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

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COASTAl SEDIMENTARY FACIES

Fine sand,

planar lamination

Fine-

nnedium sand,

longshore or

seaward

crossbedding;

Inclined planar

lamination

Fine-very fine sand,

SCS,

HCS.

Bioturbation

Interbedded fine sand

and shale, HCS in sand,

bioturbation

Shale,

interbedded

very fine sandstone,

btoturbation

- 1

- -1

-2

-3

- -4

-5

-6

•-7

Beach

foreshore

Upper

shoreface

(barrtrough)

Lower shoreface-

inner shelf

Inner shelf

Figure C63 Progradalion accomptinit'd by a tail in sea levtl (kirted

regression) occurs over a regressive surface of marine erosiun, cuf as

wave base is lowered.

NONMARINE BACIiSHOfiE

SHOREFACE/BARRIER

SHELF

RAVlNEMEhft SURFACE

Figure C61 Shoreface succession on a prograding fine-grained, non-

()ul)l)ly, moderattf-low energy coast tCretaceous, Book Cliffs, Utah).

Figure C64 Transgressive erosion caused by the lateral translation of

the shorefate (see Brunn, 19f)2). None of the Iransgressive coastal

facies are preserved, .ind the transgression is recorded by an erosional

("ravinemenl") surface overlain by deeper water (typically shelfl

deposition.

Fine sand,

root structures

Planar parallel

laminated fine sand

Bioturbated nne sand

Bioturbated

tine sand

and sets of

hummocky

cross-stratit teat ion

Bioturbaled very

fine sand

h 2

•-7

Backshore

Beach

foreshore

Upper

shoreface

Lower shoreface-

Inner shelf

Figure C62 Shoreface succession on a prograding very fine-grained,

low-energy coast (Eocene, WesI Texas).

systems. With less sediment cotitributed to the coast, marine

transgression occurs, whereby the shoreline shifts landward.

Where the sand supply is limited, landward transport by

shoaling waves and coastal winds concentrates it in a

shoreraee-t^orediine ridge system. In most wave-dotninated

settings, enough sand exists thai a stnall rise in relative sea

level is met by an upbuilding of the shoreline to a

topographically higher position, lf the sediment input is

sufficiently reduced, a basin develops behind the topographi-

cally high beach ridge and a barrier island is created. Most

present-day eoasts fronted by barrier islands are in a state of

transgression.

Barrier islands formed during transgression have a limited

potential for preservation. The landward translation of a

shoreface erases the deposits that lie landward from its base,

leaving an erosional {ravinctnent) surface overlain by shelf

faeies (Figure C64). Commonly this surface is the only

preserved expression of a transgression.

Linder conditions of rapid accommodation and sufficient

sedimentation, however, lagoonal and inlet facies are pre-

served (Kigure C65). These may occur in inverted succession

whereby the landward-most faeies (lagoonai marshes and

back-bay muds) are overlain by sandier deposits associated

with the approach of the transgressing barrier and capped by

tidal inlet facies just below the ravinement surface

(Figure C66). The thickness and organization of these

156 COASTAL SEDIMENTARY FACIES

Shelf sand/nud

Trfat inlal sand

nna backshore

[vivlv.vl Back bay mud

^^^1 Coastal

marsf\

peat

Nonmarine backshors

Figure C65 Transgressive succession produced in settings with rapid accommodation and adequate sedimentation, A lagoonal succession

develops as the lagoonal facies are buried by the landward-migrating barrier. It is capped by a "ravinement" surface overlain by shelf tacies.

Shell sand/moa

Transgressivs (—

bg '"

Tidal inlet

—;2—

Lower lagaan j ' -

tidal channels / '^

Ul»o'

tegoon i—C

tidal channels

Bactt iagoon mud

Salt marsh lignite

Backshore

J /^

=i_

RAVINEMENT

SURFACE

BAY

-FLOODING

SURFACE

Shett sand/mud /—

Transgressive

lag

I * * "*'• T '

Upper shoretace

RAVINEMENT

SURFACE

TIDAL

EflOSIOMAL

SURFACE

Figure C66 Vertical successions produced by preservation of transgressing lagoons. A. High rate of accommodation. B. Modest rate of

accommodation.

TRANSGRESSION

RAOATION

TRANSGRESSION

PHOGRADATION

TRANSGRESSION

PHOGHADATIO^

TRANSGRESSION

Figure C67 Typical pattern of stacked shoreline successions,

Shallowing-upward progradational parasequences are separated by

transgressive surfaces of erosion (ravinement surfaces). Falls in sea

leve!

prior to transgressions can produce erosional sequence

boundaries and incised-valley fili deposits at the top of the

parasequences.

transgressivc lagoonal deposits differs depending on the

balance between accommodation and sedimentation.

Although most are biii a few meters thick, some can exceed

20 m. Most display a basal "bay-Booding surface" and a

capping ravinement surface, although vigorous tidal currents

and low rates of accommodation may in some places lead to

the erosion of the lower part of the succession (Figure C66).

Most successions of coastal deposits consist of stacked sets

of progradational deposits separated by erosional (ravinement)

surfaces (Figure C67). The degree of preservation of individ-

ual progradational intervals depends on the amount of

accommodation provided during the interval between progra-

dation and transgression (Figure C68).

