-
Geological Society of America Bulletin
doi: 10.1130/B30754.1 published online 21 December
2012;Geological Society of America Bulletin
Donna L. Whitney, Christian Teyssier, Patrice Rey and W. Roger
Buck
Continental and oceanic core complexes
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Copyright 2012 Geological Society of America
as doi:10.1130/B30754.1Geological Society of America Bulletin,
published online on 21 December 2012
-
Continental and oceanic core complexes
Donna L. Whitney1,, Christian Teyssier1, Patrice Rey2, and W.
Roger Buck31Department of Earth Sciences, University of Minnesota,
Minneapolis, Minnesota 55455, USA2School of Geosciences, University
of Sydney, Sydney NSW 2006, Australia3Lamont-Doherty Earth
Observatory, Columbia University, Palisades, New York 10964,
USA
ABSTRACT
Core-complex formation driven by litho-spheric extension is a fi
rst-order process of heat and mass transfer in the Earth.
Core-complex structures have been recognized in the continents, at
slow- and ultraslow-spreading mid-ocean ridges, and at continen-tal
rifted margins; in each of these settings, extension has driven the
exhumation of deep crust and/or upper mantle. The style of
ex-tension and the magnitude of core-complex exhumation are
determined fundamentally by rheology: (1) Coupling between brittle
and ductile layers regulates fault patterns in the brittle layer;
and (2) viscosity of the fl owing layer is controlled dominantly by
the synextension geotherm and the presence or absence of melt. The
pressure-tempera-ture-time-fl uid-deformation history of core
complexes, investigated via fi eld- and mod-eling-based studies,
reveals the magnitude, rate, and mechanisms of advection of heat
and material from deep to shallow levels, as well as the
consequences for the chemical and physical evolution of the
lithosphere, includ-ing the role of core-complex development in
crustal differentiation, global element cycles, and ore formation.
In this review, we provide a survey of ~40 yr of core-complex
literature, discuss processes and questions relevant to the
formation and evolution of core com-plexes in continental and
oceanic settings, highlight the signifi cance of core complexes for
lithosphere dynamics, and propose a few possible directions for
future research.
INTRODUCTION
When the lithosphere is under extension, the brittle upper crust
breaks and is displaced along normal faults. When extension is
concentrated on one or a few faults in a narrow region, ductile
material ascends from deeper levels of the litho-sphere, resulting
in exhumation of deep crustal
rocks upper mantle in the footwall of the nor-mal fault(s). The
resulting structure is a core complex, which occurs in both
continental and oceanic lithosphere (Figs. 1 and 2). Extension is
the direct driving force for core-complex devel-opment, but in
continental settings, the far-fi eld tectonic regime may be one of
convergence, and therefore continental core complexes may occur in
orogenic settings under an overall regime of plate convergence.
As extension proceeds, heat and material are transferred from
deep (hot, ductile) to shallow (cool, brittle) levels, driving
vigorous fl uid fl ow and strongly infl uencing the location and
mag-nitude of subsequent extension. Interactions among minerals, fl
uids, and/or magma may pro-duce economically important mineral
deposits, and young extensional zones may be sources of
hydrothermal activity during and after active faulting.
Core complexes were fi rst recognized in the continents (e.g.,
Anderson, 1972; Coney, 1974, 1980; Crittenden et al., 1980; Lister
and Davis, 1989), and they have been identifi ed in the geo-logic
record from the Precambrian (Holm, 1996) to the present (Hill et
al., 1992). Core complexes were later identifi ed at slow- and
ultraslow-spreading oceanic divergent zones (e.g., Cannat, 1993;
Cann et al., 1997; Blackman et al., 1998; Tucholke et al., 1998;
Karson, 1999; Ranero and Reston, 1999; Dick et al., 2000).
Continen-tal and oceanic core complexes have similar di-mensions,
fault geometry, and kinematics (Figs. 1 and 3), and both involve
exhumation of deeper levels of the lithosphere to shallow levels
(John and Cheadle, 2010).
In this review, we discuss the origin and significance of
continental core complexes and oceanic core complexes. Although the
term metamorphic core complex is a common
For permission to copy, contact [email protected] 2013
Geological Society of America
1
GSA Bulletin; Month/Month 2013; v. 1xx; no. X/X; p. 126; doi:
10.1130/B30754.1; 13 fi gures.
E-mail: [email protected]
oceanic detachment system
gabbro
sea water
fluidflow
brittle-ductiletransition: ~600C(olivine rheology)
ridgeaxis
Oceanic core complex
~10 km
fluidflow
13
meteoric water
fluids derived from crystallization of metamorphic/igneous
rocks
brittle layer
ductilelayer
mylonitezone
meteoric water
brittle-ductiletransition: ~300400C
(quartz rheology)
continental detachment system
~15 km
melt
domi-nantlybrittle layer
ductilelayer
mylonite
sea waterserpentinite
13
A
B
fluidflow ?
Option 1Option 2 ?
?
??
shearedgabbro
Figure 1. Continental (A) and oceanic (B) core complexes dif-fer
in their primary rock types (granitic and metasedimentary rocks in
continental core com-plexes vs. gabbro and serpenti-nite in oceanic
core complexes), and therefore in the minerals that infl uence the
rheology of the complexes. Nevertheless, many fi rst-order
processes of their origin and evolution are similar, and therefore
there are many similarities in their archi-tecture. In B, the
detachment fault roots in gabbro magma at depth (option 1); option
2 con-siders a dry spreading center in which the brittle detachment
transitions to a ductile shear zone at depth (lithosphere
boudinage).
Invited Review
CELEBRATING ADVANCES IN GEOSCIENCE
1888 2013
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published online on 21 December 2012
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Whitney et al.
2 Geological Society of America Bulletin, Month/Month 2012
description of core complexes on the conti-nents, for the sake
of simplicity (and symme-try), in this review we use the terms
continental core complex (equivalent to metamorphic core complex)
and oceanic core complex. We do not discuss in any detail the
formation of continental margin core complexes, although some
locations are highlighted on a world map (Fig. 2A).
Over the past ~40 yr, interest in core com-plexes has remained
high because these struc-tures are common in extending orogens and
along slow-spreading oceanic divergent zones, and because they
record fundamental thermo-mechanical processes in extending
lithosphere. An understanding of the uplift and exhuma-tion of
ductile rocks below low-angle normal faults, as well as the
dynamics of the faults, is
relevant to crustal evolution and seismogenesis in extending
lithosphere, and core complexes have therefore been intensively
studied using a variety of techniques, e.g., fi eld-based stud-ies
(e.g., Davis and Coney, 1979; Miller et al., 1983; Bozkurt and
Park, 1994; Gessner et al., 2001), numerical modeling (e.g., Buck
et al., 1988; Lavier et al., 1999; Tirel et al., 2004, 2008; Rey et
al., 2009a, 2009b; Allken et al., 2011), and analog modeling (e.g.,
Brun et al., 1994; Tirel et al., 2006).
There remain important questions about core-complex initiation
and evolution. In this review, we integrate knowledge derived from
different types of investigations (fi eld, mod-eling) of
continental and oceanic core com-plexes and discuss some of these
unresolved issues.
DEFINITIONS
Herein, we use the following general defi ni-tion of a core
complex and the processes that drive core-complex formation:
A core complex is a domal or arched geologic structure composed
of ductilely deformed rocks and associated intrusions underlying a
ductile-to-brittle high-strain zone that experienced tens of
kilometers of normal-sense displacement in response to lithospheric
extension.
The lithospheric extension that results in core-complex
formation is commonly driven by plate divergence, such as at
mid-ocean ridges and along rifted continental margins. Extension
also occurs in plate convergence settings by slab rollback (e.g.,
the backarc of an oceanic subduc-tion zone) or by orogenic collapse
under fi xed
HhLi
HoDI
Da, Nb
Pa
Sh
Ma
Lo
DNCV
Xi
HaYOH
La SC
PoAA
SELR Ed
VeTo
SB
GK ChNi KS
ADGM
Re
LfPL
+ Antarctica: Fosdick core complex, Marie Byrd Land
N. Am
erican Cordillera
Aegean
Mid-AtlanticRidge
Ka
At
AB
SouthwestIndianRidge
Ba
No
Go
Australian- Antarctic Discordance
Nx
Rh
AMOR
M-Ca
BB
MC-Pyr
continental core complexoceanic core complexcontinental margin
core complex
Fig. 2B Fig. 2C
A
60N
0
60S
0120W 120E
60N
0
60S
0 120E
Figure 2 (on this and following page). (A) Map of the world
showing the locations of some Phanerozoic core complexes in the
continents and oceans. Key to abbreviations: AAAlpi Apuane (Italy);
ABAtlantis Bank (SW Indian Ridge); ADAma Drime (Nepal); AMORArctic
segment of Mid-Atlantic Ridge; AtAtlantis Massif (Mid-Atlantic
Ridge); BaBaja (Mexico); BBBay of Biscay; ChChapedony (Iran);
DaDayman (Papua New Guinea); DIDoi Inthanon (Thailand); DNCVDay Nui
Con Voi (Vietnam); EdEdough (Algeria); GKGrand Kabilye (Algeria);
GMGurla Mandhata (Pamirs); GoGodzilla; HaHarkin (China/Mongolia);
HhHohhot (China); HoHongzhen (China); KaKane (Mid-Atlantic Ridge);
KSKongur Shan (Pamirs); LaLaojunshan (China); LfLofoten (Norway);
LiLiaodong Peninsula (China); MaMalino (Indonesia); LoLouzidian
(China); LRLora del Rio (Spain); M-CaMid-Cayman spreading center;
MC-PyrMassif Central (FrancePyrenees, France, Spain; includes
Montagne-Noire); NbNormanby Island (Papua New Guinea); NiNide
(Turkey); NoNorway rifted continental margin; NxNaxos (Greece);
PaPaparoa (New Zealand); PLPayer Land (Greenland); PoPohorje
Mountains (Slovenia); ReRechnitz (Austria); RhRhodope (Greece,
Bulgaria); SBsouthern Brittany (France); SCSong Chay (China);
ShShaerdelan (China); SESierra de las Estancias (Spain); ToTormes
(Spain); VeVeporic (Slovenia); XiXiaoqinling (China);
YOHYagan-Onch-Hayrhan (China/Mongolia).
