Tectonophysics 468 (2009) 6–27
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Tectonophysics
j o u r n a l h o m e p a g e : w w w. e l s e v i e r. c o m / l o c a t e / t e c t o
The structure, evolution and symmetry of the magma-poor rifted margins of
the North and Central Atlantic: A synthesis
T.J. Reston
University of Birmingham, School of Geography, Earth and Environmental Sciences, Birmingham, United Kingdom
a r t i c l e
i n f o
Article history:
Received 31 August 2007
Accepted 3 September 2008
Available online 16 September 2008
Keywords:
Rifted margins
Rheological evolution
Symmetry of rifting
Serpentinization
Extension
a b s t r a c t
Magma-poor rifted margins consistently show extreme crustal thinning accompanied by normal faulting, the
serpentinization of the mantle beneath crust thinned to less than 8 ± 2 km, and the unroofing of a broad zone
of mantle within the continent–ocean transition, accompanied by the development of detachment and other
large-offset faults. These observations are the logical result of the progressive extension of cool lithosphere
away from thermal anomalies such as plumes. Although the paucity of magmatism may be explained by
depth-dependent extension of the lithosphere, and by pre-depleted sub-crustal lithosphere, rifting above
initially cool sub-lithospheric mantle also may explain the subsidence deficit observed at some margins:
synrift subsidence is buffered by the simultaneous influx of warmer oceanic asthenosphere. As it extends,
thins and cools, ductile creeping layers in the mid- and deep crust become progressively more brittle,
resulting in increased coupling between the upper and lower crust, and eventually the embrittlement of the
entire crust, faults cutting from the surface across the Moho, bringing water into the mantle and causing its
serpentinization. Increased coupling and the development of serpentine detachments predict the
development of late-stage asymmetry once the entire crust is brittle. Such detachments are imaged on
some margins and inferred on others; analysis of the crustal structure across conjugate margins shows that
these are approximately symmetric until this late stage when they become markedly asymmetric. Similar
analyses show that depth-dependent stretching of the crust is insufficient to explain the discrepancy
between the amount of the visible extension along faults and the amount of crustal thinning. Instead this
“extension discrepancy” may be related to the complex evolution of brittle deformation through multiple
phases and styles of faulting, related to the changes in the rheological character and strength of the
lithosphere as it is thinned. Complex polyphase faulting continues after complete crustal separation,
resulting in the exhumation of broad expanses of peridotitic basement, the top of which is everywhere
marked by an exhumed slip surface, similar to the corrugated surface observed at mid-ocean ridges. The
similarity in processes between mantle unroofing and seafloor spreading makes the distinction between the
COT and true oceanic crust difficult and possibly moot.
© 2008 Elsevier B.V. All rights reserved.
1. Introduction
The first order tectonic process of continental breakup initiates the
plate tectonic cycle of plate creation and destruction. Rifted margins are
the trailing edges of the continents that develop as the continents are
rifted apart (Fig. 1). In addition to their economic importance (hydrocarbons) and hazard potential (especially slope failure), they provide a
record of the processes that accompanied continental breakup.
In the past, rifted margins have been classified as volcanic or nonvolcanic margins. These terms are a bit misleading as even nonvolcanic margins exhibit magmatism, so in this paper I will prefer the
term “magma-poor” and “magma-dominated margin” (Sawyer et al.,
2007), depending whether less or more magmatism is observed than
might be expected during the rifting above normal asthenosphere.
E-mail address: [email protected].
0040-1951/$ – see front matter © 2008 Elsevier B.V. All rights reserved.
doi:10.1016/j.tecto.2008.09.002
This is undepleted mantle with a potential temperature of 1300 °C
(potential temperature Tp is the temperature it would have if brought
rapidly to the surface). During rifting, the upwelling of such mantle
leads to pressure release melting, producing 6–7 km for very rapid
rifting and hence upwelling, but perhaps 3–4 km for typical rift
durations (Minshull et al., 2001). Although some margins (e.g. West
Iberia) are so devoid of melt products that it is very likely that little
melting took place during rifting, at other margins the difficulties in
determining both the thickness of igneous addition to the lithosphere
(some may be trapped below the Moho) and precise rift durations
means that a practical definition of magma-poor margins may be
those where tectonic rather than magmatic processes dominate
during rifting. Judging by the magma-poor west Iberia margin,
tectonic rather than magmatic processes dominate when there has
been less than about 2–3 km of igneous addition prior to crustal
breakup. Note that by restricting the definition to what happens
during rifting, several margins and basins that had little magmatism
T.J. Reston / Tectonophysics 468 (2009) 6–27
7
Fig. 1. Top: Worldwide distribution of different margin types (modified after Boillot and Coulon, 1998) and the locations of the margins discussed in this paper (bold). Magma-poor, magma-dominated and transform are all approximately
equally numerous. Below: Cartoon sections summarising the architecture of magma-poor (based on observations made in this paper), magma-dominated (modified after Gernigon et al., 2004) and transform (based on Edwards et al., 1997)
margins. This paper investigates the first category.
8
T.J. Reston / Tectonophysics 468 (2009) 6–27
until breakup, although subsequently subjected to voluminous postrift
magmatism (e.g. Rockall Trough and the Norwegian Vøring (Gernigon et
al., 2004) and Møre Basins), have much in common with magma-poor
margins.
Other rifted margins (e.g. Woodlark Basin) appear intermediate
between the end-member “magma-poor” and “magma-dominated”
margins as they may exhibit approximately the amount of magmatism
expected for rapid mantle upwelling and do not show the extreme
tectonism of magma-poor margins described below. However in several
cases, the margin is simply not well enough characterised at synrift
levels to determine the mechanics of breakup. Finally, transform
margins (e.g., Edwards et al., 1997) form a completely separate class of
margin formed during continental breakup, generally exhibiting little
magmatism, but marked by abrupt change in crustal thickness.
This paper is concerned with the structure and tectonics of
magma-poor margins. By concentrating on magma-poor margins, it is
possible to avoid some of the geological and geophysical complexities
caused by magmatic addition both within the crust (affecting crustal
thickness, crustal rheology, and seismic velocity) and on top of the
crust as lava flows (causing imaging problems and thus obscuring the
tectonic structures beneath). As such it is intended that the paper
provide a counterpoint to the more magmatic theme of many papers
in this volume. In this paper, I combine observations of the structure of
the 22 best studied magma-poor margins of the North and Central
Atlantic with the predictions of the fundamental processes accompanying lithospheric extension and breakup.
2.
2. Magma-poor margins of the North and Central Atlantic and
neighbouring areas
Magma-poor margins are found in every ocean (Atlantic, Indian,
Southern, Pacific, Arctic), but those in the North and Central Atlantic
have been most intensively studied, are well constrained by
geophysical (both reflection and refraction) and ODP data, and for
the most part are unambiguously magma-poor (Figs. 2 and 3). Moving
from north to south, examples include parts of Labrador–West
Greenland (Chian et al., 1995; Chalmers and Pulvertaft, 2001), Flemish
Cap–Goban Spur (Keen et al., 1989; Bullock and Minshull, 2005), North
Biscay (Thinon et al., 2003), Flemish Cap–Galicia Bank (Funck et al.,
2003; Zelt et al., 2003; Hopper et al., 2004; Reston et al., 2007),
Newfoundland Basin–South Iberia Abyssal Plain (Krawczyk et al.,
1996; Pickup et al., 1996; Chian et al., 1999; Dean et al., 2001;
Whitmarsh et al., 2001; Van Avendonk et al., 2006), S Newfoundland
Basin–Tagus (Lau et al., 2006) and Nova Scotia–Morocco (Funck et al.,
2004; Contrucci et al., 2004; Wu et al., 2006; Maillard et al., 2006). In
addition are a number of failed rifts bordering the Atlantic where
rifting progressed so far that crustal separation may have taken place
or was at least imminent. These include Rockall Trough (O'Reilly et al.,
1996; Pérez-Gussinyé et al., 2001), Porcupine Basin (Reston et al., 2001,
2004; O'Reilly et al., 2006) and similar basins in the Mediterranean Sea:
the Ligurian Sea (Contrucci et al., 2001), the Tyrrhenian Sea, the
Alboran Sea, the South Balearic Basin and the Valencia Trough. Finally,
one example from further south is included: the Rio Muni (Equatorial
Guinea) margin (Wilson et al., 2003; Turner et al., 2003). Other margins
which appear very similar to the above either are not well enough
characterized geophysically (South Australia) or have thick sedimentary
sequences that obscure the deeper structure so that it is not always
totally clear whether they are magma-poor or not (e.g., Angola margin
in the South Atlantic).
The majority of magma-poor margins from the North and Central
Atlantic display a number of common features (Figs. 2 and 3; Table 1):
1. Extreme crustal thinning from ∼ 30 km to a few km in most cases
over a distance of 100–200 km. On most margins well-defined fault
blocks are imaged, but the amount of extension associated with
these is generally less than required to explain the crustal thinning.
3.
4.
5.
This is the so-called extension discrepancy, discussed below. The
thinning occurs most rapidly in a necking region between c 20 and
10 km, corresponding to β factors (McKenzie, 1978) between 1.5
and 3 and (1 − 1 / β) values of 0.33 to 0.67 (Fig. 4): most margins are
characterised by a broad shelf, a steep continental slope and an
extensive continental rise of very thin crust. Extension of the brittle
upper crust is accommodated by faulting, and summing the fault
heaves provides an estimate of the amount of extension. Other
parts of the lithosphere may initially deform by other means
(ductile flow or creep), although there is evidence that deformation
in parts of the lower crust and uppermost mantle is localised into
shear zones (Reston, 1988, 1993). However, as discussed below, the
deformation mechanics of the lithosphere evolve during extension
and thinning, as the initial rheological properties are modified by
strain, and by changes in temperature and pressure (PérezGussinyé and Reston, 2001).
