Geophys. J. Int. (2003) 154, 706–730
Magmatism at the west Iberia non-volcanic rifted continental
margin: evidence from analyses of magnetic anomalies
S. M. Russell∗ and R. B. Whitmarsh
Southampton Oceanography Centre, European Way, Southampton SO14 3ZH, UK. E-mail: [email protected]
Accepted 2003 March 12. Received 2003 January 8; in original form 2002 January 24
SUMMARY
We discuss the magmatic development of the west Iberia non-volcanic rifted continental margin
in the North Atlantic Ocean. So-called ‘non-volcanic’ rifted continental margins are characterized by a lack of syn-rift magmatism and are considered to be largely amagmatic. However,
this is clearly an oversimplification since seafloor spreading itself is a magmatic process and
it is implausible that seafloor spreading begins instantaneously. We concentrate our attention
on the recently described zone of exhumed continental mantle (ZECM) to investigate what
magmatic processes accompanied the breakup of the continental lithosphere and the subsequent formation of the ZECM leading to the onset of seafloor spreading. We use magnetic
anomalies supplemented by the interpretations of multichannel seismic reflection profiles and
wide-angle seismic experiments presented elsewhere. Forward and inverse modelling of a seasurface magnetic anomaly chart and of surface and deep-towed magnetometer profiles shows
that anomalies within the ZECM differ in trend, amplitude and source type from those in the
adjacent oceanic crust and thinned continental crust. The ZECM anomalies appear to be caused
by elongated source bodies within 8 km of the top of the acoustic basement aligned parallel
to the margin. We interpret such bodies as syn-extensional intrusions that increase in volume
oceanward. They eventually merge in the vicinity of a margin-parallel, basement peridotite
ridge to give rise to a continuous crust that records reversals in the Earth’s magnetic field from
the time of anomaly M4(N)–M5(R), i.e. to mark the onset of seafloor spreading. We find no
evidence for anomalies formed by seafloor spreading, at either slow or ultraslow rates, before
M5(R) (128 Ma).
Key words: magmatism, magnetic anomalies, rifted margin, west Iberia.
I N T RO D U C T I O N
The rifting of continental lithosphere to form a new ocean basin is
a fundamental Earth process. Yet the geological processes and the
sequence of events involved in the extension, and eventual breakup,
of continental crust and the subsequent onset of seafloor spreading
are only beginning to be understood (Whitmarsh et al. 2001a). This
is partly because the processes are complex; they involve the 4-D interplay of factors such as the rheology of the underlying lithosphere,
the temperature of the asthenosphere and pre-existing continental
heterogeneities. Rheology, strain rate (duration of extension) and
asthenospheric temperature appear to be among the most important factors. Direct observations (e.g. White et al. 1987; Korenaga
et al. 2000; Holbrook et al. 2001; Whitmarsh & Wallace 2001) and
numerical models (e.g. Bown & White 1995a,b; Minshull et al.
2001) suggest that the latter two factors determine the quantity and
∗ Now at: TotalFinaElf SA, 2 Place de la Coupole, La Defense 6, Paris 92078,
France.
706
petrology of syn-rift melt that invades, or underplates, the crust.
Rheology also strongly influences the tectonic evolution of rifted
margins, which involves both extension of continental lithosphere
normal to the margin (e.g. Bassi 1995; Bowling & Harry 2001), and
the propagation of a rift into the continental lithosphere subparallel
to the margin (e.g. Courtillot 1982; Morgan & Parmentier 1985).
Here we discuss aspects of the magmatic development of a nonvolcanic rifted margin in the North Atlantic Ocean. So-called ‘nonvolcanic’ rifted continental margins are characterized by a paucity
of evidence of syn-rift magmatism (seaward-dipping reflector sequences are absent and there is no indication offshore of underplating or onshore of significant magmatic activity). However, to
consider such margins as amagmatic is clearly an oversimplification
because, even if the extended continental lithosphere exhibits no evidence of syn-rift magmatism, seafloor spreading itself is obviously
a magmatic process. Thus, a key region in this discussion is the recently recognized zone of exhumed subcontinental mantle (ZECM;
formerly called the ocean–continent transition zone) off west Iberia
(Boillot et al. 1989a; Whitmarsh et al. 2001a). Therefore, in seeking
to answer the questions ‘what magmatic processes accompany the
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Magmatism at non-volcanic rifted margin
breakup of continental lithosphere and the subsequent mantle exhumation leading up to the onset of steady-state seafloor spreading?’
and ‘how are these processes expressed physically in the transition,
in space and time, between the faulted and extended continental
crust and the oceanic crust produced by seafloor spreading?’ we
shall concentrate our attention on the ZECM.
Our principal tool will be the study of magnetic anomalies supplemented by the interpretations of multichannel seismic (MCS)
reflection profiles and wide-angle seismic experiments presented
elsewhere (Whitmarsh et al. 1990; Pickup et al. 1996; Chian et al.
1999; Dean et al. 2000). In the context of a rifted continental margin,
three principal rock types can generate the magnetization contrasts
that give rise to magnetic anomalies. First, continental crust generally consists of acidic, metamorphic and meta-igneous rocks of
varying metamorphic grade. Above the Curie isotherm magnetite
is the main magnetic mineral throughout the crustal section, although it generally possesses negligible natural remanent magnetization (NRM) and has low susceptibilities (Wasilewski et al. 1979;
Frost & Shive 1986); the average induced magnetization of the most
magnetic samples is less than 1 A m−1 and susceptibilities of 10−3
SI units are commonly encountered (Shive & Fountain 1988). Secondly, although unaltered uppermost-mantle peridotites are inferred
to have an NRM of ∼1 A m−1 (Toft & Arkani-Hamed 1992) this is
a minimum value because the serpentinization of ultramafic rocks
produces a considerable amount of magnetite, which can acquire
a significant chemical remanent magnetization (CRM), as well as
an induced component of magnetization (Wasilewski et al. 1979;
Nazarova 1994). Thirdly, mafic rocks, which make up the oceanic
crust and many intrusive bodies, possess a significant NRM, which
is much greater than their induced magnetization. Studies of the
magnetism of oceanic rocks have led to many models of a layered
crust all of which possess a common characteristic, namely the principal magnetic contributions are from the pillow basalts of oceanic
layer 2 and the gabbros of layer 3 (e.g. Kent et al. 1978; Dunlop
& Prévot 1982). This general model is valid, even ∼130 Ma after
emplacement, in spite of crustal alteration that increases with age
(Bleil & Petersen 1983).
Single magnetic anomalies that are hundreds of nT in amplitude
and that can be traced for hundreds of kilometres along the outer
continental shelf and slope of rifted margins were recognized early
(Keen 1969). One of the best known is the 35–100 km wide East
Coast Magnetic Anomaly (ECMA) offshore North America that can
be traced for at least 1000 km (Keller et al. 1954). The ECMA, and
possibly other broad anomalies offshore Norway (Nunns 1983), Argentina (Rabinowitz & LaBreque 1979) and SW Africa (Simpson
& du Plessis 1968), appear to be related to thick piles of subaerial,
seaward-dipping reflectors (lava flows) that may represent part of
the first-formed oceanic crust (Talwani et al. 1995); in other words,
they are associated with volcanic margins. There are also offshore
magnetic anomalies that lie subparallel to non-volcanic margins but
have relatively low amplitudes and occur as multiple bands up to
25 km wide. Such anomalies have been studied in the Labrador Sea
(Srivastava & Roest 1995, 1999), off western Iberia (Discovery
215 Working Group 1998; Srivastava et al. 2000; Whitmarsh et al.
2001b) and in the Australian–Antarctic Basin (Koenig & Talwani
1977; Tikku & Cande 1999). Some authors have suggested that
they are the result of seafloor spreading (Keen et al. 1998; Srivastava et al. 1998, 2000; Srivastava & Roest 1999). The purpose of this paper is to investigate such anomalies offshore west
Iberia and to present an alternative geological explanation for these
anomalies.
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2003 RAS, GJI, 154, 706–730
707
T H E W E S T I B E R I A M A RG I N
West Iberia broke away from the Grand Banks of Newfoundland
in the early Cretaceous (e.g. Driscoll et al. 1995). The West Iberia
continental margin is regarded as non-volcanic because onshore
there is no evidence of significant syn-rift volcanism (Pinheiro et al.
1996) and offshore no significant amounts of syn-rift mafic rocks
have been recovered in basement drill cores (Sawyer et al. 1994;
Whitmarsh et al. 1998). Furthermore, extensive MCS reflection profiling has failed to reveal any indication of seaward-dipping reflector sequences (Mauffret & Montadert 1987; Mauffret et al. 1989;
Krawczyk et al. 1996; Pickup et al. 1996) and modelling of wideangle seismic profiles across the margin has not required the typical
high-velocity, lower crustal underplated bodies characteristic of volcanic margins (Pinheiro et al. 1992; Whitmarsh et al. 1996b; Chian
et al. 1999; Dean et al. 2000).
The West Iberia margin is commonly divided, from north to south,
into three physiographic provinces; these are the Galicia Bank margin (41◦ –43◦ N), the southern Iberia Abyssal Plain margin (39◦ –
41◦ N) and the Tagus Abyssal Plain margin (37◦ –39◦ N; Fig. 1). Multichannel and wide-angle seismic profiles across the two northern
provinces suggest that, from east to west, the margin can be divided
into: (1) thinned continental crust characterized, at least on its seaward edge, by fault blocks; (2) the ZECM, in which mantle rocks
are serpentinized to varying degrees; and (3) an acoustic basement
with the seismic and magnetic signature of oceanic crust (Sibuet
et al. 1995; Whitmarsh & Miles 1995; Pickup et al. 1996; Whitmarsh et al. 1996a; Dean et al. 2000). In the Tagus Abyssal Plain
a paucity of data hinders recognition of this threefold classification
(Mauffret et al. 1989; Pinheiro et al. 1992).
The continental fault blocks are recognized by their characteristic
velocity structure and by their association on seismic reflection profiles with seaward-dipping normal faults on their seaward side and,
less commonly, by a cap of tilted pre-rift sediment (Basile 2000).
The blocks are clearly visible on most profiles across the lower continental slope and rise, and can be traced seaward until the crust is
only ∼3 km thick (red dots in Fig. 1). There is a correlation between such fault blocks and the (inferred) continental nature of the
underlying basement wherever a block has been drilled (Whitmarsh
et al. 2000). For the purpose of this discussion the landward edge of
oceanic crust is assumed to be marked by a weakly segmented, but
otherwise continuous, peridotite basement ridge that extends over
300 km parallel to the margin (Beslier et al. 1993) (white triangles in Fig. 1). Seaward of this ridge Mesozoic seafloor spreading
anomalies have been recognized previously that suggest onsets of
seafloor spreading at the time of anomaly M11 ∼ (133 Ma; we use
the timescale of Gradstein et al. 1994, in this paper) and during M3
(∼126 Ma) in the Tagus and southern Iberia Abyssal Plains, respectively (Pinheiro et al. 1992; Whitmarsh & Miles 1995; Whitmarsh
et al. 1996a). Although oceanic crust immediately adjacent to the
Galicia Bank margin was created during the Cretaceous constant
polarity interval, so that the onset of seafloor spreading cannot be
determined there from magnetic anomalies, Boillot et al. (1989b)
recognize a late Aptian (112–117 Ma) ‘breakup’ unconformity in
drill cores. Evidently therefore rifting proceeded from south to north.
