The seismic structure of 140 Myr old crust in the western central Atlantic Ocean

Transcription

1 Geophys. J. R. astr. SOC. (1983) 72, The seismic structure of 140 Myr old crust in the western central Atlantic Ocean G. M. Purdy Woods Hole Oceanographic Institution, Woods Hole, Massachusetts 02543, USA Received 1982 May 3; in original form 1982 February 5 Summary. A detailed seismic refraction experiment using explosive sources and ocean bottom hydrophone (OBH) receivers was carried out over Mesozoic magnetic anomaly M17 about 300 km south-west of Bermuda. Amplitude and travel-time interpretations show this 140Myr old crust to be uniform, on the seismic scale, over a lateral distance of almost 1OOkm. The best estimate of the velocity structure consists of an average 700 m of sediment overlying a total thickness of igneous crust of 7.2 km. The principle components, beginning with a 5 km s-l velocity at the top of layer 2, are a 0.4 km thickness with gradient 1.1 s-', a 1.9km thickness with gradient 0.64s-', 1.7 km thickness with gradient 0.1 s-', a 3.2 km thickness with zero gradient and constant velocity of 7 km s-' below which lies a 0.5 km thick moho transition zone. The uppermost few kilometres of the upper mantle apparently has little or no velocity gradient. The normal incidence two-way reflection time through this structure agrees to better than 0.1 s with the location of the Moho reflection seen on the IPOD/USGS multichannel reflection profile that passes within 1 km of this experiment. The travel times of the converted shear wave arrivals constrain Poisson's ratio in the igneous crust to k A delay time study of a 20km radius circle of shots fired around a small (- 6 km) five-element array of ocean bottom hydrophone receivers confirms the uniformity of structure of this old crust on the scale of a few kilometres. The observed differences in shallow crustal delay time between the five OBH instruments can be wholly attributed to differences in sediment thickness beneath the receivers as determined by a deep towed hydrophone seismic reflection profile. These experiments were located clear of fracture zones or other structural anomalies in a region of well-defined linear constant amplitude magnetic anomalies. We propose that this was the primary cause of our result of lateral uniformity of structure on the scale of a few kilometres. We suggest that the reason that such a simple uniform structural model for oceanic crust is not supported by the historical seismic refraction dataset is that a large proportion of the old experiments were poorly located relative to fracture zones and other structural anomalies (unknown at the time) and the

2 116 G. M. Purdy Introduction data density and sediment thickness information available for each experiment was insufficient to overcome the errors introduced by the basement topography. The source of our present knowledge of the seismic structure of oceanic crust is the interpretation by many investigators of seismic refraction experiments that have been carried out during the last 25 years. The first classic compilation of these results by Raitt (1963) deserves repetition here (Table 1). Since that time the size of our total seismic refraction dataset has increased many-fold, instrumentation and experiment design have improved, and considerable advances have been made in interpretation techniques. The result is that significantly more is now known concerning the details of the seismic structure of oceanic crust (e.g. Spudich & Orcutt 1980) but our view of the variability in structure has, in fact, changed little if at all from that indicated by Raitt's root mean square deviations (Table 1). This is evident from more recent data compilations by, for example, Fox, Schreiber & Peterson (1973) or Christensen & Salisbury (1975). Important efforts have been made to explain the observed differences in seismic structure in terms of aging processes (e.g. Houtz & Ewing 1976; Goslin et al. 1972), tectonic province (e.g. Ludwig 1972) and errors in interpretation method (Purdy 1982a, b). The results show that we remain far from having sufficient understanding of the primary processes controlling the seismic velocity structure of oceanic crust to be able to explain this observed variability in terms of a unifying predictive model of crustal accretion and alteration. Before any further effort can be justified to quantify in detail any systematic changes in structure in the hope of relating them to known geological processes, we must determine the degree of innate variability in seismic structure (on a scale of a few kilornetres) to be sure that it is meaningful to attribute resolvable differences in observed structure to systematic processes. It could be argued that the formation and alteration of oceanic crust is sufficiently complex (on the scale resolvable by the conventional refraction technique) that the resulting seismic structure is truly random and any gross model of oceanic crustal structure involving laterally homogeneous velocities is meaningless. There are two primary reasons why this obvious problem cannot be solved unambiguously using existing data. First, there exists too great a range of data quality, type and density to be sure that any differences are not in fact due to artefacts of the interpretation. Secondly, and perhaps most importantly, is the fact that a large portion of the available dataset is poorly located within what we now know to be the tectonic framework of ocean basins. A prime example of this is shown in Fig. 1. Detrick & Purdy have shown that major structural changes may be associated with fracture zones. Fig. 1 shows that more than half of the existing seismic refraction dataset in the western central Atlantic Ocean is located on or within a few kilometres of what is now known to be the fossil traces of the many long-lived transform faults existing along the mid-atlantic Ridge. Thus it is possible that the view provided by, for example, Raitt (1963), Fox et al. Table 1. The average seismic velocity structure of oceanic crust as determined by Russ Raitt in his classic paper published in Layer Velocity Thickness (km s-') (km) r t f _ 0.24

