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A Dynamic Model of Rifting Between Galicia Bank and Flemish Cap During Opening of the North Atlantic Ocean
Dennis L. Harry1 and Shelly Grandell2
1Department of Geosciences, Colorado State University, Fort Collins, CO, 80523-1482 ([email protected])
2Department of Geology, Adams State College, Alamosa, CO. 81102
1
Abstract
A finite element model is used to simulate Late Jurassic through Early Cretaceous rifting
between the Flemish Cap and Galicia Bank continental margins. The model results shows that
variations in the thickness of the continental crust on these margins at wavelengths greater than
about 75 km can be explained as a consequence of the interaction of two preexisting weaknesses
in the lithosphere. A weakness in the crust, attributed to structural fabrics in the Variscan front,
controls the location of crustal extension during the early stages of rifting in the model. This
results in formation of a broad rift basin similar to the Galicia Interior Basin. A deep seated
weakness located 110 km further west, attributed to the thick crust beneath the central Variscan
orogen, controls the location of mantle necking. Extension in this region is initially diffuse, but
accelerates and becomes more focused with time. 13 m.y. after rifting begins, the locus of crustal
extension shifts from the region of pre-weakened crust into the region of pre-weakened mantle.
This marks the end of subsidence in the Galicia Interior Basin and the onset of subsidence in the
Flemish Cap and Galicia Bank marginal basins. Extension in these areas continues for another 12
m.y. before continental breakup. The asthenosphere does not ascend to depths shallow enough
for decompression melting to begin until < 5 m.y. before the onset of seafloor spreading. The
model predicts that all late stage syn-rift magmatism during this period is limited to within 45 km
of the rift axis, and production of melt thicknesses greater than 2 km is restricted to within 35 km
of the rift axis. Mantle potential temperatures of 1250 !C to 1275 !C, ca. 30 !C cooler than
normal, result in 3.1-4.5 km thick oceanic crust at the time of breakup, in general agreement with
the 2-4 km thick crust observed adjacent to these margins.
2
Introduction
This paper describes the results of a finite element numerical simulation of Late
Mesozoic and Early Cenozoic rifting on the Galicia Bank and Flemish Cap segments of the
Iberia and Newfoundland conjugate continental margins. The modeling goal is to match the
long-wavelength (> 75 km) crustal thickness variations on the margins, the timing and duration
of rifting, the magmatic history of the margins, timing and location of regional shifts in
extensional tectonism, and the regional subsidence history. This paper focuses on a conjugate
transect between seismic profiles ISE-1 on the Iberian margin and SCREECH-1 on the
Newfoundland margin (Fig. 1) (Funck et al., 2003; Henning et al., 2004; Hopper et al., 2004)].
The modeling technique used in this paper is based on the STRCH95 two-dimensional
finite element modeling program, which is a descendant of the STRCH program developed by
Dunbar (1988). STRCH95 is used to model extensional processes on a lithosphere scale. The
program allows for complex initial conditions to describe the pre-extension structure of the crust
and rheology and temperature of the lithosphere. The simulations discussed in this paper focus
on how such pre-existing features might have controlled the pattern of extension and shifts in the
loci of extension with time on the Galicia Bank/Flemish Cap margins, and how such extension is
accommodated without producing significant amounts of syn-rift magmatism.
In the following sections we describe the geological and geophysical observations that
constrain the numerical models discussed in this paper and the modeling targets. We then
describe the modeling algorithm. Finally, we describe the results and discuss their implications
within the context of Galicia Bank/Flemish Cap rifting in particular and rifting on non-volcanic
margins in general.
3
Model Constraints
The history and structure of the Galicia Bank and Flemish Cap continental margins is
described in detail elsewhere in this volume. Here, we briefly summarize those elements of rift
history and margin structure that contribute directly to constraining the numerical simulations
presented in this paper.
Timing - Extension on the Iberian Peninsula and Newfoundland occurred in several
distinct phases. The first phase occurred in the Newfoundland and Galicia Interior Basins during
the Late Triassic-Early Jurassic periods (Fig. 1) (Ziegler, 1989; Murillas et al., 1990; Tucholke et
al., 1989; Foster and Robinson, 1993; Tankard and Welsink, 1987). This was followed by a
period of tectonic quiescence lasting until Late Jurassic time. The second phase of extension
began in the Late Jurassic Epoch and lasted until late Valanginian time, when the locus of
extension shifted seaward of the Flemish Cap and Galicia Bank (Murillas et al., 1990; Tucholke
et al., 1989). This led to the third phase of extension, a period of lithospheric stretching and
crustal thinning west of Galicia Bank and east of Flemish Cap in Hauterivian through Aptian
time. This last phase of extension continued until continental breakup and formation of new
oceanic crust between the time of seafloor spreading anomalies M3 and M0 (Srivastava et al.,
2000; Hopper et al., 2004; Funck et al., 2003; Minshull et al., 2001).
