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Geophys. J . R. astr. SOC.(1986) 86,491-513 and Microfiche GJ86ll
Crustal structure of Atlantic fracture zones - 11. The
Vema fracture zone and transverse ridge
C. C.
R. S.
POttS and
White BuHard Laboratories, Department ofEarth
Sciences, Madingley Rise, Madingley Road, Cambridge CB3 OEZ
K. E.
Louden Department of Oceanography, Dalhousie University, Halifax, Nova
Scotia B3H 4J1, Canada
Accepted 1986January 29. Received 1986January 17; in original form 1985 September 10
Summary. The crustal structure beneath the Vema fracture zone and its
flanking transverse ridge was determined from seismic refraction profiles
along the fracture zone valley and across the ridge. Relatively normal oceanic
crust, but with an upwarped seismic Moho, was found under the transverse
ridge. We suggest that the transverse ridge represents a portion of tectonically
uplifted crust without a major root or zone of serpentinite diapirism beneath
it. A region of anomalous crust associated with the fracture zone itself
extends about 20 km to either side of the central fault, gradually decreasing
in thickness as the fracture zone is approached. There is evidence to suggest
that the thinnest crust is found beneath the edges of the 20 km wide fracture
zone valley. Under the fracture zone valley the crust is generally thinner than
normal oceanic crust and is also highly anomalous in its velocity structure.
Seismic layer 3 is absent, and the seismic velocities are lower than normal.
The absence of layer 3 indicates that normal magmatic accretionary processes
are considerably modified in the vicinity of the transform fault. The low
velocities are probably caused by the accumulation of rubble and talus and
by the extensive faulting and fracturing associated with the transform fault.
This same fracturing allows water to penetrate through the crust, and the
apparently somewhat thicker crust beneath the central part of the fracture
zone valley may be explained by the resultant serpentinization having
depressed the seismic Moho below its original depth.
Key words: explosion seismology, fracture zones, ocean bottom seismometers, oceanic crust, seismic ray-tracing
1 Introduction
Fracture zones are a prominent feature of the seafloor of ocean basins. They are
characterized by a ridge-and-trough topography, typically 10-20 km wide, which can some-
492
C. G. Potts, R. S. WhiteandK. E. Louden
times be followed for thousands of kilometres. The transverse ridges can be particularly
spectacular features, and are often associated with large fracture zones (e.g. Romanche,
Vema and S t Paul in the equatorial Atlantic). In addition to forming prominent bathymetric
features, fracture zones are the major source of crustal seismic heterogeneity (Detrick &
Purdy 1980; White et al. 1984).
Francheteau et al. (1976) conclude that the high relief across large Atlantic fracture zones
does not expose deep sections into the oceanic crust, but is generated by many small-throw
faults. Submersible studies in the Kane fracture zone (Karson & Dick 1983) and the Oceanographer fracture zone (OTTER 1984, 1986) confirm the common occurrence of gabbroic
and ultramafic rocks in the valley walls, which had led to the proposal that the crust beneath
large fracture zones must be thinner than normal crust (Fox, Detrick & Purdy 1980; Fox &
Gallo 1984).
Several seismic refraction experiments have been shot in the Atlantic to determine the
crustal structure in and around transform faults (White et al. 1984; Detrick, White & Purdy
1986). The crust beneath fracture zone valleys is typically thinner than normal oceanic
crust, with low crustal velocities and a marked absence of the oceanic seismic layer 3. Little
is known about the seismic or geological structure of the transverse ridge, and few
experiments have provided good constraints on the width of the anomalous crustal region
associated with fracture zones.
'The seismic refraction experiment described here provides seismic control over the crustal
and subcrustal structure along the Vema fracture zone and across its associated transverse
ridge. The most important results are the constraints on crustal structure beneath the transverse ridge to the south of the Vema fracture zone, and the nature and width of the
anomalous crust below the fracture zone itself.
2 The Verna fracture zone
The Vema transform (Fig, 1) offsets the Mid-Atlantic Ridge by 320 km with an age contrast
of 26 Myr across the end of the spreading centre. Currently the plates on either side of the
fault are moving with a relative velocity of 24 mm yr-' . The transform valley is typically
25-40 km wide between the crests of the north and south walls. In places the top of the
southern ridge is only 550 m below sea-level, and may recently have been at sea-level
(Bonatti & Crane 1982).
Dredge samples indicate that the north wall of the Vema, which has a relatively subdued
topography, is probably a disrupted section through an apparently representative segment of
normal oceanic crust; the southern transverse ridge is interpreted from the samples obtained
as being an uplifted section of crust (Bonatti & Honnorez 1976; Bonatti & Crane 1982;
Honnorez, Mevel & Montigny 1984). Material believed to be from oceanic layer 3 on the
lowest part of the inward facing wall of the transverse ridge appears to have been reworked
and altered, in situ.