A zone of anomalous basement between the last identifiable
continental crust and the first true oceanic crust. This unit typically
has a moderate velocity gradient (∼0.2 s− 1), passing down from
5 km/s to close to 8 km/s and weak magnetic anomalies. The unit
has been sampled by dredging, by submersible (Boillot et al., 1988)
and by drilling on the West Iberian margin (sites 637, 897, 899, 1068
and 1070 — Boillot and Winterer, 1988; Whitmarsh et al., 1996,
2001), by drilling off the Newfoundland margin (Tucholke et al.,
2007) and by dredging to the south of Australia (Nicholls et al.,
1981), and has been found to be partially serpentinized peridotites.
The anomalous crust is thus thought to be exhumed mantle, with
downward decreasing degrees of serpentinization explaining
the velocity gradient. Elsewhere, it has been identified by similar
geophysical characteristics (velocity), a lack of oceanic spreading
anomalies, and simply a non-continental appearance on the seismic reflection images.
Beneath the crust where it had thinned during rifting to less than
about 8 km, zones with velocity intermediate between those of the
crust and mantle (i.e. between 7 and 7.8 km/s) and a velocity
gradient of close to 0.1 s− 1. In most cases, the zones can be traced
laterally into regions of unroofed mantle (see above) and so are
interpreted as partially serpentinized peridotites, “undercrusting”
(Boillot et al., 1989) the crust. Where they cannot be traced laterally,
they might alternatively have been interpreted as mafic underplate
or a zone of mantle intruded with mafic intrusions, if it were not for
the paucity of magmatism at most of these margins. Furthermore,
the velocity gradient noted above is consistent with downward
decreasing degrees of serpentinization but less easily explained by
mafic underplate. It is however important to realise that nowhere
has serpentinized mantle actually been sampled beneath the
thinned continental crust.
At some margins, the unroofed mantle is covered with low velocity
and locally mounded units up to 3 km thick. As they are generally
observed atop the exhumed mantle and have seismic velocity
consistent with highly serpentinized peridotites with varying degrees of porosity, they are generally interpreted as such (Tucholke
et al., 2007): serpentine breccias and sedimentary serpentine have
been sampled at ODP sites 897 and 899 in the southern Iberia
Abyssal Plain (Whitmarsh et al., 1996). The mounded pattern may
reflect either lateral transport, such as the collapse of serpentinite
highs, or emplacement from below (Reston et al., 2001, 2004).
At some margins, the boundary between highly thinned continental crust and the underlying zone interpreted as serpentinized
mantle is marked by a sharp reflector, most clearly imaged on
depth images produced by prestack depth migration. Examples
include the S reflector west of Galicia (de Charpal et al., 1978;
Hoffmann and Reston, 1992), the H reflector in the Iberia Abyssal
Plain (Krawczyk et al., 1996), and the P reflection beneath the
Porcupine Basin (Reston et al., 2001, 2004). In each case, the
reflector does not only appear to separate crustal rocks from the
T.J. Reston / Tectonophysics 468 (2009) 6–27
Fig. 2. Sections through 14 conjugate or near-conjugate margins (including both sides of two basins) best constrained by combined seismic reflection and wide-angle data, all shown at the same scale, with no vertical exaggeration. Key velocity
boundaries and contours are marked; those in bold are used for the analyses in Figs. 8, 9 and 13. Also marked is the zone where the crust thins to the value at which it should have become brittle during rifting. This always occurs landward of the
first serpentinized peridotites that underlie the thinned crust of the deep margin.
9
10
T.J. Reston / Tectonophysics 468 (2009) 6–27
Fig. 3. Sections (no vertical exaggeration) through a further 8 margins, either non-conjugate or where full velocity control to the shelf is lacking. In all cases, the landward limit of the
serpentinites beneath the lies slightly oceanward of point at which the crust is thin enough to have become completely brittle during rifting.
underlying serpentinized mantle, but also to act as a detachment
on which the overlying fault blocks ride. Analysis of the physical
properties of the H and S reflection (Reston, 1996) have shown that
they are compatible with a relatively sharp boundary between
fractured crustal rocks and partially serpentinized peridotites,
perhaps supplemented by a thin zone of reduced velocity at the
detachment itself (Leythaeuser et al., 2005).
As many magma-poor margins display the same general characteristics, it seems likely that they also share a common mode of
formation. In the following sections I examine the processes that may
have led to the development of these characteristics at magma-poor
rifted margins and derive a generic model for the formation of
magma-poor margins, focussing on the large-scale structure of the
margins.
3. Large-scale lithospheric deformation — subsidence and melt
generation
Continental breakup occurs through lithospheric extension, thinning
and eventual rupture. Lithospheric stretching occurs at all levels, so that
the crust is highly extended, and separates during continental breakup.
Lithospheric thinning allows the underlying asthenospheric mantle to
well up towards the surface, undergoing partial pressure release melting
according to the amount and rate of upwelling, coupled with the
temperature and fertility of the asthenosphere (Bown and White, 1995).
The thinning of the crust causes subsidence, which is partly offset during
and shortly after rifting by the thermal uplift provided by the thinning of
the entire lithosphere (McKenzie, 1978) and the compression of the
lithospheric isotherms, and by the upwelling of the asthenosphere. As a
result, the distribution of both subsidence and magmatism in time and
space provides constraints on the amount of extension and thinning of
the different lithospheric levels.
A magma-poor margin is defined by a lack of significant magmatism. Excess melting at magma-dominated margins (Nielsen and
Hopper, 2002) is generally explained by a combination of high temperature asthenosphere (e.g. “hotspots”), secondary convection and
chemically enriched mantle (including the presence of water). Conversely, the melt deficiency of magma-poor margins may be explained
(Fig. 5) by a cool (Tp below 1300 °C — Reston and Phipps Morgan,
2004), or partially depleted sub-lithospheric mantle (Pérez-Gussinyé
et al., 2006), or one that does not rise completely due to incomplete
lithospheric separation, requiring a component of lithospheric-scale
depth-dependent stretching (DDS) if crustal separation has occurred
(Minshull et al., 2001). Higher viscosities resulting from cooler temperatures and lack of water are also likely to hinder secondary convection. These different explanations have different implications for
the subsidence history of the rifted margin.
Traditionally, the large-scale deformation of the lithosphere
beneath rifts has been determined from subsidence patterns (e.g.
White and McKenzie, 1988). In considering subsidence, lithospheric
thinning is usually considered in two parts: the thinning of the crust
and the thinning of the lithosphere. These generally occur at the same
time but may have a different spatial distribution. The thinning of the
whole lithosphere causes both a compression of the isotherms within
the lithosphere (note in general no part of the lithosphere actually
gets hotter, but hot rocks are brought towards the surface) and the
replacement of higher density mantle lithosphere by lower density
asthenosphere. These combine to cause a thermal uplift which decays with time after rifting as the isotherms revert to horizontal
(meaning that the uplifted and upwelled lithosphere and asthenosphere cool) giving a subsequent gradual thermal subsidence. The
isostatic response to crustal thinning is subsidence: as crustal thinning
is permanent it controls the eventual total amount of subsidence.
In principle, the subsidence at any time is that caused by crustal
thinning minus the effect of any residual thermal uplift. Thus the
Table 1
Geophysical characteristics of rifted margins characterized by mantle serpentinites and their conjugates
MCS Refraction DSDP, Final rift
ODP
duration
[m.y.]
References Underrusting — Thickness Gradient
vels [km/s]
[km]
[s− 1]
approximate
Extent Max
[km]
thickness
overlying
crust [km]
Beta factor
at landward
limit of
serpentinites
Unroofed
mantle max
thickness [km]
Velocity
[km/s]
Gradient
[s− 1]
Extent Highly
Max
[km]
serpentd
thickness
peridotites — [km]
vels km/s
Extent References
[km]
Y
OBS
30–50
6.4–7.6
b 4.5
0.27
50
4–6.5
5–7.5
7
6–7.6
0.23
30–66
∼ 4.8
2
35
W Greenland
Y
OBS
30–50
7.4–8.0
b4
0.15
54
9
3.5
7
6.2–7.4
0.17
54
∼ 4.5
3
35
Rockall
Y
OBS
40–70
7.5–7.8
3
0.1
30?
10–11.5
∼3
15
6–6.9
7.5–7.8
0.15.03
N150
∼ 4.5
2
N150?
Rockall II
(cont crust)
Porcupine
Y
OBS
40–70
7.5–7.8
3–10
0.1–0.03
200
10–11.5
∼3
N
–
–
–
N
–
–
Y
OBS
25–45
7.2–7.8
∼ 10
0.06
75
8–10
3.75–3
?
?
?
30
4.5
b3
30
Goban Spur
Y
OBS
30
7.0–7.5
b2
0.25
10
7
4.3
4.5
6–7
0.22
b39
4.5
2
b30
Armorican
Y
OBS
30
7.4–7.5
2
0.025
10
4
8
4
7.4–7.5
0.025
20
2
20
Ligurian Sea
Y
ESP
14
7.2–7.3
3
–
b50
6–9
5.3–3.6
N
–
–
–
4.6–5.0
(3C, 3D)
N
–
–
SE Flemish Cap
Y
OBS
15–25
7–6–8.0
4
0.1
24
4.5.6.5
6.5–4.5
?
?
?
–
Galicia Bank
IAP–ODP
149–173. LG12,
CAM144
Y
Y
OBS
OBS
103
149–
173
15–25
5–20
7.0–7.6
?
4
4
0.15
50
?
5–6
6.5
5–6
5–4.3
6
10
5.6–7.6
?
0.33
N10
∼100
–
∼4
–
b 3gradient
–
∼100
N Newf Basin
Y
OBS
210
5–20
–
–
–
.6
3
10
4
5–7.5
0.62
N20
N
–
–
S IAP–IAM9
S Newf Basin
Y
Y
OBS
OBS
(173)
5–20
10–25
N
7.6–8
–
6
–
0.07
–
95
–
8
–
3.4–4.3
b6.5
2.8
2.5–5
4–5
N 4.4
2
0.5gradient
100
30?
N
?
Y
OBS
OBS
OBS
10–25
b 55
7.2–7.6
7.2–7.6
N
5
4.5
0.08
0.09
25
70?
