Continental and Oceanic Core Complexes 1888 2013

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Continental and oceanic core complexes 1888 2013

CELEBRATING ADVANCES IN GEOSCIENCE

1,† 1 2 3 Donna L. Whitney , Christian Teyssier , Patrice Rey , and W. Roger Buck Invited Review 1Department of Earth Sciences, University of Minnesota, Minneapolis, Minnesota 55455, USA 2School of Geosciences, University of Sydney, Sydney NSW 2006, Australia 3Lamont-Doherty Earth Observatory, Columbia University, Palisades, New York 10964, USA

ABSTRACT rocks ± upper mantle in the footwall of the nor- Core complexes were fi rst recognized in the mal fault(s). The resulting structure is a core continents (e.g., Anderson, 1972; Coney, 1974, Core-complex formation driven by litho- complex, which occurs in both continental and 1980; Crittenden et al., 1980; Lister and Davis, spheric extension is a fi rst-order process of oceanic lithosphere (Figs. 1 and 2). Extension is 1989), and they have been identifi ed in the geo- heat and mass transfer in the Earth. Core- the direct driving force for core-complex devel- logic record from the Precambrian (Holm, 1996) complex structures have been recognized opment, but in continental settings, the far-fi eld to the present (Hill et al., 1992). Core complexes in the continents, at slow- and ultraslow- tectonic regime may be one of convergence, and were later identifi ed at slow- and ultraslow- spreading mid-ocean ridges, and at continen- therefore continental core complexes may occur spreading oceanic divergent zones (e.g., Cannat, tal rifted margins; in each of these settings, in orogenic settings under an overall regime of 1993; Cann et al., 1997; Blackman et al., 1998; extension has driven the exhumation of deep plate convergence. Tucholke et al., 1998; Karson, 1999; Ranero crust and/or upper mantle. The style of ex- As extension proceeds, heat and material are and Reston, 1999; Dick et al., 2000). Continen- tension and the magnitude of core-complex transferred from deep (hot, ductile) to shallow tal and oceanic core complexes have similar di- exhumation are determined fundamentally (cool, brittle) levels, driving vigorous fl uid fl ow mensions, fault geometry, and kinematics (Figs. by rheology: (1) Coupling between brittle and strongly infl uencing the location and mag- 1 and 3), and both involve exhumation of deeper and ductile layers regulates fault patterns nitude of subsequent extension. Interactions levels of the lithosphere to shallow levels (John in the brittle layer; and (2) viscosity of the among minerals, fl uids, and/or magma may pro- and Cheadle, 2010). fl owing layer is controlled dominantly by duce economically important mineral deposits, In this review, we discuss the origin and the synextension geotherm and the presence and young extensional zones may be sources significance of continental core complexes or absence of melt. The pressure-tempera- of hydrothermal activity during and after active and oceanic core complexes. Although the ture-time-fl uid-deformation history of core faulting. term metamorphic core complex is a common complexes, investigated via fi eld- and modeling-based studies, reveals the magnitude, rate, and mechanisms of advection of heat A continental detachment system and material from deep to shallow levels, as meteoric Figure 1. Continental (A) and water well as the consequences for the chemical and oceanic (B) core complexes dif- meteoric water σ1 physical evolution of the lithosphere, includ- fer in their primary rock types σ3 ing the role of core-complex development in (granitic and metasedimentary fluid crustal differentiation, global element cycles, rocks in continental core com- brittle flow layer and ore formation. In this review, we provide plexes vs. gabbro and serpenti- a survey of ~40 yr of core-complex literature, nite in oceanic core complexes), brittle-ductile discuss processes and questions relevant to and therefore in the minerals transition: ~300–400°C mylonite the formation and evolution of core com- that infl uence the rheology of (quartz rheology) zone plexes in continental and oceanic settings, fluids derived from crystallization ductile the complexes. Nevertheless, of metamorphic/igneous rocks ~15 km layer highlight the signifi cance of core complexes many fi rst-order processes of for lithosphere dynamics, and propose a few their origin and evolution are B oceanic detachment system possible directions for future research. similar, and therefore there are Oceanic core complex many similarities in their archi- sea water sea INTRODUCTION tecture. In B, the detachment ridge serpentinite water fault roots in gabbro magma at axis When the lithosphere is under extension, the sheared depth (option 1); option 2 con- fluid domi- brittle upper crust breaks and is displaced along gabbro nantly siders a “dry” spreading center flow σ1 brittle normal faults. When extension is concentrated in which the brittle detachment fluid layer σ3 on one or a few faults in a narrow region, ductile transitions to a ductile shear flow ? ? gabbro material ascends from deeper levels of the litho- zone at depth (lithosphere Option 1 ? ? sphere, resulting in exhumation of deep crustal boudinage). brittle-ductile Option 2 ? mylonite transition: ~600°C ductile ~10 km †E-mail: [email protected] (olivine rheology) melt layer

GSA Bulletin; March/April 2013; v. 125; no. 3/4; p. 273–298; doi: 10.1130/B30754.1; 13 ﬁ gures.

A 120°W0° 120°E

PL AMOR Lf No N. American Cordillera 60°N 60°N Fig. 2B MC-Pyr Fig. 2C SB ReVe Aegean Sh BB Po YOHHa To Rh Li AA Ni KS Hh LR Ch Lo SEGKEd Nx GM Xi Ba AD Ho At La SC DI M-Ca Ka DNCV Go 0° Ma 0°

Da, Nb Mid- Atlantic AB Ridge Pa Southwest Indian Australian- Ridge Antarctic 60°S Discordance 60°S continental core complex + Antarctica: Fosdick core complex, Marie Byrd Land oceanic core complex continental margin core complex 0° 120°E

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: AA—Alpi Apuane (Italy); AB—Atlantis Bank (SW Indian Ridge); AD—Ama Drime (Nepal); AMOR—Arctic segment of Mid-Atlantic Ridge; At—Atlantis Massif (Mid-Atlantic Ridge); Ba—Baja (Mexico); BB—Bay of Biscay; Ch—Chapedony (Iran); Da—Dayman (Papua New Guinea); DI—Doi Inthanon (Thailand); DNCV—Day Nui Con Voi (Vietnam); Ed—Edough (Algeria); GK—Grand Kabilye (Algeria); GM—Gurla Mandhata (Pamirs); Go—Godzilla; Ha—Harkin (China/Mongolia); Hh—Hohhot (China); Ho—Hongzhen (China); Ka—Kane (Mid-Atlantic Ridge); KS—Kongur Shan (Pamirs); La—Laojunshan (China); Lf—Lofoten (Norway); Li—Liaodong Peninsula (China); Ma—Malino (Indonesia); Lo—Louzidian (China); LR—Lora del Rio (Spain); M-Ca—Mid-Cayman spreading center; MC-Pyr—Massif Central (France–Pyrenees, France, Spain; includes Montagne-Noire); Nb—Normanby Island (Papua New Guinea); Ni—Niğde (Turkey); No—Norway rifted continental margin; Nx—Naxos (Greece); Pa—Paparoa (New Zealand); PL— Payer Land (Greenland); Po—Pohorje Mountains (Slovenia); Re—Rechnitz (Austria); Rh—Rhodope (Greece, Bulgaria); SB—southern Brittany (France); SC—Song Chay (China); Sh—Shaerdelan (China); SE—Sierra de las Estancias (Spain); To—Tormes (Spain); Ve— Veporic (Slovenia); Xi—Xiaoqinling (China); YOH—Yagan-Onch-Hayrhan (China/Mongolia).

description of core complexes on the conti- relevant to crustal evolution and seismogenesis DEFINITIONS nents, for the sake of simplicity (and symme- in extending lithosphere, and core complexes try), in this review we use the terms continental have therefore been intensively studied using Herein, we use the following general defi ni- core complex (equivalent to metamorphic a variety of techniques, e.g., fi eld-based stud- tion of a core complex and the processes that core complex) and oceanic core complex. We ies (e.g., Davis and Coney, 1979; Miller et al., drive core-complex formation: do not discuss in any detail the formation of 1983; Bozkurt and Park, 1994; Gessner et al., A core complex is a domal or arched geologic continental margin core complexes, although 2001), numerical modeling (e.g., Buck et al., structure composed of ductilely deformed rocks some locations are highlighted on a world map 1988; Lavier et al., 1999; Tirel et al., 2004, and associated intrusions underlying a ductile- (Fig. 2A). 2008; Rey et al., 2009a, 2009b; Allken et al., to-brittle high-strain zone that experienced tens Over the past ~40 yr, interest in core com- 2011), and analog modeling (e.g., Brun et al., of kilometers of normal-sense displacement in plexes has remained high because these struc- 1994; Tirel et al., 2006). response to lithospheric extension. tures are common in extending orogens and There remain important questions about The lithospheric extension that results in along slow-spreading oceanic divergent zones, core-complex initiation and evolution. In this core-complex formation is commonly driven by and because they record fundamental thermo- review, we integrate knowledge derived from plate divergence, such as at mid-ocean ridges mechanical processes in extending lithosphere. different types of investigations (fi eld, mod- and along rifted continental margins. Extension An understanding of the uplift and exhuma- eling) of continental and oceanic core com- also occurs in plate convergence settings by slab tion of ductile rocks below low-angle normal plexes and discuss some of these unresolved rollback (e.g., the backarc of an oceanic subduc- faults, as well as the dynamics of the faults, is issues. tion zone) or by orogenic collapse under fi xed

274 Geological Society of America Bulletin, March/April 2013 Continental and oceanic core complexes

B 120°W C Rhodope Massif COAST RANGES- Shuswap Complex Istanbul NORTH CASCADES Frenchmans Cap Thasos Thor- NORTHERN BELT Odin Okanogan 110° W Valhalla Kettle 40°N CANADA 50°N Kazdağ Massif USA Priest River Lew COLUMBIA PLATEAU is & Cla TURKEY Clearwater rk Bitterroot Anaconda IDAHO BATHOLITH Gulf of Corinth İzmir

CENTRAL BELT Pioneer Athens RO Andros Samos C SNAKE RIVER PLAIN K Tinos

40°N Albion- M Kea Menderes Massif Grouse Creek O Raft U Syros Naxos

Pacific Ruby-Humboldt River TA Paros

Ocean N GREECE

SIERRA F Sifnos O

NEVADA Snake L D Ios

Range 40°N

BATHOLITH & Milos

THR Rhodes S a n

An US Aegean Sea dr

T e a B s Fa E

u T lt

Chemehuevi COLORADO PLATEAU Crete Whipple Mts 120°W Buckskin-Rawhide Harcuvar SOUTHERN BELT Harquahala Mts 24° 26° 28° South Mts

PENINSULAR Picacho Catalina RANGES BATHOLITH Rincon USA MEXICO 30°N Figure 2 (continued). (B) Schematic map of the core complexes in the North American Cordillera. (C) Schematic map of the core complexes in the Aegean Sea and surrounding regions. Sierra Mazatan 30°N

110°W

boundary conditions or even during slow plate during exhumation of the footwall. Some de- are normal faults with signiﬁ cant displacement, convergence (Rey et al., 2001). Core complexes tachment faults record a progression from but they do not bound core complexes. occur in all of these settings. high-temperature (mylonitic) deformation to “Detachment zone,” “detachment shear In continental core complexes, the normal- much lower-temperature brittle deformation zone,” and “detachment system” are terms that sense high-strain zone corresponds to a pro- through time. have been used to describe the diffuse zone of found metamorphic and/or stratigraphic dis- The concept of “detachment” has also been strain, up to 1.5 km thick, that underlies the continuity typically called a “detachment fault” applied to normal faults that are not associated uppermost fault plane in some core complexes, (Fig. 1), which is so named because rocks directly with core-complex structures, e.g., the such as those of the northern U.S. and Canadian above and below the fault zone record different South Tibetan detachment system (e.g., Burg Cordillera (e.g., Mulch et al., 2006; Gébelin pressure-temperature-time-deformation his to- et al., 1984), which is the northern boundary of the et al., 2011). Characteristic features of detach- ries. Although the uppermost part of the fault Himalayan crystalline wedge, and normal faults ments and ideas and debates about the dynamics zone may be a prominent brittle fault surface, in exhumed subduction complexes, such as on of low-angle normal faults are discussed in later below this fault there may be a broad (hundreds the island of Crete (e.g., Ring et al., 2001). There sections. of meters) zone in which ductilely deformed are also detachment-like faults in some conti- Although “detachment” is widely used to rocks have been kinematically and thermally nental arcs, such as the Andes (e.g., Mpodozis describe core complex–bounding faults, other linked—for some or all of their defor ma tion and Allmendinger, 1993) and the North Cas- terms are also used: for example, “low-angle history—to the structurally higher, brittle fault cades (Paterson et al., 2004). These detachments normal fault” (e.g., Axen, 2007). The term

Geological Society of America Bulletin, March/April 2013 275 Whitney et al.

