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

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DIACENESIS

215

be substantially reduced by nimfuictioii and also through

partial or complete (illing by authigenic minerals (see Authi-

genesis),

a process known as

cementation.

With the possible

exception of detrital quartz and certain other minerals of high

chemical stability, primary detrital minerals are subject to

dissolution that can lead to the formation of

.secondary

iniiui^rantiUn

pores.

Secondary components, both aulhigenic minerals and pore

spaee.

are also subject to modifications during diagenesis.

Early-formed authigenic phases may dissolve to leave second-

ary pores as ambient temperature and fiuid composition evolve

to conditions unfavorable for the mineral's continued stability.

Authigenic phases precipitated during deep burial may be

prtine to dissolve when exposed to lluid fiow at low

temperatures following uplift.

Secondary pore spaces, within both primary and secondary

minerals, may become filled with authigenic minerals, in a

process known as

replaeenieni.

Some of the most notable

compositional and textural changes in diagenesis. for example

dolomitization and albitization. take place largely through

replacement. Replacement of primary or secondary minerals

by new crystals ofthe same gross mineralogy, lecrystatlization.

is notably diflicull lo document with complete certitude. This

difficulty arises because a single mineral may be present in both

primary and secondary forms or as co-existing authigenic

crystals representing multiple episodes of precipitation, yet

display only subtle heterogeneity in texture and composition.

Demonstration of a cementation versus a replacement (or.

recrystallization) origin for mierocrystalline components

(e.g..

clay minerals, mierite) is partieularly challenging because

heterogeneities in texture and composition that can prove a

primary versus an aulhigenic origin cannot be readily imaged

at submicron scales.

Physical processes in diagenesis

Physical processes are important across the entire range of

diagenetic conditions, but are prominently observed in the

early stages of burial, prior to substantial lithifieation. when

grain rearrangement by bioturbation and compaction are most

readily accomplished. Major categories of physical processes in

diagenesis include bioturbation. elutriation. compaction, and

tectonic delonnaiion.

Physical grain rearrangement in very early diagenesis of

unconsolidated sediment is an important control on the

subsequent path of chemical diagenesis because of the effects

on permeability and grain surface characteristics. Bioturhalitm

(sec Bioiicnic Sedimentary Slruclures) is responsible for

substantial grain rearrangement in certain depositional envir-

onments. Primary depositional fabrics

(e.g..

sorting, layering,

etc.) may be greatly rearranged and new fabrics established by

bioturbation. profoundly redueing the permeability of the

sediment and the rocks that develop from them if fine and

coarse detrital eomponents. for example, elay and sand,

become mixed. FJulriation is a process by which fine-grained

particles arc infiltered through pore spaces following deposi-

tion,

leading in some cases lo rocks with reduced permeability.

Thick grain coating clays introduced by eluiriatioii and

adhered tightly to grain surfaces by repeated wetting and

drying can greatly alter the path of diagenesis in the sediment

both by reducing overall permeability and also by creating a

partial barrier lo reaetion beiween pore lluids and grain

surfaces. This is one illustration of how physieal and chemical

processes in diagenesis are linked.

Compaction refers to the reduetion of primary intergranular

space during burial and deformation. Distinet from eementa-

tion.

compaction arises solely from the rearrangement of

detrital grains. In well-sorted sand-size sediments the inter-

granular volume (IGV). also known as the "minus-cement

porosity", is the sum of primary intergranular pore space and

intergranular cements. Volume loss by compaction is calcu-

lated as follows

(e.g..

Lundegard. 1992);

Compactional porosity loss =

P, - (((100- PJ*IGV)/(inO-IGV))

Al deposition, the initial porosity (P,) ranges from around 40

percent of sediment volume for sand-size materials to much

larger values (ranging to 80 percent.') for sediments with

abundant clay-size particles and sediments containing grains

with abundant primary intragranular

pores.

IGV reduced solely

by grain rearrangement is believed to reaeh

minimum value of

around 26 percent. Larger IGV reduction (to a maximum of

IGV =

is observed in both sandstones and limestones.

requiring the operation of other mechanisms of compaction

such as ductile deformation, grain crushing by brittle deforma-

tion (see Pressure Dissolution). Near-depositional values for

IGV can be preserved by early cementation, a phenomenon

that is more widespread in limestones than in sandstones.

Deibrmation at a variety of scales is increasingly recognized

as an integral element of diagenesis. In addition to prominent

grain crushing during burial compaction, grains may also

succumb to crushing by stresses applied tcctonically. especially

if the tectonism occurs while the rock is still relatively porous.

Planar zones of crushed and eemented grains known as

deformation bands arc prominent in some porous sandstones.

In rocks that are highly lithified. deformation may be through-

going,

cutting across detrital and authigenie components.

Substantial modifications of porosity, both increases and

decreases, can result from deformation. Fracture porosity is

one ot" the main categories of secondary porosity.

Chemical processes in diagenesis

Chemical processes in diagenesis involve interactions between

solid phases and fluids contained within the rock's pore spaces.

Such aqueous reactions, both dissolution and precipitation,

commence in unconsolidated sediment within Ihe depositional

environment and continue throughout ihc range of diagenetic

conditions into metamorphism. All reaction in diagenesis

manifests a relalive slowness of rate compared to reaetions

at higher temperatures. This tendency for reactions to be

sluggish has profound implications (br the kinds of mineral

assemblages that ean exist and also Ibr the kinds of textures

that ean arise in diagenesis.

It is reasonable to assert that overall chemieal equilibrium

between solid phases in rocks is not achieved in diagenesis.