Gravity and magnetic intensity profiles across the active transform of the Vema at around
44"W and 43OW have been modelled by Robb & Kane (1975). They propose a structure
which includes dense, ultrarnafic bodies at shallow depths below the walls of the fracture
zone. This solution has been challenged by other authors (e.g. Sibuet & Veyrat-Peinet 1980;
Louden & Forsyth 1982) but in the absence of other evidence, no more credible structure
could be proposed.
Ludwig & Rabinowitz (1980) reported the results of the first seismic refraction experiments in the Vema using free-floating sonobuoys (Fig. 2). No detailed geological
Vema fracture zone and transverse ridge
493
20'N
10 N
0"
10ys
.
200s
60'W
SOUTH AMERICA
50'W
40"W
30°W
20aw
1oow
Figure 1. The location of the Vema fracture zone and other major features of the Mid-Atlantic Ridge
between Africa, and South America. The seismic refraction experiment was near the western ridgetransform intersection.
interpretation was presented, but it was concluded that the seismic structure along the Vema
fracture zone is not the same as that below apparently normal oceanic crust nearby, and is
not easily reconciled with the velocity structure of deep-ocean crust observed elsewhere.
120
11'30
1'
0'30
00
Figure 2. The location of the shooting tracks from this experiment superimposed on the simplified bathymetry of Prince & Forsyth (1986). Lines A, B and C are described in this paper. Also shown by broken
lines are Detrick et al. (1982) OBS line, the six sonobuoy experiments (SLF3-6 and R45, R48) of Ludwig
& Rabinowitz (1980), and Line E2, the OBS refraction line along the spreading centre of the Mid-Atlantic
Ridge shot o n the same cruise (White et a/. 1986).
494
C. G. Potts, R. S. White and K. E. Louden
Velocity (km s-'1
2
4
6
8
10
2
4
n
E
Y
v
6
c
+Ir
P
10
12
14
Figure 3. Velocity-depth functions determined for the Vema fracture zone by Detrick et QI. (1982)
superimposed on the bounds for normal ocean crust greater than 3 M y r (from White 1984). Tau-p
solutions from two OBSs are shown, with a plane layer solution from one instrument. Vertical scale is
depth into igneous crust. (See Detrick ef al. 1982 for location of shooting track and receivers.)
In 1978 a single explosive refraction line (Fig. 2) was fired into three OBSs (Detrick et al.
1982). The data were topographically corrected using sediment thicknesses estimated from
single channel reflection profiles. Detrick et al. conclude that the crust beneath the fracture
zone has lower seismic velocities than normal oceanic crust, and is probably thinner. From
the velocity-depth functions (Fig. 3 ) it is also apparent that a typical 'layer-3' type region
(with seismic velocity of approximately 6.7 km s-', and low velocity gradients) is missing,
or extensively altered. The tau-p solutions from this experiment were used as a starting
point for interpreting the data from the OBS refraction experiment described in this paper.
3 The seismic experiment
The experiment discussed in this paper comprised three wide-angle seismic lines, A, B and
C, located on a regional map in Fig. 2 and in detail in Fig. 4. They were shot from RRS
Vema fracture zone and transverse ridge
44'30'w
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43'30'W
44'W
I
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Shot
LINE B
l k n t 7n
~
.
10'30 N
Figure 4. Detailed layout of the seismic experiment superimposed o n a simplified bathymetry. Cambridge
OBSs are prefixed by 'C', Bedford Institute of Oceanography OBSs are prefixed by 'R'. Shot locations are
small dots. Shot sizes were alternated along the lines to provide short range data for shallow structure and
long-range arrivals for deep structure. Also marked by a heavy Line is the location of the airgun reflection
profile along line B (Fig. 5 ) . A complete table of shot locations, sizes and detonation times is included in
Fig. C of the microfiche appendix (GJ86/1) .
Shackleton (leg 9/82). A fourth line along the spreading centre (labelled MAR on Fig. 2) is
described elsewhere (Louden &White 1984; White, Louden & Potts 1986).
Nine OBSs were laid. Four were Cambridge instruments (prefix 'C' on Fig. 4), used with a
hydrophone and a three-component deployed geophone package (Duschenes, Potts &
Rayner 1986). Five were Bedford Institute of Oceanography (BIO) instruments (prefix 'R'
on Fig. 4) containing two-component geophones inside the main pressure vessel (Heffler
& Barrett 1979).
Three of the Cambridge OBSs recorded good quality data. The fourth, C5, recorded only a
single channel of seismic data and was not used in the analysis. One Bedford Institute of
Oceanography OBS (R14) failed t o release from the seafloor and was lost. Clock drifts and
other details for all the instruments are listed in Fig. D of the Appendix (see Microfiche
GJ 8611).
A single-channel reflection profile, using a 2.6 1 (160 in3) airgun, was recorded along the
shooting track in the fracture zone to provide sediment thickness control. This profile shows
clearly the rugged nature of the igneous basement (Fig. 5).
