Geophys. J . Int. (1990) la,367-378
The seismic velocity structure of some NE Atlantic continental rise
sediments; a lithification index?
R. B. Whitmarsh, P. R. Miles and L. M. Pinheiro"
Institute of Oceanographic Sciences Deacon Laboratory, Brook Road, Wormley, Surrey GU8 SUB, VK
Accepted 1989 October 24. Received 1989 October 17; in original form 1989 June 14
SUMMARY
Two separate sets of experiments with digital ocean-bottom seismographs (DOBS)
and airguns, on continental rise areas off Madeira and west of Portugal, produced en
echelon second arrivals from the sediment layer on record sections. Traveltime and
synthetic seismogram modelling indicate that the arrivals represent multiplyreflected refracted phases which have undergone reflection within the sediment layer
itself. Further, although the P-wave contrast at the intrasediment reflecting horizon
is relatively small, the modelling indicates a large downward increase in S-wave
velocity from 100-250ms-' (Poisson's ratio of at least 0.42) to about 1200ms-'
(Poisson's ratio of about 0.25). A reflection event can usually be found on reflection
profiles along the refraction lines at almost exactly the same 'depth' as the
intrasediment reflector. In one case such an event can be traced to a nearby Deep
Sea Drilling Project (DSDP) borehole where it is associated with the transition from
ooze to chalk. This, and other circumstantial evidence, suggests that the intrasediment reflector marks an important increase in lithification within the sediment layer.
If so it means that, in future, straightforward OBS experiments may be used to
measure the depth of this increase without resorting to the drill.
Key words: lithification, ocean sediment, seismic reflector, seismic structure.
INTRODUCTION
The study of the seismic velocity structure of sediments in
the vicinity of the ocean floor has been an elusive and
difficult task for many years. Even with the advent of
logging in scientific drill holes the topmost 100 m or so is still
often inacessible, due to hole stability problems, and shear
wave measurements are impractical until the S-wave
velocity exceeds that of the drill fluid (Anon. 1986).
Alternatively seismic refraction methods can be used but,
unless both seismic sources and receivers are placed on or
very close to the sea-bed, the geometry of the situation is
often unfavourable. That is, the onset of arrivals refracted
within the sediments is masked either by arrivals from
acoustic basement or by strong water-borne phases. The
horizontal extent over which sediment arrivals may in fact
be seen is also curtailed by the often rather strong velocity
gradients in the near-bottom sediments. However, since
experiments with closely spaced and well-navigated bottom
shots colinear with an OBS are difficult and expensive,
and therefore rarely performed, data have mostly been
* On leave from Departmento de GeociEncias, Universidade de
Aveiro, Portugal.
collected using near-surface sources and ocean-bottom
seismographs (OBS), which occasionally provide useful
information about sedimentary structure.
This paper describes a characteristic feature of sediment
arrivals common to OBS record sections, acquired with
airguns from two widely separated continental rise environments in the NE Atlantic, and relates the inferred
velocity structure to an important increase of lithification
within the top few hundreds of metres of the sedimentary
layer.
COLLECTION OF THE DATA
The data to be described were collected during RRS
Discovery Cruise 161 in August-September 1986 from two
areas, one off Madeira and the other just west of Portugal
(Whitmarsh 1986). Digital ocean-bottom seismographs
(Kirk, Langford & Whitmarsh 1982; Peal & Kirk 1983) were
used to record signals from an array of four 16-litre airguns
fired every 2 min and towed at a speed of 4 to 5 knots.
Shot-to-DOBS ranges were calculated from direct waterwave traveltimes and soundspeed models based on
bathythermograph and water-bottle data. In addition
disposable sonobuoys were used, in conjunction with a
367
368
R . B. Whitmarsh, P. R . Miles and L. M . Pinheiro
5-litre airgun with waveshape kit, to compute interval
velocities within the sediment layer.
MEASUREMENTS ESE OF MADEIRA
The first data to be described were collected from an area
about 200 km ESE of Madeira (Fig. 1). Three experiments
were conducted along a NE-SW profile near the foot of the
continental slope off Morocco where the along-track
sea-floor gradient is 0.2 in 1000 and the across-track gradient
less than 2.5 in 1OOO.