3 Seismic structure of old Atlantic crust 1 I7 --- Fracture zone - Magnetic sea-floor spreading lineation - Seismic refroctoon shooting line, reversed a Bathyrnetric high Positive magnetic anomaly Figure 1. Compilation of seismic refraction profiles in the western central Atlantic Ocean. The location of magnetic lineations and fracture zones is from Schouten & Klitgord (1977). The seismic refraction data sources are Ewing, Sutton & Officer (1952). Officer, Ewing & Wuenschel (1952), Katz & Ewing (1956), Ewing & Ewing (1959), Hersey et al. (19593, Houtz & Ewing (1963, 1964) and Sheridan et al. (1966). Note that a significant proportion of these data cross or lie close to the fracture zone traces. The location of the work described in this paper is indicated by the black square south-west of Bermuda. 1 (1973) or Christensen & Salisbury (1975) of the variability in structure of so called normal oceanic crust is far too great and is primarily an artefact caused by this fracture zone effect. This paper describes the interpretation of seismic refraction experiments carried out using explosives and ocean bottom hydrophones that were located within an area of dense aeromagnetic coverage over the Mesozoic series of magnetic anomalies approximately 300 km

4 118 G. M. Purdy 31"33'N 31"ZdN 31"IO'N 68" 30 W 68-15, W 68'00'W Figure 2. Aeromagnetic anomaly contour chart of the principal work area from Schouten & Klitgord (1982). Stars denote the locations of the five ocean bottom hydrophone instruments. OBH 2 is located in the trough of Mesozoic magnetic anomaly M17. The solid dots denote the shot locations which are annotated with their shot number. The line is the track of the IPOD/USGS multichannel profile that is also annotated with shot numbers. south-west of Bermuda in the western central Atlantic Ocean. Because of the available high quality aeromagnetic data and the detailed interpretation of these data made by Schouten Rr Klitgord (1982) it was possible to locate the experiment with certainty over an area free from fracture zones and over Mesozoic magnetic anomaly M-I7 known to have been created during a time of constant spreading rate and direction (Fig. 2). The intent was to make a particularly careful determination of the seismic structure of what may be truly 'normal' 138 Myr old oceanic crust in the western central Atlantic Ocean and to determine its degree of variability on a scale of a few kilometres. To this end several seismic experiments were carried out in this region using airguns, explosives and ocean bottom hydrophone receivers. In this paper we shall describe the results of two of these experiments: a 100 km line and a 20km radius circle of explosive shots as received by five ocean bottom hydrophone receivers. The experiments The receiver and shot locations for the line and circle data are shown in Figs 2 and 3. The receivers were Woods Hole Oceanographic Institution ocean botton hydrophones (Koelsch & Purdy 1979) and the shots were Tovex explosive of either 10.9kg (241b) or 108.9kg (2401b). Fig. 2 shows the lookm long shooting line to be oriented perpendicular to the well-defined magnetic lineation trend. The five ocean bottom hydrophones (OBH) are located approximately in the middle of this line and they straddle the eastern flank of the