7
6.5
4–5
4–5.5
6.5
b6
N
0.3
1.33
0.5
.16–.08
0.08
N 0.07
–
N170
50
S Newf Basin
NovaScotia 1
Morocco
6–7.9
4.4–6.4
6.4–7.8
7.6–8.0
7.2–7.7
7.2–7.6
–
N20
60?
–
4.5–5
5.1
–
2–3
b2
100
60?
Nova Scotia 2
Rio Muni
?
Y
OBS
Grav
b 55
10–15
7.6–7.8
n/a
b 3.5
0.06
n/a
50
6.5
8
4–5.5
4
N
Y?
–
n/a
–
n/a
–
n/a
N? 5?
N
3?
–
35?
–
Chalmers +
Pulvertaft, 2001;
Chian et al., 1995,
Chian et al., 1995;
Reston + PérezGussinyé, 2007
Perez-Gussinye
et al., 2001;
O'Reilly et al.,
1996
O'Reilly et al.,
1996
O'Reilly et al.,
2006; Reston
et al., 2004;
Bullock and
Minshull (2005)
Thinon et al.
(2003)
Contrucci et al.
(2001)
Hopper et al.
(2004)
Zelt et al. (2003)
Krawczyk et al.,
1996; Wilson
et al., 2001;
Chian et al., 1999
Van Avendonk
et al. (2006)
Dean et al. (2001)
Lau et al.,
2006a; 2006b
T.J. Reston / Tectonophysics 468 (2009) 6–27
Labrador
Reid et al. (1994)
Funck et al.,
Maillard et al.,
2006;
Wu et al. (2006)
Wilson et al.
(2003)
Most margins are characterised by, beneath the feather edge of the continental crust, a layer of reduced mantle velocity, which passes laterally into a region of unroofed partially serpentinized mantle, in places overlain by a layer of highly
serpentinized peridotites.
11
12
T.J. Reston / Tectonophysics 468 (2009) 6–27
Fig. 4. Whole crustal thinning factors vs distance from the edge of the continental crust for the conjugate margins in Fig. 2. Some margins appear to thin rapidly and might be termed
“hard” margins (Davison, 1997), others thin more gradually and might be termed “soft”. However, there is no clear pattern of hard–soft asymmetry: some conjugate pairs are hard–
hard, some soft–soft and some hard–soft.
spatial distribution of crustal thinning controls the total tectonic
subsidence, whereas the spatial distribution of whole lithospheric
strain controls the distribution of that subsidence in time. One classic
example is the interpretation of postrift onlap onto prerift sediments
and basement as evidence that mantle extension and thinning is
distributed over a wider area than that of the crust (White and
McKenzie, 1988), consistent with the presence of outward dipping
mantle shear zones, both observed on deep seismic profiles (Reston,
1993) and predicted by numerical modelling (Harry and Sawyer, 1992;
Huismans and Beaumont, 2002). Such models of lithospheric depthdependent stretching are an effective way of suppressing melt
generation: where the crust thins more than does the lithospheric
mantle, the amount of melt generated will be less than expected for
the observed crustal thinning (Fig. 5).
Increasingly however it is being reported that rifted margins
exhibit a subsidence deficit: the deep margin subsided less during
rifting than might be expected from the observed crustal thinning (e.g.
Kusznir and Karner, 2007). This contradicts the DDS explanation for
the steer's head geometry and for the suppression of melting as a
synrift subsidence deficit would imply excessive thermal uplift and
hence more lithospheric than crustal thinning. Although, the
subsidence deficit can potentially be explained by major crustal
DDS, problems with such models are discussed below.
It may be possible to explain both the subsidence deficit and the
lack of voluminous magmatism at rifted margins if sub-lithospheric
mantle is cooler beneath the continents than the oceans, as implied by
several studies of mantle convection (e.g. Lenardic and Moresi, 2000;
Phipps Morgan et al., 1995). In support of these modelling studies,
independent evidence for cool sub-lithospheric mantle beneath
continents comes from a variety of sources, including cratonic
geotherms constructed from kimberlites and more general geotherms
deduced from seismic velocities (Reston and Phipps Morgan, 2004).
Even a slightly cool sub-lithospheric mantle (e.g. potential temperature Tp of 1200–1250 °C) will first of all lead to major melt suppression
in the same way that a “hotspot” might lead to an increase in melt
production (Fig. 5). However a ubiquitous cool sub-lithospheric mantle
cannot explain the observed 6–7 km thickness of the magmatic crust in
the oceans, which requires “normal” asthenospheric Tp of ∼ 1300 °C.
Reston and Phipps Morgan (2004) explained the switch from
suppressed to normal (oceanic) melt production through the influx
of oceanic asthenosphere (1300 °C Tp) beneath a cool continental rift
and pointed out that such an influx would also lead to thermal uplift
during the late stages of rifting leading to breakup. If the influx
Fig. 5. Melt thickness vs rift duration for mantle potential temperatures of 1300, 1250
and 1200 °C (after Minshull et al., 2001; Reston, 2007b). Magma-poor margins may be
defined as those that plot below the 1300 °C contour, due to the presence of cool or
depleted sub-lithospheric mantle, or by lithospheric depth-dependent stretching
(DDS), or as those where there is less than 2–3 km of igneous addition (shaded).
T.J. Reston / Tectonophysics 468 (2009) 6–27
occurred gradually as rifting progressed, the associated thermal uplift
would maintain the thinning crust at a higher level than might
otherwise be expected, appearing as a subsidence deficit. As the
McKenzie (1978) model assumes a constant Tp asthenosphere, this
explanation for the subsidence deficit may have been overlooked.
In general, detailed analysis of the subsidence patterns to determine
how the lithosphere was thinned is problematic. Where the margin is
well-supplied with sediment throughout its evolution, the deepest
synrift units are so deeply buried that they cannot be sampled; where
such margins are sediment-starved, most of the sediments are deposited
in deep water where the paleobathymetric controls are so poor that
meaningful estimates of tectonic subsidence after back-stripping are
impossible. Furthermore, the most reliable indicators of paleo-water
depth (e.g. erosional unconformities) may record not so much the
pattern of lithospheric deformation as the thermal structure of the
underlying mantle and specifically the passage of thermal anomalies
such as mantle plumes. Finally, it may in some circumstances be
misleading to relate paleobathymetry to global sea-level: during the
opening of the ocean basins (and the development of rifted margins)
isolated deep basins may develop in which the local sea-level may be far
below global. Such restricted basins may be characterized by lacustrine
to hypersaline conditions, leading to potentially important lacustrine
deposits (including potential source rocks) and to salt, depending on
climate and on the relative importance of influx from rivers and from the
global ocean. Thus the reported subsidence deficit at rifted margins may
also have a paleo-oceanographic explanation.
Although studies of active rifts could in principle help address these
problems (substituting heat flow measurement for postrift subsidence
as a proxy for the thinning of the lithosphere), there are few such
margins, and even fewer appropriate studies. Buck et al. (1988) did point
out that the structure and heat distribution of the northern Red Sea, once
invoked as an example of simple shear lithospheric extension
(Wernicke, 1985), could be best explained by a symmetric model in
which extension and thinning focused beneath the centre of the rift.
Thus for many ancient rifted margins, the measurable subsidence provides constraints on the pattern of whole crustal but not
lithospheric thinning. However, as discussed below, identical patterns
of whole crustal thinning can be produced by very different patterns
of lithospheric extension. The implication is that total subsidence (and
thus bathymetry), although invoked by some authors (e.g., Le Pichon
and Sibuet, 1981; Kusznir and Karner, 2007) as support for specific
models of lithospheric thinning, really only confirm the pattern of
whole crustal thinning (which also controls gravity) and not the
extension mechanism.
13
(Nova Scotia to 0.75 km/km (S Newfoundland Basin). The gradual
increase in the thinning gradient as the crust thins from ∼30 km
towards ∼ 25 km is reminiscent of a necking profile, except that the
thinning appears asymmetric relative to the edge of the continental
crust (Fig. 4). Using the terminology of Davison (1997), some margins
appear “hard” (abrupt transition from relatively normal thickness
crust) whereas at other (“soft”) margins, the change in crustal
thickness takes place far more gradually. Some conjugate pairs show
similar thinning profiles for both margins, for instance Morocco–Nova
Scotia (soft–soft), Porcupine (hard–hard), Rockall (hard–hard, with
the constraint that it is not known exactly where the COT occurs if at
all), whereas others appear markedly asymmetric (Labrador–W
Greenland; Flemish Cap–Galicia; south IAP–Newfoundland Basin).
The sort of asymmetry the last group displays might be interpreted as
implying extension by an asymmetric crustal-scale process, such as a
crustal-scale detachment, but asymmetry of the final margins does
not necessarily imply an asymmetric rifting process (Reston and
Pérez-Gussinyé, 2007). Rifted margins form through a long period of
extension and even if each rift phase is fundamentally symmetric, the
locus of rifting may change with time, resulting in asymmetric basins
and margins. The Galicia–Flemish Cap margin is a case in point: the
asymmetry here need not imply the presence of a lithospheric-scale
shear zone as has been suggested (Lau and Louden, 2007) but rather
may reflect the changing focus of rifting during prolonged extension
(Tucholke et al., 2007), resulting in an older abandoned rift basin
on one side (the Galicia Interior Basin –rifting) and a more abrupt
continental shelf on the other. This can be illustrated by modelling the
stretching across the Galicia margin as the superposition of two
symmetric rift phases (Fig. 6), each approximated by Gaussian necking
profiles: an older (Tithonian–Hauterivian) one centred on the Galicia
Interior Basin (Pérez-Gussinyé et al., 2003; Reston, 2005) but extending as far as the deep Galicia margin, and a later one (Hauterivian
or Barremian to Aptian) modelled on the shape of thinning across the
SE Flemish Cap margin, but applied to the Galicia margin to make this
second rift phase symmetric between the two margins. In applying
the curve from the Flemish Cap margin to the Galicia margin, the two
profiles are matched at the point of complete crustal embrittlement,
as discussed below. Summing the log of the stretching factors of these
two rift phases reproduces the log of the observed stretching factor
across the entire Galicia margin remarkably well, suggesting that the
main asymmetry between the Flemish Cap and Galicia margins may
4. Crustal thinning: distribution of strain
Subsidence and melt production are both related to the way the
lithosphere thins and the thermal structure of the lithosphere and the
underlying asthenosphere, with total subsidence being controlled in
particular by the way the crust thins. I now turn my attention to this
topic, and consider the distribution of crustal thinning across margins
and their conjugates, focussing in particular on symmetry.