The profile lies at depths of 4300-4400 m. It is bounded to
the SE by the foot of the roughly linear slope on the
seaward edge of the complex continental borderland of
Morocco. To the NE lies the Seine Abyssal Plain which is
separated from the work area by a low sill at about 33"N. To
the WSW the sea-floor deepens steadily towards the
Madeira Abyssal Plain and the Canary Basin (both over
5200 m deep) some 700 km away. Finally to the NW lies the
linear NE-SW slope at the foot of the Madeira archipelago.
Therefore it appears that, at least in the recent geological
past, the major non-pelagic sediment sources have been the
borderland of Morocco south of about 33"N, including the
Agadir Canyon, and the east flank of the rise on which
Madeira stands.
A continuous seismic reflection profile was obtained from
Points K to NN (Figs 1 and 2). Along the profile, oceanic
basement (Reflector e) is clearly visible everywhere and the
sediment thickness varies from 0.75 to 1.85 s, roughly
700-1700 m. The greatest basement relief (1.1 s) occurs just
SW of Point MM and this may correspond to one of the
minor fracture zones indicated by Klitgord & Schouten
(1986). Within the sediments, four distinctive boundaries
are seen. Reflectors c and d parallel each other and the
sub-c sediments appear to be stratified. Above Reflector c
there are three distinct sequences of well-stratified sediment
each having a clear angular unconformity with the
immediately underlying beds. These three sequences are
strongly influenced by turbidites of uncertain age, origin and
composition. It is likely that the sediments below Reflector c
are almost entirely pelagic and will consist of calcareous
ooze or chalk, at least towards the basement.
The interpretation of just one of the three seismic
refraction profiles is described next. The other two profiles
show essentially the same features and at the end of this
section the results of all three profiles will be presented
together. The complete interpretation was published by
Whitmarsh et al. (1988).
In the vicinity of Point MM an airgun seismic refraction
profile, which was recorded by DOBS 1 and 2 1.1 km apart,
and an airgun seismic reflection profile (Fig. 3) were
collected. The record section from DOBS 2 is presented in
Fig. 4. This shows clear basement first arrivals, with an
apparent velocity of approximately 5 km s-', and also very
clear sediment second arrivals at ranges of 9-19 km. The
sediment arrivals for DOBS 2 are shown at a larger scale in
Figure 2. Seismic reflection profile obtained with a 5-litre airgun and waveshape kit along the dotted line in Fig. 1. The positions of volcanic
basement and several reflectors within the sediments have been accentuated. Thick bars = extent of wide-angle refraction profiles shot with
four 16-litre airguns; thin bars = extent of near-vertical-incidence sonobuoy profiles shot with 5-litre airgun and waveship kit. SB = sonobuoy;
triangle = DOBS location.
Figure 3. Seismic reflection profile in the vicinity of Point MM (see Fig. 2 for symbols), together with a velocity structure based on synthetic
seismogram modelling. Note the close correspondence between the reflector (accentuated) and the major increase in S-wave velocity.
Velocity structure of continental rke sediments
369
Figure 4. Unfiltered vertical geophone record section recorded by DOBS 2 at point MM. Reduction velocity is 5 km s-'.
Fig. 5 where it is seen they have an apparent velocity of
about 2.5 km s-'. Close inspection reveals that the sediment
arrivals actually fall into three groups (8.5-10.3, 10.5-13
and beyond 13km). Each group is delayed in time by
roughly 0.2 s with respect to the previous group so that the
groups appear to be en echelon. The same delays are seen
on the adjacent DOBS 1 record section and, within a few
hundred metres, at the same range so that it may be
concluded that they are functions of the velocityldepth
structure and are not the result of lateral inhomogeneities.
In fact the amplitudes and shapes of these arrivals, observed
at about the same range, are closely similar. Similarly the
traveltimes of corresponding peaks and troughs at the same
range vary by up to 0.02 s between the two DOBS indicating
a lack of significant heterogeneity within the sediments.
There appear to be two types of velocity/depth model
which can give rise to the en echelon arrivals. The first
involves a sequence of low-velocity (or negative gradient)
zones between higher velocity layers and the second
requires the second and third groups of arrivals to be
multiply reflected versions of the first group. In the first case
the apparent velocity should increase beyond each offset
while in the second case it should decrease. An objective test
was therefore to use the pair of DOBS as an array to
8
10
12
DISTANCE
estimate the apparent velocity of arrivals for each shot by
cross-correlating between traces. The results appear in Fig.
6. Although imprecise, due to signal :noise limitations, this
figure does suggest an increase in ray parameter (decrease of
velocity) as one passes from group to group with increasing
range thereby favouring the second hypothesis.
