Middleton G.V. (Ed.) Encyclopedia of Sediments and Sedimentary Rocks
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COMPACTION (CONSOLIDATION) OF SEDIMENTS
165
1998).
Some
of the
water constituting
the
high porosity
is
iiL-lLially bound water due
to the
large surtaee area
of
smectitie
elays.
Fine-grained clay
and
mudstones tend
to
eompact slowly
when loaded
in
sedimentary basins
as
well
as
when loaded
by
eonstruetions.
The
rate
of
compaetion
is a
function
of the
gradual increase
in
effective stress
as the
pore fluid escapes
with
a
corresponding reduetion
in
pore pressure.
The eompositjon
of the
pore water
has
very signilicant
effects
on the
rate
of
compaction
In
line-grained mudstones.
The cation concentration (salinity)
in the
pore water plays
an
important role because
the
cations adsorbed
on the
mineral
surfaces. Marine clays have
a
"house
of
card structure" from
the floccuiation process,
and
may become unstable (quick)
and
collapse
if the
saline pore water
has
been diluted
by
meteoric
ground water (Roscnquist. 1958). Such clays can gain strength
and stiffncs
by
adding salt (NaCI
or
KCI). The compressibility
of
mud is
thus
a
complex tunetion oi" mineralogy, grain size
and pore water composition.
Chemical compaction
Compaction processes involving dissolution
and
precipitation
of minerals
or
other solids
may be
referred
to as
chemical
compaction.
We
then have
to
eonsider
the
chemical processes
involved
and the
thermodynamics
and
kinetics
of the
critical
reactions. Chemical compaction
is
therefore different
in
principal from mechanical compaetion.
We
distinguish
be-
tween
two
types
of
dissolution:
(a) Dissolution
of
grain
or
matrix because minerals
or
associations
of
minerals
are
ehemieally unstable,
for
example, aragonite. gypsum,
and
smectite. Kaolinite
becomes unstable
in the
presence
of
K-feldspar.
Dissolution
of
aniorphic materials like Opal
A may
eause compaction when replaced
by
Opal
CT and
quartz.
Transformation
of
solids like kerogen into fluids (petro-
leum) changes the fluid/solid ratio (void ratio). This type
of
dissolution will only result
in
compaction
if
the dissolving
grains
are
load-bearing. Dissolution
of
pore-filling kaoli-
nite
and
smectite
and
precipitation
of
illite
in
sandstones
(>120-140
C)
will normally
not
cause compaction
in
sandstones. There will however
be a
porosity increase
corresponding
to the
dehydration involved
in
mineral
reactions.
In
mudstones. however, kaotinite crystals
may
be
a
load-bearing part
of the
grain framework
and therefore cause
a
closer packing
of
grains
as
they
dissolve.
(b) Dissolution
at
grain contacts and preeipitation
of
the same
mineral
in the
pore space (pressure solution).
arc
not
under stress (Renard
et al.,
2000).
The
increase
in
solubility
has
been estimated
to be I 3
percent
for an
increase
in effective stress
of
10 MPa (Jahren
and
Ramm, 2000). Due
to
very high activation energies preeipitation
of
quartz
is
however
very slow
and at
normal
pH
values
(5 6)
quartz does
not
precipitate
in
sedimentary basins
at
temperatures below
70
•
BOX' (Bjorlykke
and
tgeberg. 1993).
Chemical compaction
of
quartz sandstones consists
of
four
steps:
(1) Dissolution
of
quartz
at
grain eontacts
or
along stylolites;
(2) Transport
of
silica
in
solution along
a
thin water film
at the
contact;
(3) Diffusion
of
silica
in the
pore water;
(4) Precipitation
of
quartz
on
grain surfaces.
In
the
case
of
highly soluble minerals with
a low
kinetic
activation energy like halite steps
1 3 may be
rate limiting
(Renard
ct af, 2001) and the
system
is
then dissolution-
or
transport-controlled.
In the
case
of
dissolution
at
contacts
between quartz grains
the
transport
of
silica along thin water
film is very complex
and is not a
clear function
of the
stress
(Renard
etaf,
2001)
and the
rate
of
pressure solution
can
then
be ignored.
The
rate
of
quartz cementation should also
be a
function
of the
degree
of
supersaturation with respect
to
quartz
in the
pore water,
and
therefore influenced
by
dissolution
and
transport
of
silica. There
is
considerable
evidence that
the
precipitation
of
cement
is
rate limiting
(Oelkers
etaf,
1996)
for
short diffusion distances
(a few
cm).
The kinetics
of the
precipitation process will therefore
be
controlled
by
temperature.
The
rate
of
quartz cementation
at
constant temperature
is
then
a
function
of the
surface area
available
for
quartz cementation (Walderhaug. 1996; Bjorkum
and Nadeau, 1998). Based
on
these assumptions, eompaetioii
and porosity loss
of
sandstones
can be
modeled
as a
function
of subsidence rates, geothermal gradients,
and
grain textures
determining
the
surface area available
(or
quartz cementation.
As
a
result
of
the larger surface area, finegrained sandstones
will eompact faster than coarse grained sandstones under
the
same temperature condition. This
is in
contrast
to
mechanical
compaction where well-sorted coarse grained sand
is
most
compressible. Compaction
of
sand
at
grain contact
or
along
stylolites requires some stress even
if
the magnitude
of
stress
is
not rate limiting. Petrographic observations suggest that
pressure solution
of
quartz does
not
require very high stress
and that dissolution preferentially occurs along eontacts
between mica
and
quartz (Bjurkum. 1996).
Surface controlled quartz cementation occurs
as a
function
of temperature
and
time
and
sandstones compact even when
there
is no
burial
or
temperature increase (Figure
C74)
(Walderhaug tVi//., 2001).
For reactions with very
low
kinetic reaction rates
at low
temperature like
in
most silicate reactions, temperature rather
than stress
at the
grain contacts
may be
rate limiting
for the
reactions
and
thereby
for the
rate
of
compaction.
Chemical compaction
of
sandstones
The stress
at the
contact between sand grains
is
very high prior
to precipitation
of
mineral cement because
of
the small contact
area. This causes increased solubility
of
minerals like quartz,
and there
is
thus
a
thermodynamic drive
for
dissolving
at
grain
contact
and
precipitation
on
mineral surfaces
of
quartz that
Relationships between compaction and porosity loss
During mechanical compaction
the
volume
of
the solid phase
remains practically constant
and the
change
in
rock volume
(shortening)
is
therefore
a
direct fuiiction
of
the porosity (lluid)
loss.
Chemical compaction implies that the volume
of
the solid
phase
may
change
due to
dehydration
of
minerals
and
other
mineral reaction. Generation
of oil and gas
from solid load-
bearing kerogen
is
another phase change that
has
conse-
quences
for
compaction because
it
causes changes
in the
fluid/
solid ratio (void ratio)
due to
temperature rather than stress.
