Download Formation of plate boundaries: The role of mantle volatilization

Earth-Science Reviews 129 (2014) 85–99
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Earth-Science Reviews
journal homepage: www.elsevier.com/locate/earscirev
Formation of plate boundaries: The role of mantle volatilization
Tetsuzo Seno a,⁎, Stephen H. Kirby b
a
b
Earthquake Research Institute, University of Tokyo, Bunkyo-ku, Tokyo, 113-0032, Japan
U.S. Geological Survey, 345 Middlefield, Menlo Park, CA 94025, USA
a r t i c l e
i n f o
Article history:
Received 23 August 2013
Accepted 18 October 2013
Available online 6 December 2013
Keywords:
Plate tectonics
Initiation of subduction
Plate boundaries
Hydration
Carbonation
Serpentine
a b s t r a c t
In the early Earth, convection occurred with the accumulation of thick crust over a weak boundary layer
downwelling into the mantle (Davies, G.F., 1992. On the emergence of plate tectonics. Geology 20, 963–966.).
This would have transitioned to stagnant-lid convection as the mantle cooled (Solomatov, V.S., Moresi, L.-N.,
1997. Three regimes of mantle convection with non-Newtonian viscosity and stagnant lid convection on the
terrestrial planets. Geophys. Res. Lett. 24, 1907–1910.) or back to a magma ocean as the mantle heated (Sleep,
N., 2000. Evolution of the mode of convection within terrestrial planets. J. Geophys. Res. 105(E7): 17563–
17578). Because plate tectonics began operating on the Earth, subduction must have been initiated, thus avoiding
these shifts. Based on an analogy with the continental crust subducted beneath Hindu Kush and Burma, we propose that the lithosphere was hydrated and/or carbonated by H2O–CO2 vapors released from magmas generated
in upwelling plumes and subsequently volatilized during underthrusting, resulting in lubrication of the thrust
above, and subduction of the lithosphere along with the overlying thick crust. Once subduction had been initiated, serpentinized forearc mantle may have formed in a wedge-shaped body above a dehydrating slab. In relict
arcs, suture zones, or rifted margins, any agent that warms and dehydrates the wedge would weaken the region
surrounding it, and form various types of plate boundaries depending on the operating tectonic stress. Thus, once
subduction is initiated, formation of plate boundaries might be facilitated by a major fundamental process: weakening due to the release of pressurized water from the warming serpentinized forearc mantle.
© 2013 Elsevier B.V. All rights reserved.
Contents
1.
2.
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Subduction initiation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
2.1.
Subduction of the continental lithosphere beneath Hindu Kush and Burma . . . . . . . . .
2.2.
The Reunion and Kerguelen hotspots . . . . . . . . . . . . . . . . . . . . . . . . . .
2.3.
Hydration/carbonation of oceanic lithosphere . . . . . . . . . . . . . . . . . . . . . .
2.4.
Collision versus subduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
2.5.
Subduction initiation in the Archean . . . . . . . . . . . . . . . . . . . . . . . . . .
3.
Serpentinized forearc mantle wedge and its role in the formation and mobilization of plate boundaries
3.1.
Divergent boundaries . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.2.
Convergent boundaries . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.3.
The San Andreas fault system (SAF) . . . . . . . . . . . . . . . . . . . . . . . . . . .
3.4.
Other continental and oceanic transform faults . . . . . . . . . . . . . . . . . . . . . .
4.
Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
5.
Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Acknowledgments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
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1. Introduction
⁎ Corresponding author: Tel.: +81 3 5841 5747; fax: +81 3 5802 3391.
E-mail address: [email protected] (T. Seno).
0012-8252/$ – see front matter © 2013 Elsevier B.V. All rights reserved.
http://dx.doi.org/10.1016/j.earscirev.2013.10.011
At present, plate tectonics operates on the Earth. This was probably
not the case during the magma ocean stage (Warren, 1993; Sleep,
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T. Seno, S.H. Kirby / Earth-Science Reviews 129 (2014) 85–99
2000). Paleomagnetic, geochemical, and isotope studies suggest that
plate tectonics had started earllier than the Archean/Proterozoic boundary (Burke et al., 1976; Campbell and Griffiths, 1992). However, even if
we accept that it had started by that time, it is not easy to describe how
it began. Because of the strongly temperature-dependent viscosity of
the mantle material (e.g. Kirby, 1983), the lithosphere is not easy to rupture. A number of studies show that subduction initiation requires high
levels of stress to disrupt the lithosphere, which is a few tens of kms
thick (e.g. McKenzie, 1977; Cloetingh et al., 1984; Mueller and Phillips,
1991). Meanwhile, the largest available plate tectonic forces are of the
order of 1013 N/m (Molnar and Gray, 1979; Parsons and Richter, 1980;
England and Molnar, 1991; Fleitout, 1991); these forces are insufficient
to cause top-to-bottom disruption of a plate. Strain-weakening is thus
required to form new plate boundaries (e.g. Bercovici, 1998; Moresi
and Solomatov, 1998; Tackley, 1998; Regenauer-Lieb et al., 2001;
Tommasi and Vauchez, 2001; Bercovici, 2003). Because plate tectonics
does not operate on Venus or Mars, it is inferred that the weakening
mechanisms should include effects of volatiles like H2O and CO2 (e.g.
Bercovici, 1998; Regenauer-Lieb et al., 2001; Solomatov, 2004; O'Neill
et al., 2007). Their effects on the thermo-mechanical properties of the
lithosphere have been considered in numerical simulations of the initiation of plate boundaries (e.g. Bercovici, 1998; Regenauer-Lieb et al.,
2001). They are, however, in general, not in the context of the formation
of each type of plate boundary during the Earth's tectonic history.
Eclogitization of granulites induced by fluid injection might have played
an important role in deforming the lithosphere (Austrheim et al., 1997;
Bjornerud et al., 2002) and could be an example of the effects of
volatiles. However, it is not yet clear how eclogitization played a role
in the initiation of plate tectonics.
Another difficulty in initiating plate tectonics is the production of a
large amount of basaltic crust at the surface by upwelling of the mantle.
This occurred due to the high mantle potential temperature during the
Archean (Sleep and Windley, 1982; Bickle, 1986; McKenzie and Bickle,
1988; Vlaar et al., 1994). The buoyant crust does not cool enough,
which prevents subduction and leads to heating-up of the Earth's
mantle and a shift back to a magma ocean (Sleep, 2000).
In this paper, we present a scenario where H2O and CO2 play a role in
the initiation of plate boundaries in the context of realistic tectonic situations. We propose that subduction was initiated by volatilization of
the underthrusting lithosphere that was once hydrated or carbonated
by upwelling plumes. The ultramafic minerals that constitute the lithospheric mantle are affected by vapors rich in H2O–CO2 released from solidifying magmas in a plume head, and are altered to minerals such as
amphibole, phlogopite, dolomite, magnesite, chlorite, and serpentine
(Menzies et al., 1987; Spera, 1987; Wyllie, 1988). If such hydrated/
carbonated lithosphere approaches a convergent zone and is underthrust, volatilization from it would occur due to increases in temperature and pressure. The released volatiles would lubricate the thrust
above, and make subduction initiation possible. This inference is obtained from the fact that subduction of the lithosphere with thick continental crust is now occurring beneath Hindu Kush and Burma according to
intermediate-depth seismicity observed there (Seno and Rehman,
2011), although the nature of the crust is chemically different from
that in the Archean. The lower portion of the lithosphere in these regions is likely to have been hydrated/carbonated by plumes when it
passed over the Reunion and Kerguelen hotspots. This will be described
later in more details.