A Northern Snake Range, Nevada (USA) B Kane “megamullion”, Mid-Atlantic Ridge in the continents: Spencer, 1984; Wernicke and Axen, 1988; Brun and van den Driessche, 1994; footwall western fault north hanging wall (3.3 m.y.) in the oceans: Lavier and Manatschal, 2006; ~10 km eastern fault MacLeod et al., 2009). Crustal fl ow beneath the (2.1 m.y.) extending upper crust has also been proposed to explain the domal structure of many conti- Sacramento Pass nental core complexes, as well as the moderate Basin Mid-Atlantic Ridge axis 2 topographic relief of the exhumed footwall de- detachment 4 spite tens of kilometers of offset and the exis- (km) tence of a fl at Moho in many highly extended 10 km 6 seafloor depth regions (e.g., Block and Royden, 1990; Buck, westSnake Range, Nevada east westKane east 1991; McKenzie et al., 2000). Exploration of the 0 2 MAR thermal and mechanical consequences of large- + + oceanic Axis magnitude crustal fl ow from deep to shallow core complex

depth (km) 4 30 crustal levels, including consideration of the re- 20 40 60 80 continental 10 20 30 40 core complex seafloor depth (km) lationship between continental core complexes west Thor-Odin (Shuswap), Canada east west Atlantis Massif east and gneiss domes (Teyssier and Whitney, 2002), shows that core complexes can be signifi cant 2 0 MAR oceanic sites for heat and mass transfer and have played migmatite Axis core complex a role in differentiation of continental crust depth (km) 4 30 20 4060 80 100 120 140 10 20 30 through geologic time (e.g., Rey et al., 2009a, seafloor depth (km) distance (km) distance (km) 2009b; Thébaud and Rey, 2012). Figure 3. Scale and surface expression of representative (A) continental and (B) oceanic core There has long been debate about the dynam- complexes. Snake Range cross section in A is modifi ed from Miller et al. (1983); the satel- ics of low-angle normal faults, primarily fo- lite image is from Google Earth. Kane oceanic core complexes images are modifi ed from cused on the question of whether faults of low Tucholke et al. (2008). MAR—Mid-Atlantic Ridge. (<30°) to very low (<10°, including subhorizon- tal) dip are mechanically possible or whether they represent former high-angle faults that have rotated to lower angles (Jackson, 1987; “décollement ” has also been used (Coney, 1980; as “Cordilleran-style metamorphic core com- Wernicke, 1981; Davis et al., 1986; Buck, 1988; Davis et al., 1980) and can be synonymous with plexes” (e.g., Aegean [Fig. 2C]—Lister et al., Scott and Lister, 1992; Axen, 2007). The ques- detachment in the context of a core complex, in 1984; West Antarctica—Richard et al., 1994; tion of the initial dip of detachments in oceanic some cases indicating a low-angle master fault Norwegian Caledonides—Steltenpohl et al., core complexes has been resolved in favor of or shear zone. “Breakaway fault” refers to the 2004; Iran—Verdel et al., 2007). high dip angles (Smith et al., 2008; Morris et al., fault identifi ed as the detachment that initially The existence of detachment faults and core 2009), but it remains unresolved for continental intersected Earth’s surface. complex–like structures was proposed for slow- core complexes. Some detachment faults are “listric”; this in- spreading mid-ocean ridges before such struc- Another debate relates to the role of fl uids dicates a curving fault with a dip that changes tures were observed in detail (Karson and Dick, in core-complex development. Deformation in from steep at shallow crustal levels to sub hori- 1983; Cannat, 1993; Tucholke and Lin, 1994). continental core complexes is strongly infl u- zontal at depth. In contrast, some detachment Fault-bounded domal structures resembling enced by the behavior of quartz and feldspars or faults (or parts of detachment faults) in both continental core complexes were later identi- calcite (particularly in the marble-rich Tethyan continental and oceanic settings are “downward fi ed in seafl oor images of the Mid-Atlantic belt of the Alpine-Himalayan orogen), as well concave” and accommodate large offset yet Ridge (Cann et al., 1997; Blackman et al., 1998; as by hydrous minerals (e.g., chlorite, white moderate topography (e.g., van den Driessche Tucholke et al., 1998; Ranero and Reston, 1999; mica) that form during fl uid infi ltration of the and Brun, 1992; Lavier and Manatschal, 2006). Tucholke et al., 2001) and confi rmed by ob- detachment zone. Deformation in oceanic core servation of samples collected from low-angle complexes is infl uenced by plagioclase feld- DEVELOPMENT AND EVOLUTION normal fault zones bounding these structures spar and olivine in gabbro, as well as by hy- OF THE CONCEPT (MacLeod et al., 2002; Escartín et al., 2003; drous minerals (e.g., talc, chlorite, serpentine, Schroeder and John, 2004; Ildefonse et al., and tremolite) that form in hydrothermally al- As a result of debate, primarily in the 1960s– 2007). Oceanic core complexes have now been tered peridotite and gabbro of the detachment 1970s, about the tectonic history of “basement recognized along segments of the Southwest zone. For both settings, an important question uplifts” (exposures of middle- and lower-crustal Indian Ridge (Dick et al., 2000; Baines et al., is whether the presence of weak minerals is a rocks) bounded by low-angle faults in the North 2003) and at other divergent zones (e.g., the prerequisite for slip or whether weak minerals American Cordillera (Figs. 2A and 2B), a model Caribbean–North American Ridge—Hayman mostly form after a fault has initiated, for exam- of core complexes as extensional structures et al., 2011; the Australian-Antarctic Discor- ple, as a result of hydrothermal activity in the was developed (Anderson, 1972; Coney, 1974; dance—Christie et al., 1998) (Fig. 2A). fault zone. Detachment faults are typically ori- Wright et al., 1974; Proffett, 1977; Davis and Many models for core-complex develop- ented at a high angle relative to the maximum σ Coney, 1979; Coney and Harms, 1984). The ment have invoked isostatic rebound beneath principal stress ( 1) in the extending region, concept has been applied to similar structures the detachment fault to explain the arching of and therefore may require a low coeffi cient of elsewhere, and some of these are described the fault and exhumation of the footwall (e.g., friction in order to initiate and slip (Wernicke,

276 Geological Society of America Bulletin, March/April 2013 Continental and oceanic core complexes

1981). However, weak zones (faults) could re- (Fig. 2B), has been used as a type locality hot crust, extension localizes in a single, large- sult from fl uid-rock interaction that increases for understanding continental core-complex offset detachment fault system that arches as pore pressure or generates phyllosilicates and dynamics of the crust and lithosphere, although the low-viscosity deep crust develops a core other weak minerals (Morrison, 1994; Boschi there are important differences in core-complex complex in the footwall (Fig. 4B). (3) In the et al., 2006), leaving open the question of development along the length of the belt. The hottest crust case, the combination of localized which comes fi rst: the faults or the weak min- Cordilleran core complexes record differences upper-crust extension and reduction of lower- erals in the fault zones (e.g., Grasemann and in the nature of the interaction between the shal- crust viscosity by partial melting results in Tschegg, 2012). low and the deep crust, as shown by variations exhumation of the deep crust; partially molten As the core-complex concept was devel- in tectonic evolution in three regions (Fig. 2B): material is exhumed nearly isothermally and oped, the relationships among core-complex (1) a southern core-complex belt (Mexico; Ari- undergoes complex deformation during ascent, formation, magmatic and/or hydrothermal zona, southern California, USA); (2) a central with contractional structures overprinted by ex- activity, and ore deposition were recognized. belt, from the Snake Range (Nevada, USA) tension (Fig. 4C). Continental and oceanic detachment faults are to the Raft River complex (Utah, USA); and common sites for metallic ore deposits, owing (3) a northern belt, from the Pioneer Mountains Continental Core Complexes: to the interaction of minerals and hot fl uids (Idaho, USA) to the Shuswap Complex (British Hanging-Wall Characteristics and Processes (Roddy et al., 1988; Beaudoin et al., 1991; Columbia, Canada). These three regions of core Smith et al., 1991). For example, Cu-Fe sul- complexes vary in age of extension (as young In most continental core complexes, hanging- fi de and oxide deposits occur in the core com- as Miocene in south, Eocene in north), but, wall rocks are present, although these typically plexes of SE California and western Arizona more importantly, in magnitude of exhuma- have been at least partially removed by tectonic (Whipple-Buckskin-Rawhide; Spencer and tion (with some exceptions: least in south and and/or erosional processes. In the North Ameri- Welty, 1986), and other core complexes are most [tens of kilometers] in north), and in the can Cordillera, hanging-wall rocks of some core associated with Au or Au-Ag deposits (South presence/involvement of partially molten crust complexes are composed of unmetamorphosed Mountain, Arizona, USA; Valhalla Complex, (least in south, most in north) (Vanderhaeghe to low-grade (meta)sedimentary and volcanic Canada; Rhodope, Greece; Massif Central, and Teyssier , 2001; Rey et al., 2009a). rocks; in other core complexes in the Cordillera, France) or uranium mineralization (Chapedony Continental core complexes are typically ellip- hanging-wall rocks are medium- to high-grade complex, Iran; Yassaghi and Masoodi, 2011). tical, with a long axis of ~10–40 km (Fig. 3A); the metamorphic rocks and intrusions in which In some oceanic core complexes, hydrothermal footwall of core complexes is typically elevated metamorphism and intrusion predated core- Cu-Zn-Co-Au–rich massive sulfi de deposits above the surrounding rocks, in some cases by complex development. In some of the Aegean are associated with ultramafi c rocks in detach- 1–2 km of relief. There are a few core complexes core complexes, the hanging wall consists of ment fault zones (Mid-Atlantic Ridge; Fouquet that are substantially larger (e.g., the Shuswap ophiolitic rocks and unmetamorphosed sedi- et al., 2010). complex, Canada/United States; the Menderes mentary rocks, in places fi lling structural basins In the following sections, we survey the pri- complex, Turkey; and some dome complexes in (Gautier et al., 1993). mary structural and petrologic features of conti- the Pamirs, central Asia), but these typically con- In some continental core complexes, supra- nental and oceanic core complexes, followed by tain several subsidiary core- complex/dome struc- detachment basins make up a signifi cant frac- discussion of the dynamics of core complexes in tures within them. tion of the hanging wall, indicating that the different tectonic settings. In the following sections, we survey structural detachment came close to the surface during ex- and petrologic features relevant to understand- tension. For example, in the southern Basin and CONTINENTAL CORE COMPLEXES ing the origin and evolution of continental core Range of the North American Cordillera, some complexes from the upper crust, through detach- detachment faults intersect the surface, and ad- Some of the most studied continental core- ment faults and shear zones, to the lower crust jacent syntectonic basins contain sediments de- complex belts are in the North American Cor- (Fig. 4). The hanging wall and footwall in core rived from the footwall of the detachment (e.g., dillera (where core complexes formed during complexes are mechanically coupled in various Miller and John, 1988; Miller and John, 1999). diachronous collapse of thickened crust), the ways, such that the geometry of the hanging In addition, the western detachment that bounds Aegean Sea/western Turkey (where core com- wall (e.g., multiple or single normal faults, gra- the Shuswap core complex (Okanagan detach- plexes formed in the backarc setting associated ben, half graben, tilted blocks) is inherently tied ment in British Columbia) is near the base of with rollback of the Hellenic subduction zones), to the ability of the deep crust to fl ow and gener- some basins and is itself cut by normal faults and Mongolia-China-Korea (where core com- ate a core complex (Block and Royden, 1990; and tilted to the east, so that the detachment plexes extend far into Eurasia and their tectonic Brun and van den Driessche, 1994; Lavier et al., here has the geometry of a thrust (Vanderhaeghe setting is not clear) (Fig. 2). Core complexes 2000; Rey et al., 2009a, 2009b). et al., 1999, 2003). The normal faults that cut have also been described along the entire length An important parameter controlling lower- the mylonitic detachment may root at depth into of the Alpine-Himalayan orogen—from the crustal viscosity—and therefore coupling of another detachment system that formed as the Pyrenees to SE Asia—as well as in the older deep and shallow crust—is the geotherm. An core complex cooled during exhumation. In this orogens of Europe (Caledonian, Variscan) and elevated geotherm appears to be necessary to case, the basins, the detachment, and a part of Asia (Fig. 2A). In addition, core complexes are the development of continental core complexes, the footwall became the hanging wall for this reported in Papua New Guinea, New Zealand, but models of the infl uence of geotherm on new, hypothetical detachment. Antarctica, and various Precambrian terranes; it core-complex generation can be divided into The basal units of supradetachment basins remains controversial whether the Appalachian three categories (Fig. 4). (1) In a warm crust, commonly record a high paleogeothermal orogen contains core complexes. the deep crust is able to fl ow, but the strong gradient during and after their deposition. For The belt of core complexes in the North coupling between deep and upper crust results example, study of coal units in a half graben American Cordillera, from Mexico to Canada in multiple upper-crust faults (Fig. 4A). (2) In a above the detachment on the north fl ank of the

Geological Society of America Bulletin, March/April 2013 277 Whitney et al.

A. WARM CRUST B. HOT CRUST BLOCK-ROTATION MODEL FOR CORE COMPLEX FORMATION ROLLING-HINGE MODEL FOR CORE COMPLEX FORMATION extensional basins 1 2

incipient normal faulting, fault rotation, lower crustal flow; incipient normal faulting, block rotation, fault 1 stops slipping, fault 2 takes over, and lower crustal flow Moho resulting in passive rotation of fault 1 Moho "domino" rotation of upper crust blocks Order of faulting: 12 3 45

flow of lower crust flow of lower crust Moho Moho

pre-extension Moho ~ 50 km pre-extension Moho

C. HOTTEST CRUST C1. CONVERGING CHANNEL FLOW OF LOW-VISCOSITY C2. CORE COMPLEX DEVELOPED AT EDGE OF OROGENIC PLATEAU (PARTIALLY MOLTEN) LOWER CRUST

PLATEAU: THICK, HOT CRUST FORELAND: COLD CRUST channel detachment incipient kinematic hinge

upper crust solidus lower crust Moho transfer of thick crust fixed boundaryfixed shearing in channel flowing toward foreland P s partially y u r d h i o molten crust l t contraction t a s o i s p h extension -T n P i a strain extension tr s

T kinematic hinge rolling-hinge detachment

us lid so solidus channel moho flow flowing partially molten crust pre-extension Moho contraction strain pre-extension Moho

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 contraction 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).

