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

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CEMENTS AND CEMENTATION

115

from hot. geochetnieally evolved, basin-derived brines. Unlike

the (cited masses of crystals that form under near-surface

conditions, mesogenetic hydt"othermal anhydrite forms as

large,

blocky crystals that may also replace the carbonate

host rock.

Silica eements ean fortn at relatively early stages of burtal

where primary opal (opal-A) is present in the form of siliceous

fossils (diatotns, radiolarians. siliceous sponges) or voleanic

ash.

Small particle sizes, high surfaee areas, high solubility nf

opal relatixe to quartz, and increasing temperatures with depth

all drive opal alteration (Williatns

etui..

W85). C)pal-A tnay

undergo diagenetic alteration to opal-CT, a transfortnation

that can yield cement

k'pispltcrc.s

(q.y.). However, the ultimate

transformation product normally consists of one or more

fortns of quartz: megaquariz, tnicroquartz (chert), or

chal-

cedony that form either as replacement of carbonate grains or

as cemetit. fibrous quartz (chalcedony) occurs in both length-

fast (chalcedonitel and length-slow (quaitzinc and/or luteeite)

forms; the presence of length-slow chalcedony ean indicate a

possible association with evaporites or fortner evaporites.

Understanding the formation of burial-stage cements

depends largely on observations of the sequenee of cement

formation (paragenesis) eoupled with geoehemieal informa-

tion.

Paragenctic dctertninations show the relative tempotal

relationships between any given cctncnt and a variety (if other

geological events, ihe absolute titning of which may or tnay not

be known. Sueh events can inelude earlier cetnentation

episodes, compaction effects (including stylolitization). uplift

and possible exposure, development of fractures or other

defortnation features, hydrocarbon emplacement, and hydro-

thertnal tnineralization.

Several types of geochetnical information are used to assist

in the recognition of burial-stage cements. Trace-element data

is useful mainly because most near-surface eements are

produced in oxic cnvirontnciits and many btirial-stage eements

are (brnied at louer Eh levels. Thus, Fe, Mn and other tnetals

concentrations typically are elevated in burial eements relative

to earlier-formed precipitates. Stable isotope geochemistry

(primarily "'O/'^O) is widely applied to the study of timing of

carbonate cementation because ''O/'^O ratios depend, in part,

on tetnperature of precipitation. Thus, burial-stage cetnents

typicalK arc cnrichctf in "'O (and, largely as a result of organic

tnatlcr alteration. '"C) relative to most tnarine precipitates.

Unfortunately, the isotopie composition of waters tVom which

precipitation occurs also strongly affects the isotopie ratios of

the cement, tnaking it impossible to determine uniquely the

temperature of formation from isotopie measurements alone

on a single precipitate. The same "light" O and C ratios that

eharacterizc burial cements are also found in many near-

surfaee meteoric precipitates. Analyzes of two or more

minerals

(e,g,,

calcite and quartz) precipitated al the same

time and frotii the same waters ean resohe the problem of

dealing with two "unknowns'", but proving sitnultaneous

precipitation is difficult.

Homogenization tneasurements on Jtuut inclusion.^

(</.

r,)

incorporated within cements can be used to directly determine

preeipitation tetnperatures. and thus, their time of formation if

the geothertnal history is well eonstrained (Goldstein and

Reynolds. 1994). In addition, freezing-point determinations on

the fluid inclusions ean provide data on the chemical

composition of precipitating fluids. Such studies have been

applied \er\ successfully to a variety of minerals, but should be

approached with care for several reasons. Inclusions can be

incorporated into crystals when they arc precipitated bul also

at later stages through the healing of fractures, clea\agcs and

other dislocations. Only "primary"' inclusions provide aceurate

temperatures and fluid chemistries of initial crystal formation;

thus,

their proper recognition is critical, but often difficult. In

addition,

some minerals, especially calcite, are unlikely to

tnaintain intact primary fluid inclusions during deeper burial

and elevated temperatures. Several studies have shown that as

little as 15 C of "overheating"' can cause primary inclusions in

calcite to rc-equilibrate. thus altering their appareni fortnation

tetnperatures (Goldstein and Reynolds. 1994),

In recent years. '^^Sr/^^'^Sr determinations have been applied

to carbonate cement studies. Compilation of a seawater

**^Sr/*^''Sr secular variation curve for the Phanerozoic (Burke

etal.. 1982) ean aid in the identifieation of unaltered tnarine

precipitates. Many burial cements, by contrast, have ^ Sr/^^Sr

ratios far more radiogenic than tnarine cements in the same

rocks, largely ihrough incorporation of stronlitun derived

from subsurface altct"ation of feldspars, clays and other elastic

terrigenous minerals.

Absolute age determinations using a variety of radiotnetric

tneihods ('•'C, Th/Th. -"'Th/^'^U/-'^U, -'^UP'Tb. and others)

have been used to constrain the timing of earbonate cetnent

formation and continue to be perfected. Additional dating

leehniques (K/Ar. AtVAr. U-Th, and others) are applicable to

non-carbonate precipitates, such as authigenic elays or

feldspars.

Uplift-stage (telogenetic) cements

carbonate rocks

Uplift typically leads to reduced tcmpcniturcs and pressures

and may be accotnpanied by substantial changes in pore fluid

composition due to deep penetration of tneteorie fluids frotn

elevated recharge areas. This generally leads to substatitial

dissolution, with generation of solution-enlarged fractures,

vugs and caves. Carbonate cements precipitated during this

stage are volutnetrically subordinate to earlier precipitates and

are difficult to distinguish from cither early meteorie or burial

eements e.xcepi by their paragenetic relationships. Thus, they

are discussed only very briefly here.

Carbonate cctncnls nortnally remain unaltered or may

undergo partial dissolution during telogenesis. but evaporite

eetnenis tnay undergo dramatic transfortnations, Atihydrite

cements, beeause of their high solubility in meteorie waters,

may be extensively dissolved and repreeipitated as gypsutn or

may be leached entirely during uplift. Calcite cements

commonly fill some of the porosity that is newly generated

by sueh dissolution (Scholle etal., 1992).

Cements

sandstones

and

other siliciclastic strata

Cementation of siliciclaslic sandstones and conglomerates can

be quite complex where siliciciastic strata have a diverse

primary tnineralogicai eotnposition, they also undergo com-

plex subsurface alteration processes. Many of the varied

detrital grains in siliciclastic strata act as substrates for the

precipitation of overgrowth cements. Other grains are unstable

in surfleial or shallow- to moderate-burial eonditions and

dissolve lo yield a variety of eomplex aqueous solutions

capable of precipitating a diverse range of cetnents.

