Middleton G.V. (Ed.) Encyclopedia of Sediments and Sedimentary Rocks
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("YCLIC SEDIMENTATION
18.5
Osleger, D.. and Read, J,F.. 1991. Relation of eustasy to stacking
patterns of meter-scale carbonate cycles. Lale Cambrian.
U.S,A,
Jiutrnallit Seditnentary Petrology. 61: 1225 1252,
Osmond.
J.k,. Carpenter. J.R., and Windom, ILL.. 1965. 2.WTh/
234U age ofthe Pleistocene corals and oolites of Florida. JottrnaloJ
Geophysical
Re.search.
70: 1843-1847,
Perkins,
R.D.. 1977. Depositional framework of Pleistocene rocks in
south Florida. Geological Stniety of
.America
Memoir 147, pp. 131-
198,
Pitman.
W.C. 1978. Relationship between eustasy and stratigraphic
sequences of passive margins. Geological Society of America Bulle-
tin.
89: 1389-1403.
Plafker. G.. 1965. Tectonic deformation associated with the 1964
Alaskan earthquake. .S'ii(7nc, 148, 1675 1687.
Ramsbottom.
W.H.C, 1979, Rates of transgression and regression in
the Carboniferous of NW Europe. Geological Society of London
Journal. 136: 147-153.
Read.
J.F., 1995. Oveniew of carbonate platform sequences, cycle
stratigraphy and reservoirs in greenhouse and ice-house worlds. In
Read,
J.F., Kerans. C. and Weber. L.J. (eds.). Milatikovtieh Sea
Level Changes. Cyeles and Reservoirs on Carbonate Platforms in
Greenhouse
and
Ice-Hou.se
IVorltls. SHPM (Society for Sedimentary
Geology) Short Course Notes no. 35. pp. I 102.
Read,
J.F., Grotzinger. J.P., Bova, J.A,. and Koerschner, W.F.,
I98(i.
Models for generation of carbonate cycles. Geolot^y. 14:
107-110.
Read,
J.F., and Goldhammer. R.K,, 1988. Use of Fischer plots to
define third-order sea-level curves in Ordovician peritidal cyclic
carbonates, Appalachians. Geology. 16: 895 899.
Reading,
H.G. (ed.), 1996, Sedimentary Environments: Processes.
Faeies
and Stratigraphy. 3rd cdn. New York: Blaekweli Scientific.
Ross.
C.A., and Ross. J.R.P., 1987. Late Paleozoic sea levels and
depositional sequences. Cushman Foundation for Foraminiferal
Research.
Special Publication 24. pp. 137 149.
Sadler. P,M,, Osleger, D.A., and Montanez, LP,. 1993. On the
labeling,
length, and objective basis of Fischer plots. Journal of
Scdimciilary Petrology. 63: 360-368.
Sattcrley, A.K., 1996. Cyclic carbonate sedimentation in the Upper
Triassic Dachstein Limestone. Austria—ihc role of patterns of
sediment supply and tectonics in a plat form-reef-bas in system.
Journal of Sedimentary
Research.
66: 307-323.
Schwarzacher, W,. and Fischer, A.G., 1982. Limestone-shale bedding
and perturbations ofthe earth's orbit. In Einsele, G., and Seilacher,
A. (eds.). Cyclic and
Event
Stratification. New York: Spinger Verlag.
pp.
72 95,
Schwarzacher. W., 1975. Sedimentation Models and Quantitative Strati-
graphy. Developments in Sedimentology 19. New York: Elsevier,
Sloss,
L.L.. 1963. Sequences in the cratonic interior of North Ameriea.
Geological Society of
.Atnerica
Bulletin. 74: 93-1 13,
Sonnenfeld,
M.D., 1996. Seqtience evolution and hierarchy within
the Lower Mississippian Madison Limestone of Wyoming. In
Longman,
M.VV..
and Sonnenteld. M.D. (eds.), FaleozoieSystemsof
the Roeky Mountain Region. Rocky Mountain Section SEPM
(Society for Sedimentary Geology), pp. 165-192.
Sonnenfeld,
M.D,, and Cross. T.A., 1993. Volumetric partitioning and
facies differentiation within the Permian Upper San .Andres
Formation of Last Chance Canyon. Guadalupe Mountains,
New Mexico. In Loticks. R.G,, and Sarg. J.F, |eds.). Carbonate
Sequenee Stratigraphy: Recent Developments and .Application.s.
American Association of Petroleum Geologists Memoir, 57, pp.
435-474.
Sprenger, A., and Ten Kate. W,G,. 1993, Orbital ftircing of calcilutite-
marl cycles in southeast Spain and an estimate for the duration of
the Berriasian stage, Geologieal Soeiety oj America Bulletin. 105:
807 818.
Stout, W.H., 1931. Pennsylvanian cycles in Ohio. Illinoi.s State
Geological
Survey
Bulletin. 60: 195 216.
Strasser, A., 1988. Shallowing-iipward sequences in Purbeckian
peritidal carbonates, lowermost Cretaceous. Swiss and French
.liLra Mountains, Sedimentology. 35: 369-383.
Suess,
E., 1888. Das Antlitzder Frder\olumc 2 (English Translation by
Sollas.
1906). Oxford: Clarendon Press, pp.
1-254.
Tinker, S.W., 199S. Shelf-to-basin facies distribtilion and sequence
stratigraphy of a steep-rimmed carbonate margin: Capitan deposi-
tional system. McKittrick Canyon, New Mexico and Texas.
Jour-
nalof Seditneniary
Re.seareh.
68, 1146 1175.
Turcotte.
D.L, und Willemann, J.H.. 1983. Synthetie cyclic stratigra-
phy. Earth and Planetary
Science
Letters. 63: 89-96.
Vaii.
P.R., 1987. Seismic stratigraphy interpretation procedure. In
Bally, A.W., ed.. Atlas of Seismic Stratigraphy I'olume
1.
American
Association of Petroleum Geologists Studies in Geology 27, Tulsa:
Oklahoma, pp. 1-11,
Vail.
P.R., Mitchum, R.M.. Jr.. and Thompson, S., III. 1977. Seismic
stratigraphy and global changes of
sea
level. In Payton. C.E., (ed.).
Seismic Stratigraphy—.Applications to Hydrocarbon E.xploration.
American Association of Petroleum Geologists Memoir 26, pp.
83-97.
Vail.
P.R., Hardenbol, J.. and Todd, R.G., 1984. Jurassic unconfor-
mities,
chronostratigraphy, and sea-leve! changes from seismic
stratigraphy and biostratigraphy. In Schlee, J.S. (ed.). Interregional
i nconformities and Hydrocarbon .Accumulation. .American .Associa-
tion of Petroleum Geologists Memoir 36, pp. 129 144.
Van der Heide. S,. 1950, Compaction as a possible factor in upper
Carbonilerous rhythmic sedimentation. Report ofthe fStli Interna-
tional
Geologie
Congress,
part 4. pp. 38 45.
Van Wagoner. J.C, Mitehum, R.M. Jr., Posamentier, H.W,, and
Vail,
P.R.. 1987. The key delinitions of stratigraphy. In Bally. A. W.
(ed.).
.Atlas ofSei.smie
Stratigraphy, lolume I. American Association
of Petroleum Geologists Studies in Geology 27, Tulsa: Oklahoma,
pp.
II 14.