Montagne Noire core complex (France) record heat fl ow is consistent with the strong gradient Continental Core Complexes: Detachment paleogeothermal heat fl ow between 150 and 180 that develops during fast extension at the con- Fault Characteristics and Processes mW/m2 (Copard et al., 2000). Similarly, thermo- tact between hanging wall and footwall (Rey chronologic studies of some continental core et al., 2009b) and shows that, although detach- In some continental core complexes, the complexes show that heat may be conducted (or ment faults represent a profound discontinuity detachment is not a single fault but is made advected via fl uids) from footwall to hanging in pressure-temperature-time history of footwall up of multiple, closely spaced and anastomos- wall (e.g., Zeffren et al., 2005), resulting in re- relative to hanging wall, rocks above and being faults (Wernicke and Burchfi el, 1982). The setting (or partial resetting) of isotopic systems low the detachment fault may share a late-stage uppermost detachment fault may have a particu- in hanging-wall rocks and minerals. Such a high thermal history. larly well-defi ned fault plane, typically dipping

278 Geological Society of America Bulletin, March/April 2013 Continental and oceanic core complexes

≤30° (Figs. 1 and 3). A region of brecciation displacement (extension) direction and the undu- of high-grade metamorphic rocks (in this case, and greenschist-facies alteration (recorded by lations have amplitudes of approximately tens in a migmatite dome). The channel detach- secondary growth of chlorite ± epidote) may to hundreds of meters and wavelengths of hun- ment, in contrast, is initially more passive and characterize the structurally highest (brittle) re- dreds of meters to tens of kilometers (John, 1987; accommodates the eastward fl ow of the crustal gions of detachment zones if suitable lithologies Richard et al., 1990; Spencer and Reynolds, channel (Fig. 4C2). The well-documented top- are present (e.g., granitic gneiss). In addition, 1991; Cann et al., 1997; Yin, 2004; Cannat et al., to-the-west shear on the western detachment of many detachment faults record an early history 2009). In oceanic core complexes, these corruga- the Shuswap core complex is consistent with as a ductile shear zone that evolved into a brittle tions have been referred to as “megamullions” the proposed eastward fl ow of underlying crust. fault during tectonic and erosional denudation, (Christie et al., 1998; Tucholke et al., 1998). This defi nition of a channel detachment applied as well as during cooling driven in part by circu- Although some continental core complexes to core complexes is reminiscent of detachments lating fl uids (Malavieille et al., 1990). The upper are bivergent, with symmetric, oppositely dip- that overlie fl owing crust in convergent orogens, brittle fault zone is typically underlain by a re- ping detachments on either side of a footwall such as the South Tibetan detachment (Burg gion of high strain (and temperature) gradient, core (e.g., Hetzel et al., 1995a), many show et al., 1984; Burchfi el and Royden, 1985). up to ~1.5 km thick in the footwall (Wernicke, structural asymmetry in their footwall, particu- In many core complexes worldwide, it is 1981; Miller et al., 1983; Mueller and Snoke, larly in those exhumed at the edge of a continen- common for detachment fault zones (brittle 1993; Wells et al., 2000; Foster and Raza, 2002; tal plateau (Teyssier et al., 2005). For example, fault and footwall shear zone) to be traced in Mulch et al., 2006). in the Shuswap core complex, British Columbia the fi eld from shallow to deep structural levels Although some of the deformation in this re- (Fig. 2B), the continental core complex is asym- (e.g., Snake Range, Nevada; Miller et al., 1983; gion of high strain may have developed during metric, with a series of gneiss domes located in Gébelin et al., 2011; Fig. 5). Detachment zones pre-extensional tectonism, the deformation his- the immediate footwall of the eastern (Columbia have also been imaged in seismic experiments tory related to core-complex development can River) detachment system. The gneiss domes (Smith and Bruhn, 1984; Gans et al., 1985; be discerned through integrated structural and contain the highest-grade metamorphic rocks Cook et al., 1992; Johnson and Loy, 1992), al- isotopic studies, including geochronology. For exposed in the core complex, including high- lowing a view of their geometry to the present example, although some shear zones bounding melt-fraction migmatite (Vanderhaeghe et al., midcrust. In some continental core complexes, Cordilleran core complexes record Mesozoic 1999) and gneiss containing sillimanite and cor- curving detachments may sole into a subhori- deformation, the majority of Cordilleran core dierite pseudomorphs after kyanite. These rocks zontal crustal shear zone at the brittle-ductile complexes show widespread exposures of my- recorded rapid near-isothermal decompression transition at ~10–15 km depth (Wernicke, 1981; lonitized Tertiary rocks that formed during ex- at ~750 °C to 800 °C from 1 GPa to <0.5 GPa Lister et al., 1984), referred to by some workers tension and core-complex development (e.g., (Norlander et al., 2002) during rapid ascent be- as the “master fault.” Other detachments may Foster and Fanning, 1997; Vanderhaeghe et al., neath the eastern, Columbia River detachment. be lithosphere-scale (Wernicke, 1981; Wernicke 1999; Wells et al., 2000; Gébelin et al., 2011). In contrast, the western detachment is relatively and Burchfi el, 1982; Govers and Wortel, 1993). In the northern Cordilleran core complexes, fl at lying and separates laterally extensive leuco- The amount of offset on typical detachment the mylonite zone is several hundred meters granite sheets in the footwall from remnants of faults in continental core complexes is estimated thick (Hyndman, 1980; Mulch et al., 2006; upper-crust klippen in the hanging wall; these at a few to tens of kilometers (Wernicke, 1981; Brown et al., 2012). It affects all lithologies but thin klippen are overlain by supradetachment Axen, 2007; Little et al., 2007). Techniques to has an affi nity for quartzite or marble if these Eocene sedimentary basins, suggesting that the evaluate displacement amount (and extension units are present in the extended crust. Sense of channel detachment initiated and developed in magnitude) include cutoffs of lithologic or shear is typically unambiguous and indicates a fl at position over the time scale of exhuma- structural contacts (Foster et al., 2010), thermo- footwall up, unless the detachment zone has tion of the continental core complex (Teyssier chronology (Foster and John, 1999), and paleo- been tilted by late normal faulting or arched as a et al., 2005). The cumulative shear strain on this magnetism (Livaccari et al., 1995). In most core result of a rolling-hinge or doming effects. detachment may vary considerably along strike complexes, displacement on the detachment Detachment zones are self-exhuming struc- from modest (Glombick et al., 2006) to very sig- cannot account for the total amount of exhuma- tures, and therefore detachment-related my- nifi cant (50–100 km; Brown et al., 2012). tion recorded by footwall rocks (Fayon et al., lonite zones display a range of fabrics, from Given this variation in geometry and de- 2004; Ring et al., 2003; van Hinsbergen, 2010), ductile to brittle. A prevailing view is that high- compression style, two types of detachments indicating that other structures and/or processes temperature mylonitic fabrics are progressively have been proposed for continental core com- are involved in the ascent of footwall rocks from overprinted by lower-temperature ductile fabrics plexes that have experienced deep-crustal fl ow: depth. These processes include fl ow of footwall followed by brittle processes such as cataclastic a “rolling-hinge detachment” (as classically material prior to incorporation into the detach- fl ow and brecciation (e.g., Davis et al., 1980). defi ned: Buck, 1988; Wernicke and Axen, ment system and/or ductile thinning of the crust. Based on the degree of refrigeration of footwall 1988; Brun and van den Driessche, 1994) and Thermochronometers with different clo- during extension, mylonitic rocks may become a “channel detachment” (Teyssier et al., 2005) sure temperatures can be used to evaluate the incorporated into the hanging wall and preserve (Fig. 4C2). These detachments coexist, and cooling rate of rocks in detachment zones a history of deformation fabrics formed at vari- each has a different role in core-complex evolu- (Fig. 6) and to determine the timing of dis- ous temperatures (Mulch et al., 2006). tion. In the Shuswap core complex, according placement on the fault zone, displacement rate, Some core complexes have planar detach- to the Teyssier et al. (2005) model, the rolling- and fault dip (Foster et al., 1990, 1993; John ment faults (e.g., Bitterroot continental core hinge detachment separates the cool foreland to and Foster, 1993; Hetzel et al., 1995b; Foster complexes), but many continental and oceanic the east from the presumed plateau domain and John, 1999; Wells et al., 2000; Ring et al., core complexes are characterized by a corrugated to the west (Whitney et al., 2004a) and there- 2011). Typically, a suite of samples is collected (undulating) detachment fault surface in which fore cuts deeply into the crust, allowing local- along the slip direction of a detachment fault the corrugation axis is oriented parallel to the ized exhumation and signifi cant decompression zone, and one or more thermochronometers

Geological Society of America Bulletin, March/April 2013 279 Whitney et al.

A 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 signifi cant fl uid circulation and hydrothermal al- B 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 interaction in detachment zones is relevant to understanding 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 C D ratios are high during deformation at temperatures suffi cient for (re)crystallization of minerals in the fault zone, and detachment zones are therefore characterized by greenschist-facies and lower-grade minerals: typically, hydrous minerals such as chlorite, white mica, and epidote in continental core complexes, and talc, chlorite, tremolite, and serpentine in oceanic core complexes, as well as metalliferous ore deposits in both settings (Smith et al., 1991; McCaig et al., 2010). Hydrous minerals and other alteration Figure 5. (A) West-dipping, Eocene Okanogan detachment zone, eastern Washington; the products (e.g., from K-metasomatism) that form detachment footwall grades downward from mylonite to migmatite across a 2–3 km sec- in the fault zone may be important in the initiation. (B) East-dipping, Miocene Raft River detachment with remnants of hanging wall on tion and subsequent structural evolution of the mylonitic quartzite; view is looking south from the Idaho-Utah state line. (C) Photomicro- core complex, such as by promoting strain lo- graph of mylonitic quartzite from the Snake Range detachment, Nevada; mica fi sh, S-C calization and/or affecting the thermal state and fabrics, and quartz crystallographic preferred orientation (c-axis, electron backscatter therefore mode and pattern of faulting of the brit- diffraction measurements of >1000 grains) are well developed and indicate top-to-the-east tle crust (Lavier and Buck, 2002). In addition, a shear; stretched quartz grains are partially to entirely recrystallized by subgrain rotation in vigorous hydrothermal system in which mete- the dislocation creep regime. (D) Photomicrograph of mylonitic quartzite from the Colum- oric water circulates through faults in the upper bia River detachment that bounds the Shuswap core complex to the east; mica fi sh and S-C 10–15 km of the extending crust (i.e., to the level fabrics indicate top-to-the-east shear; quartz grains are recrystallized by combination of of the detachment fault) drives effi cient advec- subgrain rotation and grain boundary migration in the dislocation creep regime. 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 (e.g., 40Ar/39Ar in hornblende, biotite, musco- ture of the detachment zone (e.g., position of records. Vein systems and mineralized fault vite, K-feldspar; apatite ± zircon fi ssion-track; isotherms, calculation of geothermal gradients) zones are common in the hanging wall of de- and/or apatite ± zircon U-Th/He) are used to and contribution of erosion to denudation his- tachments and indicate that minerals precipi- capture temperature-time information as a tory (Ketcham, 1996; Fayon et al., 2000). tated from hot fl uids during ascent and cooling. function of exhumation history on the detach- Thermochronology-based estimates of aver- Stable isotope (particularly hydrogen) values of ment. Some detachment zones record a slow age slip rate for detachments bounding Cordi- hydrous minerals such as white mica, biotite, cooling stage (5–15 °C/m.y.) followed by more lleran core complexes range from ~1 mm/yr to chlorite, and epidote show that the water-rich rapid cooling (70–100 °C/m.y.) (Scott et al., 12 mm/yr (Foster et al., 1993; Scott et al., 1998; fl uids that interact with minerals in detach- 1998; Wells et al., 2000). This trend has been Foster and John, 1999; Wells et al., 2000; Carter ment zones are derived from various sources, interpreted to indicate a steepening of the fault et al., 2004, 2006; Foster et al., 2010), and simi- including meteoric sources at structurally high through time or other changes in detachment lar rates have been determined for Aegean core levels and metamorphic/magmatic sources zone geometry. To interpret thermochronol- complexes (John and Howard, 1995; Brichau at lower levels (Kerrich and Hyndman, 1986; ogy data in terms of detachment evolution, care et al., 2006; Thomson et al., 2009). Studies Spencer and Welty, 1986; Kerrich and Rehrig, must be taken to understand the thermal struc- involving multiple thermochronometers are 1987; Wickham and Taylor, 1987; Baker et al.,

280 Geological Society of America Bulletin, March/April 2013 Continental and oceanic core complexes

the temperature of deformation (350–500 °C) 800 Chemehuevi- inferred from quartz-mica stable isotope pairs Sacramento Mts Thor-Odin Valhalla (Mulch et al., 2006; Gottardi et al., 2011; Gébelin et al., 2011), calculated fl ow stress is 700 Bitterroot Naxos Kazdağ typically high and quite constant over the en- granite tire mylonitic section. This behavior suggests 600 East Bitterroot that localization or delocalization of strain, cor- 1 detachment responding to increasing or decreasing strain 2 rates, respectively, is a response of the system 500 to maintain fl ow stress near the critical crustal strength, a property that is best exemplifi ed in Raft River 400 extensional systems (Mulch et al., 2006). There has been much discussion of the dy- 10°C/m.y. namics of low-angle normal faults during core- 300 complex development. Debate stems from predictions from rock mechanics theory that