Once subduction started, formation of serpentinized forearc mantle
in a wedge-shaped body above a dehydrating slab followed. Kirby
et al. (2003, 2013) proposed that the San Andreas fault system (SAF)
was mobilized by weakening of the warming and dehydrating
serpentinized forearc mantle, following the cessation of subduction of
the Farallon plate. Serpentinized wedge mantle formed in a forearc
may be found in certain tectonic settings, such as relict arcs, suture
zones, and rifted margins. If the wedge is warmed by any agent, the resultant released pressurized water will weaken the region surrounding
the wedge, and will form various types of plate boundaries, depending
on the stress regime. In the latter half of this paper, we show application
of this weakening mechanism to the creation of each type of plate
boundary during the stage after subduction had been initiated.
2. Subduction initiation
In this section, we treat the case where we consider that the initiation of subduction has not occurred. In uniform-viscosity convection,
both convergent and divergent motions occur at the surface. This may
mimic the convection that occurred during the early stages of the
Earth with a weak boundary layer sinking into the mantle (Fig. 1a and
b, Davies, 1992). However, in these conditions, crust thicker than the
modern one would form over the lithosphere (Sleep and Windley,
1982; McKenzie and Bickle, 1988). Such crust results from primordial
melt production of the mantle having a basaltic composition and a density of ~ 2900 kg/m3 (McKenzie and Bickle, 1988). It is still prevented
from subduction due to the plate being thin during the Archean
(Davies, 1992; See also Cloos, 1993). At convergent margins, the crust
is offscraped, thrust upward, and juxtaposed over crustal slivers, as in
modern collision zones (Mattauer, 1986), with the mantle boundary
layer downwelling alone, leaving the surface tectonics behaving unlike
plate tectonics (Davies, 1992; Fig. 1a and b). In such a tectonic situation,
if the average Earth's surface heat flow is larger than the radioactive
heat production, the mantle is cooled. If the boundary layer is too strong
to be subducted, there occurs a shift to stagnant-lid convection like that
on Venus (Solomatov and Moresi, 1996, 1997; Sleep, 2000). Conversely,
if the average surface heat flow is smaller than the radioactive heat
production, the mantle becomes hotter, and large-scale melting of the
mantle (i.e. a magma ocean) resumes (Sleep, 2000). Therefore, for
plate tectonics to start on the Earth, the lithosphere must be subducted
along with its thick crust as a whole, avoiding the shift to stagnant-lid
convection or to a magma ocean. However, it seems unlikely to occur
in preference to the buoyant crust piling up at the surface.
2.1. Subduction of the continental lithosphere beneath Hindu Kush and
Burma
We propose a mechanism that describes the start of plate tectonics
on the Earth. The locations where continental lithosphere with thick
crust subducts into the modern Earth give us a clue. Hindu Kush and
Burma provide evidence of the mechanism (See Rehman et al., 2011
and the references therein for the geologic history of this region, and
see Seno and Rehman, 2011 for the tectonic setting). Hindu Kush is
located within the Asian plate, southwest of Pamir and west of
Karakorum. To the south, there is the Mesozoic Kohistan–Ladakh arc,
which was welded to Asia during the late Mesozoic. India later collided
with this arc in the early Tertiary. The geological terranes associated
with the collision between India and Kohistan–the high Himalaya, the
lesser Himalaya, and the Siwaliks–are similar to those in other
Himalayan regions where India and Asia collided. However, the widths
of the high and lesser Himalayas are only ~1/2 of those in other regions.
This indicates that offscraping and accretion of the underthrusting
Indian continental margin crust have had a smaller extent south of
Kohistan.
The plate boundary where the Indian plate is currently underthrust
is located south of Kohistan (Searle et al., 2001). To the north,
intermediate-depth seismicity dips steeply northward to a depth of
~ 300 km beneath Hindu Kush and dips southward to a depth of
~ 150 km beneath Pamir (e.g. Billington et al., 1977; Searle et al.,
2001). Studies of this seismicity using focal mechanisms and seismic tomography (e.g. Van der Voo et al., 1999; Pavlis and Das, 2000) show that
the seismic zone beneath Hindu Kush is contiguous to that beneath
Pamir. Because India has been colliding against Kohistan since the
early Tertiary, no oceanic plate is recently subducting beneath this region, as pointed out by Searle et al. (2001). It is thus difficult to relate
T. Seno, S.H. Kirby / Earth-Science Reviews 129 (2014) 85–99
(a)
Hydrated/carbonated
lithosphere
(b)
Toroidal field
Convergent zone
Divergent
zone
Divergent zone
87
Hydrated/carbonated
lithosphere
Convergent zone
Thick crust
Plume
Plume
Weak boundary layer
Toroidal field
Toroidal field
Divergent
zone
Divergent zone
Volatilizing
lithosphere
Subduction
(f)
Dehydration
boundary
Plume
Supercontinent
Plume
Subduction zone
Warming serpentinized
mantle
Hydrated/carbonated
lithosphere
Convergent zone
Volatilizing
lithosphere
Toroidal field
(e)
(d)
Convergent zone
Hydrated/carbonated
lithosphere
Transforrm fault
Mid-ocean
ridge
Subduction zone
(c)
Toroidal field
Fig. 1. Schematic illustration showing the processes of formation of plate boundaries. All figures are in a plan view except (b) and (d). (a) Flow fields for convection with a weak boundary
layer in the early stage of the Earth before plate tectonics began. Divergent and convergent zones are represented by pink and gray ellipses, respectively. The portion of the lithosphere that
is hydrated/carbonated by a plume is shown in green. Features like mid-ocean ridges (MOR) appear in divergent zones, arising due to weakening by melting (Sleep, 2000). Toroidal motions occur in a diffuse way (upper and lower). (b) Convection with a weak boundary layer and thick basaltic crust (a cross-section view, modified from Davies, 1992). A weak boundary
layer is downwelling into the mantle in the convergent zone. Thick crust (brown color) is thrust over each other and piles up over the downwelling because it is difficult to subduct this
material (Davies, 1992). The portion of the lithosphere that is hydrated/carbonated by a plume is shown in green. The divergent boundary is marked by a red triangle, which indicates a
magma chamber. (c) Lithosphere that was earlier hydrated/carbonated by a plume (green color) is underthrusting the convergent zone. A zone of volatilization from the underthrusting
lithosphere is shown in blue. Other symbols are the same as (a). (d) Lithosphere that was earlier hydrated/carbonated by a plume (green color) is underthrust in the convergent zone (a
cross-section view). Volatilization from the underthrusting lithosphere (blue color) lubricates the thrust, making subduction possible. Other symbols are the same as (b). (e) Serpentinized
forearc mantle is embedded in a suture zone of a supercontinent (left) and in a rifted margin or relict arc (upper). Any thermal event like an upwelling plume shrinks the serpentinized
forearc mantle by dehydration, which weakens its boundaries by the release of pressurized water (see text). If the supercontinent overrides a divergent zone, drag forces at the base of the
continent might lead to breakup of the continent. If a shearing tectonic stress is operating in the relict arc or rifted margin, they might turn into a transform fault. (f) Once a complete set of
the three types of plate boundaries forms via the weakening processes shown in (c), (d) and (e), plate tectonics starts to operate.
the intermediate-depth seismicity, which is a phenomenon that has
only occurred within the past several Ma, with subduction of an oceanic
plate. The continental crust is thus unusually subducting in this region.