Van Wagoner, J.C, Mitchum, R.M., Campion, K.M.. and
Rahmanian.
V.D,. 1990, Siliciclaslic sequence stratigraphy in
well-logs,
cores, and outcrops. Ameriean
.A.ssociation
of
Petroleum
Geologists
Methods in F.xploration
Series.
No. 7. 55p.
Vella,
P.. 1965. Sedimentary cycles, correlation, and stratigraphic
classification.
Tran.sactimi.s
ofthe Royal Society oj Nt'\y Zealand,
3: 1-9.
Walker, R.G., and James, N.P, (eds.). 1992, Fades Models:
Re.spon.seto
Sea
Level
Change.
Geological Association of Canada.
Wanless. H.R., and Shepard, F.P., 1936. Sea level and climate changes
related to late Paleozoic cyclothems. Geological Society of America
Bulletin. 47: 1177-1206.
Wanless, H.R., and Cannon. J.R.. 1966. Late Paleozoie glaeiation.
Earth
.Science
Reviews.
1: 247-286.
Weedon.
G,P,, 1989. The detection and illustration of regular
sedimentary cycles using Walsh power spectra and filtering, with
examples t>om the Lias of Switzerland, Journal of
rlie
Geological
Societyof London. 146: 133-144.
Weller, J.M., 1930, Cyclical sedimentation ofthe Pennsylvanian period
and its significance. Journal oj
Geology.
38: 97-135.
Weller, J.M,. 1956. Argument for diastrophic control of late Paleozoic
cyclothems. .American
.-issoeiation
of
Petroleum Geoloi;ist.s
Bulletin.
40:
i7 50,
Wells.
A.J.. 1960. Cyclic sedimentation: a review. Geological
Magazine. 97: 389-403.
Wheeler. H.E,, 1964. Baselevel, lithosphere surfaee, and time-
stratigraphy.
Geological
Society oj America Bulletin. 75: 599-610.
Wheeler, ILE.. and Murray. H.H., 1957. Base level control patterns
in cyclothemic sedimentation. .Ameriean
.A.ssoeiation
of Petroleum
Geologists
Bulletm. 41: 1985 2011.
Wilkinson,
B.H., Diedrich. N.W.. and Drummond. CN., 1996. Facies
successions in peritidal carbonate sequences. Journal of
Sedimen-
lary
Research.
66:
U)(^5
1078.
Wilkinson.
B.H,, Drummond. C.N., Rothman. E.D., and Diedrich,
N.W.. 1997, Straial order in peritidal carbonate sequences, .lournal
of Sedimcntaty
Re.seareh.
67: 1068-1082.
Wilkinson,
B.H., 1982. Cyclic cratonic carbonates and Phanerozoic
caleite seas, .lournalof
Geologieal
Edueation. 30: 189-203.
Wilson.
J.L.. 1975, Carbonate
Eaeies
in Geologie History. New York:
Springer-Verlag.
D
DEBRIS FLOW
A debris flow is a type of sediment gravity flow. It is a rapid mass
movement of a concentrated mixture of sediment, organic
matter, and water that can flow like a liquid yet can stop on
sloping surfaces and form a nearly rigid deposit. Observed debris
flows commonly have been likened to flowing masses of wet
concrete, Genericaliy, debris flow includes mudflow (debris
containing mostly sand, silt, and clay), lahar (volcanic debris
flows), and till flow (debris flow from active or stagnant ice).
Debris flows are a signiflcant geomorphological and
sedimentological agent in a wide range of geologic and
climatic settings, Subaerial debris flows pose a natural hazard
than can cause fatalities, damage structures, and diminish land
productivity. Subaqueous debris flows can produce turbidity
currents, and they can transport sand from continental slopes
to abyssal plains, threatening underwater infrastructures and
creating important hydrocarbon reservoirs. Modern debris-
flow deposits that line numerous mountain channels and
blanket many subaerial and subaqueous fans, and ancient
debris-flow deposits that form distinctive strata in many
sedimentary sequences attest to their widespread sedimentolo-
gical importance. Although the discussion below describes
recent developments regarding subaqueous debris flows, it
foeuses primarily on subaerial debris flows.
Debris flows have distinctive physical characteristics. They
have bulk densities approximately twice that of clear water, and
they can easily transport clasts larger than
1
m in diameter,
Del)ris flows have volumes that range from a few tens of cubic
centimeters to more than I billion cubic meters; they have
maximum velocities that range from a few to a few tens of meters
per second; they can affect lowlands more than 100 km from
their initial source; and they can flow tens to hundreds of km
across nearly flat subaqueous plains, A single subaerial debris
flow can transport about as much sediment as that transported
annually by small mountainous rivers, and it can introduce
abundant organic matter (trees) to river channels, modifying
channel dynamics and affecting aquatic habitat.
Debris-flow origins
A debris flow forms when a mass of sediment is mobilized and
transformed into a flowing slurry. Most commonly, a debris
flow results from a landslide triggered by torrential rainfall,
rapid snowmelt, or an earthquake, and in subaqueous
environments by additional combinations of rapid sedimenta-
tion, gas release, and storm-wave loading, A debris flow can
form during a volcanic eruption when hot volcanic debris
melts and mixes with snow and ice, or when crater lakes or
near-surface groundwater are explosively expelled, A debris
flow may also form when pore water is expelled during
consolidation of a landslide deposit, A rapidly released flood
surge, resulting from a glacier outburst or from breaching of a
natural or engineered dam, may entrain a sufflcient amount of
sediment to become a debris flow.
Behavior of flowing debris
A debris flow moves in an unsteady and nonuniform manner. It
commonly flows in a series of surges having periods that range
from a few seconds to several minutes or longer. Debris surges
commonly develop bulbous heads that contain a concentration
of the coarsest particles and precede gradually waning, flner-
grained bodies, but surges can also have agitated watery flow
fronts that precede more sediment-rich waning flow. Despite its
commonly coarse grain size, a debris flow can move across very
gentle slopes. When a debris flow stops, it commonly leaves a
lobate or digitate deposit having a steep terminal snout and
coarse-grained, steep-sided lateral levees, A more watery debris
flow forms a deposit that drapes valley slopes and terraces and
leaves a thin veneer near its inundation limit.
Scientific advances in understanding debris flows
Conceptual mechanics
Concepts and analyzes of debris-flow motion and deposition
have evolved over several decades. From the 1920s through
1970s, debris-flow studies progressed from descriptive ac-
counts of flows and deposits toward efforts to model flow
DEBRIS FLOW
187
behavior and interpret deposit character in terms of material
rheology (e,g,, a Newtonian fluid. Sharp and Nobles, 1953; a
Bingham viscoplastic material, Johnson, 1970, 1984) or
particle collisions (e,g., Takahashi, 1978, 1991). Although
many studies proposed various constitutive equations that
relate shear stress to shear strain rate, support for these models
remains largely anecdotal or based primarily upon small-scale
laboratory experiments. Small-scale experiments, however, do
not necessarily scale appropriately with natural processes.
Many laboratory-scale experiments lack gravel and use
amounts of silt and clay that are disproportionate for most
debris flows. Physico-chemical effects related to cohesion,
important at the small scale, are negligible at the field scale,
and inertial effects important at the field scale are poorly
simulated at the small laboratory scale.