Temperature (°C) Temperature CMM 100°C/m.y. extension produces high-angle faults (~60°; 200 Ander son, 1942, 1951) and that low-angle faults cannot slip, consistent with the proposal ≥ 100 that moderate to large earthquakes (M 5.5; Jackson, 1987; Jackson and White, 1989) do not occur along these faults. This conclusion EastWest apparently confl icts with fi eld observations 0 20406080that low-angle, and in some cases subhorizon- tal, normal faults have been active in the brittle Time (Ma) crust (Wernicke , 1981; Reynolds and Spencer, 1985; Davis et al., 1986; John, 1987; Wernicke Figure 6. Compilation of temperature-time data for continental core complexes in the North and Axen, 1988; Scott and Lister, 1992; John American Cordillera and the Aegean/Anatolia region. Sources: Naxos/Aegean: 1—migma- and Foster, 1993; Lecomte et al., 2010); for tite core, 2—schist carapace (Gautier et al., 1993); Kazdağ/Anatolia: Cavazza et al. (2009); example, the recognition that horizontal fi eld Central Menderes Massif (CMM)/Anatolia: Ring et al. (2003); Chemehuevi/Sacramento markers are close to the detachment surface Mountains/Cordillera (south): John and Foster (1993), Foster and John (1999), Campbell- (e.g., Scott and Lister, 1992; John and Foster, Stone et al. (2000); Raft River/Cordillera (central): Wells et al. (2000); Bitterroot/Cordillera 1993), and therefore that displacement on the (north): Foster and Raza (2002); Valhalla/Cordillera (north): Gordon et al. (2008, and refer- detachment fault occurred at low dips. In addi- ences therein); Thor-Odin/Cordillera (north): Vanderhaeghe et al. (2003). tion, some seismically active, low-angle (<30°) normal faults have been inferred based on the choice of low-dip fault plane solutions (Abers, 1989; Nesbitt and Muehlenbachs, 1995; Losh, 2004). Studies that combine data from detach- 1991; Abers et al., 1997; Axen, 1999; Famin and 1997; Holk and Taylor, 2007; Mulch et al., ment-basin pairs have verifi ed that the isotopic Nakashima, 2005), indicating that some detach- 2004; Gébelin et al., 2011). Studies of detach- composition of mylonite minerals matches that ments may slip at low angles. ments that are localized in simple lithologies measured in basin strata deposited at high eleva- To account for this discrepancy, it has been (quartzite, marble) have shown that, in general, tion (Mulch and Chamberlain, 2007). proposed that some low-angle normal faults ini- meteoric water equilibrates during detachment Detachment zones typically accommodate tiate at a steep dip and subsequently rotate to- activity with neo/recrystallized hydrous phases large lateral displacement and considerable thin- ward horizontal via rotation of successive fault such as white mica at moderate temperatures ning. Tens of kilometers of displacement are ac- blocks or by isostatic adjustment (i.e., the roll- (~350–450 °C) (Famin et al., 2004). commodated by deformation in the detachment ing-hinge model; Fig. 4B; Buck, 1988; Ham- Hydrogen isotopic values in synkinematic zone over the time scale of 1–5 m.y., implying ilton, 1988; Wernicke and Axen, 1988; Axen white mica (e.g., Figs. 5C and 5D) and other high average strain rates. Quartz microstruc- et al., 1995; Lavier et al., 1999). Alternatively, hydrous minerals in detachment zones have ture in detachment zones provides information it has been proposed that conclusions based on been interpreted to indicate interaction of mica on fl ow paleostress through the analysis of re- seismic data sets may have missed earthquakes with a fl uid derived from a high-elevation catch- crystallized grain size using paleopiezometers; with long recurrence intervals, seismicity on ment (e.g., Columbia River fault at the latitude temperature of deformation can be derived detachment faults may more typically involve of Thor-Odin, British Columbia, and at the lati- from quartz-mica oxygen isotope thermometry small-magnitude events, and/or models based tude of the Kettle dome, Washington; detach- or from titanium-in-quartz thermometry. In the on fault mechanics theory may involve invalid ments in the Ruby Range and Snake Range, quartzite mylonites of some core complexes in assumptions about rock strength or stress orien- Nevada; Fricke and O’Neil, 1999; Mulch et al., the North American Cordillera, quartz dynamic tations along detachment faults (Axen, 2007). 2007; Gébelin et al., 2011). Hydrous minerals recrystallization is dominated by subgrain ro- The origin and evolution of detachment faults, in detachment mylonites may therefore contain tation, which is consistent with low recovery with specifi c reference to fault angle, are dis- the paleoelevation record over the time scale under high-fl ow stress conditions in the disloca- cussed more in a later section (Mechanics of (~1–5 m.y.) of mylonite formation (Mulch et al., tion creep regime (Figs. 5C and 5D). Whatever Core-Complex Faults).

Geological Society of America Bulletin, March/April 2013 281 Whitney et al.

Continental Core Complexes: Footwall Some domes have a double-dome pattern con- for example, zircons in gneiss yield pre-Cenozoic Characteristics and Processes sisting of two main compartments defi ned by ages, recording protolith crystallization and/or foliation divided by a steep, median high-strain later metamorphism associated with pre-exten- Footwall rocks may record a wide range of zone (Fig. 4C1); examples include the Naxos sion crustal thickening (e.g., Kruckenberg et al., ages—from pre- to synextension—of meta- (Greece) and Montagne Noire (France) core 2008). The age of extension and accompanying morphic, magmatic, and deformation events. complexes (Rey et al., 2011). Other domes exhumation of the footwall rocks is indicated by Metamorphic grade may also vary within the show nappe-like recumbent folds that overprint the youngest U-Pb zircon and monazite ages, footwall, not only as a function of structural earlier, steeper structures (e.g., McFadden et al., and is further bracketed by cooling ages deter- level exposed, but also owing to the complex 2010). These double-dome and more complex mined by thermochronometers such as 40Ar/39Ar pressure-temperature (P-T) paths that rocks fol- three-dimensional structures are signifi cant for for hornblende, micas, and K-feldspar (Fig. 6). low before and during extension. It is important understanding crustal fl ow under extension, and In migmatite-cored core complexes, melt that to determine the thermal state of the lithosphere they provide a framework for interpreting defor- collected in boudin necks or extensional shear during extension because the ambient synexten- mation features as a function of time and space. zones may provide an indication of the timing sion geothermal gradient is an important factor Numerical modeling provides insights to help of onset and/or cessation of major extension in controlling the evolution of core complexes, understand the relative infl uence of geothermal (Gordon et al., 2008; McFadden et al., 2010). as well as the role of crustal fl ow in creating and gradient, crustal thickness, and crustal fl ow in Even in cases in which the extension-related maintaining a fl at Moho, as is observed in many core-complex initiation and evolution (e.g., temperature-time path of footwall rocks is highly extended regions (Block and Royden, Tirel et al., 2008; Rey et al., 2009a, 2009b). known, it can nevertheless be challenging to 1990; Buck, 1991; Rey, 1993) (Fig. 4). For example, Tirel et al. (2008) showed that for understand the pressure (depth) evolution of The footwall of most core complexes has the model assumptions and parameters used, a footwall rocks, and in particular it is diffi cult to a domal structure (Fig. 3). The domal geom- Moho T > 800 °C is required to produce a core determine the maximum pressure experienced etry has been explained for some continen- complex in 60-km-thick crust. At T < 800 °C, a by rocks exhumed in a core complex; however, tal core complexes as resulting from uplift of strong upper mantle contributes to crustal-scale this is important information for reconstructing the detachment during isostatic rebound as the boudinage (necking). Insights from numerical pre-extension crustal thickness and unravel- hanging-wall rocks are extended, thinned, and modeling relevant to the fl ow of deep crust, as ing particle paths in footwall rocks. Some core denuded by tectonic and erosional processes well as the origin of double-domes, are dis- complexes exhume rocks from great depths, (Spencer, 1984; Buck, 1988). These models as- cussed in a later section (Dynamic Models of e.g., high-P rocks such as eclogite or kyanite- sumed that the lower-crust fl owed at fast rates, Continental Core-Complex Development). bearing gneiss, although much of the core- whereas later studies considered how weak the Although some core complexes are associ- complex footwall may consist of metamorphic lower crust must be to allow dome formation ated with regions of low heat fl ow (e.g., south- rocks recording only a low-P, high-T history (Block and Royden, 1990; Kruse et al., 1991; ern California and Arizona; Lachenbruch et al., (Rey et al., 2009a, 2011). In some cases, high-P Wdowinski and Axen, 1992; McKenzie et al., 1994), others are associated with regions of high rocks exhumed in a core complex record a much 2000; Rey et al., 2009a, 2009b, 2011). heat fl ow (e.g., the islands of Naxos and Paros, older (pre-extension) metamorphic history (e.g., Evidence for fl ow of weak crust can be seen in in the central Aegean; Gautier et al., 1993; Keay Precambrian eclogite in the Cenozoic Menderes the complex internal structure of gneiss (or mig- et al., 2001; Brichau et al., 2006; Seward et al., core complex, western Turkey; Candan et al., matite) domes, which occur within some conti- 2009). These Aegean islands contain cores of 2001), although it is likely that exhumation of nental core complexes. Some core complexes syntectonic migmatite and granite and have the high-P rocks occurred during extension, and contain one or more gneiss domes, typically been proposed as the site of a thermal anomaly. therefore the presence of these rocks is relevant beneath a carapace of high- to medium-grade However, these core complexes do not neces- to understanding core-complex dynamics. It metamorphic rocks (Brun and van den Driessche, sarily represent an anomalously hot region of is therefore particularly important to know the 1994; Vanderhaeghe and Teyssier, 2001; Whitney the crust. Instead, they may have formed in a P-T-time history of these rocks so as to be able et al., 2004b); these have been called migmatite- zone of large-scale extension or transtension to track the timing, magnitude, and paths of cored metamorphic core complexes (Rey et al., that triggered the ascent of hot, ductile crust that their exhumation. 2009a, 2009b), and their origin relates to regional fl owed from deep to shallow levels and became Some conceptual and numerical models for extension and fl ow of deep crust beneath detach- involved in the high-strain zone beneath the de- core-complex development assume that the ment faults. Studies have shown that the crystal- tachment, where isotherms collapsed (Krucken- major motion of footwall rocks is horizontal, lization of the magmatic portions of migmatite berg et al., 2011; Rey et al., 2011). According to except for some vertical motion related to arch- domes in core complexes was followed by rapid this idea, hot ductile crust was present region- ing of the footwall beneath the detachment (Brun cooling to T < 300 °C (Fig. 6); cooling ages co- ally, but only locally exhumed. and van den Driessche, 1994). Based on this as- incide with ages of synkinematic minerals (e.g., Interpretation of the P-T conditions and paths sumption, predictions are made about expected mica) in detachment fault zones (Malavieille of metamorphism relevant to core-complex de- changes (or lack of changes) in metamorphic et al., 1990; Maluski et al., 1991; Kruckenberg velopment (e.g., Buick and Holland, 1989) re- grade in footwall rocks in the direction of tec- et al., 2008). These results show that high-tem- quires knowledge of the age of metamorphic tonic transport for core complexes with different perature crustal fl ow occurred during core-com- events that affected footwall rocks. As noted, thermal histories (Gessner et al., 2007). This as- plex formation but ended when hot rocks reached footwall rocks may have experienced multiple sumption is valid for some core complexes but shallow crustal levels and cooled rapidly. metamorphic events prior to metamorphism not for those in which rocks registered a major Gneiss (migmatite) domes commonly dis- related to extension and core-complex develop- component of vertical motion, as shown by a play a relatively simple external surface but a ment, so care must be taken (particularly with record of isothermal decompression, e.g., from complex internal structure (e.g., Kruckenberg zircon) to determine the age of the last metamor- >20–30 km to <10 km depth (common in core et al., 2011) that varies with level of exposure. phic event. In some Cenozoic core complexes, complexes cored by migmatite). In many conti-