In this paper, we focus on the Late Jurassic to Early Cretaceous phases of extension.
These phases involved an initial stage of extension in the Galicia Interior Basin, a later stage of
extension seaward of the unextended continental crust on Flemish Cap and the slightly extended
crust on Galicia Bank, and, finally, the onset of seafloor spreading. The total duration of rifting
(excluding the Triassic through Early Jurassic phase in the interior basins) lasted approximately
25 m.y., beginning in Late Tithonian time (ca. 143 Ma) and ending in the early Aptian age (ca.
118 m.y.) (Hopper et al., 2004; Minshull et al., 2001). The first ca. 12 m.y. of extension
4
(Tithonian-Valanginian) involved the Galicia Interior Basin and Newfoundland Basin. The last
ca. 13 m.y. of extension was primarily focused seaward of the Galicia Bank and Flemish Cap.
Postrift crustal structure – A reconstruction of the Newfoundland and Galicia
continental margins at the time of anomaly M0, near the onset of seafloor spreading, is shown in
Fig. 2. Unextended portions of the continental crust on Iberia and Newfoundland are 30-32 km
thick (Banda, 1988; Diaz et al., 1993; Mendes Victor et al., 1980). The pre-rift crystalline crust
on the Iberia margin thins westward to less than 10 km-thick beneath the 100 km-wide Galicia
Interior Basin (Pérez-Gussinyé et al., 2003), increases to approximately 20 km beneath the
Galicia Bank (Gonzalez et al., 1999; Pérez-Gussinyé et al., 2003), and then thins progressively
westward over a 100 km-wide region between the eastern edge of the Galicia Bank and the rift
axis, where the oldest oceanic crust is emplaced. The pre-rift crust reaches a minimum
thickness of 2-5 km at its seaward limit (Gonzalez et al., 1999; Chian et al., 1999). On the
Newfoundland margin, the crust is approximately 30 km thick beneath the relatively unextended
Flemish Cap, and thins progressively eastward to ca. 1.5 km over a ca. 80 km-wide region
between the eastern edge of the Flemish Cap and the rift axis (Funck et al., 2003). Mantle
exposed in a ca. 10 km-wide region on the seafloor west of Galicia Bank is thought to have been
exhumed by late-stage low angle faulting during the last 1-2 m.y. of extension (Fuegenschuh et
al., 1998; Pérez-Gussinyé et al., 2003; Pickup et al., 1996; Reston, 1996; Sibuet, 1992).
Extension rate - Extension rates on the Newfoundland and Iberia margins vary locally at
the scale of a few tens of km, probably reflecting different timing between movements on
individual fault systems. In a regional sense, extension rates are best constrained on the Iberia
Abyssal Plain. Here, regional extension rates are estimated to be ca. 10 mm/yr during rifting
(Minshull et al., 2001) and 10-14 mm/yr at the onset of seafloor spreading (Whitmarsh et al.,
5
2001a; Russell and Whitmarsh, 2003). Assuming symmetric spreading, this corresponds to a
whole-extension rate of 20 mm/yr during rifting. This is consistent with area balancing of the
cross section in Fig. 2. Restoration of the cross section to a uniform crust thickness of 32 km
requires an extension rate of 20 mm/yr to produce the observed amount of stretching in 25 m.y.
Pre-rift lithospheric structure, rheology, and thermal constraints – Breakup between
the Galicia Bank and Newfoundland occurred within a region roughly coincident with the trend
of the Late Paleozoic Variscan orogen in this region (Capdevila and Mougenot, 1988; Ziegler,
1989). Focusing of extensional deformation in this region was most likely a result of orogenic
weakening of the lithosphere caused by the increased thickness of the crust beneath the orogen
(Fig. 3). The thickness of the crust in the central Variscan orogen near the rift axis is estimated
to have been ca. 34-35 km thick prior to extension (Pérez-Gussinyé et al., 2003), 2-5 km thicker
than the crust in the unextended parts of the Iberia Peninsula and Newfoundland. Assuming the
thermal parameters estimated by Tejero and Ruiz (2002) for the crust in the Duero Basin in the
Iberian Peninsula and a simple two-layer rheological model following empirical ductile flow
laws for quartz-diorite in the crust and wet dunite in the mantle (Table 1), this difference in
crustal thickness would weaken the lithosphere in the central Variscan orogen by up to 15% in
comparison to the surrounding regions (Fig. 3). The strength of the nominal lithosphere
predicted by this rheological model is 7.0x1012 N m-1 (Fig. 3a), which is moderately weaker than
the 8x1012 N m-1 lithospheric strength estimated by Tejero and Ruiz (2002) for the central
Iberian Peninsula and is consistent with the heat flow and thermal structure of the Iberian
lithosphere estimated by Fernandez et al. (1998).