Debris fiows undergo significant changes of state from
initiation through deposition (Iverson and Vallance, 2001), As
a result, others view debris fiows as granular materials having
flaw behavior modified by interstitial fiuid (Iverson, 1997),
Iverson's research has challenged efforts aimed at describing
and modeling a debris fiow in terms of fixed rheology. Most
debris fiows are fundamentally granular materials mixed with
water; therefore, Iverson (1997) hypothesized that a debris
fiow can be modeled most parsimoniously as a simple grain-
fiuid mixture having properties that depend on position and
time.
In his model, the Coulomb criterion for frictional
material relates shear stress to normal stress, and intergranular
pore-fiuid pressure that varies with position and time mediates
the frictional resistance of the granular debris. Rather than
treating a debris fiow as a material having a fixed rheology,
Iverson's model focuses on conserving mass and momentum in
a deformable body and allows debris-fiow "rheology" to
evolve as a fiow forms, moves, and deposits sediment, Iverson
and colleagues (e,g,, Iverson, 1997; Major and Iverson, 1999;
Iverson and Denlinger, 2001) tested Iverson's hypothesis in a
series of large-scale flume experiments using up to
10
m' of
poorly sorted sandy muddy gravel. The experiments demon-
strate that fiuid pressure, solids concentration, and mixture
agitation evolve as a debris fiow proceeds from initiation to
deposition and that the evolution of these variables strongly
influences fiow dynamics, depositional process, and the
characteristics of the deposit.
Depositional process and deposit characteristics
At any given location, debris-flow deposition can occur
abruptly or it can be prolonged over a significant time interval.
Recent field and experimental studies show that common
characteristics of debris-fiow deposits, such as poor sorting,
massive to stratified texture, and graded to ungraded bedding
can result from progressive, incremental accretion rather than
exclusively by abrupt en masse deposition. Debris fiows that
are longitudinally segregated—those having coarse fiow fronts
followed by finer-grained fiow bodies, and those that coarsen
from head to tail—can produce massively textured deposits
exhibiting normally graded or inversely graded clastic frag-
ments (Vallance, 2000), Deposits having pervasive particle
alignment, strong vertical changes in particle composition or
sorting, and clear evidence of emplacement by recurrent debris
surges also result from incremental accretion (Vallance and
Scott, 1997; Major, 1997, 1998), Unfortunately, single
outcrops of a debris-fiow deposit commonly do not display
clear evidence of depositional process. Owing to possible
accumulation by prolonged accretion, deposit thickness
cannot be used to infer rheological properties of a debris fiow,
contrary to a common field procedure that assumes deposition
by a Bingham viscoplastic material (Johnson, 1984).
Advances in theoretical and experimental studies indicate
that flow-margin dynamics play a pivotal role in debris-fiow
deposition, contrary to previous suggestions that deposition is
controlled by intrinsic rheological properties of fiow, such as
uniform yield strength, or pervasive dissipation of pore-fluid
pressure, Iverson (1997) demonstrated that leading edges of
experimental debris fiows of lOm^ exhibited little or no
positive pore-fluid pressure, whereas trailing flow was nearly
liquefied, and he successfully modeled the experimental debris
fiows by incorporating high-resistance fiow fronts dominated
by solid forces. Major and Iverson (1999) further demon-
strated that high pore-fiuid pressure developed within debris-
fiow interiors during flow initiation and acceleration persisted
during fiow deceleration and deposition. Deposition of the
experimental debris flows was therefore controlled by grain-
contact friction concentrated at flow margins that acted as
dynamic debris dams rather that by an intrinsic uniform
material strength or by pervasive dissipation of high pore-fluid
pressure. The dynamic dams that developed in these fiows
resulted from segregation and concentration of the coarsest
clastic material at surge fronts.
Flow-front dynamics affect deposition by submarine debris
flows as well as subaerial debris flows. For example, Mohrig
et al. (1998) observed that the fronts of subaqueous experi-
mental debris flows, composed of nearly equal mixtures of
sand and silt, accelerated rapidly and sometimes detached
from flow bodies if the flow fronts hydroplaned over the
substrate. Hydroplaning occurs when a debris flow cannot
displace leading ambient fiuid rapidly enough and as a result
high basal fiuid pressure develops at the fiow front. Thus, when
high fiuid pressure is developed at the fiow front, fiow
resistance markedly declines, a finding compatible with results
from subaerial debris-fiow experiments.
Sustained high pore-fiuid pressure within bodies of debris
flows can infiuence deposit characteristics as well as deposi-
tional process. This can be demonstrated by monitoring post-
depositional pore-fluid pressure in debris-fiow deposits. Major
and Iverson (1999) demonstrated that significant dissipation of
high pore-fluid pressure in a debris flow occurs only after
deposition, except in narrow regions around flow margins, and
Major (2000) demonstrated that high pore-fluid pressure in a
debris-flow deposit dissipates over minutes to months depend-
ing on deposit thickness and material diffusivity. By compar-
ing typical periods of debris surges with characteristic times of
dissipation of deposit fluid pressure. Major (2000) demon-
strated that freshly deposited debris could be remobilized
easily. Deposit remobilization may be preserved as soft-
sediment deformation (e,g,, Mohrig et al., 1999), However,
deposit remobilization may simply mute or obliterate strati-
graphic distinction among recurrent debris surges and lead to
a homogeneous, incrementally accreted deposit that appears to
be the product of a single debris surge.
Regardless of depositional process or depositional environ-
ment, all debris-fiow deposits share common sedimentological
characteristics. In general, debris-flow deposits typically are
unstratified, poorly sorted mixtures of sediment ranging in size
from microns to meters. Deposits may show little particle-size
segregation or exhibit normal or inverse grading of their
coarsest fragments. Deposit thicknesses typically range from
188
DEDOLOMITIZATION
several centimeters
to
several meters. Elongate clasts
in
deposits
may
show preferential alignment
or
imbricate
structure, particularly along deposit margins. Deposit termini,
where preserved, typically
are
lobate
and
have blunt margins
that commonly
are
studded with
the
coarsest debris. Although
deposit interiors
may
appear
to
represent deposition from
a
single debris surge, deposit surfaces
can
exhibit arcuate ridges,
clusters
and
bands
of
gravel
or
sand,
and
particles that align
along debris-surge perimeters,
all
sedimentologic features that
indicate that deposition resulted from emplacement
by
multi-
ple debris surges.
Debris-flow transformation
Although
the
composition
and
size distribution
of
sediment
in
a debris-flow deposit
is
commonly similar
to
that ofthe source
material, debris-flow character
and
volume
can
evolve during
transit
and
result
in a
deposit having characteristics that
are
different than those
of
the initiating flow. After mobilization,
a
debris flow
can
entrain channel-bed
and
bank sediment
and
increase
in
volume several fold. Such "bulking", which usually
occurs
in
steep, proximal channel reaches (e.g., Vallance
and
Scott, 1997),
can
modify
the
initial particle-size distribution
and introduce "exotic" clasts into
the
mixture.
In
more distal
reaches
a
debris flow
can
interact with streamflow
and
undergo
volume loss
and a
gradual transformation into so-called
hy-
perconcentratedstreamflow (Pierson
and
Scott, 1985; Vallance,
2000)
and
eventually into sediment-laden floodflow.
A
hyperconcentrated-streamflow deposit typically
is
composed
of massive
to
weakly stratified, moderately sorted sand,
and
has
a
sedimentologic texture that
is
intermediate between those
of debris-flow
and
alluvial deposits (Scott, 1988).