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published online on 21 December 2012
-
Continental and oceanic core complexes
Geological Society of America Bulletin, Month/Month 2012 3
boundary conditions or even during slow plate convergence (Rey
et al., 2001). Core complexes occur in all of these settings.
In continental core complexes, the normal-sense high-strain zone
corresponds to a pro-found metamorphic and/or stratigraphic
dis-continuity typically called a detachment fault (Fig. 1), which
is so named because rocks above and below the fault zone record
different pressure-temperature-time-deformation his to-ries.
Although the uppermost part of the fault zone may be a prominent
brittle fault surface, below this fault there may be a broad
(hundreds of meters) zone in which ductilely deformed rocks have
been kinematically and thermally linkedfor some or all of their
defor ma tion historyto the structurally higher, brittle fault
during exhumation of the footwall. Some de-tachment faults
record a progression from high-temperature (mylonitic) deformation
to much lower-temperature brittle deformation through time.
The concept of detachment has also been applied to normal faults
that are not associated directly with core-complex structures,
e.g., the South Tibetan detachment system (e.g., Burg et al.,
1984), which is the northern boundary of the Himalayan crystalline
wedge, and normal faults in exhumed subduction complexes, such as
on the island of Crete (e.g., Ring et al., 2001). There are also
detachment-like faults in some conti-nental arcs, such as the Andes
(e.g., Mpodozis and Allmendinger, 1993) and the North Cas-cades
(Paterson et al., 2004). These detachments
are normal faults with signifi cant displacement, but they do
not bound core complexes.
Detachment zone, detachment shear zone, and detachment system
are terms that have been used to describe the diffuse zone of
strain, up to 1.5 km thick, that underlies the upper most fault
plane in some core complexes, such as those of the northern U.S.
and Canadian Cordillera (e.g., Mulch et al., 2006; Gbelin et al.,
2011). Characteristic features of detach-ments and ideas and
debates about the dynamics of low-angle normal faults are discussed
in later sections.
Although detachment is widely used to describe core
complexbounding faults, other terms are also used: for example,
low-angle normal fault (e.g., Axen, 2007). The term
Figure 2 (continued ). (B) Schematic map of the core complexes
in the North American Cordillera. (C) Schematic map of the core
com-plexes in the Aegean Sea and surrounding regions.
zmirAthens
Istanbul
Rhodes
Crete
NaxosParos
Samos
Gulf of Corinth
Syros
Tinos
Thasos
IosMilos
Kea
Andros
Sifnos
Kazda Massif
24 26 28
40N
Menderes Massif
Rhodope Massif
Aegean Sea
GREECE
TURKEY
C
ChemehueviWhipple Mts
Buckskin-RawhideHarcuvarHarquahala Mts
Catalina
Sierra Mazatan
South Mts
Rincon
Picacho
SnakeRange
Ruby-Humboldt
Albion-Grouse Creek
SNAKE RIVER PLAIN
COLUMBIA PLATEAU
COLORADO PLATEAU
SIERRANEVADABATHOLITH
COAST RANGES-NORTH CASCADES
IDAHOBATHOLITH
PENINSULARRANGES
BATHOLITH
PacificOcean
MEXICOUSA
USACANADA
Priest River
Kettle
OkanoganValhalla
Thor-Odin
Frenchmans Cap
San Andreas Fault
Pioneer
AnacondaBitterroot
Lewis & Clark Clearwater
Shuswap Complex
NO
RTHERN
BELTCEN
TRAL BELT
SOU
THERN
BELT
RaftRiver
ROCKY M
OU
NTA
I N FO
LD &
THR U
ST BELT
B
110 W
120W
120W
110W
50N
40N
30N
40N
30N
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Whitney et al.
4 Geological Society of America Bulletin, Month/Month 2012
dcollement has also been used (Coney, 1980; Davis et al., 1980)
and can be synonymous with detachment in the context of a core
complex, in some cases indicating a low-angle master fault or shear
zone. Breakaway fault refers to the fault identifi ed as the
detachment that initially intersected Earths surface.
Some detachment faults are listric; this in-dicates a curving
fault with a dip that changes from steep at shallow crustal levels
to sub hori-zontal at depth. In contrast, some detachment faults
(or parts of detachment faults) in both continental and oceanic
settings are downward concave and accommodate large offset yet
moderate topography (e.g., van den Driessche and Brun, 1992; Lavier
and Manatschal, 2006).
DEVELOPMENT AND EVOLUTION OF THE CONCEPT
As a result of debate, primarily in the 1960s1970s, about the
tectonic history of basement uplifts (exposures of middle- and
lower-crustal rocks) bounded by low-angle faults in the North
American Cordillera (Figs. 2A and 2B), a model of core complexes as
extensional structures was developed (Anderson, 1972; Coney, 1974;
Wright et al., 1974; Proffett, 1977; Davis and Coney, 1979; Coney
and Harms, 1984). The concept has been applied to similar
structures elsewhere, and some of these are described
as Cordilleran-style metamorphic core com-plexes (e.g., Aegean
[Fig. 2C]Lister et al., 1984; West AntarcticaRichard et al., 1994;
Norwegian CaledonidesSteltenpohl et al., 2004; IranVerdel et al.,
2007).
The existence of detachment faults and core complexlike
structures was proposed for slow-spreading mid-ocean ridges before
such struc-tures were observed in detail (Karson and Dick, 1983;
Cannat, 1993; Tucholke and Lin, 1994). Fault-bounded domal
structures resembling continental core complexes were later
identi-fi ed in seafl oor images of the Mid-Atlantic Ridge (Cann et
al., 1997; Blackman et al., 1998; Tucholke et al., 1998; Ranero and
Reston, 1999; Tucholke et al., 2001) and confi rmed by ob-servation
of samples collected from low-angle normal fault zones bounding
these structures (MacLeod et al., 2002; Escartn et al., 2003;
Schroeder and John, 2004; Ildefonse et al., 2007). Oceanic core
complexes have now been recognized along segments of the Southwest
Indian Ridge (Dick et al., 2000; Baines et al., 2003) and at other
divergent zones (e.g., the CaribbeanNorth American RidgeHayman et
al., 2011; the Australian-Antarctic Discor-danceChristie et al.,
1998) (Fig. 2A).
Many models for core-complex develop-ment have invoked isostatic
rebound beneath the detachment fault to explain the arching of the
fault and exhumation of the footwall (e.g.,
in the continents: Spencer, 1984; Wernicke and Axen, 1988; Brun
and van den Driessche, 1994; in the oceans: Lavier and Manatschal,
2006; MacLeod et al., 2009). Crustal fl ow beneath the extending
upper crust has also been proposed to explain the domal structure
of many conti-nental core complexes, as well as the moderate
topographic relief of the exhumed footwall de-spite tens of
kilometers of offset and the exis-tence of a fl at Moho in many
highly extended regions (e.g., Block and Royden, 1990; Buck, 1991;
McKenzie et al., 2000). Exploration of the thermal and mechanical
consequences of large-magnitude crustal fl ow from deep to shallow
crustal levels, including consideration of the re-lationship
between continental core complexes and gneiss domes (Teyssier and
Whitney, 2002), shows that core complexes can be signifi cant sites
for heat and mass transfer and have played a role in
differentiation of continental crust through geologic time (e.g.,
Rey et al., 2009a, 2009b; Thbaud and Rey, 2012).
There has long been debate about the dynam-ics of low-angle
normal faults, primarily fo-cused on the question of whether faults
of low (
-
Continental and oceanic core complexes
Geological Society of America Bulletin, Month/Month 2012 5
1981). However, weak zones (faults) could re-sult from fl
uid-rock interaction that increases pore pressure or generates
phyllosilicates and other weak minerals (Morrison, 1994; Boschi et
al., 2006), leaving open the question of which comes fi rst: the
faults or the weak min-erals in the fault zones (e.g., Grasemann
and Tschegg, 2012).
As the core-complex concept was devel-oped, the relationships
among core-complex formation, magmatic and/or hydrothermal
activity, and ore deposition were recognized. Continental and
oceanic detachment faults are common sites for metallic ore
deposits, owing to the interaction of minerals and hot fl uids
(Roddy et al., 1988; Beaudoin et al., 1991; Smith et al., 1991).
For example, Cu-Fe sul-fi de and oxide deposits occur in the core
com-plexes of SE California and western Arizona (
Whipple-Buckskin-Rawhide; Spencer and Welty, 1986), and other core
complexes are associated with Au or Au-Ag deposits (South Mountain,
Arizona, USA; Valhalla Complex, Canada; Rhodope, Greece; Massif
Central, France) or uranium mineralization (Chapedony complex,
Iran; Yassaghi and Masoodi, 2011). In some oceanic core complexes,
hydrothermal Cu-Zn-Co-Aurich massive sulfi de deposits are
associated with ultramafi c rocks in detach-ment fault zones
(Mid-Atlantic Ridge; Fouquet et al., 2010).
In the following sections, we survey the pri-mary structural and
petrologic features of conti-nental and oceanic core complexes,
followed by discussion of the dynamics of core complexes in
different tectonic settings.
CONTINENTAL CORE COMPLEXES
Some of the most studied continental core-complex belts are in
the North American Cor-dillera (where core complexes formed during
diachronous collapse of thickened crust), the Aegean Sea/western
Turkey (where core com-plexes formed in the backarc setting
associated with rollback of the Hellenic subduction zones), and
Mongolia-China-Korea (where core com-plexes extend far into Eurasia
and their tectonic setting is not clear) (Fig. 2). Core complexes
have also been described along the entire length of the
Alpine-Himalayan orogenfrom the Pyrenees to SE Asiaas well as in
the older orogens of Europe (Caledonian, Variscan) and Asia (Fig.