282 Geological Society of America Bulletin, March/April 2013 Continental and oceanic core complexes

nental core complexes, this decompression has Continental Core Complexes: Melting complexes in general: (1) Oceanic core com- been interpreted as driven by extension (Soto and Magmatism plexes form from “new” material rising from and Platt, 1999; Viruete et al., 2000; Norlander the asthenosphere and therefore have no com- et al., 2002). Additional insights about pressure There has been much debate for both conti- plicating deformation history that predates core- history come from observations of reaction tex- nental core complexes and oceanic core com- complex formation, and (2) they are subject to tures in footwall metamorphic rocks (Krucken- plexes regarding the importance of magmatism little erosion, which can obscure structures and berg and Whitney, 2011; Goergen and Whitney, in core-complex evolution (some of the relevant fault-generated topography on continents. 2012) and from numerical modeling (Ruppel evidence and questions are summarized in the In recent years, many oceanic core com- et al., 1988; Rey et al., 2009a, 2009b). previous sections). An important question for plexes have been identifi ed along the Mid- Temperature-time paths for footwall rocks both continental core complexes and oceanic Atlantic Ridge, the Southwest Indian Ridge, during extension-related exhumation can be core complexes is whether magmatism facili- the Caribbean–North American Ridge, the deter mined by application of an array of thermo- tates core-complex formation or whether mag- Australian-Antarctic Discordance, and along chronometers with different closure tempera- matism is a result of the extension that also parts of backarc spreading centers (e.g., the tures in different minerals (cf. similar methods drives core-complex formation (or both) (Gans Godzilla Megamullion, Philippine Sea; Ohara applied to detachment zone rocks; Fig. 6). In- et al., 1989; Armstrong and Ward, 1991; Lister et al., 2001) (Fig. 2A). These oceanic core com- terpretation of these data in the context of the and Baldwin, 1993; Spencer et al., 1995; Foster plexes may comprise 50%–60% of some ridge exhumation path (i.e., changes in depth accom- et al., 2001; Buck et al., 2005; Tucholke et al., segments of the Mid-Atlantic Ridge (Smith et panying cooling) requires estimates of the geo- 2008; Olive et al., 2010). A likely mechanism for al., 2006, 2008; Escartín et al., 2008) and 70% thermal gradient during extension. Similar to weakening and strain localization involves par- of part of the Australian-Antarctic Discordance the T-time history of some detachment zones, tial melting, intrusion, and release of volatiles (Okino et al., 2004). They are generally recog- footwall rocks (below the detachment zone) associated with magmatism. These processes nized by their domal morphology and promi- may record rapid cooling (>50 °C/m.y.) fol- promote thermal weakening of the crust, lead- nent corrugated fault surfaces (Fig. 3B). lowed by a more protracted cooling history ing to orogenic collapse and strain localization, Major questions about the dynamics of (<20 °C/m.y.; Fig. 6). and can result in rotation of the orientation of oceanic core complexes concern the extent to In the northern Cordilleran core complexes the principal stresses, facilitating slip on low- which plate separation at slow-spreading ridges (Thor-Odin, Valhalla, Okanogan, Kettle, Bitter- angle normal faults (Parsons and Thompson, occurs via detachment-style faulting and core- root; Fig. 2B), U-Pb ages of zircon in leuco- 1993). Debate has nevertheless persisted in part complex formation, the degree of rotation of granite and/or migmatite (Foster and Fanning, because large-scale extension is typically ac- oceanic detachments, the role of magmatism in 1997; Vanderhaeghe et al., 1999) are within companied by magmatism, but magmatism can the development of oceanic core complexes, and 5–10 m.y. of the age of cooling through low-T precede (Foster and Fanning, 1997), coincide the importance of fl uid-rock interaction associ- thermo chronometers (40Ar/39Ar, apatite fi ssion with (Hill et al., 1995), and/or postdate (Simp- ated with detachment tectonics. To assist with track, or apatite and zircon He ages; Lorencak son et al., 1991) major extension. In addition, discussion of these questions, we summarize et al., 2001; Foster and Raza, 2002; Vander- there may be feedbacks between extension and the main features of hanging-wall, detachment, haeghe et al., 2003) (Fig. 6). In the Aegean core magmatism, as decompression driven by exten- and footwall regions of oceanic core complexes, complexes, a similar cooling path is recorded sion and crustal thinning may result in decom- with an emphasis on hydrothermal, magmatic, (Fig. 6). For example, the gneiss dome in the pression melting, further facilitating extension and rheological processes. footwall of the Naxos detachment crystallized (Teyssier and Whitney, 2002). zircon at 17 Ma (Keay et al., 2001; Martin et al., Coupled thermal-mechanical numerical mod- Oceanic Core Complexes: Hanging-Wall 2006), only a few million years prior to cooling eling can help clarify this issue (see section Characteristics and Processes through low-T thermochronometers (Brichau Dynamic Models of Continental Core-Complex et al., 2006; Seward et al., 2009). Development). In models in which kinematic In many oceanic core complexes, the corru- Near-isothermal decompression followed by boundary conditions drive extension, a core gated fault zone is observed at the seafl oor, and rapid cooling are characteristic of migmatite- complex develops whether or not melt is pres- there is little remaining hanging wall. Typically, cored complexes in which hot rocks rapidly ent in the system (in models, “melt” means a the active part of the detachment dips toward the ascended from deep to shallow crustal levels. lower-viscosity and lower-density material; Rey spreading axis, so the hanging wall (if present) This nonlinear cooling history is expected from et al., 2009b). In this case, isostasy is the driver is a block between the fault and the spreading the steady exhumation of deep crust owing to of core-complex formation. In contrast, when axis (Figs. 1B and 3). In places, smaller blocks the relative rates of material and heat advection extension is driven by volume forces alone, the of hanging wall, including volcanic infi ll, sit on (Batt and Braun, 1997). Therefore, the onset of buoyancy of the melted region facilitates the de- top of and ride along with the footwall. These rapid cooling in migmatite-cored complexes oc- velopment of a core complex. rafted or rider blocks have been recognized by curs late in the exhumation history, when hot bathymetric and seismic studies (Dannowski rocks are so close to the surface that heat ad- OCEANIC CORE COMPLEXES et al., 2010; Reston and Ranero, 2011) and vection is no longer possible and rocks cool at are similar to features seen in continental core rates on the order of 100 °C/m.y. Another impli- Owing to the logistical complexities of fi eld- complexes. cation of this exhumation-driven heat advection based studies in the oceans, there is compara- The possibility of hanging wall occurring lat- is that core complexes generate short-lived but tively less information about detachments in erally along the axis away from corrugated core very hot transient geotherms; in turn, the high oceanic core complexes as compared to conti- complexes is an open question. One possibility geotherm drives convective fl uid fl ow and asso- nental core complexes. Nevertheless, oceanic is that the detachment is replaced along-axis by ciated refrigeration of the detachment footwall core complexes offer some advantages for a series of high-angle, lithosphere-cutting faults (e.g., Gottardi et al., 2011). under standing the origin and evolution of core (Cannat et al., 2006; MacLeod et al., 2011).

Geological Society of America Bulletin, March/April 2013 283 Whitney et al.

Another possibility is that the detachment con- to thermal cracking. In shallower and cooler Studies of slow-spreading centers show that tinues along-axis, but it is covered by extrusive domains, mantle peridotite undergoes extensive hydration of mantle at mid-ocean ridges is in- rocks and rider blocks (Escartín et al., 2008; hydration reactions at temperatures <500 °C, re- timately related to deformation in the detach- Reston and Ranero, 2011). sulting in schist and cataclasite containing talc, ment zones that bound oceanic core complexes. In some regions of the Mid-Atlantic Ridge, actinolite-tremolite, chlorite, and/or serpentine Ocean drilling at MARK (Kane zone of Mid- the hanging wall is affected by intense hydro- (Schroeder and John, 2004; Boschi et al., 2006; Atlantic Ridge) has recovered samples of highly thermal circulation (Fig. 1B); this has led to the Dick et al., 2008; McCaig et al., 2010). serpentinized peridotite in which the high-tem- suggestion that hydrothermal activity at slow- In oceanic core complexes, a similar idea has perature foliation (30° dip) is overprinted by spreading systems relates to the dynamics of been proposed to that of continental core com- several generations of serpentine veins (Cannat the underlying detachment fault and does not plexes regarding the rooting, or termination at et al., 1995; Andreani et al., 2007). Serpentini- directly relate to magmatism (Petersen et al., depth, of the detachment fault: i.e., detachment zation is an important process in detachment 2009). However, it has also been proposed that faults terminate at a rheological transition, such development owing to the exothermic nature of the detachment serves as a conduit for fl uids and as the brittle-ductile transition at the diabase the reaction and the volume increase and signifi - links the fractured hanging wall and associated dike–gabbro boundary in oceanic crust or in cant mechanical weakening associated with the hydrothermal fi elds at the ocean fl oor to magma or near a melt-rich zone or magma chamber olivine to serpentine transformation. In addition, chambers at depths of ~5–10 km (e.g., Escartín (Tucholke et al., 1998; Dick et al., 2000) (Fig. geophysical results suggest that detachment sys- et al., 2003; McCaig et al., 2007, 2010; McCaig 1B). Escartín et al. (2003) proposed a classifi ca- tems allow seawater to circulate to 7–8 km depth and Harris, 2012). tion of oceanic core-complex detachment sys- and to tap heat from gabbro bodies (Escartín Intrusion and extrusion of basalt clearly add tems in which “hot” detachments terminate in et al., 2003; Canales et al., 2008). Therefore, the to the hanging wall. Extrusion is important in or near magma pathways or magma chambers, potential exists for seawater interaction to occur the formation of hanging-wall rider blocks whereas “cold” detachments terminate in a zone at ~1000 °C with gabbro magma and its crystal- (Reston and Ranero, 2011; Choi and Buck, of alteration of gabbro or peridotite (e.g., where lized equivalent as the gabbro is exhumed along 2012), and the formation of oceanic detach- amphibole, talc, serpentine, or other hydrous the detachment zone. ments may depend on the amount of dike intru- minerals are forming). Fault rocks retrieved from active and ancient sion into the hanging wall (Buck et al., 2005). The deployment of ocean-bottom seismom- oceanic core complexes show extreme hydration This topic is discussed more in the section eters has allowed monitoring of microseismicity and fl uid interaction at low temperature (serpen- “Dynamic Models of Oceanic Core-Complex in oceanic core-complex regions to help under- tinized peridotite) but also at high temperature Development.” stand detachment geometry and mechanics at (peridotite mylonite). Some peridotite mylonite depth. For example, in the hydrothermal fi eld contains fl uid inclusions and high-T (~800 °C) Oceanic Core Complexes: Detachment associated with the Trans-Atlantic Geotraverse hydrous phases that likely formed at the ridge Fault Characteristics and Processes (TAG), microseismicity defi nes a fairly planar in the root zones of the detachment system swarm of earthquakes that dips steeply (70°), (e.g., kaersutite mylonite; Campos et al., 2010; Geological sampling of oceanic detachments extends to ~7 km depth below the seafl oor, and Kurz et al., 2009). Although lower-T (<600 °C) by coring and dredging indicates that the zone projects upward to the locus of hydrothermal processes characterize the spectacular hydro- of brittle deformation in some core complexes activity (Canales et al., 2007; de Martin et al., thermal fi elds associated with mid-ocean ridges may be <200 m thick (Escartín et al., 2003); 2007). The seafl oor in the immediate footwall (serpentinization, generation of vent fl uids, min- this is underlain by a broader zone of ductile of this steeply dipping fault exposes a detach- eralization), the presence of high-T mylonite deformation. In the Atlantis Bank oceanic core ment fault that can be traced over ~5 km along on the ocean fl oor suggests that hydration of complex (SW Indian Ridge; Figs. 2A and 3B), the gentle slope of the oceanic core complex, mantle may reach deep and hot mantle at slow- an upper zone of brittle deformation (30–80 m suggesting that the detachment was exhumed and ultraslow-spreading centers. This may have thick) overlies a broader zone of ductile fabrics by a rolling-hinge process, with the active steep important implications for the trapping of water (to a current depth of 400 m below the fault sur- normal fault representing the latest structure that in the deep oceanic lithosphere, and ultimately face; Miranda and John, 2010). In the Atlantis is accommodating footwall exhumation. The for the budget of water in the Earth. Massif oceanic core complex (Mid-Atlantic active fault also controls the location of hydro- Many oceanic core complexes are located in Ridge; Figs. 2A and 3), an ~500-m-thick zone of thermal activity and allows circulation of fl uids the inside corners of ridge-transform intersec- ductile deformation in gabbro is overprinted by to considerable depth (~7 km), resulting in sig- tions (e.g., Atlantis Bank; Baines et al., 2003), cataclastic fabrics (Schroeder and John, 2004). nifi cant heat extraction. In this hydrothermal raising the question of the role of transform Peridotite mylonites recovered in the vicin- system, heat budget calculations suggest that faulting in oceanic core-complex initiation and ity of some oceanic core complexes (Dick et al., magmatic heat input is necessary and likely ac- development. The spatial association of detach- 2010) or dispersed along transform faults after counts for three quarters of the heat dissipated at ments with transform faults may refl ect the forming in oceanic core complexes (Cipriani the ocean fl oor (Canales et al., 2007). An impor- fundamental dynamics of melt migration along et al., 2009) are typically unaltered compared tant implication of this work is that detachments ridge segments (Cannat et al., 1995; Tucholke to their nonmylonitic counterparts. Limited ser- that root in a hot source (such as a magma body) et al., 2008). As melt migrates toward the cen- pentinization of peridotite mylonite indicates develop an effi cient refrigeration system that ter of a segment, tectonic rather than magmatic that mylonitization occurred at temperatures keeps the detachment zone relatively cool while processes accommodate extension at the ridge- higher than that of serpentine stability; preser- a remarkable amount of heat is rapidly trans- transform corners (Tolstoy et al., 1993; Detrick vation of high-T mylonitic fabrics (dynamically ferred to the ocean fl oor by convecting seawater- et al., 1995; Karson, 1998; Planert et al., 2009), recrystallized olivine) may be related to the ex- derived fl uids. This is analogous to the transient, and therefore oceanic core complexes are more treme reduction in grain size (three orders of detachment-related hydrothermal systems that likely to develop at these corners than in the magnitude) that makes mylonite more resistant develop above continental core complexes. ridge segment centers.