Even if a phase is in gross equilibrium

(e.g..

detrital albite in

contact with albite-saturated water), the actLial crystal lattice

may still have trace element, isotopic. and structural disequili-

brium wilh ambieni conditions. Such disequilibrium represents

a sort of driving force that can ultimately lead to dissolution.

even in the presence of fiuids that are near equilibrium with the

bulk mineral. Transient loeal, microscale equilibrium between

an individual authigenic phase and pore lluid occurs in

216

DIACENESIS

diagenesis. at least theoretically, but does not apply more

generally to the bulk rock which contains a diversity of phases

that are not in equilibrium either with pore fiuids or one

another. The equilibrium that does exist is important, in the

sense that it is the basis for using isolopic and elemental

analysis of authigenic phases to interpret the compositions of

ancient fiuids responsible for precipitation.

Dissolution

A key factor in chemical diagenesis is nietaslabilily, which

refers to the tendency of mineral crystals to persist for

geologically long periods under conditions far outside the

range of their thermodynamic stability because the rate of their

reequilibration is slow. A metastable mineral is said to have a

"kinetic barrier"" to reaetion. Amazingly, some feldspars,

originally formed by crystallization from magma, can persist

for billions of years in aqueous environments at surface

temperatures. Some detrital feldspar persists even in Archean

sandstones, having survived erosion, transport, deposition,

and a prolonged history of diagenesis.

The degree of metastability differs greatly for different

sedimentary materials (both crystalline and amorphous) and

can vary greatly even among individual components of similar

bulk eomposition. for example among different grains of

K-feldspar in a sandstone or mudroek. Hence, there is a

tremendous range in the conditions required to initiate

dissolution of primary and secondary mineral components.

Most evaporite and many earbonate minerals manifest

extensive recrystallization (dissolution followed by repreeipita-

tion of a more stable crystal lattice) over short time periods

under essentially syndepositional eondilions. Highly unstable

materials may recrystallize repeatedly through a range of

diagenetic conditions and environments, leading to extensive

and eomplex textural and chemieal modifieations. Eor

example, dolomites typically retain almost no discernible

primary detrital textures at any scale, being so highly prone

to reerystallization. In a less extreme case, albitized detrital

feldspars may retain a gross sand grain morphology as seen al

low magnifications, but as observed on the scale of a few

micrometers are revealed lo be highly microporous aggregates

of the authigenie feldspar albile. In stark contrast with

dolomite and feldspars, detrital quartz fails to reaet with pore

fiuids and maintains its primary isotopic and trace element

composition up through the incipient stages of metamorphism.

Crystal defects are an important control on the mieroseale

localization of dissolution in diagenesis. In scanni[ig electron

microscope observation, dissolution of detrital minerals is

typically observed to be localized at pits that correspond to the

intersection of defects with the crystal surface. Important

textural implications arise from the defect control on dis-

solution. At the scale of a few 10s or 100s of micrometers,

dissolution of both primary and secondary minerals is typically

observed to be highly irregular. Crystals are not generally

observed to dissolve progressively from the outer surface, but

rather in patches, at sites anywhere within the crystal that

correspond to energetically favorable defects. Secondary pores

formed in this manner show a wide range of size, varying from

sub-micron micropores to pores of lOs to 100s of microns

surrounded only by skeletal remnants of the original grain.

Complete dissolution of detrital grains can lead ultimately

to the formation of oversized pores. Compactive collapse of

oversized pores can completely erase evidenee for the former

existenee of some detrital grains. Dissolution ean ultimaiely

induce a considerable loss of information on the nature of

primary roek components. In limestones for example, arago-

nitie alloehems that may have been abundant in the deposi-

tional environment are rare in rocks of all ages and have never

been observed in rocks older than Pennsylvanian. Such

alloehems are typically replaced by low-Mg caicite early in

their diagenetic history. In the sandstones known as diagetietie

quafizaienites. detrilal quartz is the only remaining unreacted

primary mineral, all feldspars havitig been lost to dissolution.

Precipitation

The sluggishness of precipitation reactions in diagenesis is

manifested as a difficulty with nnclcation. As a result, in both

cementation and replacement, authigenic minerals typically

show a highly localized spatial distribution, an effect that is

apparent at a variety of scales. Eor example, some minerals

show a marked tendency to nucleate only on the surfaces of

pre-existing (i.e.. mostly primary) crystals, a phenomenon that

is apparent on the seale of a few microns to hundreds of

microns. In essence, precipitation of such authigenic crystals

enlarges the earlier-formed crystal, forming an outer zone

known as an oyergromli. Authigenic quartz, feldspar, calcite.

and ankerite are all notable examples of this mode of

nueleation and growth. In the absence of appropriate

substrates, these minerals may fail to grow even in the presence

of fiuids of appropriate composition. Substrate control of

cement localization is observed in all types of porosity.

Coatings on the surfaees of potential nuclei, formed through

elutriation or early cementation, can markedly inhibit the

growth of minerals in later diagenesis.

Replacement can be thought of as cementation occurring

within secondary pores (however tiny) and, thus, replacement

minerals, too. tend have highly localized distribution. Replace-

ments tend to occur only at the sites of defect-con trolled

dissolution, as described above, and thus, typically have a

highly patchy distribution within the crystal that is being

replaced. As with dissolution, replacement that is localized

near the outermost crystal surfaces (halo texture) generally is

not observed in diagenesis.

An example of authigenic mineral localization on a some-

what larger scale (typically 100s of microns to meters) is the

formation diagenetie structures known as nodules (mineral

aggregates distinct from the host rock and typically formed

very early in diagenesis) and concretions (regions of prefer-

ential cementation, formed at any time in diagenesis). As with

overgrowths, this localization reflects, at least in part, the

(Jiffieulty of nueleation at low temperatures. Nueleation of one

crystal appears to have a dampening effect on further

nueleation in its vicinity, leading to subsequent precipitation

that is preferred (as overgrowths'.') on the early-formed

nucleus. The detailed mechanics of how concretions and

nodules are localized is one of the enduring enigmas of

diagenesis.