Three precipitates accounl for the vast tnajority of

cementation in clastic terrigenous strata: various forms of

silica (quartz and opal), carbonate minerals (mainly calcite).

116

CEMENTS AND CEMENTATION

and authtgemc elay tninerals. Many other minerals, including

carbonates (dolomite and siderite). evaporites. zeolites, iron

oxides, iron sulfldes. and feldspars occur as cements in lesser

volumes or under speeialized eonditions: many of these

cements are illustrated in Scholle (1979),

Syndepositional

early diagenetic (eogenetic stage)

cements

siliciclastic deposits

It is uncommon for siliciclastic sediments to be extensively

cemented during eogenetic diagenesis, especially as compared

to their carbonate counterparts. Some hardgrounds, especially

phosphatic ones, are known frotn tnodern silieielastic shelf

sands in areas of very low sediment aecuniulation rates,

Beachrock cementation oeeurs on siliciclastic shorelines in

areas of marine carbonate formation

(e.g,.

the Red Sea and

Gulf of Eiat/Aqaba) and carbonate cementalion of sheli'sands

occurs in extensive areas of methane seepage, Aragonite and

high-Mg calcite are the predominant tnarine earbonate

cements in modern tnarine siliciclastic settings, as they are in

carbonate environments.

The most common sites of eogenetic cementation of

silieielastic deposits are surfleial to shallow subsurface meteoric

environments. In surfleial settings, calcium carbonate-bearing

waters are drawn to the surface by eapillary forces and

evaporatively concentrated. This leads to precipitation of

micritic low-Mg calcite cements forming highly lithiflcd

calcretes and caliche crusts. This proeess occurs primarily in

areas of shallow water tables in semiarid to arid settings and,

in some instanees, is substantially influcnee by transpiration

(sotnetimes yielding cemented root struetures, tertned "rhizo-

erctions'"). Dissolution and repreeipitation of silica- and/or

iron-bearing minerals by surface evaporation creates silieeous

or ferruginous surface crusts, termed silcretes and ferrieretes,

respectively. In arid regions, it is even more common to

observe near-surfaee dissolution, evaporative concentration.

and reprecipitation of sulfate evaporites, producing surficiai

layers of siliciclastic sediment that are completely cemented by

coarsely crystalline, poikilolopie gypsum (gypcretes).

Influx of calciutn carbonate-rich groundwaters in teetoni-

eally active basins oeeurs by subsurface flow from

adjacent highlands that may contain carbonate sti"ata under-

going surfleial weathering or subsurface dissolution. Surface

re-emergenee and evaporation of those waters produces

spring-fed low-Mg calcite travertines as well as interstitial

cementation and surface coating of fluvial sediments down-

stream from groundwater emergence sites.

Early diagenetic infiltration and lateral flow of gtound-

waters can also lead to localized concrelionary cementation of

siliciclastic sediments. Carbonate eoncretions are most

eotn-

mon.

but iron oxides and hydroxides, iron sulfides, silica,

barite,

and gypsum also form concretions. Concretions are

spherical or elongate masses of sediment cemented by finely

crystalline precipitates. Scptarian concretions (see Scptarian

Concretions) are precipitated carbonate nodules that also have

fractures filled with eoarse, multigeneration earbonate. In-

dividual concretions can range from millimeters to several

meters in length, and these sometimes coalesce to form larger

sheets or irregular masses. Most concretions are syngenetic to

very early diagenetic precipitates, although some have long

precipitation histories or fortn during late diagenesis.

Carbonate concretions normally are composed of tnultiple

generations of non-ferroan low-Mg calcite, reflecting the

consistently high Eh and compositional variability of near-

surface waters. However, where eogenetie redueing eonditions

exist, siderite and/or ferroan ealcite form concretionary

preeipitates. Root traces or other permeability eorridors,

loealized concentrations of organic matter, distribution of

carbonate fossil fVagmenls. and other factors play roles in

localizing sites of concretion fortnation, and the directions of

groundwater flow may control the elongation of eoncretions.

Siliceous cetnents also are fortned in early diagenetie

settings, espeeiaily in voleanielastic rocks. Amorphous silica

(opal) is a common preeipitate in sandstones that have

volcanic glass as a primary constituent. Near-surface

devi-

trilication of glass shards proceeds rapidly in meteoric waters

and the increase in aqueous silica concentrations is followed

by opal precipitation. The opal eements, however, normally

arc transfortned to one or tnore fortns of quartz (chert,

chalcedony, or megaquartz) at the increased temperatures

encountered during deeper burial. Zeolites (especially ana-

Icime,

clinoptilolite. laumontile, and phillipsite). clays (espe-

cially smectite and kaolinitc), and calcite are other common

cementing agents in early- to intermediate-stage diagenesis of

voleanogenic sediments.

Iron oxide grain coatings, once considered entirely detrital,

in many eases now are recognized to be early diagenetie

preeipitates (Walker, 1976). Amorphous or extremely finely

erystalline hetnatite and hetnatitc precursot"s (mainly goethite)

form as grain eoatings in oxidizing near-surface settings,

produeing reddened sediments (redbeds, q.v.). The iron is

derived from dissolution of unstable iron-bearing minerals,

ineluding amphiboles, olivene, and magnetite.

Burial (mesogenetic stage) cements

siliciclastic

deposits

Much more extensive cementation of silieielastic strata takes

place during moderate burial and comtnonly is dominated by

quartz precipitation, especially in strongly quarti^ose sediments

(quartz arenites, orthoquartzites), Mesogenetic sources of

silica for cementation are varied, but pressure solution of

detrital quartz grains was once considered to be the primary

soutcc. That importance is now questioned, and other sourees

of siliea are considered to be more signilicant by many authors

(e,g..

Land. 1984). Alternative silica sources include dissolu-

tion and reprecipitation of less stable forms of silica (especially

opal from radiolarians, diatoms and volcanic glass, and

extremely fine-grained quartzose eolian dust), subsurface

alteration or dissolution of amorphous aluminosilieates,

feldspars, and other silicate minerals, and diagenetic transfor-

mations of clays (especially conversion of stneetite to illite)

within sandstones or associated shales (Fiiehtbauer. 1978).

