Download Chapter 17 Basins in arc-continent collisions

Chapter 17
Basins in arc-continent collisions
AMY E. DRAUT and PETER D. CLIFT†
U.S. Geological Survey, Santa Cruz, California, USA
School of Geosciences, University of Aberdeen, Aberdeen, UK
†
ABSTRACT
Arc-continent collisions occur commonly in the plate-tectonic cycle and result in
rapidly formed and rapidly collapsing orogens, often spanning just 5–15 My. Growth
of continental masses through arc-continent collision is widely thought to be a major
process governing the structural and geochemical evolution of the continental crust over
geologic time. Collisions of intra-oceanic arcs with passive continental margins (a
situation in which the arc, on the upper plate, faces the continent) involve a substantially
different geometry than collisions of intra-oceanic arcs with active continental margins
(a situation requiring more than one convergence zone and in which the arc, on the lower
plate, backs into the continent), with variable preservation potential for basins in each
case. Substantial differences also occur between trench and forearc evolution in
tectonically erosive versus tectonically accreting margins, both before and after
collision.
We examine the evolution of trenches, trench-slope basins, forearc basins, intra-arc
basins, and backarc basins during arc-continent collision. The preservation potential of
trench-slope basins is low; in collision they are rapidly uplifted and eroded, and at
erosive margins they are progressively destroyed by subduction erosion. Post-collisional
preservation of trench sediment and trench-slope basins is biased toward margins that
were tectonically accreting for a substantial length of time before collision. Forearc
basins in erosive margins are usually floored by strong lithosphere and may survive
collision with a passive margin, sometimes continuing sedimentation throughout
collision and orogeny. The low flexural rigidity of intra-arc basins makes them deep
and, if preserved, potentially long records of arc and collisional tectonism. Backarc
basins, in contrast, are typically subducted and their sediment either lost or preserved
only as fragments in m
elange sequences. A substantial proportion of the sediment
derived from collisional orogenesis ends up in the foreland basin that forms as a result of
collision, and may be preserved largely undeformed. Compared to continent-continent
collisional foreland basins, arc-continent collisional foreland basins are short-lived and
may undergo partial inversion after collision as a new, active continental margin forms
outboard of the collision zone and the orogen whose load forms the basin collapses in
extension.
Keywords: arc continent collision; forearc basin; backarc basin; trench slope basin;
tectonic erosion
IN TR ODUCTION
The sedimentary record that accumulates around
volcanic arcs forms the basis for interpreting the
tectonic and magmatic history of active margins,
both as an end in itself and as a means to under-
stand the development of Earth’s continental crust,
much of which formed in these settings (Rudnick
and Fountain, 1995; Hawkesworth and Kemp,
2006). Although arc basins can contain the most
complete sedimentary and geochemical history of
plate convergence and continental-crust generation
Tectonics of Sedimentary Basins: Recent Advances, First Edition. Edited by Cathy Busby and Antonio Azor.
Ó 2012 Blackwell Publishing Ltd. Published 2012 by Blackwell Publishing Ltd.
347
348
Part 4: Convergent Margins
Talkeetna
Chitina/Wrangell Mtns
Urals
Connemara
-Mayo
Kamchatka
Altaids-Mongolia
Kohistan
as
At l
Hellenic
Nankai
Kuril
Aleutians
Cascadia
Tanzawa
Taiwan
Izu-BoninMariana
Okinawa
Luzon
Papua New
Guinea
Halmahera
Bismark
Solomons
Sumatra
Banda Arc
Savu Timor
Tonga
Alisitos
Lesser
Antilles
Guerrero
Peru
N. Chile
Kermadec
S. Chile
South
Sandwich
Fig. 17.1. World map showing the active arcs and accreted arc terranes discussed in the text. Subduction zones are shown
with open triangles indicating erosive margins and filled triangles indicating accretionary margins.
of any geologic province, their utility in deciphering the ancient record is hindered by their low
preservation potential relative to basins that form
away from destructive plate boundaries. Nevertheless, arc basins are preserved in some suture
zones (Fig. 17.1). Much can also be learned from
the sedimentary record of active arc basins
(Fig. 17.1), from which a more complete understanding of facies evolution can be obtained. This
in turn allows such units to be recognized more
easily in accreted terranes and used to make
paleogeographic and tectonic reconstructions.
In the 15 years since the contributions of Busby
and Ingersoll (1995), new advances in understanding basin formation and evolution have been permitted by improved and more widely applied
marine mapping techniques around active arcs.
These methods have generated high-resolution
images of seafloor features, structure, and sub-bottom stratigraphy (e.g., Gaedicke et al., 2000; Wright
et al., 2000; Laursen et al., 2002; Kopp et al., 2006).
Geodynamic modeling capabilities have similarly
expanded in recent years, leading to a greater
understanding of crustal loading and basin development (Tang and Chemenda, 2000; Londo~
no and
Lorenzo, 2004; Waltham et al., 2008). Geophysical
imaging of active and ancient orogens has greatly
improved knowledge of tectonic processes in
arc-continent collision zones (Snyder et al.,
1996; Gaedicke et al., 2000; Brown et al., 2006a;
Wu et al., 2007). Field studies of modern and
ancient arc collisions have also extended our
understanding of climate-tectonic coupling (Dadson et al., 2003; Johnson et al., 2008; Kimura
et al., 2008).
Here we review briefly the major types of basins
that form at intra-oceanic arcs and that evolve
during arc-arc and arc-continent collision, and
assess accreted arc terranes to evaluate the fate
of basins during collisions of arcs with continental
margins. Our discussion draws upon specific
examples, but is not intended to be an exhaustive
review of all accreted arcs. Arc-continent collisions not only commonly destroy pre-existing
basins but also create new ones, as we review
using as an example the ongoing collision of the
Luzon Arc with the Eurasian margin at Taiwan.
Other examples of presently active arc-continent
collision (Fig. 17.1) occur in Papua New Guinea
(Melanesian arc; e.g., Cloos et al., 2005) and Japan
(Izu-Bonin arc; e.g., Taira et al., 2001). Understanding the evolution of accommodation space,
Basins in Arc-Continent Collisions
together with constraints on lithospheric strength
throughout all phases of arc activity and collision,
maximizes what can be learned from analyzing the
sedimentary and geochemical records of arc
basins.
Ultimately, understanding arc-continent collision processes can provide valuable insight into
the formation and evolution of continental crust.
The growth of continents by accretion of arc terranes, accompanied by siliceous magmatism and
loss of dense, mafic lower crust through delamination (Bird, 1978) or convective instability (Jull
and Kelemen, 2001), is thought to be a key process
by which the continents attained their structure
and geochemical composition (Pearcy et al., 1990;
Rudnick and Fountain, 1995; Suyehiro et al., 1996;
Holbrook et al., 1999; Busby et al., 2006;
Brown, 2009; Draut et al., 2009). In addition,
most of the continental material that returns to
the upper mantle does so via subduction zones,
the dynamics of which are recorded in arc and
trench-slope basins.
