Boggs S. Principles of Sedimentology and Stratigraphy
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Strontium Isotopes
15.4 Chronocorrelation
547
Strontium has four principal isotopes (Table 15.3), of which 87Sr and 86Sr are of
conce here. The relative abundances shown in Table 15.3 are somewhat variable
because 87Sr is the daughter product of the radiogenic isotope 87Rb (Table 15.3), as
previously discussed. The present-day quantity of 87Sr is thus a fraction of e ini
tial amount of 87Sr in Earth plus the amount of radiogenic 87Sr generated from
decay of 87Rb through time (Veizer, 1989).
The amount of strontium in the ocean is derived by three mechanisms and
from three sources: (1) leaching of strontium from basalc rocks owing to hy
drothermal circulation at mid-ocean ridges (see discussion in Chapter 1), (2)
weathering of contental rocks and delivery to the ocean in river water, and (3)
dision of strontium from carbonate sediments caused by recrystallizaon dur
ing burial diagenesis (McArthur, 1998). Because the 87Sr/86Sr ratio in each of these
sources is different, different ratios of 87Sr /86Sr are thus furnished to the ocean.
Strontium in river runoff has the highest 87Sr/86Sr ratio; sontium derived by hy
drothermal circulation has the lowest. Weathering and river-runoff furnish the
greatest quantity of strontium; hydrothermal circulation at ridge crests is second
in importance.
The strontium isotope ratio of ocean water is constant throughout the ocean
at any given me but has varied through time owing to variations in stronum
contributed by these three processes. Va riations in mid-ocean ridge volcanism and
the effects of changing world climates on weathering and river ow have exerted
the most important controls. Strontium is removed from the oceans by coprecipi
tation with calcium in calcium carbonate minerals. erefore, analysis of marine
carbonates of various ages allows the strontium isotope composition of the ocean
through time to be determined.
Figure 15.14 shows variations of the 87Sr /86Sr ratio of Phanerozoic-age ma
re carbonates, reflecting vaations in these isotope ratios in ocean water since
Precambrian time owing to changing proportions of strontium contributed to the
ocean from different sources. Note a general decrease 87Sr /86Sr ratios of the
ocean from Precambrian time until the Jurassic, with several pronounced excur
sions of lowered ratios (e.g., Ordovician, Devonian, Mississippian, Permian/Trias
sic).
A general increase in the ratio since Jurassic time is evident from Figure 15.14.
Strontium isotope data can be reported in terms of deviation from a stan
dard, as in the ce of oxygen, carbon, and sulfur isotopes (McArthur, 1998; Ve iz
er, 1989); however, many research results appear to be reported simply as 87Sr/86Sr
ratios. Isotope ratios can be used to establish correlation by find� horizons with
two different stratigraphic sections that have identical 87Sr I 6Sr ratios (Fig.
15.15). Because more th one point on the sontium-isotope curve (Fig. 15.14)
can have the same 87Sr/86Sr value, independent evidence must be available (e.g.,
fossils, magnetostratigraphic data, regional geologic age relationships) to show
that the two horizons are approximately of equivalent age. Once correlation has
been established, it is possible to determine the numerical age (in millions of
years) of the correlation points from the isotopic age curve shown in Figure 15.14
(Faure, 1982; McArthur, 1994, 1998).
Correlation by strontium isotopes has been applied with particularly good
results to some Te rtiary formations. For example, Depaolo and Finger (1988) use
strontium isotopes to correlate marine sediments of the Miocene Monterey For
mation of Califoia at resolutions comparable to those of biostratigraphic meth
ods. By correlating the 87Sr/86Sr ratios in these sediments with the strontium
isotope vs. time curve for seawater derived from Deep Sea Drilling Program
(DSDP) coholes in the southweste and central Pacific, these authors deter
mined ages with resolutions ranging from <0.1 to 2.5 Ma. See Bralower et al.
548
Chapter 15 / Chronostratigraphy and Geologic Time
2
0
c
300
�
I
4
.7060
.7070
87
Sr/86
s
r
.7080
.7090
.71 00
Figure 15.14
Variations of
87
Sr/
86
Sr of seawater during Phanero
zoic time. [After Veizer, J., 1989, Strontium isotopes
in seawater through time: Annual Review Earth and
Planetary Sciences, v. 17, Fig. 9., p. 157. Reproduced
by permission.]
(1997) for an application to Cretaceous deep-sea sediments recovered during
DSDP and Ocean Drilling Program (ODP) coring.