166
COMPACTION (CONSOLIDATION) OF SEDIMENTS
Changes in the volume of solids in a volume of roek may be
due to net export (leaching) or import (precipitation) of solids
by diffusion or fiuid flow (advection). If all the pore space
between grains is filled by mineral cement where the solids
constituents arc imported from outside, the porosity may be
reduced to zero with little or no compaction. This can occur by
diffusion on a local scale and is typical of concretion, which
0 20 40 60 80 100
Compaction
time
(million years)
Figure C74 Modeling of compaction as a function of temperature and
time even if there is no further subsidt?nce. The rate of quartz
cementation and Ihereby also ihe rate of compaction is a function of
the activation energy
E
()mol 'l and the surface area A Imolcm"'s"')
available lor quartz cementation. Erom Walderhaug ct al. (2001).
can form at shallow depths and provide an indication of
porosity at shallow burial. Dissolution of chemically unstable
grains may increase the porosity when there is a net transport
of solids from the rock volume. This may be the case when
feldspar and carbonate minerals are dissolved during meteoric
water flushing.
Flow of compaction-driven pore water normally has fluxes.
which are very low and insufficient to transport and precipitate
large volumes of cement. When we eonsider larger volumes
of rocks (e.g.. >1.000ni"). the import and export of solids in
solution can be ignored except in very special settings.
Compaction of carbonate sand
Normal marine carbonate sediments consist of grains com-
posed of fossils fragments, ooids. pellets or lithie fragments
which may primarily be composed of aragonite. Mg-calcite or
calcite. In terms of mechanical compaction they are in
principal like siliceous sand and mud. Carbonate grains are
mechanically weaker than quartz grains, but the difference as
measured by experimental compaetion of carbonate sand is
not very great. Chemical compaction of carbonate sediments
carbonate is however very different from that of siliceous
sediments because of the faster reaction kinetics also at low
temperatures. Chemical compaction may therefore start at
shallow depths (low temperatures) shortly after deposition.
Strain
2.5 Km
(80-100°C)
5.0 Km
Depth
Porosity
Stress
Porosity curve
Rock strength due
to cementation .
Mechanical
compaction
X, Over consolidation
due to build-up of
overpressure
Xg
Over consolidation
due to cementation
Effective stress at
over pressure
Effective stress at
hydrostatic pore pressure
Figure C75 Schematic illustration of the effects of mechanical and chemical compaction. During progressive burial sediments may become
over-consolidated by a build up of over-pressure (X,), and by cemenlation (X^i. Further compaction then requires a stress exceeding the
over-consolidation stress clue to mechanical compacticjn, or the "pseudo over-tonsolidation" stress due to cementation.
COMPACTION (CONSOLIDATION) OF SEDIMENTS
t67
Chemical compaction in carbonate sediments includes two
dilTcrcru processes;
(1) Dissolution of thermodynamically unstable mineral. Dis-
solution of load-bearing grains that is, of aragonite may
the cause compaction at shallovi' depth {low stresses).
Dissolution of aragonite may be accompanied by pre-
cipitation of early calcite cement, which will stabilize grain
structure and thus preserve porosity during burial.
(2) Pressure solution. The driving force here is the increased
solubility of minerals like calcite due to stress at grain
contacts, and the tcmperaiure plays a minor role. The rate
of precipitation may also depend on grains surface
properties. Large ealeite erystals, tor example, from
erinoids, may provide very suitable sites (low solubilities)
for precipitation, while very small erystals and organic
coatings may retard precipitation of cement, thus influen-
cing the rate of eompaction.
Chemical compaction and rock strength
Sediments become over-consolidated when the effective stress
is reduced by uplift and erosion or overpressure and can then
be loaded up to the previous maximum stress before any
significant compaction occurs (Figure C75|. Precipitation of
cement in the pores between grains causes an increase in the
strength of the grain framework and produees a pseudo over-
consolidation effect, that is not due to high mechanical
stresses. This "over-consolidation"' may allow a rather high
inerease in overburden stress without any significant meehan-
ical compaction. In the ease of sandstones, quartz cement will
normally become significant at 2.5 km to 3 km depth (80-
100 C). Only 2 3 percent quartz cement may be enough to
strength the grain framework and virtually stop further
mechanical compaction. Fractures in quartz grains formed
by meehanical compaetion usually do not penetrate the quartz
overgrowths, indicating that mechanical compaction stops
after significant quartz cementation (Figure C75). Well-
cemented sandstones have a high rock strength. Consequently,
reducing fluid pressure during production of water or
petroleum will normally not cause mechanical eompaction
and surface subsidence.
Summary
Mechanical compaetion tests of sediments form the basis for
predicting subsidence (settlement) due to the ioading of
eonstruetions like buildings, pipelines or oil platforms.
Compaction tests at high stresses ean also simulate mechanical
compaction in the upper 2 3 km (at temperatures <80 120 C)
in elastic sedimentary basins, while chemical compaetion is
dominant in the deeper and hotter parts. In carbonate
sediments compaetion is more a funetion of stress and is less
dependent on temperature.
Modeling of chemieal compaction may now provide a basis
for quantitative prediction of porosity in deep reservoir rocks.
Compaction processes also control the rock meehanical and
petrophysieal properties, seismic velocities and attributes but
this is still poorly understood.
Knut Bjorlykke
Bibliography
Aplin. A.C. Fleet. A.J.. and Miit-qiiiiker. J.-H.-S. (eds.). 1999. Mud
and Mudstones: Physical and Fluid Flow Properties. Geological
Socioly of London. Special Publication. 158,
Alhy. L.F., 1930, Density, porosity and compaction of sedimentary
rocks.
Bulletin of the American Assoeiation of Petroleum Geologists.
14:
I -24.
Bjorkum. P.-A.. 1996. How important is pressure solution in causing
dissolution of quartz in sandstones:' JournalSedintantary Research.
66:
147-154.
Bjerkum. P.A.. and Nadeau. P.M.. 1998. Temperature controlled
porosiiy/permeahility reduction; fluid migration and petroleum
exploration in sedimentary basins. .APPEAJmtrnal. 38: 453-464.
Bjorlykke. K.. 1998. Clay Mineral diagenesis in sedimentary basins—
;i key lo tlie prediction of rock properties. Clav Minerals. 33:
15 34.
Bjorlykke. K.. and F.geberg. P.K.. 1993, Quartz cementation in
sedimentary basins. Bulletin of the American Assneiation of Petro-
leum Geologists. 77: 1538 1548.
Bjorlykke. K.. and Hoeg. K.. 1997. FJTecLs of burial diagenesis on
stresses, compaction and fluid flow in sednnentary basins. Marine
and Petroleum
Gevtoi-y.
14: 2bl-276.
Cliulian. F.A,.Kjeldstad. A.. Bjorlykke. K.. and Hoeg. K..
2002.
Porosity
loss in sand by grain crushing experimental evidence and relevance
to reservoir quality. Marine and Petroleum
Geoloi^y.
19:
39-53,
Fisher. Q.J.. Casey. M.. Clenell. M.B.. and Knipe, R.J.. 1999.
Mechanical compaction of deeply buried sandstones of the Norlh
Sen. Marine and Petroleum Geology. 16: 605-6IS^,
Giles.
M,R.. 1997. Diagenesis:
.4
Quantitative
Perspcetive.
Implications
for Basin Madellingand Roek Property Prediction. Kluwer Academic
Publishers.
fhibheri. M.K.. and Rubey. W.W.. 1959. Role of fluid pressure in
mechiuiies ofovcrthrust faulting 1. Mechanics of fluid lilled porous
solids and its application to ovorthrust faulting. Bulletin of the
Geologieal Soeiety of .Ameriea. 70: 115 166.
Jahren. J.. and Ramm. M.. 20()(), The porosity preserving effcets of
microcrystalline quartz coatings In arenitic sandstones: examples
from the Norwegian continental
siielf.