Traveltime models were constrained by the sediment
traveltimes discussed above as well as by independent
estimates, from near-vertical incidence reflections, of 0.45
and 0.95 km thick layers with interval velocities of 1.9 and
2.4 km sC1 respectively within the sediments. The traveltime
spent in the upper 1.9 km s-' layer is always so great as to
preclude the possibility of multiples being reflected at the
sea-bed. Thus reflection of refracted arrivals within the
sediment is indicated. Models of this sort were found to fit
the observed arrival times when the multiples were reflected
at the interface between the upper and lower sediments
(Fig.7). As a final check synthetic seismograms were
calculated using the reflectivity method (Fuchs & Mueller
1971). The best fitting model (see MM in Fig. 8) was able to
reproduce the en echelon arrivals, as well as their relative
amplitudes (Fig. 9), quite successfully (Fig. 5). This appears
to confirm the correct choice among the hypotheses
discussed above.
(24
16
18
--
20
*re
5. Unfiltered vertical geophone record section recorded by DOBS 2 at Point MM. Reduction velocity 2.5kms-'. Synthetic
seismograms are drawn as light dashed lines for an upper shear wave velocity of 250 m s-'. The heavy dashed lines indicate the sediment
arrivals.
370
R. B. Whitmarsh, P. R . Miles and L. M. Pinheiro
I\ 4 \I
I
10
15
20
Mean distance to DOBS (km)
F p r e 6. Phase velocity of sediment arrivals across the two-DOBS array at Point MM.The hachured rectangles mark the transition between
different groups of sediment arrivals. The dashed line indicates the mean velocity of the primary arrivals which were too close to the water
wave to be cross-correlated. The vertical bar indicates the uncertainty due to a lag error of 1.
li;-
Veloctty (Km/s)
4 ,
0.5
1.0
-----____
i
Distance (Km)
Fylre 7. (a) Calculated hodochrons which fit the observed amvals in Fig. 5 . Three groups of sediment arrivals and a set of basement amvals
are shown. The concave upwards curves represent the calculated traveltimes of wide-angle basement reflections. (b) A velocity/depth
structure, constrained by traveltimes alone, which was used to calculate the traveltimes in Fig. 7(a).
NORTHEAST
SOUTHWEST
VELOCITY (km s-' )
NN
1.5
2.0
2.5
SB6
MM
1.5
21.5
SB5
2.0
2.5
I
DEPTH
.
.
1.5t
1.5
L
SB4
2.0
2.5 1.5
I-I
2.0
2.5
1.5
2.0
2.5
I
5
2.0
NN based on DOBS and sonobuoy recordings. The continuous lines
denote models based on synthetic seismogram calculations. The discontinuous lines (dashes and dots) are models from near-vertical incidence
reflections. At SB4, a reflection from the intrasediment reflector (ISR) was not detected. One possible interpretation of the SB4 model, which
includes this reflector, is indicated. Hachured lines = top of volcanic basement. Reflector a is the topmost accentuated reflector in Fig. 2.
Figure 8. Summary of all velocity/depth structures from Point L to Point
Velocity structure of continental rise sediments
POINT M M
n
n
u.u-
~
1
4
-
15
10
20
Distance (Km)
F i r e 9. Amplitude/distance plot for sediment synthetic seismograms compared with the seismograms observed at Point MM. The
continuous line marks the envelope of the observed points.
In order to obtain sufficiently strong multiple
substantial shear wave velocity (V,) contrast' is
across the intrasediment reflecting horizon.
illustrated in Fig. 10, where a Poisson's ratio of
5
T-&
(S)
phases a
required
This is
0.25 was
371
assumed for the lower sediments (giving V, = 1.25 km s-'),
whereas V, in the upper sediments must be about
0.25 km s-l and certainly less than 0.5 km s-'. No realistic
change in density can produce the required effect instead.
The efficiency of such an interface as a reflector for rays
coming from below is indicated in Fig. 11. The reflection
coefficient peaks for angles of incidence in the range 50°-800
(which includes rays turning at depths where the velocity is
in the range 2.23-2.40 km s-l) and reaches a maximum of
0.4, for example, when V, above the interface is
0.25 km s-'. Except for the shallowest rays this reflection
coefficient exceeds even that of the sea-bed.