In western Burma, there is a belt of high-grade metamorphic rocks
with Mesozoic ophiolites (e.g. Maurin and Rangin, 2009). This ophiolite
belt was regarded as an eastward continuation of the Indus-Zangpo suture by Mitchell (1984). In Tibet, south of the ophiolite belt, the Tethys
and high Himalayas, i.e., the Mesozoic and Paleozoic rocks offscraped
from the Indian continental margin, display collision characteristics.
However, these characteristics are absent in Burma. Instead, to the
west of the ophiolite belt, there is the Indo-Burman Wedge, i.e. a Tertiary accretionary complex growing to the southwest (Maurin and Rangin,
2009). The Miocene-Quaternary calcalkaline volcanics are found to the
east of the ophiolite belt (Mitchell, 1984). The Quaternary volcanic
rocks show isotope and trace element patterns that indicate subduction
of the Indian continental lithosphere (Zhou et al., 2012). There is
intermediate-depth seismicity beneath Burma down to a depth of
200 km (e.g. Le Dain et al., 1984). These characteristics suggest that, in
the Burman arc, subduction, rather than collision, has been occurring
during the late Cenozoic. Because the Bay of Bengal and the Bengal
Basin to the north have a crustal thickness N25 km (Brune and Singh,
1986; Kaila et al., 1992), Burma is an anomalous place where continental crust is subducting with intermediate-depth seismicity, similar to
Hindu Kush.
2.2. The Reunion and Kerguelen hotspots
The unusual subduction of the Indian continental lithosphere beneath Hindu Kush and Burma is in marked contrast to the collision in
other parts of the Himalayas. Volcanics produced by the Reunion and
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T. Seno, S.H. Kirby / Earth-Science Reviews 129 (2014) 85–99
solidify near the base of the lithosphere (Fig. 3a, see also Figs. 45 and 46
of Wyllie, 1988). This would alter mantle minerals at depths b 100 km
above the plate base by permeation through pores (Mibe et al., 1999),
producing amphibole at T b 1050 °C, chlorite at T b 700 ~ 800 °C, and
serpentine/talc at T b 600–700 °C (e.g. Poli and Schmidt, 1995; Ulmer
and Trommsdorff, 1995; Wunder and Schreyer, 1997; Bose and
Navrotsky, 1998; Iwamori, 1998; Bromiley and Pawley, 2003;
Fumagalli and Poli, 2005; Till et al., 2012), in addition to other
peridotite–CO2 minerals at similar temperature–depth ranges (Brey
et al., 1983; Wallace and Green, 1988; Falloon and Green, 1989; Canil
and Scarfe, 1990; Dasgupta et al., 2007). As noted in the introduction
Kerguelen hotspots during the Mesozoic-Tertiary are distributed
forming the traces of these hotspots on the Indian plate (Fig. 2). In accordance with the peculiarity mentioned above, Seno and Rehman
(2011) found that the traces of the Reunion and Kerguelen hotspots
on the Indian plate, formed at the time of ~100–126 Ma (Storey et al.,
1989; Muller et al., 1993; Kent et al., 2002), are located in the vicinity
of the intermediate-depth seismicity beneath Hindu Kush and Burma
(Fig. 2). The Indian continental lithosphere in these regions is thus likely
to have been affected by plumes upwelling beneath these hotspots. Following Wyllie (1988), we expect that uprising plumes with partial
melts involving H–O–C release these volatiles as vapors when magmas
E
N
Hotspot-related volcanics
Reunion
Predicted hotspot traces
Kerguelen
Muller et al.(1993)
Kent et al.(2002)
90
100
Storey et al.(1989)
84
80
150
300
km
80.2
S. Tethyan suture
zone volcanics
Intermediate-depth
seismicity
Depth
83.5
72-73 Ma
Rajmahal Traps
73.6
115
118
68.5
68.5
115-126 Ma
Lamprophyres
Alkali intrusions
Deccan Traps
110
100
90
84
58.6
Ma
Reunion
80.2
Ma
Kerguelen
Fig. 2. Spatial relation between the Reunion and Kerguelen hotspot traces and the intermediate-depth seismicity beneath Hindu Kush and Burma. This provides an example of
possible hydration/carbonation of the continental lithosphere by upwelling plumes (Seno and Rehman, 2011). The epicenters of intermediate-depth earthquakes (Inter. Seism. Centre
(depth N 80 km, 1983–2004, M N 4.0)) are shown by green dots. Volcanics related to the Reunion hotspot, i.e. the Deccan flood basalts, the alkali intrusions ~3 M.y. older than the Deccan
(triangles, Basu et al., 1993), and the ~80 Ma South Tethyan suture zone volcanics in Pakistan (Mahoney et al., 2002), are shown in dark-brown. Volcanics related to the Kerguelen hotspot,
i.e. the Ninety-east Ridge, its northern extension that is buried beneath the sediments in the Bengal Fan (Curray et al., 1979), the Rajmahal Traps and the lamprophyres (triangles) in NE.
India (120–115 Ma), are shown in orange. The hotspot traces that are located using Muller et al.'s (1993) solution for the Indian plate motion over the Reunion and Kerguelen hotspots
younger than 100 Ma are shown by the blue lines with ages. The 90–100 Ma trace for the Reunion hotspot is located at the southern edge of the intermediate-depth seismicity in Hindu
Kush. The 115–126 Ma (Kent et al., 2002) and 110 Ma traces (Storey et al., 1989) for the Kerguelen hotspot are shown by large and medium-size blue stars, respectively. These are located
within the belt of the intermediate-depth seismicity in Burma.
T. Seno, S.H. Kirby / Earth-Science Reviews 129 (2014) 85–99
89
(a) Continental lithosphere with a plume below
Continental
lithosphere
Hydrated/carbonated
zone
Partial melt
Plume
Lubricated megathrust
(b) Continent-continent convergent zone
Upper crust
Lower crust
Hydrated/carbonated
zone
Volatilization
Subduction
initiation
Fig. 3. (a) Hydration/carbonation of continental lithosphere by an upwelling plume (modified from Fig. 45 of Wyllie, 1988). Lines 2 and 3 represent crossing points between the geotherm
and the solidus for peridotite–C–H–O. Within the plume, melting occurs at greater depths due to its higher temperature. Line 1 represents the lithosphere–asthenosphere boundary, above
which partial melts accumulate and form magmas. Vapors of CO2 and H2O are released from the cooling magmas and hydrate or carbonate the lower portion of the lithosphere (green
color) by permutation or hydrofracturing (see text). (b) Volatilization of hydrated/carbonated continental lithosphere in a continent–continent convergent zone (after Seno and Rehman,
2011). Generally, when continental lithosphere converges, most of the crust is offscraped, and the delaminated boundary layer sinks. If the lower portion of the lithosphere is hydrated/
carbonated by a plume as in (a) (green color, Wyllie, 1988), it heats up and volatilizes whilst being underthrust to intermediate-depths. Released volatiles (blue color) migrate upward and
reduce frictional resistance to underthrusting. This would allow the continental lithosphere to be subducted as a whole (black arrow). Intermediate-depth seismicity is produced in the
subducted lithosphere by devolatilization embrittlement of the previously hydrated/carbonated mantle. This stage corresponds to Fig. 1d.
section, we call these chemical alternation processes of lithospheric
mantle minerals “hydration and/or carbonation” in this paper. If
the lithosphere is in tensional tectonic stress, magmas, and vapors
might further migrate upward by magma-hydrofracturing, inducing
hydration/carbonation in the shallower portion (see Fig. 46 of Wyllie,
1988; Spera, 1987; Wilshire and Kirby, 1989). Fig. 3a schematically illustrates this hydration/carbonation, adapting Fig. 45 of Wyllie (1988). Because the solidus at level 3 in Fig. 3a is assumed to be ~ 1300 °C (see
Fig. 20 of Wyllie, 1988), the altered minerals are stable in the lithosphere when they are more than a few tens of kms above the solidus
(green shaded in Fig. 3a). The lower portion of the Indian continental
lithosphere, once located above the hotspots, is thus likely to have
been hydrated/carbonated by plumes.