Symmetry comes in all shapes and sizes. For instance, rifted margins
and rifts commonly are divided into domains or segments dominated by
faults dipping in one direction, but this is completely compatible with an
overall symmetric pattern of extension and thinning (e.g. Reston and
Pérez-Gussinyé, 2007). Here we are concerned instead with large-scale
asymmetry in margin structure and in the processes of lithospheric
extension. Most previous discussions (e.g. Davison, 1997) have been
rather qualitative, but the thinning profiles presented here provide a
more quantitative assessment, bearing in mind the limitations of the
seismic velocities for mapping out different crustal levels.
It is clear from both the crustal sections (Figs. 2 and 3) and thinning
factor profiles (Fig. 4) that the crust undergoes a relatively sharp
thinning from ∼ 25 to ∼10 km at a gradient of between 0.2 km/km
Fig. 6. Plot of the logarithm (base 10) of the stretching factors deduced from crustal
thinning across the Galicia and Flemish Cap margins against distance from the point
where the entire crust became brittle. In gray dash are simplified curves for the
stretching in the Galicia Interior Basin (GIB) and the Flemish Cap margin. The sum of the
two rift phases (solid grey line) fits the observed distribution of stretching across the
entire Galicia margin (solid black line), illustrating how an asymmetric margin pair can
be produced by two non-superimposed symmetric phases of rifting.
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T.J. Reston / Tectonophysics 468 (2009) 6–27
be due to the presence of a major, older rift phase on the Galicia side.
Of course, these thinning profiles are only a crude representation of
the true distribution of thinning both in time and space and should
not be taken as a true representation of the rift history between Spain
and Flemish Cap (especially give the complexities associated with the
rotation of the Cap — Sibuet et al., 2007), but do demonstrate that a
strongly asymmetric rift can develop through the occurrence of two or
more symmetric phases of rifting at different locations. Such a pattern
might arise if the locus of rifting migrates during slow rifting due
to strain hardening beneath the rift axis as the stretched mantle
lithosphere and upwelling asthenosphere both cool and increase in
strength (Kusznir and Park, 1987; Bassi, 1995).
4.1. Determining thinning at different crustal levels
It is tempting to relate an asymmetric crustal thickness distribution to an asymmetric rift process, such as the simple shear model
(Wernicke, 1985). However, in its simplest form, this model predicts
symmetric thinning of the whole crust — and symmetric total
subsidence — as long as the master shear zone passes through the
crust at a constant dip (Voorhoeve and Houseman, 1988; Fig. 7).
Observations from mature rifted margins (e.g. Le Pichon and Sibuet,
1981) that crustal thinning and total subsidence give the same
stretching factor does not mean that extension occurred by symmetric, uniform pure shear (McKenzie, 1978) as the simple shear
model predicts exactly the same relationship. Although some
conceptual models for margin development with more complex
detachment geometries (e.g. “delamination” — Lister et al., 1991) may
produce “upper plate” and “lower plate” asymmetries, these are not a
fundamental prediction of lithospheric extension along a throughgoing master shear zone. As a result, the absence of the detailed
“predictions” of these models does not exclude a simple shear origin,
even if as discussed below, such a model is rheologically unlikely.
However, asymmetric rifting processes such as the simple shear
model and depth-dependent stretching models do predict an asymmetry in the distribution of upper and lower crust (Fig. 7). Consequently, it is useful to consider the distribution of upper crustal and
lower crustal as well as whole crustal thinning. Systematic differences
between the patterns of upper crustal and lower crustal thinning may
indicate an asymmetric rifting process; the upper plate margin and
lower plate margin for a simple shear model show clear and distinguishable divergence in upper crustal and lower crustal thinning.
Thus if the thickness of different crustal layers can be mapped, they
can be used to constrain the distribution of strain during rifting. High
quality, modern wide-angle data would seem to provide a means to do
just this, but raise new problems, discussed below.
Reston (2007a) discusses the main problems with using seismic
velocity structure to investigate changes in the thickness of various
crustal layers. These are: 1) the velocity of rocks need not stay constant
during rifting, 2) the probability that crustal layers are not of uniform
thickness to start with, and 3) the accuracy of the velocity models.
Overall crustal thickness is relatively well determined, especially
when PmP reflections are recorded, but the location of intracrustal
velocity boundaries is less well resolved. Typically, high quality
modern wide-angle data can determine the large-scale velocity
structure to about 0.1 km/s (e.g. Dean et al., 2001), but are less
capable of resolving small structures (Van Avendonk et al., 2006) and
have a vertical resolution of ∼ 1 km. The implication is that trends of
the thickness of velocity layers are likely to be correct (Reston, 2007a),
although the precise degree of such changes less so.
The effect of uncertainities in both seismic velocity and in the location
of intracrustal boundaries can be estimated, as shown here for two
profiles across the Newfoundland Basin (Fig. 8). Although the uncertainty
in the stretching factor β of the upper crust and of the lower crust
become very large towards the continent–ocean transition, that in the
reciprocal (1 /β) or the thinning factor (1 − 1 /β) remain more constant.
Similar results were found for the Iberia Abyssal Plain margin (Reston,
2007a). However the overlap of the thinning factor ranges towards the
COT imply that any depth-dependent variation in thinning (depthdependent stretching) is too small to be resolved using seismic velocities.
Less quantifiable are the changes in seismic velocity of any given
rock that are expected during rifting as a result of reductions in
temperature and pressure, fracturing and alteration. The effects of
temperature and pressure tend to cancel out, but both fracturing and
alteration can lead to a net reduction in velocity. As a result, lower
crustal rocks (initially with relatively high lower crustal velocity) can
appear after rifting to be upper crust on the basis of their reduced
seismic velocity. As a result, excess thinning of the lower crust should
be viewed with suspicion. However when comparing the symmetry of
a pair of conjugate margins, similar effects might be expected on both
Fig. 7. Crustal structure predicted by extension of the lithosphere along a single throughgoing, planar shear zone, modified after Voorhoeve and Houseman (1988). A: starting model.
Lithospheric shear zone dips at an angle ϕ, cutting through crust (thickness Zc, equally divided into upper crust and lower crust (shaded)) and whole lithosphere (thickness ZL).
B: Extension sufficient to exhume mantle rocks by movement along the shear zone. C: as B, but after isostatic compensation by vertical shear. Note simple symmetric thinning of the
crust, producing a symmetric basin. D: detail of crustal structure (V.E. of 2) resulting from complete crustal separation along a single shear zone. Note overall symmetry of crustal
thinning and thus of subsidence, but asymmetry in the distribution of upper crustal and lower crustal thinning. E: plot of 1 − 1 / β for the upper, lower and whole crust as a function of
distance from the edge of the continental crust. Asymmetric rifting on a crustal scale should leave a similar pattern at rifted margins.
T.J. Reston / Tectonophysics 468 (2009) 6–27
15
Fig. 8. A, B, C: Plot of whole crustal, upper crustal and lower crustal parameters across the north Newfoundland Basin margin. A) apparent stretching factor β for the whole crust (βc),
the upper crust (βuc), and the lower crust (βlc) showing error bars for a ∼0.1 km/s uncertainty in seismic velocity and a 1 km uncertainty on the depth to the boundary between the
mid and lower crust. B) the corresponding plot of 1 / β showing uncertainties; C) the corresponding plot of thinning factors (1 − 1 / β) for the whole crust, upper crust (with
uncertainity) and lower crust (with uncertainty). Note that uncertainty is generally less than 0.1 and that at thinning factors greater than ∼0.7, the thinning factors are the same
within error for upper crust, lower crust and whole crust. D) Plot of whole crustal thinning vs upper crustal thinning (including uncertainty) at the north Newfoundland Basin margin,
showing no discernible depth-dependent stretching. E) Plot of thinning factors (1 − 1 / β) across the south Newfoundland Basin margin. Note uncertainty is generally less than 0.1 and
how at high thinning factors upper crustal, whole crustal and lower crustal thinning factors are all approximately equal within error. F) upper crustal thinning vs whole crustal
thinning across the South Newfoundland Basin, showing no discernible crustal depth-dependent stretching.
16
T.J. Reston / Tectonophysics 468 (2009) 6–27
sides, so the relative (between the conjugate margins) thinning of the
different layers should still be relevant.
Finally, it is unlikely that the crust consists of horizontal layers of
uniform thickness at the start of rifting. Furthermore, rifts are likely to
initiate or focus on inhomogeneities, e.g. terrane boundaries and
sutures across which the velocity structures are likely to change. An
example is the Galicia Interior Basin which follows at least in part a
Variscan terrane boundary (Pérez-Gussinyé et al., 2003) and as a result
has a very different velocity structure on either side. As focussing of
rifting is most likely to occur at such boundaries, it is not surprising
that the velocity structures of conjugate margins do not match. As a
result, I do not directly compare these, but rather the way the different
crustal levels appear to thin on each margin. Furthermore, I only
consider those margins where velocity layering is constrained by good
quality wide-angle data from the COT until the crust is at least 20 km
thick (i.e. at or close to the continental shelf) and concentrate on
conjugate pairs of margins where any major asymmetries could be
readily identified. Of the 22 margins shown in Figs. 2 and 3, only the 14
in Fig. 2 meet these criteria and are used for analysis of the distribution
of crustal thinning across conjugate margins.