Problems with Isotopic Chronocorlatlon
e field of stable isotope geochemistry is very complex; many questions remain
with regard to e validity of observed oxygen, carbon, sulfur, and strontium iso
tope excursions in the geologic record and use of these excursions in chronocorre
lation. A discussion of these problems and the details of correlation by stable
isotopes is outside e scope of this book. Additional information on this subject
available in several of the publicatio listed at the end of this chapter.
Further Reading
549
SECTION 1
STRATIGRAPHIC
LEVEL
SECTION 2
Fire 15.15
meters measured
0
from cli top
70
meters
0
80
100
0. 7076 0. 7078
0. 7076 0. 7078
Schematic illustration of cor
relation by
87
Sr/
86
Sr ratios
between two widely sepa
rated hypothetical strati
graphic sections. If the
sections can be established
to have approximately
equivalent ages on the basis
of fossils or other age data,
levels within the stratigraph
ic sections that have identi
ca1
87
Sr/
86
Sr ratios formed
at the same time and there
fore correlate. [After
McArthur, J. M., 1998,
Strontium isotope stratigra
phy, in Doyle, P., and M. R.
Bennett (eds.), Unlocking
the stratigraphical record:
Carbonates Clays
l Sandstone
FURTHER READING
Berggren, A., D. Kent, M-P. Aubry, and J. Hardenbol (eds.),
1995, Geochronology, time scales and global stratigraphic
correlation: SEPM Spec. Pub. 54, Society for Sedimentary Ge
ology, Tu lsa, Okla; 386 p.
Clauer, N., and S. Chaudhuri (eds.), 1992, Isotopic signatures and
sedimentary records: Lecture Notes in Earth Sciences 43,
Springer-Verlag, Berlin, 529 p.
Clifton, H. E. (ed.), 1988, Sedimentological consequences of con
vulsive geologic events: Geol. Soc. America Spec. Paper 229,
157 p.
Dickin, A. P., 1995, Radiogenic isotope geology: Cambridge Uni
versity Press, Cambridge, 452 p.
Faure, G., 1986, Principles of isotope geology, 2nd ed.: John Wiley
& Sons, New Yo rk, 589 p.
Geyh, M. A., and H. Schleicher, 1990, Absolute age determina
tion: Springer-Verlag, Berlin, 503 p.
Harland, W. B., R. L. Armstrong, A. Cox, L. E. Craig, A. G.
Smith, and D. G. Smith, 1990, A geologic time scale 1989, 2nd
ed.: Cambridge University Press, Cambridge, 263 p.
Hoefs, J., 1997, Sble isotope geochemistry, 4th ed.: Springer-Ver
lag, Berlin, 201 p.
"sr'"sr
0. Fossils
Advances in modern stratig
raphy: John Wiley & Sons
Ltd., Fig. 8.3, p. 228.]
Odin, G. S. (ed.), 1982, Numerical dang in stratigraphy: John
Wiley & Sons, ew York, Parts I and II, 1040 p.
Shiki, , S. K. Chough, and G. Einsele, 1996, Marine sedimentary
events and their records: Sedimentary Geology, v. 104, 255 p.
[Special Issue]
Snelling, N. J. (ed.), 1985, The chronology of the geological
record: Geol. Soc. Mem. 10, Blackwell, Oxford, 343 p.
Valley, J. W., and D. R. Cole (eds.), 2001, Stable isotope geochem
istry: Reviews in mineralogy and geochemistry, v. 43, Miner
alogical Society of America, 662 p.
Vertes, A., S. Nagy, and K. Silvegh (eds.), 1998, Nuclear methods
in mineralogy and geology: Techniques and applations:
Plenum Press, New York, 555 p.
Wa lliser, 0. H. (ed.), 1996, Global events and global event stratig
raphy in the Phanerozoic: Sprger-Verlag, Berlin, 333 p.
Williams, D. , I. Lerche, and W. E. Full, 1988, Isotope chronos
tratigraphy: Theory and methods: Academic Press, San
Diego, 345 p.
Basin Analysis I Tectonics I
and Sedimentation
550
16.1 INTRODUCTION
P
receding chapters of this book oduce the fundamental prciples of sedi
mentology and stratigraphy. I have reviewed those chapters the basic
processes at generate sedimentary rocks; discussed the physical, chemical,
and biological properties of these rocks; and described the various environments
in which they accumulate. Further, I have emphasized those properties of sedi
mentary rocks that can be used to subdivide strata into meaningful stratigraphic
units on the basis of liology, seismic reflection characteristics, magnetic polarity,
fossil content, and age.