The ZECM, which lies between the thinned continental crust and
the oceanic crust, is described in the next section. It has a varying
width and appears to narrow northwards from a maximum of ca.150
km under the southern Iberia Abyssal Plain to ca.15 km off the
Galicia Bank margin (Pickup et al. 1996; Whitmarsh et al. 1996b;
Discovery 215 Working Group 1998; Dean et al. 2000). The ZECM
708
S. M. Russell and R. B. Whitmarsh
44˚
CF
GB
Fig.2
SB
S2
40˚
S1
IBERIA
Fig.3a
34
J
TS
ES
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GRB
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10˚
1100
500
400
300
200
150
100
50
0
-50
-100
-150
-200
-600
Magnetic anomaly (nT)
Figure 1. Magnetic anomalies west of Iberia (Miles et al. 1996) gridded at 5 km and overlain by bathymetric contours at 1000 m intervals. Anomalies 34
and J are labelled. Note the non-linear colour palette. Boxes indicate areas covered by Figs 2 and 3(a); dashed boxes indicate areas S1 and S2 (see text). White
triangles, peridotite ridge; red dots, approximate seaward edge of thinned continental crust; the zone of exhumed continental mantle is the region between the
while triangles and red dots. Black dots represent DSDP/ODP drill sites. CF, Cape Finisterre; GB, Galicia Bank; SB, Swallow Bank; TS, Tore Seamount; ES,
Estremadura Spur; TAP, Tagus Abyssal Plain; CSV, Cape St. Vincent; GRB, Gorringe Bank. The inset shows the location of west Iberia in the North Atlantic.
is not yet clearly delimited in the Tagus Abyssal Plain (Pinheiro
et al. 1992).
THE SOUTHERN IBERIA
ABYSSAL PLAIN
We shall analyse magnetic anomalies in the southern Iberia Abyssal
Plain for evidence of any magmatism that accompanied rifting prior
to the onset of seafloor spreading. This area lies between the southern
margin of Galicia Bank and the Estremadura Spur (Fig. 1).
In the northern part of the area a grid of intersecting wide-angle
seismic lines has been interpreted in terms of regions of thinned continental crust, a ZECM and oceanic crust (Fig. 2; Chian et al. 1999).
The three principal basement ridges east of 11◦ 50 W, on which ODP
sites 1069, 900, 1067, 1068 and 1065 were drilled (Fig. 2), have the
velocity structure of thinned continental crust, which suggests they
represent the southward extension of the three ridges on the south
flank of Vasco da Gama seamount (Mougenot et al. 1986). Cores of
sediment and acoustic basement from these ODP sites also provide
evidence that all of these sites are underlain by continental crust
(Whitmarsh & Wallace 2001). The ridges appear to die out south of
40◦ 35 N. West of 11◦ 50 W the seismic velocity increases rapidly
(4.0–7.3 km s−1 ) in the top 2.0–2.75 km of basement and then
increases gradually to ca.8 km s−1 (Chian et al. 1999). No clear
Moho refraction or reflection is observed here. The structure has
been interpreted as a serpentinized peridotite basement in which
serpentinization decreases as velocity increases with depth. This
ZECM velocity structure persists at least as far west as the peridotite
ridge. ODP sites 899 and 897 both encountered serpentinized peridotite in basement cores, consistent with this interpretation (Fig. 2).
Further west, both the velocity structure, albeit less constrained
(Whitmarsh et al. 1990), and the presence of seafloor spreading
magnetic anomalies (already mentioned) strongly suggest the presence of a geophysically defined oceanic crust. Also, in the northern
part of the southern Iberia Abyssal Plain, MCS reflection profiles
roughly normal to the margin show tilted fault blocks, with high- and
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2003 RAS, GJI, 154, 706–730
Magmatism at non-volcanic rifted margin
41˚
SAR
93
897
1069
1070
TOBI91
A
899
R3
1068
1067
398
D
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898
B
709
901
900
C
R4
SAR95
40˚
-12˚
6000
7000
-10˚
8000
9000
Metres below sea-level
Figure 2. Colour-coded, depth-to-basement chart on a 500 m grid (from P. Cole, Pers. comm. 2000 after Discovery 215 Working Group 1998) and bathymetric
contours from swath bathymetry over Vasco da Gama seamount (contour interval 500 m; J.-C. Sibuet, Pers. comm.) showing regions A–D bounded by dotted
lines (see the text; from Russell 1999). Bold lines denote three SAR and TOBI deep-towed magnetometer tracks discussed in the text; other lines are wide-angle
seismic lines of Dean et al. (2000), Dean et al.’s line IAM-9 is coincident with SAR95, and Chian et al. (1999). Red triangles, peridotite ridge segments R3 and
R4; black triangles and dots, ODP/DSDP drill sites on basement highs. Rhomboidal box bounded by dashed lines represents the area within which 29 regularly
spaced, 100◦ -trending, synthetic basement (and magnetic anomaly) profiles were extracted from gridded data sets (see the text).
low-angle normal faults, which are coincident with the above three
ridges of thinned continental crust and another ridge beneath site
901. The faults merge into subhorizontal detachment faults near the
base of the crust or even within the upper-mantle (Krawczyk et al.
1996; Whitmarsh et al. 2000, and unpublished data). Further west
and south in the ZECM, tilted blocks are absent and the acoustic
basement has less relief until the peridotite ridge is reached.
In the southern part of the area a deep MCS profile and a wideangle seismic velocity model were acquired along the 330 km long
IAM-9 profile (coincident with SAR95 in Fig. 2; Banda et al. 1995).
The seismic velocity model (Dean et al. 2000) clearly reveals the
same threefold division identified further north although the en echelon overlap of the R3 and R4 segments of the peridotite ridge creates
the impression of a broader zone of peridotite ridges on this profile. Continental crust thins westwards to only ∼3 km near 10◦ 45 W
(Pickup et al. 1996). Further west, in the ZECM, the seismic velocity increases rapidly (4.5–7.3 km s−1 ) with depth into the basement
to about 2–3 km below the top of the basement and then increases
more gradually up to 8 km s−1 . No clear Moho refraction or reflection is observed in the ZECM. West of the peridotite ridge R3 the
velocity structure falls within the envelope for normal, 59–170 Ma,
oceanic crust in the North Atlantic (Whitmarsh et al. 2001b).
The MCS profile IAM-9 extends from the edge of the continental
shelf to 65 km west of the peridotite ridge. West of the basement fault
blocks of probable continental crust, which die out around 10◦ 50 W,
the broad 150 km wide ZECM is characterized by low basement relief (except in the region of overlapping peridotite ridges). Here the
unreflective character of the upper 2.5 km of basement has been
associated with extensive serpentinization (Pickup et al. 1996), exactly in accord with Dean et al.’s (2000) interpretation of the velocity
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2003 RAS, GJI, 154, 706–730
model. West of the westmost peridotite ridge R3 the basement relief
increases and a Moho reflection can be distinguished.
The concept of a threefold division of the continental margin
in the southern Iberia Abyssal Plain, which is well documented along
the two separate transects mentioned above, is further supported by
the chart of the basement relief of the acoustic basement (Fig. 2).
The chart was constructed using a relatively dense network of MCS
reflection profiles. A surface was fitted to the depths to the acoustic
basement, computed from bathymetry (including swath bathymetry
around Vasco da Gama Seamount where the basement crops out; J.C. Sibuet, personal communication) and two-way times picked from
MCS reflection profiles, and then gridded (Russell 1999; Whitmarsh
et al. 2001b). After adjustment to minimize crossover differences,
but neglecting errors in the predicted off-track depths inherent in a
surface fitted with a tension of 0.35, the uncertainty in calculated
depths is estimated to be ≤200 m.
The depth-to-basement chart was divided into four regions A–D
of contrasting relief, which will be used as a basis for later discussion in this paper (Fig. 2). In a western region A, the basement
surface consists of elongated ridges and troughs that parallel the
seafloor spreading isochrons further west. Previous modelling of
magnetic anomalies and seismic velocities indicated that this region
has seafloor spreading magnetic anomalies and the seismic velocity
structure characteristic of oceanic crust, albeit perhaps slightly thinner than normal; magnetic modelling suggested that it was formed
from the time of anomaly M3 (124 Ma; Whitmarsh & Miles 1995;
Whitmarsh et al. 1996a; Dean et al. 2000). This interpretation is
supported by the discovery at ODP site 1070 of a 119 Ma pegmatite
gabbro and gabbro veins, which were cored in the basement without encountering basaltic rocks (ODP Leg 173 Shipboard Scientific
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S. M. Russell and R. B. Whitmarsh
Party 1998; Manatschal et al. 2001; Whitmarsh & Wallace 2001).
The gabbro had an enriched mid-ocean ridge basalt (E-MORB)
source. However, the gabbro is associated with a depleted subcontinental serpentinized peridotite. Thus, although near site 1070 one
encounters the geophysical characteristics of normal oceanic crust,
the cores reveal subcontinental mantle intruded by MORB dykes exhumed at the seafloor and covered by sediment shortly afterwards. In
contrast, 170 km further west, tholeiitic pillow basalts were dredged
from the crest of Swallow Bank, a basement ridge standing above
the Iberia Abyssal Plain (Fig. 1; Matthews 1962).
The basement in region B is relatively shallow. Region B is flanked
to the west by the margin-parallel, segmented peridotite basement
ridge. This ridge may be offset near 40◦ 25 N. Peridotite breccia
was also encountered in the basement of region B at ODP site 899
(Sawyer et al. 1994). The peridotite ridge is an enigmatic feature
which, west of Galicia Bank, corresponds to an abrupt boundary
across which basement magnetization decreases by an order of magnitude from west to east (Sibuet et al. 1995). Region B, described
here as the region of offset peridotite ridges, is also associated with
post-rift sediments that were folded into two NNE-trending monoclines during Middle Miocene compression (L. North, personal
communication; Masson et al. 1994).
Seismic observations indicate that the region C basement is also
largely serpentinized mantle. However, in contrast to region B, it is
deeper, has yet to be sampled and has much less relief (Pickup et al.
1996; Discovery 215 Working Group 1998; Dean et al. 2000). Thus,
regions B and C correspond to the ZECM, although the basement
chart indicates there are differences in relief between the two. Region
D contains the southern flank of Galicia Bank; seismic observations,
dredging and drilling strongly suggest that it consists of thinned
continental crust (Capdevila & Mougenot 1988; Chian et al. 1999;
Whitmarsh & Wallace 2001).
Although seafloor spreading magnetic anomalies occur west of
the peridotite ridges other weaker magnetic anomalies, that parallel the margin, are also evident within the ZECM and its probably
southward extension into the Tagus Abyssal Plain (Fig. 1). Previously, these anomalies were only forward modelled (Pinheiro et al.