In Quartz Cementation in
Sandstones. Association of Sedimentoloszisls. Special Publication,
29.
pp. 271-280.
Lambe. T.W.. and Whitman. R.V.. 1979. Soil Mechanics. SI lersion.
John Wiley,
Nadeau. P,H,. and Bain. D,C,. 1986. Composition of some smectite
and diagentic illitic clays and implications for their origin. Clays
and Clav Minerals. 34: 455 464,
Oelkers. E.H.. Bjorkum. P.A.. and Murphy. W.M.. 1996. A petro-
graphic and computational investigations of quartz cement and
porosity reduction in North Sea sandstones. American Jmirnal of
Seieme. 296: 420 452.
Pittman. E.D.. and Larese. R.E.. 1991. Compaction of lithic sand:
experimental results and applications.
.American .Association
of Pet-
roleum Geologists Bulletin, 75: 1279-1299.
Renard. F.. Brosse. E.. and Gratier. J.P.. 2000. The different pro-
cesses involved in the mechanism of pressure solution in quartz-rich
roeks and ihcir interactions. In Woiden. R.li,. and Morad. S,
leds.).
Quartz Cementation
iti
Sandstones. International Association
of Sedimenlologists. Special Publication. 29. pp. 67 7S.
Renard. F,. Dysthe. D,. Feder. J.. Bjorlykke. K.. and Jamtveit. B..
2001.
Enhanced pressure solution creep rated induced by clay
particles: experimental evidence from sail aggregate. Geophysieal
Researeh Letters. 28: 1295-1298.
Rieke, H.H,. and Chillingarian. G.V,. 1974. Compaetion of Argillaeeous
Si'dimcfit.s. Developments in Sedimenlology. 16. Elsevier.
Rosenquist. I.Th.. 1958. Remarks lo tlie mechanical properties
of soils—water systems. Geotogiske Foreningens Stoekholm. 80:
435-457.
Sclater. J.G.. and Christie. P.A.F,. 1980. Continental siretching. An
explanation of the post Mid-Cretaceous subsidence of ihe Central
Norlh Sea Basin. Joumitl of
Geophysieal
Researeh. 85(B7): 3711-
3739.
Ter7agbi. K.. 1925. Frdhaunieckanik Auf
Bodcnphy.sikalischer
Grung-
lage. Leipzig: Deuiicke.
168
CORING METHODS, CORES
Thyberg. B.L. Jordt. H.. Bjorlykke. K.. and Falcide. J.I.. 20O0.
Relationships between sequence stratigraphy, mineralogy and
geochemistry in Cenozoie sediments of the northern Norlh Sea.
In Noitvedt, A. etal. (ed.). Dynamics of the Norwegian Margin.
Geological Society of London. Special Publication. !67. pp. 245
272.
Walderhaug. O.. 1996. Kinetic modeling of quartz cementation and
porosity loss in deeply buried sandstones reservoirs. Bulletinof the
.American
Association ofPetroletim Geologists. 80:
731
-745.
Wiilderhaug. O.. Bjtikum. P.A.. Nadeau. P.H.. and Langnes. O.. 2001.
Quantitative modelling of basin subsidence caused by temperature-
driven silica dissolution and repreeipitalion. Petroleum
Geo.seience,
7:
107-114.
Wilson. M.D., 1994. Reservoir Quality Assessment and Prediction in
Clastic Rocks. SEPM Short Coursc'30. 432p.
Cross-references
Atterberg Limits
Cements and Cementation
Diagenesis
Diagenetic Structures
Diffusion. Chemical
Fabrie. Porosity, und Penneabilily
Geophysical Properties of Sediments
Porewaters in Sediments
Pressure Solution
Slylolites
CONVOLUTE LAMINATION
A term ititrodiiced by Kiienen (1953) for irregularly wavy
laminae confined within a single sediment layer. The deforma-
tion dies out both downward, and upward, and may be
confined to a part of a bed: it does nol atTect the geometry of
the lower or upper boundaries of the bed. It is commonly
found in turbidites. but is also present in other facies
(Figure C76). The convolutions appear to develop gradually
above some level within the bed. frequently distorting
otherwise plane or ripple cross-lamination, and may show
one or more surfaces of erosional truncation before dying out
upward, generally at or just below the top of the bed. Thus
they are generally interpreted as having developed during or
immediately after deposition. Good figures are provided by
Allen (1982) and Ricci Lucchi (1995).
The actual structure was described earlier by several authors
using several different names (see Potter and Pettijohn. 1963.
p.
152; Allen. 1982. p. 349). and its abundance was noted
particularly by Rich (1950) who thought the structure had
been formed postdepositionally by intra-stratal sliding, and
had therefore developed on submarine slopes. Bouma (1962)
thought that the structure developed in. or as a replacement of.
the rippled (C) division of his ideal turbidite "sequenee."" but it
is now clear that it may affect either the plane laminated (B) or
rippled (C) divisions. The strueture has mueh in eommon with
Pillow Strueture
(</.
r.). whieh has also been misinterpreted as a
slope indieator.
Most convolute lamination is found in fine grained sand or
coarse silts, and has been reported (rarely) from limestones as
well as from siliciclastics. Allen (1982) distinguished three
types:
postdepositional. syndepositional. and metadeposi-
tional, depending on the geometry (the metadepositional
variety shows erosional truncation at the top of the bed).
The mechanism of formation is deposition that is rapid enough
to produce pore pressure exceeding hydrostatie. and conse-
quent loss of strength (partial liquefaction) of the sediment
either during or immediately following deposition. Gravita-
tional instability then produees "diapir-like"" structures, and
shear by the depositing flow commonly results in a predomi-
nant orientation and direction of overturning of the (generally
very irregular) folds. Thus, some convolute lamination can be
used as a paleoeurrent indicator.
Gerard V. Middleton
Bibliography
.'\llcn. J.R.L.. 1982. Sedimentary Struetures: Their
Character
and
Physi-
cal
Basi.s.
FIsevier. 2. pp. 349-354.
Bouma. A.H., 1962. Sedimentology of Some
Ftysch
Deposits. Elsevier.
Kuenen. Pb.H.. 1953. Significant features of graded bedding. Ameri-
can Assciation of Petroleum Geologists Bulletin. 37: 1044-1066.
Potter. P,E.. and Petlijohn. F.J.. 1963.
Puleocurrents
and Basin Analy-
sis.
Aeademic Press,
Ricci Luechi. F.. 1995. Sedimentographiea: Photographic Atlas of Sedi-
nientctry
Structures. 2nd Edn. Columbia University Press.
Rich. J.L., 1950. Flow markings, groovings, and intrastralal crum-
plings as criteria for recognition of slope deposits, with illustrations
from Silurian roeks of Wales. Anterican Association of Petroleum
Geologists Bulletin. 34: 717 741.
Figure C76 Syndepositional convolute lamination. Redrawn by Allen
(1982.
p.;i49. bis Figure 9.4) from observations made on the Pliocene
turbidites of tbe Ventura Basin, California. See Allen's book for original
source.
Cross-references
Liquefaction and Fluidization
Pillar Strueture
CORING METHODS, CORES
Cores are a very important source for the study of sediments.
The oil and gas industry prefers to eliminate or strongly reduce
coring by collecting a suite of well logs and thereby reducing
well eosts. Exploration and produetion people nevertheless
need some cores for grain size, grain shape and surface
ORINC
METHODS.