H. Edward Clifton

COLLOIDAL PROPERTIES OF SEDIMENTS 157

TRANSGRESSION

, ^ LARGE ACCOMMODATION

NONMARINE 8ACKSH0RE

SHOREFACE/BARRIER

TRANSGRESSION

NONMARINE BACKSHORE

SHOREFACE/BARRIER

[23

Figure C68 Preservation of progradational sequences as a function of

accommodation in an area of alternating regression and transgression.

Where the amount of accommodation is large (subsidence is rapid

relative to the rate of sea-level change), complete parasequences,

including nonmarine facies at their tops are preserved. Where the

amount of accommodation is small, erosion associated with the

marine transgression is likely to remove a significant part of the

parasequence deposited as part of the previous regression. Such

"beheaded"

successions could prove difficult to interpret. The location

relative to the point of maximum regression is also a factor. Near the

turnaround to transgression, much of regressive shoreline deposit is

likely to be lost.

Bibliography

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associations of the shoreface-connected ridge systems in the

German Bight (southern North Sea). Deutehes Hydrographiche

Zeitsehrift, 46: 229-244.

Augustinius, P.G.E.F., 1989. Cheniers and chenier plains: a general

introduction. Marine Geology, 90: 219-229.

Brunn, P., 1962. Sea-level rise as cause of shore erosion. Journalof

Waterways, Harbor Division, American Society of Civil Engineers,

88:

117-130.

Clifton, H.E., Hunter, R.E., and Phillips, R.L., 197t. Depositional

structures and processes in the nonbarred high-energy nearshore.

Journal of Sedimentary

Petrology,

41: 651-670.

Cobb,

J.C, and Cecil, C.B., 1993, Modern and Ancient Coal-forming

Environments. Geological Society of America, Special Publication,

286,

198p.

Demieco, R.V., and Hardy, L.A., 1994. Sedimentary Structures and

Early Diagenetic Features of Shallow Marine Carbonate Deposits.

SEPM (Society for Sedimentary Geology), Atlas Series No. 1,

265 p.

Galloway, W.E., and Hobday, D.K., 1996.

Terrigenous

Clastic Deposi-

tional Systems: Applications

Fossil Fuel and Groundwater Resources,

2nd edn. New York: Springer.

Hill, G.W., and Hunter, R.E., 1976. Interaction of biological and

geological processes in the beach and wearshore environment,

northern Padre Island. In Davis, R.A., and Ethington, R.L. (eds.).

Beach and

Nearshore

Processes.

SEPM Special Publication, 24, pp.

169-187.

Hill, P.R., Barnes, P.W., Hequette, A., and Ruz, M.-H., 1995. Arctic

coastal plain shorelines. In Carter, R.W.G., and Woodroffe, CD.

(eds.),

Coastal Evolution: Late Quaternary Shoreline Mor-

phodynamics. Cambridge University Press, chapter 9,

pp.

341-372.

Hine, A.C, 1975. Bedform distribution and migration patterns on

tidal deltas in the Chatham Harbor estuary. Cape Cod, Massa-

chusetts. In Cronin, L.E. (ed.), Estuarine

Research,

v II, Geologyand

Engineering. London: Academic Press, pp, 235-252.

Howard, J.D., and Reineck, H,-E,, 1981, Depositional facies of high-

energy beach to offshore sequence: comparison with the low-energy

sequence. American Association of Petroleum

Geologists

Bulletin, 65:

807-830.

Johnson, D.W., 1919. Shore Processes and Shoreline Development. John

Wiley and Sons.

Komar, P.D., 1976. Beach Processes and Sedimentation. Englewood

Cliffs,

New Jersey: Prentice-Hall.

McBride, R.A., and Moslow, T.F., 1991. Origin, evolution, and

distribution of shoreface sand ridge, Atlantic inner

shelf,

USA.

Marine Geology, 21: 1063-1066.

Niedoroda, A.W,, and Swift, D.J.P., 1981. Maintenance of the

shoreface by wave orbital currents and mean flow. Geophysical Re-

search Letters, 8: 337-348.

Niedoroda, A.W., Swift, D.J.P., Hopkins, 1985. The shoreface. In

Davis,

R.A. Jr. (ed.). Coastal Sedimentary Environments, 2nd edn.

New York: Springer-Verlag, pp, 533-624.

Pemberton, S.G,, Van Waggoner, J.C, and Wach, G.D., 1992.

Ichnofacies of a wave-dominated shoreline. In Pemberton S.G.

(ed,).

Applications of Ichnology to Petroleum Exploration. SEPM

Core Workshop No, 17, pp. 339-382.

Pemberton, S.G,, and Wiglitman, D,M,, 1992, Ichnologic character-

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Workshop No. 17, pp. 141-167.

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Processes,

Facies.

and

Stratigraphy. 3rd edition, Blackwell Scientific.

Reineck, H.-E., and Singh, I.B., 1973. Depositional Sedimentary Envir-

onments - with Reference to Terrigenous Clastics. Berlin: Springer-

Verlag.