Shot depths and times were determined from the bubble-pulse periods (Spudich & Orcutt
1980a, b) and are listed in the Appendix (Fig. C). Accuracy is estimated as 21.5 ni in depth
and 20.01 s in time.
Ranges from shots t o each receiver were calculated by ray-tracing both direct and
multiple water paths through a model of the water sound velocity versus depth structure
(illustrated in the Appendix, Fig. B). A consistent set of shot positions (listed in the
Appendix, Fig. C) and receiver locations (Appendix, Fig. D), was then derived by integrating
the calculated ranges with the ship's position derived from satellite fixes and log and heading
information. The absolute position of each receiver is known to only +200 m, although the
relative ranges between shots into any given receiver are considerably better defined than
this.
Figs 6-8 show three typical record sections. Other record sections used in this analysis
are illustrated in Figs E to Z of the microfiche Appendix. The record sections shown are not
496
Figure 5. Single channel 2.6 1 (160 in3) airgun reflection profile recorded along shooting track B. The
three panels join to form a continuous east-west profile (see Fig. 4 for location of profile). The rugged
basement topography is easily identified, except beneath the intersection deep.
Vema fracture zone and transverse ridge
Shotnumber
r l
7475
497
OBS 1 LINE B
70
0001
0304
0607
0990
9293
9596
9899
101
8
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E
,w
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80
60
40
20
0
Range (Km)
Figure6. Record section from OBS C1, line B, vertical (Z) geophone (low gain). Bandpass filtered
between 2.0 and 20.0 Hz. Each trace has equal maximum plotting amplitude. First arrivals can be picked
out to 80 k m range.
sediment or water path corrected, and the effect of the rugged basement topography along
the lines can be seen on Fig. 7, and more dramatically on Fig. 8, which is a section from line
A across the transverse ridge.
The first break on each hydrophone or vertical geophone trace is interpreted as the
refracted P-wave arrival. N o S-waves were identified, except for P to S conversions at the
basement just below some of the instruments. Picking accuracy varied with instrument,
Shotnumber
? 1 72 73 74 75 76
40
OBS 12 LINE 6
78 79 8 0 8 1 82 8384 85 86878889
102 103
90 91 9293 94 95Bg 97 9899 lOOIOli
I 104
20
Figure 7. Unfiltered record section from OBS R12, line B, hydrophone. Note ‘ringy’ character of seismograms which obscure amplitudes and second arrival data. Each trace has equal maximum plotth&
amplitude.
498
C G. Potts, R. S. White and K. E. Louden
OBS 15 LINE A
Shotnumber
Range (Km)
Figure 8. Unfiltered record section from OBS R 1 5 , line A, hydrophone. Each trace has equal maximum
plotting amplitude. Note the effect of the transverse ridge topography. Data quality from shots over the
ridge was consistently good.
range and shot size, and so estimated errors were assigned to each pick, and used throughout
the interpretation.
Almost all the seismograms had a ‘ringy’ character. Fig. 7 shows an example of this, on a
section recorded by the hydrophone of OBS R12 on line B, along the fracture zone. This
phenomenon is often observed in OBS data, and there is some evidence that it may be
attributable to resonance of the instruments themselves (Loncarevic 1981) or is a result of
reverberations in shallow layers. It has, however, been observed particularly in experiments
in fracture zones using OBSs of different designs, and also surface receivers such as sonobuoys (e.g Sinha & Louden 1983) and multichannel hydrophone streamers (Potts, Calvert &
White 1986). The ringy character may thus be a feature of fracture zone structure, caused by
dispersion, attenuation and reverberation of the seismic signals in the highly fractured crust
which is inevitably present, especially in the active transform region. Whatever the detailed
cause, the effect is important, since in the Vema experiment the ringing distorted the
amplitudes of early arrivals and obscured second arrivals. It proved impossible to model the
amplitudes of the data sensibly.
4 Forward modelling using seismic ray-tracing
4.1
LINE B, ALONG THE FRACTURE ZONE
Line B, which consisted of 35 shots and five instruments along the fracture zone, was raytraced first. The final model provided a starting point for ray-tracing line A which crossed
the transverse ridge. Ray-tracing was performed using S E I S I 1 , with models incorporating
vertically and horizontally varying seismic structures (Cerveny, Molotkov & PZen&‘k 1979).
Fig. 9 shows the velocity-depth structure of the starting model for line B, derived from
the results of Detrick et al. (1982). The structure was applied initially as a uniform lateral
model and then modified to produce the final solution. Ray-tracing was stopped when it
become apparent that the only method remaining to improve the fit of the calculated arrival
Vemafracture zone and transverse ridge
499
Velocity (Km/s)
Figure 9. The starting structure for forward modelling by ray-tracing along the fracture zone, line B. The
structure was derived directly from the tau-p results of Detrick et al. (1982) (Fig. 3). Vertical scale is
depth into igneous crust.
times to the observed data was artificially to alter sediment thicknesses or local upper crustal
structure beneath individual shot-points.