Although, following Helmberger, Engen & Scott (1979),
rather high Q factors (Q, = 500, Q, = 250) were assumed
when calculating the synthetic seismograms, no significant
amplitude change was detected if much lower values
(Q, = 100, Q, = 50) were used (Whitmarsh et al. 1988).
The other two refraction lines at Points L and NN (Figs 1
and 2) also exhibited en echelon sediment arrivals. Here too
synthetic seismogram modelling required the presence of an
intrasediment reflector with a strong shear wave velocity
contrast, but limited compressional wave velocity contrast
(Fig. 8); indeed at Point NN no first-order contrast in
compressional wave velocity seems to exist.
JA
'
3
2
a
I
10
12
1
18
20
DISTANCE (kml
Figure 10. Three sets of synthetic seismograms based on the best compressional wave velocity/depth model for Point MM shown in Figs 3 and
8. The synthetics differ in the shear wave velocity assumed for the sediments above the intrasediment reflector as follows; (a) 100 m s-' (b)
250ms-' (c) 500ms-'. The figure indicates that a shear wave velocity of about 250msC' is required to fit the observed record section
in Fig. 5 .
372
R. B. Whitmarsh, P. R . Miles and L. M . Pinheiro
Displacement
reflection
coefficient
MEASUREMENTS OVER THE EASTERN
IBERIA ABYSSAL PLAIN
t
Angle of incidence (degrees)
Figure 11. Calculated P-wave displacement reflection coefficients
for up-coming incident P-waves based on Zoeppritz's equations.
Dotted line, sea-bed reflection; dashed line, intrasediment reflection
for upper S-wave velocity of 250 m s-'; continuous line, as dashed
line but S-wave velocity is 100 m s-'. Other parameters as for the
Point MM synthetic seismogram model.
When the three velocity-depth models based on synthetic
seismograms are compared with the almost coincident
reflection profile (Figs 2, 3 and 8) it becomes clear that the
intrasediment reflector is associated with a 'depth' roughly
midway between Reflectors a and b, at which the turbidite
banding becomes distinctly finer. This horizon cannot be
traced everywhere (Fig. 2). Its rather indistinct nature is
compatible with the low, or even zero, first-order
compressional wave contrast predicted by the synthetic
seismogram models.
Having arrived at the interpretation discussed above it was
natural to consider whether the Madeira sediment model
might also be valid in similar locations on, or close to,
abyssal plains. By chance we also shot a series of refraction
profiles in 1986 over the eastern Iberia Abyssal Plain, as
part of a study of the ocean-continent boundary, which
recorded closely spaced airgun shots out to about 25km
range (Fig. 12). Inspection of the record sections showed no
evidence of en echelon sediment arrivals for Lines 1 and 2
but good evidence of such arrivals for Lines 3 and 4 at all
four DOBS locations which recorded airgun shots. The
eastern two lines have a maximum sea-floor gradient of
1.5-2 parts per lo00 down to the west, whereas west of
12O45'W the gradient flattens to about 0.6 parts per 1OOO.
Two E-W reflection profiles obtained in the area are
shown in Fig. 13. They show stratified sediments which are
over 2 s thick in places and lie on an irregular, and at times
poorly discerned, acoustic basement. The principal discontinuity within the sediments is a horizon up to 0.7s below
the sea-floor, below which the sediments exhibit gentle
folding and above which the stratification is essentially
horizontal.
The record section for DOBS 1 at Line 4 is shown in Fig.
14. The sediment arrivals exhibit en echelon behaviour and
have a velocity of about 2-2.5 km s-'. It is not possible with
our data to demonstrate that the en echelon behaviour is not
due to lateral inhomogeneities because the DOBS were
42
41
40
-14
- 12
39
-10
Figure 12. Location and numbering of airgun refraction profiles in the eastern Iberia Abyssal Plain which recorded en echelon sediment
arrivals (dots) and those which did not (open circles). A and B denote reflection profiles shown in Fig. 13. Square, DSDP Site 398,
Figure 13. E-W seismic reflection profiles over the eastern Iberia Abyssal Plain. See Fig. 12 for locations, Fig. 2 for symbols. Note east is to
the left.
EAST
WEST
Velocity structure of continental rise sediments
2
373
T
7
6
8
9
10
11
12
13
14
15
.