Rine,

J.M., and Ginsburg, R,N,, 1985, Depositional facies of a mud

shoreface in Surinam, South America—a mud analogue to sandy

shallow-marine deposits. Journal of Sedimentary Petrology, 55:

633-652,

Rine,

J,M., Tillman, R.W,, Stubblefeld, W.L., and Swift, D.J.P., 1991.

Lithostratigraphy of Holocene sand ridges from the nearshore

and middle continental shelf of New Jersey, USA. In Moslow, T.F.,

and Rhodes, E.G. (eds.). Modern and Ancient Shelf Clastics: A

CoreWorkshop. SEPM Core Workshop No. 9, pp. 1-72.

Swift, D.J.D.P., Niedoroda, A.W., Vincent, CE., and Hopkins, T.S.,

1985.

Barrier island evolution. Middle Atlantic

Shelf,

USA, Part 1:

shoreface dynamics. Marine Geology, 63:

331-361.

Walker, R.G., and James, N.P., 1992. Facies Models Response to Sea

Level Change. Geological Association of Canada.

COLLOIDAL PROPERTIES OF SEDIMENTS

Sediment is a mass of organic or inorganic solid fragmented

material that comes from the weathering of rock and is carried

by, suspended in, or dropped by air, water, or ice, and that

forms layers on the earth's surface as gravel, sand, silts, or

mud. In sediments, the clay-sized and the fine silt-sized

particles are classified as colloids, which are made up of

particles having dimensions of 10-10,000A. The term "col-

loids"

indicates microparticles or macromolecules that are

158

COLLOIDAL PROPERTIES OF SEDIMENTS

small enough to move primarily by Brownian motion, rather

than by gravitational settling, and are large enough to provide

a microscopic "environment" into or onto which molecules of

interest can escape the aqueous solution. (The theoretical

gravitational settling velocity of colloid particles is less than

0.01 cms"'.) Colloids found in soil and sediment are aggre-

gates of highly disperse or loosely cohering organic, mineral

and organo-mineral particles, which are capable of acting as

sorbents and as ion exchangers.

Colloids are ubiquitous, they are found everywhere in

concentrations of above lO' to 10^ particles per liter of water.

Colloidal dispersions bridge the gap between the true dissolved

solutes and true bulk phases. The different varieties of

colloidal materials found in natural system are: (i) soil colloids

(aluminosilicates; humus and colloidal humic acids; iron

hydrous oxides; polymeric coating of soil particles by humus,

by hydrous Fe(III) oxides and hydroxy-Al(III) compounds);

(ii) river-borne particles (weathering products and soil colloids,

along with colloidal Fe(III) oxides stabilized by humic and

fulvie acids; phytoplankton, biological debris, and colloidal

humic acid); and (iii) sulfide and polysulfide colloids in anoxic

sediments.

The fine size of the colloidal particles is responsible for their

very high surface area/volume ratio. Chemical and physical

reactions are promoted in colloidal suspension due to the large

area of solid-liquid interface characteristic of systems contain-

ing many small particles. The high surface area of the colloids

also means that there are many broken bonds, and

j^oth

positively and negatively charged ions and molecules will be

present on the solid surface. According to Gouy-Chapman

model, the excess ions are non-uniformly distributed in the

vicinity of the colloid surface, their concentration is the largest

at the surface, and decreasing non-linearly till they reach bulk

concentration (Stumm, 1992). The net charge on the solid

colloidal particles will be the sum of all these surface charges

together with the contributions from the adsorbed ions. The

charge will depend on the chemical and erystallographic

structure of tlie solid and on the pH of the solution

surrounding the solid (Yariv and Cross, 1979). For each type

of colloidal solid, there is a pH value at which the surface

concentrations of

H"*"

and OH~ are equal and there is no net

charge on the particle, called the point of zero charge (PZC); in

the absence of ions other than H* and 0H~, the PZC will be

at a pH referred to as the isoelectric point (IEP). At pH values

above the IEP, the colloidal particles will carry a net negative

charge, and the particles will have a net positive charge at pH

values below the IEP.

Colloidal sediments can be grouped into two distinct classes

as lypophilic (i.e., liquid loving) and lyophobic (liquid-hating).

Lyophilic colloids display strong affinity between the dispersed

particles and the liquid. The liquid is strongly adsorbed to the

surface of the particle, so the interfacial tension between the

particle and the liquid is very similar to that between the liquid

molecules, leading to an intrinsically stable system. In case of

lyophobic colloids, the liquid does not show much affinity for

the particles. If attractive forces exist between the particles,

there will be a strong tendency for them to stick togetlier when

they come in contact, rendering the system generally unstable,

and resulting in floccuiation (q.v.). A system containing

colloidal particles is said to be stable if during the period of

observation it is slow in changing its state of dispersion.

Colloidal stability can be affected by the presence of

electrolytes, adsorbates that affect the surface charge of the

colloids, polymers that affect particle interaction by forming

bridges between them or sterically stabilizing them. Colloidal

floccuiation, induced by a rise in electrolyte concentration, is

an important sedimentation process, and has significant

implications regarding dispersion of potential pollutants that

may be bound to the charged particle surface.