Summary
Researchers have made signiflcant strides developing greater
insights regarding debris-flow initiation, transport,
and
deposi-
tion since
the
1980s. Studies have moved from descriptions
of
flow behavior
and
sedimentologic properties
of
deposits
toward analyzes
of
event frequency
and
magnitude, develop-
ment
of
more realistic mechanical models,
and
quantiflcation
of
the
subtleties that influence mobilization, transport,
and
deposition
of
debris.
New
insights show that
a
debris flow
can
be modeled
as a
simple grain-fluid mixture having flow
behavior that
is
influenced
by
distributions
of
fluid pressure
and solids concentration, which evolve with position
and
time.
Characteristics
of
deposits formerly thought
to
form only
by
sudden
en
masse deposition have been shown
to
develop
through prolonged, incremental deposition
as
well.
Jon
J.
Major
Johnson,
A.M., 1984.
Debris flow.
In
Brunsden,
D., and
Prior,
D.B.
(eds.).
Slope Insiability.
New
York: Wiley
and
Sons,
pp.
257-361.
Major,
J.J., 1997.
Depositional processes
in
large-scale debris-flow
experiments. Journal of Geology, 105: 345-366.
Major,
J.J.,
1998. Pebble orientation
on
large experimental debris-flow
deposits. Sedimeniary Geology,
\ll:
151-164.
Major,
J.J.,
2000. Gravity-driven consolidation
of
granular slurries:
implications
for
debris-flow deposition
and
deposit characteristics.
Journalof Sedimentary Researeh,
70:
64-83.
Major,
J.J., and
Iverson,
R.M., 1999.
Debris-flow deposition: effects
of pore-fluid pressure
and
friction concentrated
at
flow margins.
Geological
Soeiety of Ameriea Bulletin, 111: 1424-1434.
Mohrig,
D.,
Whipple,
K.X.,
Hondzo,
M.,
Ellis,
C, and
Parker,
G.,
1998.
Hydroplaning
of
subaqueous debris flows. Geological Society
of America Bulletin, 110: 387-394.
Mohrig,
D.,
Elverhei,
A., and
Parker,
G., 1999.
Experiments
on the
relative mobility
of
muddy subaqueous
and
subaerial debris flows,
and their capacity
to
remobilize antecedent deposits. Marine
Geol-
ogy, 154: 117-129.
Pierson, T.C.,
and
Scott, K.M., 1985. Downstream dilution
of a
lahar:
transition from debris flow
to
hyperconcentrated streamflow.
Water
Resources Researeh, 21: 1511-1524.
Scott,
K.M., 1988.
Origins.
Behavior,
and
Sedimentology of Lahars and
Lahar-runout Flows
in
the Toutle-Cowlitz River system.
U.S.
Geolo-
gical Survey Professional Paper, 1447-A.
Sharp,
R.P., and
Nobles,
L.H., 1953.
Mudflow
of 1941 at
Wrightwood, southern Califomia. Geologieal Society
of
America
Bulletin,
64:
547-560.
Takahashi,
T., 1978.
Mechanical aspects
of
debris flow. Ameriean
Soeiety
of
Civil Engineers. Journal
of
the Hydraulics Division,
104:
1153-1169.
Takahashi,
T.,
1991. Debris
Flow.
Rotterdam;
A.A.
Balkema.
Vallance,
J.W.,
2000. Lahars.
In
Sigurdsson,
H.,
Houghton,
B.,
McNutt,
S.R.,
Rymer,
H., and
Stix,
J.
(eds.), Eneyehpedia
of
Volcanoes.
San
Diego: Academic Press,
pp.
601-616.
Vallance,
J.W., and
Scott,
K.M., 1997. The
Osceola mudflow from
Mount Rainier: sedimentology
and
hazards implications
of a
huge
clay-rich debris flow. Geohgicat Society
of
America Bulletin,
109:
143-163.
Cross-references
Alluvial Fans
Avalanche
and
Rock Fall
Earth Flows
Fabric, Porosity,
and
Permeability
Flume
Grading, Graded Bedding
Grain Flow
Grain Size
and
Shape
Gravity-Driven Mass Flows
Hindered Settling
Imbrication
and
Flow-Oriented Clasts
Liquefaction
and
Fluidization
Mass Movement
Slope Sediments
Slurry
Submarine Fans
and
Channels
Tills
and
Tillites
Bibliography
Iverson,
R.M.,
1997.
The
physics
of
debris flows. Reviews of Geophy-
sics,
35:
245-296.
Iverson,
R.M., and
Denlinger,
R.P.,
2001. Flow
of
variably fluidized
granular masses across three-dimensional terrain:
1.
Coulomb
mixture theory. Journal
of
Geophysical Research, 1O6(B1):
537-
552.
Iverson,
R.M., and
Vallance,
J.W.,
2001.
New
views
of
granular mass
flows. Geology,
29:
115-118.
Johnson,
A.M., 1970.
Physical Processes
in
Geology.
San
Francisco:
Freeman, Cooper,
and Co.
DEDOLOMITIZATION
Introduction
The term "dedolomitization" refers
to the
partial
to
wholesale
transformation
of
former dolomite rocks (dolostones—see
Dolomites
and
Dolomitization)
to
limestones,
or
partial
to
complete transformation
of
dolomite
to
calcite
on the
scale
of
DEDOLOMITIZATION
189
individual crystals. The transformation process includes both
neomorphic repUicement (recrystallization see Neomorphism
and Rccrystallization) of the dolomite by calcite, and dolomite
dissolution followed by calcite cementation.
Dedolomitization has been recognized as an Important
process in the diagenetic histories of a wide range of
Precambrian to Phaneio/oic marine, and to a lesser extent.
lacustrine carbonates. As with dolomitization. dedolomitiza-
lion histories can be multistagcd and protracted. Dedolomi-
tization has also been shown by geoehemical modeling to be an
ongoing and important reaction in modern-day karstic
aquifers (Plummer eial.. 1990; Raines and Dewers, 1997).
Dedolomitization process
Several aulliors have described the dedolomitization process as
the reverse but analogous situation to replacement dolomiti-
zation (e.g.. Fairbridge. 1978; Dockal. 1988). As with
dolomitization. dedolomitization involves consideration of
thermodynamics, kinetics, and fluid flux. But unlike replace-
ment dolomitization. dedolomitization can also refer to a
pore-filling process. Porosity changes associated with dedolo-
mitization are inconsistent and difficult to predict.
Dedolomitization can be represented simply by the follow-
ing reaction:
CaMgfCO,},,
The above reaction should proceed to the right if
aCa'7aMg-'>l (Abbott. 1974). Ca"' is usually supplied by
dissolution ol" a mineral source, for example gypsum or
anhydrite. Mg""^ liberated into solution may accumulate,
causing the solution to reach saturation with respect to both
calcite and dolomite, and thus halting the reaction. Alterna-
tively, dedolomitization may continue until one of the
reactants is no longer present if Mg'* is incorporated into
another diagenetic phase, or if the tluid tlux. is sulTiciently high
that dolomile saturation is not attained. The above reaction
reverses, theoretically, when aCa""^/aMg'"' <
1
and dolomite is
precipitated. The dedolomitization reaction is generally
considered to be slow under natural geological conditions,
but reactive transport modeling by hydrogcologists has
suggested that, in some instances, the process is geologically
quite rapid, occurring in hundreds of years.
von Morlot (1S47) considered the dissolution of evaporites
(gypsum or anhdrite) to be essential in the dedolomitization
process, as shown in the following reaction:
Numerous others have also suggested that evaporite
dissolution was essential in driving dedolomitization. where
the continual addition of Cn~^ from gypsum or anhydrite
results in the precipitation of calcite due to the common ion
effect (e.g.. Back cial.. 1983; Plummer eial., 1990). Decreases
in pH. calcium and carbonate from solution cause further
dissolution of dolomite. Back ei al. (1983) and others, have
also suggested that dedolomitization will not occur in low-
temperature environments unless gypsum, or some other
source of additional Ca"' ions, is present to maintain a high
ratio.