2.3. Hydration/carbonation of oceanic lithosphere
This kind of hydration/carbonation by plumes, albeit for oceanic lithosphere, has been invoked by Kirby (1995) and Seno and Yamanaka
(1996) to explain the occurrence of double-seismic zones in the
circum-Pacific subduction zones. Based on the devolatilization embrittlement mechanism for intraslab intermediate-depth earthquakes
(Raleigh and Paterson, 1965; Nishiyama, 1992; Kirby et al., 1996; Seno
and Yamanaka, 1996; Jung et al., 2004; Takahashi et al., 2009),
they interpreted double seismic zones to represent volatilization of
slabs due to increasing temperature and pressure during subduction.
The upper plane of double seismic zones can be explained by
devolatilization embrittlement of altered basalts in the subducting
oceanic crust (Kirby et al., 1996; Peacock and Wang, 1999; Hacker
et al., 2003; Yamasaki and Seno, 2003; Kita et al., 2006; Nakajima
et al., 2009). The lower plane requires the deeper half of the oceanic
lithosphere to be hydrated/carbonated prior to subduction. Kirby
(1995) and Seno and Yamanaka (1996) attributed this to be the
carbonatization and hydration, respectively, similar to that depicted in
Fig. 3a, by H2O–CO2 vapors released from magmas in plumes rising up
in the Pacific Ocean. Although direct evidence for this metamorphism
is scarce, Seno and Yamanaka (1996), Hacker et al. (2003), and
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Yamasaki and Seno (2003) showed that the dehydration locus of
serpentinites formed in the lower portion of subducting lithosphere is
consistent with the observed geometry of the double seismic zone; on
the contrary, the locus formed by infiltration of seawater in the upper
half of oceanic lithosphere by normal faulting (Peacock, 2001; Ranero
et al., 2005; Faccenda et al., 2009) is not consistent.
2.4. Collision versus subduction
It is generally believed that, when continental lithosphere approaches a convergent boundary, subduction is hindered due to the
thick buoyant crust, and therefore, collision generally occurs (e.g.
McKenzie, 1969; Molnar and Gray, 1979; van den Beukel, 1992; Cloos,
1993). However, Seno (2008) showed that buoyancy is not an essential
factor in preventing subduction, but absence of dehydration from a slab
is rather important. Continental crust does not contain metamorphic
minerals that dehydrate at intermediate-depths (Ernst et al., 1998),
which makes both pore fluid pressure ratio in the thrust and
intermediate-depth seismicity low (Seno and Yamasaki, 2003; Seno,
2007), resulting in offscraping and piling up of the crust in collision
zones (Seno, 2008). Conversely, in subduction zones, dehydration of
the slab crust and mantle (Anderson et al., 1976; Peacock, 1993; Poli
and Schmidt, 1995; Okamoto and Maruyama, 1999; Peacock and
Wang, 1999; Seno et al., 2001; Hacker et al., 2003; Fumagalli and Poli,
2005; Kameda et al., 2011; Kuwatani et al., 2011; Till et al., 2012) results
in pore fluid pressure ratios as high as 0.90–0.98 (Wang and He, 1999;
Wang and Suyehiro, 1999; Seno, 2009; Kimura et al., 2012), which render smooth subduction with very low friction at the thrust.
The unusual subduction of the continental crust beneath Hindu Kush
and Burma should be closely related to the existence of intermediatedepth seismicity in these regions. As mentioned before, the lower portion of the Indian lithosphere could be hydrated/carbonated by plumes
uprising beneath the Reunion and Kerguelen hotspots, similar to
oceanic lithosphere with a double seismic zone. After traveling over
thousands of kms to the north, the hydrated/carbonated lithosphere approaches and is underthrust beneath Asia (Fig. 2), and volatilization occurs from the lithosphere due to increases in temperature and pressure.
This would result in intermediate-depth seismicity by devolatilization
embrittlement and the released volatiles would lubricate the thrust
above, enabling subduction under Asia (Fig. 3b). The intermediatedepth seismicity beneath these regions is essentially the same as that
in the western Pacific (e.g. Billington et al., 1977; see other references
cited in Seno and Rehman, 2011), which implies that the extent of hydration/carbonation of the slab is similar to that of subducting slabs in
the circum-Pacific. Because the diameter of the temperature anomaly
and related volcanics of an uprising plume is generally of the order of
1000 km (White and McKenzie, 1989; see also the spatial dimension
of the volcanics related to the Reunion and Kerguelen hotspots shown
in Fig. 2), the induced subduction is not spot-like, but can be in a zone
extending over several hundreds of km. Although direct evidence of
subduction of hydrated/carbonated continental lithosphere in these regions is not yet available, Sumino et al. (2011) noted OIB signatures in
the 3He/4He ratio of diamond exhumed in the ultra-high pressure metamorphic belt of northern Kazakhstan, as an example of such subduction.
The Quaternary volcanic rocks in Burma also show isotope and trace
element patterns indicating subduction of the Indian continental lithosphere there (Zhou et al., 2012), as mentioned before.
2.5. Subduction initiation in the Archean
We propose that subduction initiation in the Archean occurred in a
similar manner, although, in this case, the thick crust that prevented
subduction is basaltic. Hydration/carbonation of the lithosphere by
plumes, as shown in Fig. 3a, would have occurred more frequently during the Archean, because plume activity was much higher at that time
(Campbell and Griffiths, 1992; Kroner and Layer, 1992). Conversely,
the thickness and thermal state of the lithosphere were not much different from those at present (Bickle, 1978; Burke and Kidd, 1978; Jarvis
and Campbell, 1983; England and Bickle, 1984; Bickle, 1986), due to
the efficient removal of heat from the Earth's interior by the high plate
creation rate at that time. When such lithosphere approached and was
underthrust in the Archean convergent zones, as in Hindu Kush and
Burma, increases in temperature and pressure caused by underthrusting would have induced volatilization from the hydrated/carbonated
lithosphere, and the volatiles thus released would have lubricated the
thrust (Fig. 3b). If the subducted crust was transformed to denser
eclogite as shown by Austrheim et al. (1997), it would have aided further underthrusting of the lithosphere (Ringwood and Green, 1966;
Ahrens and Schubert, 1975). Eclogitization of the lower crust in the vicinity of the thrust zone also enhanced its weakness by the interplay between fluid injection, chemical reaction, and deformation (Austrheim
et al., 1997; Bjornerud et al., 2002) and made subduction easier by overcoming the energy dissipation associated with plate bending (Conrad
and Hager, 1999). These together would have made such lithosphere
with thick crust enter the deeper mantle, resulting in subduction
initiation (Figs. 1d and 3b).