2A). In addition, core complexes are reported in Papua New Guinea,
New Zealand, Antarctica, and various Precambrian terranes; it
remains controversial whether the Appalachian orogen contains core
complexes.
The belt of core complexes in the North American Cordillera,
from Mexico to Canada
(Fig. 2B), has been used as a type locality for understanding
continental core-complex dynamics of the crust and lithosphere,
although there are important differences in core-complex
development along the length of the belt. The Cordilleran core
complexes record differences in the nature of the interaction
between the shal-low and the deep crust, as shown by variations in
tectonic evolution in three regions (Fig. 2B): (1) a southern
core-complex belt (Mexico; Ari-zona, southern California, USA); (2)
a central belt, from the Snake Range (Nevada, USA) to the Raft
River complex (Utah, USA); and (3) a northern belt, from the
Pioneer Mountains (Idaho, USA) to the Shuswap Complex (British
Columbia, Canada). These three regions of core complexes vary in
age of extension (as young as Miocene in south, Eocene in north),
but, more importantly, in magnitude of exhuma-tion (with some
exceptions: least in south and most [tens of kilometers] in north),
and in the presence/involvement of partially molten crust (least in
south, most in north) (Vanderhaeghe and Teyssier , 2001; Rey et
al., 2009a).
Continental core complexes are typically ellip-tical, with a
long axis of ~1040 km (Fig. 3A); the footwall of core complexes is
typically elevated above the surrounding rocks, in some cases by 12
km of relief. There are a few core complexes that are substantially
larger (e.g., the Shuswap complex, Canada/United States; the
Menderes complex, Turkey; and some dome complexes in the Pamirs,
central Asia), but these typically con-tain several subsidiary
core- complex/dome struc-tures within them.
In the following sections, we survey structural and petrologic
features relevant to understand-ing the origin and evolution of
continental core complexes from the upper crust, through
detach-ment faults and shear zones, to the lower crust (Fig. 4).
The hanging wall and footwall in core complexes are mechanically
coupled in various ways, such that the geometry of the hanging wall
(e.g., multiple or single normal faults, gra-ben, half graben,
tilted blocks) is inherently tied to the ability of the deep crust
to fl ow and gener-ate a core complex (Block and Royden, 1990; Brun
and van den Driessche, 1994; Lavier et al., 2000; Rey et al.,
2009a, 2009b).
An important parameter controlling lower-crustal viscosityand
therefore coupling of deep and shallow crustis the geotherm. An
elevated geotherm appears to be necessary to the development of
continental core complexes, but models of the infl uence of
geotherm on core-complex generation can be divided into three
categories (Fig. 4). (1) In a warm crust, the deep crust is able to
fl ow, but the strong coupling between deep and upper crust results
in multiple upper-crust faults (Fig. 4A). (2) In a
hot crust, extension localizes in a single, large-offset
detachment fault system that arches as the low-viscosity deep crust
develops a core complex in the footwall (Fig. 4B). (3) In the
hottest crust case, the combination of localized upper-crust
extension and reduction of lower-crust viscosity by partial melting
results in exhumation of the deep crust; partially molten material
is exhumed nearly isothermally and undergoes complex deformation
during ascent, with contractional structures overprinted by
ex-tension (Fig. 4C).
Continental Core Complexes: Hanging-Wall Characteristics and
Processes
In most continental core complexes, hanging-wall rocks are
present, although these typically have been at least partially
removed by tectonic and/or erosional processes. In the North
Ameri-can Cordillera, hanging-wall rocks of some core complexes are
composed of unmetamorphosed to low-grade (meta)sedimentary and
volcanic rocks; in other core complexes in the Cordillera,
hanging-wall rocks are medium- to high-grade metamorphic rocks and
intrusions in which metamorphism and intrusion predated
core-complex development. In some of the Aegean core complexes, the
hanging wall consists of ophiolitic rocks and unmetamorphosed
sedi-mentary rocks, in places fi lling structural basins (Gautier
et al., 1993).
In some continental core complexes, supra-detach ment basins
make up a signifi cant frac-tion of the hanging wall, indicating
that the detachment came close to the surface during ex-tension.
For example, in the southern Basin and Range of the North American
Cordillera, some detachment faults intersect the surface, and
ad-jacent syntectonic basins contain sediments de-rived from the
footwall of the detachment (e.g., Miller and John, 1988; Miller and
John, 1999). In addition, the western detachment that bounds the
Shuswap core complex (Okanagan detach-ment in British Columbia) is
near the base of some basins and is itself cut by normal faults and
tilted to the east, so that the detachment here has the geometry of
a thrust (Vanderhaeghe et al., 1999, 2003). The normal faults that
cut the mylonitic detachment may root at depth into another
detachment system that formed as the core complex cooled during
exhumation. In this case, the basins, the detachment, and a part of
the footwall became the hanging wall for this new, hypothetical
detachment.
The basal units of supradetachment basins commonly record a high
paleogeothermal gradient during and after their deposition. For
example, study of coal units in a half graben above the detachment
on the north fl ank of the
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Whitney et al.
6 Geological Society of America Bulletin, Month/Month 2012
Montagne Noire core complex (France) record paleogeothermal heat
fl ow between 150 and 180 mW/m2 (Copard et al., 2000). Similarly,
thermo-chronologic studies of some continental core complexes show
that heat may be conducted (or advected via fl uids) from footwall
to hanging wall (e.g., Zeffren et al., 2005), resulting in
re-setting (or partial resetting) of isotopic systems in
hanging-wall rocks and minerals. Such a high
heat fl ow is consistent with the strong gradient that develops
during fast extension at the con-tact between hanging wall and
footwall (Rey et al., 2009b) and shows that, although detach-ment
faults represent a profound discontinuity in
pressure-temperature-time history of footwall relative to hanging
wall, rocks above and be-low the detachment fault may share a
late-stage thermal history.
Continental Core Complexes: Detachment Fault Characteristics and
Processes
In some continental core complexes, the detachment is not a
single fault but is made up of multiple, closely spaced and
anastomos-ing faults (Wernicke and Burchfi el, 1982). The upper
most detachment fault may have a particu-larly well-defi ned fault
plane, typically dipping
~ 50 km
extensionstrain
Moho
flowing partially
molten crust
solidus
solidus
contraction strain
"domino" rotation of upper crust blocks
flow of lower crust
ROLLING-HINGE MODEL FOR CORE COMPLEX FORMATIONBLOCK-ROTATION
MODEL FOR CORE COMPLEX FORMATION
flow of lower crust
1 2 3 4 5Order of faulting:
1
CONVERGING CHANNEL FLOW OF LOW-VISCOSITY(PARTIALLY MOLTEN) LOWER
CRUST
C1. C2.
incipient normal faulting, block rotation,and lower crustal
flow
extensional basins
incipient normal faulting, fault rotation, lower crustal flow;
fault 1 stops slipping, fault 2 takes over, resulting in passive
rotation of fault 1
channel flow
Moho
Moho
Moho
Moho
moho
pre-extension Moho
pre-extension Moho
pre-extension Moho
2
kinematic hinge
PLATEAU: THICK, HOT CRUST FORELAND: COLD CRUST
channel detachment
CORE COMPLEX DEVELOPED AT EDGE OF OROGENIC PLATEAU
pre-extension Moho
upper crust
lower crust
incipient kinematic hinge
transfer of thick crust toward foreland
flowing partially molten crust
rolling-hinge detachment
solidu
s
fixed
bo
un
dar
y
T
P
solid
us
shearing in channel
contraction
extension
A. WARM CRUST B. HOT CRUST
C. HOTTEST CRUST
strai
n hi
stor
y
P-T
pat
h
Figure 4. Modes of development of continental core complexes in
warm, hot, and hottest crust. (A) Warm crust exhumes continental
core complexes in footwall of normal faults that are distributed in
upper crust; for example, exhumation by domino-style rotation of
upper-crust blocks. (B) Hot crust focuses faulting in upper crust,
lead-ing to large-offset fault and exhumation of lower crust by
development of rolling-hinge detachment (after Brun and van den
Driessche, 1994). (C1) Hottest crust has signifi cant partial melt;
the low-viscosity lower crust fl ows in channels attracted by
focused zone of upper-crust extension; channels collide and move
upward to fi ll the gap created by upper-crust extension; deep
crustal rocks record signifi cant decompression and deformation
from con-traction to extension as they are exhumed (after Rey et
al., 2011). (C2) At the edge of an orogenic plateau, partially
molten crust fl ows owing to lateral gradients in gravitational
potential energy; note expected reversal of sense of shear
(kinematic hinge) between channel and rolling-hinge detachments
(after Teyssier et al., 2005).
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published online on 21 December 2012
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Continental and oceanic core complexes
Geological Society of America Bulletin, Month/Month 2012 7
30 (Figs. 1 and 3). A region of brecciation and
greenschist-facies alteration (recorded by secondary growth of
chlorite epidote) may characterize the structurally highest
(brittle) re-gions of detachment zones if suitable lithologies are
present (e.g., granitic gneiss). In addition, many detachment
faults record an early history as a ductile shear zone that evolved
into a brittle fault during tectonic and erosional denudation, as
well as during cooling driven in part by circu-lating fl uids
(Malavieille et al., 1990). The upper brittle fault zone is
typically underlain by a re-gion of high strain (and temperature)
gradient, up to ~1.5 km thick in the footwall (Wernicke, 1981;
Miller et al., 1983; Mueller and Snoke, 1993; Wells et al., 2000;
Foster and Raza, 2002; Mulch et al., 2006).
Although some of the deformation in this re-gion of high strain
may have developed during pre-extensional tectonism, the
deformation his-tory related to core-complex development can be
discerned through integrated structural and isotopic studies,
including geochronology. For example, although some shear zones
bounding Cordilleran core complexes record Mesozoic deformation,
the majority of Cordilleran core complexes show widespread
exposures of my-lonitized Tertiary rocks that formed during
ex-tension and core-complex development (e.g., Foster and Fanning,
1997; Vanderhaeghe et al., 1999; Wells et al., 2000; Gbelin et al.,
2011).