5 Seismic structure of old Atlantic crust nT linear trough trending 046" that is the anomaly due to the Mesozoic magnetic reversal M17 (- 138Myr bp; Larson & Hilde 1975). The north-eastern corner of Fig. 2 shows the disturbed magnetic field associated with a small offset zone (which is in fact the 'northern fracture zone' referred to in Purdy & Rohr 1979). A second small fracture zone lies to the south (the 'central fracture zone' of Purdy & Rohr 1979) but is not obviously visible as it just intersects the south-western corner of Fig. 2. The shot-receiver configuration shown in this figure provides five split refraction profiles each extending approximately 50km to the north-west and south-east of the receivers which are separated by as little as 2 km from one another. In the following two sections of this paper these ten datasets will be interpreted using conventional methods. Shot and receiver locations were determined using Loran C to interpolate between the highest-quality satellite fixes (- 5 per day). Extensive conventional airgun reflection profiling along the shooting line and over the receivers provided the necessary sediment thickness data. A few tens of kilometres of reflection profile using a hydrophone towed within loom 30' 20' 10' 68"OO'W Figure 3. Bathymetry contour chart of the principal work area. The bounds are identical to those of Fig. 2. The contour interval is 10 corrected m. The track density was too great to include in the figure. Stars denote the locations of the five ocean bottom hydrophone instruments used in the interpretation of the circle shots. Solid dots denote the shot locations which are annotated with their shot number. Contoured by Erika Francis.

6 120 G. M. Purdy of the seafloor were located close to the receiver locations to provide particularly precise measurements of basement morphology (Purdy & Cove 1982). These data will be described later in this paper. A sediment thickness contour chart of this same area (Rohr & Purdy 1981; Rohr, in preparation) shows thicknesses varying from 0.5 to 0.8 s two-way reflection time ( m). The shot and receiver locations for the second dataset to be interpreted in this paper are shown in Fig. 3 which also shows a 10m contour interval bathymetry chart of the area. It can be seen that there are no dominant bathymetric trends and generally relief does not exceed a few tens of metres. In this case a 20km radius circle of 10.9 kg (241b) explosive charges were detonated around a central array of five OBH instruments, the intention being to apply delay time theory to investigate differences in structure within the small section of crust beneath the instruments (White & Matthews 1980). Both the lookm line and this circle experiments were carried out during one deployment of the OBH instruments, thus the location of instruments 8, 4 and 6 (data from which are used in both experiments) shown in Figs 2 and 3 is identical. The intent was to centre the circle of shots on OBH 4. The digitization reduction and corrections of all these data was carried out in a conventional manner (Purdy et al. 1982). Shot-receiver ranges were determined from direct water wave travel times. Travel-time and amplitude interpretations of the 100 km line are given in the next two sections of this paper followed by the analysis of the 20 km circle travel times. The 100 km line: travel-time interpretations The shot-receiver combinations shown in Fig. 2 produce ten refraction profiles which were all plotted in the form shown in Fig. 4. First arrival travel times were picked by eye, corrected for variations in sediment thickness beneath the shots and receivers and displayed as conventional T-X curves (Fig. 5). The travel-time effects of the basement topography were removed using a datum model consisting of 700 m of 2 km-' sediment overlying a 5.O km s-' basement. The correction to be added to the observed travel time is simply the appropriate difference between sediment and basement rock delay times At = (D-s) [(ut -p2)112-(ui-p2)1/2] where D is the reference sediment thickness, s is the measured sediment thickness, u1 and u2 are respectively the sediment and basement slownesses and p is the ray parameter. This approach was chosen because it is impossible to make accurate determinations of the ray parameter p using a relatively sparse dataset of this type and we have an imprecise knowledge of the uppermost basement velocity. Thus it is best to use a correction the magnitude of which is sufficiently small that the effects of unavoidable errors in its calculation are minimized. In fact this method, of the few available, results in the smallest correction, and thus is Figure 4. Three sample record sections of data collected from the line of explosive shots. Reduction velocity is 8 km SKI. Shot numbers and are kg (240 Ib), the remainder are 10.9 kg (24 Ib). The amplitudes have been normalized for charge size and amplified with range using the factor (R/R,) (Wo/W)o'6s where R,= 5 km, We= 10.9 kg (24Ib), W is the charge weight of a shot a distance R from the receiver. The travel-time curves are computed from the velocity-depth function shown in Fig. 7. The phases are identified using the nomenclature of Spudich & Orcutt (1980). The seismograms have been truncated at the time of arrival of the direct water wave and low pass filtered at 15 Hz. The seismograms have not been corrected for sediment thickness variations (as the correction is phase velocity dependent and the appropriate correction cannot be simultaneously applied to compressional and shear arrivals within a single seismogram). Note the well developed triplication point at - 30 km, the lack of P, energy, the s separation between P, and P,P at - 25 km range and the gross variability in converted shear wave amplitudes.