Geologists study sedimentary rocks to develop a critical understanding of
their geologic history or to evaluate their economic potential. Effective study re
quires utilization of all the sedimentological and stratigraphic prciples dis
cussed in this book. Because most sedimentary rocks were deposited in basins, it
is common to refer to such detailed, integrated study as basin analysis. e con
cept of basin analysis, which goes back to the 1940s, is attributed to one of geolo
gy's most famous sedimentologists, Francis Pettijohn (Kleinspehn and Paola,
1988). Basin analysis has been defined in somewhat different ways (e.g., Potter
and Pettijohn, 1977, p. 1; Allen and Allen, 1990; Miall, 2000, p. 3; Klein, 1987); how
ever, the common theme that emerges from all of these points of view is that basin
analysis is an integrated program of study that involves application of sedimento
logic, stratigraphic, and tectonic principles to develop a full understanding of the
rocks at fill sedimentary basins for e purpose of interpretg their geologic
history and evaluating their economic importance.
During the earlier years of geologic study from about 1860--1960, most geol
ogists regarded marine basins as mainly linear troughs, called geosynclines, in
which great thicknesses of predominantly shallow-marine deposits accumulated
owing to continued subsidence of the geosynclines (see review by Dott, 1974).
With the emergence of the plate tectonics concept in the late 1950s and early 1960s,
geological thinking shifted away from geosynclines. Today, geologists recognize
that there are many kinds of basins and many dierent mechanisms by which
basins form. Under the general rubric of basin analysis, geologists are concerned
about e global tectonic controls that create basins and the geologic controls (e.g.,
sea-level changes, sediment supply, basin subsidence) that govern basin filling.
Some of these factors are discussed in preceding chapters.
16.2 Mechanisms of Basin Formation (Subsidence)
551
this final chapter, attention is focused particularly on the different kds of
basins that we now recognize, the processes that form these basins, the processes
that bring about filling of basins, and the nature of the fills. Brief discussions of e
methods used to carry out basin analysis and the applations of basin analysis are
also included. 1his chapter rounds out the text coverage by providing an overall
perspective regarding the integrated nature of sedimentologic, environmental,
and stratigraphic synthesis.
16.2 MECHANISMS OF BASIN FORMATION (SUBSIDENCE)
A sedimentary basin is a depression of some kind capable of trapping sediment.
Subsidence of the upper surface of the crust must take place to form such a de
pression. Mechanisms that can generate sucient subsidence to create basins in
clude crustal thinning, mantle-lithosphere thickening, sedimentary and volcanic
loading, tectonic loading, subcrustalloading, asthenopheric flow, and crustal den
sification (Dickson, 1993). A brief explanation of each of these mechanisms is
given in Table 16.1.
Note in Table 16.1 that isostatic compensation is an important aspect of sed
imentary and volcanic loading. The concept of isostasy assumes that local com
pensation of the crust occurs as if Earth consists of a series of free-floating pistons.
Adjacent blocks of crust of different thickness and I or density structure will have
different relative relief (Angevine, Heller, and Paola, 1990). Thus, adding a load to
the crust (e.g., filling a basin with sediment) causes subsidence; removg a load
(e.g., erosion of the crust) causes uplift. It follows from this premise that a basin
originally filled with water will be deepened by the sediment load as the basin
gradually accumulates sediment. addition to the effects of loading, flexing of
the crust also occurs, to various degrees depending upon the rigidity of the un
derlying lithosphere, as a result of tectonic forces: overrusting, underpulling,
underthrusng of dense lithosphere. Finally, thermal effects (e.g., cooling of lhos
phere, increase in crustal density caused by changing temperatu/pressure con
ditions) may also be important basin formation.
Crustal thinning:
Mantle-lithospheric
thickening:
Sedimentary and
volcanic loading:
Tectonic loading:
Subcrustal loading:
Asthenospheric ow:
Crustal densification:
Extensional stretching, erosion during uplift, and magmatic
withdrawal
Cooling of lithosphere following either cessation of stretching
or heating due to adiabatic melting or rise of asthenospheric
melts
Local isostatic compensation of crust and regional lithospheric
flexure, dependent on flexural rigidity of lithosphere, during
sedimentation and volcanism
Local isostatic compensation of crust and regional lithospheric
flexure, dependent on flexural rigidity of underlying
lithosphere, during overthrusting and/ or underpulling
Lithospheric flexure during underthrusting of dense
lithosphere
Dynamic effects of asthenospheric flow, commonly due to
descent or delamination of subducted lithosphere
Increased density of crust due to changing pressure/
temperature conditions and/ or emplacement of
higher-density melts into lower-density crust
urce: Dickinson, 1993; In
g
soll and Busb
y
, 1995.