Interactions of different interfacial forces, such as electro-

static repulsion, van der Waals attraction, hydration phenom-

enon, hydrodynamic interactions on particle collision, and

forces due to adsorbed substances dictate colloidal stability,

and these factors can be categorized into two groups as (i) the

frequency of collisions between particles and (ii) the prob-

ability of particles adhering if they collide. Particles in

suspension collide with each other as a consequence of at

least three mechanisms of particle transport: (1) particles move

because of their thermal energy (Brownian motion); (2) if

colloids are sufficiently large or the fluid shear rate is high, the

relative motion from velocity gradients exceeds that caused by

Brownian effects; (3) in settling, particles of different gravita-

tional settling velocities may collide and agglomerate (Ryan

and Elimelech, 1996).

Colloids are stable when they are electrically charged. The

stability is usually explained by Derjaguin-Landau-Verwey-

Overbeek (DLVO) theory. The attractive forces and repulsive

forces between two particles are calculated as a function of

particle-particle separation distance. The curves of potential

energy versus particle-particle separation distance typically

show an energy barrier that must be overcome by colliding

particles to adhere. The primary attractive forces are disper-

sion forces, and the primary repulsive forces are long-range

electrical forces and short-range Born repulsion. The double-

layer potential energy is affected by the changes in solution

chemistry, though the attractive van der Waals forces are

independent of it. Short-range repulsive forces (up to a few

nanometers separation distance) may also be affected by

changes in solution chemistry (Ryan and Elimelech, 1996).

Figure C69 shows the plot of normalized potential energy as

a function of separation distance between a colloid and

collector. A typical total potential energy curve is character-

ized by a deep well at a very small separation distance, which is

the primary minimum (0mini), a repulsive energy barrier,

which is the primary maximum (</)max), and a shallow well at a

larger separation distance

i4>mm2)-

The attractive energy is primarily the result of London-van

der Waals dispersion forces and is related to the polarizability

of the particles and the distance of separation between the

particles. For colloidal particles, the attractive energy declines

with the separation according to

ivdW

(Eq. 1)

where A// is the Hamaker constant, which depends on the

difference in polarizability between dispersed particles and the

dispersing medium, d is the separation, and n is an exponent,

the value of which is l<n<3, depending on the geometry of the

system (flat-flat, flat-sphere, sphere-sphere, etc.). If a

colloidal particle and the dispersing solvent are of similar

polarizability, then the tendency for the particles to seek out

one another and to separate from the solvent is small and there

is little tendency for the particles to adhere. The electrical

repulsion energy is due to the repulsion of like-charged

surfaces. There are three components of surface charge

COLORS OF SEDIMENTARY ROCKS

159

300

200

100-

-200

I ic|)Born

-300

,-10

-9

-8

Separation Distance (m)

Figure C69 The resultant curve for the total potential energy, 0'°'°'

(solid curve), is the sum of the van der Waals potential energy, (/>"''"',

the double layer potential energy, </>'", which is due to the repulsion

between the particles of similar charge, and the Born potential energy,

(j)^°'",

the short range repulsion energy, kg is the Boltzmann's constant,

and T is the absolute temperature.

resulting from specific chemical interactions or specific

adsorption: (1) constant charge surfaces (e.g., clays); (2) con-

stant potential surfaces (e.g., Agl); and (3) constant chemical

affinity surfaces (e.g., oxides).

Another component of surface charge is counter charge that

is held electrostatically to balance the specific-adsorption

charge. The counter charge is usually treated through an

electric double layer based on the Gouy-Chapman-Stern

theory. According to this theory, the repulsive energy drops off

exponentially with distance from the surface according to

exp i-Kd

(Eq. 2)

where K is the reciprocal thickness of double layer, which is

proportional to the square root of the ionic strength, /, for a

z:z electrolyte, K OC /"^.

Another repulsive interaction that has not been widely

documented is steric repulsion. This repulsion occurs when

particles that are coated with large polymers collide and the

polymeric coatings overlap. Factors affecting the extent of the

steric repulsion are the similarity of the coating to the solvent

and to the particle, and the expansion or contraction of the

polymer according to the properties of the solvent.

Finally, it can be said that colloids represent the most

valuable soil fraction, as colloids can retain mineral nutrients

(including trace elements) in the mobile state, and retain

moisture and support formation of soil structure favorable to

aeration. It is the colloidal component that dominates the

leachable chemistry component of the sediments.

Sandip and Devamita Chattopadhyay

Bibliography

Ryan, J.N., and Elimelech, M., 1996. Colloid mobilization and

transport in groundwater.

Colloids

and Surfaces A: Physicochemicat

and Engineering Aspects, 107: 1-56.

Stumm, W., 1992. Chemistry of the Solid-Water Interface: Processes at

the Mineral-Water and

Particle—Water

Interface in Natural Systems.

New York: John Wiley & Sons, Inc.

Yariv, S., and Cross, H., 1979. Geochemistry of Colloid Systems for

Earth Scientists. Berlin: Springer-Verlag.

Cross-references

Clay Mineralogy

Floccuiation

COLORS OF SEDIMENTARY ROCKS

The colors of sedimentary rocks can have complex origins and

in cases are secondary. However, colors are commonly

primary and reflect important aspects of depositional environ-

ments including redox conditions and rates of deposition of

organic matter. These parameters in turn affect fauna and thus

a strong correlation exists between color and biofacies

patterns, including macrofaunal distributions and burrowing

type and depth (Leszcynski, 1993), particularly in deposits of

oxygen-stratified basins (Myrow and Landing, 1992). Color

may be useful for the interpretation of variations in such

factors as relative sea level, oceanic circulation, sedimentation

rate and primary productivity.