Oxidation of pyrite has also been proposed
as a source of aqueous (SO4)" in some situations (e.g.,
Zenger. 1973). Others have indicated that the presence
of (S04)^" is not a prerequisite for dedolomitization to occur
(e.g., Stoessel elal., 1987).
Numerous field, petrographic and geoehemical studies have
supported Ihe conclusion that dedolomitization is a surficial
phenomenon, related to post-burial surface weathering and
other near-surface processes associated with subaerial expo-
sure,
including the development of unconformities and karstie
aquifers. Dedolomite often occurs within the most readily
weathered parts of the rock and grades into "fresh" dolomite
away from weathered surfaces. Dedolomites from near-
surface diagenetic settings may show negative carbon and
oxygen isotopic values retlectlng soil-gas infiuences in meteori-
cally-derived water, including fluids derived from buried
paleoaquifers.
Dedolomitization can occur during deeper burial diagenesis
and may be associated with stylolites. hydrocarbons and other
late diagenetic mineral phases, including fracture or vug-filling
caicites (e.g., Budai ciat.. 1984). The influence of hydrocarbons
may be recognized in light carbon isotopic values of the
dedolomite. Light oxygen isotopic values should refiect
precipitation at elevated temperatures characteristic of the
burial environment. Other authors have noted dedolomiti/a-
tion associated with mixing zones, eontact metamorphism and
hot Ca-rich brines derived from albitization of plagioclase in
nearby sandstones.
Recognition of dedolomite
Limestone formed by dedolomitization may have very little in
common with the parent dolostone. or its precursor limestone.
In the field, the occurrence o\' limestone in otherwise dolomitic
strata in the vicinity of faults, or alotig weathered surfaces and
below unconformities, may indicate dedolomitization.
A range of microscopic labrics can be used to identify
dedolomitization (Figure Dl). although not all are reliable
indicators, Fmpty rhomb-shaped pores indicate dissolution of
dolomite and perhaps dedolomitization if it can be shown that
calcite once occupied these pores. Calcite partially or entirely
filling rhomb-shaped pore space, or replacing dolomite
crystals, constitutes the most straightforward evidence of
dedolomitization. Differentiation of calcite precipitated as
cement versus replacement is made on the same type of
textural evidence used to interpret aragonite to calcite
transformations. Isolated calcite rhombs may represent very
selective dedolomitization because calcite does not commonly
develop a rhombohedral habit in limestones. However,
exceptions have been reported.
Mosaics of calcite crystals showing rhombic patterns of
ferric oxides or other inclusions could represent the palimpsest
remnants or "ghost" outlines of former dolomite crystals or
internal zoning in these crystals, ln some dedolomite rocks, the
dedolomite matrix is characterized by a clotted "grumeleuse"
texture resembling peloidal grainstone. This fabric results from
the retention of impurities from the cloudy, inclusion-rich
cores of the original dolomite crystals in the newly formed
dedolomite crystals, which otherwise obscure the former
dolomite crystal boundaries (Evamy. 1967; Kenny, 1992).
Ferroan dolomites appear to be more susceptible to
dedolomitization due to their metastability in the surface
environment, resulting in the oxidation and hydration of the
ferrous iron previously bound in the dolomite lattice, and a
190 [HhORMATION
OF
SEDIMENTS
polycrystalhne calcite
pseudomorph
after dolomite
single calcite crystat
pseudomorph afler dolomite
± Fe-oxides
or
other inclusions
dedoiomile
nm
dedoiomite
zone
dedolomite
core
1
1
1
•
—•
DOLOMITE
POROSITY
CALCITE
(DEDOLOMITE]
CEMENTATION
REPLACEMENT
CEMENTATION
Ot
REPLACEMENT
Figure
Dl
Schem<itiL surnm-iry
of
common dedolomite microl'jbrics.
close association
of
hematite
and
goethite with
the
newly
formed dedolomite. These iron-bearing minerals
are
usually
responsible
for the
rusty weathering color
of
many dedolomi-
tized rocks,
and
they
may
also retain
a
seeondary ehemieal
remanent magnetization useful
for
dating diagenetic events.
Dedolomitization
ean
also
be
zone-selective, typically after
ferroan
or
calcian internal zones, rims
or
cores
of
dolomite
crystals, leaving more stoichiometric dolomite relatively
unaltered (Figure
Dl).
Terminology
Modern usage
of
the term dedolomite typically infers caleite
as
the mineral occupying
the
rock volume formerly taken
by
dolomite,
but
other minerals, such
as
quartz, eelestite,
and
iron
oxides,
are
also known
to
replace dolomite.
As
terms, dedolo-
mite
and
dcdolomiiizaiioti have heen criticized
as
being
ambiguous
and
inconsistent with
the
nomenclature
for
other
types
of
mineral replacements, where names
of
newly formed
minerals were based
on
mineralogy ofthe replacement, rather
than
the
replaced phase (Smit
and
Swett. 1969). Dcdolomitiza-
tkm
has
also been used
as an
all-eneompassing definition
to
refer
to any
alteration affecting dolomite (Fairbridge. 1978).
However,
the
terms dcdoloniiic
and
dcdnliiinitizatidn.
as
used
in this article, are well-entrenched
in the
modern literature
and
their current meanings
are
clear.
Mario Conielio
Bibliography
Abbott. P.L,. 1974, Calcitization
of
Edwards Group dolomites
in the
Btilcones fault
/one
aqtiifer. soiilli-centnil Texas. Geology.
1:
359-362.
Back,
W.,
Hanshaw,
B.B..
PItLmtncr.
L.N..
R;ihii,
P.M..
Righlmirt-.
C.T..
and
Rubin,
M..
1983. Process
and
rate
of
dedolomitization:
mass transfer
and C
dating
in a
regional earbonate aquifer. Cn'o-
liigiculSinietvol Anicricii Bulletin. 94: 1415-1429.
Budai. .I.M., Loiimann. K.C..
and
Ouen. R.M.. 19K4. Burial dedolo-
mite
ill the
Mississippian Maiiison Limestone, Wyoming imd Utah
thrust belt. Jouriuilof SediDtoiuiry Petrology.
54: 276 288.
Dockal.
J.A..
I98fi. Thertiiodynamic
and
kinetic descriplion
of
dolomiiizalion
of
calciie
and
calcitization
of
dotomite (dedoiomi-
tization). Ctirhoiuili'.saiidEvaparate.s.
3:
125-141.
Evamy,
B.D., 1967.
Dedolomilizalion
and the
development
of
rhombohedral pores
in
limestones. Joiiriui!
of
Scdiiuciiturv Peiml-
of;y.