Interrelated physical and chemical processes

Some diagenetic processes lake place through intimately linked

chemical and physical processes. Eor example, a weak

connection between physieal and chemieal processes is seen

in localization of cementation in fraeture porosity (vein

(brmation) which takes place partly as a result of generation

DIACENESIS

217

of fresh, uncoated surfaces favorable to nueleation and partly

as a result of enhanced fluid fiow. A stronger link between

ehemieal and physical processes is seen in the ease of pressure

solution (see Pre.ssureDis.wliiiion) including both intergranular

pressure solution and the formation of stylolites. In this

instance, dissolution is localized along certain crystal contacts

(in evaporites. carbonates, and quartz-rich sandstones),

lead-

ing to prominent porosity loss and developmeni of sutured

boundaries. A variety of mechanisms have been proposed for

the dissolution step in the pressure solution phenomenon, bul

the stress field is, in every ease, the fundamental determinant of

form and orientation of the dissolution surfaee. In simple

burial eompaciion sutures are oriented parallel to bedding

(perpendieular to the principal, compactional stress), whereas

in tectonic deformation, stylolite orientation may be deter-

mined by the orientation of tectonic stresses. Pervasive

tectonieally driven stylolite development in late diagenesis.

grades, ultimately, into metamorphic (bliation. Another

intriguing example of closely linked physical and chemical

proeesses is dissolution driven by force of crystallization. In

eertain circumstances it appears that nueleation of one

crystalline phase can actually induce the dissolution of

surrounding crystals, an alternative mechanism for over-

coming the kinetic barrier to mineral dissolution, leading to

replacement (Sorby. 1863; Weyl, 1959; Maliva and Siever.

1988).

Fluid flow and diagenetic heterogeneity

An on-going controversy in diagenetie studies concerns the

degree to which diagenetic processes of dissolulion. cemcnta-

titin.

replacement, and recrystallization are accomplished in so-

called "open"" versus "closed"' chemical systems. The opening

volley of this debate was the paper by Land and Dutton (1978)

which demonstrated a disparity of tremendous magnitude

between the amount of quartz cement in a Pennsylvanian

sandstone and the amount of silica available from pore fluids.

given the fluid fiow limitations ofa simple burial compaction

model.

Similar conundrums immediately became apparent in

connection with many instances of cementation, dissolution,

and replacement in all types of sedimentary rocks, espeeially

with regard to the transport of eomponents such as Al which

are relatively insoluble in pore lluids and also components of

volumetrically significant reactions in the deep subsurface. In

general,

studies that focus on petrographic documentation of

authigenic reactions find evidence for a greater degree of

"openness'" in terms of the amounts of material transport

required to accomplish the observed amounts of cementation

and replacement. Studies that focus on rock permeability, fluid

fiow. and fiuid compositions generally Hnd potential for lesser

amounis of material transport.

One of the obstacles to resolving this controversy concerns

the tendency of diagenetic processes to manifest a high degree

of spatial heterogeneity, making it difficult to obtain statisti-

cally meaningful data on sedimentary rock compositions.

Within a given rock unit, the degree of cementation or the

degree of replacemenl can vary due to the controls on reaction

localization as discussed above as well as from historical

vagaries in primary composition and tkiid flow. As a resull of

compositional heterogeneity it is difficult to ascertain quanti-

tatively whether materials apparently lost from one volume of

rock have been gained by another. Likewise, some workers

have called inlo question whether our understanding of

permeability in the subsurface is sufficiently complete. Studies

to determine the appropriate scale for rock property observa-

tions are continuing as are efforts to better understand the role

of endolithic microbes in element mobilitv in the subsurface.

Microhial influences in diagenesis

Microbial activity is recognized as an important element o(

diagenetic processes in a wide range of subsurface environ-

ments

(e.g..

Banfield and Nealson. 1997). Mineral surfaees

within endolithic pore spaces provide attachment sites and

nutrient sources to bacteria, archaea. and fungi across

environments extending (Vom ncar-syndepositional lo several

kilometers depth. Microbes affiliated with mineral surfaces

can,

either actively or passively, be a significant factor in

overcoming kinetic barriers to both dissolution and precipita-

tion in diagenesis. The mere presence of microbial commu-

nities within diagenetic environments may exerl controls on

the microscale morphology and distribuiion of authigenic

precipitates because biofilms exuded by living eells change

mineral surfaces in ways that can promote, inhibit, or modify

nucleation.

There is currently vigorous research activity focused on the

microscopy and geochemistry of microbe/mineral interactions

seen in laboratory and field experiments, in modern surface

and subsurface environments, and in ancient roeks. These

studies are beginning to contribute a view in which ihe

diagenetic realm is also a vital region of the biosphere on the

Earth,

and perhaps, other planets.

Boundaries hetween diagenesis and depositional

processes, weathering, and metamorphism

No sharp division between diagenesis and metamorphism. nor

between diagenesis and weathering, can be readily specified.

Historically, various temperature/pressure conditions, textural,

and mineral assemblage criteria have been applied in making

distinctions between these dilTerent realms of rock alteration.

All of these apprt)aches have been rendered somewhat obso-

lete by the growing recognilion of the role of initial

composition on textural modification and the fact that many

reactions and mineral assemblages once placed within the

zone of metamorphism

(e.g.,

zeolite precipitation, albitization)

are now known lo occur within rocks that are clearly

sedi-

mentary. Rocks might be considered to be "metamorphic"

once their sedimentary tcxturai components (grains, cements,

and pore spaces) have been "erased"". Development of

through-going metamorphic fabries sueh as pressure shadows

or foliation is clearly a marker for ihe transition out of

diagenesis. Of eourse such degrees of textural alteration are

achieved under very different conditions for materials of

different starting compositions. Eor example, limestones and

espeeially dolomites may manifest a high degree of textural

and chemical modification by one kilometer of burial (<30'C)

whereas an interbedded quartz-rich sandstone may survive to

nearly 300 C with clearly sedimentary textures intact. The

method of observation is another control on such a

determination

(e.g.,

rocks that appear as metamorphic

"quartzitcs"" in light microscope observation, may still appear

as "quartz-cemented sandstones" in cathodoluminescence).