552
Chapter 16 / Basin Analysis, Te ctonics, and Sedimentation
Millions of Yea Before Present
140 120 100 80 60 40 20 0
0
E
2
�
Figure 16.1
Example of a geohistory diagram showing, in this case, burial history
of a stratigraphic horizon in COST B- 2 Well, Atlantic passive margin
of the eastern United States. Note that subsidence owing to tecton
ism is distinguished from subsidence caused by sediment and water
loading. [After Watts, A. B., 1981 , The U.S. Atlantic continental mar
gin: Subsidence history, crustal structure and thermal evolution, in
Bally, A. W., et al., Geology of passive continental margins: MPG Ed
ucation Course Note Series 19, Fig. 14, p. 2-33.]
0
3
4
5
1
c
m
i1
0
0
0
yr
-
-----
0� ',
"
Total sediment
o
accumulation
Geologists, particularly petroleum geologists searching for oil and gas, are
commonly interested in developing an understanding of the history of basin sub
sidence to determine the relative importance of the various basin-subsiding mech
anisms. Subsidence history is often studied by means of backstripping, which
attempts to remove the effects of sedimentation in order to analyze the manner in
which a basin would have subsided if sediments had not been deposited (e.g.,
Slingerland, Harbaugh, and Furlong, 1994, p. 12; Wa tts, 1981). The results of back
stripping are generally displayed in geohistory diagrams, such as that shown in
Figure 16.1. These diagrams allow the eects of sediment loading to be distin
guished from those owing to tectonic subsidence. Data required to carry out geo
history analysis include a stratigraphic column showing present-day thickness of
stratigraphic units, types of lithologies, ages of stratigraphic horizons, estimated
water depths at the sedimentation site at the time of deposition (paleodepth), and
sediment porosity (Angevine, Heller, and Paola, 1990). In constructing geohistory
diagrams, a correction must be made for sediment compaction. The relative im
portance of the basin subsidence mechanisms listed in Table 16.1 to the formation
of different kinds of sedimentary basins (discussed in the succeeding section) is
indicated schematically in Figure 16.2.
16.3 PLATE TECTONICS AND BASINS
I have made numerous references throughout this book regarding the eects of
tectonics on sedimentation pattes and stratigraphic characteristics. Plate tecton
ics provides a rst-order control on sedimentation through its influence on e
sediment source area. e kinds of basins in which sediments accumulate are also
directly related to tectonic processes (e.g., Frostick and Steel, 1993; Ingersoll and
Busby, 1995; Miall, 2000, Chapter 7). For example, some basins form as a result of
tectonic processes that produce faulting; others are sag basins created by crustal
cooling and subsidence or other tectonic processes. any case, tectonic forces
control the shape, and location of the basins. Tectonic processes, together
with sediment loading, further determine the rate of basin subsidence and thus
the space available (accommodation) for sediment accumulation.
Plate-tectonic processes bring about major changes in continental masses
and ocean basins through time. Continents break up and drift apart to create
Dominant Important
Minor
I
Basin T es
Terrestrial Rift Valleys
Proto-Oceanic Rift Troughs
Continental Rises and Terraces
Continental
E
mbankments
� Intracratonic Basins
. Continental Platforms
�
z
Active Ocean Basins
Oceanic Islands, Aseismic Ridges/Plateaus
Dormant Ocean Basins
Trenches
Trench
-
Slope Basins
Forearc Basins
� Intra-arc Basins
Backarc Basins
� Retroarc Foreland Basins
z
8 Remnant Ocean Basins
Peripheral Foreland Basins
Piggyback Basins
Foreland Intermontane Basins
16.3
Plate Tectonics and Basins
553
Subsidence Mechanisms
·
l
0
0
:
�
l
l
:
:
�.�
':
O :
c
·;
0
:
O
Ql
: u
i
0
0
·-
·
E
:
i
·-
0
Ql
i
0
�
� !