In the northern Cordilleran core complexes, the mylonite zone is
several hundred meters thick (Hyndman, 1980; Mulch et al., 2006;
Brown et al., 2012). It affects all lithologies but has an affi
nity for quartzite or marble if these units are present in the
extended crust. Sense of shear is typically unambiguous and
indicates footwall up, unless the detachment zone has been tilted
by late normal faulting or arched as a result of a rolling-hinge or
doming effects.
Detachment zones are self-exhuming struc-tures, and therefore
detachment-related my-lonite zones display a range of fabrics, from
ductile to brittle. A prevailing view is that high-temperature
mylonitic fabrics are progressively overprinted by
lower-temperature ductile fabrics followed by brittle processes
such as cataclastic fl ow and brecciation (e.g., Davis et al.,
1980). Based on the degree of refrigeration of footwall during
extension, mylonitic rocks may become incorporated into the hanging
wall and preserve a history of deformation fabrics formed at
vari-ous temperatures (Mulch et al., 2006).
Some core complexes have planar detach-ment faults (e.g.,
Bitterroot continental core complexes), but many continental and
oceanic core complexes are characterized by a corrugated
(undulating) detachment fault surface in which the corrugation axis
is oriented parallel to the
displacement (extension) direction and the undu-lations have
amplitudes of approximately tens to hundreds of meters and
wavelengths of hun-dreds of meters to tens of kilometers (John,
1987; Richard et al., 1990; Spencer and Reynolds, 1991; Cann et
al., 1997; Yin, 2004; Cannat et al., 2009). In oceanic core
complexes, these corruga-tions have been referred to as
megamullions (Christie et al., 1998; Tucholke et al., 1998).
Although some continental core complexes are bivergent, with
symmetric, oppositely dip-ping detachments on either side of a
footwall core (e.g., Hetzel et al., 1995a), many show structural
asymmetry in their footwall, particu-larly in those exhumed at the
edge of a continen-tal plateau (Teyssier et al., 2005). For
example, in the Shuswap core complex, British Columbia (Fig. 2B),
the continental core complex is asym-metric, with a series of
gneiss domes located in the immediate footwall of the eastern
(Columbia River) detachment system. The gneiss domes contain the
highest-grade metamorphic rocks exposed in the core complex,
including high-melt-fraction migmatite (Vanderhaeghe et al., 1999)
and gneiss containing sillimanite and cor-dierite pseudomorphs
after kyanite. These rocks recorded rapid near-isothermal
decompression at ~750 C to 800 C from 1 GPa to
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Whitney et al.
8 Geological Society of America Bulletin, Month/Month 2012
(e.g., 40Ar/39Ar in hornblende, biotite, musco-vite, K-feldspar;
apatite zircon fi ssion-track; and/or apatite zircon U-Th/He) are
used to capture temperature-time information as a function of
exhumation history on the detach-ment. Some detachment zones record
a slow cooling stage (515 C/m.y.) followed by more rapid cooling
(70100 C/m.y.) (Scott et al., 1998; Wells et al., 2000). This trend
has been interpreted to indicate a steepening of the fault through
time or other changes in detachment zone geometry. To interpret
thermochronol-ogy data in terms of detachment evolution, care must
be taken to understand the thermal struc-
ture of the detachment zone (e.g., position of isotherms,
calculation of geothermal gradients) and contribution of erosion to
denudation his-tory (Ketcham, 1996; Fayon et al., 2000).
Thermochronology-based estimates of aver-age slip rate for
detachments bounding Cor di-lleran core complexes range from ~1
mm/yr to 12 mm/yr (Foster et al., 1993; Scott et al., 1998; Foster
and John, 1999; Wells et al., 2000; Carter et al., 2004, 2006;
Foster et al., 2010), and simi-lar rates have been determined for
Aegean core complexes (John and Howard, 1995; Brichau et al., 2006;
Thomson et al., 2009). Studies involving multiple
thermochronometers are
able to detect possible changes in slip rate with time, such as
a signifi cant increase in slip rate proposed for some detachments
in the southern Basin and Range (Carter et al., 2004, 2006). The
increase has been ascribed to a change in the regional tectonic
regime, e.g., presence of a slab window beneath part of the Basin
and Range at ca. 15 Ma.
Detachment fault zones may be sites of sig-nifi cant fl uid
circulation and hydrothermal al-teration (Bartley and Glazner,
1985; Kerrich and Hyndman, 1986; Kerrich and Rehrig, 1987; Kerrich,
1988; Fricke et al., 1992; Famin et al., 2004; Person et al.,
2007). Fluid-mineral inter-action in detachment zones is relevant
to under-standing the chemical, thermal, and physical evolution of
detachment systems, the origin and location of ore deposits, and
the interpretation of low-temperature thermochronometry
results.
Locally, and likely transiently, water-rock ratios are high
during deformation at tempera-tures suffi cient for
(re)crystallization of miner-als in the fault zone, and detachment
zones are therefore characterized by greenschist-facies and
lower-grade minerals: typically, hydrous miner-als such as
chlorite, white mica, and epidote in continental core complexes,
and talc, chlorite, tremolite, and serpentine in oceanic core
com-plexes, as well as metalliferous ore deposits in both settings
(Smith et al., 1991; McCaig et al., 2010). Hydrous minerals and
other alteration products (e.g., from K-metasomatism) that form in
the fault zone may be important in the initia-tion and subsequent
structural evolution of the core complex, such as by promoting
strain lo-calization and/or affecting the thermal state and
therefore mode and pattern of faulting of the brit-tle crust
(Lavier and Buck, 2002). In addition, a vigorous hydrothermal
system in which mete-oric water circulates through faults in the
upper 1015 km of the extending crust (i.e., to the level of the
detachment fault) drives effi cient advec-tive removal of heat
(Morrison and Anderson, 1998; Famin et al., 2004; Person et al.,
2007).
The effects of fl uid circulation in extended upper crust can be
seen in fi eld and isotopic records. Vein systems and mineralized
fault zones are common in the hanging wall of de-tachments and
indicate that minerals precipi-tated from hot fl uids during ascent
and cooling. Stable isotope (particularly hydrogen) values of
hydrous minerals such as white mica, biotite, chlorite, and epidote
show that the water-rich fl uids that interact with minerals in
detach-ment zones are derived from various sources, including
meteoric sources at structurally high levels and
metamorphic/magmatic sources at lower levels (Kerrich and Hyndman,
1986; Spencer and Welty, 1986; Kerrich and Rehrig, 1987; Wickham
and Taylor, 1987; Baker et al.,
A
B
C D
Figure 5. (A) West-dipping, Eocene Okanogan detachment zone,
eastern Washington; the detachment footwall grades downward from
mylonite to migmatite across a 23 km sec-tion. (B) East-dipping,
Miocene Raft River detachment with remnants of hanging wall on
mylonitic quartzite; view is looking south from the Idaho-Utah
state line. (C) Photomicro-graph of mylonitic quartzite from the
Snake Range detachment, Nevada; mica fi sh, S-C fabrics, and quartz
crystallographic preferred orientation (c-axis, electron
backscatter diffraction measurements of >1000 grains) are well
developed and indicate top-to-the-east shear; stretched quartz
grains are partially to entirely recrystallized by subgrain
rotation in the dislocation creep regime. (D) Photomicrograph of
mylonitic quartzite from the Colum-bia River detachment that bounds
the Shuswap core complex to the east; mica fi sh and S-C fabrics
indicate top-to-the-east shear; quartz grains are recrystallized by
combination of subgrain rotation and grain boundary migration in
the dislocation creep regime.
as doi:10.1130/B30754.1Geological Society of America Bulletin,
published online on 21 December 2012
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Continental and oceanic core complexes
Geological Society of America Bulletin, Month/Month 2012 9
1989; Nesbitt and Muehlenbachs, 1995; Losh, 1997; Holk and
Taylor, 2007; Mulch et al., 2004; Gbelin et al., 2011). Studies of
detach-ments that are localized in simple lithologies (quartzite,
marble) have shown that, in general, meteoric water equilibrates
during detachment activity with neo/recrystallized hydrous phases
such as white mica at moderate temperatures (~350450 C) (Famin et
al., 2004).
Hydrogen isotopic values in synkinematic white mica (e.g., Figs.
5C and 5D) and other hydrous minerals in detachment zones have been
interpreted to indicate interaction of mica with a fl uid derived
from a high-elevation catch-ment (e.g., Columbia River fault at the
latitude of Thor-Odin, British Columbia, and at the lati-tude of
the Kettle dome, Washington; detach-ments in the Ruby Range and
Snake Range, Nevada; Fricke and ONeil, 1999; Mulch et al., 2007;
Gbelin et al., 2011). Hydrous minerals in detachment mylonites may
therefore contain the paleoelevation record over the time scale
(~15 m.y.) of mylonite formation (Mulch et al.,
2004). Studies that combine data from detach-ment-basin pairs
have verifi ed that the isotopic composition of mylonite minerals
matches that measured in basin strata deposited at high eleva-tion
(Mulch and Chamberlain, 2007).
Detachment zones typically accommodate large lateral
displacement and considerable thin-ning. Tens of kilometers of
displacement are ac-commodated by deformation in the detachment
zone over the time scale of 15 m.y., implying high average strain
rates. Quartz microstruc-ture in detachment zones provides
information on fl ow paleostress through the analysis of
re-crystallized grain size using paleopiezometers; temperature of
deformation can be derived from quartz-mica oxygen isotope
thermometry or from titanium-in-quartz thermometry. In the
quartzite mylonites of some core complexes in the North American
Cordillera, quartz dynamic recrystallization is dominated by
subgrain ro-tation, which is consistent with low recovery under
high-fl ow stress conditions in the disloca-tion creep regime
(Figs. 5C and 5D). Whatever
the temperature of deformation (350500 C) inferred from
quartz-mica stable isotope pairs (Mulch et al., 2006; Gottardi et
al., 2011; Gbelin et al., 2011), calculated fl ow stress is
typically high and quite constant over the en-tire mylonitic
section. This behavior suggests that localization or delocalization
of strain, cor-responding to increasing or decreasing strain rates,
respectively, is a response of the system to maintain fl ow stress
near the critical crustal strength, a property that is best
exemplifi ed in extensional systems (Mulch et al., 2006).