Fig. 10 shows all the rays generating arrivals into the five OBSs along line B with the
associated calculated and observed travel times for three of the instruments. Rays traced into
individual receivers can be identified more clearly on the Appendix (Figs CC and EE), with
corresponding large-scale plots of the observed and calculated travel times in Figs DD and
FF.
Water depths were picked to *I0 m from the echo-sounder profile recorded while firing.
Sediment thicknesses were determined from the airgun reflection line, shot approximately
along the shooting track (Figs 4 and 5). The inaccuracy of control over sediment thicknesses
from the reflection profile is the major source of uncertainty along the whole of line B, and
dictates that the minimum time-error is k0.075 s for all arrivals, however sharp their
apparent onset may be. Weaker arrivals were sometimes assigned larger picking errors.
4.2
COMMENTS ON THE FINAL MODEL FOR LINE B
An isovelocity plot of the final model is shown in Fig. 1 1. The major features are:
(a) Crustal structure. Crustal velocities along the fracture zone are low compared with
normal oceanic crust of the same age. Fig. 12 shows the velocity-depth functions from
three sites along line B plotted against the bounds for normal oceanic crust greater than 3
Myr (from White 1984). Along nearly 60 km of the line, from OBS R13 to within 10 km of
R1 1 , the crustal structure is fairly constant. At the western end of the line the crustal
structure changes relatively abruptly, between OBS R11 and C1. The velocity-depth
function (number 3 ) between the instruments indicates that a different velocity profile has
emerged.
(b) Moho depth and subcrustal structure. The seismic Moho beneath the fracture zone is
marked by a large velocity increase and along most of the .line is shallower than under
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Figure 10. (Top) Calculated arrival times (dots) generated by ray-tracing through the final model along
line B, for OBS C 2 , R12 and C1. Shown for comparison are the observed arrivals (rectangles). The
dimensions of the rectangles are estimates of the time (vertical) and distance (horizontal) errors o n the
data. The distance uncertainty is assumed t o be + 2 0 0 m, the estimated uncertainty of the receiver
location..The location of each OBS is marked with a solid triangle. (Bottom) All the raypaths constraining
the final model for line B, along the fracture zone. Similar plots for OBS R11 and R 1 3 are shown in Figs
EE and FF of the Appendix.
501
Vema fracture zone and transverse ridge
0
Sea level
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WEST
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T
I
10
20
30
40
50
Distance (Km)
60
70
80
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EAST
Figure 11. Isovelocity contour plot of final model for line B, d o n g the fracture zone. The structure is
constrained by the raypaths shown in Pig. 10. Contour interval is 0.2 km SKI.The numbered vertical
lines indicate the locations of the velocity-depth functions shown in Fig. 12.
normal crust of the Same age. Sub-Moho velocities start a t approximately 8.0 km s-l. The
crust is thinnest (4.3 km) below the intersection deep (between OBSs C2 and R13), but
remains thin for approximately 50 km westwards along the fracture zone. The thickening of
the crust at the ends of the line is poorly constrained, compared with the control over Moho
depth from the dense network of intersecting raypaths in the middle of the section (see Fig.
10). Thickening at the eastern end of the line, below the edge of the intersection deep is
accompanied by lowering of the seismic velocities throughout the crust. However, at the
western end of the line the velocities near the top of the crust increase as the seismic Moho
deepens.
Crusta: structure along the Vema fracture zone is similar to Sinha & Louden’s (1983)
‘type A’, since it exhibits low seismic velocities and is moderately thin, and to the crust
found by Whitmarsh & Calvert (1986) in the Charlie-Gibbs fracture zone.
4.3
L I N E A , ACROSS THE T R A N S V E R S E R I D G E
Line A, consisting of 21 shots and three OBS across the transverse ridge, was ray-traced from
three different starting points.
The structure of line B at its intersection with line A was the first starting model tried
(Fig. 13, profile number 1). After 1 0 iterations, it became apparent that n o realistic adjustment of this model would reproduce the observed arrival times.
Even though this attempt to model line A was unsuccessful, two features of the modelling
were useful in establishing a new starting structure:
(a) The shallow crustal structure over the ridge was definitely not the same as the shallow
structure at the fracture zone line intersection.
(b) To reproduce the observed arrivals, a region of high seismic velocity with low velocity
gradients must be present under the ridge, 2-3 km beneath the seafloor.
Layer 3 exhibits consistent high velocities (6.7-7.2 km s - ’ ) and low gradients (less than
0.15 s-’) (Spudich & Orcutt 1980b, White 1984), and its introduction into the structure
502
C. G. Potts, R. S. White and K . E. Louden
Velocity (km s-’1
2
4
6
8
1
2
4
n
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Y
6
Y
r
-P a
Q)
n
10
12
14
Figure 12. Velocity-depth profiles below three sites on line B, from the final model after ray-tracing (see
Fig. l l ) , superimposed on the velocity bounds for normal crust greater than 3 Myr (from White 1984).