__ _l
Figure 16.2
Transtensional Basins
Transpressional Basins
Transrotational Basins
Intracontinental Wrench Basins
0
Aulacogens
� lmpactogens
I
Successor Basins
I 1
Suggested subsidence mecha
nisms for all kinds of sedimentary
basins. [From Ingersoll, R. V., and
C. J. Busby, 1995, Tectonics of
sedimentary basins, in Busby, C.
)., and R. V. Ingersoll (eds.), Te c
tonics of sedimentary basins:
Blackwell Science, Cambridge,
Mass., Fig. 1.1, p. 8.]
ocean basins as much as 500 wide, which can subsequently close again as
ocean-floor crust is subducted trenches. The opening and closing of an ocean
basin is referred to as a Wilson c
y
cle (after Wilson, 1966). Wilson cycles begin with
the
formation of rift basins (floored by continental crust), which subsequently
evolve into proto-oceanic troughs (parally oored by oceanic crust), and eventu
ally into ocean bass, floored by oceanic crust and bordered by passive continen
tal margins. After tens of millions of years or more, subduction zones develop
around the ocean margins and the ocean begins to close. Closure culminates with
continental collision and the formation of an orogenic belt. The entire process of
basin formation and destction requires perhaps 50 to 150 million years. The ge
ologic record suggests that there have been many Wilson cycles in the history of
each continent. Thus, few sedimentary basins remain unchanged with time, or in
fixed positions, except perhaps some basins located on cratons well within conti
nental margins.
554 Chapter 16 I Basin Analysis, Te ctonics, and Sedimentation
During the opening phases of a Wilson cycle, tectonic plates are moving apart
(by rifting) to form divergent (passive) continental margins. The closing stages of
a Wilson cycle are characterized by plates moving toward each other, as oceanic
crust is subducted (consumed) in trenches. Continental margins formed during a
stage of ocean closing are called convergent (active) margins. During opening or
closing of an ocean basin, some parts of plates may slide past each other without ei
ther diverging or converging. Such a setting is referred to as a transform margin.
During a Wilson cycle, various kinds of sedimentary basins form in divergent, con
vergent, and transform settings, as well as in intraplate settings.
16.4 KINDS OF SEDIMENTARY BASINS
We now recognize that the origin of sedimentary basins is related in some way to
crustal movements and plate-tectonics processes. Several tectonic classifications
for basins have been proposed (e.g., Dickinson, 1974; Bally and Snelson, 1980;
Kingston, Dishroon, and Williams, 1983; Mitchell and Reading, 1986; Klein, 1987;
Ingersoll, 1988; Ingersoll and Busby, 1995). Ingersoll and Busby (1995) point out
that sedimentary basins can form in the four types of tectonic settings discussed
above (divergent, intraplate, convergent, transform) as well as in some hybrid set
tings (Table 16.2). Different kinds of basins are recognized within these various
settings on the basis of (1) the type of crust on which the basins rest, (2) the posi
tion of the basins with respect to plate margins, and (3) for basins lying close to a
plate margin, the type of plate interactions occurring during sedimentation (e.g.,
Dickinson, 1974; Miall, 2000, p. 468). Detailed discussion of all of these kinds of
:'
�.1o.� ·. · ��jor.��:�:����f
�
�s�:
d:��.t�·to�e ·���
Divergent Settings
Intraplate Settings
Te rrestrial rift valleys:
Proto-oceanic rift troughs:
Continental rises and terraces:
Continental embankments:
Intracratonic basins:
Continental platforms:
Active ocean basins:
Oceanic islands, aseismic
ridges and plateaus:
Rifts within continental crust commonly associated with
bimodal volcanism. Modern example: Rio Grand Rift
(New Mexico).
Incipient oceanic basins floored by new oceanic crust
and flanked by young rifted continental margins.
Modern example: Red Sea.
Mature rifted continental margins in intraplate settings
at continental-oceanic interfaces. Mode example: East
coast of USA.
Progradational sediment wedges constructed off edges
of rifted continental margins. Mode example:
Mississippi Gulf Coast.
Broad cratonic basins floored by fossil rifts in axial
zones. Mode example: Chad Basin (Africa).
Stable cratons covered with thin and laterally extensive
sedimentary strata. Mode example: Barents Sea (Asia).
Basins floored by oceanic crust formed at divergent plate
boundaries unrelated to arc-trench systems (spreading
still active). Mode example: Pacific Ocean.
Sedimentary aprons and platforms formed in
intraoceanic settings other than magmatic arcs. Modem
example: Emperor-Hawaii seamounts.