Many proeesses identiiied as "diagenetic""

(e.g..

albitization.

quart7 and feldspar preeipitation in fractures, clay mineral

218

DIAGENESIS

formation) also take place within crystalline rocks at relatively

low temperatures.

Both weathering and metamorphism encompass physical

and chemical processes, involve aqueous reactions, and assign

great importance to melaslability. ln metamorphism there is a

role for solid state ditTusion that is not observed in diagenesis

(halo textures do occur prominently in metamorphic roeks)

and also a degree of solid phase equilibrium. In weathering

there is a physieal decomposition of roeks thai is not affiliated

with diagenesis.

Given that the boundaries between diagenesis and meta-

morphism and between diagenesis and weathering are so

obscure, it is especially difficult to subdivide diagenesis in a

meaningful and consistent way (Fairbridge, 1967). Diagenetic

processes are often divided loosely on the basis of their timing

relative to eompaetion. those processes occurring nearest the

depositional interface, prior to substantial compaction, and at

temperatures

KSO'C

being designated "early", whereas pro-

cesses occurring in deeper burial, tbllowing substantial

eompaetion. and at temperatures around 80 C and higher,

are designated "late"'.

In general, chemical processes are assigned to diagenesis

more broadly than physical processes. Eor example, dissolu-

tion of unstable minerals during transport or through

bioerosion (chemical boring) would be widely considered as

elements of diagenesis whereas bioturbation affecting uncon-

solidated sediments within the depositional environment is

typically studied by experts familiar with deposition rather

than diagenesis. Similarly, growth of clay minerals through

replacement of feldspars in weathering would be considered

diagenetic whereas physieal disaggregation of rock by wind or

ice would be most eommonly thought of as weathering.

Methods

for the study of diagenesis

Documentation of the primary versus secondary character of

sedimentary rock components is accomplished through the

integrated use of microscopical and geoehemical techniques.

The pervasive chemieal disequilibrium that characterizes rocks

throughout diagenesis lessens the importance of bulk methods

of chemical analysis in diagenetic studies in comparison to

studies of igneous and metamorphie systems. Diagenetic

studies over the past several decades have benefited greatly

from advances in imaging and spatially resolved chemical

analysis. Such methods allow characterization of the micro-

scale chemical heterogeneities that persist throughout the

realm of diagenesis and permit a correlation of such

heterogeneities with information on the sequential develop-

ment of rock components

(e.g..

eement stratigraphy).

A technology of particular importance in diagenetie studies

is the use of colored impregnation media, first applied in the

1970s, that allows reliable discrimination of natural versus

artificially induced porosity in thin sections. Widespread use of

this technique in limesiones and sandstones led to a revolution

in understanding ofthe importance of grain dissolution in the

subsurfaee and also inspired a renaissance of data collection

that greatly improved understanding of compaction. Imaging

with scanning electron microscopes and electron microprobes

created a vastly greater appreciation of the heterogeneity that

exists for any given mineral phase across a range of

components (grains, cements, and replacements) within

sedimentary rocks. Eor example, scanned cathodolumines-

eence imaging and back-scattered electron imaging may permit

discrimination of grains, cements, and grain replacements all

composed of the same mineral phase, but having small

differences in defect structure or trace element composition.

Likewise, different generations of

single authigenic phase can

be identified on images that arc sensitive to subtle differences

in trace element eomposition. Such distinctions may be highly

obscure in other modes of imaging such as transmitted light

microscopy that only allows discrimination between phases

or dilTerent crystal orientations of a single phase. Diagenetic

studies typically employ a combination of imaging (for

historical aspects) and analysis (in order to quantify T and fluid

compositions). Important petr{igraphic methods in diagenetic

studies include transmitted plane- and cross-polarized miero-

scopy. secondary electron imaging, back-scattered electron

imaging.

X-ray mapping, and cathodoluminescence. Important

analytical approaches include X-ray diffraction, major and

trace element analysis by electron microprobe, and analysis of

the stable isotopes of O. C, and Sr. Sampling of detrital and

authigenic phases by microdrilling teehniques has been widely

employed in combination with isotopic analyzes.

Diagenetic

studies

Published work on diagenesis is extensive and is summarized in

numerous textbooks, review papers, and reprint collections

(e.g..

Pettijohn. Potter el at.. 1973; Bathurst. 1975; Peltijohn.

1975:

Berner. I9S(); Eolk, 1980; Land. 19K5; Horbury and

Robinson.

1993; Giles. 1997: Land. 1997). Diagenetic s'tudies

are found primarily in mainstream geological journals as well

as in publications in the areas of petroleum and environmental

engineering, chemistry, and physics.

Summary

Diagenesis encompasses all physical and chemical modifica-

tions that affecl sediments and sedimentary rocks following

deposition and prior to the advent of weathering or

metamorphism. The tendency of minerals to manifest sluggish

reequilibration reactions under diagenetic conditions is re-

sponsible lor the preservation in sedimentary rocks ol'

considerable historical information concerning both the

origins of delrital grains and the progress of chemical reactions

involving dissolution, cementation, and replacement. Even

though reactions in diagenesis are comparatively slow, modern

diagenetic studies have concluded that most primary compo-

nents in sediments are ultimately subject to profound

modifications during burial as metastable components react

with the surrounding pore fiuids. The microbial role in

diagenetic reactions is an exciting area of research lhat is

eontribuling important new perspectives on the nature of

diagenesis.

No absolute boundaries are assignable between diagenesis

and either weathering or metamorphism. but rocks in the

diagenetic realm are characterized by pervasive chemical non-

equilibrium between the various solid phases that comprise the

rock and, lo a high degree, nonequilibrium between the solid

phases in the rock and the contained pore fiuids. With the

exception of so-called chemical sediments that are constituted

almost entirely of authigenie precipitates, sedimentary roeks in

the diagenelic realm manilest a clearly sedimentary fabric with

distinct grains, pores, and cements. Erasure of such fabrics is a

hallmark ofthe transition into metamorphism.

