bohor and

Triplehorn (1993) with greater emphasis on the formation of

the clay minerals. There is textural evidence from bentonites

that clay minerals sometimes formed relatively rapidly. The

rate does depend on a number of factors including ash

composition and the pore solution chemistry, which is also

influenced by the depositional environment and the perme-

ability of the enclosing sediments.

In the earlier literature there is an emphasis on low pH for

the formation of kaolinite. The work of Garrels and Christ

(1965) and others shows low ionic activities of component

elements are also important. In seawater, activities are higher

and smectite is stable. For pore solutions to remain within the

kaolinite stability field during diagenesis the system must

remain open.

Use of bentonites

Those bentonites consisting of smectite are of economic value

and have many commercial uses (Grim and Guven, 1978).

Tonsteins are generally not of such value with the exception of

some deposits in China.

Bentonites are comparable to tephra layers in that they are

chronostratigraphic markers and are very important in many

parts of the stratigraphical column. Tonsteins came to

prominence over a hundred years ago because of their

stratigraphic values, and correlation of coal seams in the

paralic coalfields of mainland Europe would have been

inconceivable without them. The survival of pyrogenic

minerals in bentonites creates excellent opportunities for

absolute age dating, often from parts of the column where

other suitable materials are absent. The time consuming task

of separating a sufficient mass of minerals for analysis is now

redundant with current instrumentation (SHRIMP) capable of

producing age profiles within grains.

Summary

Bentonites are clay-altered volcanic ash layers. They are

formed during explosive volcanism and as a consequence have

a wide distribution. Bentonites are characteristically preserved

as thin beds in low energy depositional environments, such as

shales and coals, and have considerable stratigraphic value.

Many bentonites in Mesozoic sediments consist of smectite.

Some authors would restrict the term bentonite to smectite

rich deposits irrespective of origin. As the diagenetic level

increases the smectite is converted to illite via a mixed-layer

clay. These are the K-bentonites. Another group were

subjected to greater alteration, which produced kaolinite and

these are the tonsteins. The terms K-bentonite and tonstein are

well established in the literature and should be retained. If

smectite-bentonite were generally adopted potential confusion

could be avoided.

Bentonites are recognized as altered volcanic ashes on bed

form, structures, textures, mineralogy and trace element

geochemistry. The latter two also enable the original ash

composition and the tectonic setting of the vulcanism to be

determined. Although tonsteins result from the greatest level

of alteration their classification and that of tuffs have many

features in common.

Alan Spears

Bibliography

Bohor, B.F., and Triplehorn, D.M., 1993. Tonstein: altered volcanic

ash layers in coal-bearing sequences.

Geological

Society of Atnerica

Special

Paper,

285, 40p.

Bouroz, A., Spears, D.A., and Arbey, F. 1983. Essai de synthese des

donneesacqiiises sur la

getie.ieet levolution des

marqueurspetrographi-

quesdansleshassinshouilters. Memoire XVI, Societe Geologique du

Nord.

Diessel, C.F.K., 1985. Tuffs and tonsteins in the Coal Measures of

New South Wales, Australia. 10th Congres International de Strati-

graphic

de Geotogiedu

Carhonifere.

Madrid.

1983,

4: 197-210.

Fisher, R.V., and Schmincke, H.U., 1984. Pyroclastic Roeks. Springer

Verlag.

Garrels, R.M., and Christ, C.L., 1965. Solutiotu, Minerals and Equili-

bria. Freeman, Cooper and Co.

Grim, R.E., and Guven, N., 1978. Bentonites: Geology, Mineral-

ogy. Properties and Use. Developments in Sedimentology, 24:

Amsterdam: Elsevier.

Huff,

W.D., 1983. Misuse of the term "bentonite" for ash beds of

Devonian age in the Appalachian basin. Discussion and reply.

Geological

Society of Ameriea Bulletin, 94: 681-683.

Huff,

W.D., and Morgan, D.J., 1990. Stratigraphy, mineralogy and

tectonic setting of Silurian K-bentonites in Southern England and

Wales. In Farmer, V.C., and Tardy, Y. (eds.).