37:
1204-1215.
Tairbridge,
R.W.. 1978.
Dedolomitization.
In
Fairbridge,
R.W..
and Bourgeois.
J.
(eds.).
The
EiiLyhpedia
of
Sedinwntolo^y.
Siroiidsburg. Pennsylvania: Dowden, Hutchison
and
Ross
Inc..
pp.
233-235.
Kenny,
R.,
1992. Origin
of
disconformity dedolomite
in the
Martin
Formation (Lale Devonian, northern Arizona). Sedinieniurv
Gealogv. 78:
137 146.
von Morloi. A.. 1847. Ueber Dolomit und seine kunslliche Darstellung
aus Kalkstein. SiitiintisseiisthaftlUhf
.•ihlicnulluiigcii
Gesammelt.
und dunk
Stih.seriptivn
I.iiisii,
van
Wilhehii
Haidinf^cr.
1:
305-315.
Plummer,
L-N.,
Busby.
J.F.. Lee. R.W.. and
Hanshaw.
B.B.. 19%.
Gcochemiea! modeling
of the
Madison Aquifer
in
parts
of
Monlana. Wyoming,
and
Soiilh Dakota. Hiiter Rcsoiirce.s
Re-
search.26:
1981 2014.
Raines. M.A..
and
Dewers, T.A,, 1997, Dedolomitization
as a
driving
tnechanism
for
karst generation
in
Permian Blalne Formation,
sotilhwesteinOklahoma. USA.
CcirhiiniitesuinlEvuporiles.Xl:
24
31.
.Smit. D.E..
and
Swell.
K..
1969. Devaluation
of
""dedolomiti/;ilion".
JinirnahifSedinienuirv Pvtnilogv.
39:
379-380.
Stoessell.
R.K..
Klimetitidis.
R.E.. and
Prezbindowski,
D,R., 1987,
Dedoiomilization
in Na
Ca-CI brines from
100' to 200 al 301)
bars,
(ifoehiniiaiel
Co.siiioehimieti
.4etci.
51: 847-855.
Zenger.
D.H..
1973. Syniaxial caleile borders
on
dolomite crystals.
Liltle Falls Formation lUpper Cambrian).
New
York. Journal<if
ry
Hetrolof^y.
43: I 18-124.
Cross-references
Ancienl Karsl
Anhydrite
and
Gypsum
Dolomiles
and
Dolomiti/ation
Magnetic Properties
of
Sediments
Neomorphism
and
Recrystalli/aiion
Porewaters
in
Sediments
DEFORMATION
OF
SFDIMFNTS
Introduction
Most sediments undergo some degree
of
deformation after
deposition. Burial under later sediment leads
to
expulsion
of
pore fluids
and to
adjustments
in
grain packing, involving both
physieal and chemical processes. The main result
is a
reduction
in thickness
of
sedimentary units
and., in
some cases,
the
development
of
horizontal fabrics
of
platy and elongate grains.
In
the
case
of
fine-grained sediments, containing clay minerals
and deposited with high initial porosities, such compaetiona!
reduction
in
thickness
can be
large
so
thai depositiona!
lamination
and
bedding
are
compressed beyond recognition.
For coarser sediments,
the
depositional grain framework
usually involves lower initial porosities and
a
greater resistance
to compaction.
In the
extreme, organic-rich sediments such
as
peat
may
undergo compaction
to
only
10
pereent
of
initial
thickness whilst,
for
well-sorted sands, thickness changes
are
negligible. Where sediments
of
contrasting grain size
are
interbedded
in a
lenticular fashion, differential compaction
can
OFFORMATION OF SEDIMENTS
191
disioiT depositional geometries (see CoiJipadion). Sediments
ih;U iuivc undergone eompaelion and olher diagenetic changes
are typically lithified. In this more rigid state, they may be
deformed by teetonic processes, most typically compressional
Iblding and thrusting or extensional faulting, often at great
depths where temperatures are high and rates of deformation
shnv. These processes fall uithin the realm of tectonics and arc
not dealt with here.
Prior to significant compaction, sediments may be de-
formed,
in appropriate circumstances, soon afler deposition
and this early, soft-sediment deformation is the subject of this
article.
Some early deformation is virtually contemporaneous
with deposition whilst some occurs because of shallow burial
and early compaction. In all cases, the sediment experiences a
loss of strength and behaves, for a short period, as a plastic or
a
lluid.
In some cases, loss of strength may result in total
remobilization of the sediment and transformation into a
gravitational mass movement, leading to a phase of re-
sedimentation. This section concentrates on situations where
the loss of strength is sufficiently short-lived for sediment to
regain its strength and preserve struetures that record more or
less in situ deformation. However, these distinetions are not
always clear-cut. Tectonic structures are often geometrically
identical with syndepositional struetures and boundaries
between tectonic stress regimes and more local, sediment-
induced regimes may be blurred.
Deformation within a body ol" sediment occurs when inter-
granular forces are unable to resist apphed stresses, which are
usually, but not exclusively, of gravitational origin. The shear
strength of sediment is normally expressed by the equation:
T
^ C + (o - /?)tan<p
uhere r is shear strength. C is grain cohesion, o is pressure
normal to shear, /» is excess pore fluid pressure, and tp is the
angle of internal friction. This means that sediment will fail
when the applied stress exceeds r and such a condition will be
favored by lower cohesive strength C. by changes in grain
packing to reduce lan
(p
or by pore fluid pressure /) increasing
beyond some critical value. Such conditions can occur tor
several reasons, which are discussed below. Failure, when it
does occur, may be isotropie. e.\tending throughout the body
of sediment or it may be concentrated on diserete surfaces.
Loss of sirength
Loss of strength and sediment deformation are most likely in
sands and finer-grained sediments. They are rare in conglom-
erates unless these are very poorly sorted with a high
proportion of tine-grained matrix. The presence of fine-
grained,
especially muddy sediment tends to increase the role
of cohesive strength, which is largely a function of grain size:
in belter sorted, mud-free sediment, tVictional forces dominate.
ln both cohesive and non-cohesive sediments, elevated pore-
water pressure is the main means by which strength is lost but
that eondition can have different causes and durations in
contrasting sediments.
In fine-grained, muddy sediments, excess pore fluid pres-
sures are created by relatively rapid deposition. Low perme-
abilities inhibit escape of pore water and hence sediment
compaction, giving over-pressured conditions. In addition, the
decay ol' organic matter within a body of sediment can
generate gases, which, if unable tu escape, also contribute to
pore pressure. Over-pressured eonditions can occur close lo the
sediment surfaee where highly mobile mud may be susceptible
to deformation or ilowage. They can also develop at a larger
scale and on a longer timescale within a thick pile of sediments
sueh as that produced by a prograding delta or deltaic margin
(e.g..
Mississippi. Niger). In such cases, increasing pore-water
pressure wilh burial may eventually e.\ceed lithostaiic pressure,
leading to failure. This may result in plastic flowage (Fig. D2)
of a thick, over-pressured layer or it may result in the
development of discrete slip surfaces, which cut up section as
listric faults or the bounding surfaces of sediment gravity
slides. Such conditions, and the resultant movements, develop
and persist over long periods of time, in the case of the largest
deltaic margins, over millions of years.