INTRA-OCEANIC ARC BASINS AND
THEIR FATE IN COLLISION ZONES
Intra-oceanic arc systems were discussed in detail
in Busby and Ingersoll (1995). We review briefly
the major basin types formed along intra-oceanic
arcs and discuss what is known from accreted arc
terranes about how each basin type may evolve
during arc-continent collision. Our discussion is
restricted to basins that form around intra-oceanic
arcs, and which eventually collide with continents; continental arcs and their basins are beyond
the scope of this work and are well reviewed
elsewhere (Wilson, 1991; Jordan, 1995; Fildani
and Hessler, 2005; Centeno-Garcia et al., 2008;
Lamarche et al., 2008). Here we consider “intraoceanic” arcs to be those that formed as a result of
subduction initiation within oceanic lithosphere
and do not include pre-existing continental crust
within their basement prior to the collision event.
Oceanic subduction zones assume a geometry
that is a function of plate convergence rate, slab age
and dip, as well as the rheology and buoyancy of
the upper plate (Turcotte and Schubert, 1982;
Jarrard, 1986; Billen and Gurnis, 2003). Crosssections of erosional and accretionary intraoceanic subduction systems are shown in Figure 17.2,
with representative sedimentary profiles from various
arc basins in Figure 17.3. Approximately three
349
quarters of modern subduction zones are of the
erosional type (Fig. 17.2A), whereas one quarter
are of the accretionary type (Fig. 17.2B; von Huene
and Scholl, 1991). Among the major arc-basin
types, it has long been recognized that geodynamic
mechanisms of basin formation differ in important
ways. Trenches and trench-slope basins form along
the boundary between the two plates in response to
flexure of the subducting slab and extension of the
upper plate driven by basal tectonic erosion
(Fig. 17.2; von Huene and Scholl, 1991; Underwood and Moore, 1995; Underwood et al., 2003).
Forearc basins form on the upper plate between the
arc and the outer-arc high (trench-slope break;
Dickinson, 1995). Intra-arc basins form among volcanic edifices on the arc platform, as part of the arc
massif (Smith and Landis, 1995). Backarc basins
can originate by rifting of the arc, associated with
high heat flow as new crust forms, or can be deepwater features formed by trapped old oceanic crust
that pre-dates the adjacent arc (Karig, 1971; Taylor
and Karner, 1983; Tamaki and Honza, 1991;
Clift, 1995; Marsaglia, 1995). Deep basins can
also form very close to the arc in response to
flexural loading by the volcanic arc edifice
(Waltham et al., 2008), as well as from arc extension. Each basin type is characterized by different
sedimentary facies, structural controls, and underlying basement composition, and each responds
differently to the processes of arc-continent collision. Basin types differ in their potential for preservation, metamorphism, or destruction as the arc
in which they formed collides with a continental
margin. Basin sedimentation in arc-continent collisions can be thought of as rapid but short-lived;
collision, orogeny, and orogenic collapse span
typically 5–15 My. Nonetheless, explosive volcanism in the arc coupled with rapid erosion of the
uplifting orogen results in voluminous sediment
supply into the basins around the collision zone.
Trenches and trench-slope basins
Outboard of the arc and forearc, sediment accumulates in the trench and usually form an accretionary prism if the thicknesses exceeds 1 km
(Clift and Vannucchi, 2004; Fig. 17.2B). Intraoceanic arcs that collide with a passive (continental) margin typically form an accretionary prism as
collision nears, even if the open-ocean phase of arc
activity is typified by subduction erosion
(Fig. 17.2A). In either case, however, the preservation potential of trench sedimentary sequences is
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Part 4: Convergent Margins
(A)
Arc volcanic front
Backarc
basin
Intra-arc
basin
Forearc
Outer forearc
basin
high
Trench slope
basins
Sedimentary
wedge
Tectonically eroded
upper plate material
(B)
Backarc
basin
Arc volcanic front
Intra-arc
basin
Perched trench
slope basin
Forearc basin
Landward
dipping
thrust
Frontally accreted
trench sediment
Underplated
material
Arc-derived sediment
Oceanic crust
Accreted trench sediment
Arc crust
Fig. 17.2. Schematic cross sections through (A) a typical western-Pacific-type intra-oceanic arc undergoing subduction
erosion, showing basins between the arc and the outer arc high, backarc basin development, as well as extension associated
with subduction-erosion tectonics; and (B) an accretionary intra-oceanic arc system (e.g., Luzon or Lesser Antilles) in which
sediment subduction has formed a frontal accretionary prism and is being underplated beneath the forearc. Intra-arc basins,
shown here inboard of the volcanic centers, commonly occur along the axis of volcanism. Source: Diagrams modified after
Clift and Vannucchi (2004). Reproduced with permission of the American Geophysical Union.
limited. Structurally controlled basins on the
trench slope of subduction zones (Fig. 17.2A)
form smaller depocenters; these can occur in
both erosional and accretionary subduction
zones. These perched trench-slope basins, which
in Tonga are as large as 10 50 km, have complex
stratigraphy reflecting their dynamic structural
setting (Fig. 17.3; Clift et al., 1998). They are formed
by extension of the brittle, deformed crust in the
cold forearc region. Because the flexural rigidity of
the trench slope is low owing to strong deformation, these basins may be quite deep and yet narrow
in their aspect. In some arcs the trench and trench-
slope basins receive sediment from a broad, elevated forearc platform (delivered via submarine
canyons), but not necessarily from the arc itself
because the platform and forearc basin separate
the arc front from the trench (e.g., Tonga, Hellenic,
and Lesser Antilles arcs; Scholl and Vallier, 1985;
Marsaglia and Ingersoll, 1992; Draut and
Clift, 2006). High-resolution bathymetric images
have shown, for example, the forearc Tongan
Platform and canyons that incise it (Wright
et al., 2000). The preservation potential of
trench-slope basins is very low; in collision
they are rapidly uplifted and eroded, and at
Basins in Arc-Continent Collisions
(A)
Backarc
deep-water
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(B)
(C)
(D)
(E)
(F)
(G)
Backarc
shallow-water
Intra-arc
basin
Forearc
shallow-water
Forearc deep-water
erosional
Forearc deep-water
accretionary
Trench-slope
basin
Shale
Pelagic carbonate
Lapilli tuff
Accretionary prism
Airfall tuff
Fine sand
Conglomerate
Slumping
Lava
Coarse sand
Volcanic breccia
Fig. 17.3. Representative vertical profiles of sediment in various types of basins in intra-oceanic arc settings. (A) Distal, deepwater backarc basin; volcanic basement overlain by breccia and then pelagic carbonates, volcanic sands, and airfall tuffs.
Proportion of carbonate increases upsection as the basin widens. (B) Proximal, shallow-water backarc basin; sediment
generally similar to distal backarc basin, but coarser. Includes proximal volcanic products with siliceous composition,
including lavas, tuffs, and lapilli tuffs. (C) Intra-arc basin; volcanic basement with overlying shallow marine carbonates
indicating subsidence; dominated by proximal volcanic products, breccias, tuff (pumice and lapilli tuff), occasional
turbidites and carbonates. (D) Proximal, shallow-water forearc basin; similar to proximal backarc basin, essentially debris
aprons around volcanoes, with breccias (mass-flow deposits) interbedded with turbidites. (E) Distal, deep-water forearc basin
at an erosional margin; volcanic basement overlain by breccia, deepwater carbonates, and channelized volcaniclastic
sediment from the arc, including submarine fan complexes made possible by great lateral continuity of basin. (F) Distal
forearc basin at accretionary margin; basement is deformed, uptilted accretionary prism material overlain by deepwater shale,
flysch turbidite sequence (sands and shales), occasional airfall tuff; background sedimentation is dominated by shale as
opposed to carbonate, because the thickest accretionary prisms form where there is terrestrial runoff from continental rivers,
implying delivery of mud will dominate over background pelagic carbonate sedimentation. (G) Trench-slope basin, shown
here with volcanic basement as in the erosive margin at Tonga, although in accretionary settings such as Nankai the basement
would be deformed sediments of the accretionary prism; shale-dominated basin fill with sandstones and turbidites including
channel cut-and-fill conglomerates; steep slopes and active faulting causes slumping that reworks deposits.
erosive margins even in intra-oceanic subduction
they are progressively destroyed by subduction
erosion.