7 12 Seismic structure of old Atlantic crust iy L SHOT NOS ip I L to RZWf /kmd

8 122 G. M. Purdy the least sensitive to errors from any source. The sediment thicknesses used in these calculations were picked at the estimated point of ray entry beneath each shot and receiver from conventional airgun seismic reflection profiles that were located along the shooting line. Smooth polynomial curves were fitted to the data constrained to decrease gradient with range and to minimize the root mean square deviation of the observed travel times from the curve. A prominent feature of these curves was their lack of upper mantle type velocities (- 8 km s-'). Even at ranges of 50km or so apparent velocities of 7 km s-l were common. Closer study of the record sections showed that the phase observed as an apparent first arrival at these longer ranges is in fact a wide-angle reflection from the moho. The true P, I 2w 1 I 58 \+ cl 4.O 58 %\" k %+ L %\" k 4.O 58 F E j./ c 2E 8E F 4~ 1 %\" d KMS KMS Figure 5. Travel-time curves for all ten record sections available in this study. Sediment thickness corrections have been applied in the manner described in the text, Polynomials constrained to reduce their gradient with increasing range and minimize their deviation from the data points are shown. Note the lack of high phase velocities: arrivals picked beyond - 35 km are probably super-critical moho reflections.

9 Seismic structure of old Atlantic crust 123 phase, refracted through the moho or upper mantle is extremely weak and is in fact only observable on two of the profiles. Thus this travel-time interpretation alone can place no meaningful constraints on the total crustal thickness of this region nor on the upper mantle velocity. The travel-time curves shown in Fig. 5 are not well constrained. Shot spacings of km are completely inadequate to define unambiguously the high vertical velocity gradients present in the uppermost 2-3 km of the crust. One way to proceed with these data would be to apply one of the commonly available inversion methods (e.g. Kennett & Orcutt 1976; Dorman & Jacobson 1981) to the curves shown in Fig. 5 and show that the poorly constrained T-X curves result in a set of velocity-depth bounds that are insignificantly different from one another. With this dataset we thought it more enlightening to proceed as shown in Fig. 6 and compile all the data at the earliest opportunity, at the uncorrected T-X stage, and investigate the significance of any differences between the profiles at that stage before they are influenced by subjective curve fitting processes and the enforced assumptions of lateral homogeneity. The travel-time data for all ten profiles is shown combined in Fig. 6 both before and after correction for sediment thickness variations. In the range window 0-35 km a smooth curve may be fitted through the 183 corrected data points, constrained to reduce gradient with increasing range with a root mean square deviation of s. This we consider to be within the bounds of uncertainty due to the limited bandwidth of the data, errors in knowledge of, and correction for basement topography, and basic reduction errors due to, for example, inaccurate shot instant corrections. Thus we conclude that on the basis of travel-time data alone there exists no detectable difference in' 4."1 h t 4.0 ' - : /,..:...,...,..:..., '_.. *, ".....,. i 4.O ~ RANGE lkrnsl ~Wre 6. CompiIation of the data shown in Fig. 5. (a) shows the data before correction for sediment thickness variations beneath the shots and receivers and (b) shows the data after these corrections were applied in the manner described in the text. The polynomial constrained to decrease gradient with range fits the 183 data points that lie at ranges of less than 35 km with a root mean square deviation of s.

10 124 G. M. Purdy 5 VEL oc/ i- Y (kms/secj TRAVEl TIME INVERSION Figure 7. The inversion of the polynomial curve out to 35 km range shown in.fig. 6(b) results in the velocity-depth function shown by the dashed lines. The sediment and water column thicknesses and velocities are those of the datum model used for the sediment thickness corrections. The inversion of the travel-time curve results in a velocity-depth function which can be most simply represented by three layers with linear gradients of 1.1, 0.64 and 0.1 s-l and thicknesses respectively of 0.4, 1.9 and 1.7 km. The velocity at the top of layer 2 is 5 km s-'. The solid lines are the compressional ( Vp) and shear ( Vs) wave velocity-depth functions used to compute the reflectivity synthetic seismograms shown in Fig. 9, and represent our best solution. The shear wave velocity depth function represents a constant Poisson's ratio of 0.28 throughout the hard rock crust. The shear wave velocity in the sediment column was assumed to be 0.5 km s-l. i "PA structure beneath the ten refraction profiles and the best estimate of the mean crustal structure in this region should be determined from an inversion (Purdy 1982a) of the curve shown in Fig. 6(b). The resulting best solution for the velocity-depth function is shown in Fig. 7. The T-X curve was truncated at 35km range to avoid the mistaken use of moho reflection times in the inversion. In order to make an estimate of total crustal thickness in this region from travel times alone we use the normal incidence moho reflections observed on the IPOD/USGS multichannel profrle (Grow & Markl 1977) which passes within 2 km of this refraction data. The