Colors are generally controlled by accessory minerals and

compounds of iron and organic carbon (see reviews by

Pettijohn, 1975; Potter et al, 1980; and Myrow, 1990).

Compounds of other transition metals (e.g., Ti, Mn, Co, Cu,

and Zn) also impart a pigment in some cases. For the most

part the colors of sediment and sedimentary rock fall within

two spectra: green-gray to red and olive-gray to black

(Figure C70). The first color spectrum is controlled by the

oxidation state of iron, specifically the Fe"''^/Fe"'"'^ ratio in

various minerals, and not the total iron content (Tomiinson,

1916;

McBride, 1974). In particular, the color is a function of

the mole fraction representing the proportion of iron in the

-t-2 state (Fe"'"^/Fe"'"'^ + Fe"""') per gram of rock. In relatively

unoxidized strata, green colors result from the presence of

iron-bearing phyllosilicates such as chlorite, illite, and in cases,

glauconite. Kaolinite and smectite provide for white to light

neutral colors. Red coloration is due to the presence of

hematite, whereas less common yellows and browns generally

result from limonite and goethite, respectively. Red hematitic

staining is generally an early diagenetic phenomenon resulting

from: (1) dehydration reactions in which limonite stains on

detrital particles are altered to hematite; (2) dissolution of iron

silicates and precipitation of the released iron; and (3) direct

oxidation of magnetite and ilmenite grains (Hubert and Reed,

1978).

Very low levels of hematite can impart a deep red

coloration to a rock. The conversion from red to green colors

occurs by reduction of the iron, which is either carried away in

solution or reprecipitated as iron-rich clay minerals such as

chlorite (Thompson, 1970). The strength of red or green color

is a function of grain size, with fine-grained rocks having

higher iron content and thus more intense color. This

160

COLORS OF SEDIMENTARY ROCKS

0.4-

Time

Figure C70 Graph relating color of rock, organic contenl and oxidation state of iron to time. "Time" refers to the length of lime ihat pore fluids

interact with sediment prior to lithification. Dark colors, olive-gray to black, are controlled by orgtinic content. Green-gray to red colors,

controlled by the oxidation state of

iron,

are reversible. The color history of sediment is dependent upon the Eh condilions. Three representative

path5 (high, moderate and low Eh) are shown for sedimeni with initially high organic content. Figure from Myrow (1990).

relationship is due to the presence of amorphous or poorly

crystalline iron oxides attached to clays (McPherson, 1980).

Changes in the oxidation state of iron, and hence the color,

result from interactions with altering fluids at any time in the

postdcpositioniil history of a unit, and are thus controlled in

part by permeability. Oxidation or reduction spots are due to

incomplete diagenetic alteration of a layer.

The grccn-gray to black color spectrum in rocks is a

function of total organic carbon (TOC). with darker colors

corresponding to higher carbon content. This relationship is

empirically confimied in studies of modern sediment (e.g..

Sheu and Presley, 1986). Main controls on organic carbon

content include accumulation rate of organic matter, sediment

accumulation rate, rate of decay of organics, and oxygen

levels (Potter etal.. 1980). Diagenetic loss of carbon, which

leads to lighter colors, is an irreversible process. The carbon

reservoir acts to shield the iron-bearing minerals from

oxidation until most of the carbon is oxidized, at which point

the green to red color spectrum is found. Oxidative loss of

carbon, and red coloration, occurs when bottom water and/or

pore-fluids are highly oxygenated. Green colors result from

deposition of sediment with low organic content and weakly

reducing to oxidizing conditions. Gray lo black colors are

associated with high TOC and dysaerobic to anaerobic bottom

waters.

Gray colors also occur as a result of the presence of

disseminated pyrite. During burial diagenesis rnetastable

minerals such as mackinawite (FeS) and greigite (FeS4) are

transformed into submicron sized framboids of pyrite (Berner.

1971).

Thermal maturity may also be an important factor in

some rock. For instance, Lyons (1988) demonstrated that the

black color of some limestone is due to carbonization of very

small quantities of organic matter (<0.06 percent TOC).

Darker colors may also be imparted by Mn oxides or Fe-Mn

oxides such as pyrolusite. Shale with siderite is gray to bluish

on fresh surfaces but altered to brown tones by weathering

(Pettijohn, 1975).

Paul Myrow

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COMPACTION {CONSOLIDATION) OF SEDIMENTS

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organic carbon and iron suUides in anoxic sediment from the Orca

Basin. Gulf of Mexico. Marine Geology. 70: 103-118.

Thompson. A.M.. 1970. Geochemistry of color genesis in red-bed

sequence. Juniata and Bald Eagle formation:!. Pennsylvania.

J<nir-

nalof Sedimentary Petrology. 40: 599-615.

Tomiinson. C.W.. 1916. The origin of red beds. Journalof

Geology.

24:

153-179.