284 Geological Society of America Bulletin, March/April 2013 Continental and oceanic core complexes

Oceanic core-complex detachments gener- be locally mylonitic to mildly deformed, and tion (Morris et al., 2009). These high rotation ally display a domal structure as well as shorter- some also contain diabase dikes (Escartín et al., values about a nearly horizontal axis are consis- wavelength corrugations with axes parallel 2003; Canales et al., 2008; Dick et al., 2008). tent with a rolling-hinge detachment model and to the extension direction; these corrugations Gravity and seismic characteristics of oceanic suggest that the oceanic core-complex footwall have amplitudes up to several hundred meters core complexes have led to various interpreta- rotates quite rigidly during exhumation, raising (Cann et al., 1997; Tucholke et al., 1998, 2008; tions of crustal structure, including the position the question of the nature of the material. Ranero and Reston, 1999; Cannat et al., 2006). of the Moho, the degree of serpentinization of In continental core complexes, rolling-hinge The longer wavelengths, which generate domal the crust and mantle, and the location of magma motion is made possible by an isostatic compen- oceanic core complexes, may be related to mag- bodies and other heterogeneities relative to sation of low-viscosity (lower-crustal) material matic versus tectonic extension, and the shorter detachments. that fl ows to fi ll the gap created by rotation of wavelength (corrugations) may be related to Seismic-refraction and wide-angle experi- the detachment footwall. In oceanic core com- variations in both space and time of melt bodies ments (Canales et al., 2007; Planert et al., 2009) plexes, candidates to solve the space problem that feed the spreading axis (Lin et al., 1990) and as well as other geophysical investigations (e.g., are (1) the addition of magma though successive that ultimately control the rheology of detach- Mallows and Searle, 2012) on segments of the gabbro bodies that pool at the root of the de- ment root zones. Mid-Atlantic Ridge have uncovered fi rst-order tachment before they are exhumed in the detach- Fault rocks associated with detachment relations in crustal structure. For example, south ment footwall (Olive et al., 2010; Grimes et al., faults in some oceanic core complexes record of the Kane fracture zone, the imaged Moho 2011), (2) the fl ow of ductile mantle (cf. fl ow similar temperatures to those in continental faithfully follows the oceanic core-complex of deep crust toward the detachment footwall in core complexes: ~550–300 °C (amphibolite domal structure, showing that the crust beneath continental core complexes), or (3) a combina- to greenschist facies; Karson, 1999); in others the detachment is thin. Relief on the Moho over tion of solid-state fl ow and magmatic addition. (Atlantis Bank, Southwest Indian Ridge), 20 km across the oceanic core complex in that The retrieval of foliated and mylonitic gabbro defor ma tion conditions ranged from granulite region is as much as 4 km (Dannowski et al., and peridotite from oceanic core complexes to greenschist facies (from >900 °C to <300 °C; 2010), and the Moho also displays high relief suggests that a combination of ductile fl ow and Schroeder and John, 2004; Miranda and John, in the direction parallel to the ridge axis. Simi- magma intrusion likely compensates for the sig- 2010). Thermochronology studies show that lar relief of ~4 km on the Moho is present at nifi cant rotation of detachment footwalls. footwall and detachment zone rocks in oceanic the Mid-Atlantic Ridge region at ~5°S, both on core complexes experience rapid cooling at a ridge-parallel and ridge-perpendicular sections Oceanic Core Complexes: Melting rate that is an order of magnitude faster than (Planert et al., 2009). This three-dimensional and Magmatism in continental core complexes. For example, in variability in crustal thickness may impose the Atlantis Bank oceanic core complex, foot- particular patterns of fl ow in the mantle that Several sets of observations suggest that mag- wall gabbro cooled from ~850 °C to <350 °C underlies oceanic core complexes, and perhaps matism plays an important role in the formation at ~800 °C/m.y. (John et al., 2004). At several even links to the processes of melt generation of oceanic core complexes. In a remarkable oceanic core-complex sites along the Mid- and extraction that make oceanic crust at slow- study of bathymetry, magnetism, and gravity Atlantic Ridge, the difference in ages between spreading centers. of 39 oceanic core complexes in a >2000 km2 U-Pb (high closure temperature) and (U-Th)/He The rate at which footwall exhumation oc- region of the eastern Southwest Indian Ridge, zircon ages (low closure temperature) yields curs is largely known from average plate mo- Cannat et al. (2006) identifi ed different modes cooling rates of ~1000–2000 °C/m.y. over the tion, but the motion of footwall rocks during of spreading and demonstrated systematic rela- ~650 °C interval between thermochronometers detachment tectonics depends to a large extent tionships between these modes and the appar- (Grimes et al., 2011). on the geometry of the detachment fault or shear ent magma budget along parts of the spreading As in continental core complexes, there ap- zone. In oceanic core complexes, detachment center. Three distinct morphologies were pears to be a diversity of detachment zone styles systems tend to be convex upward and to reg- recognized: (1) volcanic seafl oor, defi ned by that exhume oceanic core complexes. However, ister seismicity consistent with the active steep numerous scarps and volcanic cones, (2) cor- observations converge toward a fairly general fault reaching signifi cant depth (7–8 km) below rugated seafl oor, and (3) smooth seafl oor with model in which oceanic core-complex detach- the seafl oor and possibly rooting in a magma no vol canic feature or corrugations (Fig. 7). A ments are convex upward, evolve according to chamber (Fig. 1B). The active fault is typically systematic relationship is seen between modes a rolling-hinge mechanism (Buck et al., 2005; connected to the domes that are bounded by the of accretion and the gravity and topography Smith et al., 2008), and rapidly exhume their corrugated detachment fault and defi ne the core signatures for conjugate pairs of Southwest high-T fabrics and footwall. Detachments pro- complex. Indian Ridge seafl oor crust (Cannat et al., vide a pathway between the ocean and the deep Paleomagnetism has been used to evalu- 2006). Where the residual gravity is low and the roots of oceanic core complexes, with implica- ate the rotation history that footwall rocks ex- topography is high, the crust is thick and cor- tions for mineral hydration/reactions and advec- perienced during their exhumation and to test responds to volcanic regions with high rates of tive cooling. whether a rolling hinge may accommodate magma input. Corrugated zones correspond to footwall exhumation. At the TAG hydrother- intermediate rates of magmatism, and smooth Oceanic Core Complexes: Footwall mal fi eld (Mid-Atlantic Ridge), a rotation of seafl oor forms in regions with the lowest levels Characteristics and Processes ~50° has been documented based on the re- of inferred magma supply. construction of magnetic vectors (Garcés and A study of global occurrences of corrugated Geologic sampling and seismic data indi- Gee, 2007). Using a technique involving cor- seafl oor also shows that depth of the seafl oor cate that oceanic core-complex footwalls are relation between core structures and borehole may correlate with crustal thickness and there- composed primarily of gabbroic and ultramafi c wall imagery, paleomagnetic data of cores from fore magma supply (Tucholke et al., 2008). Oce- rocks, many of which are altered. These may the Atlantis Massif indicate 46° ± 6° of rota- anic core complexes occur where the adjacent

Geological Society of America Bulletin, March/April 2013 285 Whitney et al.

ration rate is not accommodated by magma- In the rest of this section, we fi rst use model A VOLCANIC-VOLCANIC SEAFLOOR PAIR volcanic seafloor tism in the form of dike opening, as suggested results that focus on the brittle layer to explore by Schouten et al., 2010). However, any rate- the mechanics that lead to the development of axial valley dependent weakening (or healing, as in Buck detachment faults. We then examine more com- 0 et al., 2005) leads to only one fault being active plete lithospheric models in which the coupling crust dike at a given time. The inference that low values of rheologically realistic layers is investigated 5 of magma supply result in a complex pattern of (brittle, temperature-dependent viscous, par- mantle moderate-offset, crosscutting faults argues for tially molten) for the case of continental core lithosphere moho there being one dominantly active fault at a time complexes, followed by discussion of the dy- ~20 km in the axial region (Cannat et al., 2006). namics of oceanic core complexes. Regard- However, as little of the seafl oor is mapped ing the broader geodynamic settings in which with suffi cient resolution to detect oceanic core continental core complexes form, we then ad- B CORRUGATED-VOLCANIC SEAFLOOR PAIR complexes, it is likely that they are more com- dress two cases: (1) the role of mantle wedge axial mon than the Tucholke et al. (2008) study in- dynamics in Cordilleran-type orogens, where valley volcanic layer dicates. For example, the correlation of oceanic continental core complexes develop in the con-

crust core complexes with higher than normal rates of tinental overlying plate; and (2) the transition corrugated 0 seafloor seismicity and hydrothermal activity may indi- between an orogenic plateau and its foreland. mantle 5 cate that ~50% of a long (75–100 km) section lithosphere of the Mid-Atlantic Ridge could involve detach- Mechanics of Core-Complex Faults moho ments (Smith et al., 2006; Escartín et al., 2008). Mohr-Coulomb fracture mechanics imply ~20 km CORE COMPLEXES AND that faulting occurs most easily at an angle of C SMOOTH-SMOOTH SEAFLOOR PAIR LITHOSPHERE DYNAMICS ~30° to the maximum principal stress (e.g., Anderson, 1942). Assuming an “Andersonian” Core complexes form in regions of extension extensional stress fi eld in which the minimum 0 driven by surface forces at plate boundaries, or principal stress is horizontal, normal faults in crust 5 volume forces in relation to lateral variation of the brittle upper crust should initiate at dips gravitational potential energy, or both. World- ~60° (Anderson, 1942) and should be active moho wide, most core complexes form in regions of at dips of no less than 30° (e.g., Sibson, 1985). shear zones extension in collapsed orogens and at relatively Several authors have suggested that normal mantle lithosphere ~20 slow-spreading mid-ocean-ridge systems. Oro- faults could initiate with low dips if particular melt km genic collapse may occur under free boundary loads reorient the tectonic stress fi eld (Spencer (after Cannat et al., 2006) conditions, such as driven by slab rollback at and Chase, 1989; Yin, 1989; Parsons and convergent margins, or under a fi xed bound- Thompson, 1993). However, Wills and Buck Figure 7. Sketches of axial regions for three ary, owing to spreading of thick, weak crust (1997) showed that even with these purpose- proposed modes of slow to ultraslow spread- (Rey et al., 2001). In these settings, extension built stress fi elds, low-angle faults would not ing, shown in order of inferred decreas- is lithosphere scale, but the major expression of form before high-angle faults. ing melt supply: (A) volcanic seafl oor and extension may be in the crust because normal An alternative to slip on low-angle normal (B) corrugated seafl oor occur at both slow- faulting of the brittle upper crust is coupled with faults is that the upper parts of some actively and ultraslow-spreading ridges, whereas ascent of the ductile crust, resulting in regions slipping high-angle normal faults rotate to shal- (C) smooth seafl oor occurs only at ultraslow- of extreme thermal and strain gradients across lower dips. Spencer (1984) suggested that the spreading ridges. Horizontal dimensions are detachment zones (meters to ~1.5 km thick). isostatic response to offset of a low-angle nor- ~80 km across axis and ~40 km along axis During lithospheric extension, strain is natu- mal fault would tend to decrease the dip of the (modifi ed from Cannat et al., 2006). rally partitioned in the crust into weak fault fault. Although there is ample fi eld evidence zones in the brittle layers, and shear zones and to suggest that detachment faulting occurs on homogeneous strain in the lower crust. Strain low- angle faults (e.g., Scott and Lister, 1992; spreading center is between ~3500 and 4800 m localization may result from physical processes John and Foster, 1993), large rotation of high- depth. Shallower seafl oor is likely to have too involving thermodynamic energy fl uxes even angle faults is consistent with structures seen in large a magma supply, and deeper seafl oor is in the absence of any particular rheological many continental core complexes (Hamilton, likely to have too little magma. Fast-spreading anomaly. Whatever its origin, strain localiza- 1988; Wernicke and Axen, 1988; Buck, 1988). ridges have very thin lithosphere (~1 km thick), tion in the upper crust is essential to initiate More recent paleomagnetic studies of oceanic so even a modest supply of magma could ac- core complexes, and it is easily achieved around core complexes have demonstrated a minimum commodate all lithospheric spreading. For oce- rheological and/or density anomalies such as a of 40°–50° rotation of footwall rocks below anic core complexes to form, the magma supply preexisting fault in the brittle upper crust (Buck, the detachment (Morris et al., 2009; MacLeod apparently cannot be too high or too low (i.e., 1993; Lavier et al., 1999; Koyi and Skelton, et al., 2011), validating the rolling-hinge model “the Goldilocks condition” of Tucholke et al., 2001; Gessner et al., 2007), a rheological and/or for low-angle normal and fault formation asso- 2008). If faults on opposite sides of the axis are density anomaly in the lower crust (Brun et al., ciated with plate spreading at slow and ultraslow active simultaneously and slip at the same rate, 1994; Tirel et al., 2004, 2008), or a strong den- mid-ocean ridges. then a core complex can form on both sides of sity discontinuity along the brittle-ductile transi- In their conceptual rolling-hinge models, the ridge (even in cases in which the plate sepa- tion (Wijns et al., 2005). Wernicke and Axen (1988) assumed local

286 Geological Society of America Bulletin, March/April 2013 Continental and oceanic core complexes

isostasy, whereas Buck (1988) calculated the ters of offset. Such a fault could only initiate on opment (Poliakov and Buck, 1998; Buck and fl exural response of lithosphere to loads caused a preexisting very weak zone. Poliakov, 1998); these models showed that a by the offset of a high-angle normal fault. Nor- The inclusion of realistic yield stresses de- sequence of high-angle faults might form and mal fault offset is supported regionally by fl ex- creases the wavelength of fault and footwall accommodate extension at a simple mid-ocean- ure, and since the Moho remains fl at in many bending and radically lowers the work that ridge structure. When more strain weakening continental core complexes owing to fl ow of results from fault-related topography (Buck, is allowed, large- offset faults develop (Lavier low-viscosity lower crust, the elastic lithosphere 1993). A reasonably weak fault may accommo- et al., 1999, 2000) (Fig. 8). If the amount of in which bending occurs is restricted to the date an offset suffi ciently large that the inactive fault strain weakening is dependent on the upper crust (Masek et al., 1994). By analogy part of the fault rolls over and becomes fl at. A thickness of the brittle layer being extended, with bending at subduction zones (e.g., McAdoo fault that is not suffi ciently weak is replaced then large-offset faults only develop when a et al., 1978; Bodine et al., 1981), applying a by another fault before large offset develops layer is thinner than a critical thickness. reasonable rock yield strength (the stress needed (Buck, 1988, 1993). Faults may weaken as they The behavior and characteristics of core com- to break and slip on a fault) based on labora- record slip, and the amount of fault weakening plex–bounding faults may inform not only how tory measurements for an ~10 km brittle layer needed to allow large fault offset increases lin- much faults weaken, but also how they weaken gives the observed range of domal wavelengths early with brittle layer thickness (Buck, 1993). with offset (Lavier et al., 2000). If faults weaken and topographic relief (Buck, 1988). For oce- For example, a fault affecting a 10-km-thick too fast, the entire layer shatters, and bending anic core complexes, thin crust combined with a layer would have to weaken by ~10 MPa to de- around a fi rst fault gives rise to secondary faults small density contrast between crust and mantle velop a large offset. that delocalize extension. If faults weaken too reduce the importance of crustal thickness varia- Analog models are useful for simulating the slowly with offset, then the initial fault does not tions compared to topographic variations. early, small-offset stage of fault development become suffi ciently weak before the extra resis- The energetics of fault offset and layer bend- (e.g., Brun et al., 1994; Tirel et al., 2006), but tance to slip owing to topography makes it easier ing (Forsyth, 1992) predict that the initial orien- they cannot easily simulate the thermally con- for another fault to break elsewhere. A thin layer tation of a fault corresponds to the least friction trolled strength evolution that is likely to affect can generate a large-offset fault, while multiple on the fault per unit of horizontal displace- large-offset faults. Early two-dimensional nu- faults develop in a thicker layer (Fig. 9; Lavier ment. This energy or work approach yields the merical studies (e.g., King and Ellis, 1990) and Buck, 2002). This is consistent with the ob- same initial orientation for a fault as the clas- assumed a preexisting weak normal fault em- servation that large-offset normal faults are only sic Ander sonian stress analysis, but when the bedded in a purely elastic layer and solved for seen in areas of higher-than-average heat fl ow, fault records signifi cant slip, work is done in the the topographic relief and stress changes when where one expects thin brittle crust. In areas of bending of the displaced layer. An analytical es- the fault recorded slip. Given the potential high heat fl ow and thick crust, the lower con- timation of this extra work using an approxima- importance of the fi nite brittle yield strength tinental crust may fl ow easily (e.g., Block and tion of the lithosphere as a thin, perfectly elastic (plastic deformation), early models treated the Royden, 1990; Buck, 1991), allowing a single layer fl oating on an inviscid substrate (Forsyth, lithosphere as a viscous-plastic layer and were large-offset fault (detachment) to accommodate 1992) showed that an initially low-angle normal not concerned with fault localization (Braun signifi cant extension. fault could accumulate much more offset than and Beaumont, 1989; Bassi, 1991). Subsequent Recent work suggests that relatively small a high-angle fault. If this elastic plate model is extension models of a viscous-elastic-plastic features of core complexes may help bound the correct, the initial dip of a normal fault has to be layer incorporated strain weakening as a func- amount of fault weakening with offset. Kilome- extremely low to accommodate tens of kilome- tion of strain to promote normal fault devel- ter-scale allochthonous rider blocks that are cut

STRAIN TOPOGRAPHY (m) 1500 0 plastic strain 0 5 vertical exaggeration 3.0 : 1 0 5 10 -1500

depth (km) -50distance (km) 0 +50 10 vertical exaggeration 1.5 : 1 1500 0 0 5 -1500 depth (km) 10 -50distance (km) 0 +50

1500

Increasing extension 0 0 5 -1500 depth (km) 10 -50distance (km) 0 +50

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 footwall 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).