11 at Pennsylvania State University on September 19, 2016 Downloaded from Figure 8. A section of the IPOD/USGS multichannel reflection profile (Grow & Mark1 1977) that passes over the seismic refraction line described in this paper (see Fig. 2 for shot point locations). The discontinuous reflector at approximately 10 s is interpreted to be the moho. (It is not understood why there is not better correlation between the topography of the basement and the topography of this reflector.) location of this track relative to our experiment is shown in Fig. 2 and the relevant section of the multichannel profile is shown in Fig. 8. If the reflection at s two-way time is associated with the moho discontinuity, and if layer 3 is assumed to have a constant velocity of 7.0 km s-l, then the total hard rock crustal thickness we infer is 7.2 km (Fig. 7). This is simply a result of computing the two-way reflection time through the velocity model determined from the refraction data, then calculating the additional thickness of 7.0 km s- material required to increase that reflection time to that observed in Fig. 8. This is an uncertain estimate but does provide a useful starting point for the amplitude modelling that will be described in the next section. The lookm line: amplitude interpretation There are a number of clear and consistent amplitude patterns observed on the refraction profiles which are invaluable in defining the details of the velocity depth structure shown in Fig. 7. A combination of WKBJ method (Chapman 1978) and reflectivity method (Fuchs & Muller 1971) synthetic seismogram models were used to converge on the best solution (also shown in Fig. 7). In Fig. 9 we show the reflectivity method synthetic seismograms computed for this model compared with the observed seismograms for shots south east of OBH4. The choice of profile 4E for this comparison was arbitrary. There are five specific features of the velocity-depth function that are defined by the amplitude pattern: (1) The lack of refracted energy from the moho transition zone or upper mantle is a prominent feature of all the refraction profiles. This is not an artefact of poor signal-tonoise ratio data. The ambient noise level in this area (determined by comparison to the instruments 10Hz calibration pulse, Koelsch & Purdy 1979) is higher than that we have experienced elsewhere in the world s oceans: - 2 pbar at 10 Hz compared with < 1 pbar (e.g. Purdy & Detrick 1978; Detrick & Purdy 1980). However, although the majority of the shots were only 10.9 kg (241b) in weight there are several instances where kg (2401b) were detonated within - 40km of a receiver. This lack of moho and upper mantle refracted energy requires that there be both zero vertical velocity gradient in the uppermost few kilometres of the upper mantle and a moho transition zone sufficiently thin to prevent significant lateral propagation within it. The upper mantle P,, phase has been widely observed in the western central Atlantic Ocean but generally at ranges greater than Ewing & Ewing 1959). One possible explanation may be that our experiment did not

12 126 G. M. F'urdy RANGE /kmsl Figure 9. A comparison of one of the observed record sections with synthetic seismograms calculated using the velocity-depth function shown in Fig. 8. The synthetic seismograms were calculated using the reflectivity method (Fuchs & Muller 1971) with a phase velocity window of kms-' and a frequency bandpass of Hz. The source signal used was an analytically determined wavelet 0.3 s long centred about a frequency of 7 Hz. The reflectivity zone started at the top of layer 2. The travel-time curves were computed from the velocity-depth functions in Fig. 7. Amplitudes of both observed and synthetic seismograms have been weighted for range. The observed seismograms have not been corrected for sediment thickness variations. extend to a sufficient range to allow observation of energy turning in a small gradient that begins a few kilometres beneath the moho. (2) At ranges of about 25 km the moho reflection (PmP) and the crustal refraction (P3) are clearly observed as two discrete arrivals separated in time by s. This separation is most straightforwardly controlled by the total thickness of layer 3 and thus is particularly helpful in constraining the total crustal thickness. A particularly thick layer 3 (- 4km) is required to produce this discrete separation of the P3 and PmP phases at this range. This amplitude characteristic is observed in nine of the available refraction profiles, the exception being OBHl east. (3) The relative amplitude of these same P, and PmP phases together with the amplituderange relationship of the PmP phase are particularly diagnostic of the lower crustal velocity structure. The amplitude-range relationship of the PmP phase is most simply modelled using a moho transition zone less than 500 m thick. This is consistent with, but not necessarily