Convergent Settings
Transform Settings
Hybrid Settings
Dormant ocean basins:
Trenches:
Trench-slope basins:
Fore-arc basi:
Intra-arc bins
Back-arc basins:
Retro-arc foreland basins:
Remnant ocean basins:
Peripheral foreland basins:
Piggyback basins:
Foreland intermontane basins
(broken forelands):
Tr anstensional basins:
Transpressional basins:
Transrotational basins:
Intracontinental wrench basins:
Aulacogens:
16.4 Kinds of Sedimenta Basins
555
Basins floored by oceanic crust, which is neither
spreading nor subducting (no active plate boundaries
within or adjoining basin). Modem example: Gulf of
Mexico.
Deep troughs formed by subduction of oceanic
lithosphere. Modem example: Chile Trench.
Local structural depressions developed on subduction
complexes. Modem example: Central America Trench.
Basins within arc-trench gaps. Modern example:
Sumatra.
Basins along arc platform, which includes superposed
and overlapping volcanoes. Modern example: Lago de
Nicaragua.
Oceanic basins behind intraoceanic magmatic arcs
(including interarc basins between active and remnant
arcs), and continental basins behind continental-margin
magmatic arcs without foreland fold-thrust belts.
Mode example: Marianas.
Foreland basins on continental sides of
continental-margin arc-trench systems (formed by
subduction-generated compression and/ or collision).
Modern example: Andes foothills.
Shrinking ocean basins caught between colliding
continental margins and/ or arc-trench systems, and
ultimately subducted or deformed within suture belts.
Modern example: Bay of Bengal.
Foreland basins above rifted continental margins that
have been pulled into subduction zones during crustal
collisions (primary type of collision-related forelands).
Modern example: Persian Gulf.
Basins formed and carried atop moving thrust sheets.
Modern example: Peshawar Basin (Pakistan).
Basins formed among basement-cored uplifts in foreland
settings. Mode example: Sierras Pampeanas basins
(Argentina).
Basins formed by extension along strike-slip fault
systems. Mode example: Salton Sea (California).
Basins formed by compression along strike-slip fault
systems. Mode example: Santa Barbara Basin
(California) (foland).
Basins formed by rotation of crustal blocks about
vertical
axes within strike-slip fault systems. Mode
example: We ste Aleutian fore-arc (?).
Diverse basins formed within and on continental crust
owing to distant collisional processes. Mode example:
Quaidam Basin (China).
Former failed rifts at high angles to continental margins,
which have been reactivated during convergent
tectonics, so that they are at high angles to orogenic
belts. Mode example: Mississippi Embayment.
556
Chapter 16 I Basin Analysis, Te ctonics, and Sedimentation
Impactogens:
Rifts formed at high angles to orogenic belts, without
preorogenic history (in contrast with aulacogens).
Modern example: Baikal Rift (Siberia) (distal).
Successor basins:
Basins formed in intermontane
following
cessation of local orogenic or taphrogenic activity.
Mode example: Souern Basin and Range (Arizona).
Basin classification-modified after Dickinson, 197 4, 1976, and Ingersoll, 1988. Source: Ingersoll, R. , and C. J. Busby, 1995 Tectonics of sedimentary
basins, in Busb
y
, C. J., and R. V. Ingersoll (eds.), Tec tonics of sedimtary basins: Blackwell Science, Table 1.1, p. 3, Table 1.2, p. 5. Reproduced by permission.
Figure 16.3
Schematic representation of
selected kinds of tectonical
ly formed basins. [After
Dickinson and Yarborough,
1976; Kingston, Dishroon,
and Williams, 1983;
Mitchell and Reading,
1986; Einsele, 1992; Inger
soll and Busby, 1995.]
basins is beyond the scope of this book; however, a few of the more interesting and
important types of bass are illustrated schematically Figu 16.3 and furer
discussed below. Interested readers should consult Busby and Ingersoll (1995),
Einsele (2000), and Miall (2000) for additional insight.
A. Rift graben
Continental crust
C. Intracratonic basin
Spreading
Remnant
center
arc
Forearc
Volcanic island arc
E. Subduction-related basin
G. Tra nsform-fault-related basin
crust
crust
B. Passive-margin basins (continental
rises and terraces)
D. Basins floored by oceanic crust
Peripheral foreland
.
basin
vedi
ng plate
F. Collision-related basins
Oceanic crust
H. Aulacogen