‘M’ marks the seismic Moho interface.
beneath the ridge is sensible, considering point (b) above. The inclusion of an intermediate
layer allowed more flexibility in the ray-tracing modelling, and the first reasonable fit of
calculated arrivals t o the observed data was achieved. Fig. 13 shows the ray-paths through
the model, called ‘solution one’, with the calculated arrival times, and Fig. 15(a) is an
isovelocity contour plot.
Due to the rough topography over the ridge, water depths from the wide-angle echosounder cannot be determined better than +I00 m. This inaccuracy, combined with the lack
of off-line bathymetry, was the largest consistent source of error, and causes a minimum
uncertainty in the travel times of 20.1 s (rectangular symbols on Fig. 13). There was no
reflection profile along line A, so sediment thicknesses in the fracture zone valley at the
north end of the line were extrapolated from the line B data and the isopach map of Prince
& Forsyth (1986).
Only long-range shots into OBS R13 (> 15 km) were used. This was because R13 did not
lie exactly on the shooting track for line A, but was approximately 4.6 km to the east, and
at short ranges the deviation from the plane of the experiment becomes more serious. The
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Figure 13. The first model, ‘solution one’, for line A, across the transverse ridge, assuming a fixed velocity
structure at the line intersection, and a layer-3 region beneath the ridge. Shown (top) are the arrivals
generated by the model (dots) compared with the observed values (rectangles). The dimensions of the
rectangles are determined by estimates of the errors o n the observed data. Bottom are all the ray-paths
constraining the model into the three OBSs.
C.G. Potts, R. S. White and K . E. Louden
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A R R I V A L TIMES L I N E A (SOLUTION T W O )
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Figure 14, The second model, ‘solution two’, for line A, across the transverse ridge, with n o imposed
velocity structure. Top are the observed (rectangles) and computed (dots) travel times, and bottom are
the rays constraining the structure. The dimensions of the rectangles are determined by estimates of
the errors on the observed data.
Vema fracture zone and transverse ridge
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Figure 15. Isovelocity contour plots of the two solutions for line A. Solution one has very thin crust
beneath the abrupt ‘transition’ zone, whereas solution two allows the normal crust to alter more
gradually, as layer 3 thins out. Contour interval 0.2 k m s - l .
fracture zone structure from line B at the line intersection was kept fixed throughout the
modelling. To maintain this fixed structure, and fit the observed arrival times, a ‘transition’
zone evolved beneath the edge of the transform valley in which crustal layers thinned to
only 2.5 km and seismic velocities were reduced.
The nature of the transition zone was dictated by the choice of the line B structure as a
constraint at the line intersection. To test the sensitivity of the modelling process, this
C. G. Potts, R. S. White and K . E. Louden
506
constraint was removed. Once again, a reasonable fit to the observed data was obtained
quickly after adjustment of the structure (Fig. 14). The isovelocity contour plot of this
model, called 'solution two', is shown in Fig. 15(b).
4.4
COMMENTS O N T H E F I N A L M O D E L S FOR L I N E A
Solution one (Fig. 15a) is the best constrained structure modelled for line A since it
incorporates the structural tie with line B. The main difference between the two final
solutions is the inclusion in solution one of the transition zone.
Fig. 16 shows velocity-depth functions from selected sites along both models.
Summarized below are observations on the nature of the two structures:
(a) Both models indicate that a layer-3 type region exists beneath the transverse ridge and
that, as the fracture zone is approached, this region thins and alters (Fig. 15). Both models
LINE A Solution One
Velocity ( k m 6-9
Velocity (km 9.')
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Velocity
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Figure 16. Velocity-depth profiles from the two solutions to line A , across the transverse ridge. Shown,
top, are profiles from solution one, and bottom, from solution two. The common trends observed are
crustal thinning, disappearance of layer 3 and lowering of crustal velocities. The main difference is the
inclusion of very thin crust in solution one beneath the southern flank of the fracture zone valley.
Vema fracture zone and transverse ridge
507
also suggest that the crust beneath the transverse ridge is normal in character, with velocitydepth curves falling within the bounds for normal ocean crust (Fig. 15).
(b) There is evidence in both models that crustal thinning commences u i e r 20 km away
froin the fracture zone: in solution one, thinning from approximately 4 t o 2 km; in solution
two, thinning less dramatically from approximately 5 to 3.5 km. This overall crustal thickness change is superimposed on the disappearance of layer 3, which is vertically thinned and
laterally altered in solution one, or simply thinned out in solution two. I n both cases the net
crustal velocities decrease while velocity gradients increase.
(c) In solution one the thinnest crust, approximately 2 km thick, lies beneath the flanks
of the fracture zone, with a thickening of the seismic crust beneath the middle of the
fracture zone valley. Solution two is more homogeneous, including thin, but not
exceptionally so, crust.
The results of forward modelling using real data are not unique, and the final models are
influenced by the initial assumptions of the starting structure. The two models presented
here satisfy the observed arrival times, and are both plausible geological structures.