13 Seismic structure of old Atlantic crust 127 required by the nearby observation of normal incidence reflections from the moho on the IPOD/USGS multichannel reflection profile (Fig. 8). If the thickness of the moho transition zone is increased to, say, 1 km then the triplication point maximum amplitude moves to too great a range. This range can be reduced, of course, by thinning layer 3 but this is inconsistent with the previously noted s separation of P3 and PmP at - 25 km range. If the nature of the moho transition is fixed, then the relative amplitude of the P3 and PmP phases at this range (- 25 km) is then primarily controlled by the vertical velocity gradient in layer 3. Modelling showed that a gradient of 0.1 s-' was far too great in this regard. An extremely small gradient was necessary within layer 3 severely to reduce the P3 amplitude relative to P,P at this range. Our preferred model has zero velocity gradient below the 7.Okms-' velocity contour. (4) The single most variable amplitude characteristic of the ten profiles being studied was the amplitude of the converted shear phases (Fig. 4). Their phase velocity is impossible to measure with any great confidence but is in the range 3.6 to 3.8km s-'. Their travel times are consistent with conversion P to S and S to P at the sediment-basement interface beneath the shot and receiver respectively but the possibility of this conversion taking place at the top of the limestones just above the basement cannot be ruled out. On some profdes (e.g. OBH 8E) they were strong and coherent but on others they were hardly usable (e.g. OBH 4W). Significant shot-to-shot variation in amplitude is observed too, notably on OBH 4E. We explain this variability in terms of small changes in the velocity structure in the uppermost loom of layer 2 which in turn control the conversion efficiency P to S beneath the shots or S to P beneath the receivers. White & Stephen (1980) show that changes in the nature of this transition zone at the top of the basement on the scale of a few tens of metres have a dramatic effect on the efficiency of P to S conversion (or vice versa), but basement topography must, of course, also play a role in this mechanism. Models with a simple discontinuity at the sediment-basement interface resulted in converted shear wave amplitudes three to four times greater than those observed and two to three times greater than the first arrival P-wave. However, by introducing a loom thick transition zone as shown in Fig. 7, the converted shear wave amplitudes are considerably reduced to a level similar to some of the observations. The variability in amplitudes of the shear phase from profile to profde and from shot to shot (and our lack of knowledge of Poisson's ratio in the uppermost crust) make it impossible for us to define the nature of this transition zone in any detail. We can say little more than the basement-sediment interface is not a first order discontinuity but is a m thick transition zone that is laterally variable in velocity structure. This variability in conversion efficiency disallows any use of the shear wave amplitude distribution to constrain the shear wave velocity depth profile. One exception to this may be the following general observation. (5) The last feature of the amplitude patterns, though not as common a characteristic as those previously mentioned but still worthy of note, is the relative amplitude of the S3 and SmS phases at ranges of about 30 km. Where adequate conversion has taken place beneath shot and receiver to allow these phases to be observed then, as in the case of the P-waves, the SmS amplitude is many times that of S3 (Figs 4 and 9). Thus the shear wave velocity gradient in layer 3 must also be zero or very small. Though not well established by these data, this would suggest that Poisson's ratio remains approximately constant through layer 3 and into the upper mantle. These five paragraphs are our primary justification for representing the velocity-depth function in Fig. 7 as our best estimate of the crustal structure in this region. It has a total hard rock crustal thickness of 7.2 km and is characterized by a 2.3 km thick layer 2 topped by a thin l00m transition zone but otherwise with velocity gradients of 1.1 and 0.6s-',