Geological Society of America Bulletin, March/April 2013 287 Whitney et al.

Thin lithosphere Thick lithosphere 4 4

0 0

topography (km) topography –5 –5 0.7 2.1 0.7 2.1 Total Strain Total Strain 1.4 2.8 1.4 2.8 0 0 250°C 250°C 500°C 10 10 500°C 20 700°C 20 700°C depth (km) 30 30 0Distance (km) 50 100 150 0Distance (km) 50 100 150

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). off from the hanging wall and related basin in- specifi ed by the fraction M of the plate separa- ever, the fi xed temperature structure assumed by fi ll, and are transported on top of the footwall, tion rate that is accommodated by dike opening. Buck et al. (2005) is questionable when there is are common in both continental and oceanic For M = 0, dikes account for none of the plate little magma input. If the temperature structure core complexes (e.g., Reston and Ranero, 2011). spreading; for M = 1, they accommodate all of is strongly affected by the advection associated Rider blocks superimposed on large-offset nor- it. Dikes may supply much of the heat that keeps with fault offset, then the weakest part of the mal faults may not form if a fault loses too much the axis hotter and the axial lithosphere thinner lithosphere migrates with the active fault. Since strength (Choi and Buck, 2012); a narrow range than the lithosphere farther from the axis. the sensible and latent heat of intruding dikes of fault weakening relative to intact surrounding A simple geometric argument shows how should dominate over this effect, the simple rock allows the formation of large-offset faults faults and dikes interact at a ridge with a fi xed symmetric strength structure assumed in Buck with rider blocks. These blocks form when the position of diking and fi xed thermally defi ned et al. (2005) should be valid for M >~0.4. dominant form of weakening is by reduction strength structure. If one fault forms due to Tucholke et al. (2008) included a more realis- of fault cohesion, whereas faults that weaken lithospheric stretching, it should initially cut tic, evolving temperature structure that consid- primarily by friction reduction do not generate the thinnest axial lithosphere on one side of the ered advection and diffusion of heat and latent distinct rider blocks. Either a lack of infi ll or an axis. If the fault moves away from the axis into heat of crystallization. This study showed that extreme reduction of friction by serpentinization thicker lithosphere, it becomes more diffi cult large-offset faults could form for a range of M of exhumed mantle rocks may explain the lack for this fault to accommodate deformation, even values between ~0.3 and 0.6 (M = 0.5 shown in of rider blocks on some oceanic core complexes. though it is weaker than the surrounding litho- Fig. 10). As noted earlier, several sets of obser- sphere. Eventually it will be easier to form a vations indicate that oceanic core complexes de- Dynamic Models of Oceanic new fault cutting the axis, and the fi rst fault will velop when the supply of magma to a spreading Core-Complex Development be replaced by a new fault. center is not too high and not too low. Olive et al. Models of large-offset normal faulting have (2010) found that the formation of large-offset Magmatism and associated hydrothermal evaluated the infl uence of M on fault behavior faults and oceanic core complex–like structures effects are likely to be signifi cant factors in (Lavier et al., 1999). For M = 0.5, the hanging depends more on the rate of dike opening than oceanic core-complex development, but the in- wall of the fault does not move away from the on the rate of gabbro intrusion. There appears clusion of magmatism in numerical models is axis, so the fault could build up potentially un- to be a reasonable correspondence between the at an early stage. The input of limited amounts limited offset. For M close to 1, the maximum “M-based” models and the inference of similar of gabbro may be needed to allow alteration fault offset can be small, whereas for M close to modes of spreading based on geological and of peridotite to weak minerals, as silica-rich 0.5, the offset should be very large. A numeri- geophysical data from slow-spreading ridge magmatic fl uids interact with peridotite (e.g., cal model of diking and stretching (Buck et al., systems (Cannat et al., 2006, 2009). Ildefonse et al., 2007). Such weakening would 2005) tested this conceptual model using the nu- A potential problem with these simple mod- allow oceanic core-complex detachments to merical approach of Poliakov and Buck (1998). els concerns the rate of oceanic detachment slip at lower than normal levels of stress, but Results show that normal fault offset varies faulting that has been estimated using magmatic so far such weakening has not been included in greatly as a function of M; large fault offset and thermochronologic data. For example, spreading center models, and most modeling occurs when M = 0.5; for M = 0.95, the model in the Atlantis Massif, much of the full plate effort has focused on dikes. generates a fairly symmetric pattern of mainly spreading rate was taken up on the oceanic de- One simple way of treating the effect of dike inward-dipping, small-offset faults, and a sym- tachment (Grimes et al., 2008). Estimates of the intrusion in numerical models of long-term metric axial valley. For values of M < 0.5, the rate of oceanic detachment slip range from ~14 lithospheric extension was developed by Buck active fault gradually moves across the spread- km/m.y. at Atlantis Bank (Baines et al., 2008), et al. (2005). The rate of dike opening was ing axis and may be cut by later faults. How- to ~24 km/m.y. at Atlantis Massif (Grimes

288 Geological Society of America Bulletin, March/April 2013 Continental and oceanic core complexes

magma M = 0.7 injection zone 0

depth (km) 600°C 0.9 m.y. 10

M = 0.5 model original seafloor magmatic accretion seafloor faulted surface seafloor 0

strain rate (s–1) 5 –16 –14 –12

depth (km) 600°C 0.9 m.y. 10

M = 0 0

5 depth (km) 600°C 0.8 m.y. 10 30 20 10 0 10 20 30 distance (km)

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 (10–16 to 10–12 s–1); 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. et al., 2008), and up to ~38 km/m.y. at Godzilla domes within continental core complexes (Teys- gravitational potential energy. Under high ex- Megamullion (Ohara et al., 2001). John and sier and Whitney, 2002; Whitney et al., 2004b). tension rate and low surface heat fl ow, the duc- Cheadle (2010) concluded that, during the time Physical experiments show that a strongly tile crust is too strong to respond to gravitational period when the oceanic detachments were ac- coupled upper and lower continental crust leads stresses over short time scales. In this case, tive, they accounted for 60%–100% of the total to distributed surface extension, whereas me- two-dimensional models predict that the lower plate spreading. This would give M values be- chanical decoupling between weak lower crust crust extends and thins rather homogeneously tween 0.0 and 0.4; these values are lower than and much stronger upper crust leads to localized (Fig. 11, a1, with TMoho = 600 °C), redistributing predicted to produce large-offset faults. surface extension that favors continental core- stresses uniformly in the upper crust and upper complex development (Brun et al., 1994). The mantle, where numerous normal faults develop. Dynamic Models of Continental magnitude of coupling is strongly dependent That is, extension under relatively cool geother- Core-Complex Development on the geothermal gradient. In a cold to normal conditions results in more normal faults in mal geothermal gradient, the lower crust is me- the brittle upper crust (graben, half graben, and Localized thinning of the continental upper chanically coupled to the upper crust and upper rotated blocks) and relatively homogeneous crust by extension is isostatically compensated mantle. In this case, strain in the lower crust is extension in the ductile crust (Fig. 11A). The by the fl ow of deep crust into core complexes a response to plate boundary forces. In contrast, ductile crust thins and fl ows to accommodate (Block and Royden, 1990; Wdowinski and under hotter conditions, the lower crust is weak the topography of the brittle-ductile transition Axen, 1992). Partially molten crust is particu- and fl ows in response to both tectonic stresses and that of the Moho. Major normal faults in larly likely to facilitate core-complex develop- and gravitational stresses. the upper crust evolve into strong shear strain ment during extension owing to the dramatic Numerical modeling that integrates tempera- gradients in the lower crust (Figs. 11A–B, a1–a2, decrease in viscosity associated with the pres- ture-dependent viscosity (Fig. 11) addresses the b1–b2). In the ductile crust, fl at foliations domi- ence of melt. In the case of fl uid-absent melt- competition between plate boundary extension nate, although some thin regions of vertical ing reactions, which can be encountered during rate and the rate at which ductile crust can fl ow. foliation occur underneath major normal faults heating and/or decompression, a positive feed- The fl ow of ductile crust is driven by gravita- (cf. Fig. 11, a2, b2). back between extension-induced decompres- tional stresses and is controlled by the viscosity When the geotherm is warmer, the lower sion and partial melting may generate migmatite (temperature) of lower crust and the gradient in crust is able to fl ow over short time scales in

Geological Society of America Bulletin, March/April 2013 289 Whitney et al.

ABVARIABLE TMoho (600–900°C) – CRUSTAL THICKNESS 60 KM VARIABLE CRUSTAL THICKNESS (45–60 KM)

a Stress 1 faults b1 air Temperature air brittle crust

ductile crust Viscosity 45 km TMoho = 600ºC TMoho = 617ºC mantle

b a2 2

50 km TMoho = 700ºC TMoho = 702ºC

b a3 3

55 km TMoho = 800ºC TMoho = 767ºC

a4 b4

TMoho = 900ºC solidus 60 km

TMoho = 845ºC

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 extension (1.13 cm/yr at both sides, i.e., 2 × 10–15 s–1). (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 homogeneously distributed. (a3, a4) In contrast, warmer geotherms lead to more strongly localized extension in the upper crust and a more localized 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.

response to gravitational stresses (Fig. 11, a2, folds, nappes) beneath the zone of upper-crust for material that is exhumed from the lower with TMoho = 700 °C). Higher geothermal gra- extension (Rey et al., 2011). Therefore, exten- crust. These deformation stages occur at vari- dients result in more strongly localized ex- sion in the upper crust is coeval with contrac- ous P-T conditions and include shearing in the tension in the upper crust, driving fl ow in the tion in the deep crust. The two-dimensional lower-crust channel, horizontal contraction in deep crust. This illustrates that the amount of structure beneath the dome-shaped detachment the domain of collided channels, and horizontal coupling between the upper and lower crust fault consists of a subvertical high-strain zone extension when material moves up beneath the is controlled by the crustal-scale temperature that separates two subdomes, creating a double detachment zone. distribution (Block and Royden, 1990; Lavier dome (Figs. 4C and 11). In high-geotherm settings, upward fl ow of and Buck, 2002). As the lower-crust viscosity Double domes (or more complex dome geom- low-viscosity material is localized, and there- diminishes (for example, when melt fraction etries in three dimensions) generated in this fore the lower crust infl uences upper-crustal increases), the lower crust fl ows upward be- way may explain the presence of relict high- stress and strain only in the extended region neath the region of upper-crustal necking; the P metamorphic rocks (typically dismembered (necking of upper crust). Upward advection of ductile crust is hot and suffi ciently weak to be layers and pods) in domes that otherwise record heat promotes surface fl uid circulation (Mulch decoupled from the upper crust. This upward low-P, high-T emplacement. Two-dimensional et al., 2004; Person et al., 2007; Gébelin et al., fl ow generates horizontal fl ow in two channels model results predict that deep crustal rocks 2011), which contributes to further strain local- that converge toward the zone of extension and (eclogite, granulite) and possibly lithospheric ization in the upper crust. In the brittle upper move material in a direction opposite to the mo- mantle will be entrained and carried upward crust, extensional strain is strongly focused, and tion of the brittle upper crust and upper mantle. in steep high-strain zones. These simple dy- the number of normal faults is limited; typically This strongly partitioned fl ow results in the namic models are run under steady extension a single large-offset detachment fault accom- formation of contractional structures (upright but produce a complex deformation sequence modates upper-crustal extension (Lavier et al.,