There has been much discussion of the dy-namics of low-angle
normal faults during core-complex development. Debate stems from
predictions from rock mechanics theory that extension produces
high-angle faults (~60; Ander son, 1942, 1951) and that low-angle
faults cannot slip, consistent with the proposal that moderate to
large earthquakes (M 5.5; Jackson, 1987; Jackson and White, 1989)
do not occur along these faults. This conclusion apparently confl
icts with fi eld observations that low-angle, and in some cases
subhorizon-tal, normal faults have been active in the brittle crust
(Wernicke , 1981; Reynolds and Spencer, 1985; Davis et al., 1986;
John, 1987; Wernicke and Axen, 1988; Scott and Lister, 1992; John
and Foster, 1993; Lecomte et al., 2010); for example, the
recognition that horizontal fi eld markers are close to the
detachment surface (e.g., Scott and Lister, 1992; John and Foster,
1993), and therefore that displacement on the detachment fault
occurred at low dips. In addi-tion, some seismically active,
low-angle (
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Whitney et al.
10 Geological Society of America Bulletin, Month/Month 2012
Continental Core Complexes: Footwall Characteristics and
Processes
Footwall rocks may record a wide range of agesfrom pre- to
synextensionof meta-morphic, magmatic, and deformation events.
Metamorphic grade may also vary within the footwall, not only as a
function of structural level exposed, but also owing to the complex
pressure-temperature (P-T ) paths that rocks fol-low before and
during extension. It is important to determine the thermal state of
the lithosphere during extension because the ambient synexten-sion
geothermal gradient is an important factor in controlling the
evolution of core complexes, as well as the role of crustal fl ow
in creating and maintaining a fl at Moho, as is observed in many
highly extended regions (Block and Royden, 1990; Buck, 1991; Rey,
1993) (Fig. 4).
The footwall of most core complexes has a domal structure (Fig.
3). The domal geom-etry has been explained for some continen-tal
core complexes as resulting from uplift of the detachment during
isostatic rebound as the hanging-wall rocks are extended, thinned,
and denuded by tectonic and erosional processes (Spencer, 1984;
Buck, 1988). These models as-sumed that the lower-crust fl owed at
fast rates, whereas later studies considered how weak the lower
crust must be to allow dome formation (Block and Royden, 1990;
Kruse et al., 1991; Wdowinski and Axen, 1992; McKenzie et al.,
2000; Rey et al., 2009a, 2009b, 2011).
Evidence for fl ow of weak crust can be seen in the complex
internal structure of gneiss (or mig-matite) domes, which occur
within some conti-nental core complexes. Some core complexes
contain one or more gneiss domes, typically beneath a carapace of
high- to medium-grade metamorphic rocks (Brun and van den
Driessche, 1994; Vanderhaeghe and Teyssier, 2001; Whitney et al.,
2004b); these have been called migmatite-cored metamorphic core
complexes (Rey et al., 2009a, 2009b), and their origin relates to
regional extension and fl ow of deep crust beneath detach-ment
faults. Studies have shown that the crystal-lization of the
magmatic portions of migmatite domes in core complexes was followed
by rapid cooling to T < 300 C (Fig. 6); cooling ages co-incide
with ages of synkinematic minerals (e.g., mica) in detachment fault
zones (Malavieille et al., 1990; Maluski et al., 1991; Kruckenberg
et al., 2008). These results show that high-tem-perature crustal fl
ow occurred during core-com-plex formation but ended when hot rocks
reached shallow crustal levels and cooled rapidly.
Gneiss (migmatite) domes commonly dis-play a relatively simple
external surface but a complex internal structure (e.g.,
Kruckenberg et al., 2011) that varies with level of exposure.
Some domes have a double-dome pattern con-sisting of two main
compartments defi ned by foliation divided by a steep, median
high-strain zone (Fig. 4C1); examples include the Naxos (Greece)
and Montagne Noire (France) core complexes (Rey et al., 2011).
Other domes show nappe-like recumbent folds that overprint earlier,
steeper structures (e.g., McFadden et al., 2010). These double-dome
and more complex three-dimensional structures are signifi cant for
understanding crustal fl ow under extension, and they provide a
framework for interpreting defor-mation features as a function of
time and space.
Numerical modeling provides insights to help understand the
relative infl uence of geothermal gradient, crustal thickness, and
crustal fl ow in core-complex initiation and evolution (e.g., Tirel
et al., 2008; Rey et al., 2009a, 2009b). For example, Tirel et al.
(2008) showed that for the model assumptions and parameters used, a
Moho T > 800 C is required to produce a core complex in
60-km-thick crust. At T < 800 C, a strong upper mantle
contributes to crustal-scale boudinage (necking). Insights from
numerical modeling relevant to the fl ow of deep crust, as well as
the origin of double-domes, are dis-cussed in a later section
(Dynamic Models of Continental Core-Complex Development).
Although some core complexes are associ-ated with regions of low
heat fl ow (e.g., south-ern California and Arizona; Lachenbruch et
al., 1994), others are associated with regions of high heat fl ow
(e.g., the islands of Naxos and Paros, in the central Aegean;
Gautier et al., 1993; Keay et al., 2001; Brichau et al., 2006;
Seward et al., 2009). These Aegean islands contain cores of
syntectonic migmatite and granite and have been proposed as the
site of a thermal anomaly. However, these core complexes do not
neces-sarily represent an anomalously hot region of the crust.
Instead, they may have formed in a zone of large-scale extension or
transtension that triggered the ascent of hot, ductile crust that
fl owed from deep to shallow levels and became involved in the
high-strain zone beneath the de-tachment, where isotherms collapsed
(Krucken-berg et al., 2011; Rey et al., 2011). According to this
idea, hot ductile crust was present region-ally, but only locally
exhumed.
Interpretation of the P-T conditions and paths of metamorphism
relevant to core-complex de-velopment (e.g., Buick and Holland,
1989) re-quires knowledge of the age of metamorphic events that
affected footwall rocks. As noted, footwall rocks may have
experienced multiple metamorphic events prior to metamorphism
related to extension and core-complex develop-ment, so care must be
taken (particularly with zircon) to determine the age of the last
metamor-phic event. In some Cenozoic core complexes,
for example, zircons in gneiss yield pre-Cenozoic ages,
recording protolith crystallization and/or later metamorphism
associated with pre-exten-sion crustal thickening (e.g.,
Kruckenberg et al., 2008). The age of extension and accompanying
exhumation of the footwall rocks is indicated by the youngest U-Pb
zircon and monazite ages, and is further bracketed by cooling ages
deter-mined by thermochronometers such as 40Ar/39Ar for hornblende,
micas, and K-feldspar (Fig. 6). In migmatite-cored core complexes,
melt that collected in boudin necks or extensional shear zones may
provide an indication of the timing of onset and/or cessation of
major extension (Gordon et al., 2008; McFadden et al., 2010).
Even in cases in which the extension-related temperature-time
path of footwall rocks is known, it can nevertheless be challenging
to understand the pressure (depth) evolution of footwall rocks, and
in particular it is diffi cult to determine the maximum pressure
experienced by rocks exhumed in a core complex; however, this is
important information for reconstructing pre-extension crustal
thickness and unravel-ing particle paths in footwall rocks. Some
core complexes exhume rocks from great depths, e.g., high-P rocks
such as eclogite or kyanite-bearing gneiss, although much of the
core-complex footwall may consist of metamorphic rocks recording
only a low-P, high-T history (Rey et al., 2009a, 2011). In some
cases, high-P rocks exhumed in a core complex record a much older
(pre-extension) metamorphic history (e.g., Precambrian eclogite in
the Cenozoic Menderes core complex, western Turkey; Candan et al.,
2001), although it is likely that exhumation of the high-P rocks
occurred during extension, and therefore the presence of these
rocks is relevant to understanding core-complex dynamics. It is
therefore particularly important to know the P-T-time history of
these rocks so as to be able to track the timing, magnitude, and
paths of their exhumation.
Some conceptual and numerical models for core-complex
development assume that the major motion of footwall rocks is
horizontal, except for some vertical motion related to arch-ing of
the footwall beneath the detachment (Brun and van den Driessche,
1994). Based on this as-sumption, predictions are made about
expected changes (or lack of changes) in metamorphic grade in
footwall rocks in the direction of tec-tonic transport for core
complexes with different thermal histories (Gessner et al., 2007).
This as-sumption is valid for some core complexes but not for those
in which rocks registered a major component of vertical motion, as
shown by a record of isothermal decompression, e.g., from >2030
km to
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Continental and oceanic core complexes
Geological Society of America Bulletin, Month/Month 2012 11
nental core complexes, this decompression has been interpreted
as driven by extension (Soto and Platt, 1999; Viruete et al., 2000;
Norlander et al., 2002). Additional insights about pressure history
come from observations of reaction tex-tures in footwall
metamorphic rocks (Krucken-berg and Whitney, 2011; Goergen and
Whitney, 2012) and from numerical modeling (Ruppel et al., 1988;
Rey et al., 2009a, 2009b).
Temperature-time paths for footwall rocks during
extension-related exhumation can be deter mined by application of
an array of thermo-chronometers with different closure
tempera-tures in different minerals (cf. similar methods applied to
detachment zone rocks; Fig. 6). In-terpretation of these data in
the context of the exhumation path (i.e., changes in depth
accom-panying cooling) requires estimates of the geo-thermal
gradient during extension. Similar to the T-time history of some
detachment zones, footwall rocks (below the detachment zone) may
record rapid cooling (>50 C/m.y.) fol-lowed by a more protracted
cooling history (
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Whitney et al.