CORES
169
r=igur~
en
lwo
t)~
of
L'Ufing
devices.
CAl
Vol"
u.atl·n tube
ust.'C!
ofl~
in
W;lt
up
to 10m deq). 1'hev.llve
an
be
SWIJOR
to
the-ski
Jnd
the
core pushed
out,
unl
S J
liner
is
used.
Three
meIer
lon~
l'xh."f'ISlon
rods
make
II
po~ihle
10
wen
from
a
sm.ill
bual.
(8)
PISlon
(Of(Of
operatJoo.
I,
m.\in able; 2.
lIigg<.'f
devil(! wire-cl:unp;
J.
triW""
;trm;
4,
stabilizing
farK;
S.
'~ejgh15;
6, piSlon;
7,
l(iggCf
\\-
ighl.
(SI)
lowering.. 182) inlliotl penetration;
plSlon
Slot)"
just ahov(' Ihe
bolt
m.
connected properly.
the
piston should SLOP a soon
as
It
IS ncar
the
""tier
sediment Interface:. while
the
tore: pipe slid
ovcr
lhe piston. penetrating the
Sl..""<fimcnt
O-iGure
7·82). Because
of
the
pi
Ion a light \'aeuum
is
created undcntC'J.th
it
whi b
reduce! the friction
or
sediment
\\
itllln tbe liner.
This
type
of
pislon corer. weighing about one ton. and having a pipe length
of
12-1 m. mil)' collect :t core
up
to
12
m
lonl.
linc
melers I
more onllnon.
If
lhe length sctting Wit wrong.
or
the
clamping devi
OPCIU
too early. Ihc piston
an
suck in
scdim('f\t. ch<lfacterized
by
Row-in
structures (Figure 7S·IJ).
A
Giant
Pi
ton
orer
was de\e1oped
at
the Woods Hole
Oce-dnogmphlc
In
thudoR with the goal to collect
SOm
long
C'or~
The
lighter Jumbo Piston Corer(2.300 kg) has been used
Sit ccssfuUy
in
lhe Gulf
of
Mexico. Recovered
rorC'S
wcre up
10
19.201
in
length. relriC\'oo from 800 2.600m water
depth
( ih'a
and
Dry.lIlt. 20(0). he extended length
is
necessary for
geological
and
gooHx:hnical siudie
10
bettet
understand Ihe
processes
of
slumping. sliding. debris
tlO\\S
and turbidit}'
('urfCntS.
Anything longer
lhan
the dc\rices mentioned abO\e has
10
be
donc
with rotary drilling front a special
\'c.~1
SII
has
opei..l1ed
b)' the Deep Sro Drilling
Proj\.'Ct.
hs
uccessor thc Ocean
Drilling
Project.
or
Ihe drilling. \'
Is
working for the oil
indu.\lry.
DL-cp
drilling on hon:
is
simpler
bL'culIsc
Ihl: rig is
stable
rather
Ihun heming
and
other
motions.
oring
;11
M:a
or
land can
proout"C
cores 8
10
J!
"(20.S
109(,'01) in dillmC:ler wilh
n
COUlmon
diameter
of
~
.. (lOent).
The
mining induslry. lind
most
other
shallow drilling
rL'<Iuircmen~
uses
Ihe
limh
Ie
\ersion.
producingcorts
tlllH
nre 2
11t
to 1
1
-
"(6.3
103.
em,
in
diameter
~ilh
a common sire
of
(711
..
(4.7<:m).
Th
sc:
ngs
lin:
INmAl
PENETRATION
6.
'
6
1
=
•
A
•
~
A
"---.-/
chJtfllClenstics. mincrnlog)! nnd dingcncsis. lind for
the:
culibnnion
of
seismic lind
well
log
<lilla,
MO~1
'iL-dimenll"If)'
gwlogist working
:"
unin:rsilics.
govcTnrocI11
und the
minin'
indus!')' prefer
10
denl with cores. Because
of
the
\'nricly
ill
interests lind tbe Inan)' types und hardncsscs
of
sediment. an e.(tcnsivc rtrruy
of
coring methods has been
dc\'doped.
The)' mogc from implc h{md-opcralcd samplers
on
lund lind hallow walcr.
to
large
rmd
expenslve drilling
opemliol1s.
The specific
nC'l"d
and the
llVllilublc
bud!-'CI
dicUHc
the
~ilc
and type
or
OJ
selccted coring
dc,~.
1I
i also
po
iblc to
c~lcgori7c
coring equipmellt by .he
de
ired
length
lind
diameter
of
nocdcd core. the
hard"
s
of
the sediment, ilnd if
the coring occurs
on
bore
or
ofT~horc.
The
'mull h:tlld-
opcnHcd lools h:t\·c the
IldV~tnHIgC
that they
an:
light lind
CUR
be
tntnsporccd
evC'~'hcn:.
The
large
\,<.ncty
of
coring
dcvil"CS
611s
more
than
.1
book
and
therefore onl)"1
few
major
t}'pes
will
be
discusstd. Typic.lll)' a core i a long. c)'lindricHI
objcct.
~
c\'crthcl the
tenn
"core" refer also to
other
sh21pc:s,
such
as
box core. box sample. squun: long box core.
and some olhcrs.
Sampling unconsolidated sediments
on
land
or
in
wading
depth
is
easy
:.md
inc.xpensh·e a long a a
shan
sample
is
sunicient.
MCIlII
pipes. plastic
or
PVC tubes.
or
even mcl.II
COtns
with a hole
in
the bottom
to
let air
or
water
OUI.
have
prOVided
@rt':u
result.
or
work
ofT;t
smull
bo,\!
ne
C~tn
11
C
lllel'll. pl:lstic
or
PV tube with exten ion rods
(8
uma. 1969:
Pigures C77·A. C78-A).
In
01051 casc a core
e.1t
her i
required. A second pair
of
hands
facilitates the
operation
and
pJ'(.\"Cnts
hitLing
the
boat
5C\'eral times. wlueh
onen
result in
core
los.
Sm<l1J
bo~es.
such as the Senckcnbcrg box (Reineck.
1957)
do
well on
hUld
and
nt ","ding depth. c\'cn
in
sands.
l'cLri
dishes
arc
being used ucccssrully because
the
S3mplc
Co1"
be
rndiogf'oll,hed while
lill
in
the
di
h.
or
pholoG,r3phed. dried.
imprcg.mlled. etc.
Mo~ing
into deeper
\\alcr
requil"C5
.i
ship. lhe sUe
deptnding
on waler
dtpth.
Bnd
the size
BIIlI
weight
of
both the cortr
and
the winch.
The
010S1
common
corers :tre gravity corers
and
piston
corer.
Both C'.Jltgories collect the core
in
II
pire. In
Inosl
ca
e-s
Olle
uses
It
mel;,1 pipe
Blld
It ptastie liner Inside.
.M)u\clill\(.'S
bOlh
C'.ln
be replaced h)'
one
P C tube. A bevel d
coring
nose.
WiLh
1_
core
cat her inside. i secured to lite base
of
the pipe.
The
lOp
of
tht
pipe fils in;;idt a stiglllly wider pipe.
which
tub a hook (bail)
:11
the
tOp
wilh tnbilizing tins
underneath
and
"ei
'Ill belo\'.· the
fin~.
A \'ah'e
on
lOp
of
the upper pipe
doses
directl)'
after
the corer 1 nninnles
penetration.