Kitty L. Milliken

DIACENFTir STRUCTURES

219

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2001-2025.

Cross-references

DIAGENETIC STRUCTURES

The results

diagenetic alterations

arc

visible whenever

examine sedimentary rueks

in the

field. Indeed,

the

very

hardness

of the

rocks

is the

result

of the

diagenetic process

lilliitkation.

whieh loose sediment

transformed into solid

rock

cementation

and

compaction. However, there

arc

numerous structures that

can be

readily observed

outcrop

that result wholly from diagenetic alterations. These diagenetic

structures include features such

concretions, cone-in-cone

structure, geodes. liesegang bands, reduction features, sand

crystals,

and

slylolites.

these, concretions

and

stylolites

are

fur the

most cotnmon

and

have been intensively studied,

where;ts

ihe

other features, although locally ubundanL

are

relatively rare

and

have received less attention

in the

literature.

Concretions

are

hard masses

sedimentary

and.

more rarely,

fragtnental volcanic rock, that form

by the

preferential

preeipitation

minerals (cementation)

localized portions

ofthe rock. They

are

commonly subspherieal (Figure

DIS).

but frequently form

variety

other shapes, including disks,

grape-like aggregates,

and

eomplex shapes that defy descrip-

tion. Concretions often form around

nucleus, such

as a

fossil

fragment

piece

organie matter. This occurs because,

for a

variety

reasons,

the

nucleus created

more favorable site

for

cement precipitation than other sites

in the

roek. Concretions

are usually very noticeable features, because they have

strikingly different color and/or hardness than

the

rest

of the

rock.

The

term nodule

often used

as a

synonym

for

concretion,

but

some workers prefer

reserve

the

term nodule

for concentrations

authigenic minerals that

do not

enclose

detrital grains (Selles-Martinez,

1996) or are

relatively small

and largely

replacement origin (Bates

and

Jackson, 1987).

Concretions often preserve features

of the

original sedi-

ment such

burrows, fossils,

and

sedimentary layering

that cannot

seen

in the

rest ofthe rock (e.g,, Kidder. 19^5).

This occurs because concretions usually begin

form

relatively early, before sediment compaction

and

other

processes disrupt

the

original sediment. Thus, cementation

Biogenic Sedimentary Struclures

C'alhodoluminescence

Cements

and

Cemenlalion

Compaction (Consolidation)

Sediments

Dedolomilizalion

Deformallon

Seditnenis

Diagenetic Structures

Dolomites

and

Dolomiiizalion

Feldspars

Sedimentary Rocks

Geodes

Isolopic Methods

Scdimcnlology

Maturaiion, Organic

Micritizalion

Neomorphism

and

Recrystallization

Porewaters

Seditiienls

Seplarian Concretions

Styloliies

Figure

D18

Contrefion

shale.

LL'HS

<.<ip

lor

stale.

220

fJIACENETIC STKLK TURES

the concretions "freezes" the early sediment structure, forming

rigid mass that resists later chemical and mechanical alterations.

Concretions most commonly result from carbonate pre-

cipitation,

although a variety of minerals also can form

concretions, including pyrite, apatite, iron oxides, and silica (as

opal,

chert, or chalcedony). Ofthe carbonate minerals, calcite

most commonly fomis concretions, followed by dolomite, and

siderite.

Carbonate and iron oxide concretions arc most

commonly found in shales and sandstones, whereas apatite

and pyrite concretions are mainly found in shales, and silica

concretions normally form in limestones and shales. Carbo-

nate concretions, given their abundance, and the ease of

obtaining traee-elemcnt and isotopic data from them, have

been studied in much greater detail than any other type. These

studies have provided a wealth of information concerning the

nature of early diagenetic reactions, and have conclusively

demonstrated the importance of the microbial o.xidation of

organic matter in the early diagnetic environment

(e.g.,

Raiswell and Fisher. 2000).

Concretions grow either by precipitation between grains in

the host rock (usually the case for carbonates and apatite) or

by replacement of the host rock (usually the case for chert). In

either case, the conventional model for concretion growth is

concentric, in which the authigenic mineral is progressively

added (by replacement or direct precipitation) to the outer

edge ofthe growing concretion. Consequently, early zones are

at the center and late zones near the outer edge

(e.g..

Raisweil.

1971).

Most studies of concretions have assumed this mode of

growth.

However, recent studies indicate that concretion

structure and growth often does not follow the concentric

model.

Small-scale zonation incompatible with such a mode of

cementation has been observed in caleite, siderite, dolomite,

and chert concretions (Mozley, 1996: McBride

elal.,

1999;

Raiswell and Fisher. 2000). In some cases, it is clear that late-

stage and early-stage cements precipitated throughout the

concretion,

indicating that early cementation was incomplete,

and that signilicant porosity and permeability remained to

allow precipitation of late-stage cements in the concretion

interior.

Some concretions are cut by a radiating network of fractures

(septaria) filled with coarsely crystalline calcite and other

minerals. Such concretions are known as septarian concre-

tions,

and a variety of hypotheses have been proposed for the

formation ofthe fracture systems, including: (1) Shrinkage of

the interior of the concretion, a result of the dehydration or

"chemical desiccation" of either a clay-mineral-rich center

(Richardson,

1919; Raiswell. 1971). or dehydration of a gel-

like precursor (Lippmann, 1955). (2) Expansion of the

concretion interior and expansive growth of the fractures

(Davies. 1913;

Todd.

191.1), presumably accomplished by the

force of crystallization. (3) Development as tensile fractures in

response to stresses induced during burial (Astin. 1986). (4)

Fracturing induced by propagation of seistnie waves through

poorly lithified sediment (Pratt. 2001). Of these various

hypotheses, the shrinkage hypothesis remains the most

favored.