12 Geological Society of America Bulletin, Month/Month 2012
Another possibility is that the detachment con-tinues
along-axis, but it is covered by extrusive rocks and rider blocks
(Escartn et al., 2008; Reston and Ranero, 2011).
In some regions of the Mid-Atlantic Ridge, the hanging wall is
affected by intense hydro-thermal circulation (Fig. 1B); this has
led to the suggestion that hydrothermal activity at slow-spreading
systems relates to the dynamics of the underlying detachment fault
and does not directly relate to magmatism (Petersen et al., 2009).
However, it has also been proposed that the detachment serves as a
conduit for fl uids and links the fractured hanging wall and
associated hydrothermal fi elds at the ocean fl oor to magma
chambers at depths of ~510 km (e.g., Escartn et al., 2003; McCaig
et al., 2007, 2010; McCaig and Harris, 2012).
Intrusion and extrusion of basalt clearly add to the hanging
wall. Extrusion is important in the formation of hanging-wall rider
blocks (Reston and Ranero, 2011; Choi and Buck, 2012), and the
formation of oceanic detach-ments may depend on the amount of dike
intru-sion into the hanging wall (Buck et al., 2005). This topic is
discussed more in the section Dynamic Models of Oceanic
Core-Complex Development.
Oceanic Core Complexes: Detachment Fault Characteristics and
Processes
Geological sampling of oceanic detachments by coring and
dredging indicates that the zone of brittle deformation in some
core complexes may be
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Continental and oceanic core complexes
Geological Society of America Bulletin, Month/Month 2012 13
Oceanic core-complex detachments gener-ally display a domal
structure as well as shorter-wavelength corrugations with axes
parallel to the extension direction; these corrugations have
amplitudes up to several hundred meters (Cann et al., 1997;
Tucholke et al., 1998, 2008; Ranero and Reston, 1999; Cannat et
al., 2006). The longer wavelengths, which generate domal oceanic
core complexes, may be related to mag-matic versus tectonic
extension, and the shorter wavelength (corrugations) may be related
to variations in both space and time of melt bodies that feed the
spreading axis (Lin et al., 1990) and that ultimately control the
rheology of detach-ment root zones.
Fault rocks associated with detachment faults in some oceanic
core complexes record similar temperatures to those in continental
core complexes: ~550300 C (amphibolite to greenschist facies;
Karson, 1999); in others (Atlantis Bank, Southwest Indian Ridge),
defor ma tion conditions ranged from granulite to greenschist
facies (from >900 C to
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Whitney et al.
14 Geological Society of America Bulletin, Month/Month 2012
spreading center is between ~3500 and 4800 m depth. Shallower
seafl oor is likely to have too large a magma supply, and deeper
seafl oor is likely to have too little magma. Fast-spreading ridges
have very thin lithosphere (~1 km thick), so even a modest supply
of magma could ac-commodate all lithospheric spreading. For
oce-anic core complexes to form, the magma supply apparently cannot
be too high or too low (i.e., the Goldilocks condition of Tucholke
et al., 2008). If faults on opposite sides of the axis are active
simultaneously and slip at the same rate, then a core complex can
form on both sides of the ridge (even in cases in which the plate
sepa-
ration rate is not accommodated by magma-tism in the form of
dike opening, as suggested by Schouten et al., 2010). However, any
rate-dependent weakening (or healing, as in Buck et al., 2005)
leads to only one fault being active at a given time. The inference
that low values of magma supply result in a complex pattern of
moderate-offset, crosscutting faults argues for there being one
dominantly active fault at a time in the axial region (Cannat et
al., 2006).
However, as little of the seafl oor is mapped with suffi cient
resolution to detect oceanic core complexes, it is likely that they
are more com-mon than the Tucholke et al. (2008) study in-dicates.
For example, the correlation of oceanic core complexes with higher
than normal rates of seismicity and hydrothermal activity may
indi-cate that ~50% of a long (75100 km) section of the
Mid-Atlantic Ridge could involve detach-ments (Smith et al., 2006;
Escartn et al., 2008).
CORE COMPLEXES AND LITHOSPHERE DYNAMICS
Core complexes form in regions of extension driven by surface
forces at plate boundaries, or volume forces in relation to lateral
variation of gravitational potential energy, or both. World-wide,
most core complexes form in regions of extension in collapsed
orogens and at relatively slow-spreading mid-ocean-ridge systems.
Oro-genic collapse may occur under free boundary conditions, such
as driven by slab rollback at convergent margins, or under a fi xed
bound-ary, owing to spreading of thick, weak crust (Rey et al.,
2001). In these settings, extension is lithosphere scale, but the
major expression of extension may be in the crust because normal
faulting of the brittle upper crust is coupled with ascent of the
ductile crust, resulting in regions of extreme thermal and strain
gradients across detachment zones (meters to ~1.5 km thick).
During lithospheric extension, strain is natu-rally partitioned
in the crust into weak fault zones in the brittle layers, and shear
zones and homogeneous strain in the lower crust. Strain
localization may result from physical processes involving
thermodynamic energy fl uxes even in the absence of any particular
rheological anomaly. Whatever its origin, strain localiza-tion in
the upper crust is essential to initiate core complexes, and it is
easily achieved around rheological and/or density anomalies such as
a preexisting fault in the brittle upper crust (Buck, 1993; Lavier
et al., 1999; Koyi and Skelton, 2001; Gessner et al., 2007), a
rheological and/or density anomaly in the lower crust (Brun et al.,
1994; Tirel et al., 2004, 2008), or a strong den-sity discontinuity
along the brittle-ductile transi-tion (Wijns et al., 2005).
In the rest of this section, we fi rst use model results that
focus on the brittle layer to explore the mechanics that lead to
the development of detachment faults. We then examine more
com-plete lithospheric models in which the coupling of
rheologically realistic layers is investigated (brittle,
temperature-dependent viscous, par-tially molten) for the case of
continental core complexes, followed by discussion of the dy-namics
of oceanic core complexes. Regard-ing the broader geodynamic
settings in which continental core complexes form, we then ad-dress
two cases: (1) the role of mantle wedge dynamics in
Cordilleran-type orogens, where continental core complexes develop
in the con-tinental overlying plate; and (2) the transition between
an orogenic plateau and its foreland.
Mechanics of Core-Complex Faults
Mohr-Coulomb fracture mechanics imply that faulting occurs most
easily at an angle of ~30 to the maximum principal stress (e.g.,
Anderson, 1942). Assuming an Andersonian extensional stress fi eld
in which the minimum principal stress is horizontal, normal faults
in the brittle upper crust should initiate at dips ~60 (Anderson,
1942) and should be active at dips of no less than 30 (e.g.,
Sibson, 1985). Several authors have suggested that normal faults
could initiate with low dips if particular loads reorient the
tectonic stress fi eld (Spencer and Chase, 1989; Yin, 1989; Parsons
and Thompson, 1993). However, Wills and Buck (1997) showed that
even with these purpose-built stress fi elds, low-angle faults
would not form before high-angle faults.
An alternative to slip on low-angle normal faults is that the
upper parts of some actively slipping high-angle normal faults
rotate to shal-lower dips. Spencer (1984) suggested that the
isostatic response to offset of a low-angle nor-mal fault would
tend to decrease the dip of the fault. Although there is ample fi
eld evidence to suggest that detachment faulting occurs on low-
angle faults (e.g., Scott and Lister, 1992; John and Foster, 1993),
large rotation of high-angle faults is consistent with structures
seen in many continental core complexes (Hamilton, 1988; Wernicke
and Axen, 1988; Buck, 1988). More recent paleomagnetic studies of
oceanic core complexes have demonstrated a minimum of 4050 rotation
of footwall rocks below the detachment (Morris et al., 2009;
MacLeod et al., 2011), validating the rolling-hinge model for
low-angle normal and fault formation asso-ciated with plate
spreading at slow and ultraslow mid-ocean ridges.
In their conceptual rolling-hinge models, Wernicke and Axen
(1988) assumed local
crus
t
moho
mantlelithosphere
melt
shear zones
crus
t
axial valley
moho
corrug
ated
seafloo
rvolcanic
layer
mantlelithosphere
axial valley
crus
t
moho
0
5
~20 km
volcanic seafloor
dike
mantlelithosphere
0
5
~20 km
0
5
~20km
A VOLCANIC-VOLCANIC SEAFLOOR PAIR
B CORRUGATED-VOLCANIC SEAFLOOR PAIR
C SMOOTH-SMOOTH SEAFLOOR PAIR
(after Cannat et al., 2006)
Figure 7. Sketches of axial regions for three proposed modes of
slow to ultraslow spread-ing, shown in order of inferred
decreas-ing melt supply: (A) volcanic seafl oor and (B) corrugated
seafl oor occur at both slow- and ultraslow-spreading ridges,
whereas (C) smooth seafl oor occurs only at ultraslow-spreading
ridges. Horizontal dimensions are ~80 km across axis and ~40 km
along axis (modifi ed from Cannat et al., 2006).
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Continental and oceanic core complexes
Geological Society of America Bulletin, Month/Month 2012 15
isostasy , whereas Buck (1988) calculated the fl exural response
of lithosphere to loads caused by the offset of a high-angle normal
fault. Nor-mal fault offset is supported regionally by fl ex-ure,
and since the Moho remains fl at in many continental core complexes
owing to fl ow of low-viscosity lower crust, the elastic
lithosphere in which bending occurs is restricted to the upper
crust (Masek et al., 1994). By analogy with bending at subduction
zones (e.g., McAdoo et al., 1978; Bodine et al., 1981), applying a
reasonable rock yield strength (the stress needed to break and slip
on a fault) based on labora-tory measurements for an ~10 km brittle
layer gives the observed range of domal wavelengths and topographic
relief (Buck, 1988). For oce-anic core complexes, thin crust
combined with a small density contrast between crust and mantle
reduce the importance of crustal thickness varia-tions compared to
topographic variations.