290 Geological Society of America Bulletin, March/April 2013 Continental and oceanic core complexes

2000). In this tectonic setting, buoyancy of the Figure 12. Two-dimensional weak lower crust plays a second-order role, and Ellipsis models of the effect of 10 cm/yr core complexes form even when the lower crust mantle wedge melting on the 10 cm/yr is denser than the upper crust. development of Cordilleran core complexes. (A) Convergence of Subduction Dynamics and Continental oceanic and continental plates. A Core Complexes Compared to no-melt case (B), presence of melt in mantle 10 cm/yr 10 cm/yr In active ocean-continent plate margins, the wedge (buoyancy) favors crustal no melt in wedge state of stress in the upper plate can rapidly extension in overlying plate switch from contractional to extensional de- (modifi ed from Rey and Müller, pending on the interplay among (1) trench-nor- 2010). In this set of experiments, mal velocity, (2) friction along the subduction the continental lower crust is B interface, (3) gravitational forces in the thick made of diabase (density 2800 crust of the overriding plate, and (4) traction kg m–3). Despite its higher den- 10 cm/yr imposed at the base of the overriding plate by sity, this lower crust is advected 10 cm/yr the buoyant mantle wedge (e.g., Billen, 2008). upward into the continental core melt in wedge The effect of trench rollback on the stability of complexes, similar to what is ob- the overriding plate has been studied in detail served in some core complexes of C (e.g., Faccenna et al., 2007; Becker and Fac- the North American Cordillera. cenna, 2011), particularly for the Aegean region (slab rollback and breakoff; Jolivet and Brun, 2010). Here, we focus on the role of the mantle toward the trench drives necking in the brittle IMPLICATIONS AND wedge because the transformation of strong upper crust, allowing rapid formation of conti- UNRESOLVED QUESTIONS lithospheric mantle into weaker and more buoy- nental core complexes through upward fl ow of ant mantle, driven in part by the release of fl uids weak lower crust (Fig. 12C). Documenting the history of core-complex from the subducting plate and partial melting of development in the continents and oceans has the wedge, must have a signifi cant effect. Plateau-Foreland Dynamics and consequences beyond understanding exten- Under dynamic equilibrium (no deforma- Continental Core Complexes sional modes in the lithosphere. Extension of tion in the overriding plate), there is a balance the lithosphere leads to exhumation of mid- to between driving forces (ridge push, slab pull, Numerical modeling shows that the lateral lower crust and upper mantle in continental and and gravitational forces) and resisting forces fl ow of crust from beneath a continental pla- oceanic core complexes and is therefore a major (generated by friction along the subduction in- teau toward a foreland may be diverted upward process in lithospheric differentiation and sta- terface as well as other viscous forces). Gravi- to form a continental core complex (e.g., Rey bilization, element cycling, and heat and mass tational forces stored in the thick overriding et al., 2010). At plateau margins, gravitational transfer. In this section, we highlight a few of plate and the buoyant mantle wedge can shift stresses initiate stretching and necking in the these broader implications of core-complex for- this dynamic balance, for example, through the plateau upper crust, introducing lateral pressure mation and mention some questions for possible slowing of trench-normal velocities that result gradients in the lower crust. In the absence of future research. in extension of the overlying plate (e.g., Rey preexisting faults or weak regions in the plateau and Müller, 2010) and the development of core upper crust, necking develops preferentially Core-Complex Formation and complexes. above the Moho ramp that underlies the plateau Global Element Cycles Numerical models can simulate a scenario in margin (Fig. 13). Part of the lower crust fl ows which both the subducting and overriding plates horizontally into the adjacent foreland lower Along with volcanism and erosion, exhuma- converge above a subduction zone (Fig. 12A). In crust, resulting in lateral growth of the plateau tion is one of the most signifi cant processes by this two-dimensional model of a 200-km-long (e.g., Clark and Royden, 2000; Rey et al., 2010). which “fresh” rock is exposed at Earth’s surface subducting slab, a reference model in which the However, because of the local pressure gradient and made available for reaction with the atmo- mantle wedge is hydrated but has no partial melt in the vicinity of the upper-crustal necking re- sphere and/or hydrosphere. In the continents, (reference density of the wedge is 3350 kg m–3, gion, part of the lower crust fl ows upward into erosion and tectonic denudation may cooperate i.e., 20 kg m–3 less than the density of the sur- the necking region to form a continental core in a positive feedback relation; and in both con- rounding mantle) is compared with the same complex (Fig. 13). tinents and oceans, magmatism accompanies model with both hydration and partial melting Flow in the weak lower crust is therefore exhumation. Nevertheless, the magnitude of in the wedge. In the experiment with no melt in partitioned into horizontal extrusion into the ad- decompression of deep rocks under extension, the mantle wedge (Fig. 12B), the trench-normal jacent foreland and upward fl ow into the conti- primarily in core complexes, indicates that tec- velocity of 20 cm/yr maintains the stability of nental core complex. In such a tectonic setting, tonic processes are a main driver of exhumation the overriding plate (limited surface extension). buoyancy of the lower crust plays an important in the continents and oceans. When partial melt is allowed (Fig. 12C), provid- role. Indeed, when the weak lower crust has The amount of material exhumed can be es- ing more buoyancy to the mantle wedge, the no buoyancy in the numerical experiments, timated from the magnitude of vertical motion orogen becomes unstable and collapses under horizontal fl ow into the adjacent foreland domi- during extension. Vertical motion (decompres- the buoyancy traction imposed by the buoyant nates, inhibiting the formation of continental sion) may be tens to many tens of kilometers, mantle wedge. The surge of the overriding crust core complexes. and in some cases, signifi cantly more than can

Geological Society of America Bulletin, March/April 2013 291 Whitney et al.

Initial conditions plateau the extension-driven exhumation processes involved in core-complex formation contribute to 100 km foreland the global ﬂ ux of crustal and mantle material from depth to the surface?

Temperature 300 K Core-Complex Formation and Heat Transfer

1000 K The rapid advection of deep crust and upper mantle from depth to surface also involves the transfer of large amounts of heat, analogous to Fixed foreland, melt with no buoyancy, crust with depth-independent density heat transfer in the crust by magmatism. This is clearly a signiﬁ cant factor in the heat budgets of Time = to+12 m.y. extensional zones, whether at oceanic divergent zones or in collapsing orogens, and is a major Total strain driver of hydrothermal activity and therefore mineral-ﬂ uid interaction and ore deposition associated with core-complex formation. This raises the question: How much do the extension- Temperature 330000 K driven processes involved in core-complex formation contribute to the heat budgets of active plate boundaries in the oceans and continents? 1000 K Core-Complex Formation and Oblique Tectonics Fixed foreland, melt with buoyancy

Time = to+12 m.y. To have reliable estimates of material and heat fl ux, it is essential to be able to reconstruct Total strain the magnitude and paths of crustal fl ow (vertical and lateral) and the temperature-time history of core-complex footwall rocks to answer the question: What are the three-dimensional trajectories of advected material through time? Temperature 300300 K This is another unresolved issue: “Out-of- plane” motion of material may be a major mode by which crust and mantle are displaced in 1000 K extensional settings, but the role of oblique di- Figure 13. Two-dimensional Ellipsis models of continental metamorphic core-complex de- vergence and strike-slip (transform) faults is an velopment at the margin of an orogenic plateau. Top panel: Initial conditions; plateau Moho incompletely understood aspect of the dynam- T = 790 °C (>crustal solidus), and foreland Moho T = 560 °C. Middle panel: Experiment ics of extension and core-complex development. with depth-independent density in the crust, even in region of partial melting. After 12 m.y., Many continental and oceanic core complexes the plateau lower crust has fl owed 150 km laterally into the foreland crust and keeps fl owing are associated with strike-slip (transform) faults at a velocity of 1 cm/yr. The foreland upper crust has thickened by faulting. Above the pla- (Cann et al., 1997; Tucholke et al., 1998; Ranero teau ramp, a dome has formed, but, at the surface, exhumation is moderate. Bottom panel: and Reston, 1999; Baines et al., 2003; Foster Experiment with buoyant partial melt (2366 kg m–3). The lateral fl ow of the plateau lower et al., 2007; Whitney et al., 2007; McFadden crust is smaller (120 km), and a fully developed dome has formed, exhuming deep-crustal et al., 2010; Gasser et al., 2011), indicating that rocks to the surface. the dynamics of core-complex formation may be enhanced by transform motion and oblique divergence. be accounted for by displacement on detachment UHP rocks and formation of coeval basins (e.g., faults. It is common that the continental rocks re- the Western Gneiss Region, Norway—Séranne Broader Perspective cording the highest metamorphic pressures (e.g., and Séguret; 1987; Kylander-Clark et al., 2012; >1 GPa) in an orogen are in migmatite-cored Qinling-Dabie-Sulu, China—Faure et al., 2003; Extension creates an instability (exhuma- core complexes (Whitney et al., 2004b; Rey Hacker et al., 2006). tion) that ultimately results in stabilization of et al., 2011), indicating the signifi cance of these The processes, rates, and conditions of the crust, in part because the fl ux of hot, deep structures for exhuming deep crust. It is also sig- extension-driven exhumation of rocks from crust to surface and near-surface levels results in nifi cant that some exhumed ultrahigh-pressure mantle or lower-crustal depths to Earth’s sur- isostatic adjustment and rapid cooling. In oro- (UHP) complexes that formed during continental face require more investigation. Therefore, a genic settings, this marks the end of orogeny, subduction into the mantle contain extensional question relevant to core-complex research in but an open question is whether this process domains that were involved in exhumation of the the continents and oceans is: How much do drives the end of orogeny or is a consequence

292 Geological Society of America Bulletin, March/April 2013 Continental and oceanic core complexes of orogenic demise caused by other (perhaps Armstrong, R.L., 1972, Low-angle (denudation) faults, Anatolia: Journal of the Geological Society of London, external) forces. hinterland of the Sevier orogenic belt, eastern Nevada v. 151, p. 213–216, doi:10.1144/gsjgs.151.2.0213. and western Utah: Geological Society of America Bul- Braun, J., and Beaumont, C., 1989, A physical explanation of Core complexes provide a glimpse of the type letin, v. 83, p. 1729–1754, doi:10.1130/0016-7606 the relation between fl ank uplifts and the breakup un- of crustal fl ow that characterizes the deep struc- (1972)83[1729:LDFHOT]2.0.CO;2. conformity at rifted continental margins: Geology, v. 17, Armstrong, R.L., and Ward, P., 1991, Evolving geographic p. 760–764, doi:10.1130/0091-7613(1989)017<0760: ture of all continents. By redistributing material patterns of Cenozoic magmatism in the North Ameri- APEOTR>2.3.CO;2. and spreading it over an increased surface area, can Cordillera: The temporal and spatial association of Brichau, S., Ring, U., Ketcham, R.A., Carter, A., Stockli, orogenic collapse and the formation of core magmatism and metamorphic core complexes: Journal D., and Brunel, M., 2006, Constraining the long-term of Geophysical Research, v. 96, p. 13,201–13,224, evolution of the slip rate for a major extensional fault complexes represent the ultimate process of doi:10.1029/91JB00412. system in the central Aegean, Greece, using thermo- continental growth and thermal and mechanical Axen, G.J., 1999, Low-angle normal fault earthquakes chronology: Earth and Planetary Science Letters, stabilization. and triggering: Geophysical Research Letters, v. 26, v. 241, p. 293–306, doi:10.1016/j.epsl.2005.09.065. p. 3693–3696, doi:10.1029/1999GL005405. Brown, S.R., Gibson, H.D., Andrews, G.D.M., Thorkelson, In the oceans, core-complex formation, with Axen, G.J., 2007, Signifi cance of large-displacement, low- D.J., Marshall, D.D., Vervoort, J.D., and Rayner, N., its density, viscosity, and bathymetric signa- angle normal faults: Geology, v. 35, p. 287–288, doi: 2012, New constraints on Eocene extension within the 10.1130/0091-7613(2007)35[287:RFSOLL]2.0.CO;2. Canadian Cordillera and identifi cation of Phanerozoic tures, may also exert some control on ridge Axen, G.J., Bartley, J.M., and Selverstone, J., 1995, Struc- protoliths for footwall gneisses of the Okanagan Valley dynamics, fault evolution, hydrothermal activ- tural expression of a rolling hinge in the footwall of shear zone: Lithosphere, v. 4, p. 354–377, doi:10.1130 ity, and generation of new oceanic crust. The the Brenner Line normal fault, eastern Alps: Tectonics, /L199.1. v. 14, p. 1380–1392, doi:10.1029/95TC02406. Brun, J.-P., and van den Driessche, J., 1994, Extensional oceanic lithosphere produced at slow-spreading Baines, A.G., Cheadle, M.J., Dick, H.J.B., Scheirer, A.H., John, gneiss domes and detachment fault systems—Structure centers differs from Pacifi c-type lithosphere B.E., Kusznir, N.J., and Matsumoto, T., 2003, Mechanism and kinematics: Bulletin de la Société Géologique de in terms of structure, hydration, and depth of for generating the anomalous uplift of oceanic core com- France, v. 165, p. 519–530. plexes: Atlantis Bank, Southwest Indian Ridge: Geology, Brun, J.-P., Sokoutis, D., and van den Driessche, J., 1994, mineralogical changes. A question for further v. 31, p. 1105–1108, doi:10.1130/G19829.1. 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Miocene M2 metamorphism on Naxos, Greece: Geo- Buck, W.R., 1993, Effect of lithospheric thickness on chimica et Cosmochimica Acta, v. 53, p. 2037–2050, the formation of high- and low-angle normal faults: ACKNOWLEDGMENTS doi:10.1016/0016-7037(89)90323-2. Geol ogy, v. 21, p. 933–936, doi:10.1130/0091-7613 Bartley, J.M., and Glazner, A.F., 1985, Hydrothermal (1993)021<0933:EOLTOT>2.3.CO;2. We thank our many colleagues and students who systems and Tertiary low-angle normal faulting in Buck, W.R., and Poliakov, A.N.B., 1998, Abyssal hills formed have worked with us on core complexes over the the southwestern United States: Geology, v. 13, by stretching oceanic lithosphere: Nature, v. 392, years, and we thank David Foster and an anonymous p. 562–564, doi:10.1130/0091-7613(1985)13<562: p. 272–275, doi:10.1038/32636. HSATLN>2.0.CO;2. Buck, W.R., Martinez, F., Steckler, M.S., and Cochran, J.R., reviewer for valuable comments that helped us im- Bassi, G., 1991, Factors controlling the style of continen- 1988, Thermal consequences of lithospheric exten- prove the paper. Whitney and Teyssier acknowledge tal rifting; insights from numerical modeling: Earth sion: Pure and simple: Tectonics, v. 7, p. 213–234, support from National Science Foundation (NSF) and Planetary Science Letters, v. 105, p. 430–452, doi:10.1029/TC007i002p00213. grants EAR-0911497 and EAR-1050020, and Teys- doi:10.1016/0012-821X(91)90183-I. Buck, W.R., Lavier, L.L., and Poliakov, A.N.B., 2005, sier acknowledges support from NSF grant EAR- Batt, G.E., and Braun, J., 1997, On the thermomechani- Modes of faulting at mid-ocean ridges: Nature, v. 434, 0838541. Buck thanks Javier Escartin, Mathilde cal evolution of compressional orogens: Geophysical p. 719–723, doi:10.1038/nature03358. 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