Substantial differences are expected between
trench and forearc evolution (and associated
sedimentation) in tectonically erosive versus
tectonically accreting margins (Fig. 17.2), both
before and after collision. Along accretionary margins such as North Luzon, Nankai, Sumatra, or the
Lesser Antilles, accommodation space may be limited and sediment can be deposited in perched
basins overlying the accretionary wedge (Beaudry
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Part 4: Convergent Margins
and Moore, 1985; Underwood et al., 2003). Uplift
can culminate in emergence and recycling of sediment from the accretionary prism that can mix
with arc-derived volcaniclastic material forming
shallow-water facies. Emergence is most likely to
happen where the forearc wedge is affected by
substantial underplating and where thrust and
antithetic thrust faults shorten the wedge. In contrast, trench slopes along tectonically erosive margins (e.g., western Pacific; Fig. 17.2A) subside
fairly rapidly in response to basal erosion and
thinning of the forearc crust, creating abundant
accommodation space near the trench, and are
areas of slow, deep-water sedimentation. Trenches
and their sedimentary record are progressively
destroyed as subduction erosion causes migration
of the arc toward the trench (Collot et al., 2008).
Therefore, the preservation of trench sediment and
trench-slope basins following arc-continent collision is biased toward margins that were tectonically accreting for a substantial length of time
before collision. If accretionary prism material is
preserved in an orogen, it may be difficult to distinguish from volumetrically less important
trench-slope basin deposits (Dickinson, 1982),
although the accretionary prism might be expected
to be more deformed.
Trench sediment from tectonically accretionary
margins (Fig. 17.2B) is found in a number of
ancient arc terranes; accretionary prism metasedimentary rock tends to be more altered and
deformed than other parts of accreted arcs, having
been metamorphosed even before collision. This is
a result of burial below the forearc, where the
accreted, underplated material can be deformed
and heated (Sample and Moore, 1987; Moore
et al., 1991). Frontal accretionary prisms, in contrast to underplated material, tend to be deformed
but not substantially metamorphosed. The metasedimentary Westport Complex and South Connemara Group of western Ireland, for example, are
interpreted as frontal accretionary melange material associated with an Ordovician island arc that
collided with the Laurentian continental margin
ca. 470 Ma (Dewey and Ryan, 1990; Ryan and
Dewey, 2004). A rapid transition from tectonic
erosion in an oceanic setting to collision of the
arc with a passive margin results in zones of subduction m
elange, often ophiolitic in composition,
only hundreds of meters thick interposed between
the accreted arc and the continent, for example in
the Shyok Suture between Kohistan and Eurasia
(Robertson and Collins, 2002), and the Lichi
Melange in Taiwan (Chang et al., 2000). This
melange represents the contents of the subduction
channel at the time of collision, frozen in place as
plate motion effectively ceased.
The Magnitogorsk accreted arc terrane in the
Ural Mountains, the product of diachronous arccontinent collision in Devonian to early Carboniferous time (summarized by Puchkov, 2009)
includes metasedimentary rocks of the Southern
Urals accretionary complex. This Devonian accretionary complex provides important constraints on
the timing of arc-continent collision (AlvarezMarron et al., 2000; Brown et al., 2006b). There,
greywacke turbidites 5–6 km thick have been
deformed, brecciated, and ductilely sheared into
a large synform, with metamorphic grade ranging
from anchizone to greenschist facies. Brown
et al. (2006b) interpreted the end of trench sedimentation in this accretionary prism (Frasnian;
385–375 Ma) as indicating uplift and widespread
basin instability caused by the arrival of the full
thickness of the Baltica continental crust at the
subduction zone.
Deformation of the Southern Urals accretionary
complex is thought to resemble processes in modern arc-continent collisional settings of the western Pacific and southeast Asia at New Guinea,
where northward subduction of the oceanic part
of the Australian plate began ca. 30 Ma (Quarles
van Ufford and Cloos, 2005). In Timor, the 6-kmthick accretionary wedge in the Timor Trough
(trench to the Banda Arc) is emplaced southward
over the Australian continental margin as basement-involved thrusting imbricated the accretionary prism (Snyder et al., 1996). To the east, in a
related collision between the Bismarck Arc and
Australian margin, turbidites and pelagic deposits
of a frontal accretionary prism are exposed as the
Erap Complex in the Finisterre Mountains of
Papua New Guinea (Abbott et al., 1994), interpreted as sedimentary material scraped off the
lower (Australian) plate. The Erap Complex has
been sliced into imbricate thrust sheets as crustal
thickness has doubled; the accretionary prism
itself is being overthrust by the Bismarck Arc
along out-of-sequence thrusts as both are emplaced
over part of the Australian continental margin
(Abbott et al., 1994). Sedimentary material that
was formerly underplated beneath this same arc
is exposed as the Derewo metamorphic belt of
northern Papua New Guinea (Warren and
Cloos, 2007). Arc-continent collision in the New
Guinea region has involved the preservation of
Basins in Arc-Continent Collisions
several syn-collisional basin sequences; one, the 7km-thick siliclastic Aure Group, is interpreted as
having filled the trench (Aure Trough) of the Oligocene Papuan subduction zone as the early arccontinent-collisional orogen underwent unroofing
and erosion; the preserved trench fill was uplifted
and deformed during mid-Miocene time (Quarles
van Ufford and Cloos, 2005).
It is important to note that accretionary complexes are generally constructed in the final stages
of collision between an arc and a passive continental margin, whereas open-ocean subduction is
more commonly associated with subduction erosion. This difference largely reflects the relative
flux of sediment to the trench in each setting. As an
arc that faces a passive continental margin nears
collision with that margin (as in Taiwan; Fig. 17.4),
accretionary processes will dominate because the
sedimentary section entering the trench greatly
exceeds 1 km thickness. Because arc-continent collision is commonly oblique, with the active collision area migrating progressively along strike,
much of the trench sediment that accretes prior
to suturing is actually reworked from the active
orogen along strike (i.e., material that accreted onto
the margin and was subsequently eroded and redeposited along strike in the trench). In cases
where the arc collides by backing into an active
continental margin (as opposed to facing a passive
continental margin), sedimentation would not
increase substantially at the colliding arc’s trench,
and so a large accretionary complex is less likely to
form (e.g., the Jurassic Talkeetna arc, discussed
below; Fig. 17.5).