Magmatic history – Both the Flemish Cap and Galicia Bank margins are flanked by
unusually thin oceanic crust, ranging from 2.5-4 km thick (Hopper et al., 2004; Whitmarsh et al.,
6
1996), suggesting a low volume of melt production immediately following continental breakup.
On the Galicia margin the oceanic crust increases to normal thickness (approximately 7 km)
within 15 km seaward of the oldest oceanic crust (Whitmarsh et al., 1996). At the estimated
spreading half-rate of 10-14 mm/yr (Whitmarsh et al., 2001a), this suggests that the thermal
regime of the mantle rapidly evolved to that of a typical mid-ocean ridge system within ca. 1-1.5
m.y. after the onset of seafloor spreading. The melt production history after the onset of seafloor
spreading on the Newfoundland margin is more complicated. Here, the oldest oceanic crust is
ca. 3-4 km thick and thins seaward to < 1.3 km over a ca. 50 km wide region (Hopper et al.,
2004), indicating a magma starved environment during the first ca. 3.5-5 m.y. of seafloor
spreading. Near-normal oceanic crust (> 6 km thick) located 3 km seaward of the thinnest
oceanic crust on the Flemish Cap margin (Funck et al., 2003; Hopper et al., 2004) indicates that
the transition from magma starved seafloor spreading to normal seafloor spreading occurred in
about 0.2-0.3 m.y. Recovery of late syn-rift volcanic and plutonic rocks seaward of the oldest
oceanic crust on both the Newfoundland and Iberian margins suggests that the last stages of
rifting were accompanied by minor amounts for magmatism, just prior to the onset of seafloor
spreading [Leg 210 Shipboard Scientific Party, 2003; Whitmarsh et al., 2001b; Manatschal and
Bernoulli, 1999).
Regardless of the complexities of late rift-stage and early spreading phase melt
production, there is broad consensus that for most of the rift history the margins were
amagmatic, or at least involved very little melt production. Volcanism that occurred prior to the
onset of seafloor spreading appears to have been limited in volume and space, being confined to
within 20 km or so of the location where oceanic crust is first produced.
7
Modeling Targets
On the basis of the above discussion, our numerical simulations of rifting between
Galicia Bank and Flemish Cap are constrained by 1) the pre-rift structure of the crust, including a
presumably weak and heterogeneous crust in the Variscan orogen embedded between older
strong Precambrian crust of Canada and the Iberian peninsula, 2) the pre-rift crustal thickness,
taken to be 32 km by comparison to the thickness of the modern unextended crust beneath Iberia
and Canada, 3) radiogenic heat production of the crust, taken to be similar to that of basement
rocks exposed in Iberia (Table 1). The models seek to reproduce 1) the post-rift crustal thickness
variation across the margins at length scales greater than ca. 75 km, particularly moderately
extended crust beneath the Galicia Interior Basin, unextended crust beneath the Flemish Cap and
slightly extended crust beneath Galicia Bank, and highly extended crust beneath the marginal rift
basins, 2) the lack of magmatism during rifting, 3) the duration of rifting, 4) the amount of
magmatism required to produce thin oceanic crust at the time of breakup, and 5) the subsidence
history of the margin, including Late Jurassic through late Valanginian subsidence in the Galicia
Interior Basin and late Valanginian through early Aptian subsidence seaward of Flemish Cap and
Galicia Bank. The finite element model formulation used in these simulations does not explicitly
account for formation of detachment faults, and so does not reproduce exhumation of the
peridotite ridge west of Galicia Bank.
Modeling Method
The finite element modeling program used in this research, STRCH95, is the most recent
version of the FORTRAN 90 computer program STRCH developed by Dunbar (1988). The
STRCH95 algorithm successively solves the two-dimensional heat
conduction/advection/generation equation
8
Equation 1: ATKtTc "#$#%&&'
and the Navier-Stokes equation for flow in a visco-plastic lithosphere
Equation 2: uPFtdu 2#"#(%&
)'' ,
where u is the particle velocity, T is temperature, ' is the density of the rock, c is the specific
heat, K is thermal conductivity, A is the volumetric heat production rate, F is body force per unit
mass, P is pressure, and ) is viscosity. The physical properties ', c, A, and K for each element
in the finite element mesh are specified by the user. Viscosity is defined by the rheology of each
element, and is pressure, temperature, and strain rate dependant. Viscosity is determined by
empirical relations derived from experimental rock deformation studies. Initially, a uniform
strain rate,*! , is assumed throughout the model. An effective viscosity, )eff , for each element is
determined from the estimated stress + and strain rate:
Equation 3: *+)!