R:Hher tit;]"
IClIing
the corer frte-fall
OH:r
1he
la't
3-
m,'l
tripping device is
u"Cd
Ihal gives a
pre-n1C3surc,,"(1
rrec·fall
(Doul11~l.
1969).
It
con:sisLS
of
nn 11ml
CQllnt:CtL"(l
to
the mnin
cltble
vin
a c,lblc..elumping device (Figure 7·0I.
1)2).
t the
end
of
lite
3nn
is
11
C1tble
(corer length + frec.falllol1g) with a
weight. soon ;1$ Ihe weight hits the sediment
it
unhook the
cubic
nllh~
chunring
d<::\'icc:
and
frcc-fall
~Inrts.
Du~
to
friction
between the core pipe
and
Ihe sediment.
one.
seldom
collL'(;1
tI
corc over
2-
m
in
length.
The
friction along the in.side
of
the
liner
fCSuh
in
a
core:
shortening. nmging from
10
erln.
The
piMon
(.'Orer
is \'C:f)' sitnilM
10
the:
gmwty
COrer.
excepl
that a pislon i inserted
Itt
the
b",sc
of
the liner.
The
main tble
i eonncctc:d
10
the piston. A
loop
of
cilble
(frec.r~llllcngth)
is
rolled
up
between
.he
fin
( igure C77·81), The trigger cable
htls
lhe lenglh
or
the corer plus
frcc:-rnll.
If e\'erything
is
170
CROSS-STRATIFICATION
Figure C78 (A) Part of a core (2.Scm wide) collected by hand with a
short lube. Point bar. Mississippi River. Larye current ripples. Courtesy
|.M.
Coleman. (Bl Piston core in plastic liner. Upper 1/3 normal core,
lower 2/3 flow-in. Core
5
tm wide. Surinam shelf, water depth 72 m.
for land use and typically are mounted on a medium-sized
truck.
Cochran (1959) and Bouma (1969) have described several
varieties of coring devices. Amateurs with some experience can
handle small devices. Professionals should handle the larger
tools to avoid serious aeeidents and poor core recovery.
Arnold H. Bouma
Bibliography
BoLima. A.li.. 1969. Methods
Fin
The Study Qf Sedimentary Structures.
New York: Wiley-InlLTsciL-ncc. (Can be obtained from Krieger
{'tiblishing
Co.. Huntington. New York. 197'-).)
Cochran. J.B,. 1959. SamplingTechniipie.s. 2nd edn. Wik-v and Sons.
Reineck. H.E,. 1957. Stechkaslen und Det-kwcisz. Hilfsmiltcl des
Meeresgcologen- Naturlblk. 87: 132-134.
Silva. AJ,. and Bryant. W.R.. 2()()(). Jumbo pislon coring in deep Wiiler
Gulf of Mexico for seabed geohazard and geoteehnieal investigii-
tions.
In
Tenth
International
Offsliore
and
Ptitar
Engineering Confer-
ence, pp. 424 433.
Cross-references
Bedding and Inlernal Structures
X-ray Radiography
CROSS-STRATIFICATION
Introduction
Cross-strata are layers of sediment that are inclined relative to
the base and top of the set in which the inclined layers
are grouped. Each group is called a set of cross-strata or a
Figure C79 Ex^imple ot eolian cross-slratification. Bed "A" (hetween
heavy white linesl is a cro^s-stralified bed or set of cross-strata.
Because the cross-strata (such as bed "B". between thin white lines)
are themselves tross-stratified (bed "C", for example), this structure is
called compound cross-stratification.
cross-stratified bed (Figure C79). Individual cross-strata can
be classified (MeKee and Weir. 1953) as eross-laminae (<1 cm
thick) or cross-beds (>lem thick). In general, each set of
cross-strata is deposited by a migrating bedform. Thin sets are
deposited by small migrating bedforms sueh as ripples, small
dunes.,
or small antidunes, and thiek sets are deposited by
larger dunes, antidunes, bars, or other large bedforms (Allen,
1962;
Harms and Fahnestock, 1965). Cross-strata are a natural
record of transported sediment and arc therefore useful for
understanding the behavior of modern bedforms and for
interpreting environments in ancient deposits.
Beeause most cross-strata are deposited by migrating
bedforms, study of cross-strata is inextricably entwined with
studies of bedforms. sediment transport, and lee-side proeesses
including granular fiow. Advances in the understanding of
eross-stratifieation include recognition of cross-laminae de-
posited by ctirrent ripples (Sorby. 1859), recognition of
climbing eolian dune deposits (Barrell, 1917: Shotton, 1937),
experimental studies of graintiows (Bagnold. 1941). discovery
of large subaqueous dunes using sonar, recognition that
subaqueous dunes deposit cross-strata (Allen, 1962, 1963a:
Potter and Pettijohn, 1963: Harms and Fahnestock, 1965),
recognition of stratifieation types in eoHan dunes (Hunter.
1977),
and the use of 3-D computer graphics to visualize the
migration of eomplex bedforms (Rubin, 1987).
Origin of cross-stratified beds
Deposition of cross-stratitied beds requires two eonditions:
migration of geomorphic surfaces that are inelined relative to
the generalized depositional surface and deposition during
migration. The inclined surfaees ean either be individual
landforms (such as delta-like bodies) or can oceur within a
train (bedforms): most sets of cross-strata are produeed by
trains of migrating bedforms.
Bedforms can only leave deposits where deposition occurs
while the bedforms are migrating. This combination is defined
as bedform climbing, because deposition causes the bedforms
to move upward (climb) while they migrate. Where the rate of
deposition is rapid enough (relative to the rate of migration),
both the stoss and lee slopes of the migrating bedforms are
preserved, resulting in complete preservation of the original
bedforms (supercritical climb of Hunter, 1977). In other cases,
the stoss sides of bedforms ean undergo erosion while the
lower lee sides arc prefcrcnlially preserved (subcritical elimb).
CROSS-STRATIFICATION 171
Origin of individual cross-strata
Sediment grains within cross-strata are deposited by a variety
of processes: settling out from the water or air, avalanehing
(grainllow). or deposition by superimposed bedforms. On
steep lee slopes, initial deposition is primarily by grains settling
out of the flow. In both air and water, sediment can be sorted
by grain size or density as il sellles out on the lee slopes. On
wind ripples, segregation by size ean mask or obseure eross-
laminiHion. Instead ol" producing visible cross-lamination,
wind ripples typically deposit inversely graded laminae,
because the grains on the lee sides of the wind ripples are
sorted from coarse at the crest to fine at the trough (Hunter.
1977).
Cross-stratifieation deposited by grainfiow is eommon on
steep lee slopes of bedforms. Grainflow oeeurs beeause
sediment that is transported aeross a bedform crest tends to
be deposited at higher rate on the upper lee side than the lower
lee side. Consequently, the lee slope can steepen until the slope
exceeds the angle of repose, reaches the higher angle of initial
yield, and causes the slope to fail hy grainllow. In air,
grainflow is discontinuous in time and space, producing
grainflow deposits that are lenticular in horizontal cross-
section. In contrast, subaqueous grainflow can occur more
continuously in both time and space (Hunter and Koeurek.
19S6),
In both air and uater. grainflow sorts grains by size,
wiih the finest grains tending to sink to the base of the flow
(Bagnold. 1941). On low-angle lee slopes, cross-strata are
commonly deposited by superimposed bedfonns.