Hot waters from deep basinal sources that eool as they move

upward ean also preeipitate quartz, because silica solubility

deereases with decreasing temperature. Although such basin-

scale fluid movements may be slow, the tens to hundreds of

millions of years over whieh burial diagenesis can act allows

extensive quartz cementation of siliciclastic strata (Sibley and

Blatt. 1976),

Quartz cements form pritnarily on detrital quartz grains and

almost always grow as syntaxial overgrowths

(i.e,.

in optical

continuity with the substrate). Monoerystalline quartz grains

have single-crystal overgrowths; polycrystalline detrital grains

can have more complex overgrowths. A variety of sedimentary

or early diagenetic coatings on detrital grains (including iron

CEMENTS AND CEMENTATION

117

oxides or hydroxides, organic liltns, and clays) can inhibit

formation of quartz ovci growths and help to preserve porosity

in quartz-rich sandstones, but such coatings rarely completely

cover the surface of grains. Overgrowths thus may be initiated

at breaks in the coatings, initially forming as stnall protrusions

frotn the detrital grain surface. Multiple initial growths tnay

eventually coalesee into a complete overgrowth with euhedral

erystal outlines or may completely fill pore spaees. forming

eompromise boundaries with overgrowths on adjacent grains.

Extensive overgrowth prccipitatioti yields a completely cemen-

ted fabric of interlocked quartz grains that can erroneously

appear to be the result of extensive compaction and grain-

boundary suturing (Figure CI8A), The presence of visible

grain coatings surrounding the detrital quartz grains, where

they have not inhibited eemeniation. are a great aid in the

proper identifieation ol" overgrowths as the cause of porosity

destruction. They can clearly delineate the contacts between

dctrit;tl cores and overgrowths, even though both act as a

single erystal. Where sueh "dust rims"' are absent, eathodolu-

minescenee mieroseopy may aid in the recognition of over-

growths (Figure C18B), through accentuation of trace-element

variations between detrital and authiizenic constituents (Sippel.

1968;

Miller, I9S8).

Other forms of silica tnainly chalcedony and ehcrt

(microquartz) can also cement siliciclastic strata in tneso-

gcnctic settitigs. especially where detrital quartz substrates are

relatively scarce or where silica precipitation rates are very

high.

Chalcedony is a fibrous growth form that typically

forms palisade rims and radiating, spherulitic growths in

interstitial pores. Microquartz fortn as tiny, equant. often

inclusion-rich erystals. with anhedral to subhedral tnorpholo-

gies,

commonly inctcasing in crystal size from pore wall to

pore center.

As wiih carbonate eetnents. fluid inclusion geothermomett"y

and stable isotope geochemical studies are very useful in

determining the temperature and geochemieal conditions of

quartz cement formation. Because ofthe physical strength and

lack of cleavage in quartz, fluid inclusions studies of this

mineral are far more reliable than those done on caleite.

dolomite or tnost evaporite minerals.

Burial-stage calcite is an itnportant tnesogenetic cetnenting

agent in tnany silieielastic strata and can form throughout the

history of mesogenetic alteration. Early-stage ealeites fonn

ptcdominatuly where a signifleant volume of unstable earbo-

nate (commonly aragonitic shell fragments) was present in the

original sedimeni. These grains are readily dissolved during

early diagenesis in the presence of nonmarine or low-salinity

(mixed) pore fluids. The presence of calcitic shells or other

carbonate grains ean provide is<ilated nucleation sites for the

growth of calcite cements from carbonate-saturated pore

waters. These isoiaied sites of cement formation allow

nucleated crystals lo grow unitnpeded by eotnpetition from

neighboring precipitates. This commonly produces coarsely

crystalline cements and even poikilotopic fabrics in which

single calcite crystals engulf several detrital grains. The force of

crystallization of such early mesogenetic calcite can lead to

displacctncnl or fracturing of detrilal grains atid produce

apparetitly undercompacted or "floating" fabrics.

Late-stage calcite cements lypicully postdate quartz over-

growths and iiuthigcnic clays. They result from the f^act that

calcite solubility decreases at increased tempet"atures and

basinal brines may have elevated pH and high Ca and COi

concenttations. In addition, under most deep burial

conditions, waters arc strotigly reducitig, and latc-stagc ealeites

therefore are characterized by elevated iron and tnanganese

concentrations.

Mesogenetie dolomite and siderite cetnents can be locally

itnportant in siliciclastic deposits. Dolomite cements can be

anhedral to euhedral, generally are zoned and tend to be

ferroan,

reflecting reducing conditions during formation.

Siderite cements tend lo be finely crystalline with yellow-

brown flattened rhombs; they loo require reducing conditions

for formation.

Clay minerals iti siliciclastic rocks were onee eonsidered to

be entirely detrital. In recent decades, however, numerous

studies have demonstrated that authigenic clays are extremely

common and precipitate under a variety of conditions (Blatt.

1979), in some cases substantially affecting the pctrophysical

properties of sandstones (Bjorlykke, 19S8t, Illite, stneclite, and

mixed-layer smectitc-illite and smectitc-chlorite and kaolinite

ali occur as porc-fllling and grain-coating cements and ean be

difficult to distinguish from detrital counterparts. Some

authigenic clays, espeeiaily stnectite-illite grain eoatings,

elearly predate all other cements and may be of eogenetic

origin.

As sueh. these elays can inhibit later quartz or calcite

eemeniation and thus help to preserve sandstone porosity.

Other clays fortn later in the cementation history, Kaoliniie

(or its polymorph, dickitc) occurs as authigenic vermicular

stacks ("books" of mica-like plates) that generally postdate

eogenetic cetnents, Becattse kaoiinite precipitation is favtired

by low pH eondilions. its formation sometimes is associated

with deeotnposition of organie matter or meteoric water influx.

Alteration and recrystallization of both detrital and authigenic

clays during progressive burial can further complicate

recognition of primary versus secondary origins or litning of

precipitation.

A number of other cetnents, iticluding authigenie feldspar

overgrowths, pyrite and other iron sulfides. barile, fluoriie,

sphalerite, and other hydrothermal minerals, ean fbrm during

\arious stages of siliciclastic diagenesis. However, they

normally have only local importance in controlling rock

properties and so will not be detailed here.