Proceedings.

9th In-

ternational Clay

Conference.

Strasbourg

1989.

Sci. Geol. Mem., 88:

33-42.

Knight, W.C., 1898. Mineral soap. Engineering Mining Journal, 66:

481.

Lyons, P.C., Spears, D.A., Outerbridge, W.F., Congdon, R.D., and

Evans, H.T. Jr., 1993. Euramerican tonsteins: overview, magmatic

origin and depositional-tectonic implications.

Palaeogeography.

Pa-

taoclimatology, Palaeoecology, 106: 113-139.

Pettijohn, F.J., 1975. Sedimentary Rocks. 3rd edn. Harper and Row.

Roen, J.B., and Hosterman, J.W., 1982. Misuse of the term

"bentonite" for ash beds of Devonian age in the Appalachian

basin.

Geological

Society of Ameriea Bulletin, 93: 921-925.

Ross,

C.S., and Shannon, E.V., 1926. The minerals of bentonite and

related clays and their physical properties. Journal of the American

Ceramic Society, 9: 77-96.

Weaver, C.E., 1963. Interpretative value of heavy minerals from

bentonites. Journalof Sedimentary Petrology, 33: 343-349.

Winchester, J.A., and Floyd, P.A., 1977. Geoehemical discrimination

of different magma series and their differentiation products using

immobile elements. Chemical

Geology,

20: 325-343.

Cross-references

Diagenesis

Kaolin Group Minerals

Mixed-Layer Clays

Mudrocks

Smectite Group

BERTHIERINE

Introduction

Berthierine is an iron-rich, aluminous, 1 :

1-type

layer silicate

belonging to the serpentine group (Bailey, 1980; Brindley,

1980,

1981). The name honors Berthier, who reported the first

studies of this mineral in 1827. In the past, this clay mineral

was also known as septechlorite, septechamosite, chamosite,

7A-chamosite, and 7A-chlorite, but berthierine is now the

accepted name (Brindley et al., 1968). Berthierine was

considered rare until recently. The failure to recognize

berthierine arose mostly from its chemical similarity to Fe-

rich chlorite (chamosite); these phases share many significant

X-ray powder diffraction (XRD) lines. The ease with which

berthierine can be mistaken for kaolinite in XRD patterns of

oriented crystal aggregates has also contributed to its

misidentification (Moore and Reynolds, 1997, p. 139).

Berthierine is chemically and structurally similar to odinite.

However, odinite is now formally recognized as a distinct

0.7 nm phase (Bailey, 1988a), and exhibits some key differences

from berthierine, the most significant being the abundance of

ferric iron. Odinite is a member of the verdine facies, a group

of green clay minerals present on many shallow, marine shelves

and reef lagoonal areas in tropical locations (Odin, 1988). It

has been previously identified as berthierine or 7A-chamosite

in some localities (e.g., Porrenga, 1967; Rohrlich etal., 1969).

Berthierine is best known from shallow marine environ-

ments, particularly ironstones (Van Houten and Purucker,

1984).

However, it is increasingly recognized in brackish

to nonmarine environments. These include laterites, fiint

clays and other pedogenic settings, and estuarine/deltaic

sandstones, where it may have precipitated directly or formed

by alteration of a precursor (kaolinite, aluminous goethite,

odinite). Berthierine and odinite themselves may be precursors

to diagenetic chamosite (Bailey, 1988a, b; Hornibrook and

Longstaffe, 1996).

Structure and chemistry

Berthierine occurs in two structural forms, orthohexagonal

and monoclinic, the former being most abundant. Both can be

present in the same sample. The cell parameters of berthierine

are a = 0.540 nm, b = 0.936 nm, c sin/J = 0.170 nm, and /? =

104.5° and 90° (Bailey, 1980). Disorder within the structure is

common, arising from variations in layer stacking and

chemistry (Brindley, 1980).