Rapidly deposited sands, lacking significant line-grained
sediment, eommonly have rather loose grain packing. They are
particularly sLisceptible to shock, such as might be generated
by an earthquake, by heavy wave action, by further rapid
deposition or by a sudden rise in water level. Shock breaks
grain contacts and induces tighter packing and. as a result.
excess pore water will be present, raising the pore-water
pressure and causing temporary liquefaction (Fig. D2). The
duration of liquefaction will be determined by the rate at
which pore water escapes. In the case of sands, high
permeability usually means that liquefaction is short-lived.
Whilst liquefied, the sediment-water mixture deforms readily
as a Newtonian
fiuid.
in response to gravitational and olher
applied stresses. As pore fluid escapes, usually upward to the
sediment surface, grain contacts are re-established as a rising
front of reconsolidation "freezes" the deformed sediment.
Deformation is usually visualized as the distortion of original
depositional lamination but. in the case of extreme or
sustained liquefaction and deformation, depositional lamina-
tion is destroyed and structureless sand or slurried textures are
produced.
Upward movement of
escaping
pore water may also
create deformation, both by distorting depositional laminae
and by creating new vertical structures such as dishes, pipes
and sheets within which sediment-water mixtures move as
fluidized flows
(Fig.
D2).
Deforming forces
Virtually all the forces lhat act upon sediment weakened by the
above processes are gravitational in origin. These ean be
divided into two major classes, a down-slope component of
gravity and the action of gravity on inverted density gradients.
In addition, deforming forces ean also be produeed by shear at
the sediment surface caused by flowing water, ice and sediment
gravity flows.
Where muddy sediment occurs on a slope, the down-slope
gravitational component may be sufficient to overcome its
cohesive strength, This can lead to both detachment of a slab
of sediment on a basal shear surface and to internal
deformation within the moving layer. The style of deformation
will vary depending upon the position within the moving sheet.
Stretching, particularly in the up-slope end. may lead to tensile
failure and to the development of brittle e.\tensional structures
whilst compression, commonly in a down-slope setting, eauses
plastic folds. These styles of deformation are associated with a
family of mass movements (see Mass Movcmenl: Slide
and Slump Structures). Also partly driven by down-slope
gravitational components are sedimentary growth faults.
These are related to failures in deeply buried over-pressured
muds.
Extension on the faults is also driven by lateral flowage
192
DEFORMATION
OF
SEDIMENTS
Deformation
style
Properties
Plastic
Liquefied
Fluidized
Yield strength
Relative pore fluid velocity
Flow structure
%
Water
Viscosity
Rate of water escape
Primary structures
Elutriation of fine grains
Intrusions
Significant
<Uo
Laminar
Preserved deformed
—
Negligible >^
Generally concordant
Negligible
—
>Uo
—
Turbulent
Not preserved
Minor
— Significant
—
Generally discotdani
Dish structures
Figure
D2
Main processes
ot
sediment deformation,
the
associated characteristics
and the
resultant effects
and
products.
Based
on
Lowe (1975).
of those muds
due to
vertieal loading
by the
overlying
sediment pile.
Density inversions occur commonly,
as a
result both
of
depositiona! processes
and of
early loss
of
pore water.
A
sand
layer, deposited
on mud.
will commonly have
a
higher density
than
the mud and. if the mud is
sufficiently weak,
the
instability
may
lead
to
loading
of
sand into
the mud (see
Load
Structures).
Where
the
sand
is of
uneven thickness,
as
with
ripple lenses
or
erosional sole marks, loading
may be
concentrated where
the
sand
is
thickest. Extreme cases
of
loading show detached balls
of
sand, usually eneased
in
slurried
mud or
silt.
Within liquefied sand layers, which commonly result from
shock, patterns
of
density inversion
may be
more subtle.
Internal depositional lamination
may be
contorted
in con-
volute lamination
U/.v.)
some
o(
which reliccts
the
upward
escape
of
water
or
low-density sediment-water nii.xtures. Very
rapidly deposited sands
may
lack depositional lamination
and
undergo liquefaction
as
part
of the
depositional process.
De-
watering
of
such sands
and the
associated etutriation
of
fines
can lead
to the
development
of
subtle structures sueh
as
dish
structures
and
water-escape pillars
and
pipes.
Deeper burial
and
associated differential compaction
due to
gravitational loading
may
lead
to
deformation
of
larger-scale
depositional geometries.
The
cross-sections
of
channel sand-
bodies,
in
particular,
are
susceptible
to
change IVom
planar-eonvex forms
to
concavo-convex patterns where
encased
in
muds. Where channel sands
are
encased
in mud.
as
in
some deep-water depositional settings, de-watering ofthe
muds during burial
can
lead
to
high pore-water pressures
in
the sands. This
can
eventually lead
to
rupture
of the mud
seal
and
to
injection
of
liquefied sands
as
sills
and
dykes
(see
Clas-
tic (Ni'plunianj Dykes and
Silts).
When injection occurs dose
to
the sediment surface, liquefied sand
may be
extruded
as
slurry
flows
or as
sand volcanoes.
Where sands
are
deposited
by
powerful currents
as
dunes
orsandwaves. they
may be
susceptible
to
liquefaction as
a
result
mainly
of
rapid deposition
on the lee
slope
of the bed
form.
If
this happens whilst
the
current
is
still active,
the
cross-bedded
sands,
developed
by
lee-side deposition,
may be
sheared
and the
foresets dragged into folds, giving overturned eross-bedding.
The movement
of ice
over unconsolidated sediment also
generates large-scale
and
intense deformation
of
those sedi-
ments. This occurs near
the
front
of a
glacier, where glacial
advance operates
as a
bulldozer
to
fold
and
thrust proglacial
deposits.
In
addition, basal layers
of a
mobile glacier
may be
rich
in
sediment
and
shearing across
the
basal layers
can
lead
to intense deformation though
the
rheology
has
more
to do
with
the
plastic deformation
of ice
than with processes
of
liquefaction
or
sediment cohesion.
John
D.
Collinson
DEFORMATION STRUCTURES AND CROWTH FAULTS
193
Bibliography
Boulton, G.S., and Hindmarsh, R.C.A., 1987. Sediment deformation
beneath glaciers: rheology and geological consequences. Journal
of
Geophysical
Research, 92: 9059-9082.
Maltman, A. (ed.), 1994. The Geological Deformation of Sediments.
London: Chapman & Hall.
Jones,
M.E., and Preston, R.M.F. (eds.), 1987. Deformation of Sedi-
ments and Sedimentary Rocks. London: Geological Society of
London, Special Publication, 29.
Lowe, D.R., 1975. Water escape structures in coarse grained
sediments. Sedimentology, li: 285-308.
Structures of mud. Some occur on otherwise flat bedding
surfaces whilst others accentuate erosional or depositional
irregularities. In extreme cases, sand may be detached from its
source bed and occur as isolated load balls or pseudonodules
(^.v.)
within highly disturbed muds or silts. Internal lamination
within both sand and mud is highly contorted. Loading is
caused by gravity acting on an unstable density stratification
where the sand layer is denser than the muds and where both
layers are weakened by excess pore fluid.