Summary

Many cements and eement fabrics are known frotn carbonate

and siliciclastic strata. In combination, they occlude 20 40

percent porosity in many sedimcntaty rocks and contribute

substantially to sediment lithiflcation. The fabrics and para-

genetic relationships of cements provide important clues to

the time and conditions under which they were fortned.

Those interpretations are strengthened when done in combina-

tion with trace-element analyzes, isotopie studies, and

fluid-

inelusion geothermomeiry. Absolute age determinations on

cetnents are becotning tnore precise and as techniques are

perfected for analyzing small, in situ samples, such determina-

tions will substantially improve our understanding of cetnen-

tation.

Future studies may elarify the poorly understood role

of microbial activity in mediating cetnentation. especially in

the subsurface, Likewise, rclining our understanding of the

factors that prevent or retard cetnentation retnains of

fundamental importance, because it aids in predicting the

loeation and quality of possible uncotiventional oil atid gas

reservoirs and economic mineral deposits.

Peter A, Scholle and Dana Ulmer-Seholle

118

CFMENTS AND CEMENTATION

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Spt'ciniscopy: Qualilaliyeand Qiiantilalive

Application.'..

Tulsa. OK:

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Tables.

Amsterdam;

Elsevier,

Sandbetg, P,A,, 1983, An oscillating trend in Phanerozoic non-skeletal

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Schmoker, J,W,. and Hailey, R,B,, 1982, Carbonate porosity

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.4s.sociaiion

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2570,

Scholle, P,A,, 1978,

Color flluslraletlGuide

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Association of Peiroieum Geologists Memoir 27,

Scholle, P,A,, 1979, A Color Illustrated Guide to Constituents.

Tiwttires.

Cements, and

Porosities

Sand.stone.s

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.AssociatedRoek.s.

Tulsa, OK: American Association of Peiroletim Geologists

Metnoir 28,

Scholle. P,A,. and Halley, R,B,, I9S5, Burial diagenesis: out of sight,

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pp,

.309-334,

Schollc. P,A,, Lllmer, D.S,. and Melim. LA,. 1992, Laie-slage ealeiies

in the Pennian Capitan Formalion and its equivalents, Delaware

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Scoffin, T,P,. and Stoddarl, D,R,. 1983, Beachrock and inlertidal

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andCeomorphology. London: Academic Press, pp, 401 425,

Sibley. D,F,. and Blalt. H,. 1976. Intergranular pressure soltition and

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Sippel. R,F,, 1968, Sandstone petrology, evidenee from luminescence

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Waiker. T,R., 1976. Diagenetic origin of continenial red beds. In

Falke, H, (ed.). The Continental Permian in Central

West

and Soiiih

Europe. Dordrecht (Netherlands): D. Reidel Publishing Co,.

pp,

240 282,

Williams. L.A,, Parks. G,A,. and Crerar, D.A., 1985, Silica diagenesis, I.

Solubility controls, JournalofSediinenturyPenology. 55: 301 311,

Wollast. R,. I97L Kinelic aspects of the nucleation and growth of

calcite from aqtteous solutions. In Bricker. O,P, (ed.). Carbonate

Cements. Baltimore. MD: Johns Hopkins Press, pp, 264-273,

Cross-references

Ankerite

Beachrock

Caliche Calcrete

Cathodoluminescence

CHALK

119

Chalk

Chen. Chalcedony and Flint

Compaction (Consolidation) of Seditnents

Diagenetic Structures

Difftision, Chemical

Dolomiles and Dolomitization

Evaporiles

Glaucony and Verdine

Gypsum

Hardground (and 1-irmground)

Hydrocarbons in Sediments

Kaolin Ciroup Minerals

Laleriles

Lepisphere

Micritization

Porewalers in Sediments

Pressure Solution

Ouiirtz. Detrilal

Red Beds

Sabkha, Salar, Sail

l-lal

Sand and Sandstones (Siliciclaslics)

Seawater

Septarian Concretions

Silcrete

Siliceous Sediments

Stneetite Ciroup

Speleothems

Styloliles

Travertine

Zeolites in Sedimenlarv Rocks

CHALK

The Oceati Drilling Program dcilnes chalk as a Hrtn pelagic

sediinetit coniposcd predotninaiitly ofcalcareous pelagic grains

(e,g,. Kroenke i'/(//.. 1991. see Oceanic Sediments). Calcareous

pelagic grains comprise calcareous nannot^ossils of silt size and

calcareous microfossils of mm size, as well as fine fossil debris.

The term chalk ciin also refer to a calcareous tiaiinofossil-rich

facies irrespective of iiiditration. The grains are mixed into

sediments

ot"

mudstone. wackestotie. or packstone texture (see

ChtssiticationofSeiliittents and Sedimentary Rocks). Calcareous

nannotbssils include coceoliths which originate from spores of

golden brown algae, and are found in sediments ranging from

Late Triassic to present (Siesser and Winter. 1994), Micro-

Ibssils arc typically foraniinifers.

Chalk facies sediments contaiti varying atnounts of siliceous

tnicro- and nannofossils wich iipoti burial tnay form chert (see

Siliceous Sediments). Similarly, chalk also contains varying

amounts of clay which upon burial may fortn stylolites or

flaser structures (Garrison ttnd Ketinedy. 1977, see Diagenetic

Slructttres).

Oceanic chalk deposits form in areas where terrestrial input

is low. and where the seafloor is above the carbonate

compensation depth, typically on deep sea plateaus (see Ocea-

nic Sediments). These seditnenls are almost exclusively of

Cretaceous age or younger, Lipon the large transgression in

Upper Cretaceous, chalk facies seditnents accumulated on the

continental shelf (liakansson etal.. 1974), The Chalk Group of

the North Sea Basin is pritnarily composed of chalk sediments

of Upper Cretaceous and Paleocene age.

Freshly deposited chalk facies sediments have porosities

around 70 percent: and porosities of around 35 percent tnay be

maintained to an effective burial of c.lkm (Scholle, 1977,

Figure C19). This preservation of porosity probably reflects

that the fossils constituting the chalk are mainly composed of

chemically stable low Mg calcite (sec Carbonate Mineralogy

und Geochemistry].

Chalk is utilized as a source of lime and for improvement of

acid soils. In North West Europe it is a reservoir tor

groundwater (Downing ctat., 1993) and in the Central North

Sea for hydrocarbons (Andersen, 1995). Chalk is used for

rinsing smoke after coal combustion in power plants and as a

raw material for production of pure calcite or aragonite for

industrial purposes: pigment in paper, plastic and paint.