%eff .
The stress in the above equation is taken to be the weaker of either a ductile (power law creep) or
plastic (a linear pressure-dependant yield criterion) rheology:
Equation 4 (ductile): nRTQn
ceA
//1
,-.
/01%*+!
Equation 5 (plastic): BzS "% 0+
where z is depth and R is the Universal Gas Constant. Once the effective viscosity is
determined, the Navier-Stokes equation is solved. This results in an updated estimate of the
strain rates in the model, which are used to revise the estimated effective viscosity of each
element. The process is repeated until the strain rate estimates on successive iterations agree to
9
within a convergence criterion prescribed by the user. The strain rates are used to determine
velocities at each node in the mesh, and the geometry of the model mesh is updated by stepping
forward the locations of each node in the mesh a finite amount of time (time steps of 10,000
years duration are typical). The heat equation is then solved to update the thermal structure,
using the new mesh geometry as a boundary condition and the thermal state of the model at the
previous time step as an initial condition. The new model temperatures and the strain rates from
the past time step are used to calculate a revised estimate of the element viscosities to begin the
next time step.
Accuracy of the model solution is assessed empirically by comparing model results
obtained with coarse and fine meshes and coarse and fine time step sizes. For a given family of
models, a coarse mesh and time step size is initially selected. The mesh size and time step size
are then halved and the model re-run. Refinement of mesh size and time step size continues until
successive models produce the same result. If the model becomes highly deformed (individual
elements develop large aspect ratios), a singularity develops in the linear system of equations
that define the finite element problem being solved. When this occurs, the model is remeshed.
During remeshing, domain boundaries that define different materials used in the model (e.g., the
model edges and the contacts between different rock types) are preserved. Within each domain,
new elements are defined that have an aspect ratio as close to unity as possible. Temperatures,
pressures, and strain rate information within the model are preserved during the remeshing
process.
Model Description
A basic assumption behind the modeling strategy used here is that pre-existing strength
heterogeneities in the lithosphere determine the location and distribution of loci of extension
10
during rifting (e.g., Dunbar and Sawyer, 1989). Weaknesses in the upper and middle crust
(crustal weaknesses) tend to produce short-lived extensional provinces during the early stages of
rifting, whereas deep seated weaknesses (mantle weaknesses) control the location where
lithospheric necking develops and, hence, the location of eventual continental breakup. If these
weaknesses are offset from one another, interior rift basins (controlled by the crustal weakness)
often form landward of the deep offshore marginal basin (controlled by the mantle weakness)
(e.g., Harry and Sawyer, 1992).
The models presented here include both of these types of weakness in the lithosphere
(Figs. 3 and 4). Mantle weaknesses are simulated as regions of thickened crust, which is
attributed to the root beneath the Variscan orogen prior to rifting. This is represented with a
simple triangular shaped region of increased crustal thickness. Key variables describing the
mantle weakness are the width of the region of thickened crust and the amount that the crust is
thickened. A crustal weakness is simulated by a region in which both the plastic and ductile
yield strengths are decreased relative to the surrounding crust. In the plastic regime, this
involves a reduction of S0 (equation 5). For ductile behavior, both the temperature dependant
terms (Qc/n) and temperature independent terms (A-1/n) are reduced by similar amounts (equation
4). The crustal weakness may be attributed to lithologic variations in the upper and middle crust
or to pre-existing faulting near the Variscan front. Key variables describing this weakness in the
model are the amount by which the yield strength is decreased, the width and thickness of the
weakened crust, and the position of the crustal weakness relative to the mantle weakness. All of
these variables were iteratively adjusted to produce a model that best fits the geometry and rift
history of the Flemish Cap and Galicia Bank margins. Other modeling parameters, including
heat production, specific heat, thermal conductivity, extension rate, and nominal thickness of the
11
crust (outside the Variscan orogen) were held fixed (Table 1). The criterion used to determine
when rifting ends in the models is the requirement that the crust thin to ca. 2 km at the rift axis,
which is similar to the thinnest continental crust on the Galicia Bank and Flemish Cap margins
(Chian et al., 1999; Funck et al., 2003; Gonzalez et al., 1999).