DISTANCE (km)
Figure 14. Unfiltered record section of DOBS 1 from the west end of Line 4, reduced to 2.5 km s-'. The light dashed lines are synthetic
seismograms for the model in Figs 15 and 17. The heavy dashed lines indicate the sediment arrivals.
deployed singly. However lateral homogeneity is a
reasonable assumption in view of the seismic reflection
profile associated with this line (Fig. 13) and the
observations near Madeira which were made i n . a similar
sedimentary environment. Traveltime modelling of the
DOBS 1 sediment arrivals required a very similar
compressional wave model to those observed near Madeira
with multiple reflections occurring at an intrasediment
reflector about 450 m below the sea-floor. Similar velocity
models were obtained from the other DOBS record sections
in the eastern Iberia Abyssal Plain where en echelon
sediment arrivals were observed (Fig. 15). The depths of the
intrasediment reflector lie very close to the unconformity
previously recognized on the profiles in Fig. 13.
Finally synthetic seismograms were calculated using the
reflectivity method (Fuchs & Mueller 1971) to model the
amplitudes of the sediment arrivals seen in Fig. 14. This
confirmed the need for an intrasediment reflector with a
relatively strong shear wave velocity contrast (Fig. 17).
Indeed the upper sediments probably have a mean velocity
close to lUOms-', and certainly less than 250ms-'. The
relative amplitudes of the observed and synthetic seismograms are compared in Fig. 16. Fig. 17 also demonstrates
the apparent coincidence of the 'depth' of the reflector
noted in Fig. 13 and the intrasediment reflector required by
the seismic modelling.
(hB 4 )
VYEST
4w
3s
4c
EAST
Q
0
0.3
1.0
1.5
3.0
.-
3.5
.-
4.0
--
4.5
-
Figure 15. Summary of traveltime models, and one synthetic seismogram model (4W), which fit all the record sections in the eastern Iberia
Abyssal Plain which exhibit en echelon sediment arrivals. 4W/4C/4E=west endlcentreleast end of Line 4; 3S=south end of Line 3;
ISR = intrasediment reflector.
374
R . B. Whitmarsh, P. R. Miles and L. M . Pinheiro
0.9 1.0
0.8
-
0.7
-
DOBS 1
LINE 4
’D
3
5
0
0.6
-
0.5
-
0.4
-
0.3
-
_J
0.2 0.1
- _ -0
-x
Calculated
Observed
-
0.0
6.0
I
I
I
I
I
I
I
I
1
I
I
I
7.0
8.0
9.0
10.0
11.0
12.0
13.0
14.0
15.0
16.0
17.0
18.0
Distonce
(km)
Figure 16. Relative amplitudes of the differentsediment phases seen on the observed and synthetic seismogram record sections form DOBS 1
on Line 4. Synthetic seismograms were based on the model in Figs 15 and 17.
DISCUSSION
The three principal features of the seismic velocity models
for the sedimentary sequences in two separate continental
rise or abyssal plain environments are as follows.
(i) An intrasediment reflector is present 100 to 600m
below the sea-bed which, at wide angles, gives rise to
multiply reflected compressional waves with a clear en
echelon arrangement on record sections.
(ii) The main seismic characteristic of the reflecting
horizon is the relatively large contrast in shear wave
velocity. The synthetic seismograms require the overlying
sediments to have a shear wave speed of about
100-250 m s-’ (implying a Poisson’s ratio of 0.47-0.42) and
the the underlying sediments a velocity of about 1.25 km s-l
(0.25 Poisson’s ratio). The first-order contrast in compressional wave velocity however need not be greater than
0.4 km s-’.
(iii) On reflection profiles, at the ‘depth’ of the reflecting
horizon, a clear change in reflection character is usually
discerned although this is sometimes laterally discontinuous,
probably due to a very small P-wave velocity contrast. The
varying ‘depth’ of the horizon on reflection profiles is
matched by similar changes in the calculated velocity
structures.
The shear wave velocity of a sediment depends principally
on its rigidity modulus. Rigidity is acquired by deep sea
sediments by the physical interlocking of grains and/or by
chemical bonding between them. These processes depend
on age, compaction and diagenesis but are not well
understood. An additional factor is the role of pore pressure
on effective stress (Schultheiss 1981); if the pore pressure
increases the resulting reduction in effective stress will
reduce the shear wave velocity and vice versa.