However, the geoehemical conditions necessary for

significant dehydration of clay mineral centers have not been

quantitatively explored. The gel-like precursor model has the

problem that the nature of the gel is poorly-defined (some

workers have proposed an organic soap), and such gels have

never been observed in modern concretions. Thus, the origin of

septarian fractures is currently a topic of debate and ongoing

research.

Recently hydrologists have becotne interested in elongate

concretions. These concretions range in size from pencil- and

cigar-like bodies to those that resemble large fallen logs. They

are thought to form from flowing groundwater. with the long

axis of the concretion oriented parallel to the groundwater

flow direction (McBride

etal.,

1994; Mozley and Davis, 1996).

Elongate concretions have even been found in faults, where

they reeord the past flow orientation of groundwater in the

fault zone (Mozley and Goodwin, 1995).

Cone-in-cone structure

Cone-in-cone strueture consists of nests of cones of authigenie

minerals, one inside another, which form either in thin beds, as

veins,

or at the margins of carbonate concretions (Figure DI9).

Concretions that consist almost entirely of cone-in-cone

carbonate are also observed. In such concretions, the apices

of the cones invariably point toward a shale core that bisects

the concretions

(e.g..

Franks. 1969). Individual cones are often

separated from one another by a thin layer of clay. The

structure mainly forms in shales, but has been reported for

argillaceous sandstones as

well.

Most cone-in-cone is com-

posed of calcite. but it has also been observed in. siderite.

gypsum and pyrite (Carstens. 1985).

The carbonate material making up the cones has variable

characteristics. It is often fibrous calcite (6 microns or less in

diameter, but up to

long;

Franks, 1969), however.

microcrystalline calcite has also been reported. This miero-

crystalline calcite is generally interpreted to represent recrys-

tallized fibrous ealeite, based upon the presence of relict

textures (Franks. 1969). When present in sandstone, the cone-

in-cone structure is often thinner and less distinct than in

shales.

A variety of mechanism have been proposed for the origin

of cone-in-cone structure. These include; (I) Volume increase

accompanying inversion from aragonite to calcite (Tarr, 1922;

Figure D19 Cone-in-cone structure. Coin for scale.

DIACENETIC STRUCTURES

221

Cjilman and Metzger. 1967), in which expansion of conical

aragonite pushed cones apart and allowed clay intrusion;

(2) burial-induced pressure solution, with clay films originat-

ing as insoluble residues (Tarr. 1932); (3) fracturing of

crystalline aggregates that grow in overpressured chambers,

with the fractures induced by a decrease in pore pressure

(Selles-Martine/, 1994); and (4) formation during early

diagenesis by expansive mineral growth

(i.e.,

force of crystal-

lization), in which the cones are produced by the growth of

cone-shaped aggregates of fibrous calcite (Franks, 1969), and

the clay films originate as the crystals displace and disturb the

original clay-rich sediment. By the latter mechanism, the

orientation of the growing crystal is assumed to be controlled

by the local stress configuration (direction of easiest crystal

growth). Of these hypotheses, the displacive crystal growth

mechanism has proven to be the most popular, thought the

debate is by no means settled.

The inferred displacive nature of the crystal growth, and

common deflection of host rock laminae around cone-in-eone

concretions has led most authors to conclude that much ofthe

precipitation occurs very early, during shallow burial. How-

ever, some authors have concluded, largely on the basis of "^O-

depteted values of some ofthe eone-iii-eone material, that they

can form later, at perhaps hundreds of meters of burial depth

(overpressuring would be necessary to allow displacive growth

al such depths': Marshall. 1982; Pirrie

eial..

1994),

Geodes

Cieodes arc subspherieal accumulations of authigenic minerals

(mainly silica minerals and to a lesser degree caleite and

eetestite) whieh occur in a variety of rock types, particularly

limestone and shale (Figure D20). They are superficially

similar to ci>neretions. However, unlike concretions, which

contain considerable detrital material, they consist almost

entirely of authigenic minerals. Geodes typically have a thin

outermost layer of dense chalcedony and a hollow interior, in

which well formed crystals project into a central cavity. They

can also be completely filled with authigenic minerals (some

authors use the term "filled geode'" for such features),

including a variety of exotic minerals, such as millerite.

Figure D20 Silica geode. Eihrous mineral in interior is millerile. Coin

tor scale.

anatase, ehalcopyrite. and galena (Goldstein and McKenzie.

1997). Some geodes even eontain oil in the central cavity

(Anonymous. 1938).

The term geode is sometimes applied to features in volcanic

rock, including basalt (precipitated in gas cavities; Garlick and

Jones, 1990). though some authors indicate that such features

are more properly termed amygdutes and that the term should

be restricted to features observed in sedimentary rocks

(Chowns and Elkins, 1974). Thunder eggs are related

structures of spherical or ellipsoidal shape, which eontain

cores of chalcedony and sometimes quartz. They are com-

pletely filled with minerals

(i.e.,

contain no internal cavity) and

are found exclusively in welded tuffs and rhyolites (Staples.

1965).

Numerous theories have been proposed to explain the origin

of geodes. including: (I) Precipitation within dissolved fossil

material,

particularly sponges (Van

Tuyl,

1916); (2) a complex

process of replacement and solution/precipitation alteration of

calcite eoncretions (Hayes, 1964); and (3) dissolution and

replacement of evaporite nodules. Of these, hypothesis-3,

alteration of evaporite nodules, is the most commonly cited

meehanism. Indeed, many workers use the term "silicified

evaporite nodule" rather than geode to describe these

structures

(e.g.,

Milliken. 1979; Ulmer-Scholle et cd., 1993;

Ulmer-SchoHe and Scholle. 1994). Geodes have been reported

in rocks that originated in sabkha environments (Chowns and

Elkins. 1974). however, they can also form in "normal marine""

sediments that have experienced circulation of concentrated

brines (Maliva. 1987). Geode formation

(i.e.,

silicification) can

apparently proceed during early diagenesis

(e.g.,

prior to

compaction; Chowns and Elkins, 1974), as well as during late-

stage diagenesis in association with meteoric influx and

hydrocarbon migration (Ulmer-Scholle

eial..