Indurated chalk has been used as builditig stone (a beautiful

example is the cathedral of York, England),

Until the 1960s chalk was seen as a uniform white pouder,

possibly formed by precipitation of calcite in the seawater.

and when studied in exposures, the purpose was mainly

extracting macrofossils. With the introduction of electron

Porosity [%]

0 50 100

200-

400-

600-

800-

1000 -

1200 •

1400

Neogene

Paleogene

Cretaceous

Figure C19 Chalk facies sediments on the Ontona lava Plateau at a

water depth of 2,805 m, investigated by the

(^cean

Drilling Pr()gr.im

(Kroenke e( ,iL, 1991), Laboratory porosity versus depth in meters

below sea floor (nibst'l. With burial, the soft calcareous ooze grades

into firmer chalk and this aj^ain to hard limestone. Compaction

experiments indicate that fhe gradual porosity loss through ooze and

chalk is a consequence of physical compaction (Lind,

199'i,

see

Compaction ot

Sediments),

whereas geothemical data indicate that

the limestone is formed via porosity reducing precipitation of calcite

caused by dit'iusion ol' ions tVum where flinl or stylolites form or from

the basement iDelaney et

al.,

1991; see

Cemcntfi

dnd

Cement^Uion}.

The small scale porosity variation in the chalk section pritnarily reflects

varying proportion of carlxinate mud and hollow micrtjtosBifs (Masters

et al., 1993), whereas the more dramatic porosity shift within the

limeslone reflects an up|)er wdckestone interval ri( h in lemenled

microlossils overlying an interval nf re-crystallized c-irbonate

mudslone (Borre and Fabricius 1948; see F.ibric,

Porosity

and

Permeability).

The interval from 970mbsf to

lOOmbsf is rich in flinl

beds isee Siliceous

Sediments).

120 CHALK

Figure C20 While chalk containing about 5-10"/(i ti.iy minerals,

Maa^Irichlian, form ihe quarry "Dania"

northern lutland, Denmark.

The chalk has been sectioned vertically and smoothed. The surtace has

been lifjhtly oiled with common lubricating

oil

(see Bromley, 1'58I

The treatment greatly increases the visibility ot'lhe sediment structures.

The sediment has been tully bioturi><Ued

burrowing animals.

The

Irace fossils represented are:

(T)

ThiiLissinoides,

(Z)

Zoopiiycos,

(P)

PUnolites

and (C)

Chondrites

(see Biogenic

Structurusi.

R. G. Bromley.

FigureC21 Electron

mii

rngr<i|in m

M.I.ISIIH

ii.in

(

H.IIK

iroiii ilic

iMilr.il

North Sea. The samf)le has

mudstone lexture,

talcile content

98.8 wt.'/o and

porosity,

O, of

377i..

The mud

composed

coccoliths and fossil debris.

All

calcite particles are apparently

thoroughly

re-crystal I Ized

and because the cLiy content

low,

the

specific surface, s, as measured

BET

is relatively low

.f>m"'f;'|,

and

the resulting gas permeability 4.3 mD. Chalk

pelrophysically

exceptionally simple sediment which obeys Kozeny's equation relating

permeability, specific surface, and porosity:

k = C' 0

/S", where

the specific surface relalive

bulk volume,

k is

the Klinkenberg

corrected gas permeability (in m") and C

near constant, varying from

0.21

for

a porosity

20%

to 0.24 for

a porosity

40% (Mortensen

etal.,

1998).

microscopy it became clear ihal Ihc while powder is composed

ofnantiorossils and fossil debris: and the expansion of drilling

activity in the oceans and in the North Sea directed the

scientific interest to the study of nanno- and microfossils with

the purpose of establishing ;i hiostratigraphy Lisabie on drilling

cores and cuttings. In the sixties and seventies scientists

intensified the study of sedimentary structures and it appeared

that chalk as a rule is hioturhated {Figure C20). It was also

found that chalk may contain re-sedimented chalk-clast

bearing intervals (e.g., in the central North Sea. Nygaard

et al., 1983). Chalk may exhibit cyclicity in silica content

(frequently seen as series of chert beds) or clay content

(rellcctcd as series of marl beds). By some researchers this was

seen as a signature of geoehemical segregation from an

originally more homogeneous sediment, but from studies of

the pattern of bed thickness variations and oxygen isotopie

variations (see Isotopic Meihods in Sedimeniology) it became

more and more probable thai the cyclicity reflects an original

sedimentary variation which from the eighties have been

related to paleo-elimatie fluctuations (e.g.. Gale, 1989: and see

CyclicSedinienlalidii. Milaiikovinh Cycles).

Major constructions in chalk (e.g.. the Channel tunnel

between England and France) and the search for and

production of hydrocarbons frotn chalk reservoirs has in the

nineties directed research attention to the tiiechanical. hy-

draulic and petrophysical properties o\' chalk (e.g.. Burland

(7 (//.. 1990: Maliva and Diekson. 1992: Mcgson and

Hardmann 2001: Figure C21). Specific current research

questions arc how the porosity loss proceeds in geologic time

in the seditnentary hasin, and the t"ock mechanical hehavior

during hydrocarhon production and hy the change in pore

fluid from oil to waler.

Ida L. Fabricius

Bibliography

Andersen.

M.A.. 1995.

Pi'troleitm Ri'seanli

North

Sea

Cliiilk.

Stavanger, Norway: RF-Rogaland Research.

Borrc,

M., and

Fabricius.

I.L.. 1998.

Chemical

and

mechanical

processes during buriul Jiiigencsis

chalk:

intLTpreialion biised

on specific surface dalii

deep-sea sediments. Scdimciilolo^v,

45:

755-769.

Bromley.

R.G,.

1^81. F.nhanccment

visihility

structures

marly

chalk: modification

of the

Bushinsky

oil

technique. Bulletin iif ihe

Ccoh^^kal

Society

Dfimnirk.

29:

111-118.

Buikmd. .I.B. Clai. 1990. Cluilk. London: Thomas Telford.

Delaney.

M.L.. and

Shiphoard Scientific Party.

1991.

inorganic

Gcocheislry Summary.

Kroenke. I..W.. Berger.

W.H. and

.lanecek.

T.R.

(eds.). i'lvcffillii^.sofilu'

Occtiii

Diilliiiii Programliiilt-

al Reports

130.