Forearc basins
Accommodation space in forearc basins is controlled by the geometry of the arc massif, the
trench-slope break, and their position relative to
sea level (Dickinson, 1995). Uplift and faulting of
the outer-arc high can modify basin shape over
time (Ryan and Scholl, 1989). It has been proposed
that structural relief and subsidence of forearc
basins, particularly those of continental arcs, are
functions of crustal thinning by basal subduction
erosion (Oncken, 1998; Clift et al., 2003; Wells
et al., 2003). Although recent numerical modeling
indicates that forearc basins can form as a natural
consequence of intra-plate coupling at subduction
zones and do not require focused subduction erosion (Fuller et al., 2006), both subduction erosion
and underplating beneath the forearc may still
353
contribute to vertical tectonic motions and thus
to forearc-basin and outer-arc-high topography
(Sample and Moore, 1987; Moore et al., 1991).
Forearc basins do not require active subsidence
to form, but can reflect relative uplift of structural
highs that allows sediment to pond between the
uplifted region and the arc itself. Forearc basin
geometries can vary greatly depending on the
mechanical character of the basin lithosphere.
Low flexural rigidity (relatively weak lithosphere)
would be anticipated for basins overlying tectonized accretionary complexes, whereas basins
overlying mature, igneous, oceanic forearc crust
would be less likely to deform because of their
cold, stiff lithospheric roots (Fig. 17.6); the latter
are more likely to survive collisional orogenesis,
even if juxtaposed against weaker accretionary
complexes.
Sediment in the forearc basin before collision
(Fig. 17.3D–F) includes proximal mass-flow
deposits from the steep slopes of volcanic centers,
distal volcaniclastic turbidites, reworked ophiolitic debris from the trenchward side of the basin (the
outer-arc high), with pelagic sedimentation significant only at the outer forearc, the only area likely
to record distal tephra fallout without disruption
by mass wasting (Dickinson, 1974; Larue
et al., 1991; Marsaglia and Ingersoll, 1992; Underwood et al., 1995; Ballance et al., 2004; Draut and
Clift, 2006). Submarine canyons can transfer sediment into forearc basins (Kopp et al., 2006), just as
canyons feed trenches elsewhere. Forearc basins
may contain a sedimentary record of much longer
duration than the time the arc front has been active
in any one place, as in the Mariana arc, which has
been split twice by rifting while providing sediment to the same forearc basin for 45 My. As discussed above for trenches, tectonically erosive and
accreting margins (Fig. 17.2A and 17.2B, respectively) differ in the accommodation space available
for sediment storage, with greater water depths
maintained in forearc basins of erosive margins
(Fig. 17.3E) because of ongoing basement
subsidence.
Forearc basins not only can be preserved during
arc-continent collision, but also may even continue to subside and accumulate sediment
throughout the collision. This partly reflects the
strong mechanical character of oceanic forearc
lithosphere (Fig. 17.6). Ryan (2008), using the
Ordovician Grampian orogen of western Ireland
as an example, showed this to be possible if eclogite formation in the subducting outer continental
354
Part 4: Convergent Margins
Northeastern Taiwan
Exhumation and
formation of Ilan Basin
NW
SE
Lishan Fault
detachment
Taipei Basin
Ilan Basin
sea level
0
Depth (km)
Okinawa Trough
20
40
60
Central Taiwan
Lishan Fault
thrust
NW
SE
sea level
awa
Okin gh
Trou
20
n
isio
ly
ecoallr
Depth (km)
0
enic
orogapse
coll
col
l
oroisiona
ge l
n
Compression and
metamorphism
North
Luzon
40
oceanic arc
North
Luzon
Trough
60
South of Taiwan
Active oceanic subduction
Chinese passive
margin
Accretionary
Complex
North Luzon
Trough
Luzon
Arc
NW
SE
sea level
Depth (km)
0
20
40
Arc lower crust
and mantle
oceanic crust
continental crust
60
Arc lower crust
Arc crust
Continental crust
Accretionary Complex
Oceanic crust
Continental margin sediment
Mantle
Fig. 17.4. Schematic depiction of arc-continent collision at Taiwan, in which the Luzon arc collides with the passive
continental margin of Eurasia. Inset map shows locations of the three cross-sections. The origin of the Ilan Basin is interpreted
as a result of gravitational collapse of the Taiwan Central Ranges during a SW-migrating collision of the Luzon Arc and
mainland China. Source: Clift et al. (2008). Reprinted with permission of the Geological Society of America.
Basins in Arc-Continent Collisions
180 Ma: Oceanic subduction
N
Oshetna
Series
eroded
forearc
Oceanic subduction
S
tectonic erosion
lower crust
172 Ma: Pre-collisional
migration of volcanic front
tectonic erosion
155 Ma: Syn to post-collisional
Naknek Formation
back-tilting of Talkeetna Arc
Proto-Border Ranges Fault
intermittent accretion
Fig. 17.5. Schematic diagram showing a collision in which
an arc backs into an active continental margin (continental
arc), using the example of the Jurassic Talkeetna arc in
south-central Alaska. Deformation is concentrated between
the colliding arcs, while the intra-oceanic trench to the
south remains open and active. Source: Clift et al.
(2005b). Re-printed with permission of the Geological
Society of America.
margin sufficiently reduces buoyancy of the lower
plate and causes concurrent forearc subsidence. In
the Grampian orogen, the South Mayo Trough
comprises a 9-km-thick forearc-basin succession
that includes, with no major unconformities, pre-,
syn-, and post-collisional sedimentary and volcanic deposits (Dewey and Ryan, 1990; Draut and
Clift, 2001). Modeling by Ryan (2008) suggests that
forearc basins in the hanging wall of a subduction
zone are unlikely to survive a collision unless their
topography is isostatically suppressed, such as by
substantial volumes of eclogite forming in the
footwall (Cloos, 1993; Dewey et al., 1993). Otherwise, even if forearc basins are not destroyed during orogeny, basin sedimentation most likely
ceases owing to uplift and loss of accommodation
space during collision.
355
The uplifted, tilted Miocene South Savu Basin,
Indonesia, is one such basin, where rapid uplift
(5 mm/yr) and subaerial exposure are attributed to
underthrusting of rigid forearc basement by buoyant continental crust of the Australian margin (van
der Werff, 1995). Part of that forearc basin remains
seismically active now (North Savu Basin), but it is
progressively losing accommodation space as arccontinent collision proceeds.
Alternatively, forearcs and their basins can be
subducted and destroyed during arc-continent collision. Physical experiments by Shemenda (1994)
and Boutelier et al. (2003) indicated that if collision
stress ruptures the upper plate at the arc front, the
forearc region could be entirely overthrust by the arc
massif and subducted, leaving perhaps only the
most trenchward part of the forearc crust and
basin to be obducted and incorporated into an orogen. Although such a scenario has been invoked for
the intra-oceanic Achaivayam-Valaginskaya arc terrane that collided obliquely with continental crust
at Kamchatka in Paleocene to Eocene time (Konstantinovskaia, 2001), the field evidence supporting
this model is unclear. Crustal thickness and density
analyses by Cloos (1993) showed that some western
Pacific arcs have crust thin enough (<17 km for
basaltic crust, <15 km for granitic arc crust) for
the arc potentially to be subducted in its entirety.