Three-dimensional gravity modelling over the Vema fracture zone conducted independently by Prince & Forsyth (1986) produced a structure which indicated very thin crust
beneath both flanks of the fracture zone with rather thicker crust under the central part of
the valley. This result is in general agreement with the structure under the fracture zone
valley suggested by our solution one. Prince & Forsyth’s results also, however, suggest the
presence of a thick crustal root beneath the transverse ridge on the south wall of the fracture
zone which we definitely do not find. Since the gravity dataset is rather limited in this region
the discrepancy between seismic and gravity interpretations of the transverse ridge may be
explained by poor gravity data coverage.
5 Discussion
5.1
F R A C T U R E ZONE T R A N S V E R S E RIDGES
Line A, across the Vema fracture zone, is the first detailed seismic refraction experiment t o
investigate directly the structure of a large transverse ridge. Both solutions modelled (Fig.
15) indicate that the ridge is a section of oceanic crust which has been uplifted and warped
by an unknown process. The oceanic crust, however, was itself originally formed in the
vicinity of a ridge-transform intersection. Thus the crustal deformation is superimposed on
the lateral variations in crustal structure due to the proximity of the fracture zone. In both
models the crust under the ridge contains a region analogous to layer 3, albeit with somewhat higher velocity gradients (0.2-0.25 s-‘) than are normally observed (0.15 s - l , Spudich
& Orcutt 1980b), and a recognizable upper crustal layer 2. Overall crustal thickness
decreases northwards and layer 3 disappears, as the structure changes from nearly normal to
fracture zone ‘type A’ crust (as defined by Sinha & Louden 1983). Our model of the
structure across the transverse ridge supports the earlier conclusion from rock sampling that
the ridge consists of a section of uplifted oceanic crust formed near a fracture zone
(Honnorez et al. 1984; Bonatti 1978).
Two main mechanisms to explain crustal uplift of fracture zone transverse ridges have
been proposed. The first is that the extensive fracturing in the region of an active transform
allows water to penetrate through the crust to the upper mantle, where serpentinization
occurs (Francis 1981). Serpentinization is the reaction between water and olivine below
5OO0C, which results, amongst other things, in a rock volume increase. The low density
serpentinite is buoyant, rises diapirically under the flanks of the fracture zone valley, and
508
C. G. Potts, R . S. White and K . E. Louden
uplifts the crust. Serpentinized upper mantle material exhibits reduced seismic velocities
(Coleman 1971). The second suggested mechanism (Bonatti 1978) is that compressional
stresses caused by the re-orientation of the transform fault following changes in spreading
direction produce vertical tectonic movement of blocks of crustal material adjacent to the
fracture zone. Bonatti & Hamlyn (1 978), for example, attribute the considerable vertical
movements suggested by the chemistry of rocks dredged from the transverse ridge of the
Owen fracture zone to mantle upwelling below the spreading ridge at the time of crustal
formation; the rocks are later exposed tectonically by uplift along the fracture zone. Large
transverse ridges are often found where slow-spreading mid-ocean ridges change direction,
as for example near the Vema and Romanche in the central Atlantic, the Charlie-Gibbs in
the North Atlantic and along the south-west Indian Ridge in the south Indian Ocean. At
such locations tectonic compression is likely to occur.
The seismic results reported here cannot, alone, discriminate between these two
mechanisms. Both would result in the upwarped crustal section which is observed. If
extensive serpentinization has occurred, a large body of low velocity, low density material
may be emplaced under the ridge; unfortunately, this could prove to have a similar seismic
character to oceanic layer 3 (Francis 1981). The simplicity of the final crustal cross-section
produced by modelling the seismic data, however, betrays no trace of the structural disruption which extensive intrusion of serpentinite must cause. The only aspect of the crustal
model that might indicate the presence of a large serpentinized body is the anomalously high
velocity gradient found in layer 3. This value could, however, be explained as being symptomatic of a change in the process of formation of the crust in the vicinity of fracture zones
(White e t al. 1984). The suggestion that the uplift was due to tectonic compression in the
Vema is considered preferable.
5.2
C R U S T A L S T R U C T U R E BENEATH THE F R A C T U R E ZONE V A L L E Y
Fig. 11 shows an isovelocity contour plot of the final model along line B, with velocitydepth profiles for three sites shown on Fig. 12. Crustal velocities are low compared with
normal oceanic crust of the same age, and the crust along most of the line is significantly
thinner than normal crust (4-4.5 km compared with 5-8 km). There is an absence of a
layer-3 type region along most of the line. The observed sub-Moho velocities (7.9-8.0
km s-l) and gradients (less than 0.1 s-') are not unusual, compared to normally observed
values.