Berthierine is characterized in XRD by a (001) diffraction at

0.7 nm, which defines its 1 : 1 layer structure, and a (060)

diffraction at 0.1555 nm, which indicates its dominantly

trioctahedral character. The XRD identification of berthierine

is described in Bailey (1980, 1988b) and Brindley (1980).

Identification by XRD alone is difficult, however, because of

the presence of two berthierine polytypes and the similarity to

chamosite. Chlorite ought to be distinguishable from berthier-

ine in XRD by its characteristic 1.4nm (001) diffraction.

However, this diffraction is very poorly developed or absent in

chamosite, while its (002) diffraction at 0.7 nm overlaps that of

the berthierine (001). Berthierine also can occur with

chamosite (as a physical mixture and/or an interstratification),

which further complicates identification. Unambiguous recog-

nition of berthierine commonly requires direct measurement of

layer thickness using TEM (transmission electron microscope)

lattice-fringe images (e.g., Jiang etal., 1992).

The generalized formula of berthierine is (R'^'''aR''^bnc)

(Si2.xAlx)O5(OH)4, where R^"'"(Fe, Mg, Mn) and R^"'"(Fe, Al)

are cations in octahedral positions, D describes possible

vacant octahedral sites, and a4-b-)-c =

(Brindley, 1982).

Berthierine is dominantly trioctahedral, with a typical com-

position of (Fe^+i.5Mgo.375Al|.ono.l25)(Sil.25Alo.75)05(OH)4

(after Moore and Reynolds, 1997). However, a wide range in

composition is known (Brindley, 1980, 1982). For example,

end-member Fe-berthierine [(Fe^"^2Al)(SiAl)O5(OH4)] occurs

in Cretaceous laterite deposits of Minnesota (Toth and Fritz,

1997).

By comparison, di-trioctahedral berthierine is present

in oil-sands from Alberta, (Fe'^'^i.oiAlo.82Mgo.46Fe"^'''o28

Mn<o.oino.43)(Sii 74Alo.26)05(OH)4 (Hornibrook and Long-

staffe, 1996). Its composition suggests the possibility of a series

between berthierine and odinite. Odinite is similar to

berthierine in that it is Al- and Fe-rich, has a 1 : 1 layer

structure, and shares many of the same XRD lines. However, it

is richer in Fe^""", Mg and Si, and contains less Al than most

berthierines. It has a representative formula of (Fe'^g

Mgo8Alo53Fe^"'"o3no57)(Si| 8Alo 2)O5(OH)4 (after Moore and

Reynolds, 1997).

Occurrence and formation

Marine

Berthierine is most commonly associated with marine-oolitic

ironstone formations, which are generally considered to have

formed in low energy, marginal marine shelf environments in

tropical to subtropical regions. There, berthierine is typically a

component of pisoids and ooids, consisting of concentric

sheathes around a nucleus of clay, ferric oxide, other ooids,

fossil fragments or quartz grains (Bhattacharyya, 1983; Van

Houten and Purucker, 1984; Maynard, 1986; Siehl and Thein,

1989).

Van Houten and Purucker (1984) summarized possible

origins for such berthierine and other 'chamositic' minerals.

Neoformation of berthierine is possible where grain-coatings,

pore-linings and void-fillings are observed. However, the

general consensus favors early diagenetic, anoxic modification

of a precursor, which had accumulated as granules on an

oxygenated seafioor. Numerous identities for this early

substrate have been proposed. Bhattacharyya (1983) showed

that remobilization of iron from hydroxides could react with

detrital kaolinite to form berthierine in seawater. Odin (1988)

and his coauthors suggested that odinite could evolve to

berthierine during early diagenesis, producing a better

organized 0.7 nm structure with only moderate chemical

modification. Morphological relationships are usually cited

as the main difficulty in accepting odinite as a precursor to

berthierine. Odinite in Recent sediments almost always

occurs as peloids of presumed fecal origin, whereas there

are no reports of grain-coating or pore-lining odinite.

Reworked, transported and transformed lateritic minerals

are another possible precursor for berthierine in shallow

marine sediments.

Nonmarine

Nonmarine occurrences of berthierine are reported with

increasing frequency, including laterites (Toth and Fritz,