Arc subduction appears to have occurred in
several collision zones, for example the northern
Izu-Bonin arc subducting under Honshu
(Otsuki, 1990) and the Aleutian arc being underthrust below Kamchatka (Scholl, 2007). However,
whether the arc crust is actually subducted to great
depth is far from clear, as it may instead be underplated below the tectonized crust of the overriding
arc. Mass-balancing arguments for the continental
crust suggest that major arc crustal subduction
cannot be common (Clift et al., 2009a). In the
case of Izu, it may be argued that most of the arc
crust is accreted and underplated to the edge of the
colliding Japanese margin. In the Izu-Honshu collision zone, the lower crust of the Izu arc is exposed
in the Tanzawa Mountains (Tani et al., 2007) and
the forearc basin is exposed in the Miura Peninsula
(Ogawa et al., 1985), indicating that at least some of
the arc is being accreted. Arc thrusting is also
known in the Molucca Sea, where the Sangihe
arc is thrust over the Halmahera arc (Hall and
Smyth, 2008). The same process is inferred in
Ordovician accreted terranes in Newfoundland,
where partial subduction of one arc under another
is inferred to have caused rapid subsidence,
356
Part 4: Convergent Margins
(A) Backarc basin
(B) Forearc basin
Stress differential (GPa)
0
Stress differential (GPa)
2
0
0
Moho
Moho
50
Depth (mk)
Depth (mk)
2
0
50
100
100
150
150
Quartz-bearing upper crust
Plagioclase-bearing lower crust
Mantle lithosphere
loading, and deposition of black shales in a newly
formed basin above the subducting arc and backarc
(Zagorevski et al., 2008). The younger and thinner
the arc crust, the more likely it is to be subducted;
therefore, basins of young, thin arcs are less likely
to be preserved in orogenic belts than basins of
older, thicker arcs (Cloos, 1993). Whether they are
ultimately subducted or preserved in an orogen,
forearc crust and basin material can be substantially deformed and offset by strike-slip faulting in
collision zones, as in the Kuril-Hokkaido Arc since
the Late Miocene (Kusunoki and Kimura, 1998)
and in the oblique collision of the Aleutian arc with
continental crust at the Mys Peninsula, Kamchatka
(Gaedicke et al., 2000).
Another fate for pre-collisional forearc basins
can be seen in the active example of Taiwan
(Figs. 17.4 and 17.7). In this area the accretionary
complex grows as the Luzon arc progressively
collides with the passive margin of China
(Suppe, 1981; Huang et al., 2000). Imbrication of
Chinese passive-margin sediment, as well as Taiwan-derived sediment fed into the Manila Trench,
leads to formation of a large accretionary ridge and
eventually the Coast Ranges onshore (Fig. 17.7).
Although the dominant thrust direction is toward
the trench, backthrusts also push accretionary
material over the older North Luzon Trough forearc
basin, resulting in its eventual burial and cessation
of sedimentation (Lundberg et al., 1997; Hirtzel
Fig. 17.6. Diagram showing where the
strength lies in the lithosphere during
arc-continent collisions. Basins located
close to the arc volcanic front or in backarc
rifts (A) tend to be weak (Watts et al., 1982;
Watts, 2001), while forearc basins (B) have
stronger underlying mantle because they
are thermally mature, i.e., cooler (Lin and
Watts, 2002; Crawford et al., 2003). Forearc basins are thus less likely to be
deformed
and
destroyed
during
orogenesis.
et al., 2009). Basin remnants are exposed onshore,
but their width is much reduced and the sediment
preserved is commonly strongly deformed
(Fig. 17.7). Although the arc locally overthrusts
the basin, most of the arc crust lies at depth under
the mountains and is efficiently accreted
(obducted) onto the margin of Asia (Clift
et al., 2009a). The preservation potential of sediment in the forearc basins south of Taiwan is
modest compared with the foreland basin and
post-collisional collapse basins, which are nonetheless filled owing to along-strike transport of syncollisional sediment.
Preserved forearc basins provide valuable data
concerning the timing and dynamics of arc-continent collision. The aforementioned South Mayo
Trough has yielded a detailed record of volcanic
geochemical changes during arc-continent collision, indicating the degree and timing of continental-crust subduction (Draut and Clift, 2001). In
the Urals, the Devonian Aktau Formation represents 5-km-thick forearc basin deposits of cherty
and volcaniclastic sediment (Brown et al., 2006b).
As the passive continental margin arrived at that
trench, forearc basin sedimentation increased and
arc-derived volcaniclastic turbidites were deposited across what had been the subducting slab.
Rapid sedimentation, soft-sediment deformation
of poorly consolidated forearc sediment, and
emplacement of large olistostromes (10 km2)
Basins in Arc-Continent Collisions
357
120º
Okinawa
Trough
(A)
Ta
St iwa
ra n
it
25º
W
IP
LF
Volcanic Arc
Mawuku
Intra-arc
basin
Tungho
Limestone
Taiyuan
Forearc Basin
Cen
tral
Coa
Ran
stal
ge
Ran
ges
Coastal Range
Lichi
Melange
Chengkung
Intra-arc
basin
24º
E
A
23º
HB
1 km
22º
Two-way Travel time (sec)
(B)
0
Two-way Travel time (sec)
MT
2
Henchung
Peninsula
North Luzon Forearc Basin
B
C
NLT
Huatung
Ridge
0
1
2
1
W
E
3
2
3
4
4
1 km
5
(C)
5
W
E
Accretionary prism
3
2
North Luzon Trough
Arc flank
3
4
4
5
5
6
6
7
2 km
8
7
8
Fig. 17.7. Cross sections through the North Luzon Trough (NLT), the forearc basin to the Luzon Arc. Insert map shows the
locations of the cross sections. HB ¼ Huatung Basin, MT ¼ Manila Trench, IP ¼ Ilan Plain, LF ¼ Lishan Fault. (A) The
telescoped forearc basin onshore in central Taiwan (after Huang et al., 2000). (B) The basin just south of Taiwan, showing a
deeper fill and much wider extent, and (C) in deep water remote from the collision zone but showing its western edge already
impinged by the Manila Accretionary Prism.
are attributed to seismicity caused by continental
crust underthrusting the arc for 3–5 My (Brown
and Spadea, 1999).
In southern Alaska, where various arc and
microcontinental terranes have been accreted to
the active margin of North America since Jurassic
time (Plafker and Berg, 1994), sedimentary basins
that recorded the collisional history are exposed
subaerially, allowing exceptionally detailed facies
analysis and reconstruction of sediment-transport
pathways (Trop et al., 2002; Fig. 17.5). Syn-
collisional forearc basin deposits there include
the Upper Jurassic Naknek Formation, a >900-mthick belt exposed for 1200 km total length along
the margin, that recorded crustal-scale compression and exhumation of plutonic sources during
collision of the Talkeetna-Chitina arc either with
the North American continent or with smaller terranes during amalgamation of the Wrangellia Composite Terrane before its final accretion to North
America (Trop et al., 2005). Naknek Formation
facies indicate deposition on a steep basin floor,
358
Part 4: Convergent Margins
with submarine mass-flow conglomerates and
proximal fan-delta units transitioning to deeperwater turbidites and muddy pro-delta deposits
(Fig. 17.8). An upper unconformity reflects subaerial uplift and erosion following arc accretion.