Classifications of cross-stratification
McKee and Weir (1953) classified cross-stratification on the
basis of descriptive features sueh as the shape of the set and
shape of individtial cross-strata. This is a useful approach,
because some of these attributes are related to the kinds of
bedforms that deposited the eross-strata (Allen. 1962). For
example, sets of cross-strata with a planar base and uniform
dip azimuth are deposited by two-dimensional bedfonns
(siraight-crcsted bedforms without seour pits in Iheir troughs),
whereas trough-shaped (festoon) sets of eross-strata are
deposited by more sinuous bedforms with seour pits in their
troughs (Figure C8()). Rubin (1987) also followed a geometric
approach and classified cross-stratilication based on properties
that relleet whether the strata are constant along-strike (i.e.,
deposited by two-dimensional bedforms). constant down-dip
(i.e.,
deposited by bedforms that maintained the same shape
while migrating), and properties of the eross-stratifieation that
indicate whether the original bedforms were transverse,
oblique, or parallel to the resultant sediment-traEisport vector.
Other classifications have been proposed for cross-stratifiea-
lion and bounding surfaees deposited by specific subaqueous
or eolian bedforms.
Kinds of cross-stratification
Compound cross-s[ratitication
Some cross-stratification is complicated by containing cross-
strata that are internally cross-stratified (Figure C79), a
structure defined as eompound cross-stratitieaiion {Harms
etal.. 1975), double cross-bedding, or furious cross-bedding
(Reiche, 1938). Compound cross-stratification is formed where
Figure C80 Trough cross-stratificatiun. (A) Simulated bedform surface
and stratification. Each irnugh-shaped set of strata is deposited by a
migrating scour pit in the bedtorm trough. Arrow shows migration
direction.
(B) Horizontal section through structure in (A). IC) Eotian
example of struclure shown in (A) and (B). Dunes and scour pit
migrated left-to-right out of the outcrop.
the iee side of a bedform undergoes changes in shape (often
accompanied by erosion). Shape can change in response to
changes in flow (such as ebb-flood reversals in tidal eurrents or
annual changes in wind direction), or by migration of
bedforms superimposed on a larger bedform as described by
Allen (1963b, epsilon cross-stratification) or MeCabe and
Jones (1977). Some workers refer to bounding surfaces within
sets of cross-strata as reactivation surfaces, but if "reactiva-
tion"" is taken literally (he term is more appropriate to surfaces
produced by flucluating-flow, because superimposed bedforms
ean produce bounding surfaces under steady conditions (no
reactivation required).
Compound cross-strata deposited by bedform assemblages
are particularly useful for determining the flow direction
172
rROSS-STRATIFlCATION
relative to the bedform crest. For example, the best method of
identifying an oblique bedform is to demonstrate that
bedforms superimposed on the lee side of the main bedform
systematically migrated along the crest of the main bedform in
a preferred direction (Rubin and Hunter, 1983).
Cyclic cross-stratification
Some cross-stratifleation exhibits cyclic changes in grain size
(such as mud drapes in tidal bundles), dip azimuth or
inclination, stratification type (e.g.. alternations of grainflow
and grainlall), or spacing of bounding surfaces. In herringbone
eross-stratification (attributed to reversing tidal flows) succes-
sive sets dip in opposing directions. Scalloped cross-bedding
eontains cyclically truncated trough-shaped subsets. Some
cyclic eross-strata arise from flow cycles, and others arise from
the passage of superimposed bedforms.
Deformed cross-stratification
Because cross-strata arc deposited on slopes, they are
particularly susceptible to slumping. Slumping can disturb
individual beds or a sequence of beds near the free surfaee.
Bedforms can also fail by the crestal region of a bedforms
subsiding and sediment flowing into trough regions. This kind
of failure may include the lower set boundary as well as
sediment in the subsurfaee.
Relation of set thickness to bedform size
Where bedforms in a train are similar to each other and have a
relatively uniform trough elevation, the thickness of the
preserved sets depends on the bedform wavelength and the
angle of elimb (Rubin and Hunter, 1982). In such situations,
the thickness of cross-stratified beds ean be a very small
fraction of the original bedform height. For example, eolian
dunes in large deserts are commonly lOOm in height, whereas
eolian eross-stratified beds in the rock record are typically a
full order of magnitude smaller.
On the other extreme, where bedforms have irregular trough
elevations, the thickness of cross-stratified beds depends on the
particular sequence of troughs passing a point on the bed
(Paola and Borgnian, 1991). In this case, the thickness of cross-
stratitied beds ean be a substantial proportion of the original
bedtorm height.
Relation of dip direction to flow direction
Dip directions of cross-stratilication are routinely used to infer
paleoeurrent directions. Although this approach probably
gives reasonable results when sufficient measurements are
made, the results can be biased in Individual cases, particularly
where bedforms are oblique to flow (Rubin and Hunter, 1987).
Trough axes may be a more representative measure of the flow
direction, but even trough axes can diverge from the actual
transport direction. For example, in oblique bedforms, both
the cross-strata dips and trough axis trends can diverge
substantially from the transport direction. A more general rule
ean be demonstrated Ibr compound cross-stratification depos-
ited by bedform assemblages: the transport direction lies
between the dip direction of the bounding surfaees scoured by
the superimposed bedforms and the dip direction of foresets
deposited by superimposed bedforms (Rubin and Hunter,
1983;
Rubin, 1987).
Techniques
Several new techniques are particularly useful for studying the
geometry of cross-stratification. Ground-penetrating radar has
proven successful at mapping internal stratification within
eolian dunes, and borehole imagery can determine vertical
profiles of dip directions. Three-dimensional graphics modeling
is essential for reconstructing the morphology and behavior of
bedforms that are too complicated to be visualized otherwise.
Future studies
Perhaps the greatest advance in understanding eross-stratifiea-
tion will eome with the development of coupled flow and
sediment-transport models that ean aeeurately predict the
morphology, migration, and evolution of bedforms. including
interaction between successive bedforms in a train. Such a
model would allow computational experiments to recreate
specific bedding geometries from simulated flow sequenees.
David M. Rubin
Bibliography
.Allen. .I.R.L.. 1962. Asymmetrical ripple marks and the origin of cross-
slralification: ^'aturc. 194: 167-169.
Allen, J.R.L.. 1963a. Asymmclrical ripple marks and the origin of
water-laid cosets of cross-strata. Liverpool Manchester Geologieal
.lournal. 3: 187-236.
.Allen. .I.R.L., 1963. The classification of cross-stratified units with
notes on their origin. Si'dimentology. 2: 93-114.
Allen. J.R.L., 1981. Palcotidal speeds and ranges eslimated from cross-
bedding sets wiili mud drapes. Nature. 293: 394 396.
Bagnold. R.A.. 1941. The Physies of Blown Sand and Desert Dunes.
London: Metheun.
Barrell. J.. 1917. Rhythms and the measurement of geological time.
Butle/in of the
Geological
Soeiety of Ameriea. 28: 45-904.
Harms. J.C. and Fahnestock, R.K.. 1965. Stratification, bed forms,
and flow phenomena, with an example from the Rio Grande. In
Middieton. G.V. (ed.). Prituary Sedimentary
Structure.s
and Their
Hydrodynamic Interpretation. SEPM Special Publication. [2, pp.
84-11
<.
Harms, J.C. Southard, J.B.. Spearing. D.R.. and Walker, R.G., 1975.