3. Serpentinized forearc mantle wedge and its role in the formation
and mobilization of plate boundaries
In uniform-viscosity convection, divergence can occur by ductile extension of a boundary layer, without any mid-ocean ridge (MOR) like
feature (Davies, 1992). Toroidal motions, if any, can be accommodated
by diffuse deformation between the loci of convergence and divergence
(Fig. 1a and c; see also Fig. 1 of Bercovici, 1996), rather than by transform faults (Fig. 1f). As an exception, partial melting in the divergence
zone may weaken the boundary layer and form a linear divergent
boundary with a MOR-like feature (Fig. 1a and c, Sleep, 2000). Other
types of plate boundaries, however, do not form easily in the lithosphere
as it cools and becomes strong. In this section, we propose the following
mechanism to cause strain localization, which leads to the formation of
new plate boundaries in such strong lithosphere (Fig. 1e and f).
Initiation of subduction discussed in the previous section may form
serpentinized forearc mantle in subduction zones. In the modern
Earth, recent geophysical evidence shows that serpentinized mantle
occurs in cold, shallow, wedge-shaped regions of forearcs above
dehydrating slabs in several well-studied subduction zones (Kamiya
and Kobayashi, 2000; Bostok et al., 2002; Brocher et al., 2003;
Hyndman and Peacock, 2003; Blakely et al., 2005; Matsubara et al.,
2005; Seno, 2005; Wang et al., 2006). Shallow serpentinized forearc
mantle forms because water is released into the forearc from
downgoing, warming, and thus dehydrating slab crust and mantle,
and also because the descending slab cools the forearc mantle nearest
the trench and slab, resulting in a large region where serpentinite is stable (e.g. Hyndman and Peacock, 2003). The volume of water potentially
stored in the forearc is huge, about 200 km3/km of margin or more,
depending on the degree of serpentinization of the forearc mantle
(Kirby et al., 2013). The inboard and outboard wedge boundaries of
the serpentinized mantle are controlled by the thermal structure of
the forearc, the plate interface, and the forearc crust–mantle boundary.
If the subducting slab gets progressively warmer (e.g. as a MOR approaches the trench, or the rate of subduction slows) or subduction
stops, the forearc mantle will begin to warm and the region of serpentine stability will shrink, releasing water along the wedge boundaries.
Kirby et al. (2013) calculated the release rate of water for the postsubduction case pertinent to the California Coast Ranges. Most of the
~ 1014 kg/km water stored in the wedge mantle is released as free
water by 10–25 M.y. after subduction ceases; its longevity depends on
specific model assumptions on the wedge size and the warming mechanism. In particular, these calculations were conducted for the subduction zone where the warm Farallon slab was subducting. In cold
subduction zones, a volume of serpentinized forearc mantle could be
T. Seno, S.H. Kirby / Earth-Science Reviews 129 (2014) 85–99
much larger, and consequently the water release time could be longer.
During the period of dehydration of such serpentinized forearc mantle,
the released pressurized water weakens its boundary by the mechanisms described below in more detail. If a regional shearing tectonic
stress operates, the dehydrating boundary could become a transform
fault (Fig. 1e and f), as suggested for the Coast Ranges by Kirby et al.
(2013).
Such formation of plate boundaries may not be restricted to active or
relict forearcs. In some cases, serpentinized forearc mantle that overrides active margins may be amalgamated into a collision zone (Burke
et al., 1977; Dewey, 1977). Utilizing the result that water release occurs
over a time scale of tens of M.y., some of this serpentinized mantle may
survive and be stabilized into suture zones, if the period of orogeny is
short enough. Furthermore, since the Proterozoic, rifts disrupting a supercontinent often occurred along suture zones (e.g. Burke et al.,
2003). Disrupted smaller continents amalgamated again to form a
new supercontinent. This is called a Wilson cycle or a supercontinent
cycle, after Wilson (1966) pointed out that a proto-Atlantic (Iapetus
Ocean) closed during the Middle–Upper Paleozoic and the Atlantic reopened in the early Mesozoic. It is now known that such cycles have occurred a few times since the Proterozoic (McKerrow and Ziegler, 1972;
Dalziel, 1995; Rogers, 1996). For example, West Gondwana collided
with Laurasia at ~ 750, 490, and 370 Ma since the dispersal of Rodinia
(Dalziel, 1995). Following breakup of a supercontinent, rifted margins
surrounded the ocean that had opened. Some of the serpentinized
forearc mantle may survive and remain in such rifted margins. Relict
arcs, suture zones, and rifted margins are thus possible loci for storage
of serpentinized mantle. Any agents warming them would release
water and weaken the dehydration boundaries. This strain localization,
with sufficient tectonic stress, would create new plate boundaries at
these localities.
Possible mechanisms for the released water to weaken the lithosphere are (1) reduction of fault strength by increased pore fluid
pressure (Hubbert and Rubey, 1959), (2) hydrofracturing, and
(3) reaction-enhanced ductility. The effects of increased pore fluid
pressure are similar to those for active faults in the crust. Fluid production rate arising from the interplay between dehydration, permeability
of the host rock, pore creation by faulting, pore compaction by ambient
stresses, and shear heating controls the temporal change in the
pore pressure and thus, the strength of the dehydrating boundary
(e.g. Sleep and Blanpied, 1992; Segall and Rice, 1995; Sibson, 1996).
However, application of this weakening mechanism to realistic serpentinized forearc mantle awaits future studies. Hydrofracturing occurs
under the following four conditions for the dehydrating material: positive volume change (ΔV N 0) through the dehydration reaction, low
permeability, low strain rate (Nishiyama, 1989), and small differential
stress with σ Hmin b 0 (Secor, 1965; Etheridge, 1983). For antigorite,
ΔV N 0 is satisfied by dehydration that occurs at depths b 75–125 km
(Ulmer and Trommsdorff, 1995; Wunder and Schreyer, 1997;
Bromiley and Pawley, 2003; see also Fig. 4b of Yamasaki and Seno,
2003), which is the depth range of the serpentinized mantle of our interest. Enhancement of ductility of the lower crust by transformation
to eclogite from granulites through fluid injection (Austrheim et al.,
1997; Bjornerud et al., 2002) may be in the category of reactionenhanced ductility, although there is evidence for the existence of brittle deformation because pseudotachylites are observed simultaneously
for this transformation (Austrheim and Boundy, 1994). In the following
subsections, we depict the processes of formation for each type of plate
boundary in terms of these weakening mechanisms.
3.1. Divergent boundaries
In this subsection, we consider the formation mechanism of a divergent boundary within thick lithosphere. Generally, it is difficult for slab
pull or suction forces to disrupt continental lithosphere. Previous studies suggested that plumes play important roles in the breakup of a
91
continent or a supercontinent. Morgan (1971) pointed out temporal coincidence between flood basalts and continental breakup and suggested
a connection between plume activity and the onset of seafloor spreading. Duncan (1984) also noted that hotspots have been overridden by
spreading ridges in the central Atlantic. White and McKenzie (1989)
showed that huge volcanic provinces are located in rift zones, which
may be associated with uprising plumes (i.e. volcanic rifted margins).
Courtillot et al. (1999) further advanced the Morgan's (1971) idea on
the temporal coincidence between plume activities and breakup.