College Station

TX:

Ocean Drilling Progratn.

pp.

549-551.

Downing. R.A.. Price.

M., and

Jones. G.P., 1993. TheHydrogeologyof

ihc Chalk of

Nofth-We.'il

Europe. Oxford: Claredon Press.

Gale.

A.S.. 1989. A

Milankovitch scale

for

Cenomanian time. Term

Nova, 1:420-425.

Garrison.

R.H.. and

Kennedy.

W.J,,

1977. Origin

solution seiisnis

and tlascr structures

Upper Cretaceous chalks

Southern

England. ScdimcnturyGeology.

19: 1(17 137.

HAkansson.

E..

Bromley,

R.. and

Perch-Niclscn.

K.. 1974.

Maas-

trichtian chalk

north-uesl Europe—a pelagic shelf sediment.

H^ii.

K.. and

Jcnkyns.

W.C

(eds.),

Pekigic

Seilimcnis: on Lamlnml

L'liderihe.Sen. SperlnlPiiNiraiion

1.4S,

I, pp.

211-233.

Kroenke.

L.W,.

Berger, W.H.. Janecck.

T.R.

(•/(//..

1991. Proc. ODP,

Initial Repts. 130. College Stalioii

TX:

Ocean Drilling Program.

Lind.

I.L.. 1993.

Loading experiments

ciirbonaic

oo/e

anti chalk

from

Leg 130.

Oiilonj; Jav:i Plateau.

Berger.

W.H..

Kroenke,

CHARCOAL IN SEDIMENTb 121

L.W . ;nul Ma\er. I..A,

ledsj.

Proceedings of the Ocean Drilling

Program, Scientific Results 130, College Station TX: Ocean

Drilling Program, pp. 67.^ 6X6.

Maliva. R.G , and Diekson, J.A.D.. 1992. Mieroliicies and DiagenL-tic

Control of Porosity in Crctaecous/Tertiary Chalks, Eldfisk Field.

Norwegian North Sea. .-iAPGBidletiii. 76: 1825 IS38.

Masters. J.C. Rcsig. J.M.. and Wiicoxon, J.A., l')93. Relationship

between physical propcriics and microtbssil content and preserva-

tion iti CLilcareoits sediments of the Ontong Java Pliitcau. In Berger.

W.H.. Kroenke. L.W., atid Mayer, L.A~ (eds), Proc. ODP Sci.

Results I.W. College Station TX: Ocean Drilling Progratn, pp.

641 652.

Mcgson.

.1.,

and Hardman. R..

2001.

Exploration for and development

of hydrocarbons in the Chalk of Ihe North Sen: a low permeability

systctii. Pelnileiim(ieii.sciciuc.l\ 3 12.

Mortensen. J.. IZngstrom. I-., and Lind, L 1998. The relation among

porosity, permeability, and specific .siirfiice of chulk from the Gorm

field, banish North'Sca. SPF

31062.

SPE Reservoir Evaluation &

Fngiiieeriii^. 1. pp. 245 251.

Nygaard. R.. Lieberkind. K.. and Frykman. P, 19S.i. Sedimentology

and reservoir parameters of the Chalk Grottp in the Danish Central

Graben. Gcolof^ieen

.Mijiilwttw.

62: 177 190.

Scholle. P.A., 1977. Chalk diagenesis and its relationship to petrolctnn

exploration: oil trom chalks, a modern miracle' A.APGBulletin, 61:

982 1009.

Sicser. W.G.. and Winter. A.. 1994. Composition and morphology of

coccolithophote skeletons. In Winter. A., and Sieser. W.G. (eds.t,

Coccolilliiifihoie.s.

t~ambridge Llniversity Press, pp. 51-62.

Cross-references

Biogenic Sedimentary Structures

Carbonate Mineralogy and Geochemistry

Cements and Cementation

ClassiCication of Sediments and Sedimentary Rocks

Compaction (Consolidation) of Sediments

Cyelie Sedimentation

Diagenesis

Diagenetie Structures

Fabric. Porosity attd Pertneabitily

Notopie Methods in Sedimentology

Milankoviteh Cycles

Oceanic Sediments

Siliceous Sediments

CHARCOAL IN SEDIMENTS

Definition

Wheti plant material is burnt by lire there is often an

incomplete combusteLi residue ktiovvn as charcoal. All plant

organs, frotn leaves lo Howers to wood, may be preserved as

chafcoal. During the clnirring process wood tnay break up into

characlerislic sizes and shapes. It has been shown that wood

tragtnents may typically be aboui I cm cubed. The charcoal is

brittle (getting hands dirty upon touch) and relatively inert

with the cell walls being converted to almost pure carbon

(Scott, 1989). Charcoal usually shows exceptional anatotnical

detail when seeti. for example, under the scanning electron

microscope (Scott and Jones. 1991) (Figures C27. C28. C29).

Hveii small flowers may be preserved as charcoal (Scotl

ctal.. 2000).

Transport

During wild fires small particulate fragments (around 1

IO).tm. including charcoal and soot particles) may be blown

away from the fire sites. These small particles may enter

lakes and other wetland environments or into the sea and

he incorporated into accumulating sediments (Clark and

Patterson. 1997). In contrast, larger charcoal fragments may

either, in the case of fires within mires, be incorporated into

peat or else be transported by water {Scott. 2000). Severe ftres

may destroy plants which help bind soil and it is possible that

rainstorms following major fires tnay result in significant

erosion of the fire site. Sediment and charcoal may be washed

into a depositional site (Nichols and Jones, 1992). Il has heen

shown, for example, that major alluvial fans may be triggered

by fires, such as in Yellowstone National Park (Meyer elal.,

1992).

Macroscopic charcoal fragments usually float in water. The

length of time for a charcoal fragment to be waterlogged

depends on a series of variables from size, tetnperature of

charring, organ and species (Nichols vial.. 2000).

Deposilion

It has been argued that larger charcoal fragments in a

scditnent implies a nearby fire (Clark and Patterson. 1997).

Whilst this may be the case for classes of windblown

tnicroscopic charcoal, however, it is not necessarily trite for

larger macroscopic charcoal fragments. Charcoal may be

incorporated into bed load sediments. Flume experiments have

shown that charcoal may he entrained in sand ripples of

velocities around 35cms ' (Nichols ct al.. 2000). Marine

seditnents may also yield charcoal fragments (Bird, 1997;

Seott. 2000).