1993).

Liesegang batiding

Eiesegang bands were iirst described by Eiesegang (1896) in

experimental precipilates. His original experiment involved

placing a drop of silver nitrate solution on a surface of

potassium dichromate gel. This resuitcd in the formation of

concentric rings of silver dichromate. Eiesegang and subse-

quent workers noted similar features in rock, which are now

referred to as Eiesegang bands or rings. These features consist

of bands of authigenic minerals arranged in a regular repeating

pattern.

They typieally cut across-stratification and occur in a

variety of rock types including sandstone and chert. Despite

the frequent occurrence of Eiesegang bands in rocks (Merino.

1984), only a few authors have examined their mineralogy and

texture in any detail (Fu

eial.,

1994). Eiesegang bands are

frequently cited as prime exainples of geoehemical self-

organization, in that their distribution in the roek often does

not seem to be directly related to preexisting features (Chen

etal.,

1990).

The exact mechanism of band formation is poorly under-

stood and remains an area of active research

(e.g.,

Krug

etal..

1996). A commonly cited meehanism is precipitation in

a Ostwald-Eiesegang supersaturation-nucleation depletion

cycle.

By this inechanism dilTusion of reactants leads to

supersaturation and nucleation; this precipitation results in

loealized band formation and depletion of reactants in

adjacent zones (DeCelles and Gutschick, 1983; Sultan

etal..

1990).

222 DIAGENETIC STRUCTURES

Reduction features

Gray-green reduction features are common in many redbeds.

They occur as isolated, subspherieal to irregular reduction

spots,

and as reduetion zones associated with fractures, faults,

and permeable sandstone beds (Miller. 1957; Shawe. 1976;

Chan etal., 2000). These features result from localized reducing

conditions in the pore water, most commonly resulting from

the presence of organic matter (in the case of reductions spots).

or migration of reducing groundwater along zones of higher

permeability (tn the case of reduction /ones). The exact origin

of the reduced areas is poorly understood. Some authors

ascribe the phenomenon mainly to reduction of ferric to

ferrous iron, without net transport of iron from the reduced

area (e.g.. McBride. 1974; Mykura and Hampton, 1984).

Others note either no change in ferric/ferrous ratio across

reduced areas and contend that they are the result of Fe

transport (e.g.. Manning, 1975). or they describe changes in

both ferric/ferrous ratio and total iron content (Keller. 1929;

Picard, 1965; Durranee etal.. 1978). Significant mineralogical

differences occur in association with these features, including

dissolution of iron oxides and ilmenitc in the reduced portion

and preeipitation of authigenie feldspars and biotite (Miller.

1957;

van de Poll and Sutherland. 1976). Some traee-element

results suggest that Fe, Ca. Cr, and Nb are leached, and Co.

Ni,

V, and Y may be added during production of reduced

areas (Shawe, 1976).

Sand crystals

Satid crystals are large (1 to >50cm long) euhedral to

subhedral poikilotopic crystals of authigenic minerals (com-

monly barite. gypsum or calcite) that formed in sandstone

(Figure D21). Such crystals commonly have well-developed

crystal faces despite the fact that they can eontain greater than

50 perceni included sand grains. The percentage of sand wilhin

the crystals is variable, and refiects the mode of growth of the

crystal. Passive precipitation of poikilotopic crystals results in

intergranular volumes (IGVs) thai are eomparable to that of

uncompacted or slightly compacted sand (i.e.. 40 50 percent).

Expansive crystal growth results in much higher IGVs and

disrupted fabrics (Dapples. 1971; McBride etal., 1992). In

extreme eases, expansive growth ean completely exelude sand

grains from the crystal's interior. Some crystals contain

alternating zones of greater and lesser sand percentages, which

are interpreted to reflect alternating passive and displacive

growth. This is thought to result from alternating fast and slow

crystal growth, with exclusion of sand grains occurring during

periods of slow growth (McBride etal.. 1992; Shearman. 1981).

Stylolites

Stylolites are Irregular serrated surfaces in sedimentary rock.

They frequently consist of interlocking finger- and column-like

extensions that are subparaliel to one another and perpendi-

eular to the trend of the surf^ice. This can give stylolites an

irregular zig-zag appearance when viewed in cross section

(Figure D22). They form by a process of pressure dissolution.

in which portions of minerals under relatively high stress, sueh

as at grain contacts, undergo preferential dissolution due to a

pressure-induced solubihty increase. Stylolites form perpendi-

cular to the maximum compressive stress, which is commonly

vertically oriented, due to overburden pressure, resulting in

horizontal, bedding parallel stylolites. In areas that have

experienced tectonism. however, the principal compressive

stress is commonly horizontally oriented, resulting in sub-

vertical stylolites that cut across-bedding. Sutured grain con-

tacts are similar to stylolites, but unlike sutured contacts,

stylolites are large-scale features, which traverse numerous

grains and often entire outcrops. They are most commonly

observed in carbonate rocks, which are particularly susceptible

to pressure solution, but are found in sandstones and other

rock types as well. Undulatory surfaces that show evidence for

pressure solution, but arc not serrated in cross section, are

refc-rred to as solution seams (or wispy pressure solution seams).

Lithologic heterogeneity appears to be the main variable

controlling stylolite morphology (Andrews and Railsback.

1997).

If the rock is relatively homogeneous, dissolution

proceeds equally along the stylohte and a linear or undulatory

seam develops. High heterogeneity results in unequal dissolu-

tion, with variable amounts of dissolution occtirring along the

stylolite due to differences in solubility. This results in a '"zig-

zag"" columnar appearance. As dissolution proceeds, the

nature of the heterogeneity may too. resulting in changing

stylolite morphology through time.