Cross-references

Black Shalcb

Chlorite in Sediments

Red Beds

Siiilide Minerals m Sediments

COMPACTION (CONSOLIDATION) OF

SEDIMENTS

Compaction may be defined as the process by which the

sediment volume is reduced and the sediment density

increased. Sediments change their physical properties after

deposition due to the stress from the overburden (gravitational

compaction) and as a result of biological or chemical reactions

involving dissolution and precipitation of minerals. In some

sediments compaction of amorphous materials like organic

material (kerogen) luid biogenic silica (Opal A) also play an

important role. All natural processes modifying sediments

after deposition are in geological literature referred to as

diagenesis ((/.v.) and this term also includes cotnpaction

processes. Soil engineers use the term compaction to describe

densifications on the laboratory and in the field by vibrations

of mechanical equipment to reduce the pore volume and

consolidation for the expulsion and flow of water in soils.

Geologist tends to use the term consolidation to describe a

strengthening of the rocks, often by precipitation of cements

(lithification).

Compaction of sediments and rocks (soil and rock

mechanics) has been developed us a quantitative engineering

subject since Terzaghi (1925) from the need to predict the

response of sediments due to ehanges in stress caused by

different types of constructions. Laboratory tests determining

the mechanical properties of different types of sediments (soils)

have formed the basis for such quantitative prediction (Lambe

and Whitman, 1979). Sueh tests measure the friction coeffi-

cients between grains and the mechanical strength of grains.

Later Terzaghi's principles were applied to petroleum en-

gineering and structural geology (Hubbert and Rubey, 1959).

The terminology used in the engineering literature is rather

different from that used by most geologists and that may cause

confusion.

Sedimentologists and petroleum geologists have taken an

increasing interest in sediment compaction. This is stimulated

by the need to predict the porosity of reservoir rocks as a

function of burial history in sedimentary basins (Wilson.

1994).

The densities of the different lithologies are required to

compute subsidence and backstripping in basin modeling

(Giles, 1997; Sclaterand Christie. 1980). There is also a need to

understand the distribution of rock strength, seismic velocities

and attributes in sedimentary basins.

Until recently, sediment compaction has been mostly treated

as mechanical compaction as a function of effective stress.

Compaction processes involving chemical dissolution and

precipitation of grains (chemieal compaction) are principally

different since chemical proeesses then control the rate of

compaction and the strength of the rocks.

It is much more difficult to study chemical compaction in

the laboratory, particularly in silicate rocks (i.e.. sandstones)

due to the slow kinetic reaction rates involved in the

compaction processes at low temperatures.

Clastic sediments are granular materials and the grains may

be roek fragments, minerals or amorphous tnatcrials. The

fluids between the grains may be water, oil or gasses or a

mixture of these lluids.

Porosity ((p) is the volume fraction of the fluids in the

sediment. The volume fraction of the solids then becomes

!-<p.

The ratio between the volume of the pore space (void)

filled with fluids and the solid grains and matrix (mainly clay

minerals) is also referred lo as the void ratio (iv);

-</>)

(Eq. I)

Stresses in sedimentary basins

The effective stress of the overburden is an average of the stress

transmitted through the solid phases through a network of

load-bearing grains. The stress at grain contacts may be very

high because of small contact areas and focused on some

grains, which may then bend or fracture. With increasing

compaction the rearrangements or breakage of grains will

increase both the contact area and the number of grain

contacts. Precipitation of cement in the grain surfaces increases

the area of contact further so that the stress per contact is

reduced enough to prevent further grain crushing or mechan-

ical compaction all together.

The total overburden stress (tr,.) at a certain depth in

sedimentary basins is equal to the weight of the overlying rock

and fluids per unit area and is often referred lo as the

lithostatic stress.

At a certain depth (Z,) below the seafloor in a sedimentary

basin the tolal vertical stress is:

= Z,

•

PI, + 7,, (Eq. 2)

1b2

COMPACTION (CONSOLIDATION) OF SEDIMENTS

Here the average bulk density of the sedimentary rocks over a

depth range Z,, is

/)/,.

The water depth above the seafloor is Z,,

and the density of water (/)„.).

The bulk density (ph) of the rocks varies as a function of the

porosity

((p),

the density of Ihe fluid (/),) in the pore space, and

the density of the solid phase

(p,,,).

which are mainly minerals:

Ph =

p,<(>

+ /',„(' ~ (/>) (Eq. 3)

Usually the density of the mineral matrix in sandstones and

shales is close to 2.65 2.70 g/cm' because this is the density of

the most common minerals. If there are significant contents of

denser minerals such as siderite or pyrite ihe bulk density will

be higher. Smectite and mixed layer minerals have variable but

generally lower densities. The fluid density also varies with the

composition of water and petroleum. The weight per unit area

of the rock column including fluids is called the lithostatic

pressure (stress).

The total overburden stress (load)

((T,)

can be obtained by

integrating the sediment density ip) over the entire overburden

sequence (Z) or by estimating an average density (pJ. This

stress is partly carried by the grain framework (solid phase)

which is called the eft^ective stress (n^,) and by the stress

(pressure) earried by the fluid phase which is pore pressure

(/^^,)

(Eq. 4)

(T|.

- OyC

The effective stress is thus

-f = a,. -

(Z).