In general, the source of sediment to basins
involved in arc-continent collision depends
greatly on the tectonics of the collision. In an
arc/passive-margin collision in which the forearc
faces the collision zone (as in Fig. 17.4), topography appears to be built largely from imbricated,
deformed, and metamorphosed sedimentary rocks
of the passive continental margin. This means that
syn- and post-collisional forearc-basin-filling sediment, such as the Paliwan Formation of Taiwan
(Dorsey, 1992) or the Rosroe Formation of South
Mayo, Ireland, is largely derived by erosion of the
deformed continental margin (Fig. 17.8). Sediment
provenance from South Mayo (Clift et al., 2009b)
shows that the arc itself must have formed only
modest topography, as in modern Taiwan. In contrast, when an arc collides by backing into an active
continental margin, such as in Honshu-Izu, Kohistan, or Talkeetna (Fig. 17.5), sedimentation is
dominated by erosion of the colliding and exhuming arc complex. Evidence from these regions suggests that the forearc basin of the oceanic arc can
remain relatively undeformed as it continues to
face an open ocean basin (Clift et al., 2000). Sedimentation in such a forearc basin usually continues, albeit with a change in provenance as the
composition of the arc volcanism changes, while
the backarc region (which faced the collision zone)
is strongly deformed and usually subducted (discussed below).
Other forearc-basin sedimentary sequences preserved in arc collision zones include the Sarhro
Group of Morocco, where folded greywacke turbidites and volcanic rocks several kilometers thick
record arc-continent collision in the Anti-Atlas
orogen ca. 660 Ma (Thomas et al., 2002), the Eocene
to Miocene Sarawaget Formation of the Finisterre
Range, Papua New Guinea (Abbott et al., 1994), and
possibly Upper Permian distal turbidites with arc
provenance representing a closing ocean basin
between the Altaids of southeastern Mongolia
and the North China microcontinent (Johnson
et al., 2008).
Intra-arc basins
Basins can form within the arc edifice, typically
covering less area and with shallower water depths
than forearc and backarc basins, and can be
“perched” between bathymetric highs (Smith
and Landis, 1995). The structural origin of intraarc basins, which can be bounded either by volcanic centers or by faults (Busby, 2004), is complex
and diverse among basins. These basins can form
by rifting, which may occur multiple times during
the activity of an intra-oceanic arc (as in the Alisitos accreted arc terrane of Baja California; Busby
et al., 2006) and may produce new oceanic crust
when opening intra-arc basins within a continental arc (as during the evolution of the complex
Guerrero Composite Terrane of Mexico; e.g.,
Centeno-Garcia et al., 2008). Other intra-arc basins
develop as transpressional or transtensional basins
caused by strike-slip faulting (Sarewitz and
Lewis, 1991) or large-scale block rotation within
the arc as plate motions change through time (Geist
et al., 1988). An impediment to preservation of
intra-arc basins is the fact that close to the arc
volcanic front, flexural rigidity is low because of
the lack of a major lithospheric root, such that
basins may be readily deformed, inverted, and
destroyed.
Intra-arc sedimentation consists of proximal volcanic and volcaniclastic material in debris aprons
derived from the arc massif that fine away from the
eruptive centers into turbidite and deep-water drift
facies (Fig. 17.3C; Draut and Clift, 2006). Proximal
Cretaceous intra-arc-basin facies exposed in Baja
California include silicic subaqueous pyroclastic
flow deposits, tuffs and tuff turbidites, hyaloclastite breccias, and evidence for mixing of basaltic
and silicic lavas with wet marine sediment; the
volcanic material in those basin deposits are intercalated with rudist reef deposits and marine mudstones, reflecting changes in relative sea level
during basin sedimentation (Busby et al., 2006).
During arc-continent collision, intra-arc basins can
fill with orogen-derived clastic material and, if
preserved, can record tectonic and geochemical
changes in the active margin during arc collision.
Huang et al. (1995) interpreted subaerially
exposed Pliocene and Pleistocene intra-arc basins
in Taiwan’s Coastal Ranges to have formed because
of strike-slip transtension faulting during active
arc-continent collision. These syn-collisional
basins are of similar size to those of modern
intra-arc basins not yet involved in collision
(1.5–10.0 km wide, 40 km long), and are predicted
to be short-lived as collision progresses between
the Luzon Arc and Eurasian margin, resulting in
their rapid inversion soon after formation and
Basins in Arc-Continent Collisions
359
Fig. 17.8. Field photographs of sediment deposited in arc collisional settings. (A) Submarine fanglomerates (clasts are
metasedimentary rocks) and (B) sandy turbidites from the early syn-collisional Paliwan Formation of Taiwan; these
sediments, which contain material eroded from a mature orogen in late Pliocene time, are exposed in the Coast Ranges.
(C) Conglomerates (most clasts are igneous) and (D) channelized sandstones of the early post-collisional Ordovician Rosroe
Formation, South Mayo, Ireland. Provenance of clasts in the Rosroe Formation is dominantly from the metamorphosed
continental passive margin, but with up to 20% influx from the accreted arc (Clift et al., 2009b). (E) and (F) show sedimentary
basin deposits of the syn-collisional Jurassic Naknek Formation, southeast Alaska (photos in E and F courtesy of J.M. Trop):
(E) Proximal fan-delta forearc strata (person at left for scale) dominated by poorly to moderately sorted, inversely graded
conglomerate (with granitic clasts whose Middle Jurassic isotopic ages match those of nearby Talkeetna arc plutons)
deposited by marine sediment gravity flows. (F) Distal pro-delta forearc strata (person for scale) characterized by thinly
bedded, sharp-based, normally graded sandstone turbidites (resistant units) and massive, black mudstone deposited by
suspension fallout. Sandstones are dominated by quartzofeldspathic detritus and granitic rock fragments. Mudstones yield
Jurassic ammonites, radiolarian, and foraminifera fossils.
360
Part 4: Convergent Margins
filling. Basin facies there include deep-water
flysch overlying shallow marine reef carbonate
atop volcanic basement, indicating formation by
rapid subsidence. These intra-arc basins developed, filled, and were inverted rapidly with sedimentation lasting a mere 0.8–3.1 My.
Backarc basins
Basins commonly occur behind intra-oceanic arcs
(Fig. 17.2), originating either from rifting and
extension after arc development (as has occurred
twice in the Mariana arc and twice in the Tonga arc;
Hawkins, 1974; Taylor, 1992) or as older ocean
floor that pre-dates the arc and subduction zone
(Karig, 1971; Taylor and Karner, 1983). Sedimentation in basins behind arcs is dominated by volcanic and volcaniclastic products of the arc, with
facies indicating increasing water depths with distance from the arc. Where backarc basins have
formed by arc rifting, products of silicic volcanism
can be abundant in the backarc stratigraphy (e.g.,
Izu-Bonin arc; Nishimura et al., 1992; Iizasa
et al., 1999; Fiske et al., 2001).