Results
Results of a typical model are shown in Figs. 5-7. The model simulates in general the
present structure of the Galicia Bank/Flemish Cap conjugate margins. In particular, the model
predicts extreme crustal thinning beneath conjugate offshore rift basins that form during the late
stages of rifting, formation of an interior basin on the eastern shelf during the early stages of
rifting, and a region of less extended crust between the offshore and interior rift basins (Fig. 6).
Most of the models examined in this study produced these basic attributes. Exceptions were
models in which either the crustal or mantle weakness was very small (typically involving a
reduction of yield strength in the crustal weakness of less than 5% or a change in crustal
thickness in the mantle weakness of less than 2 km). In those models (not shown here), the
dominant weakness controls rifting from the outset, resulting in formation of a nearly symmetric
rift basin centered on the dominant weakness. All other models that were examined, involving
changes in yield strength in the crustal weakness ranging from 5-30% and changes in crustal
thickness in the mantle weakness ranging from 2-5 km, produced results that were generally
similar to that shown in Figs. 5-7. However, changes in the relative locations and magnitudes of
the two weaknesses produces a wide variation in the widths and depths of the interior rift basin
and deep offshore basins, the duration of extension in the interior basin, the time elapsed before
breakup, the amount and pattern of crustal thinning, and the presence or absence of a region of
relatively unextended crust between the deep offshore and interior rift basins. As discussed
12
below, changes in these parameters also have a significant affect on the magmatic history of the
model.
The model shown in Figs. 5-7 is based on the variables that provide the best fit to
observations on the Galicia Bank/Flemish Cap conjugate margins (Table 2). During the first 10
m.y., extension results in rapid thinning of the crust in the region encompassing the crustal
weakness (Figs. 5 and 6a). Crustal thinning also occurs above the mantle weakness, but at a
much slower rate. Tectonic subsidence in the region of weakened crust begins almost
immediately after the onset of extension, marking the first stages of formation of the interior rift
basin (Fig. 6b). The region above the mantle weakness, where the deep offshore basin ultimately
forms, does not subside below sea level until approximately 5 m.y. later. The two basins are
separated by a structural high that undergoes moderate crustal thinning (from 32 km to ca. 20
km) during the first 5 m.y. of extension, and only a minor amount of thinning (ca. 3 km)
afterward. This pattern of crustal thinning and subsidence can be understood in terms of the
distribution of extension, illustrated by the model strain rates (Fig. 7). Extension in the crust is
broadly distributed across the region encompassing the crust and mantle weaknesses during the
first 10 m.y., but is concentrated most strongly in the area of the crustal weakness (Fig. 7).
Strain in the mantle is more evenly distributed across the pre-weakened region, and is coupled to
deformation in the middle and upper crust by narrow regions of relatively high strain. Offsets in
the initially vertical element boundaries show that the high strain region in the lower crust
behaves as a subhorizontal detachment during this time (Fig. 5). Because extension in the
mantle is initially distributed over a wide region, asthenospheric upwelling is broad and the rate
of lithospheric thinning is relatively slow. As extension progresses, strain in the mantle begins to
focus strongly beneath the mantle weakness. After 10 m.y., this deformation begins to propagate
13
upward, creating two regions of focused extension in the upper crust (the interior rift basin and
marginal basin) that are separated by a region of relatively slow strain forming the structural high
between the two basins. At this time, extension and subsidence is waning in the interior rift
basin and accelerating in the marginal basin. By 15 m.y., extension throughout the lithosphere
begins to focus exclusively in the offshore region (Fig. 7). Crustal thinning and subsidence
ceases in the interior rift basin and on the structural high (Fig. 6). Because extension is more
focused, lithospheric necking accelerates (Figs. 5 and 7). The lithosphere beneath the rift axis
thins to approximately half its initial thickness 20 m.y. after the onset of extension, and proceeds
rapidly to breakup in the next 5 m.y.
Discussion
The thickness of the crust in the model generally agrees with the long wavelength (> 75
km) crustal thickness variations on the Flemish Cap and Galicia Bank margins along the
SCREECH-1 and ISE-1 seismic profiles (Fig. 6a). The duration of the main extensional episode
in the interior rift basin is approximately 13 m.y., in agreement with the late Tithonian through
Valanginian phase of rifting in the Galicia Interior Basin (Murillas et al., 1990; Tucholke et al.,
1989). The model predicts that deep water depths (> ca. 1 km) did not become established in the
offshore rift basin seaward of Galicia Bank until ca. 12 m.y. after the onset of extension.
Breakup is predicted in early Aptian time (ca. 118 Ma), 25 m.y. after the onset of extension, in
agreement with the estimated onset of seafloor spreading at around anomaly M0 time (Srivastava
et al., 2000; Hopper et al., 2004; Funck et al., 2003; Minshull et al., 2001).