The rigidity of a sediment is likely to be disturbed by most
sampling techniques either due to physical shock or by
changes in effective stress during recovery (Schultheiss
1982). Thus probably only in situ measurements of shear
wave velocity are reliable. A variety of such measurements,
made in oceanic depths, appears in Table 1. Although only
four measurements pertain to a depth of over 150m below
the sea-bed, and one of these yields unusually high values,
velocities as low as 100 to 250ms-’ at depths of 100 to
600 m are not inconsistent with the values in the Table. The
only two detailed in situ measurements of velocity gradient
indicate a rapidly decreasing velocity gradient with depth (in
the top 20 m) and perhaps a linear increase in shear modulus
(Davies 1965; Whitmarsh & Lilwall 1982). Simple linear
extrapolation of the shear modulus to 300m depth, for
example, implies a shear wave velocity of about 750 m s-’.
This is likely to be an upper limit as, in the absence of
significant diagenesis, rapid acquisition of rigidity by the
interlocking of grains is probably achieved just below the
sea-bed and is unlikely to persist at the same rate at depths
of several hundred metres.
It is difficult to explain the contrast in shear wave velocity
across the intrasediment reflector. Because the reflected
P-wave energy is predominantly at 7 Hz the change from the
upper to the lower sediments must be relatively sharp, i.e. it
occurs within 70 m (a quarter wavelength). It seems unlikely
that such a relatively abrupt change could be the result of
effective stress changes alone. Alternatively the upper and
lower sediments may have experienced very different
depositional or consolidation histories. Results from nearby
DSDP boreholes were inspected to try to resolve the
question.
Eastern Iberia abyssal plain
To date no DSDP or Ocean Drilling Program (ODP)
boreholes have been sited on an abyssal plain in the eastern
North Atlantic Ocean. However Site 398 was located in
3910 m water depth about 200 km NE of the seismic profiles
reported here (Fig. 12; Sibuet, Ryan et al. 1979). The site lies
near the foot of the continental rise on the southern margin
of the Galicia Bank block but above the level of the Iberia
Abyssal Plain. However, some seismic reflectors identified
in the region around the site, which were drilled and cored
by DSDP, can be traced laterally via a network of
intersecting reflection profiles to our DOBS profiles (D. G.
Masson, personal communication). Thus the intrasediment
reflecting horizon of our seismic models, which we have
shown to correlate with a marked reflector on our reflection
Figure 17. Seismic reflection profile along the refraction profile recorded by DOBS 1 on Line 4. The velocity structure was derived from
synthetic seismogram modelling. Note the near coincidence of the abrupt velocity increases in the model at 0.5 s with a reflector seen on the
reflection profile. Triangle = DOBS 1 location.
Velocity structure of continental rise sediments
375
Table 1. In situ measurements of sediment shear wave velocity in oceanic depths.
Authors
Water
depth
Measurement
technique
Sub-bottom
depth
Sediment
type
(m)
(m)
Hamilton
4410
Scholte waves
Oavies (1965)
9 (1970)
986
submersible
b probes
foraminifera1
lutites b turbidites
Shear-wave
velocity
(m. s-‘)
0-20
NU lndian Ocean
50-190
0
clayey silt
Location
88
off San Diego, E. Pacific
sediment (S-P)
times
4000
?
0-400
400
off California. E. Pacific
sediment (S-P)
times
2850
7
0-150
150
G u l f of California,
Uhitmrsh b
Lilwall (1982)
Scholte waves
5260
Au b Clowes (1984)
sediment (S-P)
times
Sutton
(
1971)
Lewis 6 McClain (1977)
Toksoz
9
(1985)
Sauter
Ouennebier
(1986)
9 (1987)
E. Pacific
?
1410
full wave-form
modelling
3800
Scholte waves
sediment (S-PI
times
5500
calcareous lutites
b turbidites
0-50
off Madeira, C. Atlantic
25- I70
?
0-1000
500
calcareous 6
siliceous ooze
6 sand layers
0-135
135-322
850
Baltimore canyon,
900
NU Atlantic
322-553
1000
Ot
35+
(3.2 s
?
siliceous clay
0-360
210
Nootka fault zone,
NE Pacific
)
off San Diego,
E. Pacific
Hole 581C, NU Pacific
Bars over velocities indicate averages over the given depth intervals.
profiles, can be linked to the ‘green’ reflector which was
cored at Site 398 and placed there at a depth of 380 m. This
reflector was linked to a number of mainly physical changes
in the cores which are tabulated in Table 2. The depths of
some of these changes are approximate due to gaps in coring
(9.5 m long core intervals were drilled at 318.5, 366, 394.5
and 413.5111 below the sea-floor). Nevertheless it seems
clear that at Site 398 the ‘green’ reflector is associated with
the increased induration of the sediment from ooze to chalk
which occurs over a range whose greatest extent is
319-395 m. Such induration is expected to lead to a strong
increase in shear wave velocity such as we have observed. A
Table 2. Characteristics associated with the ‘green’ reflector at Site 398.