Sedimentary distribution

Macroscopic charcoal when encountered in fossil deposits is

often called fusain (Scott. 1989). This is one of the four

fundamental lithotypes of coal described by Marie Stopes. In

the Quaternary there has been a number of significant studies

of microscopic charcoal frotn lake and peat sediments to help

unra\cl more recent fire histories (Patterson ctal., 1987). In

marine clastic sediments also most research has centered upon

tnicroscopic charcoal and soot particles which are likely to

record larger and tnorc fYequent regional fires (Bird, 1997). In

contrast, studies of Pre-Quaternary charcoal has centered

upoti larger charcoal fragments that are visible in rocks (Scott,

2000).

There is a good fossil record of charcoal from the latest

Devonian and evidence of extensive wildfire from the Early

Carboniferous (Scott. 2000).

Most often a geologist will encounter charcoal (fusain)

fragtnents in terrestrial clastic rocks and also cotntnonly they

may be abundant in estuarine. deltaic and shallow marine

deposits. These may be reeognized as small black cubes with a

lustrous sheen which marks the finger upon touch

(Figures C22, C23). This is in contrast to coalified plant

material, which is often black, shiny, having a concoidal

fracture and is clean to touch. Charred leaf material may also

be recognized (Jones and Chaloncr, 1991) (Figure C24).

Smaller charred plant fragtnetits, which may include flowers,

tnay be difficult to see on a bedding surface. In these cases

only bulk maceration, using HF, or sieving of unconsolidated

sediment, will yield the stnaller charcoal fragments

(Figures C25, C26).

Charcoal is often encountered within coal seams

(Figitre C22) where it may represent 10 50 percent of the

122 CHARCOAL IN SEDIMENTS

Figure C22

WOIKI

wvi

fK'w

Figure C25 Charcoal (fusain), macerated from lacustrine clays.

Figure C26 Charcoal (tusainl macerated trom

toal.

Figure C23 Wood charcoal iUi^.iinr m ikjvi<il s.inrlsTnni

lii^^

Figure C27 Reflectance micrograph of Recent charcoal.

Figure C24 tern (.h,ir<."oal ilu^ain) in liuvi.il '^iitstiine.

Implications

seam.

Charcoal tnay also be encoutitcrcd in volcanic rocks. The recording of charcoal (fusain). as opposed to other forms

especially pyroclastic deposits where charcoal has been of organic matter, has considerable implications for the

entombed in hot lava or ash (Scott, 2001). Charcoal is often intcrpretion of paleoenvironmetits and diagenesis. In addition,

not recorded by seditnentologists in the

Held.

microscopic and isotopic studies of

charcoal

tnay yield data on

CHEOKITE IN SEDIMENTS

Figure C28 SFM ot Cretaceous conifer wood chartoal.

Figure C29 SEM of Carbonilerous lytophyte wood charcoal.

liie type, frequency, clitnate. vegeiation, terrestrial ecosystem

dytiatnics as well as an atmospheric composition (both O2 atid

CO2) (Chiilotier, 1989; Scott. 2000).

Andrew C. Scott

Bibliography

Bird. M.I.. 1997. Fire in the earth sciences. Episodes. 20: 223 226.

Chiiloner, W.G., 1989. Kossil charcoul as an indiciitor of palaeo-

atmospheric oxygen. Jourmd of the Geological Society of London,

146:

171-174.

Clark. J.S.. and Patterson III, W.A.. 1997. Background and local

charcoal in seditnents: scales of tire evidence in the fossil record. In

Clark,

.I.S.,

Citchicr, H.. Goldammcr, J.G., and Stocks. IJ. (cds,).

Sediment Records ol Hiotnass Burning anil Glohitl Change. NATO

ASI Series, Volutne 151. Berlin: Springer, pp. 2.'' 48.

Joties,

T.P.. and Chiiloticr, W.G.,

1991

.Kossil charcoal, its recognition

and palaeoatmospheric significance.

Palaeogeography.

Palacoclima-

toh'iiv. Piihieociology iGlohnl and Planetary Change Section). 97:

39 50.

Meyer, G.A., Wells. S.G., Balling. R.C. atid Jull. A.J.T,, 1992.

Response of alltiviiil syslcni.s to (ire iitid climate change in

"iV-Ilowslone National Pa'rk. \aliire. 357: 147 15(1.

Patterson III, W.A.. Edwards, K.J., and Magiiire, D.J.. 19S7.

Microscopic charcoal as an indicator of tire. Quaternary Scicnei'

Reviews. 6: 3 23.

Niehols. G., Cripps. J., Collinson, M.F., and Scott, A.C.. 2(100.

E.\perimental laphonomy of charco;il and inipticatioiis lor the

palaeoenvironmental inteiptetalion ol" fossil charcoal deposits.

PalaeoReogrtipiiy.

Palacoclimatology. Palaeociology. 164: 43 56.

Nichols. Ci.J.. and Jones, T.P.. 1992. I'usain in Carboniferous shallow

marine sediments. Donegal, Ireland: ihe scdimcntoloijici elTects of

wildfire. Scdimeniolugy, 39: 487-502.

Scott. A.C., 1989. Observalions oji the nature and origin ol' fusain.

International.lournalof Coal (icology. 12: 443 475.

Scott, A.C., 2(H)(). The pre-Quatcrnary history of (ire. Palaeogeography.

Palaeoclimaiolo^y. Palaeoecology, 164: 281 339.

Scott. A.C.. 200!'. Preservation by (ire. In Briggs, D.E.G., and

Crowther, P.R. (eds.). Palacohiology II.

O\ford:

Blackwells, pp.

277-280.

Scott, A.C.. and Jones, T.P., 1991. Microscopical observations of

Recent and (bssil charcoal. Mivro.scop\and.\i}a!y.\is. 24; 1.3 (5.

Scott, A.C, Cripps, J.. Nichols, G., and Collinson. ME., 2()()(). The

taphonomy of charcoal following a recent healhkmd lire: and some

implications for the inierpretation of (bssil charcoal deposits. Pa-

lueogeography.

Palaeoclimatology. Palaeoeeology. 164: I 31.