The magmatic consequences of rifting in the model are assessed in terms of the timing,
volume, and spatial distribution of melt produced by adiabatic decompression melting in the
upwelling asthenosphere (McKenzie and Bickle, 1988). We neglect conductive heat loss from
14
the rising mantle, which may be significant in slow spreading systems such as the
Newfoundland-Iberia rift (Bown and White, 1995; Pederson and Ro, 1992; Whitmarsh et al.,
2001a; Whitmarsh et al., 2001b), so our melt calculations should be considered to be upper
bounds. Key parameters in the melt calculation are the shape of the ascending asthenospheric
diaper, which is dictated by the pattern of lithospheric thinning in the model, and the mantle
potential temperature, which is a variable in the model. A series of models were developed using
mantle potential temperatures ranging from 1250 !C to 1300 !C as a basal boundary condition
(Fig. 4). Other parameters in these models were identical to those in Tables 1 and 2. Each of
these models produced patterns of extension and subsidence very similar to that shown in Figs.
5-7. The only significant differences were the predicted melt production history. Maximum
melt thicknesses range from 3.1 to 6.5 km at the rift axis at the time of breakup (Fig. 8). The
models using a 1250 !C and 1275 !C mantle potential temperature produce a melt thickness of
3.1 km and 4.7 km at the time of breakup, respectively. This is in good agreement with the 3-4
km thick oceanic crust adjacent to the Newfoundland and Galicia Bank margins (Hopper et al.,
2004; Whitmarsh et al., 1996), suggesting a mantle potential temperature during rifting
somewhere between these two values. Since syn-rift conductive cooling of the asthenosphere
was neglected, this should be considered a minimum estimate of the asthenospheric potential
temperature prior to and during rifting. It seems unlikely, then, that the mantle potential
temperature was more than ca. 30 !C cooler than the 1280 !C global average calculated by
McKenzie and Bickle (1988). In any case, the presence of normal 7 km-thick oceanic crust
within about 15 km of the oldest oceanic crust on the Galicia Bank margin (Whitmarsh et al.,
1996) requires that mantle potential temperatures reached ca. 1280 !C within about 1-1.5 m.y.
after breakup, assuming a 10-14 mm/yr seafloor spreading half rate (Whitmarsh et al., 2001a;
15
Russell and Whitmarsh, 2003). We deem it most likely that the asthenosphere potential
temperature was close to the global average prior to rifting. Moderate conductive cooling (less
than ca. 30 !C) of the upwelling asthenospheric diaper would be sufficient to account for the
unusually thin oceanic crust at the time of breakup, and the near-normal mantle potential
temperature would explain the rapid transition to normal midocean ridge melt production soon
after seafloor spreading began. This scenario is generally consistent with the spatial and
temporal distribution of magma production observed on the margins. In the 1250 !C and 1275
!C potential temperature models, melt production does not begin until 2.5 m.y. or 5 m.y. before
the onset of seafloor spreading, respectively. No melt is predicted further than 45 km landward
of the rift axis in either model, and melt thicknesses in excess of 2 km is restricted to within 35
km of the rift axis. Minor late-stage syn-rift magmatic episodes have been documented further
south on the Iberian margin (Whitmarsh et al., 2001b; Manatschal and Bernoulli, 1999) and
Newfoundland margin (Tucholke et al., 2004), and similar evidence in the form of small MORB-
like volcanic edifices and gabbroic intrusions emplaced on and within exhumed subcontinental
mantle has been reported on the Adrian fossil nonvolcanic rifted margin exposed in the Swiss
Alps (Manatschal and Bernoulli, 1999; Muntener et al., 2000; Muntener and Piccardo, 2003).
We surmise that a minor amount of late stage magmatism, within the range of volumes predicted
by the models here, is likely to be typical of nonvolcanic rifted margins such as Newfoundland-
Iberia.
Summary
Finite element models simulating rifting between Galicia Bank and Flemish Cap invoke
pre-existing weaknesses in the lithosphere to account for the shift in the locus of extension from
16
the Galicia Interior Basin during the first half of the rifting episode to the deep offshore rift basin
during the second half of the rifting episode. Two forms of weaknesses were examined: an
upper mantle weakness, created by excess crustal thickening over a 120 km wide region that is
attributed to the central part of the pre-rift Variscan orogen, and a crustal weakness created by
reducing the yield strength of the upper and middle crust in an 80 km wide region that is
attributed to preexisting structural fabrics in the eastern Variscan front. The model results were
found to be robust over excess crust thicknesses ranging from 2-5 km and upper and middle crust
weakening ranging from 5 to 30%. Under these circumstances, all models produced an early
interior rift basin followed by a shift in extensional deformation to a deep offshore basin where
continental breakup ultimately occurs.