Bearing in mind these concerns, it is nevertheless instructive to
discuss the apparent thinning of the different crustal levels toward
the margin. To remove the effect of initial layer and crustal thickness it is
necessary to normalise the data by plotting the thinning factor (1
− 1 / β) rather than thickness, an approach also adopted by Lau and
Louden (2007). The uncertainty in the determination of seismic
velocity of ∼ 0.1 km/s in the crust where velocity typically varies by
∼ 1 km/s represents a 10% error in the location of any given velocity.
The vertical resolution of 1 km provides an additional 1 km uncertainty. Using these bounds, maximum and minimum estimates of
the location of the boundary between the upper and lower crust
have been computed for the SCREECH2 and SCREECH 3 examples
and used to estimate errors on the thinning factors shown (Fig. 8).
Fig. 9 shows the variation in (1− 1 /β) for the whole crust (βc — solid),
the upper crust (βuc — long dash) and the lower crust (βlc — short dash)
for all margins in Fig. 2 where the structure from the COT toward the
shelf is constrained by modern wide-angle data. In most cases a midcrustal boundary or velocity contour (marked in bold in Fig. 2) has been
used to split the crust into upper and lower halves. Margins to the west
are shown black; those to the east in gray, allowing conjugate or nearly
conjugate margins to be displayed on the same plots The likely errors for
each margin are shown once even if the margin is plotted twice to avoid
unnecessary overlap with their conjugates: error bars for Porcupine
Basin and Rockall Trough are similar to the margins shown but are not
shown as thinning appears so close to symmetric that the curves would
overlap and hence obscure each other.
4.2. Evidence for asymmetry of rifting processes
As described above, the variation in the relative thickness of
different crustal layers, particularly when the converse can be
observed on the conjugate margin, may constrain the asymmetry of
the rifting process, which I define as an asymmetric distribution of
strain associated with one episode of rifting, such as that caused by
movement along asymmetric structures such as lithospheric-scale
shear zones. I begin by plotting thinning factors against distance from
the feather edge of the continental crust, chosen as an easier point to
define in most cases than the edge of the unthinned crust landward.
Apparent divergence between upper and lower crustal thinning
can be observed for IAM9, for the Galicia Interior Basin (GIB) part of the
Galicia profile, for the landward portion of the Nova Scotia margin
(SMART2) and for the Moroccan margin. In each of these cases,
the pattern might indicate a degree of asymmetry during a phase of
rifting.
The structure of the Galicia Interior Basin (on Galicia profile at
∼ 200 km from the COT) is dominated by east-dipping faults (Pérez-
Gussinyé et al., 2003; Reston, 2005, 2007b), which although not
lithospheric or even crustal shear zones do result in a local offset of
upper crustal and lower crustal extension. Overall though the basin is
close to symmetric. The conjugate SE Flemish Cap margin shows no
significant divergence between lower and whole crust, and fluctuations in the thinning of the upper crust appear to be related to
individual fault blocks but do not form a consistent asymmetric
pattern.
Probably the most pronounced example of heterogeneous crustal
extension comes from the Moroccan margin. Here, the lowermost
crust pinches out well over 120 km from the COT. The distribution of
thinning of the crust of the Moroccan margin has been interpreted by
Maillard et al. (2006) as evidence for a lithospheric-scale shear zone
and the pattern does resemble the “upper plate” margin expected
above such a shear zone, although the shear zone would have to be
very low-angle (b15°) to explain the distance between the last lowermost crust and the COT. However the conjugate Nova Scotian margin
(SMART1 — Funck et al., 2004) switches from possible “lower plate”
characteristics (the upper crust thinned more than the lower crust
between 100 and 200 km from the COT) to more upper plate characteristics (lower crust apparently absent, but a thin upper crustal
remnant) over the 50 km leading to the edge of continental crust.
Furthermore, along strike the asymmetry on the Nova Scotian margin
(SMART2 — Wu et al., 2006) also resembles an upper plate margin.
Thus the structure of these margins does not provide compelling
evidence for a simple asymmetry in the rifting process, but rather for a
more complicated pattern of non-uniform extension.
IAM9 (Fig. 2) most closely resembles the predictions of the simple
shear model, in this case for a “lower plate” margin (Fig. 9) below a
shear zone dipping at ∼ 15° to the west. However, the conjugate
Newfoundland Basin margin (SCREECH 2) shows only limited evidence for asymmetry. The fluctuations between upper plate and lower
plate characteristics 100–200 km from the COT represent fluctuations
in the location of the velocity contours: fine, largely unresolved
variations in velocity produce large apparent variations in the
thickness of the upper and lower crust. The last 100 km of continental
crust (0–100 km from the edge of continental crust) appear to show
more lower crustal than upper crustal thinning, consistent with an
“upper plate” margin, but these variations may be beyond the
resolution of the seismic velocities. A more prosaic explanation for
the structure of the IAM9 margin may be a combination of poorly
resolved velocity of the less extended crust landward (mainly
constrained by gravity data — Dean et al., 2001), leading to incorrect
estimates of the relative degree of thinning of the upper and lower
crusts, and the loss (or reworking as sediment) of some upper crust
through mass-wasting at the steep continental slope. It should also be
noticed that the heterogeneity of the IAM9 crust appears to disappear
as the crust thins to below ∼ 8 km (β N 4; 1 − 1 / β N 0.75).
The eastern margin of Porcupine also could be interpreted as a
typical upperplate margin, with apparently more rapid thinning of the
lower crust. However, as discussed below, the detachment imaged at
this margin cuts to the west, making the eastern margin a lower plate
margin if anything. Furthermore, the western margin displays similar
characteristics, although less pronounced. The thinning across the
Rockall Trough shows almost identical patterns of lower crustal
thinning on both sides, somewhat similar whole crustal thinning, but
rather different upper crustal thinning. This does not resemble a largescale simple shear but does perhaps indicate a degree of heterogeneity
of extension with the locus of upper crust thinning and separation
being shifted somewhat.
In summary, the relative distributions of upper crustal and lower
crustal thinning across these margins may indicate that extension is
heterogeneous and not uniform, symmetric pure shear, but does not
provide any evidence for asymmetry on the scale predicted by the
Wernicke model, particularly when the problems in using seismic
velocity as a proxy for crustal structure are borne in mind. The error
T.J. Reston / Tectonophysics 468 (2009) 6–27
17
Fig. 9. Plots of whole crustal, upper crustal and lower crustal thinning deduced from the velocity structure of 14 margins in Fig. 2. (S)NB — (South)Newfoundland Basin; SIAP — South
Iberia Abyssal Plain. Conjugate and near conjugates are plotted on same graph: black for western side, grey for eastern. Thinning of whole crust is solid line; upper crust long dash line
with vertical stripe error bounds, lower crust (whole crust–upper crust generally) as short dash line with grey error bounds. Lowermost crust for Moroccan margin plotted as short
bars. All are plotted against distance from the oceanward limit of continental crust except for Porcupine Basin where they are plotted from the centre of the basin. Bottom right: at
same scale thinning factors expected for shear zones cutting through 30 km thick crust at angles of 15, 30 and 45° for the upper plate margin and 15° for the lower plate margin. See
text for discussion.
bars for both lower crustal and upper crustal thinning are likely to
have a value of (1 − 1 / β) ∼0.1 (Fig. 8), sufficient to cast doubts on any
perceived asymmetry in the distribution of upper crustal and lower
crustal thinning. Interestingly, the margins tend to resemble upper
plate margins more than lower plate margins. This might represent a
systematic depth-dependent stretching (Driscoll and Karner, 1998),
but may also be influenced by the decrease in the seismic velocity as a
result of continuing extension (fracturing and alteration as described
above and by Reston, 2007a). The upper crustal velocities noted for the
Porcupine, Newfoundland Basin, Labrador, southern Iberia Abyssal
Plain (drilling transect) and northern Nova Scotia margin are
completely compatible with the presence of highly fractured lower
crustal rocks: basement has only been sampled on one of these
margins (the southern IAP), recovering dominantly lower crustal rocks
formed during the Variscan orogeny and uplifted through the Ar–Ar
closing temperature of feldspar of 200 °C at 137 Ma (Whitmarsh and
18
T.J. Reston / Tectonophysics 468 (2009) 6–27
Fig. 10. Illustration of the rheological evolution of the upper lithosphere during pure shear stretching. Initial weak zones within the mid-crust and lowermost crust gradually
disappear as the crust thins and creeping rocks become brittle (Pérez-Gussinyé and Reston, 2001). Once the entire crust becomes brittle, water can reach the mantle, causing
serpentinization. UC upper crust; LC lower crust; CMB crust–mantle boundary; WQ wet quartz; AN anorthite; OL olivine.
Wallace, 2001), that is during rifting, but with a seismic velocity of
∼ 5.5 km/s (Krawczyk et al., 1996).
The sections in Fig. 9 do however appear to show a large-scale
asymmetry between conjugate margin pairs of Labrador–West Greenland, Galicia–SE Flemish Cap, Newfoundland Basin–Iberia Abyssal
Plain, and a weaker asymmetry between Nova Scotian and Moroccan
margins, with in each case the latter margin being thinned more
abruptly that the former. Such an asymmetry may indicate the way
rifting focussed prior to breakup. However it is in part a function of
plotting the thinning curves vs distance from the edge of the continental crust. One possible influence on the development of
asymmetry may be the changing rheology of the lithosphere, and in
particular when the crust becomes brittle (Bassi, 1995).
5. Rheological evolution: embrittlement, serpentinization and
coupling
Apart from the thinning of the crust to zero and the lack of
magmatism at magma-poor margins, one of their other key characteristics (Figs. 1–3) is the presence of serpentinized mantle, both
within the continent–ocean transition (COT) oceanward of the last
continental crust, and beneath the feather edge of the crust. The
formation of these serpentinites results from the reaction of mantle
peridotites with water and thus requires that sufficient volumes of
water can be brought into the mantle. As Pérez-Gussinyé and Reston
(2001) pointed out, this can only be along faults linking the peridotites
with the surface, in turn requiring the whole crust to be brittle. As an
entirely brittle crust does not have decoupling ductile zones, whereas
serpentinites are relatively weak, serpentinization may be also be
linked to changes in tectonic style.