Bott (1982) further proposed that a rising plume warms and thins
the lithosphere, which may help disrupt it under slab suction forces.
Even if, however, the mechanical thickness of the lithosphere is halved,
its strength is still an order larger than the magnitude of plate tectonic
forces. Therefore, pre-existing heterogeneities in the lithospheric structure are necessary for breakup, in addition to plate tectonic forces and
thinning by plumes (Dunbar and Sawyer, 1989; Courtillot et al., 1999;
Tommasi and Vauchez, 2001). In fact, breakup often occurs along sutures (e.g. Burke et al., 2003), and Tommasi and Vauchez (2001) attributed the weakness of suture zones to lattice preferred orientation of
olivine. However, reduction in strength produced by this mechanism
is less than two thirds of magnitude (see Fig. 5 of Tommasi and
Vauchez, 2001) and is still short of that needed for plate tectonic forces
to cause breakup.
In Fig. 4a, we assume that 200-km-thick cratonic lithosphere has a
temperature of ~ 1200 °C at the base (adapting values of Figs. 7 and 8
of Wyllie, 1988), serpentinized forearc mantle is embedded in the suture zone, and subduction occurs from both sides of the continent
(Bott, 1982). Antigorite is stable at depths b 150 km in such lithosphere
(see Fig. 4b of Yamasaki and Seno, 2003). Even if the temperature at the
base is by ~ 100 °C higher during the Archean (Jarvis and Campbell,
1983), antigorite is still stable at a similar depth range. We now assume
that plumes rise beneath some locations in the suture. Thermal anomalies in the regions adjacent to the plumes would significantly reduce the
strength of the lithosphere by dehydration (Fig. 4b). This mechanism,
together with thinning of the lithosphere (Bott, 1982) and weakening
due to lattice preferred orientation (Tommasi and Vauchez, 2001),
might lead to breakup of a supercontinent, which is then succeeded
by oceanfloor spreading (Fig. 4c). In Fig. 1e, we assume a different
tectonic situation where a supercontinent in a divergent zone, without
subduction from both sides, overrides an upwelling plume. In this
case, basal drag forces, with a dehydration weakening mechanism
similar to that shown in Fig. 4b, may lead to breakup of a continent.
Supporting this inference, serpentinized mantle is often found in
non-volcanic rifted margins in the Atlantic (Fig. 5, Billot et al., 1989;
Whitmarsh et al., 1993; Reid, 1994; Chian et al., 1995; O'Reilly et al.,
1996). This may be comprised of relics of serpentinized forearc mantle
embedded in suture zones, whose dehydration boundaries acted as
weak zones for continental breakup. In volcanic margins, we do not observe such serpentinized mantle, but high velocity layers in the thickened lower crust, which are interpreted as igneous material added at
the time of rifting (White et al., 1987). We speculate that serpentine is
dried out in such volcanic margins beneath which plumes of high
temperature rose up (White and McKenzie, 1989).
This explanation for the serpentinized mantle found in non-volcanic
rifted margins is an alternative to downward infiltration of seawater
through brittle fractures of the entire crust proposed by others
(O'Reilly et al., 1996; Perez-Gussinye and Reston, 2001). Note that
such infiltration is only possible after extension of the rifted crust is
largely accomplished and already thinned (Perez-Gussinye and Reston,
2001). The fact that serpentinized mantle is observed at continent–
ocean boundaries, not beneath centers of abyssal plains (Billot et al.,
1989; White et al., 1990; Pinheiro et al., 1992; Whitmarsh et al., 1993;
Reid, 1994; Chian et al., 1995) seems to rule out such infiltration. Isotope
studies by Skelton and Valley (2000) in the Iberia Abyssal Plain also
show that a massive volume of the mantle was serpentinized before it
was exhumed. The mass of water stored in the serpentine beneath the
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T. Seno, S.H. Kirby / Earth-Science Reviews 129 (2014) 85–99
Suture zone
(a)
Suction force
200 km,
1200 °C
Serpentinized
mantle
Continental litosphere
in tension
(b)
Dehydration
Rifting
Plume
(c)
Initiation of
spreading
Non-volcanic margin
Plume
Fig. 4. Possible mechanism of continental breakup (modified from Fig. 1 of Bott, 1982). Subduction occurs from both sides of a supercontinent. Fsu = trench suction, Frp = ridge push.
These forces are not enough large to cause continental breakup. (a) In this case, we assume that a suture zone was formed in the middle of the supercontinent by the closure of an intervening ocean basin, and that serpentinized forearc mantle, formed during this earlier phase of subduction, is embedded in the suture zone. Note that antigorite is stable in the upper 3/4 of
the lithosphere if the temperature at the base is ~1200 °C (Wyllie, 1988). (b) In the succeeding stage, we assume that a plume begins rising beneath the suture (White and McKenzie,
1989). Dehydration of the serpentinized forearc mantle occurs as it is affected by such a plume, making the dehydration boundary very weak. (c) Along with non-volcanic margins heated
by the plume, this would make it easier for the continent to rift and break up. The above tectonic situation is different from Fig. 1e (left), but the proposed mechanism of formation of a
divergent boundary is the same.
Rockall Trough amounts to 0.6 × 1014 kg/km (O'Reilly et al., 1996),
which is on the same order of the mass of water estimated for the
California Coast Ranges (Kirby et al., 2013), suggesting that our interpretation is plausible.
3.2. Convergent boundaries
Since the initiation of plate tectonics, a number of subduction zones
must have been newly created. Stern (2004) classified these into induced nucleation subduction zones (INSZ) and spontaneous nucleation
subduction zones (SNSZ). In INSZ, transference and polarity reversal
types are distinguished. Transference INSZ move the new subduction
zone outboard of the failed one. The on-going development of a plate
boundary in the Indian Ocean, southeast of India, in response to the
Indo–Asian collision (Wiens et al., 1985; Neprochnov et al., 1988;
Stein et al., 1989) is a typical example of transference INSZ processes.
Polarity reversal INSZ also follow collision, but in these cases, a new subduction zone forms behind the magmatic arc. As discussed in the
Introduction, a force one order larger than plate tectonic forces is required to initiate a new subduction zone. This is the same for INSZ. A
low pore fluid pressure ratio at a collision zone thrust due to lesser
dehydration from underthrusting continental crust might produce
such a large tectonic force (Seno, 2007, 2008). The fact that no feature
like a subduction zone has yet developed in the Indian Ocean (e.g. Bull
and Scrutton, 1990; Van Orman et al., 1995) indicates, however, that
the magnitude of such a collision force would not be overwhelming.
SNSZ results from gravitational instability of oceanic lithosphere,
either at a passive margin or at a transform fault/fracture zone.
Cloetingh et al. (1984), however, showed that spontaneous collapse of
old lithosphere at a passive margin is unlikely due to its strength. In
fact, there is no example of SNSZ at passive margins during the Cenozoic
(Stern, 2004). Therefore, collapse under a tensional or strike–slip tectonic setting has been invoked (Turcotte et al., 1977; Erickson, 1993;
Kemp and Stevenson, 1996). Accordingly, numerical simulations show
that lithosphere in contact with another one along a transform fault/
fracture zone could be unstable if there is enough difference in age
(e.g. Fujimoto and Tomoda, 1985; Hall et al., 2003), resulting in SNSZ.