Details of the width and depth of the basins, the time at which deformation shifts from
the interior rift basin to the offshore basin, and the pattern of crustal thickness at the time of
breakup are determined by the relative positions, shape, and magnitudes of the weaknesses
imposed in the model. Model parameters in Tables 1 and 2 produced the general features of the
last 25 m.y. of rifting on the Flemish Cap and Galicia Bank continental margins, including 13
m.y. of extension in the Galicia Interior Basin, 12 m.y. of extension in the Galicia Bank and
Newfoundland Cap marginal basins, and formation of moderately extended crust on Galicia
Bank. The model geometry at the time of breakup approximates structural features of greater
than about 75 km wavelength on these margins.
The nonvolcanic nature of rifting, ultimate production of thin (2-4 km thick) oceanic
crust, and rapid transition to generation of normal thickness (7 km) oceanic crust within 1-1.5
m.y. after breakup requires a mantle potential temperature of 1250 !C to 1275 !C, roughly 30 !C
cooler than the global average at midocean ridges. This modest amount of cooling is attributed
17
to conductive cooling of the ascending asthenosphere during the late stages of rifting. At these
mantle temperatures, the model predicts production of 3.1-4.5 km new oceanic crust at the time
of breakup. Syn-rift magmatism is limited to the last < 5 m.y. of extension and to within 45 km
of the locus of continental breakup. Melt thicknesses of greater than 2 km are restricted to within
35 km of the locus of continental breakup.
Acknowledgements
This research was supported by National Science Foundation grant OPP 0408475. We
thank the organizers of the InterMargins Modeling of Extensional Deformation of the
Lithosphere workshop held in Pontresina, Switzerland July, 2004 for providing a venue for
discussions that led to this modeling study, and the many participants of the workshop who
provided data used to constrain the models. Shelly Grandell received support from the Ronald E.
McNair Postbaccalaurate Achievement Program for this research.
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Figure Captions
Figure 1. Location Map. A) Newfoundland margin. B) Iberia margin. Solid lines are
locations of model transects. Dots are DSDP and ODP drillsites. Shaded regions are interior rift
basins: GIB, Galicia Interior Basin; LB, Lusitania Basin; FCG, Flemish Cap Graben; FPB,
Flemish Pass Basin; JDB, Jeanne d’arc Basin; WB, Whale Basin; SWD, South Whale Basin; HS,
Horse Shoe Basin; OB, Orphan Basin. Contour interval is 1000 m.
Figure 2. Cross-section of the Newfoundland-Iberia rift at the time of anomaly M0 along
the trend of seismic transects SCREECH-1 on the Newfoundland margin (Funck et al., 2003;
Hopper et al., 2004) and ISE-1 on the Galicia Bank margin (Henning et al., 2004). After Funck
et al. (2003).
Figure 3. Yield strength envelopes for lithosphere based on rheological and thermal
properties in Table 1 calculated at a strain rate of 10-15 s-1. A) Lithosphere with nominal 32 km-
thick crust. B) Lithosphere with 35 km-thick crust. C) Lithosphere with weakened upper crust.
D) Steady state geotherm. See text for discussion.
Figure 4. Finite element model construction. Numbers at top and sides indicate constant
temperature and constant extension rate boundary conditions.
Figure 5. Finite element mesh for model that best fits Galicia Bank/Flemish Cap rifting.
Figure 6. A) Crustal thickness variation in the finite element model. Bold line indicates
observed thickness in cross section of Fig. 2. B) Elevation predicted by the finite element model
(does not include syn- and post-rift sediment loading).
Figure 7. Second invariant of the strain rate in the finite element model. Strain rates
range from 10-15 s-1 (dark blue) to 10-10 s-1 (red).
Figure 8. Melt production predicted by models using mantle potential temperatures of
1300 !C, 1275 !C, and 1250 !C.