Sub-bottom
Depth
Characteristic
Change observed
lm)
319-366
Lithology
Transition from ooze t o chalk
376-395
Sedimentation
Hiatus between Upper and Middle
Miocene; zones N.14, N.15 missing,
probably erosional.
370-400
Seismic velocity o f cores
Increases markedly (1.6 t o
1.7 k m 5 - l )
370-400
Density of cores
Increases markedly (1850 t o
1950 kg m-3)
395
Drilling rate
Major decrease
424
Lithology
Appearance o f upto 30% biogenic
silica in lithological Unit IC.
376
R. B. Whimzarsh, P. R . Miles and L. M . Pinheiro
clear increase in compressional wave velocity at the ‘green’
reflector is also evident from interval velocities of the
formations above and below the ‘green’ reflector calculated
from multichannel data (Groupe Galice 1979).
ESE of Madeira
The closest well-cored DSDP site is Site 370 (32”50’N,
10’47’W) situated about 400km ENE of our seismic
measurements (Lancelot et al. 1977). This site lies at the
foot of the continental slope. of Morocco at 4214m depth.
Even if it were possible to Link this site to our measurements
via a network of reflection profiles, any correlation of
reflectors made in this way would probably be tenuous. The
difficulty is that there is currently a sill between the Seine
Abyssal Plain (which is adjacent to, and 400 m deeper than,
Site 370) and the area of our measurements thereby
implying a different sedimentation history in detail at the
two sites. However if the intrasediment reflector here is due
to an important increase in lithification, as at Site 398, then
the data from Site 370 may be relevant in that the same
regional sedimentation conditions (surface productivity and
access to temgenous sources) probably apply.
At Site 370 only five cores were taken above 426m
downhole but thereafter about one core was cut every 19 m.
A Mid-Eocene reflector, apparently associated with an
increase in induration, occurs at a depth of about 475m
(0.5-0.6s). To the west of Site 370, adjacent to and
under the Seine Abyssal Plain, this reflector has a similar
appearance to the intrasediment reflector ESE of Madeira
(Lancelot et al. 1977), viz. above it the record is relatively
transparent and below it, well stratified, but sometimes
laterally discontinuous. At the site the overlying sediments
are more stratified. Indications of the increase in induration
at Site 370 are listed in Table 3.
The fact that the transition from clay to claystone appears
at about 465m at Site 370, that this depth Lies within the
range of depth of the intrasediment reflector ESE of
Madeira (Fig. 8) and that a reflection of similar character is
associated with both the important increase in lithification
and the intrasediment reflector is suggestive of a common
cause. Without further evidence however this conclusion
must remain tentative.
Lithifieation at other Atlantic drillsites
A search was made of the DSDP and ODP Initial Reports
for other examples of important increases in sediment
induration or lithification associated with reflecting horizons.
Nine examples were found at sites which lay on Atlantic
continental rises or on adjacent abyssal plains and where
coring had been reasonably continuous and the seismic
reflection records clearly showed an intrasediment reflector.
Sites on highs were excluded. The chosen sites are listed in
Table 4. Although the depths and types of lithification
changes are variable, the range of depths (210-540m) is
encompassed by the depths at which our OBS observations
show an intrasediment reflector to lie. Table 4 provides
circumstantial evidence in support of the hypothesis that
intrasediment reflectors detected at several hundred metres ,
depth by OBS may be due to important increases in
induration or lithification.
The reflectors at three holes in Table 4 were logged with
sonic and porosity tools. In Hole 417D there is a fairly sharp
increase in velocity and decrease in porosity within 20 m of
the best estimate of the ‘lithification’ horizon (Donnelly,
Francheteau et al. 1977). Better logs at Holes 646 and 647
show very abrupt changes in velocity and porosity over just
a few metres (Srivastava et al. 1987). This evidence suggests
that generally ‘lithification’ reflectors may be particularly
sharp.