This may be a useful approximate expression for

compaction trends in sedimentary basins and it is useful in

mathematical modeling. However there is nothing in the

mechanical and chemical compaction processes that should

cause the porosity to be a negative exponential function of

depth or the effective stress. Sedimentary basins consist of

sedimentary layers with different compressibility, and the

compaction curves showing observed porosity loss with depth

may be very complex.

In the laboratory we can measure the induced deformation

(strain) in a sample of sediment or rock for a given effective

stress and construct stress-strain curves. Experimental data

show degree of time dependent compaction (creep) at constant

stress.

In sedimentary basins the rates of increase in effective

stress due lo loading are very slow compared to the

experimental loading rates in the laboratory and it is some-

what uncertain how significant mechanical creep is in a

sedimentary basin.

If one assumes that the grain structure may be considered

linearly elastic and isotropic for very small deformations, one

may define the following deforniational characteristics:

Young's modulus (£) is the ratio between increase in normal

etTective stress (An) and the resulting strain (A;;-) in the stress

direction (r), when there is no change in the orthogonal normal

stresses.

E =

^(T-J^e.-_

(Eq. 8)

(Eq. 5)

The bulk modulus [K) is the ratio between the increase in equal

iill-round stress (Aa) and the resulting volumetric compression.

This was first formulated by Terzaghi in 1925 and it formed

the basis for modern soil and rock mechanics.

If the average sediment density (pJ in the upper 3 km in a

sedimentary basin is 2.0g/cm , the total stress at 3 km depth is

60 MPa. At hydrostatic conditions (no overpressure) the pore

pressure {P^} is equal to the average density of the pore fluid

multiplied by tlie height (Z) of

water column in the sediments

and any water column above (water depth). Saline water may

have a density of 1.033 g/cm"' and the pore pressure at

depth is then 31 MPa and the effective stress (tr,.) is then

29 MPa at

km depth. For very fine-grained sediments and

rocks with low porosity and poorly connected pores, the fluid

phase will carry a smaller IVaction of the overburden load

following poro-elastic theory. A constant a (Biot constant) was

therefore introduced in Ihe Terzaghi's law:

fT,.

= fT - otPn (Eq. 6)

For porous and compressible sediments

may be taken to be

I, but for stiff sedimentary rocks a may have significantly

lower values meaning that more of the total load is carried by

the solid phase.

Geomechanical properties

The reduction in porosity with burial depth may be described

as an exponential function following Athy (1930):

^ae-'''

(Eq. 7)

Here the porosity (<^) is an exponential function of an initial

porosity

{/)(,,

a compressibility parameter c and the burial depth

K = An/AEy,,i

(Eq.

The bulk compressibility (c) is the inverse of the bulk modulus

K and is a parameter describing the rate of compaction as a

function of stress. In sedimentary basins with only minor

tectonic stress, the vertical stress

(Oy)

is normally the principle

stress as in uniaxial compression tests. The compressibility of a

sedimentary layer may the be expressed as the shortening (AL)

over a bed thickness L and the strain (c) is then (ALIL).

The ratio between the lateral expansion (lateral strain

J:/)

and

vertical shortening (axial strain r,J is expressed as the Poissons

ratio (v):

v^ ((;,/(;„) (Eq. 10)

If the lateral expansion in both directions is equal to half the

shortening, there is no chanee in volume. This corresponds to

= 0.5.

For uniaxial compaction (one dimensional compaction)

which may be assumed for wide sedimentary basins, the lateral

strain is zero and a lateral effective stress will be built up. For

linearly elastic and isotropic grain structure, the lateral

effective stress increase is:

An,,, = AfT,,, • y/{\ - v)

The compaction processes detemiine rock properties such as

rock strength and the mechanical responses to changes in

stress during drilling and production of petroleum.

COMPACTION (CONSOLIDATION)

SEDIMENTS

163

Compaciioii processes

sedimentary rocks also determine

well

log

responses

and the

velocity

seismic signals

and

their

attributes.

Types of compaction

Modeling

sediment compaction requires that

under-

stand

the

processes involved.

clear distinction tnust

made between purely mechanical compaction

and

chemieal

compaction:

(1) Mechanical compaction

(a) Rearrangement (sliding

and

reorientation)

grains.

(b) Bending

and

ductile deformation

grains.

grains (grain crushing).

(2) Chemical compaction

(a) Dissolution

contacts between grains

along

stylolites (pressure solution)

and a

correspond-

ing precipitation

minerals (cement)

in the

pore

space.

(b) Dissolution

thermodynamically unstable grains that

are load-bearing.

Mechanical compaction

The vertical stress

normally

the

maximum stress (major

principle stress)

sedimentary basins thai

arc not

subjected

strong tectonic compression, extension

uplift. During

progressive burial

hydrostatic pore pressures (drained

conditions),

the

sedimentary rocks always experience increas-

ing effective stress,

and

such sediments

are

referred

to as

normally consolidated. Sediments that previously have experi-

enced higher effective stresses than

present

are

referred

to as

over-consolidated.

The term under-consolidation

sometimes used

char-

acterize poorly compacted sediments, particularly

mud

with

high pore pressure. This term should however

a\oidcd since

the degree

consolidation always refers

lo the

maximum

etTective stress sediments have experieneed,

and

they

can

therefore

not be

under-consolidated.