Low crustal velocities and thinner than normal crust are commonly observed in seismic
refraction surveys over fracture zones (e.g. Detrick et al. 1982; Cormier, Detrick & Purdy
1984; Sinha & Louden 1983; White et al. 1984; Whitmarsh & Calvert 1986). The results
from the Vema confirm these observations. The low crustal velocities are attributed mainly
to the intense fracturing and brecciation caused by faulting in the active transform. Bowen &
White (1986) observed faulting around the intersection deep of the Vema fracture zone,
which would also produce a final phase of disruption of the crust as it passed into the
inactive domain. Karson & Dick (1983), reported significant thicknesses of basaltic rubble in
the Kane fracture zone, presumably debris from the fracture zone walls. Similar rubble was
observed during submarine investigations of the Oceanographer transform valley (OTTER
1984, 1986) and also encountered during drilling in the Vema by DSDP (Perch-Nielson et al.
1977). The accumulation of' this low velocity material on top of fractured and faulted crust
would further reduce the apparent velocity of the crust.
The almost universal absence of a layer-3 type region in fracture zone crust is harder to
explain. Even after 60-77 Myr, in the Tydeman fracture zone (Calvert & Potts 1985), no
Vemofracture zone and transverse ridge
509
typical layer 3 is observed. Its absence, similarly, from most young sections of fracture zones
studied (e.g. Vema, Charlie-Gibbs, Oceanographer) indicates that the effect must be a
primary feature of crust formed in, or near, fracture zones.
One of the most interesting features of the results from line A, across the fracture zone
arid transverse ridge, is the possibility, from the correspondence between the seismic solution
one (Fig. 15a) and the gravity profile of Prince & Forsyth (1986), that the thinnest crust, or
shallowest seismic Moho, occurs beneath the flanks of the fracture zone valley. Results from
line MAR, along the spreading centre, support this hypothesis (Louden &White 1984; White
et aE. 1986). Other authors (e.g. Sinha & Louden 1983; Cormier et al. 1984) have reported
limited regions of apparently very thin crust along fracture zones (i.e. Sinha & Louden 1983,
type ‘B’ crust). It is conceivable that such results represent sampling by the seismic experiments of the flanking thin crust, or excursions of thin crust across the fracture zone valley,
or both. It is not yet possible t o decide whether these crustal variations represent real shortterm changes in the ridge-transform crustal accretion processes, or later variations in hydrothermal alteration of the crust and mantle or, again, both.
It can no longer be doubted, however, that fracture zones create significant lateral
variability on the scale of a few kilometres as well as the regional effects, discussed below in
Section 5.3, on a scale of a few tens of kilometres.
5.3
THE W I D T H O F A N O M A L O U S F R A C T U R E Z O N E C R U S T
Estimates of the full width of’ fracture zones from geophysical methods are scattered
between 7 k m (White & Matthews 1980) and 30-40 k m (White et al. 1984). The dataset is
limited, and as yet n o obvious correlation has been found between the apparent width of
anomalous crust and the age, offset, spreading rate or any other variable feature of a fracture
zone. The estimate of 10-20 km total width for the low velocity crustal and subcrustal slab
from the Tydeman fracture zone (Potts et al. 1986) is in agreement with the results reported
by several other authors (e.g. Sinha & Louden 1983; Cormier et al. 1984). The width of
anomalously thin crust under the Blake-Spur fracture zone, determined by reflection
profiling across an old section (135 Myr), is approximately I 5 km (Mutter et al. 1984).
Several authors (e.g. White et al. 1984; Sinha & Louden 1983; Corniier e f al. 1984;
Ambos 1984) suggest that anomalous crust along fracture zones is a combination of (i> largescale gradual crustal thinning, (ii) smaller scale lowering of crustal velocity, in conjunction
with more marked thinning of the crust (i.e. Sinha & Louden’s 1983, type A crust) and
(iii) occasional sections of very thin crust (i.e. Sinha & Louden’s 1983, type B crust). The
results from the Vema strongly support these proposals; the structure modelled under the
ridge shows that crustal thinning is effective u p t o 20 km away from the axis of the
transform, and low crustal velocities become important approximately 1 0 kni from the axis.
This indicates that the overall width of affected crust could be 40 kni (crustal thinning) but
the highly anomalous region may overall be only 20 km wide or less.
Surveys which d o not sample oceanic crustal structure over significant distances (>20
km) clear of a fracture zone may detect only the region of low velocity crust, and not
resolve the larger-scale variation in crustal thickness.
6 Conclusions
(a) The crust beneath the flanking transverse ridge of the Vema fracture zone is essentially
normal in its seismic velocity structure and its overall thickness. The seismic Moho is upwarped under the transverse ridge. No major zone of serpentinite diapirism was detected
510
C.G. Ports, R. S. White and K . E, Louden
under the ridge. This suggests that the ridge represents a relatively normal section of oceanic
crust uplifted by tectonic forces associated with the Vema transform, rather than by
diapirism.
(b) Gradual crustal thinning occurs over a region of up to 20 km on either side of the
transform fault.
(c) The thinnest crust is found under the flanks of the transform valley, and near the
ridge-transform intersection. The crust appears to be somewhat thicker (though still thinner
than normal oceanic crust) beneath the central part of the valley. This effect could be
attributed to depression of the seismic Moho due to serpentinization by water penetrating
down through the crust in the extensively faulted region along the middle of the transform
valley (i.e. in the vicinity of the transform fault).