Stylolites commonly contain a concentration of clay-rich

material, which is assumed at least in part to be a build-up of

Figure D21 Cdlcile s^nd crystals, ("oin lor

Figure D22 Cr<)s'^-s('( tiniial vit'w ol several stylnlites. Pencil for scale.

DIACFNFTIC STRUCTURES

insoluble minerals from the host roek that underwent

dissolution. Thus, as dissolution proceeds along the stylolite.

this insoluble residue becomes thicker. One cannot simply

assume that all the insoluble material along a given stylolite

originated in this way, because in some cases much ofthe clay

may have originated as a depositional clay seam, which

subsequently became a stylolite. Some authors have proposed

that clays promote pressure solution by providing a pathway

for ion diffusion (Weyl. 1959). Andrews and Railsback (1997).

however, demonstrated that stylolites can form in limestones

with as little as 2.1 wt percent insoluble residue and con-

cluded that abundant elay is not a necessary prerequisite for

styloiite development. Meike and Wenk (I98f^) provided TFM

evidence suggesting that some of the clay in these seams is

authigenic.

There are two competing hypotheses for how pressure

solution proceeds along the stylolite: the undercuiling hypoth-

esis and the solution film hypothesis. The undercutting

hypothesis (Bathurst. 1975) proposes that dissolution occurs

initially at point contacts resulting in preferential dissolution at

the contact, which eauses the grains to beeome closer to one

another and transforms the point contact to a surface contact.

This is followed by undercutting ofthe surface contact at the

edges (water being absent along the surface contact), which

results in undercutting and eventually in the formation of

another point contact. The solution film hypothesis proposes

that a contiguous filrn of liquid separates the two grains in a

manner similar to that operative during displacive crystal

growth (Weyl. 1959). This water is assumed to be under higher

pressure at grain contacts and selective dissolution proceeds

along the contact as a result. The second hypothesis (thin film)

is currently in favor with most workers (Morse and Mackenzie,

1990).

As dissolution proceeds, ions migrate away from the site of

dissolution cither along the stylolite or away from the stylolite

surface. Bathurst (1975) notes that much ofthe material must

leave the stylolite, migrating into overlying or underlying host

rock in order for dissolution to proeeed. This material

commonly reprecipitates near the stylolite

(e.g..

Wong and

Oldershaw. 19X1). The combination of cementation adjacent

to stylolites. and the build-up of low-permeability ciay-rich

insoltible residue (discussed above) can produce a very low-

permeabilily zone, that can act as a barrier to petroleum

migration and groundwater flow.

Loss of material by dissolution along stylolites can have a

significant impact on the sedimentary record. Bathurst (1975)

indicates that reductions of vertical thickness due to stylolite

growth of 20-35 pereent is common, and that crustal

shortening accomplished by tectonically generated stylolites

ean exceed that caused by folding and other deformation

mechanisms.

Summary

Diagenetic structures are macroscopic features that are the

result of diagenelic alteration ofthe host rock. The principal

diagenetic structures are concretions, cone-in-cone structure,

geodes, liesegang bands, reduction features, sand crystals, and

slylolites.

Concretions are concentrations of authigenic minerals

(mainly carbonate, silica, apatite, and iron oxides) that most

commonly form in shales, sandstones, and limestones. They

originate either by filling pores between framework grains or

by replacement of the host roek. Concretions form either

concentrically, in which concretion growth initiates at the

eoncretion core and material is progressively added to the

exterior of the growing concretion; or in a complex manner,

involving precipitation at multiple nueleation sites throughout

the concretion body. Septarian concretions contain internal

radial fracture networks that may result from internal

shrinkage. Elongate concretions have been used to map t)ut

palcogroundwater flow orientations.

Cone-in-cone structures consist of nested cones of authi-

genic minerals, most commonly calcite, that are separated

from one another by thin films of elay. They occur mainly in

shales, but have also been reported in argillaceous sandstones.

Numerous explanations have been put forth to explain these

structures, with the tnost popular being formation through

expansive crystal growth ofthe eones.

Geodes arc subspherieal concentrations of authigenic

minerals (most commonly silica) that oecur mainly in shales

and limestones. They typically have a tine-grained external

layer of chalcedony and an interior of coarser crystals that

project into a cavity. Most appear to have formed by

dissolution and silicification of evaporite nodules, though

other origins are possible.

Liesegang bands are nested zones of atilhigeiiic minerals

arranged in a regular repeating pattern. Their origin remains

poorly understood, however, a commonly cited mechanism is

precipitation in a Ostwald-Liesegang supersaturalion

nucleation depletion cycle.

Reduction features occur in redbeds as gray-green sub-

spherieal "'spots'* and elongate "zones"" that are commonly

associated with fractures. They form due to the presence of

locally reducing pore waters.

Sand crystals are large, well-formed poikilotopic crystals in

sandstone. They are most commonly composed of gypsum,

barite,

or calcite. Some grow passively, simply filling available

pore space in the sand, whereas others grow displacively,

excluding sand grains and disturbing the detrital fabric.

Stylolites are irregular, serrated surfaces that originate

ihrough pressure-induced dissolution of the host rock, most

commonly a limestone. The surfaces lypically contain a

buildup of clay-rich material that is commonly assumed to

represent an insoluble residue. The morphology ofthe surfaee

is largely governed by the degree of heterogeneity of the host

rock, with more heterogeneous lithologies having more

serrated and irregular stylolites. Stylolites form perpendicular

to the principal eompressive stress axis, resulting in subhor-

izontal stylolite formation during normal burial and subver-

tical stylolites of tectonic origin. Stylolite formation accounts

for the loss of considerable thicknesses of host rock by

dissolution. The material removed by dissolution is precipi-

tated adjaceni to the stylolite in some eases, but ean also

migrate great distances before reprccipitating.

Peter Mozley

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224 DIACENETIC STRUCTURES

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