Depo.sitional
Environments as Interpreted from Primary Sedimentary
Siriietun's (aid Stratifieation Sequenees. SEPM Short Course 2.
Dallas: 161 pp.
Hunler. R.E.. 1977. Basic types of stratification in small eolian dunes.
Sedimentology. 24: 362-387.
Hunter. R.E., 1985. A kinemalie model for the structure of lee-side
deposits. Sedimentology. 32: 409-422.
Hunter. R.E.. and Ki>curek, G., 1986. An experimental study of
subaqueous slipfaee deposition, .lournal of Sedimentary Petratogy,
56:
387-394.
Hunler. R.E., and Rubin, D.M., 1983. Inierpreling cydie cross-
bedding, wiih an example from the Navajo Sandstone. In
Brookfield, M.E.. and Ahlbrandt. T.S., (eds.), Folian Sediments
and
Proce.sses.
Amsterdam: Elsevier, , pp. 429-454.
Jopling, A.V., 1965. Laboratory study of the distribution ofgrain sizes
in eross-beddcd deposits. In Middleton. G.V., (ed.). Primary Sedi-
mentary Siruetitrvs and their Hydrodynaniie Interpretation. SEPM
Special Publication, 12, pp. 53-65.
McCabe. P,J.. and Jones, C.M., 1977. Formalion of reaciivation
surfaees within superimposed deltas and bedforms. Journalof Sedi-
mentary
Petrology.
47: 707-715.
MeKce, E.D., and Weir, G.W.. 1953. Terminology for stratification
and cross-slratification in sedimentary rocks. Bulletin of the Geolo-
gieal Soeiety of Ameriea. 64: 381 389.
CYCLIC SEDIMFNTATION
173
Paola.
C.
and Boryman.
E..
1991.
Reconstructing random topography
from preserved slratifieation. .Sedintentology.
38: 553 565.
Potter,
P.E., and
Petlijohn,
F.J..
1963,
Paleoeurrents
and Basin Analy-
.T/'.v.
2nd edn.
1977. Springer-Verlag.
Reiehe,
P.. 1938. An
anaiysis
of
cross-lamination:
the
Coconino
Sandstone. Journal of Geology. 46: 905-932.
Rubin. D.M.. 1987.
Cros.s-Bedding.
Bedforms,
and Paleoeurrents. SEPM
Concepts
in
Sedimeniology
and
Piileonlology,
I, 187p.
Rubin. D.M.,
and
Hunter, R.E., 1982. Bedform climbing
in
theory and
lUitiire. Sedinu-ntology.
29: 121 138.
Rubin,
D.M. and
Hunter,
R,E., 1983,
Reconstruciing bedform
assemblages from eompound erossbedding.
In
Brooklield.
M.E,,
and Ahlbrandl.
T.S..
leds.). Folian .Sediments
and
Proees.ses.
BLsevier. Amsterdam,
pp.
4(17-427.
Rubin.
D.M,. and
Hunler.
R.E.. 1987.
Bedform lilignment
in
directionally varying fiows. .Science. 237: 276-278.
Sorby. H.C.. 1859.
On the
struetures produced
by
ihe currents present
during
the
deposition
of
stratified rocks. TheGeologist. 2: 137-147.
Shotlon,
F.W., 1937. The
louer Bunter Sandstones
of
North
Worehestershire
and
Easl Shropshire. Geological Magazine.
74:
534-553.
Cross-references
Angle
of
Repose
Bedding
and
Inlernal Structures
Bedsel
and
Eaminaset
Convolute Lamination
Cyclic Sedimentation
Deformation
of
Sediments
Dunes. Eolian
Eolian Transport
and
Deposition
Grain Flow
Hummocky
and
Swaley Cross-Stratifieatiori
Paleoeurrent Anaiysis
Ripple. Ripple Mark,
and
Ripple Structure
Sediment Transport
Sedimentologists
Turbidites
CYCLIC SEDIMENTATION
Introduction and background
Cyclic
or
rhythtnic sedimentation
is an old
stratigraphic
problem that
has
received considerable attention
in the
past
(Suess,
1888;
Grabau,
1913;
Wanlcss
and
Shepard,
1936:
Wheeler and Murray, 1957; Wells, 1960; Merriani, 1964; Vella.
1965;
Dull
('/(//..
1967; Elam
and
Chuber, 1972; Schwarzacher,
1975),
and
remains today
an
unresolved controversial topic
(Wilkinson,
1982;
James. 19t^4; Algeo
and
Wilkinson,
1988;
Finsele etal.. 1991; Goldhammer
c/t//..
1993; Drutnmond
and
Wilkinson. 1993a,b.c;
de
Boer
and
Smith,
1994;
Read,
1995;
Grammer
etaf. 1996:
Wilkinson
etaf, 1996;
Miall, 1997).
Most,
if
not ail, stratigraphic successions display repetitions
of
strata,
at
different scales, that reflect
a
succession
of
related
depositiona! processes
and
environmental conditions that
arc
repeated
in the
same order.
The
repetition
of
such events
is
termed cyclic
or
rhythmic sedimentation,
and
this leads
to the
formation
of
stratigraphic successions that display repetitive,
orderly arrangement
of
dilTerent kinds
of
sediments
or
rock
types.
Studies
of
high-frequency cyclicity span
all of
geologic
time from
the
Proterozoic (e.g., Grotzinger,
1986;
Eriksson
and Simpson, 1990; Ciough
and
Goldhanimer, 2000)
to the
Pleistocene (e.g.. Mesolella etal.. 1969; Blootn
etaf.
1974).
The essential concept
of
cyclic sedimentation
i.s
that
of
repetition
of a
series
of
events rather than their mere
recurrence
or
random distribution. Thus,
the
term "cycle"
refers
to a
scries
of
connected events which return
to a
starting
point. When applied
to a
sedimentary sequence, "cycle" refers
to
a
distinctive series
of
lithologies that
are
arranged vertically
in
a
predictable pattern
in
which
at
least
one
rock type,
which
is
regarded
as the
starting point,
is
repeated (Schwar-
zaeher, 1975). Accompanying
the
concept
of
"lithologic
cyclicity"
is the
notion
of
"time cyclicity" (Schwarzaeher,
1975).
The
term "cycle" should refer
to
regularities
in
rock
pattern., whereas
the
term "rhythm" should
be
reserved
for
identical rock types that have been formed
at
equal time
intervals. However, often this distinction
is not
made (Wells,
1960),
or is
impossible
to
judge.
In (his
review 1 will address
both lithologie cycles
and
temporal rhythms,
at the
macro-
scale (i.e., outcrop-scale through seismic-scale),
and
exclude
varve-scale laminations (mm-scale; see Einsele etal.. 1991
for a
summary).
Types
of
cycles: allocycles versus autocycles
Vertical facies successions
and
their lateral distribution
observed
in the
stratigraphic record
are the
results
of
changes
in
the
environment
of
deposition through time,
and
these
ehanges may
be
induced
by
autocyclic
or
allocyclic phenomena
(Figure
C81;
Beerbower,
in
Merriam, 1964). Autocyclic
processes
are
internal
to a
sedimentary system, totally
independent
of
external influences.