The development of an entirely brittle crust comes about through
changes in rheology accompanying lithospheric extension, changes
which have implications for the coupling between crust and mantle as
well as for crustal hydrogeology. It is generally well understood that the
interplay between downward increasing pressure and temperature and
changes in the dominant mineralogy of the lithosphere from granitic
(quartz-dominated), through mafic-gabbroic (anorthosite dominated)
to ultramafic (olivine dominated) results in a rheological stratification
and the development of weak zones where creep probably dominates in
the mid-crust and at the base of the crust, and intervening strong zones
including the seismogenic brittle upper crust which is 8–10 km thick
(Kusznir and Park, 1987; Jackson and White, 1989). As a result, initial
extension in the uppermost crust is likely to be by brittle faulting, in the
weak middle and lowermost crust by ductile creep, with perhaps
boudinage of any intervening strong layer. However these zones do not
remain fixed: as the crust thins and cools, the reduction in overburden
pressure and temperature means that rocks which originally deformed
by plastic creep gradually become brittle and fracture. The result is that
the initial weak zones in the mid-crust and deep crust disappear and the
entire crust becomes brittle (Fig. 10).
Fig. 11. Summary of numerical modelling results (Huismans and Beaumont, 2002), showing how deformation of a decoupled lithosphere (weak lower crust) results in overall
symmetric extension. However for a coupled lithosphere, extension is asymmetric as throughgoing structures develop.
T.J. Reston / Tectonophysics 468 (2009) 6–27
The embrittlement of the crust is perhaps counter-intuitive as
there is a common misconception that lithospheric extension brings
hot rocks up towards the surface and as a result, the lithosphere is
heated and rocks move from being brittle to ductile. However, even if
the vertical distance between the surface and the base of the
lithosphere decreases so that the geothermal gradient increases, the
temperatures of individual pieces of the lithosphere (rocks) do not
(see Reston, 2007b for a complete discussion).
The rheological evolution culminating in crustal embrittlement
has two important consequences. First, complete crustal embrittlement implies that brittle faults can for the first time cut down from the
surface into the mantle, bringing aqueous fluids that can serpentinize
substantial volumes of the mantle (Pérez-Gussinyé and Reston, 2001).
Second, Huismans and Beaumont (2002) show that the large-scale
symmetry of the rifting process is largely controlled by whether the
different lithospheric levels are coupled or decoupled. In the simplest
coupled models, through-going shear zones may develop, but in more
realistic models for lithospheric strength, the presence of weak
decoupling layers in the mid-crust and just above the Moho make
single through-going shear zones unlikely. As a result, deformation is
initially distributed among several structures and overall is approximately symmetric (Fig. 11). However, as extension proceeds and the
weak zones disappear, the strong layers become ever more tightly
coupled to another, and in particular the crust becomes coupled to the
upper mantle, allowing deformation at a late stage to become strongly
asymmetric (Reston and Pérez-Gussinyé, 2007).
The stretching factor at which the entire crust should become
brittle (βb) is to some extent controlled by rift duration and the
original lithospheric configuration (Pérez-Gussinyé et al., 2001 —
Fig. 12); changing the mineralogy and rheology of the lower crust in
particular may shift the curve up and down respectively within the
shaded area. The end-member lower crustal rheologies (dry quartz
and anorthite) probably underestimate and overestimate lower
crustal strength respectively, so also shown is an intermediate rheology calculated using polymineralic flow laws (Tullis et al., 1991) for a
50:50 mix of the two.
19
Although young hot orogens (e.g. Woodlark Basin) behave quite
differently (Pérez-Gussinyé et al., 2001), the results for both cooled
post-orogenic and cratonic lithosphere are quite similar (Fig. 12).
Increasing the rift duration leads to crustal embrittlement at lower
stretching factors as the lower crust has time to cool; moving from a
strong (anorthitic) to weak (dry quartz) lower crust increases the
stretching factor at which embrittlement occurs. It can be seen that
the stretching factor at which the entire crust becomes brittle is
dependent on both rift duration and the rheology of the lower crust,
itself dependent on lower crustal mineralogy and hence composition.
In general, the modelling predicts that the entire crust should
become brittle at stretching factors between about 3 and 5, depending
on rift duration. As noted by Pérez-Gussinyé and Reston (2001),
this corresponds reasonably well with the crustal thickness at the
landward limit of serpentinized mantle for several margins. The range,
based on uncertainties in rift duration and lower crustal composition, at
which the crust should have become brittle during rifting (Figs. 2 and 3)
tends to occur just landward of the landward limit of serpentinized
mantle, emphasising the correlation between crustal embrittlement and
serpentinization. At some margins however serpentinization only starts
a long way further oceanward, if at all. This probably reflects other
influences in the hydrogeology and permeability of the crust, including
the spacing of faults and the presence of thick sequences of sediments
especially evaporites (e.g. on the Moroccan margin).
5.1. Change from symmetry to asymmetry — the effect of embrittlement
It has been shown that conjugate margins appeared quite asymmetric when thinning profiles were plotted as a function of distance
from the edge of the continental crust (Figs. 4 and 9). Bearing in mind
the possible importance of coupling and decoupling in controlling rift
(a)symmetry, a more useful way of aligning the curves may be at the
point at which the crust becomes brittle as this might be expected to
mark a transition from symmetric (decoupled lithosphere) to a later
asymmetric extension (Whitmarsh et al., 2000; Reston and PérezGussinyé, 2007) when the development of large throughgoing shear
Fig. 12. Plot showing the stretching factor at which the entire crust should become brittle (βb) as a function of rift duration for two different lithospheric models (Pérez-Gussinyé et al.,
2001). Broad grey band — the width represents variations depending on the relative importance of feldspar (lower limit) and dry quartz (upper limit) in the lower crust. Grey dashed
line is for a 50:50 dry quartz/feldspar aggregate. Estimates of rift duration for margins shown (Nova Scotia — Funck et al. (2004), Wu et al. (2006); Morocco — Maillard et al. (2006);
Porcupine Basin — O'Reilly et al. (2006); Goban Spur — Bullock and Minshull (2005); Armorican — Thinon et al. (2003); South Newfoundland Basin — Reid (1994),Lau et al. (2006);
Galicia Bank — Mauffret and Montardet (1988), Reston (2007a,b); SE Flemish Cap — Tucholke et al. (2007); Iberia Abyssal Plain — Wilson et al. (2003); Minshull et al. (2001);Tucholke
et al. (2007); Rockall Trough — Pérez-Gussinyé et al. (2001); West Greenland, Labrador — Chalmers and Pulvertaft (2001); Reston and Pérez-Gussinyé (2007). As βb varies more with
lower crustal rheology than with uncertainty in rift duration at any given margin, the former is used to give first estimate of βb uncertainty (Galicia is illustrated) used to plot
predicted embrittlement in Figs. 2 and 3.
20
T.J. Reston / Tectonophysics 468 (2009) 6–27
zones is possible (coupled lithosphere — Huismans and Beaumont,
2002).
To separate out such late-stage asymmetry from any early
asymmetry, Fig. 13 shows the thinning curves of individual margins
shifted laterally (arrows) so that they do not coincide at a thinning
factor of 1 (the edge of the continental crust), but rather at a thinning
factor of between 0.67 and 0.8, corresponding to the stretching factor
βb of 3–5 at which the margin in question should become brittle
(Fig. 12 — Pérez-Gussinyé and Reston, 2001). Much of the apparent
asymmetry between margin pairs (hard vs soft — Fig. 4) is removed:
the margins are symmetric (curves overlap) from thinning factors
of zero (little or no extension) until βb is reached (thinning factor of
1 − 1 / βb). Only at thinning factors above this do the conjugate margin
pairs become asymmetric (Fig. 13). This implies that the asymmetry
for example between Labrador–W Greenland, between Newfoundland Basin (north and south)–Iberia Abyssal Plain, and between Nova
Scotian and Moroccan margins only developed once the entire crust
had become brittle and coupled, exactly as predicted by Huismans
and Beaumont (2002). The asymmetry between SE Flemish Cap and
Galicia Bank is reduced as the thin crust above the S reflector is
ignored in aligning the thinning curves, leaving an asymmetry that
can easily be explained by the overlapping of two rift phases (Fig. 6).
In each case where an asymmetry at a stretching factor greater
than βb can be detected, the margin not shifted to the right (i.e. the
margin with a broader expanse of very thin crust) appears to have the
seismic structure of an upper plate margin. However, as discussed
above, these characteristics are equally compatible with the reduction
of velocity (and hence the transfer of seismically defined lower crust
into upper crust) accompanying faulting, fracturing and pressure
reduction. Thus the velocities may simply be confirming that highly
thinned crust is also highly fractured and faulted.
5.2. Development of serpentine detachments
As discussed above, crustal embrittlement leads to both increased
coupling between the crust and the mantle and the serpentinization
of the mantle beneath the thinned and fractured crust. As serpentinites have a friction coefficient considerably below that of the most
rocks, they might be expected to decouple deformation between the
crust and the mantle. However, the serpentinites develop where fluids
can penetrate the mantle, that is along faults and fractures and so
instead of forming a general decoupling zone may tend to form
local detachments where major faults cut across the crust–mantle
boundary. As detachment faults are fundamentally asymmetric
Fig. 13. Thinning factor curves plotted against distance from point where entire crust should have become brittle (βb, filled black circle — based on rift duration and starting model).
Undercrusting serpentinites occur in the region marked by bold lines. Most of the asymmetries in Fig. 9 are removed, leaving variable widths of highly extended, completely brittle crust to
the left of the embrittlement point. This implies that most of the apparent asymmetry in crustal structure across conjugate margin pairs may develop after crustal embrittlement and
increased coupling between crust and mantle. The implication is that most of the extension is approximately symmetric, but that breakup and crustal separation are asymmetric.