Table 1. Fixed model parameters.1 Constraint Pre-Rift
Model Value References
Continental Crust Thickness 32-35 km Perez-Gusenye et al. (2003) Oceanic Crust Thickness 3-4 km Tucholke et al. (2004) Lithosphere Thickness Fernandez et al. (1998) Surface Heat Production, A0 3.1 !Wm-2 Fernandez et al. (1998) Heat Decay Exponent, D 12 km Fernandez et al. (1998) Thermal Conductivity, K (crust) 2.5 W m-1 K-1 Tejero and Ruiz (2002) Thermal Conductivity, K (mantle) 3.4 W m-1 K-1 Tejero and Ruiz (2002) Specific Heat, Cp (crust) 875 J kg-1 K-1 Tejero and Ruiz (2002) Specific Heat, Cp (mantle) 1250 J kg-1 K-1 Tejero and Ruiz (2002) "hermal Expansion Coefficient, # (crust) 3.1x10-5 K-1 Tejero and Ruiz (2002) Thermal Expansion Coefficient$# (mantle) 3.1x10-5 K-1 Tejero and Ruiz (2002) Extension Rate 20 mm/yr Russell and Whitmarsh (2003) Brittle Yield Strength, S0 60 MPa Bowling and Harry (2001) Slope of Brittle Failure Strenght, %$ 5 MPa/km Bowling and Harry (2001) Ductile Creep Coefficient, A (crust) 5x10-18 Pa-n s-1 Bowling and Harry (2001) Ductile Activation Energy, Qc (crust) 219 kJ mol-1 Bowling and Harry (2001) Ductile Creep Exponent, n (crust) 2.4 Bowling and Harry (2001) Ductile Creep Coefficient, A (mantle) 4x10-25 Pa-n s-1 Bowling and Harry (2001) Ductile Activation Energy, Qc (mantle) 498 kJ mol-1 Bowling and Harry (2001) Ductile Creep Exponent, n (mantle) 4.5 Bowling and Harry (2001) 1Symbols are defined in text.
Table 2. Variable model parameters. Parameters Best Fit Model Value
Mantle Weakness Width 120 km Mantle Weakness Amplitude 3 km Crustal Weakness Width 80 km Crustal Weakness Thickness (edges) 12 km Crustal Weakness Thickness (center) 26 km Strength Reduction in Crustal Weakness 25% Distance between Center of Weaknesses 110 km Temperature at Bottom of Lithosphere 1250 !C, 1275 !C, 1300 !C
FlemishCap
FPB
JDB
Newfoundland
WB
FCG
CB
SWB
HB
OB
54 W 52 W 50 W 48 W 46 W 44 W 42 W
42 N
44 N
46 N
48 N
50 N
-4000
-4000-2000
-2000
16 W 14 W 12 W 10 W 8 W 6 W
38 N
40 N
42 N
44 N
-400
0
-4000-4000
-400
0
-2000
-200
0
LB
GIB
a)
b)
Fig. 1 - Harry & Grandell
0
10
20
30
0-100 100 200 300Distance from rift axis (km)
De
pth
(km
)
West East
FlemishCap Iberia
GaliciaBank
Galicia InteriorBasin
RiftAxis
Mantle
SerpentinizedMantle
Pre-rift crust Syn- and Post-rift
Fig. 2 - Harry and Grandell
Fig. 3. - Harry and Grandell
0 500 1000 1500Temperature ( C)
0 100 200 300Stress (MPa)
Crustalweakness
Net Strength6.2 x 10 N-m12
a) b) c) d)
0 100 200 300Stress (MPa)
35 km thickcrust
Net Strength6.0 x 10 N-m12
0 100 200 300Stress (MPa)
0
20
40
60
80
100
120
Dep
th(k
m)
32 km thickcrust
Net Strength7.0 x 10 N-m12
T = 0°C
T = Tm
Isostatic Bouyancy Pressure
Ux Ux
mantle
crust
Mantle weakness
Crustal weakness
Fig. 4 - Harry and Grandell
V.E. 5:1
0 m.y.
5 m.y.
10 m.y.
15 m.y.
20 m.y.
25 m.y.
0
32 km
122 km
Weak upper crust
Variscan orogen IberiaNoviaScotia
Crust
Mantle
500 km
Fig. 5 - Harry and Grandell
0510152025303540
-200 -100 0 100 200 300 400
Distance from rift axis (km)
Cru
st
Th
ickn
ess
(km
)
-1
0
1
2
3
4
5
-200 -100 0 100 200 300 400
Ele
vati
on
(km
)Fig. 6 - Harry and Grandell
05
10
15
20
25 m.y.
0 m.y.
510
15
20
25
a)
b)
V.E. 5:1
0 m.y.
5 m.y.
10 m.y.
15 m.y.
20 m.y.
25 m.y.
0
32 km
122 km
Weak upper crust
Variscan orogen IberiaNoviaScotia
500 km
Fig. 7 - Harry and Grandell
Fig. 8 - Harry and Grandell
0
2
4
6
8
Distance From Rift Axis (km)
Melt
Th
ickn
ess
(km
)
-200 -100 0 100 200 300 400
1300 °C, 25 m.y.
1275 °C, 25 m.y.
1250 °C, 25 m.y.
1300 °C, 20 m.y.