1992; Whitmarsh & Miles 1995; Srivastava et al. 2000). Here we
apply a suite of inverse methods, which are inherently less subjective than forward modelling, to magnetic anomaly charts and
near-bottom and surface magnetometer profiles and examine the results in the light of the magmatism that may have accompanied the
transition from continental rifting to seafloor spreading.
CHARACTERISTICS OF
M AG N E T I C S O U RC E S
Trends of sources
In the southern Iberia Abyssal Plain, we assume that rifting generated principally linear, margin-parallel tectonic features and synrift intrusions that are potential magnetic sources. Here a relatively
high-quality magnetic anomaly chart is available (Fig. 3a), which
was reduced to the pole (Schouten & McCamy 1972, Figs 3b–e),
thereby shifting anomalies close to their sources. A variety of NRM
vectors, which cover the likely range of magnetizations of the basement rocks, were used in the reduction. These range from the current
field direction I = 57◦ , D = −9◦ (for continental rocks with induced
magnetization) to I = 46◦ , D = −43◦ (for rocks that acquired a remanent magnetization before the anticlockwise rotation of Iberia
during the opening of the Bay of Biscay). Most of this rotation had
ended around M0 time (121 Ma; Galdeano et al. 1989; Van der Voo
1990; Juarez et al. 1998), although some seafloor spreading continued in the Bay of Biscay until ∼80 Ma (Sibuet & Collette 1991).
The chart was reduced to the pole in steps of 15◦ of declination, the
minimum step that produces a noticeable difference in the modified contours. It is clear from Fig. 3 that the greatest linearity and
‘sharpness’ of the contours correspond to a basement magnetization
vector of I = 46◦ and D = 0◦ or −15◦ . Assuming the anomalies
are caused by margin-parallel syn-rift intrusions, this suggests that
the magnetization of the basement rocks within the area of Fig. 3(a)
was acquired either at or after the time of M0 or is induced; it is
unlikely that a significant remanent magnetization, if present, was
acquired much before M0 time. Fig. 3(b) reveals that, west of the
peridotite ridge (X < 100 km), J and its flanking anomalies have a
strong ∼015◦ trend. Further east (100–200 km) in the ZECM the
trend is slightly more northerly (∼010◦ ). Even further east the trend
becomes ∼350◦ .
We computed directional power spectra (Bhattacharyya 1966)
over different regions of the chart of observed magnetic anomalies.
For such computations a relatively large area is required in order
to capture the longer-wavelength anomalies. Over continental crust
(Fig. 2, region D), within the wavenumber bandwidth representative
of the basement, the energy from N–S and E–W anomalies is of the
same order of magnitude, implying that either any linear anomalies
have no preferred direction and/or that the anomalies are better
represented by point, or vertically aligned, sources. Over the ZECM
(regions B and C) the E–W-trending anomalies have about 50 times
less energy than the N–S-trending anomalies; this emphasizes the
importance of the margin-parallel (∼010◦ ) trend of the anomalies
in this zone.
Magnetic contribution of the basement surface
Any modelling of magnetic anomalies needs to take account of the
relief of the basement surface where potentially there is a strong
magnetization contrast with the overlying, very weakly magnetized,
post-rift sediments (Zhao et al. 2001). Therefore, before determining the depth of magnetic sources within the acoustic basement it
was useful to ascertain to what extent relief of ∼2–3 km on the upper
surface of the acoustic basement (Fig. 2) contributes to the observed
magnetic anomalies. Preliminary calculations indicated that a uniformly magnetized (0.5–1.5 A m−1 ) basement, with relief typical
of that under the southern Iberia Abyssal Plain, could produce seasurface anomalies of ∼100 nT comparable to the largest anomalies
observed.
Two techniques were used to estimate the contribution of basement relief. The first, and more approximate, used 3-D Euler deconvolution solutions (Reid et al. 1990). We assumed a structural
index (SI) of 0.5 appropriate to finite 2-D step-like features (steep
slopes) on the basement surface. Using progressively overlapping
10 km windows, in which a solution is estimated at the centre
of each window, we found that only 10 per cent of all the calculated solutions were judged acceptable. Solutions were spaced by
1 km. This result is based on selecting only those solutions lying
below the basement surface with calculated uncertainties being less
than 1 per cent of their calculated depths (Fig. 4); we are able to
compute an uncertainty on each solution because of the overdetermined nature of the inverse problem. We justify the use of a
10 km window based on the width of anomalies and the anticipated
depth extent of sources through forward modelling of anticipated
geometries. We found for geologically plausible cases that these
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2003 RAS, GJI, 154, 706–730
Magmatism at non-volcanic rifted margin
711
0
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(b)
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150
500
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100
100
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(d)
D=-30˚
150
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(e)
D=-43˚
150
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250
0
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0.01◦ )
Figure 3. (a) Observed magnetic anomalies (contour interval 25 nT, grid interval
in the southern Iberia Abyssal Plain (for location see the box in
Fig. 1). Dashed red lines denote borders of regions A–D in Fig. 2. Anomaly J is the high-amplitude anomaly at ca.13◦ 10 W. (b) magnetic anomalies in larger
white box in (a) reduced to the pole with parameters I M = 46◦ , DM = 0◦ , I G = 57◦ and DG = −9◦ (I = inclination, D = declination, subscript M = rock
magnetization, subscript G = geomagnetic field). Black lines are trends of principal linear anomalies estimated by eye; (c)–(e) As (b) except DM = −15◦ , DM
= −30◦ and DM = −43◦ , respectively.
procedures are valid, although we recognize a trade-off exists in
that some valid solutions may be eliminated at the expense of acquiring a coherent distribution of clustered solutions. Thus we have
confidence in our solutions as many of those with underestimated
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2003 RAS, GJI, 154, 706–730
or overestimated depths, including those towards the tails of the
typical concave-down arc-like patterns associated with Euler solutions, have been removed by the harsh tolerance filter of 1 per cent.
The depth distribution of solutions suggests that many depths were
S. M. Russell and R. B. Whitmarsh
8
41N
6
4
Euler solutions
(circles) depth
below top
basement (km)
712
2
1
c
40N
13W
12W
0
11W
Figure 4. Accepted Euler deconvolution solutions (coloured circles) for SI = 0.5 presented as depth below the top of the basement for the area enclosed by
the smaller box in Fig. 3(a). Window size 10 km × 10 km and 1 per cent acceptance level. Solutions are superimposed on the gradient of basement relief
illuminated from the west. About 25 per cent of the solutions lie within 0–1 km of the top of the basement are assumed to be associated with basement relief
(black arrows). Other symbols as in Fig. 2.
underestimated but, nevertheless, those that lie below the top of the
basement do follow the trend of some steep basement surfaces. It
is estimated qualitatively that ∼25 per cent of the acceptable solutions coincide with significant basement relief. Thus, the remaining
acceptable solutions probably represent sources lying within the
basement.
The second technique relied on forward modelling anomalies
along 29 parallel synthetic profiles spaced 2.5 km apart and oriented
at 095◦ across the southern Iberia Abyssal Plain between 40.0◦ –
40.9◦ N and 11◦ –13◦ W (rhomboidal box in Fig. 2). The profiles were
chosen to be approximately normal to the principal basement trends
and margin-parallel magnetic anomalies. Magnetic anomaly and
basement profiles were extracted along coincident profiles across
the sea-surface anomaly chart and the depth-to-basement chart, respectively. 2-D models were set up with a horizontal base at 14 km
depth and with either remanent (±1 A m−1 ) magnetizations, according to the vectors of Galdeano et al. (1989) and Van der Voo (1990),
or with solely an induced magnetization.
Two independent statistical comparisons of observed and forward
computed anomaly profiles were made that are largely independent
of the assumed intensity of magnetization of the models. First, in
the space domain, linear correlation coefficients and residuals were
computed (Fig. 5). All but one of the correlations were positive but
most were <0.2 and therefore not considered statistically significant. Only three profiles exhibited marginally significant correlations. The root mean square residuals vary systematically between
profiles (Fig. 5a). Smaller (15–30 nT) residuals that exist south of
profile 15 are better explained by basement topography than those
larger values (40–60 nT) to the north, which possibly reflect the
effect of higher relief, but weakly magnetized, continental crust on
profiles 15–29. The residuals along each profile were normalized
by 15 nT (the smallest rms residual), contoured and compared with
the above four regions A–D (Fig. 5b; Table 1). The results suggest
that sources other than basement relief contribute to the observed
magnetic anomalies in regions A, B and D. The result in region C,
where the basement is on average ∼2 km deeper and anyway has
less relief, is ambiguous.
The second statistical comparison tested the correlation of magnetic anomalies with basement features as a function of wavenumber
(and hence wavelength). The squared coherence (γ 2 ot ), between the
observed and calculated anomalies, was plotted against wavenumber for each profile and contoured (Fig. 6). For wavenumbers in the
band >5 km there is no significant correlation (i.e. γ 2ot < 0.5), at
any wavenumber along any profile, and typically squared coherence
values are <0.2. The exception is an area where γ 2ot > 0.4 in the band
equivalent to wavelengths of 12.5–25 km (profiles 11–14), in which
a significant coherence (i.e. γ 2ot > 0.5) occurs at a wavelength of ∼15
km (profile 12). This area corresponds to the three N–S basement
ridges in regions A and B, which are spaced ∼22 km apart.
Therefore, based on the assumption of a uniformly magnetized
basement, we conclude that the relief of the basement surface alone
plays a minor role in determining the amplitude and location of
the anomalies. Hence we attribute the observed magnetic anomalies
in the southern Iberia Abyssal Plain largely to magnetic contrasts
within the acoustic basement. Nevertheless, we recognize the importance of including basement relief in magnetic models as such relief
may be important when its effect is coupled with that of geological
variation.
Intensity of magnetization of the source rocks
The intensity of magnetization of the source rocks can be estimated
from the inversion of magnetic anomaly charts and profiles. Lateral magnetization variations, assuming a constant-thickness source
layer, were determined by the inversion of the observed sea-surface
magnetic anomaly chart in area S2 (Fig. 1) using the 3-D method
of Parker & Huestis (1974). We ran two inversions with a source
layer either 4 or 7 km thick, comparable to the thickness of thin or
‘normal’ oceanic crust, respectively, which possessed the relief of
the basement surface. The data were first filtered to exclude wavelengths outside a 5–200 km band. The rejection of variations <5 km
avoided false oscillations inherent in the inverse Fourier method, but
it also had the consequence that if seafloor spreading had occurred
C
2003 RAS, GJI, 154, 706–730
Magmatism at non-volcanic rifted margin
-1.0 -0.8 -0.6 -0.4 -0.2 0.0 0.2 0.4 0.6 0.8
normalized residual (scale bar)
correlation
-0.5
0.0
0.5
1.0
-1.0
713
1.0
30
profile number
20
20
rms
residua
15
15
correlation
10
5
10
a
b
line number
25
25
5
0
0
20
40
60
rms residual (nT)
80 0 20 40 60 80 100 120 140 160
distance (km)
Figure 5. (a) Correlation coefficients (thin lines) and root mean square residuals (thick lines) between the observed magnetic field and a forward calculated
magnetic field along each of 29 profiles extracted from the area of the box bounded by dashed lines in Fig. 2. Each forward calculated field was computed
from a model for which the upper surface was extracted from the depth-to-basement chart (Fig. 2) along the coincident line. Each model was described by a
lower surface at 14 km depth and a magnetization vector from Van der Voo (1990); (I M = 46◦ , DM = 0◦ ; solid line) or Galdeano et al. (1989)); I M = 58◦ ,
DM = −43◦ ; dashed line) and an induced magnetization vector (I G = 57◦ , DG = −9◦ ). (b) Residuals (as a function of distance along profile) normalized by
the minimum rms residual of profile 6 (15 nT) and contoured (interval 0.5) over the region of profiles 1–29. Varying the assumed magnetization vector does
not alter the main results. The figure shows results using the vector of Van der Voo (1990).