Not many such examples are, however, known in geologic history. The
Izu–Bonin–Mariana arc is often referred to as a typical example of
SNSZ formed by such collapse along a transform fault (Hilde et al.,
1977; Hawkins et al., 1984; Wells, 1989; Stern and Bloomer, 1992).
Seno and Maruyama (1984), Seno (1988), Hall et al. (1995),
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Deschamps and Lallemand (2002), and Taylor and Goodlife (2004),
however, discounted such an origin, because the Bonin arc has been active at least since 48 Ma, with an arc trend oblique to the hypothesized
transform fault.
Therefore, although it has been generally conceived that a passive
margin or a transform fault/fracture zone turns into SNSZ, the mechanism is still disputed. Similarly, the amount of tectonic force required
for nucleation of INSZ is yet unresolved. Even if thermal cracking
weakens oceanic lithosphere (Korenaga, 2007), the lower half of the
lithosphere might remain strong, as seen by the dominant reverse
fault-type earthquakes beneath mid-oceans, outer-rises, and passive
margins (Stein et al., 1979; Forsyth, 1982; Wiens and Stein, 1985;
Seno and Yamanaka, 1996).
We propose, similar to the formation of divergent boundaries, that
dehydration weakening of serpentinized forearc mantle may have
played an important role in initiating subduction zones in the Earth's
history after plate tectonics had started. As stated before, serpentinized
forearc mantle may be embedded in non-volcanic rifted margins or in
relict arcs where subduction has stopped (Fig. 1e). If any thermal
event occurs in its vicinity under a regional compressional stress, the released pressurized water will weaken the dehydration boundaries of
the mantle, and possibly induce INSZ or SNSZ. Upwelling plumes may
be responsible for such thermal events, and may behave as a control factor of subduction zone nucleation. This would be much more efficient in
weakening lithosphere along a rifted margin or a relict arc than a simple
interaction between lithosphere and plumes (Burov and Cloetingh,
2010) or sediment loading (Cloetingh et al., 1984; Erickson, 1993). An
extraordinary amount of collision forces is not required for INSZ in
this case.
3.3. The San Andreas fault system (SAF)
The SAF was formed by disappearance of the Farallon plate and
succeeding contact of the Pacific plate with the North American plate
(Atwater, 1970). Because the fault system developed a few hundred
km east of the relict trench axis (Fig. 6a, McCulloch, 1989), we cannot
apply, in a straightforward manner, Atwater's (1970) idea (Fig. 6b) for
the initiation of the SAF. As already mentioned, Kirby et al. (2003,
2013) proposed a model suggesting that the mobility of the SAF was enabled by dehydration weakening of the serpentinized forearc mantle
beneath the Coast Ranges. As a margin-parallel shear stress was applied
along the mantle wedge by the Pacific-North American relative motion
(Fig. 6c), the weak zones in the dehydrating mantle would have faulted
aseismically, and the crust located above would have deformed by
strike–slip faulting (See Kirby et al. (2013) for more details). Warming
occurred in this case by a slab gap (Dickinson and Snyder, 1979;
Furlong et al., 1989) or by a stalled slab (Bohannon and Parsons, 1995;
ten Brink et al., 1999) after the cessation of subduction of the Farallon
plate. Fulton and Saffer (2009) supported this idea with a numerical
simulation that showed that such a release of water could maintain a
pore fluid pressure high enough to mobilize the SAF. In other regions,
similar release of pressurized water could occur along a relict arc or a
rifted margin, if they contain serpentinized forearc mantle warmed by
some agents. If a shearing tectonic stress operates as in the Coast
Ranges, it will induce transcurrent motion along the dehydrating
boundary, which might lead to initiation of a transform fault (Fig. 1e
and f).
3.4. Other continental and oceanic transform faults
A spectacular example of formation of major strike–slip faults in collision zones is the North Anatolian fault (NAF) of northern Turkey. The
NAF, whose activity was initiated by the collision of Arabia with Eurasia
(Sengor, 1979), is located along a suture zone formed by the earlier collision between Laurasia and Gondwana (the Pontide–Sakarya collision
during the Late Cretaceous) (e.g. Sengor, 1979; Yilmaz et al., 1995).
93
Three metamorphosed units are observed in the western NAF, between
the two branches (north and south); their protoliths are the Sakarya
continent, ophiolites, and the Pontide fragment from south to north
(Yilmaz et al., 1995). Yilmaz et al. (1995) regarded the metamorphosed
ophiolites as the fragments of the Neotethys. We suggest an alternative
interpretation that these rocks represent the serpentinized forearc
mantle formed by the subduction of the Neotethys beneath the Sakarya
continent, because the serpentinized peridotites occur within the Sakarya continent (Figs. 2, 3, 6, and 7 of Yilmaz et al., 1995). If this is the case,
its dehydration due to warming in recent times after the cessation of
subduction of the Neotethys might have induced the strike–slip motion
of the NAF. This, however, remains speculative at present, and evidence
for this should be sought from geological data in this region.
The above mechanism may also allow strike–slip faulting to occur at
low stresses far into continental interiors. There are many geological examples of multiple suture zones in East Asia and North America, where
fragments of continents and island-arcs were amalgamated during the
Mesozoic–early Tertiary (e.g. Ben-Avraham et al., 1981; Klimetz, 1983;
Silver and Smith, 1983; Davis and Plafker, 1986; Scholl et al., 1986).
These terrane collisions and accretions, in association with the seaward
stepping of trenches, suggest the possible existence of serpentinized
forearc mantle in such suture zones. Some of them may have survived
and mobilized huge continental strike–slip faults by later thermal
events. The strike–slip Denali fault north of the Alaska Trench and the
Altyn Tagh, Kunlun, and Red River faults in interior Asia might be such
examples, although further investigations on their origins are necessary.
Ridge–ridge transform faults developing in mid-oceans following
continental breakup or trench–trench transform faults along fracture
zones may not be relevant to the dehydration weakening proposed in
this study. In these transform faults, infiltration of seawater along the
faults may form serpentine (Calvert and Potts, 1985; Macdonald and
Fyfe, 1985; Potts et al., 1986), whose low coefficient of friction
(Reinen et al., 1991; Moore et al., 1997) may play an important role in
reducing the shear strength of these faults (e.g. Escartin et al., 1997).
Even for ridge–ridge transform faults in their initiation stage, however,
there may be embryos within the serpentinized forearc mantle, because
continental breakup often occurs along suture zones (e.g. Burke et al.,
2003). Trench–trench transform faults may also originate from arcs
that might contain serpentinized forearc mantle. In such cases, dehydration of the warming serpentinized forearc mantle may provide an important clue to elucidate the evolution of this type of transform fault.
4. Discussion
Initiation of plate tectonics by the formation of plate boundaries is a
difficult, but nevertheless important, problem in Earth sciences. The difficulty arises from two contradictory aspects of plate tectonics: a plate
(lithosphere) must be strong and, at the same time, it must be mobile
in order to extract heat efficiently from the Earth's interior. Moving
plates require weak plate boundaries, but the strength of plates makes
the formation of these boundaries difficult. The thick crust formed at
the Earth's surface also prevents subduction. The problem is diverse
and vast, covering chemical alteration, rheology of the surface and interior of the Earth, and its tectonic history. As means to address this diverse problem, we propose that volatilization of an underthrusting
hydrated/carbonated slab and dehydration of serpentinized forearc
mantle are major agents for strain localization and weakening of lithosphere to initiate plate boundary formation.