14 128 G. M. Purdy overlying a 4.9km thick layer 3 with a velocity gradient less than 0.1 s-l, overlying a sharp moho (< 500m thick) that marks the transition into an 8.2 km s-l upper mantle that has no velocity gradient in its uppermost few kilometres. In the previous section, it was shown that travel-time data could not resolve any significant changes in structure along the refraction line. Here we have shown that the amplitude patterns are consistent to a large degree and perhaps suggest the existence of a significantly homogeneous structure. The following section will investigate the possibility of lateral homogeneity on a smaller scale. The circle experiment The experiment was designed to determine the difference in delay time beneath a cluster of five OBH instruments located within a - 6 km square. No absolute determinations of structure are possible from these data but they provide a precise view of the variability invelocity structure of the uppermost - 3 km of the crust over a small area. The shot-receiver configuration used to collect these data has been shown in Fig. 3. Fig. 10 shows the seismograms of the circle shots as received by OBH4. The range to this receiver changed by only - 1 km throughout the experiment as can be qualitatively appreciated from the small changes in direct water wave travel times observed on this record section. However, the waveform of the first refracted arrival is seen to vary considerably around the circle, as is the amplitude of the converted shear phase observed s after the first arrival. The analysis of these data is based on the assumptions of delay time theory. Let the total travel time from the ith shot " 090" 100' 270" AZIMUTH OF SHOT FROM R C IVER Figure 10. Circle shots as received by OBH 4. The seismograms are unfiltered and are plotted against azimuth of the shot from the receiver. Shot numbers are shown (see Fig. 3 for locations), their charge sizes being a constant 10.9 kg (24 lb). Shot numbers 3456 and 3457 failed to detonate (hence the gap at - 030" azimuth). The large amplitude direct water wave arrival overloads the analogue recording system of the OBH instrument at s travel time. The waveform of the fist refracted arrival at -7.5 s travel time is seen to be highly variable from shot to shot even though the shot-receiver range changes by less than 1 krn throughout the experiment. We interpret this to be due to the focusing and defocusing effects of the basement topography (Purdy 1982a). The arrival at -10 s is the converted shear phase propagating in layer 3 (SJ.

15 to the jth receiver be represented by: Seismic structure of old Atlantic crust 129 where xij is the shot receiver separation, V is the refractor velocity, rj is the delay time beneath the shot and rj is the delay time beneath the receiver. It is our intention here to compute the delay time differences between pairs of receivers using: Tj+i - Tij + (Xij - Xjj +,)/V. Equation (2) can, of course, be solved for each shot around the circle and the five OBH receivers result in ten possible receiver pair combinations. The results of these calculations are presented in Table 2, but before these are described and discussed in more detail it is important to review the assumptions involved with the use of equation (2). Delay time theory assumes a structure to consist of a stack of isovelocity layers. We have seen in the previous section that this is not in fact a poor assumption for energy received in the range window km because the velocity gradient in layer 3 is negligibly small. Precise knowledge of what the velocity V should be for the travel paths used in the calculations is difficult to obtain. This is one of the rationales for shooting in a circle: it provides many observations in which the value of (Xjj-Xjj t 1) is minimized (this being the only place where V enters into the calculation). Equation (2) also contains the assumption that the delay time beneath the ith shot is constant for all receivers. The validity of this assumption is dependent primarily on the scale of roughness of the basement topography beneath the shot. Because equation (2) is concerned with the difference in travel time of one shot to two receivers then it is immune to many common errors, most notably those involved with water depth corrections and shot instant corrections. Using all 42 shots to the ten possible receiver pair combinations a total of 420 delay time differences may be calculated. Although the range differences between receiver pairs were at most a few kilometres, the refractor velocity was still an important unknown in these calculations. If the incorrect velocity is used in the computation of these differences then it introduces an apparent azimuthal dependence of the delay differences with the maximum 5 Table 2. Results of the analysis of the circle shots as described in the text. For each of the ten possible OBH pair combinations this table lists the mean range of the N shots to the two instruments. The rms deviation of this range provides a qualitative estimate of the sensitivity of the computed delay time difference to the chosen value of phase velocity. The listed mean delay time differences were calculated using the phase velocity that minimized the rms deviation about the mean. This phase velocity was either 6.4 or 6.5 km SKI. OBH Mean UrmS Delay arms N pair range range diff. delay O<m) (km) (ms) (ms)