To
develop cyclic deposits,
these processes have
to be
sell-regulating and include
a
built-in
feedback mechanism. Excellent examples
of
autocyclic
pro-
eesses which lead
to
clastic cyclic deposits
are: (i)
delta lobe
switching
due to
repeated progradation
and
retreat perhaps
due
to
local compaction
of
pro-delta deposits
or
changes
in
rates
of
sediment supply:
and (ii)
migration
and
superposition
of channel
and
lobe systems
in
fluvial environments. Other
examples include non-periodic storm
bed
deposition
and
deep-
water turbidite deposition.
As
these processes
are
limited
arcally
to the
basin depocenter
itself,
resultant "cycles" will
be
of limited lateral extent. Additionally,
the
time period
of
autocyclic deposits
is
often less than that
of
allocyclic deposits
(Einsele <'f<//., 1991).
Allocyclic processes
are
external
to the
depositional system
and they impart
a net
effect
on
stratigraphic accumulation.
Fundamentally, external variations
are
driven
by
changes
in
climate and/or regional
(or
global) tectonics. Climate
can
affect sea-level changes through
the
advance
and
retreat
of
continental glaciers,
and
obviously influences deposition
of
evaporites
in
closed (lacustrine)
or
semi-enclosed marginal
marine settings. Teetonic movements influence sea-level
changes which
in
turn effect accommodation.space, that
is the
vertical space available
for
sediment accumulation,
as
well
as
water depths which influence facies.
In
contrast
to
autocyclic
processes, allocyclic packages may extend over great distances
within
a
basin,
as
well
as
from
one
basin
to
another,
and
their
time periodicity
is
often greater than that
of
autocycles
(Einsele (•/«/., 1991).
Terminology, cyclic hierarchies, and orders
of
cyclicity
Perhaps
the
largest impediment
to
furthering
our
under-
standing
of
cyclicity
is the
development
and
utilization
of a
174
CYCLIC SEDIMENTATION
(A)
AUTOCYCLIC MECHANISM
SWITCHING DELTA LOBE
DUETO REPEATED PHO-
GRADATION AND RETREAT
DELTA PLAIN
(LAGOONS, MARSH, ETC
1
DELTA SLOPE
SUBSIDENCE
STORM WAVES
MUD SLOPE
TURBIDITES
TEMPESTiTES
SAND
ACCUMULATION
(B)
ALLOCYCLIC MECHANISM
CLIMATIC CHANGE AND TECTONICS
IN SOURCE AREA
CYCLIC
LAKE SEDIMENTS
GLOBAL
SEA LEVEL
FLUCTUATroNS
VARIATIONS
IN
SEDIMENT INPUT
INTO DIFFERENT BASINS
VARIATIONS
IN
ORGANIC
PRODUCTIVITY
TRANSGRESStVE
-
REGRESSIVE
SEQUENCE
CHANGING
WATER
CHEMISTRY
Figure C81 Autocyclk
(A)
and allucyclic
IBI
mec h.inisms (alter
Einsele e^L?/-, 1991).
consistent terminology which satisfies
a
person's perception
of
the nature
of
cyclicity
in the
stratigraphie record.
As
advances
are made
in the
accuracy
of
age-dating., cycle duration will
become better constrained
and
this will certainly assist
in a
more widely adopted terminology
and
classification scheme.
At present, however, absolute
age
dates fall short
of the
required precision,
and
thus geologists continue
to
wrestle with
a consistent terminology. Perhaps
a
more severe problem
is
the
fad
that workers have difTerent criteria regarding
the
recognition
of
cycles
and
particularly
the
surfaces that
separate them.
For example,
the
high-frequeney building block
of
shallow-
marine clastic deposits
is
lenned
the
parasei/ncnce defined
by
Van Wagoner
et al.
(1990)
as "a
relatively conformable
succession
of
genetically related beds
or
bedsets bounded
by
marine-tlooding surfaces
or
their correlative surfaces".
A
marine-tlooding surface
is
defined
as "a
surface separating
younger from older strata across which there
is
evidence
of
an
abrupt increase
in
water depth. This deepening commonly
is
accompanied
by
minor submarine erosion
or
nondeposition
(but
nol by
subaerial erosion), with
a
minor hiatus indicated".
This definition
of ihe
highest frequency building block
in
shallow marine clastic systems, which many workers would
equate with
the
term "cycle", emphasizes
the
physical
bounding surface
to the
point
of
excluding
the
internal
succession
of
rock types
or
facies from
the
definition.
By contrast,
in
shallow iTiarine carbonate systems.
Goldhammer etal.
(in
Franseen
ctaf,
1991; 1993). following
Wilson (1975). James (1984)
and
Hardie
and
Shinn (1986).
define
the
high-frequency building block
as a
cycle which they
define
as:
a relatively conformable succession
of
genetically related
subtidal subfacies bounded
by
peritidal subfacies, subaerial
exposure surfaces, and/or marine flooding surfaces (very thin
intervals characterized
by
slow rates
of
deposition). Cycles
are
the thinnest recognizable allocyelic
or
autoeyclic depositional
unit
and
they may
be
progradational and/or aggradational
and
thus subfacies within cycles shoal upward. Individual cycles
typically contain both
a
"transgressive"
and
"regressive"
component,
and the
relative proportion
of
these components
within
any
given cycle
is a
function
of the
cycle's position
within
a
lower frequency cycle
or
sequence.
Cycles
by
this definitioti
are
highly accurate depictions
of
laterally coeval facies
or
subfacies associations
in
accordance
with Walther's
Law.
Goldhatnmer etal.
(in
Franseen etal.. 1991. 1993) recognize
two types
of
shallow marine platform carbonate cycles, based
on physical bounding surfaces
:
cycles bounded
by
features
indicative
of
subaerial exposure (e.g.. mudcracked peritidal
Iaminites. caliches, etc), tertned "exposure cycles'"
or
"peritidal
cycles",
and
those that
do not
shoal
to
fill level which
are
bounded
by
marine Hooding surfaces across which subfacies
deepen, termed "subtidal cycles"
(see
also Osleger
and
Read,
1991).
Below.
I
briefly review
the
approaches
to a
consistent
terminology
and
classification, attempting
to
provide some
historical context.
Hierarchical subdivision based
on
duration
of the "cycle"
The sedimentary record throughout
all of
geologic history
is
characterized
by
stratigraphic cycles
of
different orders
of
magnitude, both
in
terms
of
thickness, duration
and
regional
extent. Suess (1888) early
on
appreciated
the
notion that
several orders
of
cyclicity existed within
the
stratigraphic
record
and
Barrell introduced
the
term "composite cycle"
to
successions
in
which cycles
of
different wavelengths
are
superimposed upon each other (Schwarzacher. 1975). This
concept
is
now well-established
as
indicated
by
F,xxon's global
cycle charts (Vail
etaf.
1977;
Haq etaf.
1987) constructed
for
the Phanerozoic, which consist
of
three superimposed orders
of cyclicity. Extending Sloss' original idea
of
regional
"unconformity-bounded sequences'" Vail
and his
Exxon
co-
workers proposed
the
existence
in the
stratigraphic record
of
"global cycles
of
relative
sea
level change"
and
they presented
a hierarchy
of
global cycles based
on
their inferred temporal
duration (Figure
C82): (1)
first-order cycles with durations
of
200
300 m year;
(2)
second-order cycles with durations
of 10-80myear;
(3)
third-order cycles with durations
of
l-lOmyear (Vail etal.. 1977).
At
that time,
in the
Exxon
scheme, cycles referred strictly
to
relative sea-level changes
and
the term "sequence"
was
reserved
for the
stratigraphic record