Table 1. Proportions, by area, of each
of the four regions (A–D) where the contoured normalized residual exceeds 1 and
0.5.
Region
A
B
C
D
>1.0
(per cent area)
>0.5
(per cent area)
50
25
0
20
90
75
15
60
at half-rates <10 mm yr−1 some crustal blocks narrower than 5 km,
if present, would not have been resolved.
The resulting magnetization contrasts for the 4 km thick crust
(Fig. 7) are different either side of a NNE boundary that parallels,
but lies ∼15 km east of, the peridotite ridge segment R3; it appears
that this boundary is not offset close to 40◦ 26 N where the peridotite
ridge is offset eastward. West of the boundary the magnetization,
regardless of sign, is commonly >0.5 A m−1 and often exceeds
1.5 A m−1 ; the contoured magnetization shows a strong NNE–SSW
trend. The normally and reversely magnetized blocks can plausibly
be explained by a sequence of seafloor spreading anomalies from
M4(N), even from M5(R) in the south, to M2(R) (Fig. 7). The region
east of the boundary is characterized by a uniformly and weakly
magnetized layer (<0.5 A m−1 ) with the exception of four isolated
highs, two of which (H2 and H4) are artefacts related to the locally
poorly constrained basement surface. In the central ZECM (11◦ 20 –
12◦ 10 W) linear magnetization contrasts are relatively weak but also
C
2003 RAS, GJI, 154, 706–730
show an east of north to N–S trend. Further east the contrasts trend
consistently west of north.
High H1, which has a maximum magnetization >1.5 A m−1 in
Fig. 7, includes the location of ODP hole 899B where magnetization, regardless of sign, of serpentinized peridotite core samples
ranges from 0.29 to 2.46 A m−1 (average ∼1.50 A m−1 ), the volume
magnetic susceptibility is ∼3 × 10−2 SI units and the mean Koenigsberger ratio is 2.1 (Zhao et al. 2001). Several core samples retained
nearly half their NRM intensity after demagnetization to 400◦ C,
suggesting that the remanence of the serpentinized peridotite body
under site 899 can contribute significantly to the magnetic anomaly.
Because of the good agreement between the NRM of drill cores
and the modelled magnetization we propose that the magnetic layer
thickness is ∼4 km at site 899.
Near-bottom observations can provide higher-resolution constraints on the sources of magnetic anomalies than surface observations, especially in water depths of 5 km or more, as encountered
in the southern Iberia Abyssal Plain. Thus, three profiles were acquired using three-component fluxgate magnetometers mounted on
the TOBI and SAR systems of IOS and Ifremer, respectively (Jarry
1986; Flewellen et al. 1993) towed approximately 200 m above the
seabed at an almost constant depth of 5 km below the sea surface.
Therefore, the SAR93 and TOBI91 profiles and the SAR95 profile were acquired at mean heights above the basement of ∼2.5 and
∼4.0 km, respectively.
Inversion of these near-bottom total field magnetometer profiles
by the same technique as above, but in two dimensions (Parker &
Huestis 1974), and with the same filter parameters, yields a similar
pattern of higher magnetization amplitudes along profile TOBI91
714
S. M. Russell and R. B. Whitmarsh
0.50
0.40
0.30
0.20
squared coherency
0.10
0.00
0.2
20
15
0.2
profile number
25
0.4
10
5
0.050
0.075
0.100
0.125
0.150
0.175
0.200
Wavenumber (km-1)
Figure 6. Squared coherence between observed and calculated magnetic anomalies along the 29 profiles contoured (interval 0.1) in profile number/wavenumber
space. Forward calculated anomalies from a model for which the magnetization vector is defined by Van der Voo (1990; I M = 46◦ , DM = 0◦ ) and geometry as
in Fig. 5. No significant correlation exists (all values <0.5) between observed magnetic fields and variations in basement topography at any wavelength on any
line, with the exception of 15 km wavelength on line 12 (values >0.5).
west of about 12◦ 10 W (30 km east of the peridotite ridge segment R3; Fig. 8a). Peak-to-trough amplitudes are 3.0–3.5 A m−1
and wavelengths are typically ∼25 km. East of 12◦ 10 W solutions
are characterized by peak-to-trough amplitudes of ≤1 A m−1 but
similar wavelengths. The same twofold division into areas of strong
and weak magnetization is exhibited by the inversions of SAR93
(west of 12◦ 10 W; 90 km) and SAR95 (east of 12◦ 10 W; 15 km).
In addition SAR95 exhibits a pair of magnetization peaks of ∼1.5
A m−1 (peak-to-trough) between 70 and 120 km. When shorterwavelength (2.5 km) magnetization variations are included in this
inversion a better fit between observed and forward calculated magnetic anomalies is achieved (Fig. 8b). On profile TOBI91 most of
the additional high wavenumber energy occurs east of 12◦ 10 W. On
profile SAR93 additional high wavenumber energy is distributed
along the entire profile, whereas on profile SAR95 little change is
apparent. This may (1) suggest that lateral variations in magnetic
sources beneath profile SAR95 occur on length-scales of ≥5 km,
whereas such variations beneath profiles TOBI91 and SAR93 occur on length-scales of ∼2.5 km or (2) reflect the greater distance
above the magnetic basement of SAR95 compared with TOBI91
and SAR93 or (3) imply that sources are more deeply buried be-
neath SAR95 or (4) indicate that some combination of these factors
is present.
Depths of sources from an analysis
of the sea-surface anomaly chart
Depths of magnetic sources were estimated by applying two techniques (inversion of power spectra and 3-D Euler deconvolution)
to the sea-surface magnetic anomaly chart. Because the top of the
basement typically lies 5–9 km below sea level (b.s.l.) it is necessary
to analyse anomalies over relatively long profiles or over relatively
large areas to adequately represent the broadest possible bandwidth
of interest.
Euler solutions were computed for a range of structural indices,
from 0 to 3 in steps of 0.5, and the distributions of solution depths
were compared with the depth ranges of the underlying acoustic
basement in regions A–D. We used a 5 km × 5 km window. An
Euler solution was accepted if its calculated depth uncertainty did
not exceed 1 per cent of its depth. As mentioned before, such a
rigorous acceptance criterion minimizes the effects of under- and
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2003 RAS, GJI, 154, 706–730
Magmatism at non-volcanic rifted margin
13W
12W
715
11W
41N
1
-0.5
0
H4
0
0.5
0
0.5
M3
(R
0.5
)
M5
(R
H3
-0.
5
M2
0
H2
0
0
(N)
)
M4(N
0
)?
-1
5
0.
-0.5
M2
(R)
0.5
-0.5
H1
1
40N
Figure 7. Magnetization solutions (0.5 A m−1 contour interval, grid interval 5 km) from 2-D inversion of the sea-surface magnetic anomaly chart (smaller
box in Fig. 3a). The upper surface of the magnetic layer was the top of the acoustic basement and the lower surface was 4 km deeper. Source magnetization
direction was assumed to be I M = 46◦ , DM = 0◦ . Thick black lines mark the positions of sharp magnetization contrasts which can be modelled as the edges of
seafloor spreading blocks M2(N)–M5(R). Black dots = ODP sites; red triangles = peridotite ridge segments; dashed red lines mark boundaries of regions B–D.
Dashed white line marks a boundary between regions of high (to the west) and low (to the east) magnetization. Solid white lines are estimates by eye of linear
trends in magnetization contours. Rms residual between observed and inverse calculated anomalies is 15 nT. Features labelled H1–H4 are discussed in the text.
overestimating depths in the distribution. We statistically analysed
the whole suite of solutions as a function of structural index. We
inspected the histograms in Fig. 9 as a subjective means of determining the SI that best represents the sources. In a qualitative way we
assume that a significant jump in a histogram may indicate a means
of selecting an appropriate SI since, for a homogeneous model, the
statistics for Euler solutions should vary smoothly with SI value.
Given the known varied geological nature of basement rocks
within this area we zoned the solutions according to the regions
A, B, C or D in which they fall (Fig. 9). In region A (oceanic crust)
an SI of 1.0 seems most appropriate because the depth histogram
is least skewed and the average solution depth coincides with the
average depth (∼7 km) to the top of the basement. This result and
the probable existence of a single population of solutions supports
the concept that the main sources of anomalies in this oceanic crust
are thin dykes or the edges of sills or lava flows near the top of the
crust. The distribution of solutions in region B is similar to that in
region A for all SI except for one important difference. A second
population of solutions exists about 7 km deeper than the first for
SI ≥ 0.5, although the absolute depth depends on the choice of SI.
An SI of 1.0 is the lowest acceptable index; it places the shallower
solutions at the top of the basement (∼7 km) and the deeper solutions at about 13 km depth. Region C differs from region B in
the character of basement relief and in having lower amplitude and
less strongly lineated magnetic anomalies. This difference is paralleled in solution depths, which are on average deeper than those
encountered in region B. For SI = 1.0 depths cluster around 10 km
and for SI =1.5 around 13 km. For SI ≥2.0 two populations appear
about 6 km apart. The preferred interpretation (see later) depends
C
2003 RAS, GJI, 154, 706–730
on the geological processes assumed to have controlled the nature
of sources as regions C, B and A formed sequentially.
Finally, region D has depth distributions that are very different
from the other three regions. There is a strong bias towards shallow
solutions (<4 km) even at high SI values. Although the main source
of magnetic anomalies lies at depths of about 2–3 km, other sources
may lie as deep as ∼15 km. Region D has a source distribution
similar to that of a larger area of continental crust (area S1; Fig. 1).
In area S1 the mean, median and modal depths computed by Euler
deconvolution for all SI all lie in the range 5–18 km b.s.l., a highly
probable depth range for magnetic sources within the continental
crust. Discontinuities in the frequency of solutions may indicate
that the most appropriate SI is 1.0 (thin dyke, sill or flow edge) or
1.5 (a mix of edges of thin dykes, sill or flow edges and vertical
volcanic conduits). If we assume an SI of 1.0 the depth histograms
suggest two families of source depths centred at 7 and 12 km with
likely variations of ±2.5 km.