The

term under-compac-

tion

may

also cause confusion.

The degree

compaction (porosity loss) varies greatly with

the lilhology

at the

same effective stress (Figure C71). During

uplifi

and

erosion

or a

build

up of

over-pressure (higher than

hydrostatic pore pressures)

the

effective stress

may be

reduced

so that

the

sediments become over-consolidated

and

brittle.

Sediments

are

likely

respond

in a

ductile manner also

Well 34/7-1

Sequence Gamma

log

API

20 iO 60 80 100 1.0

velocity

log

km/s

2.0

3.0 0

Clay minerals

2 ^m)

XRD %

50 100

Pliocena-

Pleistocene •css-al

Miocene

Oligocene

•

- -

•

depths ate In meters below KB

Figure C71 The velncily t'rom well logs

lor

seismic: riatai

may be

a good indicator

of the

degree

compaction. The velocity

in a

well

penetrating lertiary sediments trom

the

Northern North Sea show that Pliocene and Pleistocene muds show a normal compaction trend (higher

velocities) with depth. Smectite rich Eocene and Oligocene muds have very much lower velocities and

less compdcled despite grealer

burial depth. This may be partly due

overpressure

but is

mostly due

the poor compressibility

smectitic clays. Modified from Thyberg et al.

(2000).

164 COMPACTION (CONSOLIDATION) OF SEDIMENTS

(A)

Mean Grain size

•

V51 mm

0.6S mm

0.71 mm

0.33 mm

C.19 mm

0.10 mm

0 5 10 15 20 25 30 35 40 45 50

Vertical effective stress (MPa)

Figure C72 Experimental unitixial compactiun of well sorted

quariz rith and showing that coarse sands have a much higher

porosity loss than fine grained sand. From Chuhan e( al. (2002).

Mono-quartz sand

40 MPa (0.63 mm

response to leclonit: stress if the strain rates are very low so

that the rocks can deform by chemical compaction (Bjorlykke

and Hoeg. 1997).

Experimental compaction of sand

Compaction of sand depends on the compostion and primary

sorting of grains. Sands with lithic fragments show more

compaction at low to moderate stresses than well-sorted sand

(Pittman and Larese, 1991).

Experimental uniaxial compression at high stress (50MPa)

on well-sorted mono-quartz rich sands shows that stress-strain

curves are non-linear over the entire stress range. Initial

compaction is due to grain sliding rotation and rc-oricntation.

A stable grain framework is then established and further

compaction requires grain breakage. Reeent experimental

work (Chuhan etal.. 2002) show that eoarse. well-sorted sands

have a higher porosity loss than fine-grained sand subjected to

the same stress (Figure C72). Coarse sands also show evidence

of more frequent grain fracturing ihan fine sands in sandstone

reservoirs, probably because the stress per contact increases

with increasing grain size (Figure C73). Fracturing reduces the

grain size and gradually distributes the load onto more grain

contacts. At higher stresses porosity/stress curves for different

sands approach similar slopes indicating similar compressi-

bility. The porosity at higher stresses is therefore mostly a

result of the differences in compaction up to 20-25 MPa or

2.0-2.5 km burial depth.

The grain crushing can thus be considered a part of an

adjusting (self-organizing) process where the sediment is

adapting itself to higher stress. Grain crushing is particularly

intense in sandstones where chlorite or micro-quartz coatings

retard the precipitation of quartz overgrowth (Fisher etal.,

1999).

This is also true in cold basins (<20 C/km) where the

onset of quartz cementation may start at 5-6 km. Fracturing

of grains may destroy grain coatings, thus opening fresh

surfaces for quartz cementation (Figure C73(B)).

Compaction of mudstones

Mudstones compact mechanically as function of the effective

stress and the compressibility of the mud (Aplin

etaf.

1999).

Figure C73 Fracturing ol sand grains: (Al Experimental tracluring of

sand al 40MFa stress. This is equivalent to the effective stress below

3..S-4l<m of overburden; (B) Natural fracturing of sand grains in

lurassic reservoir rocks at 4,800m depth, Lavrans Field, Haltenbanken,

offshore Norway. The fractures were produced by mechanical

compaction prior to quartz cementation, and subsequently healed by

authifjenic quartz, which is darker on the cathoduluminescence

microphotgraph. The present burial depth (4,800

is probably the

maximum, and the reservoir pressure is hydrostatic. Chlorite coating

the grains delayed quartz cementation until fractures produced a

suitable substratum.

Laboratory tests show that pure clay minerals may retain a

relatively high porosity even at high stresses. Smectite which is

the most tine grained clay retains porosities of about 45

percent at stresses equivalent to about

km of overburden

(Rieke and Chillingarian, 1974) while illite and kaolinite which

form larger crystals compact more readily.

Smectitic clays consist of very small partieles. down to loA

thick, and have very large specific surface areas (Nadeau and

Bain. 19S6). This implies that the stress is distributed on a very

large number of grain contacts so that the stress per grain and

unit area is very low. This may explain the low compressibility

of smectitic shales down to 2-3 km burial depth (Bjurlykke,