Backarc basins are unlikely to be preserved after
arc-continent collision. Arc-passive margin collision (the type of collision shown in Fig. 17.4 in
which the arc faces the continent) is usually followed by formation of a new subduction zone
outboard of the accreted arc, i.e., the former backarc basin becomes the locus of a new trench and
begins to be subducted. Examples of such a progression have been identified multiple times in the
geologic record (Dewey and Ryan, 1990;
Teng, 1990; Konstantinovskaia, 2001; van Staal
et al., 2007; Dickinson, 2008). Where oceanic
arcs collide with active margins via closure of a
backarc basin, as occurred during accretion of the
Jurassic–Cretaceous Alisitos arc terrane of Baja
California (Busby, 2004; Busby et al., 2006) and
the Jurassic Talkeetna-Chitina arc terrane of southern Alaska (e.g., Plafker and Berg, 1994; Fig. 17.5),
the backarc region is necessarily subducted before
collision even begins. Although much of the backarc basin and its sedimentary fill would thus be
destroyed, remnants of the backarc may be preserved as deformed, offscraped fragments within
the new suture zone. In the deforming backarc,
oceanic arc plutons may be exposed, eroded, and
material from them locally deposited as thick turbidites and conglomerates in slope fan complexes
(e.g., Ashigara Basin of Japan; Soh et al., 1998). In
such cases where arcs accrete onto active margins
by backarc-basin closure, a situation that requires
more than one subduction zone (or at least largescale thrusting and convergence between continent and backarc, in addition to the principal
subduction zone), there is substantial potential
for large-scale subduction and recycling of part
of the arc massif (e.g., Suyehiro et al., 1996;
Busby, 2004).
Despite the generally low preservation potential
of backarc basins, several accreted arc terranes do
contain kilometers-thick backarc stratigraphy (e.g.,
Fig. 17.3A, B). In the Alisitos accreted arc of Baja
California, the Jurassic Gran Canyon Formation
comprises volcanic and volcaniclastic backarc
material that directly overlies a supra-subduction-zone ophiolite (Kimbrough, 1984; Busby,
2004). With a hydrothermally altered base interpreted to reflect deposition on hot rifted arc crust,
this backarc sequence consists of deep marine
pyroclastic deposits dominated by lapilli tuff
and tuff breccia; proximal dacitic pyroclastic
flows formed graded beds tens of meters thick
(Busby, 2004). Multiple episodes of arc rifting
are inferred from these backarc sequences that
would have isolated the backarc from the active
arc front as the backarc topography broke up into a
series of fault-bounded blocks and basins. Backarc
deposition was capped by fine-grained epiclastic
volcaniclastic and lithic sandstone and siltstone as
the proximal sediment flux from the arc was interrupted by rifting (Busby, 2004).
Large accreted basins that had been situated
behind a Jurassic arc system also occur in southern
Alaska. Their sedimentary fill spans the change
from a pre- to a post-collisional setting over tens of
My (Trop and Ridgway, 2007). The Nutzotin and
Wrangell Mountains Basins (Trop et al., 2002;
Manuszak et al., 2007) were both situated behind
the intra-oceanic Talkeetna-Chitina arc and
received Middle to Late Jurassic arc-derived sediment; facies include marine volcaniclastic sandstone, mudstone, chert, and tuff. Similar
depositional environments (distal, deep-water
sedimentation behind the arc) were also inferred
for Lower Jurassic Talkeetna Formation rocks
exposed in the nearby Horn Mountains, although
in this latter case the transition from intra-oceanic
activity to collision is not well preserved (Clift
et al., 2005a). After the Talkeetna-Chitina arc
was incorporated into the Wrangell Composite
Terrane (WCT) and ceased activity (Late Jurassic
time), the Wrangell Mountains basin remained
active within the WCT and was positioned as a
Basins in Arc-Continent Collisions
forearc basin to the newly active Chisana arc
(between the Chisana and extinct older arc)
(Trop et al., 2002, 2005; Trop and Ridgway, 2007).
Sedimentary rocks in the Wrangell Mountains
Basin then recorded uplift and exhumation of
the Talkeetna-Chitina arc and the WCT. Concurrently, the Nutzotin Basin remained active as a
foreland basin behind the Chisana arc, accumulating 3 km of upward-coarsening sediment that
represents submarine fan and overlying marine
shelf deposition (Manuszak et al., 2007). (Note
that this basin predated the Chisana arc behind
which it sat; a basin’s position behind any given arc
does not require that that basin formed in the same
way as extensional backarc basins in the Mariana
and Tonga arcs or backarc basins of the Alisitos arc,
Baja California, that were the products of arc rifting
and seafloor spreading; Clift, 1995; Busby, 2004.)
Sedimentation in basins both behind and in front
of the Chisana arc continued to evolve through the
Cretaceous; the Nutzotin and Wrangell Mountains
basins were uplifted in Late Cretaceous to Tertiary
time, and folded coevally with major dextral faulting (Trop and Ridgway, 2007).
BASINS FORMED DURING
ARC-CONTINENT COLLISION
In addition to short-lived intra-arc basins forming
in an active arc-continent collision zone (Huang
et al., 1995), arc collisions cause foreland basins to
form by crustal loading, much as they would in a
continent-continent collision zone. These basins
involve lithosphere that is generally some of the
most thermally mature (i.e., cooler) and thus strongest in the oceans, with high flexural rigidity
resulting in broad basins. The major distinction
of arc-continent collision forelands is that because
orogenesis and subsequent collapse generally
occur rapidly (5–15 My; Abbott et al., 1997;
Friedrich et al., 1999), the foreland basins form,
fill quickly, and then may be inverted as the flexural load is reduced when an active continental
margin becomes established outboard of the collision zone. This is because re-establishment of
subduction is often accompanied by both extensional collapse of the collisional orogen and rapid
erosion, both of which combine to reduce the load
and lessen flexure. Foreland-basin formation is
ongoing in the collision between Australia and
Papua New Guinea, where the Leron Formation
comprises a fan delta complex at the base of the
361
Finisterre Range (Abbott et al., 1994). Similarly,
the Timor-Aru Troughs represent young foreland
basins formed by flexural loading of the Australian
continental margin as it collides with the Banda arc
(Snyder et al., 1996; Londo~
no and Lorenzo, 2004).
Because the Australian continental margin formed
in Jurassic time, it is relatively mature and rigid
compared to arc collision zones involving younger
crust resulting in a broad foreland basin.
Brown et al. (2006b) described the Southern
Urals accretionary complex as a thrust stack that
developed in a foreland basin on the continental
side of the Devonian arc-continent collision zone
in central Asia, which received sediment while at
the same time the old forearc basin was filled
from the same sources. Using Taiwan as the
type example, it is apparent that flexural foreland
basins reach their maximum extent and depth
at the height of collision, after the inversion
and closure of many (if not all) pre-existing arc
basins. They are filled by sediment derived across
strike from the orogen, and may contain both continental and shallow marine facies, with deep
marine flysch in the earlier stages of mountain
building, shallowing as the basin fills (Figs. 17.7
and 17.8).
As the active zone of oblique arc-continent collision moves along the margin (as in Taiwan), the
post-collisional orogen progressively collapses
during reversal of subduction polarity when the
compressive tectonic forces of collision are
removed. Collapse involves tectonic extension of
weak, hot, orogenic crust, and forms basins where
sedimentation can resume (Teng, 1996). Such
basins tend to form near the suture zone and do
not affect the broader foreland. However, the redistribution of mass away from a central edifice
might be expected to lessen the load on the subducting passive margin and potentially allow a
partial unflexing of the plate and inversion of the
foreland basin. The clearest modern example of a
post-collisional collapse basin is the Ilan Plain/
Okinawa Trough of northern Taiwan (Fig. 17.4). In
the Urals arc-continent suture, shallow-water carbonates unconformably overlying the arc units
record post-collisional collapse of the orogen
(Brown and Spadea, 1999). Likewise, after arccontinent collision in the Anti-Atlas orogen of
Morocco, inferred post-orogenic collapse (580–
550 Ma) led to extension forming faulted molasse
basins in which acidic volcanic rocks and volcaniclastic sediment were deposited (the Ouarzazate
Group; Thomas et al., 2002).