Depths of sources from an analysis of near-bottom profiles
A more detailed investigation of the depths of magnetic sources
within the acoustic basement was carried out using the previously
mentioned near-bottom magnetometer profiles (Figs 2 and 10). Profile TOBI91 shows that west of 12◦ 10 W (75 km on the profile)
there are long-wavelength (∼20 km) and relatively high amplitude
(up to 600 nT) anomalies, whereas east of this point anomaly amplitudes are much reduced. Individually the SAR93 and SAR95 profile are too short to show the same contrast. Nevertheless, SAR93,
which intersects TOBI91 at its east end, also shows high-amplitude,
716
S. M. Russell and R. B. Whitmarsh
distance (km)
0
400
200
0
-200
-400
20
40
60
80
100
SAR93
4
2
0
-2
-4
899
898
OCT
400
200
0
-200
-400
4
2
0
-2
-4
898
(1069)
OCT
0
900
1067
1068
(1065)
100
50
0
-2
-4
-6
-8
-10
400
200
0
-200
-400
SAR95
4
2
0
-2
-4
0
-2
-4
-6
-8
-10
OCT
0
20
40
60
80
100
120
A m-1
(1070)
nT
TOBI91
km
897
nT
(1070)
A m-1
0
-2
-4
-6
-8
-10
km
km
A m-1
nT
(a)
140
distance (km)
Figure 8. (a) Inversion of the magnetic field intensity (5–200 km wavelengths) along three near-bottom profiles SAR93, TOBI91 and SAR95. The profiles are
aligned along a common longitude. The upper surface of the magnetic source layer was extracted from Fig. 2. A layer thickness of 4 km and the magnetization
direction of Van der Voo (1990; I M = 46◦ , DM = 0◦ ) were assumed. Upper panels show observed (solid line) and computed magnetic anomalies (dashed line).
Centre panels show magnetization solutions. Lower panels show the top of the acoustic basement (magnetic source layer; solid line), fish track (dashed line),
peridotite ridge (triangles), ODP sites (sites projected on to the tracks are in parentheses) and the width of the ZECM (labelled OCT). (b) Same as (a) but the
inversions are computed for wavelengths between 2.5 and 200 km.
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2003 RAS, GJI, 154, 706–730
Magmatism at non-volcanic rifted margin
0
(b)
distance (km)
40
60
80
100
SAR93
A m-1
nT
400
200
0
-200
-400
20
4
2
0
-2
-4
0
-2
-4
-6
-8
-10
(1070)
897
899
898
km
OCT
(1069)
OCT
0
900
1067
1068
(1065)
100
50
0
-2
-4
-6
-8
-10
A m-1
400
200
0
-200
-400
SAR95
OCT
0
20
40
60
80
100
distance (km)
Figure 8. (Continued.)
C
2003 RAS, GJI, 154, 706–730
120
140
4
2
0
-2
-4
A m-1
898
0
-2
-4
-6
-8
-10
km
(1070)
km
4
2
0
-2
-4
nT
400
200
0
-200
-400
nT
TOBI91
717
718
S. M. Russell and R. B. Whitmarsh
40
SI=0.0
SI=0.5
SI=1.0
SI=2.0
SI=1.5
SI=2.5
SI=3.0
Region A
0
40
frequency distribution (%)
Region B
0
40
Region C
0
40
{
{
Region D
0
0
10 20 30 0
10 20 30 0
10 20 30 0 10 20 30 0
source depth (km)
10 20 30 0
10 20 30 0
10 20 30
Figure 9. Histograms of Euler solution distributions (shaded bars) and basement depth distributions (open bars) in regions A–D as a function of the structural
index. Frequency distributions are given as a percentage of the total number of solutions in the given region and for the given SI. Arrows mark inferred source
clusters, parentheses indicate regions characterized by a broader distribution of source depth. Region A is characterized by a symmetrical distribution of source
depths centred on 7 km for SI = 1.0. Region B is clearly defined by two populations of sources 5 km apart although absolute depths depend on the value of
SI. The character of region C falls midway between that of regions A and B; a single source cluster at ∼10 km (SI = 1.0) splits into two populations with
increasing SI. Solutions in region D are heavily biased to shallow depths (<5 km), although for most SI the distribution extends to depths over 15 km.
long-wavelength anomalies along its entire length and SAR95, east
of 11◦ 55 W (25 km on the profile), shows very low-amplitude
(<50 nT) anomalies.
2-D Euler deconvolution of the three profiles suggests that three
regions, characterized by different sources, were traversed. We used
a window size of 2 km and only accepted solutions for which the
depth uncertainty did not exceed 20 per cent. First, there is a zone of
high solution density with solution patterns that are ∩-shaped and
occur west of ∼85 km on profile SAR93, west of ∼75 km on profile
TOBI91 and west of 15 km on profile SAR95 (Fig. 10). ∩-shaped
patterns of Euler solutions commonly result from the interference of
closely spaced sources and/or sources that are strongly magnetized.
Because the lateral error introduced by any interference is unknown,
a vertical magnetic boundary is assumed to lie near the apex of each
∩-shaped pattern. Second, a zone of intermediate solution density
occurs east of 85 km on SAR93, east of 75 km on TOBI91 and
at 15–30 km on SAR95. ∩-shaped patterns are poorly developed,
which suggests that either neighbouring boundaries are sufficiently
far apart that little interference occurs and/or that source body magnetization contrasts are weaker than in the first zone. Third, a zone
of very sparse solution density is present east of ∼30 km on SAR95.
Here, clusters of solutions at 78, 96 and 112 km suggest the presence
of weak or deeply buried magnetic contrasts.
The depths of the Euler solutions west of the peridotite ridge
(region A) nearly coincide with the top of the acoustic basement. In
contrast, further east, the tops of the blocks from the peridotite ridge
to a positive magnetic anomaly labelled B1(N) (region B, Fig. 10)
lie up to 2 km below the basement surface (assuming SI = 0.0).
East of B1(N) (region C) the ∩-shaped patterns (SI = 0.0) are less
strongly developed and an SI = 1.0 (dyke-like contact) is probably
more appropriate; in this case most solutions lie within 8 km of the
basement surface.
The above interpretation of source depths is supported by
the spectral character of the three near-bottom profiles. Power
C
2003 RAS, GJI, 154, 706–730
Magmatism at non-volcanic rifted margin
719
distance (km)
80
100
400
200
0
-200
-400
0
SAR93
A
B
897
(1070)
899
898
-10
B1(N)
PR
M3(R)(?)
M3(R)
M2(N)
M1(N)
M1(R)
-20
TOBI91
A
B
C
898
D
(1069)
1068, 900
1067
400
200
0
-200
-400
0
nT
60
(1065)
-10
km
40
nT
20
km
0
??
B1(N)
PR
M4(N)
M3(R)
M2(N)
M1(R)
M1(N)
J
?
-20
100
50
400
200
0
-200
-400
0
km
nT
0
?
SAR95
D
C
-10
0
B1(N)
?
-20
?
20
40
60
80
100
120
140
distance (km)
Figure 10. Euler deconvolution of magnetic anomalies (upper panels) along deep-tow profiles SAR93, TOBI91 and SAR95 aligned along a common longitude.
Profile tracks are given in Fig. 2. Window length = 2 km. The approximate locations of the boundaries between regions A–D are shown by vertical bars. The
solutions (lower panels) are for structural indices of 0.0 (semi-infinite contact; red circles), 0.5 (step contact; green diamonds) and 1.0 (dyke-like contacts; blue
squares). ODP site numbers, peridotite ridge (triangle), fish tracks (dashed lines) and acoustic basement are shown in the lower panels. Boundaries of magnetic
blocks were determined by the positions of apices of solution clusters (SI = 0.0); contacts of seafloor spreading blocks and other magnetic contacts (including
B1(N)) are labelled.
spectra, radially averaged in the horizontal plane (Bhattacharyya
1966; Spector & Grant 1970), were computed and source
depths were roughly estimated by fitting straight lines to bandlimited segments (Fig. 11). The results are relatively clear
C
2003 RAS, GJI, 154, 706–730
and consistent among the three profiles, suggesting that they
are reasonable estimates. Along SAR93 and TOBI91 sources
lie ∼2.5 km below the magnetometer and along SAR95
they lie at ∼4.0 km, i.e. near the top of the basement
720
S. M. Russell and R. B. Whitmarsh
(Fig. 10). In addition, SAR95, which extends much further east
coincident with the MCS profile IAM-9 (region C), shows a source
at ∼6.5 km below the top of the basement. Sources 29 km below
the magnetometer are attributed to a component of the regional field
that was not removed entirely in attempting to separate the regional
and residual fields.
14
SAR93
12
29 km
ln(Power)
10
8
6
S U M M A RY O F M A G N E T I C T R E N D S ,
I N T E N S I T I E S A N D S O U RC E D E P T H S
4
2
2.5 km
0
-2
14
TOBI91
12
29 km
ln(Power)
10
8
6
2.5 km
4
2
0
-2
14
SAR95
12
ln(Power)
10
29 km
8
6
10.5 km
4
4 km
2
0
-2
0.0
0.2
0.4
0.6
0.8
wavenumber (rad km−1)
1.0
Figure 11. Logarithms of the power spectrum calculated from the total
field anomaly along deep-tow magnetometer profiles SAR93 (top), TOBI91
(centre) and SAR95 (bottom). First-order regional trends were removed from
profiles before the computation of each spectrum. Each spectrum has been
modelled using Spector and Grant’s (1970) method of fitting straight lines
to band-limited segments. On profiles SAR93 and TOBI91 identical and
distinct source depths of 29 and 2.5 km (below the observations level) are
successfully modelled. Each fish was towed, on average, ∼2.5 km above the
top of the basement. Therefore, source depths are shown to be consistently
at, and 26.5 km below, the top of the basement in the vicinity of the ODP
drilling transect. On profile SAR95, three sources are modelled successfully
and found to be at the top of the basement and ∼6.5 and ∼25 km below the
top of the basement. An average fish height of ∼4 km above the basement was
used. The deepest source modelled in each case is attributed to the regional
field component of the total field anomaly along each of the near-bottom
profiles SAR93, TOBI91 and SAR95.
We see a difference in the magnetic source characteristics in the
four regions A–D. The sources in region D (continental crust) are
either lineated, but with no preferred orientation, or are point bodies
or bodies with a vertical axis of symmetry. If all structural indices
are considered, the sources could be in the basement at any depth
<38 km b.s.l., but if the sources are vertical structures, possibly
dykes, they lie at depths in the range 4–15 km. The magnetic sources
may be interpreted as being uniformly magnetized continental crust
and dyke-like intrusions within, or at the base of, thinned continental
crust.
The sources in regions B and C (ZECM) have been shown to
be preferentially lineated ∼010◦ . 2-D and 3-D Euler deconvolution
and 2-D power spectra show that the sources differ in depth between
regions B and C.
In region C the basement depth is typically >
∼8 km. Euler deconvolution shows that sources lie within ∼2 km of the top of the
basement and could be represented by a layer-type geometry (SI
= 1.0) consistent with a thin layer of relatively magnetic serpentinized peridotites. This layer of sources could account for the 25
per cent correlation of magnetic anomalies derived from a uniformly
magnetized basement. A deeper set of sources underlies region C,
which has been variously estimated at ∼8 km below the top of
the basement, with a vertical pipe-like geometry (SI = 2.0), dykelike, sill or flow edge geometry (SI = 1.0), or indeed something
in between, for example, a non-semi-infinite dyke. Power spectra
from near-bottom profiles independently indicate sources at ∼6.5
km below the top of the basement. The magnetic anomaly trends,
that appear after reduction to the pole of the observed magnetic field,
suggest the sources possess a remanent magnetization that was more
than likely acquired after the initiation of the rotation of Iberia as
the Bay of Biscay opened or that sources possess solely induced
magnetization.