CHLORITE IN SEDIMENTS

Chlorite is a common phyllosilicate mineral found in all kinds of

seditnents and sedimentary rocks. In fact the name chlorite.

derived frotn the green color of most specimens, cncompas.ses a

group of minerals characierized by a wide range of chetnical and

structural variation (Bailey, 1988a). In sediments, much chlorite

falls by definition into the category of minerals known as "clay

minerals" (sec Clay Mincratoiiy). Small amounts of chlorite are

cotnmon in many sediments and it tnay be both detrital or

fonned diagenetically. Overall, diagenetic chlorite is probably

rarer than other diagenetic clay minerals but it is common in

certain kinds of seditnentary rocks. To undersland some types

of diagenetie chlorite its relationship to serpentine minerals and

to mixed-layer chlorite minerals must be considered (Reynolds,

1988).

Indeed, the petroloyy of chlorite in the sedimentary

realm is diverse. This contribution attempts to summarize

hal

is known about the characteristics, occurrence, and origins of

various kinds of chlorite in sediments.

Structure and composition

.All varieties of chlorite have a common structure consisting of

a negatively charged 2:1 layer and a positively charged

'intcrlayer" octahedrally coordinated hydroxide sheet (Bailey.

1988a). Most chlorites arc trioctithcdral and as such they are

normally rich in Fe and Mg. Less commonly, chlorites may be

dioctahedral in which case they tend to be enriched in

alutninum. For many years the notnenelature and classification

of trioetahedral chlorites was plagued by the existence o'i

nutnerous species natnes. This has been greatly simplified with

Fe-rich chloriies classified as chamosites, and Mg-rich chloriies

as clinochlores, wilh suitable adjective modifiers used if

necessary (Bayliss. 1975). The chemieal cotnposiiions o'(

chlorites are complex, but a generalized structitral formula

distinguishing octahedral and leirahedral cations and encom-

passing both trioetahedral and diociahedral species can be

written as:

(R^+,_3^Rjt:vny (Si4 ^R^ ) O|i,[OH)^

Iti this fortn. R""" represents divalent cations, mainly Fe and

Mg. R'* represents trivalenl cations mainly Al. and D

124

CHLORITE IN SElllMFNTS

represents a vacant octahedral site. Fttlly tt"ioctahedral

chlorites do not have octahedral vacancies so that the

octahedral cation total in the above formula is six. Di-

trioctahedral chiorites. such as the species sudoite, have a

dioclahedra! 2:1 layer, and a trioclahcdral intcrlayer hydro-

xide sheet. Thus in the general formula given above they have

one vaeant stte and an oetahedral cation total of five. Sitiiitarly

a chlorite in which both the 2:1 layer and the interlayer

octahedral sheets were both dioctahedral would have two

vaeant sites and thus only four octahedral eations.

In terms of crystal structure the 2:

layer and the intcrlayer

hydroxide sheet in chlorites may be stacked together in

different ways so defining different chlorite polytypes. For

sotne titne it has been recognized that type 1 polytypes are

common in sedimentary rocks formed at diagenetic tempera-

tures whereas the lib polytype is the cotnmon polylypc for

high temperature metamorphic chlorites. Recent summaries

are provided by de Caritat cial. (1993) and Walker (1993).

In contrast to metatnorphie chlorites formed at higher

temperatures, recalculation of chemical analyzes to structural

formulae for trioetahedral chlorites from sedimentary rocks

otten indicates (he presence of octahedral vacancies, i.e.,

octahedral totals fall short of six cations. It is not yet clear,

however, if such vacancies are present in the structure of

sedimentary chlorites or are simply apparent due the assump-

tions made about the (usually utianalyzed) anions and/or

analytical problems inherent in the analysis of clay minerals,

e.g.. avoiding contamination (Hillicr and Vclde. 1991: Jiang

elal., 1994).

Occurrence and origin

Much of what is known about chlorite in sediments can be

rationalized into a number of broad categories or associations,

each probably with some degree of genetic significance

(Figure C30). These categories include detrital and weathered

chlorite; the odinite-berthierinc-chamositc shallow-marine as-

sociation; the Mg-rich chlorite evaporite/dolomite association;

Detrital

chlorite

Berthierine (Fe2+) 7A

Oditiite(Fe3+)7A

the aluminous dioctahedral chlorite red-bed association; the

lithic/volcaniclastic sediment association, and the shale diag-

enesis/low grade-metamorphism association. However, the

nature of chlorite in seditnents is diverse, and undoubtedly

many occurrences will not fall neatly into one of these

associations; others, such as chlorites in limestones, have not

yet received the study they deserve. Nonetheless, the use of

these associations provides a conceptual framework to help

understand the occurrence of chlorite in sediments in many

instances.

Detrital and weathered chlorite

Detrital chlorite in sediments reflects provenance. In fact,

chlorite is very easily weathered being one of the first tninerals

to disappear frotn soil profiles, even in temperate climates

(Bain, 1977; Banihiscl atid Bertsch, 1989). As a result it is quite

unusual to find abundant detrital chlorite in seditnents, unless

they were deposited in close proximity to chiorite-rich parent

materials. Nonetheless, small amounts of detrital chlorite can

be Ibund in many sediments. In modern oceanic sediments,

chlorite is largely confined to. and concentrated in, sediments

from high latitudes (Chamley, 1989; Weaver, 1989). This is

undoubtedly a direct reflection of the fact that chlorite is such

an easily weathered mineral. Indeed, estimates of its abun-

dance in the clay size fraction of high latitudes oceanic

sediment are only rarely more than 20 30 percent (Windom,

1976).

The effect of weathering on the distribution of detrital

chlorite in marine sediments is also evident from studies of

oceanic sediments from dated cores of the western Atlantic

region, which record a general increase in chlorite content

from the Eoeene to the Quaternary (Chamley, 1989). This

appears to reflect the general pattern oi global cooling over

this titne and consequent decrease in the intensity of chemical

weathering in many source areas. To some extent it may also

reflect a general increase in relief resulting frotn Alpine

tectonistii with the associated stripping o'i more mature soils

and more widespread exposure of low-grade (ehlorite-rich)

Shallow marine sandstoties

Pore litiing Fe-rich chlorite 14A

Red-heds

Al-rich chlorite

tosudite

Shale chlorite

by-product

of iliitization

Evaporites and

carbonates

Mg-rich chlorites

corrensite

Ferro-magnesian/'volcaniclastic

Sandstoties Fe-chlorite, Mg-chlorite

corrensite

Deep burial pore

filling chlorites

Fe-rich

Figure C30 Common occurrences of chlorite and related minerals in sediments.