362
Part 4: Convergent Margins
CASE STUDY: SYNCHRONOUS BASIN
D ES T R U C T I O N A N D F O R M A T I O N
DURING ARC-CONTINENT COLLISION
A T T A I WA N
The very clear tectonic relationship in which the
oceanic Luzon arc is in a state of progressive and
migrating collision with the passive margin of
China (Suppe, 1984; Teng, 1990) reveals the different styles of sedimentary basins associated with
different phases of collision. It is important to note
that the pre-, syn- and post-collisional basins are
all filled by sediment eroded largely from the more
metamorphosed high mountain range in the orogenic core. Taiwan is one of the greatest modern
sediment-producing areas on Earth (Milliman and
Syvitski, 1992), a result of its rapid rates of rock
uplift, friable lithologies, and the erosive tropical
climate. Sediment is transported southward along
strike via the Manila Trench and the long axis of the
forearc (North Luzon Trough), bringing orogenic
sediment from the mountains into the pre-collisional arc region (Fig. 17.7). Similarly, sediment is
transported along strike by the Lanyang River,
itself guided by the Lishan Fault into the Okinawa
Trough, which is interpreted as a post-collisional
extensional collapse basin (Figs. 17.4 and 17.7;
Clift et al., 2008). Thus, because of substantial
along-strike sediment transport, the erosional history recorded by the basins around Taiwan (and
presumably other arc-continent collisional orogens) is not generally synchronous with the stage
of tectonic development in which the basin itself
formed.
As described above, the pre-collisional basins
have the lowest preservation potential, although
some remnant of the North Luzon Trough does
survive between the Central and Coast Ranges.
Indeed, fluvial and alluvial fan sedimentation continues in the Longitudinal Valley of Taiwan
between the two ranges, despite the rapid uplift
on either side. Because the deepwater sediment of
the North Luzon Trough is tectonically overthrust
and buried, those deposits may survive collision
and later be exhumed. Early syn-collisional sedimentary rocks of the Paliwan Formation
(Fig. 17.8B), composed of proximal eroded metasedimentary rocks of late Pliocene age, are exposed
in the Coast Ranges adjacent to the core of the
mountains (Dorsey, 1992). The metamorphosed
character of clasts in the Paliwan Formation
shows that they were eroded from a mature orogenic belt during the late Pliocene. Because
material effectively migrates north through the
system as the collision progresses, we infer that
these sedimentary rocks were deposited in a basin
south of the highest ranges at that time, but, like the
present system, contain material recycled from a
proto-Central Range source located to the paleoNorth. Since their deposition, the Paliwan conglomerates have been deformed and uplifted.
The largest sedimentary mass in the Taiwan
collision zone lies beneath the Taiwan Straits:
the foreland basin, where the Oligocene South
China Sea rifts have been buried under 2 km of
sediment, locally >5 km thick (Lin et al., 2003).
The flexural rigidity of the margin lithosphere is
not particularly high (Te ¼ 8–13 km; Lin and
Watts, 2002), resulting in a relatively narrow
basin (70 km), but one with a high preservation
potential. In contrast, sediment eroded from the
Coast Range that is deposited in the backarc basin
of the Luzon Arc (i.e., the Huatung Basin of the
Philippine Sea) is likely to be subducted and lost as
subduction polarity reverses. The Huatung Basin
(Fig. 17.7) receives significant sediment flux from
the orogen but this crust is in turn subducted under
the newly created continental margin of the Ryukyu arc. Although some material may be offscraped
into the new accretionary prism of the continental
margin, the majority will likely be lost and subducted to depth under the arc (Clift and
Vannucchi, 2004).
SUMMARY
Arc-continent collisions occur commonly in the
plate-tectonic cycle and result in rapidly formed
and rapidly collapsing orogens, often spanning
just 5–15 My. Accretion of new material onto continents via arc-continent collision (accompanied
by siliceous magmatism and loss of dense, mafic
lower crust) is thought to be a major process governing evolution of the continental crust through
time. Rapid uplift rates at arc-continent collision
zones generate substantial amounts of sediment,
which is transported along the colliding margin in
pre- and post-collisional basins, and in the foreland. Collisions of intra-oceanic arcs with passive
continental margins (a situation in which the arc,
on the upper plate, faces the continent) involve
substantially different geometry than collisions of
intra-oceanic arcs with active continental margins
(a situation requiring more than one convergence
zone and in which the arc, on the lower plate, backs
Basins in Arc-Continent Collisions
into the active continental margin), with variable
preservation potential for basins in each case. In
the latter situation, there is substantial potential for
large-scale subduction and recycling of part of the
arc massif.
Substantial differences are also expected
between trench and forearc evolution (and associated sedimentation) in tectonically erosive versus
tectonically accreting margins, both before and
after collision. During intra-oceanic arc activity
before collision, the greater water depths that typify erosive margins lead to slow formation of deeper-water sedimentary facies than occur in
accretionary margins. The preservation potential
of trench-slope basins is very low; in collision they
are rapidly uplifted and eroded, and at erosive
margins especially in intra-oceanic subduction
they are progressively destroyed by subduction
erosion. After arc-continent collision, preservation
of (deformed and, if underplated, metamorphosed)
trench sediment and trench-slope basins is biased
toward margins that were tectonically accreting for
a substantial length of time before collision.
Forearc basins in tectonically erosive oceanic
arcs are usually floored by strong lithosphere
and may well survive collision with a passive
margin, sometimes continuing sedimentation
throughout collision and orogeny, driven by eclogite formation in the underlying slab. Sedimentation rates in surviving forearc basins can be rapid;
these basins contain material derived from the
deformed and metamorphosed passive margin
and exhumed accreted arc, and are characterized
by turbiditic, mass-wasted, and sometimes fan
facies. Forearc basins perched on weak, deformed,
accretionary-complex basement are more likely to
be destroyed in collision zones because they are
more likely to be deformed, uplifted, and eroded.
The low flexural rigidity of intra-arc basins makes
them deep and, if preserved, potentially long
records of arc and collisional tectonism. Whereas
pre-collisional forearc and intra-arc basins may be
filled and then survive the collision, backarc basins
are typically subducted and the sedimentary cover
either lost or preserved only as fragments in
m
elange sequences.
A substantial proportion of the sediment derived
from collisional orogenesis ends up in a foreland
basin that forms as a result of collision, and may be
preserved largely undeformed. Compared to continent-continent collisional foreland basins (e.g.,
Zagros and Himalaya), arc-continent collisional
basins are short-lived and may experience partial
363
inversion after collision as the orogen collapses
during the establishment of a new, active continental margin outboard of the collision zone.
ACKNOWLEDGMENTS
We thank Jeff Trop for insightful discussions on the
tectonic evolution of Alaska. Cathy Busby, Ken
Ridgway, Holly Ryan, and Dave Scholl are thanked
for their helpful and constructive reviews of this
manuscript.
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