In region B the basement depth is typically <7.5 km. Euler deconvolution suggests that sources lie both at the top of the basement (∼7
km) and ∼6 km deeper. We interpret sources as likely to be dykes or
the edges of sills or flows (SI = 1.0). The magnetic anomaly trends
that appear after reduction to the pole of the observed magnetic field
suggest that if the sources possess a remanent magnetization then
this was acquired after the initiation of the rotation of Iberia.
In region A (oceanic crust) the anomalies have a dominant 015◦
trend. The anomalies are sharpest after reduction to the pole is applied with a declination of 0◦ , suggesting magnetization was acquired after most of the rotation of Iberia had occurred. Euler deconvolution leads to source depths that are best described by SI =
1.0 and plot near to the top of the basement (∼7 km). The shape
of the depth histograms from Euler deconvolution for region A are
more like those of region B than region C.
We have demonstrated that because basement relief alone plays
only a minor role in creating the magnetic anomalies observed over
the continental margin off west Iberia that magnetization variations must lie within the basement itself. Furthermore, because the
four regions of different basement relief exhibit different magnetic
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2003 RAS, GJI, 154, 706–730
Magmatism at non-volcanic rifted margin
properties, this suggests that the nature of the sources and possibly
their origin vary across the margin. In particular, the ZECM (regions
B and C) exhibits different trends of linear magnetic anomalies, and
different source depths and source character from the oceanic crust
(region A).
The objective of the next sections is to further investigate the
magnetic sources using forward modelling and inverse techniques,
to establish the probable extent of magnetic anomalies generated
by seafloor spreading and to produce constraints on the geological
processes that generated any non-seafloor spreading anomalies.
DISCUSSION
Seafloor spreading
Seafloor spreading is characterized by sequences of linear, subparallel magnetic anomalies that can be matched by anomalies computed
400
300
200
100
0
-100
-200
-300
-400
rms = 112.47
cc = 0.6526
nT
(a)
721
TOBI91
depth (km)
0
5
M5(R)
M4(N)
R3
M3(R)
M2(N)
M1(N)
10
J
15
0
50
100
150
distance (km)
nT
(b)
100
0
300
rms = 26.68
nT
200
cc = 0.9495
observed
100
0
-100
SAR95
calculated
-200
0
acoustic
basement
M3(R)
R3
M1(R)
M1(N)
M2(N)
J
10
M0(R)
depth (km)
13.5 mm yr-1
5
L1
L2
15
20
IAM-9
A m-1
25
5
0
-5
inverse solution
approximate block model
0
50
100
150
200
250
Figure 12. (a) Magnetic models along deep-tow profile TOBI91. Upper panel, observed (solid line) and computed (dotted line) anomalies, cc = crosscorrelation coefficient; lower panel, magnetic model showing crustal layer and seafloor spreading blocks (assuming ±0.8 A m−1 ), except for a non-magnetic
block R3, and magnetization of Van der Voo (1990; I M = 46◦ , DM = 0◦ ). (b) Magnetic models along deep-tow profile SAR95 coincident with MCS profile
IAM-9. From bottom to top, panel 1, inverse derived magnetization solution; panel 2, magnetic models showing base of ‘crustal’ layer (dashed line), seafloor
spreading blocks (±0.8 A m−1 ), continental crust (0.1 A m−1 , stippled) and undefined magnetic bodies (−0.8 A m−1 , grey); panel 3, observed and calculated
anomalies; panel 4, difference between the anomalies in panel 3. Other parameters as in (a).
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2003 RAS, GJI, 154, 706–730
722
S. M. Russell and R. B. Whitmarsh
Table 2. Computed widths of constant polarity intervals compared with observed block widths (largely from Fig. 7) and estimated spreading
rates (timescale of Gradstein et al. 1995). NO = not observed.
Block
M0(R)
M1(N)
M1(R)
M2(N)
M3(R)
M4(N)
M5(R)
M5(N)
M6(R)
M1(R)–M5(R)
Computed
width (km)
@ 6 mm yr−1
Computed
width (km)
@ 8 mm yr−1
Computed
width (km)
@ 10 mm yr−1
Computed
width (km)
@ 12 mm yr−1
Observed
width (km)
3.6
16.1
2.2
4.1
12.1
5.7
3.1
0.8
0.7
5.8
21.5
3.00
5.4
16.1
7.6
4.1
1.1
0.9
6.0
27.0
3.7
6.8
20.1
9.5
5.1
1.4
1.1
7.2
32.3
4.4
8.2
24.1
11.4
6.1
1.7
1.3
NO
NO
∼10
∼6.5
∼14
∼11
∼8.5
NO
NO
∼50
from subsets of the geomagnetic reversal timescale, assuming a
constant spreading rate. The recognition of seafloor spreading by
this means is severely limited, when either the lengths of the subsets are constrained by few (≤ four) reversals or the spreading rate
changes, because the computed profile lacks uniqueness. From the
2-D inversion results alone (Fig. 7), magnetic boundaries west of
a point ∼20 km east of peridotite ridge segment R3 can plausibly
be equated with a sequence of seafloor spreading anomalies from
M5(R) to M2(R) (Table 2). The mean spreading rate from M5(R)
to M1(R) is 11.1 mm yr−1 and the mean weighted by the estimated
mean width of the individual blocks, because wider blocks will give
better estimates, is 14 mm yr−1 . Profile TOBI91 can be modelled
over the same time period at a rate of 10 mm yr−1 , provided that the
peridotite ridge is excluded from the reversal sequence (Fig. 12a).
A model of the sea-surface profile along IAM-9 was adjusted to
include the thickness of the seismically defined oceanic crust, and
of the serpentinized peridotite basement in the ZECM, from Dean
et al. (2000). The best fit was obtained for a spreading rate of 13.5
mm yr−1 from M3(R) to M0(R) (Fig. 12b).
2-D Euler deconvolution of the near-bottom profiles indicated
the existence of vertically sided boundaries to the source bodies
(Fig. 10), which have been tentatively assigned to reversal boundaries in the geomagnetic timescale. This analysis provides independent confirmation that the oldest seafloor spreading anomaly along
TOBI91 is M4(N) and along SAR95 is M5(R) (Fig. 10). The block
labelled B1(N) on TOBI91 and SAR93 appears to coincide with the
isolated and atypical basement high where hole 899B encountered
strongly magnetized peridotite breccia (Zhao et al. 2001).
A final technique, which computes a function called the intensity
of spatial differential vectors (ISDV) (Korenaga 1995; Seama et al.
1993), was applied to the two near-bottom profiles that recorded
all three orthogonal components of the Earth’s magnetic field. The
maxima in ISDV, above a noise threshold of 50 nT km−1 , were used
to estimate the location of, assumed vertical, magnetization boundaries. The technique identified 22 and 11 magnetization boundaries
on profiles TOBI91 and SAR95, respectively (Fig. 13). On comparing the positions of ISDV maxima and a model of seafloor spreading, from the time of M4(N) until M1(N), boundaries 1–6 on profile
TOBI91 match the positions of reversal boundaries if spreading rates
of 10 and 6 mm yr−1 are assumed, respectively, west and east of block
M3(R). No other boundaries can be matched on either profile at any
spreading rate.
Estimated
spreading rate
(mm yr−1 )
27
10
7
12
17
11.1
Onset of seafloor spreading
A variety of independent inverse and forward modelling methods
used to investigate the magnetic anomalies in the southern Iberia
Abyssal Plain support the concept of continuous sea-floor spreading starting around the time of M5(R) or M4(N), i.e. 128–127 Ma,
at a spreading rate of 10–14 mm yr−1 . The earliest block that can be
modelled lies ∼20 km east of peridotite ridge segment R3. Although
magnetic polarity intervals M4–M15 lasted on average for a much
shorter time (∼260 ka) than the succeeding Mesozoic intervals,
and therefore are unlikely to be resolved on sea-surface profiles,
they should have been resolvable along near-bottom profiles such
as SAR95 for spreading rates ≥4 mm yr−1 . It is very clear, however, that any such anomalies earlier than M5(R) are absent from
SAR95.
Isotopic dating of basement drill cores in the southern Iberia
Abyssal Plain suggests that the mean rate of extension increased
from ∼7 to ∼10 mm yr−1 from the exhumation of the upper mantle
in the ZECM to the onset of seafloor spreading west of the peridotite ridge (Manatschal et al. 2001; Whitmarsh et al. 2001a). Thus,
we attempted to explain the observed magnetic profiles by seafloor
spreading using models with varying but increasing spreading rates.
One attempt to model the near-bottom profile SAR95 with spreading
rates incrementing from 6 to 10 mm yr−1 (starting at M19(N) time),
and source depths according to the Euler deconvolution results, only
partially succeeded. Anomalies between 220 and 265 km do match
but anomalies between 165–220 km are not matched (Fig. 14a;
Russell 1999). Another attempt to model the near-bottom profile
TOBI91 (Srivastava et al. 2000), using blocks reaching to the top
of the basement and spreading rates between 6.2 and 9 mm yr−1
(beginning before M11 time; Fig. 14b), emphasized the isolated
and atypical subcircular anomaly near site 898. This anomaly is
apparently associated with the unusually strongly magnetized serpentinized peridotite breccia at nearby site 899 (Sawyer et al. 1994).
The second attempt failed to satisfactorily match the amplitudes and
phases of anomalies in the ZECM between sites 900 and 898.
Thus, we conclude that steady-state seafloor spreading prior to the
time of anomaly M5(R) appears unlikely. This is consistent with both
the unusual, intrabasement location of many sources in the ZECM
and with the corresponding seismic velocity structure, which largely
lies outside the bounds of normal Atlantic oceanic crust. In the next
section we discuss the possible alternative geological processes that
C
2003 RAS, GJI, 154, 706–730
nT
0
0
500
-1000
-500
nT
500
1000
-1000
-500
nT
500
1000
-1000
-500
0
X
M3(R)
7
15
16
X
11.5
900
1067
1068
(1065)
11.0
Basement
Fish track
50 km
17 18 19 20 2122
threshold
12.0
longitude (degrees W)
898
14
(1069)
111213
6 mm yr-1
X
8 9 10
TOBI91
M16(N)
M4(N)
M1(N)
-10
-8
-6
-4
0
500
-500
0
500
-500
0
500
-500
0
500
1 2
4
5
12.0
3
6
7 8
9
ISDV
N component anomaly
11
Basement
Fish track
SAR95
11.5
11.0
longitude (degrees W)
10
E component anomaly
Z component anomaly
50 km
10.5
Figure 13. Vector anomaly (and gradient, dotted line) and ISDV profiles along deep-tow profiles TOBI91 (left) and SAR95 (right) (note the different ordinate scales). Positions of magnetic boundaries are indicated
by maxima in ISDV (numbered arrows). Along TOBI91, boundaries correlate reasonably well with theoretical magnetic contacts west of boundary number 6 (start of M4(N)) for the spreading rates shown. Along
SAR95, boundaries do not correlate with magnetic contacts derived from seafloor spreading models so no block model is shown.