The energetics of fault offset and layer bend-ing (Forsyth,
1992) predict that the initial orien-tation of a fault corresponds
to the least friction on the fault per unit of horizontal
displace-ment. This energy or work approach yields the same initial
orientation for a fault as the clas-sic Ander sonian stress
analysis, but when the fault records signifi cant slip, work is
done in the bending of the displaced layer. An analytical
es-timation of this extra work using an approxima-tion of the
lithosphere as a thin, perfectly elastic layer fl oating on an
inviscid substrate (Forsyth, 1992) showed that an initially
low-angle normal fault could accumulate much more offset than a
high-angle fault. If this elastic plate model is correct, the
initial dip of a normal fault has to be extremely low to
accommodate tens of kilome-
ters of offset. Such a fault could only initiate on a
preexisting very weak zone.
The inclusion of realistic yield stresses de-creases the
wavelength of fault and footwall bending and radically lowers the
work that results from fault-related topography (Buck, 1993). A
reasonably weak fault may accommo-date an offset suffi ciently
large that the inactive part of the fault rolls over and becomes fl
at. A fault that is not suffi ciently weak is replaced by another
fault before large offset develops (Buck, 1988, 1993). Faults may
weaken as they record slip, and the amount of fault weakening
needed to allow large fault offset increases lin-early with brittle
layer thickness (Buck, 1993). For example, a fault affecting a
10-km-thick layer would have to weaken by ~10 MPa to de-velop a
large offset.
Analog models are useful for simulating the early, small-offset
stage of fault development (e.g., Brun et al., 1994; Tirel et al.,
2006), but they cannot easily simulate the thermally con-trolled
strength evolution that is likely to affect large-offset faults.
Early two-dimensional nu-merical studies (e.g., King and Ellis,
1990) assumed a preexisting weak normal fault em-bedded in a purely
elastic layer and solved for the topographic relief and stress
changes when the fault recorded slip. Given the potential
importance of the fi nite brittle yield strength (plastic
deformation), early models treated the lithosphere as a
viscous-plastic layer and were not concerned with fault
localization (Braun and Beaumont, 1989; Bassi, 1991). Subsequent
extension models of a viscous-elastic-plastic layer incorporated
strain weakening as a func-tion of strain to promote normal fault
devel-
opment (Poliakov and Buck, 1998; Buck and Poliakov, 1998); these
models showed that a sequence of high-angle faults might form and
accommodate extension at a simple mid-ocean-ridge structure. When
more strain weakening is allowed, large- offset faults develop
(Lavier et al., 1999, 2000) (Fig. 8). If the amount of fault strain
weakening is dependent on the thickness of the brittle layer being
extended, then large-offset faults only develop when a layer is
thinner than a critical thickness.
The behavior and characteristics of core com-plexbounding faults
may inform not only how much faults weaken, but also how they
weaken with offset (Lavier et al., 2000). If faults weaken too
fast, the entire layer shatters, and bending around a fi rst fault
gives rise to secondary faults that delocalize extension. If faults
weaken too slowly with offset, then the initial fault does not
become suffi ciently weak before the extra resis-tance to slip
owing to topography makes it easier for another fault to break
elsewhere. A thin layer can generate a large-offset fault, while
multiple faults develop in a thicker layer (Fig. 9; Lavier and
Buck, 2002). This is consistent with the ob-servation that
large-offset normal faults are only seen in areas of
higher-than-average heat fl ow, where one expects thin brittle
crust. In areas of high heat fl ow and thick crust, the lower
con-tinental crust may fl ow easily (e.g., Block and Royden, 1990;
Buck, 1991), allowing a single large-offset fault (detachment) to
accommodate signifi cant extension.
Recent work suggests that relatively small features of core
complexes may help bound the amount of fault weakening with offset.
Kilome-ter-scale allochthonous rider blocks that are cut
1050
plastic strain
vertical exaggeration 1.5 : 1
-50 0 +50distance (km)
1500
-1500
0
1500
-1500
0
1500
-1500
0
vertical exaggeration 3.0 : 1
dept
h (k
m) 0
5
10
dept
h (k
m) 0
5
10
dept
h (k
m) 0
5
10
Incr
easi
ng e
xten
sion
STRAIN TOPOGRAPHY (m)
-50 0 +50distance (km)
-50 0 +50distance (km)
Figure 8. Left panels: Results of numerical model of extension
of a fl oating brittle Mohr-Coulomb layer with a single seeded
normal fault (top panel). Right panels: Corresponding topographic
profi les with vertical exaggeration. Progressive extension is
suffi cient to allow foot-wall to rotate; abandoned parts of fault
rotate to, and even past, horizontal (bottom-left panel). Cohesion
loss of fault is function of strain (decrease of 1/3 of initial
brittle yield strength of layer) and occurs linearly with fault
offset up to 1.5 km (from Lavier et al., 2000).
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Whitney et al.
16 Geological Society of America Bulletin, Month/Month 2012
off from the hanging wall and related basin in-fi ll, and are
transported on top of the footwall, are common in both continental
and oceanic core complexes (e.g., Reston and Ranero, 2011). Rider
blocks superimposed on large-offset nor-mal faults may not form if
a fault loses too much strength (Choi and Buck, 2012); a narrow
range of fault weakening relative to intact surrounding rock allows
the formation of large-offset faults with rider blocks. These
blocks form when the dominant form of weakening is by reduction of
fault cohesion, whereas faults that weaken primarily by friction
reduction do not generate distinct rider blocks. Either a lack of
infi ll or an extreme reduction of friction by serpentinization of
exhumed mantle rocks may explain the lack of rider blocks on some
oceanic core complexes.
Dynamic Models of Oceanic Core-Complex Development
Magmatism and associated hydrothermal effects are likely to be
signifi cant factors in oceanic core-complex development, but the
in-clusion of magmatism in numerical models is at an early stage.
The input of limited amounts of gabbro may be needed to allow
alteration of peridotite to weak minerals, as silica-rich magmatic
fl uids interact with peridotite (e.g., Ildefonse et al., 2007).
Such weakening would allow oceanic core-complex detachments to slip
at lower than normal levels of stress, but so far such weakening
has not been included in spreading center models, and most modeling
effort has focused on dikes.
One simple way of treating the effect of dike intrusion in
numerical models of long-term lithospheric extension was developed
by Buck et al. (2005). The rate of dike opening was
specifi ed by the fraction M of the plate separa-tion rate that
is accommodated by dike opening. For M = 0, dikes account for none
of the plate spreading; for M = 1, they accommodate all of it.
Dikes may supply much of the heat that keeps the axis hotter and
the axial lithosphere thinner than the lithosphere farther from the
axis.
A simple geometric argument shows how faults and dikes interact
at a ridge with a fi xed position of diking and fi xed thermally
defi ned strength structure. If one fault forms due to lithospheric
stretching, it should initially cut the thinnest axial lithosphere
on one side of the axis. If the fault moves away from the axis into
thicker lithosphere, it becomes more diffi cult for this fault to
accommodate deformation, even though it is weaker than the
surrounding litho-sphere. Eventually it will be easier to form a
new fault cutting the axis, and the fi rst fault will be replaced
by a new fault.
Models of large-offset normal faulting have evaluated the infl
uence of M on fault behavior (Lavier et al., 1999). For M = 0.5,
the hanging wall of the fault does not move away from the axis, so
the fault could build up potentially un-limited offset. For M close
to 1, the maximum fault offset can be small, whereas for M close to
0.5, the offset should be very large. A numeri-cal model of diking
and stretching (Buck et al., 2005) tested this conceptual model
using the nu-merical approach of Poliakov and Buck (1998). Results
show that normal fault offset varies greatly as a function of M;
large fault offset occurs when M = 0.5; for M = 0.95, the model
generates a fairly symmetric pattern of mainly inward-dipping,
small-offset faults, and a sym-metric axial valley. For values of M
< 0.5, the active fault gradually moves across the spread-ing
axis and may be cut by later faults. How-
ever, the fi xed temperature structure assumed by Buck et al.
(2005) is questionable when there is little magma input. If the
temperature structure is strongly affected by the advection
associated with fault offset, then the weakest part of the
lithosphere migrates with the active fault. Since the sensible and
latent heat of intruding dikes should dominate over this effect,
the simple symmetric strength structure assumed in Buck et al.
(2005) should be valid for M >~0.4.
Tucholke et al. (2008) included a more realis-tic, evolving
temperature structure that consid-ered advection and diffusion of
heat and latent heat of crystallization. This study showed that
large-offset faults could form for a range of M values between ~0.3
and 0.6 (M = 0.5 shown in Fig. 10). As noted earlier, several sets
of obser-vations indicate that oceanic core complexes de-velop when
the supply of magma to a spreading center is not too high and not
too low. Olive et al. (2010) found that the formation of
large-offset faults and oceanic core complexlike structures depends
more on the rate of dike opening than on the rate of gabbro
intrusion. There appears to be a reasonable correspondence between
the M-based models and the inference of similar modes of spreading
based on geological and geophysical data from slow-spreading ridge
systems (Cannat et al., 2006, 2009).
A potential problem with these simple mod-els concerns the rate
of oceanic detachment faulting that has been estimated using
magmatic and thermochronologic data. For example, in the Atlantis
Massif, much of the full plate spreading rate was taken up on the
oceanic de-tachment (Grimes et al., 2008). Estimates of the rate of
oceanic detachment slip range from ~14 km/m.y. at Atlantis Bank
(Baines et al., 2008), to ~24 km/m.y. at Atlantis Massif
(Grimes
0 50 100 150
10
30
20
0
dep
th (k
m)
Distance (km)
250C500C
700C
0 50 100 150
10
30
20
0
Distance (km)
250C
500C
700C
0
5
4
0
5
4
top
og
rap
hy
(km
)
Total Strain0.7
1.4
2.1
2.8Total Strain
0.7
1.4
2.1
2.8
Thin lithosphere Thick lithosphere
Figure 9. Results of numerical model calculations for extension
of thin and thick brittle lithosphere. Rheology is
viscous-elastic-plastic; viscous strength depends strongly on
temperature. Total strain is the square root of the second
invariant of the strain tensor. Hydrothermal circulation cools the
shallow lithosphere and infl uences the temperature fi eld.