Although previous studies have discussed various weakening mechanisms of lithosphere, including the effects of volatiles (Bercovici, 1998;
Moresi and Solomatov, 1998; Tackley, 1998; Regenauer-Lieb et al.,
2001; Bercovici, 2003; O'Neill et al., 2007), their treatments are all general. In this paper, we provide a practical hypothesis on the roles of volatiles in the formation processes of new plate boundaries such as
initiation of subduction, rifting/spreading, and transform faults. Because
the mechanisms we propose are put in the context of the Earth's history,
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(a)
Greenland
Labrador
Atlantic
(b)
Iberia
Atlantic
T. Seno, S.H. Kirby / Earth-Science Reviews 129 (2014) 85–99
(a)
W 125°
124°
123°
121°
122°
95
120°
N
40°
40°
NA
39°
39°
38°
38°
Relict trench
SA
37°
37°
F
PA
36°
36°
35°
35°
125°
124°
123°
(b)
121°
122°
120°
(c)
Trench axis
Trench axis
New segment of SAF
NA
NA
NA
FA
FA
PA
NA
PA
New segment of SAF
Dehydrating
boundary
PA
Serpentinized forearc wedge
Fig. 6. (a) Tectonic setting of the California Coast Ranges: the San Andreas fault system (SAF) and various terranes (stippled area) (after McCulloch, 1989). The location of the SAF is
100–200 km landward of the relict trench in which the Farallon plate was subducting during the Tertiary. (b) The Farallon plate (FA) vanished beneath the North American plate
(NA), and the contact of the Pacific plate (PA) to NA became the SAF along the relict trench, according to the theory of Atwater (1970). (c) Dehydration from the warming serpentinized
forearc mantle occurs (Kirby et al., 2003, 2013) at the boundaries of the wedge. The SAF could initiate 100–200 km landward of the relict trench along such a boundary. The forearc wedge
behaves as part of PA after the formation of the SAF.
they are more easily tested than the mechanisms proposed in previous
studies. Obviously, however, we present only a rough sketch of the
mechanisms in this paper, and do not claim to have proven them to
be correct. A lot of work remains to be done to examine these mechanisms quantitatively based on geological or geophysical data in each
region.
Sleep (2000) presented a scheme of transition between a magma
ocean, plate tectonics, and stagnant-lid convection for terrestrial
planets. In some planets, transitions between these modes may have occurred several times. In this scheme, however, mechanisms of the transition to plate tectonics, except for the role of partial melting in the
divergent zone, were not presented. In the future, the Earth will
transition to stagnant-lid convection as it is cooled efficiently by plate
tectonics. After this, it is expected that radioactive heat production
will heat up the interior, and plate tectonics will resume (Sleep, 2000).
For this to occur, however, mechanisms that enable the initiation of subduction (similar to those presented in this study) are necessary; otherwise the Earth will return to a magma ocean. It is noted that both the
convective vigor and volatile extent of the Earth's interior change temporally, interacting with each other in a complicated manner. There is
no guarantee that the condition for resumption of plate tectonics will
be satisfied in the future. The weakening mechanisms we propose
will, nevertheless, provide a useful key to conjecture about shifts between the modes of convection for terrestrial planets.
Fig. 5. Examples of serpentinite bodies in non-volcanic rifted margins. (a) Map of the Labrador Sea, showing the zones of high-velocity (7.2–7.7 km/s) material on each margin (shaded),
which is inferred to be serpentinized peridotite (Chian et al., 1995). The grey lines are fracture zones and the solid lines are the lines of seismic experiments. (b) Map of the region offshore
of the western continental margin of Iberia (contours of ocean depth in meters), showing zones of thin crust underlain by serpentinized peridotite (shaded area, Whitmarsh et al., 1993).
A.P. denotes abyssal plain. The solid lines are the lines of seismic experiments. These maps show that the serpentinized peridotites are located at the ocean–continent boundaries of the
non-volcanic rifted margins, not at the center of the abyssal plains.
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Our hypothesis also has implications for the formation and exhumation of high pressure (HP)/ultrahigh pressure (UHP) metamorphic
rocks and emplacement of some ophiolites and ultramafic rocks in subduction zones. Belts of HP or UHP metamorphic rocks are found in
forearcs of active subduction zones and collision zones. Exposure on
the Earth's surface of such coesite- and diamond-bearing HP or UHP
rocks implies exhumation from a source in the mantle deeper than
100–150 km (Coleman and Wang, 1995; Ernst and Liou, 1999;
Chopin, 2003). Although the origin of these rocks is still disputed
(Brueckner, 1998; Terry and Robinson, 1999; Guillot et al., 2001), it is
very likely that their exhumation involves tectonic movements not
only at the mid-lower crustal level but also deep in the mantle. It is
also episodic, i.e., not a steady-state process. In many cases, HP/UHP
belts are sandwiched between lower metamorphic-grade units in fault
contact (Worrall, 1981; Burg et al., 1984; Wheeler, 1991; Kaneko
et al., 2003), which implies that they were brought to the surface largely
by faulting or strain localization (Burchfiel and Royden, 1985;
Maruyama et al., 1996). These features, i.e., episodicity, strain localization, and mantle-scale processes, are common to those of plate boundary formation discussed in this study, suggesting that a mechanism
similar to dehydration weakening and mobilization of huge thrusts in
relict subduction or collision zones might be operating in exhumation
of UHP/HP rocks. Subduction of continental materials into the deep
mantle may provide sources for UHP rocks via the process we propose
for subduction of lithosphere with thick crust. This subject, however,
will be treated in a separate paper.
5. Conclusions
In the early Earth, convection would have occurred with thick basaltic crust accumulating over a soft boundary layer downwelling into the
mantle (Davies, 1992). This might have transitioned to stagnant-lid
convection as the mantle cooled (Solomatov and Moresi, 1997) or
back to a magma ocean as the mantle heated (Sleep, 2000). For plate
tectonics to operate on the Earth, subduction of lithosphere with thick
crust must have been initiated, thus avoiding these shifts. Based on an
analogy with the continental plate subduction that is currently occurring in Hindu Kush and Burma, we propose that volatilization from underthrusting lithosphere, once hydrated/carbonated by H–O–C rich
vapors released from magmas in upwelling plumes, would have made
it possible. Once subduction has been initiated, serpentinized forearc
mantle may form in a wedge-shaped body above a dehydrating slab.
Therefore, in relict arcs, suture zones, or rifted margins, any agent that
warms and dehydrates the serpentinized mantle would weaken the
region surrounding it, and form various types of plate boundaries
depending on the operating tectonic stress.
Thus, once subduction initiates, formation of plate boundaries might
be facilitated by a single fundamental process (except for the MOR
formed by partial melting in the divergent zone): weakening due to
the release of pressurized water from warming serpentinized forearc
mantle. The fundamental ideas presented in this paper are simple, and
details of geological and geophysical facts in each appropriate region
should be examined more quantitatively. However, we feel that the insights that we have obtained from a preliminary application of these
concepts will have important implications for interpreting geological
histories in mobile belts, and are sufficiently encouraging to report
and put them into other people's discussion.
Acknowledgments
We thank H. Austrheim, an anonymous reviewer, and Manfred
Strecker, whose critical comments are very useful to improve the manuscript, and Satoru Honda and Tadashi Yamasaki for their comments
and discussion on the early version of the manuscript.
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