12.5
(897)
5 6
M4(N)
-10
13.0
-8
10 mm yr-1
X
1 2 3 4
ISDV
N component anomaly
E component anomaly
Z component anomaly
M3(R)
-6
-4
0
nT km-1
2003 RAS, GJI, 154, 706–730
M17(R)
depth (km)
C
500
1000
Magmatism at non-volcanic rifted margin
723
724
S. M. Russell and R. B. Whitmarsh
(a)
200
nT
100
0
aZ(R)
aY(N)
aM4(N)
-100
-200
0
50
100
150
200
250
300
350
300
350
depth (km)
5
10
15
10 mm yr-1
8 mm yr-1
6 mm yr-1
20
0
50
100
150
distance (km)
200
250
Figure 14. (a) Observed (fine line) and calculated (solid line) magnetic anomalies (upper panel) across the southern Iberia Abyssal Plain along deep-tow
profile SAR95 according to the block model shown in the lower panel. The model assumes the timescale of Gradstein et al. (1995), the magnetization direction
of Galdeano et al. (1989; I M = 58◦ , DM = −43◦ ) and intensities of ±1 A m−1 . The observed profile was truncated at the J anomaly (Fig. 1) for clarity. (b)
Part of Fig. 4(b) of Srivastava et al. (2000) is reprinted with permission from Elsevier. Model of TOBI91 labelled with Iap (DT) varying spreading rate.
gave rise to the ZECM and the observed distribution of magnetic
anomalies.
Geological interpretation of results
The magnetic sources in the region of thinned continental crust
(region D) can be distinguished from those adjacent to, and within,
the basement of the ZECM further west. The sources are broadly
distributed in depth and in plan view they generally possess no
dominant trend or are subcircular. Some region D anomalies are
only moderately lineated in a 350◦ direction, mostly towards the
western edge of thinned continental crust.
The longitude (11◦ 15 W) at which anomaly trends change from
∼350◦ to ∼010◦ does not coincide exactly with the seaward edge of
the thinned continental crust defined above (Fig. 3). Indeed, some
lineated ∼010◦ ZECM anomalies continue as far north as site 1065,
i.e. ∼33 km north of, and ∼50 km east of, the suggested edge. It
is therefore unlikely that a distinct discontinuity in magnetic character occurs at the proposed edge of the continental crust. Because
the basement in the vicinity of site 1065 is unequivocally thinned
continental crust, but further south it is unequivocally not continental crust, this suggests that the sources of the ∼010◦ anomalies lie
within the upper mantle in agreement with source depths derived
from Euler deconvolution. They cannot therefore be typical seafloor
spreading anomaly sources.
The origin of the lineated ∼350◦ anomalies within the westernmost thinned continental crust and easternmost ZECM are therefore likely to be related, and are possibly a result of the same geological process. Furthermore, these anomalies are similar to the
anomalies in most of the ZECM, in amplitude and wavelength,
but not trend, so that the sources of these anomalies may be more
akin to those of the ZECM, rather than to those of the continental
crust.
Reduced-to-the-pole anomalies over the majority of the ZECM
are linear with a trend of ∼010◦ , roughly parallel to the seafloor
spreading anomalies yet slightly oblique to those observed over
thinned continental crust. Even though some of these anomalies
extend over many kilometres, even into the Tagus Abyssal Plain
(Fig. 1), Euler solution positions and local magnetization zero crossings, based on deep-tow profiles, do not successfully retrieve the
boundaries of geomagnetic reversal sequence blocks if seafloor
spreading is assumed. In fact, Euler deconvolution and spectral analyses support a model in which the sources lie up to ∼5 km below
the top of the basement. In particular, Euler deconvolution solutions
for linear magnetic sources within region C of the ZECM indicate
depths of ∼11 and ∼13 km for SIs of 1.0 and 1.5, respectively, and
both ∼10 and ∼16 km for an SI of 2. These estimates lie within ∼7
km of the top of the basement. Within region B, where basement
depths average 7 km, a bipolar depth distribution is indicated with
depths separated by ∼5 km; for dyke and sill-like geometries the
upper set of solutions falls at the top of the basement. In contrast,
C
2003 RAS, GJI, 154, 706–730
Magmatism at non-volcanic rifted margin
many solutions for region A (oceanic crust) cluster around the top
of the basement.
Nature of the source bodies
Based on the above conclusions we propose that, depending on location, the magnetic source bodies are either syn-extensional gabbroic
intrusions or products of seafloor spreading magmatism. In particular, we propose that regions C, B and A represent a progression from
late rift/breakup-related magmatism to the organized, normally and
reversely magnetized, crustal blocks associated with seafloor spreading. Thus, the intrusions in the western part of region C, and certainly
those in region B, may be regarded as protooceanic-type bodies because, at least locally and discontinuously, they match predicted
seafloor spreading anomaly patterns. Unfortunately, no cores exist
to support or refute this hypothesis. Currently, the region C basement
remains unsampled and elsewhere in the ZECM ODP Legs 149 and
173 did not sample unequivocal in situ Early Cretaceous intrusions.
The mafic rocks (metagabbro, tonalite and amphibolite) retrieved
from sites 900, 1067 and 1068 appear to date from Variscan time
and are not ‘syn-rift’ (Manatschal et al. 2001), even though they
were exhumed around 136 Ma (Féraud et al. 1996). However, it is
well known that the lower oceanic crust can contribute significantly
to observed anomalies (e.g. Harrison 1987), probably through the
remanent magnetization of doleritic (isotropic), non-cumulate gabbros (e.g. Dunlop & Prévot 1982). Thus, when intruded into unaltered peridotite, which typically has a relatively weak NRM (Toft
& Arkani-Hamed 1992), or even into peridotite of opposite polarity, such gabbros can lead to magnetization contrasts that produce
substantial anomalies.
An alternative hypothesis, that the sources of the linear anomalies
are associated with some form of inhomogeneity within the serpentinized peridotite, upper-mantle rocks, is based on the observation
that magnetite is a common product of the reaction between olivine
and/or pyroxene and water under the conditions required for serpentinization to take place. Furthermore, Coleman (1971) showed
that the magnetization intensities of serpentinization products are
strongly related to the degree of serpentinization. The most serpentinized peridotites tend to have the highest susceptibilities and such
rocks are most likely to occur in the uppermost basement. Results
from ODP cores from the Iberia Abyssal Plain show that the uppermost basement serpentinized peridotites possess relatively high
NRMs and susceptibilities (Table 3).
However, there are three principal arguments against this hypothesis. First, a homogeneously and highly serpentinized upper basement, as proposed by several authors (Pickup et al. 1996; Chian
et al. 1999; Dean et al. 2000), should give rise to anomalies that
correlate strongly with basement relief, but no such correlation is
observed on any appreciable length-scale (Fig. 6). Second, localized
serpentinization controlled by water flowing preferentially along
linear fault zones could cause linear anomalies but examination of
our unpublished, closely spaced, MCS profiles across the southern
Iberia Abyssal Plain (Whitmarsh et al. 2000) reveals that individual
faults cannot be traced along strike for more than ∼10 km. On the
other hand, the linear anomalies within the ZECM are continuous
for many tens of kilometres and are therefore unlikely to be related
directly to specific individual faults or fault zones. Third, sources
demonstrated to lie ∼5 km below top basement, and blanketed today
by 2–3 km of sediment, are unlikely to be associated with the most
serpentinized, and therefore most strongly magnetized, peridotites.
Oxygen isotopes strongly suggest that serpentinization occurred at
temperatures <200◦ C (Agrinier et al. 1996) or even <50–150◦ C
(Skelton & Valley 2000), associated with a downward flow of water
implying that in general the greatest serpentinization, and highest
magnetization, will occur near the top of the basement.
A conceptual geological model
Thus, the linear magnetic anomalies in the ZECM of the southern Iberia Abyssal Plain may be principally the product of a series
of syn-extensional gabbroic bodies intruded into the serpentinized
peridotite basement and, in the western part of the ZECM, these
bodies first produced anomalies that can be modelled by continuous seafloor spreading. These conclusions support the following
2-D model briefly described by Whitmarsh et al. (2001b). This
model focuses on continental crust that has already been thinned to
∼10 km, probably during the final rift episode in which the crust
is dissected by numerous low-angle normal faults and eventually
breaks (Whitmarsh et al. 2000), and on the uppermost lithosphere
seaward of this crust (Fig. 15). Although the model is 2-D in reality it expresses a snapshot, in time and space, of the situation that
developed as rifting propagated from south to north along the west
Iberia margin (Whitmarsh & Wallace 2001).
The linear ∼350◦ -trending anomalies in the vicinity of the edge
of thinned continental crust probably represent the first significant mantle-derived melt products during the final stage of rifting
(∼130 Ma) on this part of the margin, since only small amounts of
130–144 Ma transitional dolerites, gabbros and diorites have been
recorded on land in the Lusitanian Basin (Ribeiro et al. 1979; Martins 1991). However, these marine anomalies do not parallel the NNE
and ESE trends of the dykes on land nor do they parallel the seafloor
spreading anomalies further offshore. They cannot be modelled by
seafloor spreading either. We explain these anomalies as follows.
In the model, rifting starts with pure shear deformation of a fourlayered brittle–ductile continental lithosphere with localized simple
Table 3. Magnetic properties of serpentinized peridotite cores from the top 200 m of acoustic basement in the Iberia Abyssal Plain.
ODP Leg
Site
103
149
149
149
637
897
899
899
173
173
1068
1070
∗ Estimated
C
Lithology
Serpentinized peridotite
Serpentinized peridotite
Serpentinized breccia
Serpentinized peridotite,
microgabbro and sediment
Serpentinized peridotite
Serpentinized peridotite
from a graphical plot.
2003 RAS, GJI, 154, 706–730
725
Mean
susceptibility
(SI units ×10−3 )
Mean NRM
(A m−1 )
Source
–
19.0
28.5
36.3
1.2
0.28
1.28
1.33
(Boillot et al. 1987)
(Zhao 1996; Zhao 2001)
(Zhao 2001)
(Zhao 1996; Zhao 2001)
1.5∗
28.4
0.05∗
1.39
(Shipboard Scientific Party 1998)
(Zhao 2001)
726
S. M. Russell and R. B. Whitmarsh
steady-state
seafloor
spreading
~10 mm/yr
3
4
R
R
Site
900 (?)
s
d
nin e tal ward
th nen to DG
ti
v
oc n rust
c
drilling transect
IAM-9
transect
0
M3(R)
M2(N)
M1(R)
depth below top
basement (km)
5
0
N
quasiM4(N)(?)
unaltered mantle peridotite
highly sepentinized
mantle peridotite
L1 (?)