T.J. Reston / Tectonophysics 468 (2009) 6–27
21
Fig. 14. Asymmetric P detachment beneath the Porcupine Basin (SW Ireland — A) and the S detachment at the west Galicia margin (depth image — B). Note that both detachments
developed when the crust was already thin, cut across the CMB and follow the boundary between crust and serpentinized mantle. Postrift sediments are partially blanked out and
crustal basement lightened to help emphasise fault block structure. C–F: illustration of the possible sequential evolution of the S detachment and overlying fault blocks during
polyphase faulting at the west Galicia margin. S may have started as a steep fault only becoming a low-angle detachment once the entire crust was brittle and mantle serpentinites
had formed. Note that S forms between blocks III and II and that block III is dismembered by 2 generations of later faults detaching onto S (D, E and F). Block III is the footwall to TBF
(“top basement fault”) which is cut by later faults and forms top basement to several small fault blocks above S (D–F).
structures, their presence (Fig. 14) is direct evidence for the
development of late-stage asymmetry once the crust had been prethinned to less than 8 ± 2 km (stretching factors between 3 and 5 —
Fig. 12) and thus become entirely brittle. The best known example is
the S reflector observed on the Galicia Bank margin; this structure cuts
to structurally deeper levels to the west. Its continuation may be
interpreted on the Flemish Cap margin (SCREECH 1 — Hopper et al.,
2004) where it follows the CMB, cutting to depth to the west (Reston
et al., 2007), but no east-cutting mirror-image is seen. A similar
detachment (P) imaged beneath the Porcupine Basin (Reston et al.,
2001) is also an asymmetric structure, cutting from the east flank of
the basin across the crust before following the CMB in the centre of the
basin and beneath its western flank (Fig. 14).
Reconstructions (Reston, 2005; Reston et al., 2007) suggest that
S developed as a high angle fault and only became a low-angle
detachment once the crust had thinned to about 7 km (Fig. 14),
compatible with the prior onset of serpentinization. Relationships
between wedges of synrift sediment and the block-bounding faults that
detach onto S indicate that S was active at angles down to or below 15°
(Reston et al., 2007). This can be explained by the extremely low friction
coefficients associated with serpentine and related minerals, or by the
transient development of high fluid pressures (Reston et al., 2007). The
development of detachments thus follows the onset of serpentinization,
itself a result of crustal embrittlement and the hydrogeology of the
fractured crust and is one stage in a long and complex polyphase faulting
history. This appears to have resulted in a focussing of extension towards
22
T.J. Reston / Tectonophysics 468 (2009) 6–27
the site of eventual crustal separation: the faulting above the S reflector
occurred later than the development of the larger fault blocks further to
the east (Reston, 2005) and the fault blocks overlying S developed
through the dismemberment of such larger, earlier fault blocks (Reston
et al., 2007). Similar polyphase faulting and focussing of extension
towards the COT is likely to have occurred at other rifted margins and
represents a simple focussing of strain at the weakest point of the rift,
that is the rift axis.
6. The extension discrepancy and depth-dependent stretching
Rifted margins form by the extension and eventual separation of
the crust and lithosphere. Extension results in thinning, apparent on
Figs. 2 and 3, and is accommodated by a combination of brittle faulting
on scales ranging from detachment faults through to sub-seismic
fractures, and ductile shear, flow and creep. As all lithospheric levels
are pulled apart, albeit possibly not in a uniform manner, the total
amount of extension going from the continental interior to the onset
of seafloor spreading should be equal at all levels. However, it is
commonly reported that the extension of the upper crust, as determined from fault geometries (βf), in much less than that required to
explain the crustal thinning (βc) either observed from wide-angle
velocity structure (Figs. 2 and 3) or deduced from subsidence (Fig. 15 —
Sibuet, 1992; Ziegler, 1983; Davis and Kusznir, 2004). This is the
extension discrepancy, expressed by rifted margins plotting in the
lower right quadrant of a (1 − 1 / βc) vs (1 − 1 / βf) plot, whereas
less extended rift basins plot along the diagonal (no extension
discrepancy — Kusznir and Karner, 2007).
There are two end-member explanations for the extension
discrepancy. First that crustal, as well as lithospheric, extension is
strongly depth-dependent and that at rifted margins the extension of
the upper crust is far less than the extension and thinning of the whole
crust. However, as noted above, the total amount of extension across
the whole margin must be equal at all lithospheric and crustal levels,
so that somewhere extension and thinning of the upper crust must
exceed that of the lower crust, although this has not yet been
observed. Alternatively, the other end-member explanation is that not
all the brittle extension has been measured, discussed later.
One possibility is that the margin exhibiting the extension discrepancy forms above a major shear zone, which transfers the deeper
crust out from beneath a virtually unfaulted upper plate margin.
However, as extension discrepancies are observed at virtually all
margins, this would imply that all are upper plate margins, the socalled upper plate paradox (Driscoll and Karner, 1998). Instead, more
symmetric patterns of depth-dependent stretching have been proposed, in which the lower crust on both sides of the incipient ocean
has somehow been displaced, leaving behind thin but apparently little
extended upper crustal section.
The sections in Fig. 2 and the thinning profiles in Fig. 9 can be used
to assess whether crustal depth-dependent stretching (as opposed to
lithospheric-scale depth-dependent stretching described above) is a
major process. The velocity structure and resulting thinning profiles in
Figs. 9 and 13 do show that many margins appear to exhibit upper
plate characteristics, although not consistent with movement along a
major crustal-scale detachment. However, such upper plate characteristics may simply represent the tendency for rifting to reduce seismic velocity and so give lower crustal rocks upper crustal velocities, as
implied by the presence of prerift lower crustal rocks with upper
crustal velocities on the southern IAP margin.
A degree of crustal depth-dependent stretching is to be expected as
it unlikely that the crust would extend perfectly uniformly, but the
problem with invoking DDS to explain the extension discrepancy is
the extreme amount of DDS required. By assuming that βf = βuc, and
that the extension discrepancy is caused by DDS, it is possible to invert
the βf (= βuc) and βc values and reconstruct a section, constrained
further by the condition that the upper crust can nowhere be thicker
than the whole crust (Fig. 15). The result is a section where the entire
crust thins from ~ 25 km (right) to at most 5 km (left), while the upper
crust maintains a maximum 5 km thickness. Given that the thickness
of the seismogenic brittle upper layer is 8–10 km (Jackson and White,
1989), this seems problematic.
The variation in the apparent thinning of the different crustal
layers shown in Figs. 9 and 14 does suggest a degree of heterogeneous
crustal extension (DDS), but does not show that this is sufficient to
explain the extension discrepancy. On a crossplot of upper crustal vs
whole crustal thinning (Reston, 2007a), significant crustal DDS should
Fig. 15. The extension discrepancy. A: The stretching factor deduced from fault geometries (βf) is considerably less than that deduced from crustal thinning (βc) across the Goban Spur
(A) and West Galicia Bank (B) margins. This means that these and other rifted margins plot below and to the right of the diagonal on a plot of thinning factors deduced from crustal
thickness and fault geometry data (C). This is the extension discrepancy. If fault geometries genuinely represent the amount of extension of the brittle upper crust, the data can be
inverted and thickness of the brittle upper crust and of the entire crust reconstructed and compared (D). As the upper crust can nowhere be thicker than the whole crust, the two
curves cannot cross, constraining the brittle upper crustal thickness to ∼3 km, which is unreasonably low. Even by maximising the thickness of the brittle upper crust (E), it only
reaches ∼ 5 km, still lower than evidenced by earthquake and rheological data. An alternative is that βf does not record all the extension of the brittle upper crust, meaning there no
need for depth-dependent thinning of the crust and that the reconstructions in D and E are incorrect.
T.J. Reston / Tectonophysics 468 (2009) 6–27
be represented as a systematic deviation away from the diagonal: if
crustal DDS is the cause of the observed extension discrepancy, the
data should plot in the lower right quadrant. However, all 14 margins
that provide sufficient control from the shelf to the COT plot close to
the diagonal (Fig. 16). Although several show some evidence for minor
crustal DDS (e.g., SCREECH 3 or Labrador where upper crust appears to
continue out over serpentinized mantle), the DDS is insufficient to
explain the extension discrepancy. Even where apparently observed
(e.g. SCREECH 3 and Labrador), the upper crust is also very strongly
thinned.
23
The problems in using velocity structure to infer the thinning of
different crustal levels do not help, as the general effect of rifting
would be to reduce velocity through fracturing and alteration, thus
giving lower crust rocks an upper crustal velocity and meaning that in
places the thin upper crustal layer such as observed on SCREECH 3
may actually be lower crustal rocks. This is borne out by the result of
ODP legs 149 and 173 across the west Iberia margin: although the
crustal blocks typically have velocities well below 6.5 km/s, the basement rocks sampled at Site 1067 and 900 were tonalites, amphibolites,
anorthosites and mafic granulites, i.e. mid-crustal to lower crustal
Fig. 16. Plots of upper crustal (1 − 1 / βuc) vs whole crustal thinning factors (1 − 1 / βc) for conjugate margin pairs from the North and North-Central Atlantic magma-poor margins. In
each case, the western half of the transect is shown black and the eastern half bold gray. In each case, the data plot close to the diagonal, indicating no significant depth-dependent
thinning resolvable from the velocity structure. Local deviations correspond to fault block topography, e.g. the lettered points which correspond to the letters on the corresponding
profiles (Fig. 2). A plot of the whole crust vs upper crustal thinning from all the margins (bottom right — including error bounds) emphasises that there DDS does not cause the
extension discrepancy.
24
T.J. Reston / Tectonophysics 468 (2009) 6–27
rocks which formed during the Variscan orogeny and which were
exhumed through 200 °C at 135 Ma, during early rifting (Whitmarsh
et al., 2000). Thus the geophysical “upper plate” characteristics of the
highly thinned margin near the COT are misleading and cast doubts on
the interpretation of other highly thinned crust as being exclusively of
upper crustal origin and evidence for crustal DDS on the scale required
to explain the extension discrepancy.