Extension of hot, thin lithosphere leads to single large-offset
fault; extension of cooler and thicker layer results in multiple
normal faults (from Lavier and Buck, 2002).
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Continental and oceanic core complexes
Geological Society of America Bulletin, Month/Month 2012 17
et al., 2008), and up to ~38 km/m.y. at Godzilla Megamullion
(Ohara et al., 2001). John and Cheadle (2010) concluded that,
during the time period when the oceanic detachments were ac-tive,
they accounted for 60%100% of the total plate spreading. This would
give M values be-tween 0.0 and 0.4; these values are lower than
predicted to produce large-offset faults.
Dynamic Models of Continental Core-Complex Development
Localized thinning of the continental upper crust by extension
is isostatically compensated by the fl ow of deep crust into core
complexes (Block and Royden, 1990; Wdowinski and Axen, 1992).
Partially molten crust is particu-larly likely to facilitate
core-complex develop-ment during extension owing to the dramatic
decrease in viscosity associated with the pres-ence of melt. In the
case of fl uid-absent melt-ing reactions, which can be encountered
during heating and/or decompression, a positive feed-back between
extension-induced decompres-sion and partial melting may generate
migmatite
domes within continental core complexes (Teys-sier and Whitney,
2002; Whitney et al., 2004b).
Physical experiments show that a strongly coupled upper and
lower continental crust leads to distributed surface extension,
whereas me-chanical decoupling between weak lower crust and much
stronger upper crust leads to localized surface extension that
favors continental core-complex development (Brun et al., 1994).
The magnitude of coupling is strongly dependent on the geothermal
gradient. In a cold to nor-mal geothermal gradient, the lower crust
is me-chanically coupled to the upper crust and upper mantle. In
this case, strain in the lower crust is a response to plate
boundary forces. In contrast, under hotter conditions, the lower
crust is weak and fl ows in response to both tectonic stresses and
gravitational stresses.
Numerical modeling that integrates tempera-ture-dependent
viscosity (Fig. 11) addresses the competition between plate
boundary extension rate and the rate at which ductile crust can fl
ow. The fl ow of ductile crust is driven by gravita-tional stresses
and is controlled by the viscosity (temperature) of lower crust and
the gradient in
gravitational potential energy. Under high ex-tension rate and
low surface heat fl ow, the duc-tile crust is too strong to respond
to gravitational stresses over short time scales. In this case,
two-dimensional models predict that the lower crust extends and
thins rather homogeneously (Fig. 11, a1, with TMoho = 600 C),
redistributing stresses uniformly in the upper crust and upper
mantle, where numerous normal faults develop. That is, extension
under relatively cool geother-mal conditions results in more normal
faults in the brittle upper crust (graben, half graben, and rotated
blocks) and relatively homogeneous extension in the ductile crust
(Fig. 11A). The ductile crust thins and fl ows to accommodate the
topography of the brittle-ductile transition and that of the Moho.
Major normal faults in the upper crust evolve into strong shear
strain gradients in the lower crust (Figs. 11AB, a1a2, b1b2). In
the ductile crust, fl at foliations domi-nate, although some thin
regions of vertical foliation occur underneath major normal faults
(cf. Fig. 11, a2, b2).
When the geotherm is warmer, the lower crust is able to fl ow
over short time scales in
M = 0
M = 0.5
M = 0.7
0102030 10 20 30
0
10
5
dept
h (k
m)
0
10
5
dept
h (k
m)
0
10
5
dept
h (k
m)
distance (km)
magma injection zone
strain rate (s1)14 1216
magmatic accretion seafloor faulted surface seafloormodel
original seafloor
600C
600C
600C
0.9 m.y.
0.9 m.y.
0.8 m.y.
Figure 10. Snapshots of modeled fault behavior and seafl oor
morphology for values M = 0, 0.5, and 0.7; model allows thermal
evolution throughout run (after Tucholke et al., 2008). Structural
interpretation is superimposed on modeled distribution of strain
rate (1016 to 1012 s1); model time is indicated in panels at lower
right; dashed white line at bottom is 600 C isotherm and
approximates the brittle-ductile transition; dashed seafl oor is
original model seafl oor, red seafl oor is that formed dominantly
by magmatic accretion, and solid bold seafl oor is fault
surface.
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Whitney et al.
18 Geological Society of America Bulletin, Month/Month 2012
response to gravitational stresses (Fig. 11, a2, with TMoho =
700 C). Higher geothermal gra-dients result in more strongly
localized ex-tension in the upper crust, driving fl ow in the deep
crust. This illustrates that the amount of coupling between the
upper and lower crust is controlled by the crustal-scale
temperature distribution (Block and Royden, 1990; Lavier and Buck,
2002). As the lower-crust viscosity diminishes (for example, when
melt fraction increases), the lower crust fl ows upward be-neath
the region of upper-crustal necking; the ductile crust is hot and
suffi ciently weak to be decoupled from the upper crust. This
upward fl ow generates horizontal fl ow in two channels that
converge toward the zone of extension and move material in a
direction opposite to the mo-tion of the brittle upper crust and
upper mantle. This strongly partitioned fl ow results in the
formation of contractional structures (upright
folds, nappes) beneath the zone of upper-crust extension (Rey et
al., 2011). Therefore, exten-sion in the upper crust is coeval with
contrac-tion in the deep crust. The two-dimensional structure
beneath the dome-shaped detachment fault consists of a subvertical
high-strain zone that separates two subdomes, creating a double
dome (Figs. 4C and 11).
Double domes (or more complex dome geom-etries in three
dimensions) generated in this way may explain the presence of
relict high-P metamorphic rocks (typically dismembered layers and
pods) in domes that otherwise record low-P, high-T emplacement.
Two-dimensional model results predict that deep crustal rocks
(eclogite, granulite) and possibly lithospheric mantle will be
entrained and carried upward in steep high-strain zones. These
simple dy-namic models are run under steady extension but produce a
complex deformation sequence
for material that is exhumed from the lower crust. These
deformation stages occur at vari-ous P-T conditions and include
shearing in the lower-crust channel, horizontal contraction in the
domain of collided channels, and horizontal extension when material
moves up beneath the detachment zone.
In high-geotherm settings, upward fl ow of low-viscosity
material is localized, and there-fore the lower crust infl uences
upper-crustal stress and strain only in the extended region
(necking of upper crust). Upward advection of heat promotes surface
fl uid circulation (Mulch et al., 2004; Person et al., 2007; Gbelin
et al., 2011), which contributes to further strain local-ization in
the upper crust. In the brittle upper crust, extensional strain is
strongly focused, and the number of normal faults is limited;
typically a single large-offset detachment fault accom-modates
upper-crustal extension (Lavier et al.,
TMoho = 600C
TMoho = 700C
TMoho = 800C
TMoho = 900C
45 km
50 km
55 km
60 km
TMoho = 617C
TMoho = 702C
TMoho = 767C
TMoho = 845C
Temperature
Viscosity
Stressa1 b1
b3
air
b2a2
a3
a4
air
brittle crust
ductile crust
mantle
faults
A B
solidus
VARIABLE TMoho (600900C) CRUSTAL THICKNESS 60 KM VARIABLE
CRUSTAL THICKNESS (4560 KM)
b4
Figure 11. Infl uence of the geotherm/rheological profi le on
strain distribution in continental core complexes, investigated
using two-dimensional modeling (Ellipsis) to explore the effect of
different geothermal gradients (expressed here as different Moho
temperatures) on core-complex development. (A) A 60-km-thick
continental crust, in variable thermal state, is submitted to 1.5
m.y. of symmetric exten-sion (1.13 cm/yr at both sides, i.e., 2
1015 s1). (a1, a2) Extension under a cool geothermal regime (i.e.,
without partial melting) leads to homo geneously distributed
extension: There are more normal faults in the upper crust, and
deformation in the lower crust is more homo-geneously distributed.
(a3, a4) In contrast, warmer geotherms lead to more strongly
localized extension in the upper crust and a more local-ized fl ow
in the lower crust. The temperature regime controls the amount of
coupling between the upper and lower crust. (B) Infl uence of
crustal thickness on strain distribution. In this series, a
continental crust of increasing thickness is submitted to extension
(same velocities as A). The geotherm (steady state when crustal
thickness = 40 km) is allowed to evolve for 30 m.y. before
extension begins; extension lasts for 2 m.y. This experiment confi
rms the dominant role of rheological layering on the strain fi
eld.
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-
Continental and oceanic core complexes
Geological Society of America Bulletin, Month/Month 2012 19
2000). In this tectonic setting, buoyancy of the weak lower
crust plays a second-order role, and core complexes form even when
the lower crust is denser than the upper crust.
Subduction Dynamics and Continental Core Complexes
In active ocean-continent plate margins, the state of stress in
the upper plate can rapidly switch from contractional to
extensional de-pending on the interplay among (1) trench-nor-mal
velocity, (2) friction along the subduction interface, (3)
gravitational forces in the thick crust of the overriding plate,
and (4) traction imposed at the base of the overriding plate by the
buoyant mantle wedge (e.g., Billen, 2008). The effect of trench
rollback on the stability of the overriding plate has been studied
in detail (e.g., Faccenna et al., 2007; Becker and Fac-cenna,
2011), particularly for the Aegean region (slab rollback and
breakoff; Jolivet and Brun, 2010). Here, we focus on the role of
the mantle wedge because the transformation of strong lithospheric
mantle into weaker and more buoy-ant mantle, driven in part by the
release of fl uids from the subducting plate and partial melting of
the wedge, must have a signifi cant effect.
Under dynamic equilibrium (no deforma-tion in the overriding
plate), there is a balance between driving forces (ridge push, slab
pull, and gravitational forces) and resisting forces (generated by
friction along the subduction in-terface as well as other viscous
forces). Gravi-tational forces stored in the thick overriding
plate