5
25 km
unaltered mantle peridotite
10
10
Figure 15. Schematic 3-D model of the proposed geological structure beneath the southern Iberia Abyssal Plain. During prolonged extension of the lithosphere
continental crust is thinned to breaking point and continental mantle peridotite (red) is exhumed at the seafloor and subsequently serpentinized (white layer over
red layer). The ZECM lies between the thinned continental crust (yellow; VDG = Vasco da Gama Seamount) and crust formed by seafloor spreading (alternating
black and white blocks labelled M3(R)–M1(R)). Reduced-to-the-pole magnetic anomalies over thinned continental crust and the most landward region of the
ZECM in the vicinity of MCS profile IAM-9 trend ∼350◦ (blue lines with ticks). These anomalies are more linear and similar in amplitude to those further
east (over continental crust), yet are oblique (∼010◦ ), although of similar amplitude, to seafloor spreading anomalies and other anomalies within the remainder
of the ZECM to the west (black lines with ticks). The linear ∼350◦ -trending anomalies result from the syn-extensional intrusion of dyke- and/or sill-like
gabbroic bodies (blue/grey) into exhumed mantle under rift-related stress conditions. Serpentinization of exhumed mantle (white) prevents other intrusions
(grey) from rising buoyantly above subbasement depths of ∼3 km. Anomalies trending ∼010◦ over thinned continental crust are evidence of subcontinental
intrusions, which post-date continental breakup and the pre-existing stress conditions. Through time and space (oceanwards) increased focusing of melt creates
quasi-oceanic and normal oceanic crust (black and white blocks) possibly aided by greater differentiation of melt and/or less serpentinization of the mantle.
shear in the upper crust and upper mantle (e.g. Brun & Beslier
1996; Pickup et al. 1996; Whitmarsh et al. 2000, 2001a). Continued extensional thinning of the crust and lithosphere at ≥3.5
mm yr−1 (Whitmarsh et al. 2001a) results in the upward migration
of the asthenosphere and crust–mantle boundary until the whole
crust and uppermost mantle fall entirely within the brittle regime
(Whitmarsh et al. 2000; Whitmarsh et al. 2001a; Perez-Gussinye &
Reston 2001).
Eventually, as extension continues, adiabatic decompression
melting of the asthenosphere occurs and small amounts of melt rise
buoyantly towards the surface. Such melts may penetrate the thinned
continental crust but in the ZECM may become trapped below the
top basement where their density equals that of the surrounding serpentinized peridotite (e.g. Herzberg et al. 1983). Once continental
crust has broken, mantle rocks are exhumed at the seafloor by a
series of low-angle, concave-downwards, detachment faults, analogous to the rolling-hinge faults of Lavier et al. (2000), which act
to bring mantle up to the seafloor above, and on the flanks of, a rising and ever narrowing ridge of asthenosphere and at an increased
(≥7.2 mm yr−1 ) horizontal extension rate (Whitmarsh et al. 2001a).
This dynamic system will lead to the generation of greater volumes
of melt, which will ascend towards the surface, again subject to their
buoyancy relative to the surrounding rocks. For example, a basaltic
melt density of ∼3.0 Mg m−3 , equivalent to a seismic velocity of ∼7
km s−1 (White & McKenzie 1989), would match the seismic model
of Dean et al. (2000) at 2–3.5 km below the top of the basement
along profile IAM-9 and place the melt well within the predicted
source depth range computed from the coincident SAR95 profile
(<8 km). As extension increases we speculate that the production
of greater volumes of melt leads to more closely spaced but still isolated intrusions and eventually, because of magmatic differentiation
principally by fractional crystallization and possibly vesiculation,
lower-density melts that are emplaced at higher levels in, or even
over, the serpentinized peridotite basement layer. This is consistent with the observation of magnetic sources both at the top of
the basement and ∼5 km deeper in the western part of the ZECM.
However, at this stage, the intrusions do not yet form a continuous
oceanic layer 3 and wide-angle seismic observations probably cannot resolve individual magnetic source bodies (Dean et al. 2000),
even assuming that such bodies have a distinct seismic signature.
Nevertheless, the observed apparent increase in horizontal extension rate at this stage of exhumation of continental mantle may be
explained partially by the increased volume of melt intruded into
the basement. Finally, oceanic crustal accretion becomes the dominant process in the vicinity of peridotite ridge segment R3. Seafloor
spreading anomalies can be modelled at spreading rates of ∼10–
14 mm yr−1 with an oceanic crust that has an almost normal thickness. Source depths are focused at or near the basement surface
because the magnetically dominant feature is now the top of the
oceanic crust whether capped by extrusions or, in their absence as
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2003 RAS, GJI, 154, 706–730
Magmatism at non-volcanic rifted margin
at site 1070, by intrusions. The above sequence of events is shown
diagrammatically in Fig. 15.
countered during rift propagation, such as pre-existing weak zones
in the continental lithosphere and their influence on horizontal strain
rates, temperature and stress fields and melt production, which are
beyond the scope of this paper. Published studies of the conjugate
Newfoundland margin are less detailed. For example, we only tentatively recognize several features of the magnetic anomalies on that
margin in common with the west Iberia margin.
Our model provides an answer to the question, posed at the outset,
‘what magmatic processes accompany the breakup of continental
lithosphere and the subsequent mantle exhumation leading up to
the onset of steady-state seafloor spreading?’. We have described
how we think magmatism occurs, in space and time, in the zone
of exhumed continental mantle, which lies between the faulted and
extended continental crust and the oceanic crust produced by steadystate seafloor spreading.
The magmatism that we propose, largely from the results of inverting magnetic observations, is based on inferences that require
direct observation for their confirmation. This can only be realized
by future deep drilling into basement rocks on rifted continental
margins.
Concluding remarks
Our qualitative model of the extensional processes active during
continental breakup, continental mantle exhumation and the onset
of seafloor spreading is strongly supported by extensive geophysical and geological observations, which are principally quantitative.
The model is also supported by a recent quantitative investigation
of melt production within the ZECM. Based on the results of numerical modelling Minshull et al. (2001) concluded that the lack
of significant melt production within the ZECM could be explained
in part, but not completely, by lateral heat loss to the adjacent continental lithosphere, by depth-dependent stretching (such that the
observed infinite stretching factor for the crust is not representative
of the whole continental lithosphere) or even by an anomalously
low mantle potential temperature at breakup time. However, significantly, they introduced a further explanation, adopted above and
by Whitmarsh et al. (2001a), that invokes a transitional state, between the end of continental extension and the onset of steady-state
seafloor spreading, during which mantle upwelling is less focused
than at normal spreading centres. By analogy with spreading ridges,
when a mantle viscosity structure is chosen that makes the passive
upwelling relatively unfocused then melt production decreases at
slow spreading rates (e.g. Shen & Forsyth 1992), similar to the rates
of extension inferred off west Iberia.
The 2-D model in Fig. 15 is diagrammatic and somewhat speculative. It omits some observed 3-D features that relate to the south-tonorth rift propagation along the west Iberia margin and it is based
on observations from only half of a conjugate pair of rifted margins. An attempt to explain such along-the-margin features is made
diagrammatically in Fig. 16. Fig. 15 also ignores factors such as
variations in the width of the ZECM along the margin and the fact
that the oceanward edge of the thinned continental crust does not
parallel the margin. These factors may reflect inhomogeneities en(a)
(b)
TC
C O N C LU S I O N S
Our main conclusions are that off W Iberia:
(1) Within the zone of thinned continental crust magnetic anomalies are either linear with no preferred direction or subcircular. Magnetic source bodies are scattered throughout the crust and, if they
are vertical dyke-like structures, lie at depths of 7 ± 2.5 and 12 ±
2.5 km below sea level. Towards the oceanward edge of continental
crust linear anomalies tend to trend ∼350◦ .
(2) Magnetic anomalies change their character within the zone of
exhumed continental mantle (formerly called the ocean–continent
transition zone). Landward there are linear ∼350◦ anomalies some
of which can be followed northwards along strike over thinned continental crust. Therefore, we infer that the sources of such anomalies
over continental crust exist within the upper mantle. Here the sources
(d)
(c)
TC
TC
TC
UM
UM
W
W
727
TC
TC
UM
2
1
W
TC
UM
4
3
2
1
W
TC
UM
Figure 16. Cartoon maps and sections across their southern edges (not to scale) showing the progressive stages in the rift-to-drift development of the southern
Iberia Abyssal Plain continental margin. The model is based upon the propagating rift model of Martin (1984). W = water; TC = thinned continental crust; UM
= upper mantle. Fault tick sizes denote relative degrees of extension (heave). Stippled objects are syn-extensional gabbroic-type intrusions, black and white
bars represent blocks of normal oceanic crust produced by steady-state seafloor spreading. The arrow indicates the direction of propagation of the rift. Dashed
lines mark successive locations and the shape of the top of the asthenosphere. (a) Thinning of the lithosphere by tectonic extension; strain is manifested by fault
blocks in the upper continental crust; (b) prolonged lithospheric extension results in exhumation of the upper mantle at the seafloor. Penetration of seawater
into the mantle (serpentinization) may be aided by brittle deformation. Very small amounts of syn-rift melt intrude into the mantle but their ascent is halted by
the lower-density serpentinized peridotite layer. Trends of magnetic anomalies, generated by intrusive bodies, are dominated by the continental rift stress field;
(c) the combination of the upwelling asthenosphere becoming more focused and shallower with the increasing dominance of the oceanic rift stress field as the
rift propagates northwards allows intrusions that produce magnetic anomalies oblique to those within thinned continental crust. Anomalies that straddle the
oceanward edge of continental crust must be sourced within the mantle. (d) Focusing of melt, typical of a mid-ocean ridge, creates oceanic crust that thickens
as the seafloor spreading zone moves away from the cool continental lithosphere.
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2003 RAS, GJI, 154, 706–730
728
S. M. Russell and R. B. Whitmarsh
are isolated bodies that lie ∼5 km below the top of the basement.
Further oceanward, adjacent to the peridotite ridge, the anomalies
have an ∼010◦ trend and sources exist both near the top of the basement and at ∼6–8 km depth. These anomalies can be modelled by
seafloor spreading, at least since the time of anomaly M3 (124 Ma),
and locally since M4(N) or M5(R), at rates of ∼10–14 mm yr−1 .
Seafloor spreading before the time of M5(R) (128 Ma) is unlikely.
(3) West of the peridotite ridge, magnetic anomalies and basement relief are strongly lineated in an ∼015◦ direction. The anomalies are consistent with seafloor spreading and the velocity structure
is similarly characteristic of oceanic crust. Magnetic sources lie at
the top of the basement, probably corresponding to thin dykes or
the edges of sills or lava flows near the top of the crust.
(4) We explain the above observations by a model (Figs 15 and
16), partly supported by numerical modelling of melt production, in
which syn-extensional magmatism increases across the whole margin from the last stages of continental rifting and breakup, through
the exhumation of subcontinental mantle to the onset of steady-state
seafloor spreading. We speculate that this change is brought about
largely by the increased focusing, and ascent, of a ridge of passively
upwelling asthenospheric mantle.
ACKNOWLEDGMENTS
SMR thanks the Natural Environment Research Council for a
Research Studentship and the Challenger Division, Southampton
Oceanography Centre for additional funds and facilities. SMR also
thanks Professor R.C. Searle for facilities provided in the Department of Geological Sciences, University of Durham. The Generic
Mapping Tool (GMT) software of P. Wessel and W. H. F. Smith
was used to construct several figures. We thank the two reviewers,
Tom Hildenbrand and Hans-Christian Larsen, for their extensive
comments which helped us to greatly improve this paper.
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