Cross-references
Bedding and Internal Structures
Clastic (Neptunian) Dykes and Sills
Compaction (Consolidation) of Sediments
Convolute Lamination
Debris Flow
Deformation Structures and Growth Faults
Dish Structure
Flame Structure
Fluid Escape Structures
Gases in Sediments
Liquefaction and Fluidization
Load Structures
Pillar Structure
Ball-and-Pillow (Pillow) Structure
Pseudonodules
Slide and Slump Structures
Syneresis
DEFORMATION STRUCTURES AND
GROWTH FAULTS
This section deals with so-called soft-sediment deformation. It
documents how different types of deforming stress, acting on
sediment that has been temporarily weakened, lead to a range
of deformational structures. The physical processes are
described in more detail in the section on Deformation of
Sediments. The structures themselves do not fall naturally
into well-defined classes but Figure D3 shows a scheme, which
helps to relate structures to processes. This diagram illustrates
interactions between different causes of loss of strength and
deforming forces and shows that processes are gradational
with tectonic deformation and with re-sedimentation. It shows
the wide range of scales across which deformation can occur,
from micro-loading to growth faults at the scale of continental
slopes and highlights differences in timing and duration of
deformation. Structures that form during deposition, such as
overturned cross-bedding, contrast with structures dependant
upon significant later burial, such as growth faults. The most
common deformation structures are described below, roughly
in order of increasing scale but moderated by considerations
of process. Several are described in more detail in other
sections.
Load casts and pseudo-nodules
These distortions occur at the interface between a mud layer
and an overlying sand layer. Downwards-facing bulbous load
casts of sand are separated by upward-pointing flame
Convolute lamination
This commonly occurs within sand or silt units and involves
distortion of depositional lamination by folding of variable
intensity. In some cases, the folding is chaotic whilst in other
cases upward movement of escaping water has dragged
lamination into upright folds. Cuspate folds with sharp
anticlines and more rounded synclines are typical. In many
cases,
depositional lamination can still be identified, commonly
cross-bedding or cross-lamination. The good sorting and high
porosities of eolian sands makes them particularly susceptible
to liquefaction {q.v.) when water table rises rapidly through
them.
Convolute lamination (^.
v.)
is a common feature of turbidite
sandstones, often associated with ripple laminated (Bouma C)
interval of internal lamination. Some of the largest examples
occur in eolian sands where a rising water table has caused
instability. All examples reflect liquefaction of the sand, which
can happen for a variety of reasons, including seismic shock,
rapid sedimentation, breaking waves or a shift of water table.
Where seismic shock was the cause of liquefaction, deformed
layers may have very widespread distribution. Upward escape
of excess pore water, internal density inversions and down-
slope components of gravity may all contribute to deforma-
tion. De-watering leads to reconsolidation of the sediment
from the bottom up, "freezing" the deforming laminae.
Protracted liquefaction may lead to total homogenization.
Overturned cross-bedding
Overturned cross-bedding is a special case of convolute
lamination where liquefaction occurred in an actively migrat-
ing bed form. Current shear on the sediment-water interface
dragged the liquefied sand in a down-current direction in an
essentially laminar style. An upward migrating front of
reconsolidation allowed higher parts of the sediment to be
sheared longer, giving down-current facing recumbent folds.
Dish and pillar structures
These structures, which are most common in thick, otherwise
massive sands, result from rapid de-watering following rapid
deposition. Dish and pillar structures (^.
v.)
are direct products
of de-watering and are not distortions of pre-existing lamina-
tion. Dish structures are thin concave-upward, sub-horizontal
zones of slight clay enrichment produced by local filtering out
of the elutriated fines. Their upturned edges merge with pillar
structures, which are vertical conduits of water escape and
record fluidized or near-fluidized conditions.
194
DEFORMATION STRUCTURES AND CROWTH FAULTS
~~~~-~-~...l.oss of strength
Applied -....,___^
Force
~~~~-~-»...^
Gravitational body force on slope
Gravitationally
unstable
density
gradient
Shear
stress
Physical
qual confining load
Continuous
Within a single layer
Multiple layer,
not pierced
Multiple layer,
pierced
Current drag
Vertical
Physical
Exceed Strength
Internal Tensile
(Brittle)
Slides
Growth
fauifs
Soft
sediment
faults
Liquidize
Internal Cohesive
(Plastic)
Slumps
External Surface
Cohesive (Plastic)
Slumps
and slides
Loaded rippies
Shale ridges and diapirs
Shale ridges
and mud diapirs
Liquidize
Liquefied Fluidized
Debris flows
Loaded ripples and sole marks
Convolute
lamination
Clastic dykes
Sand volcanoes
Dish Water escape
structures pipes and pillars
Bedding surface load casts
Ball and pillow/pseudonodules
Isolated load balls
Overturned
cross-bedding
Water escape
pipes and pillars
Figure D3 Styles of sedimentary deformation classified in terms of the interaction between the cause of loss of strength and the applied
deforming force. Adapted from Owen (1987).
Sand injection structures
Where a sandbody is encased in mud and progressively buried,
compaction and de-watering of the muds can lead to
overpressure in the sands. If the fluid pressure exceeds the
tensile strength of the overlying mud, rupturing occurs and
complex networks of dykes and sills of liquefied sand are
intruded into the muds. This is quite common in channel
sandbodies in deep-water and slope settings though outcrop
examples are rare.
Small-scale injection dykes occur in mudstones of inter-
bedded sandstone and mudstone successions. These may be
folded ptygmatically because of differential compaction. Sand
within dykes is typically structureless though some examples
show foliation parallel with dyke margins.
Sand volcanoes and extruded sheets
Some upward-intruding liquefied sand, both dykes and pipes,
penetrates to the free sediment surface where it is extruded.
Where the escape conduits are pipes, sand volcanoes form, at
scales from a few centimeters to a few meters. Where the
intrusions are dykes, fissure extrusion occurs and liquefied
sand may flow as sheets before de-watering and "freezing".
Where sand dykes penetrate muddy debris fiows, extruded
sand may become inter-folded with the upper levels of the
debris fiow.
Desiccation and other cracks
When surface layers of mud or silt dry out they contract and
create isotropie horizontal tension. If this exceeds the tensile
strength of the mud, polygonal (commonly hexagonal) cracks
develop which may be filled with wind-blown or water-lain
sand. Dimensions of polygons are a function of mud layer
thickness, and several scales may coexist. Sub-aqueous
shrinkage of surface mud layers also leads to cracks though
these often form less regular patterns. They are thought to
result from volume changes of clays within the surfaee layer.
Large-scale polygonal cracks also develop through protracted
freeze-thaw processes in periglacial settings.
Slumps and slides
Gravitational body forces acting on sediments lying on a slope
lead to slumps and slides where cohesive (plastic) and tensile
(brittle) strengths are exceeded. In both types of movement,
most shearing is concentrated on discrete basal surfaces. They
differ principally in their internal deformation. Slides are
essentially undeformed but may have internal faults whilst
slumps show plastic folding. The two styles commonly co-exist
in the same mass movement. Internal deformation falls short
of the total mixing and homogenization that characterizes
debris fiows but both slumps and slides may develop into such
fiows.
Mud diapirs
Overpressure develops where muds are buried quite rapidly, as
in a prograding delta. The resultant loss of strength and the
significant overburden lead to both vertical and horizontal
fiowage of the muds. Where the overlying sediment is denser
than the mud, as in an upward-coarsening deltaic succession,
mud may rise vertically as a diapir or mudlump, pushing aside
or penetrating mouth-bar sands. In the Mississippi delta,
mudlumps rise to sea level and create short-lived islands