(d) Under the transform valley the crust is not only somewhat thinner than normal
oceanic crust, but is also highly anomalous in its velocity structure. It exhibits a marked
absence of a seismic layer 3 , which is one of the most characteristic features of normal
oceanic crust, and generally has rather low seismic velocities. The absence of seismic layer 3
indicates that accretionary magmatic processes in the transform are far from normal, whilst
the low seismic velocities are attributed to the extensive faulting and fracturing in the crust
of the transform zone, together with the occurrence of talus and rubble layers in the upper
section.
Acknowledgments
The success of the field-work at sea was due in large part to the support and hard work of
our colleagues on both sides of the Atlantic, the technical staff from Research Vessel
Services of the NERC and the officers and crew of RRS Shackleton. We thank them all.
Beverley Smith processed the words and Phyl Fisher drafted most of the diagrams. The
Atlantic Geoscience Center of the Bedford Institute of Oceanography kindly loaned six of
their ocean bottom seismometers for use on this experiment and we are especially indebted
to Brian Nichols for his efforts and indispensable contributions whilst digitizing the BIO
OBS data, and to Don Locke for running the BIO instruments at sea. Roger Prince and Don
Forsyth provided bathymetric information which we gratefully acknowledge. Funding was
from the Natural Environment Research Council under grants GR3/4157 and GR3/4403 to
RSW and from the National Science and Engineering Research Council grant A8459 to KEL.
Department of Earth Sciences, Cambridge, contribution number ES 638.
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Appendix on microfiche
A microfiche accompanies this paper (GJ 86/1). It contains copies of the seismic record
sections with normalized trace amplitudes for each ocean bottom seismometer from each of
the three shooting lines used in the interpretations reported here. Tabulations of the shot
and receiver data are included, as are enlarged versions of the ray-tracing fits to the line B
arrivals. The contents of the Appendix on microfiche are as follows.
A. Map showing positions of all shots and receivers. Cambridge OBSs are prefixed by ‘C’,
Bedford Institute of Oceanography OBSs are prefixed by ‘R’. Shot locations are small dots.
B. Water sound-velocity structure in the Vema fracture zone, used in range calculations.
Structure down to 2500 m (shown by solid line) is from direct measurement using a soundvelocimeter lowered on a conducting cable. Structure below 2500 m (broken line) is from
existing nearby oceanographic measurements.
C. Shot data showing shot size, depth, time, latitude, longitude, and seafloor depth.
D. Receiver data showing latitude, longitude, water depth and clock drift.
E. Record section from line A, receiver C1, horizontal geophone.
F . Record section from line A, receiver C2, vertical geophone.
G. Record section from line A, receiver C4, vertical geophone.
H. Record section from line A, receiver C4, horizontal geophone.
I. Record section from line A, receiver R13, hydrophone.
J. Record section from line A, receiver R15, hydrophone.
K. Record section from line B, receiver C1, vertical geophone.
L. Record section from line B, receiver C1, vertical geophone (charge scaled).
M. Record section from line B, receiver C l , hydrophone.
N. Record section from line B, receiver C2, hydrophone.
0. Record section from line B, receiver C4, vertical geophone.
P. Record section from line B, receiver R11, vertical geophone.
Q. Record section from line B, receiver R12, hydrophone.
R. Record section from line B, receiver R13, vertical geophone.
S. Record section from line B, receiver R15, horizontal geophone.
T. Record section from line C, receiver C1, hydrophone.
U. Record section from line C, receiver C2, vertical geophone.
V. Record section from line C, receiver C4, vertical geophone.
W. Record section from line C, receiver R l l , horizontal geophone.
X. Record section from line C, receiver R12, hydrophone.
Y. Record section from line C, receiver R13, horizontal geophone.
Verna fracture zone and transverse ridge
513
Z. Record section from line C, receiver R I S , horizontal geophone.
AA. True scale plot of seafloor topography and sediment depths across the transverse
ridge (line A) and along the fracture zone valley (line B).
BB. Last-square fits of straight lines to uncorrected first arrivals along line B. Vertical
scale is arrival time reduced at 8.0 km s-' .Horizontal scale is range in kilometres.
CC. Raypaths through the final model for line B, along the fracture zone, into OBS C1,
C2 and R12.
DD. Calculated arrival times (ctosses) generated by final model for line B shown in Fig.
CC. Boxes show observed arrival times. Solid circles are location of OBS C1, C2 and R12.
The dimensions of the rectangular boxes correspond to the estimated errors on the data.
EE. Raypaths through the final model for line B, along the fracture zone, into OBS RI 1
and R13.
FF. Calculated arrival times (crosses) generated by final model for line B shown in Fig.
EE. Boxes show observed travel times. The dimensions of the rectangles correspond to the
estimated errors on the data.