Absence of en ecbelon arrivals at some sites off Iberia
Figure 12 shows that en echelon sediment anivals were not
observed by the three easternmost DOBS. The reason for
this is unclear. On our profiles the ‘green’ reflector is much
fainter at the east end of Profile B (Fig. 13) and on its
extension towards the south end of Line 1. This reflector is
also generally shallower (less than 0.5 s) on Profile B east of
12’40’W than it is to the west. The ‘green’ reflector is also
shallow (0.3s) and weak at the north end of Line 2. It is
plausible therefore that the increase in lithification
Table 3. Characteristics associated with the reflector at 0.5-0.6 s TWT, Site 370.
Sub-bottom
Depth
Characteristic
Change observed
(m)
Shear strength
All values > 50 k P a ; above
limit of shipboard equipment.
ca.460
Velocity of cores
Increases notably to at least
2.2 km s-‘.
ca. 4 6 0
Porosity of cores
Decreases to 46% associated with
increase in cementation of silty
clays and porcellanite.
Lithology
Transition from clay to claystone
(see core descriptions)
> 326
465
Velocity structure of continental rise sediments
377
Table 4. Important changes in lithification or induration associated with reflectors at Atlantic abyssal plains and
continental rises.
DSDP o r ODP
SiteILeg
Locat ion
Water
depth(m)
Ref lector
depth(m)
112/12
S. Labrador Sea
3657
315
early Oligocene
stiff clay to hard clay
646/105
S. Labrador Sea
3451
335
early Pliocene
biogenic silica to biogenic calcite
6471 105
S. Labrador Sea
3862
235-251
early Oligocene
siliceous biogenic claystones to
calcareous biogenic claystones
118112
C. Bay of Biscay
4901
400-500
late-middle Miocene
clay t o sandstone
418Al102
W. Atlantic Ocean
551 I
240
late Cretaceous
clay to claystone
4170/51-53
W. Atlantic Ocean
5482
286
late Cretaceous
clay t o claystone
354139
Ceara Rise
4052
2 10- 240
middle Miocene
ooze t o chalk
367/41
off W. Africa
4748
300-320
late Eocene
first chert
515/72
Brazil Basin
4250
5 10-540
late Oligocene
mud t o mudstone and chert
associated with the ‘green’ reflector is depth controlled and
insufficiently developed at the three eastern sites to give rise
to intrasediment wide-angle reflections. At Site 398 on the
other hand, the ‘green’ reflector is associated with the
transition from ooze to chalk but lies at a depth of 0.42s.
The sea-bed is also 1OOOm shallower there, implying a
different set of physical, and perhaps chemical, conditions.
Lithological change
(5) If the above hypothesis is correct, and since the
increases in lithification/induration of sediment probably
depend on a combination of different physical and chemical
effects, which may not always be achieved concurrently at
the same time horizon, it follows that ‘lithification’ reflectors
may be diachronous.
ACKNOWLEDGMENTS
CONCLUSIONS
The following conclusions may be reached from the data
presented above.
(1) Ocean-bottom seismographs, recording closely spaced
airgun shots in two separate continental rise areas in the
eastern North Atlantic, have detected en echelon sets of
sediment arrivals.
(2) Modelling using both traveltimes and synthetic seismograms indicates that these sets of arrivals are refracted
phases multiply reflected at a horizon within the sediments.
A characteristic of this horizon, required by the amplitude
of the offset arrivals, is that it possesses a substantial
contrast in shear wave velocity even though the contrast in
compressional wave velocity is less than about 0.4km s-’.
(3) Correlation with a DSDP hole via a common reflector
in one case (off Iberia) suggests that the horizon is caused by
the acquisition of appreciable rigidity accompanying an
important increase in lithification (ooze to chalk). In the
other case (off Madeira) the correlation is tenuous although
the horizon and the increase in lithification (clay to
claystone) appear to be at comparable depths.
(4) It is possible that similar studies with ocean-bottom
seismographs elsewhere may be used to detect important
increases in lithification and may even be useful for studies
of diagenesis. It is unlikely however that surface sonobuoys
will detect this phenomenon because of the masking effect
of the direct water wave.
We gratefully acknowledge the support of the Admiralty
Research Establishment, Portland which enabled us to
conduct the work near Madeira. We also thank the officers
and crew of Discovery 161 who helped us to collect the data
reported here. Bob Kirk and Martin Saunders made
important contributions to the recording and processing of
the OBS data. L.M.P. thanks the Senisos Geologicos de
Portugal for financial support and IOSDL for providing
research facilities.
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