The plot in Fig. 16 only covers the 14 conjugate margins where the
velocity structure can be traced far enough to estimate the relative
thinning of the crustal layers. However, the sort of DDS required to
explain the extension discrepancy (Fig. 15) is also unlikely to apply at
several other deep margins as these exhibit velocity structures that are
comparable with those were crustal DDS can be ruled out (Fig. 16). For
instance, both the thin Goban Spur section and the deep Armorican
margin (Fig. 3) have upper crustal velocities around or below 6 km/s
underlain by lower crustal velocities close to 7 km/s, very similar to
the other margins where such layers can be traced landwards. At both
margins, the velocity structure suggests the presence of lower crust
right out near the COT. Although several margins (Labrador,
Porcupine, S. Newfoundland Basin, do show a thin layer of upper
crustal velocities extending tens of km further oceanward than the
higher velocity deeper layers, in each case the upper crustal layer is far
thinner than its equivalent further landward, requiring far more
extension than apparently observed: the data plot in the top right not
lower right quadrant of the thinning factor crossplot (compare Figs. 15
and 16). Furthermore, the only margin where this anomalous layer
with only upper crustal velocities has been sampled (the IAP margin
along the drilling transect) has been found to contain both upper
crustal and lower crustal rocks shuffled together by extensional
tectonics; the seismic velocity of the lower crustal lithologies has been
reduced by intense fracturing and alteration.
In summary, although it is quite possible that significant crustal
DDS does occur at other, warmer margins where lower crustal flow is
possible (Hopper and Buck, 1996), there is no evidence that the
extension discrepancy is caused by crustal DDS at any of the magmapoor margins of the North and Central Atlantic. The extension
discrepancy thus needs another explanation. One possibility (Reston,
2005, 2007a) is unrecognised faulting, such as polyphase faulting and
large unrecognised top basement faults. Put another way, the
extension discrepancy can be summarised as βf ≪ βc. In the DDS
explanation, it was assumed that βf ∼ βuc and that as a result βuc ≪ βc.
However, above we have shown that βuc ∼ βc, and thus by implication
that βf ≪ βuc. This means that the seismic images are somehow not
measuring all the extension and thinning of the upper crust.
The amount of brittle extension may be underestimated if either
distributed deformation accompanies the observed faulting, if the
extent of the observed faults has been substantially underestimated, or if
the observed faults are not the only phase of faulting to have occurred.
Davis and Kusznir (2004) show that distributed deformation is
insufficient to explain the extension discrepancy and also argue that
the discrepancy cannot be caused by a phase of faulting that post-dates
the one imaged. However, there is strong evidence at some margins (e.g.
Galicia) for faulting and pre-thinning of the crust prior to the
development of the latest fault blocks (Reston, 2005). There are good
reasons why complex polyphase faulting (Proffett, 1977; Jackson and
White, 1989) may be difficult to image, especially on time sections
(Reston, 2007a), and the same is true for faults that have flexurally
rotated to follow top basement (Reston, 2007a) so that their extent can
be underestimated. In short, there are ample reasons why earlier phases
of faulting are not always recognized at rifted margins (Reston, 2005,
2007a). However, a long and complex rifting and faulting history can be
deduced from field studies (Manatschal, 2004), is predicted by
numerical modelling (Lavier and Manatschal, 2006) and is completely
compatible with the structure and rift history of rifted margins (Reston,
2005, 2007a). A single generation of faults should not accommodate
more than about 100% extension (stretching factor of 2) before rotating
to low-angle and locking up, so that the highly extended deep margins
should have undergone extension along several generations of faults
(Reston, 2005). Indeed, Lavier and Manatschal explicitly recognise three
phases of extension during the formation of a rifted margin, each with
distinctive modes of brittle extension. Fig. 14 shows how simple
polyphase faulting can both closely reproduce the sort of geometries
observed on the seismic images, and be difficult to recognize. It should
not be considered a restoration as there is too little calibration by well
data to confirm the interpretation.
7. Unroofed mantle within the COT
A final feature of many of the margins discussed here is the
presence of unroofed mantle peridotites within the COT (Figs. 2, 3
and 17). These appear to be the geophysical continuation of the
reduced velocity mantle undercrusting the feather edge of the
Fig. 17. A: cartoon after Pickup et al. (1996) showing exhumation of a broad expanse of mantle peridotites within the continent–ocean transition. However this illustration fails to explain
how the peridotites were unroofed, i.e. how the overlying rocks were removed B: detail of IAM 9 (approximate position marked by box in A) showing mounded layer (serpentine breccias —
Reston et al., 2004) above serpentinized basement, interpreted as being topped by series of top basement faults that can be traced to depth as landward-dipping reflections.
T.J. Reston / Tectonophysics 468 (2009) 6–27
continental crust. However geophysics cannot discriminate between
sub-continental lithospheric mantle that has been exhumed within
the COT, and exhumed asthenospheric mantle. These distinctions are
important to choose between the different explanations for magmapoor margins and to constrain the large-scale kinematics of the
stretching process. If the unroofed mantle is new lithosphere resulting
from the cooling of asthenosphere, it must have either been cool or
depleted to explain the lack of melting (Reston and Phipps Morgan,
2004; Pérez-Gussinyé et al., 2006). If it is old sub-crustal lithosphere
pulled out from beneath the continent, then lithospheric-scale
extension must have been strongly heterogeneous, helping to explain
some if perhaps not all, of the melt-deficit at magma-poor margins
(Reston, 2007b). Although the prevailing view is that the exhumed
mantle was sub-continental lithosphere in nature (e.g., Müntener and
Manatschal, 2006), others dispute this (e.g., Abe, 2001). Further
sampling work is needed to distinguish between these models.
In either case, the mantle must have been exhumed through the
removal of overlying portions of the lithosphere. As final exhumation
to the surface must have occurred in the brittle regime, everywhere
this unroofed mantle must be topped by faults which may predate
the structures that offset the top of the mantle basement and which
define the fault block topography within the COT (Fig. 17). The
process of mantle exhumation thus is the result of continued tectonic extension as the crust thins through zero and faults exhume
and expose at the seafloor the underlying mantle (Manatscha et al.,
2001).
Finally, the distinction between unroofing mantle within the
continent–ocean transition and seafloor spreading is partly a matter of
degree. Slow-spread crust is characterised by mantle unroofing,
suppressed magmatism, and detachment faulting (e.g., Tucholke and
Lin, 1994; Reston et al., 2002), features also apparent within the COT.
25
Although magmatism may be more important during the formation of
oceanic crust, probably due to increased focussing of mantle
upwelling and hence melting beneath the spreading centre (Minshull
et al., 2001), perhaps the most meaningful definition of the boundary
between continent and ocean is the boundary between old and new
lithosphere. This has implications for the nature of the COT: the thin
oceanic crust interpreted by Hopper et al. (2004) on profile SCREECH 1
may certainly be thin magmatic crust, but if it is floored by unroofed
sub-continental lithosphere may not be truly oceanic as it would then
not be underlain by oceanic lithosphere.
8. Conclusions
This paper has focussed on the large-scale structure and symmetry
of magma-poor rifted margins whereas the structural evolution has
been discussed more elsewhere (Reston, 2007b). Their overall
structure is the logical result of progressive extension, summarised
in Fig. 18. A number of key elements and stages in the evolution of
such margins can be identified
1. The extreme thinning of the crust (Fig. 18A–D) is the result of
multiple phases of extensional tectonics: that not all of these are
recognised is to be expected, leading to the extension discrepancy.
Substantial depth-dependent stretching of the crust is neither
allowed by the velocity structure of the margins nor expected for a
cool lithosphere with an initially substantial brittle lid.
2. Rifting appears to be a largely symmetric process on a crustal-scale
until the entire crust has become brittle (Figs. 15 and 18C,D),
although the distribution of extension with time may lead to
asymmetric rifts. When this occurs, the complete coupling
between the crustal and mantle allows the development of single
Fig. 18. Schematic model for lithospheric extension resulting in the formation of conjugate magma-poor margins modified from Reston (2007b). A: Initial lithospheric extension by
∼100% (β of 2). Rift focusses above axis of mantle thinning: polyphase faulting in the crust thins this to less than a quarter of its original thickness, until the entire crust is brittle.
(B). Weak serpentinites develop at the base of the thinned crust (B,C), allowing the development of serpentine detachment (C) as observed west of Iberia (S, H) and in the Porcupine
Basin (P). D: Movement along the serpentine detachment and on subsequent detachments results in complete crustal separation and the unroofing of mantle rocks within the COT.
26
T.J. Reston / Tectonophysics 468 (2009) 6–27
large asymmetric faults and shear zones which leave much of the
highly thinned crust on one side of the conjugate margin pair,
producing a late-stage asymmetry. Early asymmetry, such as
between Flemish Cap and Galicia is probably due to the distribution
of different rift phases rather than due to a large-scale asymmetry
in the rifting process.
3. The undercrusting of the crust by serpentinized mantle is the
logical result of the complete embrittlement of the crust once a
stretching factor of about 4 has been reached (Fig. 18B–C).
Serpentinites are only observed where the overlying crust is so
thin that should have become brittle during rifting.
4. The serpentinized mantle beneath the crust is laterally continuous
with serpentinized mantle that forms the basement between the
last continental crust and the first unambiguous oceanic crust.
Mantle unroofing is the logical result of continued extension once
the entire crust has become brittle and detachment faults develop
in the weak mantle serpentinites.
5. The transition to seafloor spreading (Fig. 18D) is marked by a
transition to more magma-dominated and less tectonic divergence,
but may be most rigorously defined as a transition from old to new
lithosphere, marked by the transition from exhumed mantle
peridotites were originally part of the continental lithosphere to
those that represent younger lithosphere formed by the cooling
and depletion (through melt extraction) of asthenosphere.
One important control on the development of such margins
appears to be the rheological evolution of the lithosphere, with crustal
embrittlement both allowing mantle serpentinization and controlling
the transition from a decoupled to a coupled lithosphere with a
related change from symmetric to late-stage asymmetric extension.
Acknowledgements
This paper has benefited from numerous discussions with
colleagues, in particular Marta Pérez-Gussinyé, Jonathan Turner and
Gianreto Manatschal. It is the result of a long programme of support by
funding agencies, in particular the DFG.
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