16 130 G. M. Purdy 31" 22" 31"20' I I 68"18'W 68"16' 68" 14' 68"12' 68" 10' 31" 18' Figure 11. Detailed view of location of OBH instruments used in circle experiments, and the deep towed hydrophone reflection profile tracks used to make determination of sediment thickness beneath the receivers. The fine lines are the ship track as determined by smoothed 1 min interval Loran C fixes and the thicker lines are the deep towed hydrophone track annotated at each hour. The data from the east-west line are shown in Fig. 12 (from Purdy & Gove 1982). occurring along the azimuth of maximum range difference. For each receiver pair the velocity was computed that minimized this azimuthal dependence (or, in fact, the deviation from the mean). For each receiver pair separately the shots that gave delay time differences greater than one standard deviation from the mean were rejected and a new mean and rms deviation for that receiver pair were calculated using the depleted dataset (Table 2). This was an arbitrary step that we preferred to any decimation of the dataset based on subjective judgements of data quality. The intention was to reduce the error estimate for delay time differences beneath the receivers to a minimum by imposing the assumption that the largest '00 Figure 12. Deep towed hydrophone reflection profile that passes over OBH 8, 4 and 6. These data have been corrected for hydrophone motion, deconvolved for the source signature and stacked five fold simply by summing consecutive shots. Vertical exaggeration is approximately 4:l. The common western central Atlantic reflectors At, A, and A* (Tucholke et al. 1979) are indicated as are the basement identifications beneath the three instruments. The sediment thicknesses beneath OBHs 8, 4 and 6 determined from this profile were respectively 0.58, 0.78 and 0.9 s two-way reflection time (from F'urdy & Gove 1982).

17 Seismic structure of old Atlantic crust 131 deviations from the mean are caused by shot dependent factors (e.g. extremely steep scarp in the basement beneath a shot or poor arrival time pick) and not to true structural differences beneath the receivers' random bandwidth-limited arrival time errors. This step reduced the rms deviation by approximately a factor of two. A crude measure of the relative quality of the delay time difference determination in terms of their dependence upon V may be gained from the mean range differences listed in Table 2. A generalized summary of the results in Table 2 is that no resolvable and significant difference exists in the upper crustal structure beneath instruments 3,4 and 5. OBH 6 has an upper crustal delay time that is s greater than this group and OBH 8 has an upper crustal delay time that is s smaller than this group. As our primary interest is in the hard rock crustal structure, and as these delay times include the sediment column, then a necessary first step in attaching significance to these results is the accurate determination of sediment thickness beneath the receivers. For this we use an airgun reflection profile collected with a hydrophone towed -100m above the seafloor (Purdy, Ewing & Bryan Table 3. The sediment thicknesses beneath the OBH instruments in seconds of two-way travel time were: OBI3 8 = 0.58 f 0.02; OBH 4 = 0.78 t 0.02; OBH 6 = 0.90 f 0.02; OBH 3 = 0.80 f 0.04; OBH 5 = 0.08 f All determinations with the exception of OBH 3 were made from deep tow hydrophone reflection profiles (see Fig. 12). The errors attributed to these thicknesses are subjective estimates. This table lists the observed sediment thickness differences for the ten instrument pairs in milliseconds of two-way reflection time. It also lists sediment thickness differences computed under the following assumptions: (1) The delay time differences listed in Table 2 are wholly attributable to differences in sediment thickness. That is, these computed thickness differences are those that would, by themselves, result in the observed delay time differences. (2) The compressional wave velocities are 2.0 km s-' for the sediment, 5.0 km s-' for the basement, and the phase velocity used was 6.5 km s-i. It can be seen that the observed and computed differences agree, within their errors. The errors in the values of computed sediment thickness are these resulting from the delay time difference rms deviation shown in Table 2. OBH Pair Sediment thickness difference (ms) Observed Computed -1OOf 45-12Of 28-2Of f f f f lo? t f f f * f f f f 29

18 132 G. M. Purdy 1980; Purdy & Cove 1982) that passes over four of the OBH instruments. The track is shown in Fig. 11. The research vessel was navigated by smoothing 1 min interval Loran C fixes (Purdy & Cove 1982). (Confidence in the relative accuracy of heavily averaged Loran C data in this area was gained from the comparison of the OBH instrument separations determined from these data with those determined from water wave travel times. The discrepancies did not exceed 150 m.) The hydrophone track shown in Fig. 11 was obtained by further smoothing of the ships track and calculation of the horizontal separation of airgun and hydrophone as described by Puray & Cove (1982). As the point of reflection on the basement is m distant from the hydrophone location but perhaps as much as 1 km astern of the source, it is the hydrophone track that is used for making sediment thickness determinations. In Fig. 12 the reflection profie is shown after corrections for hydrophone motion were made, source deconvolution and a five shot stack were carried out (Purdy & Cove 1982). The location of OBHs 8,4 and 6 are shown on the profile. Data of similar quality were collected over OBH 5 but unfortunately not over OBH 3 in which case conventional surface reflection data were used. The best estimates for sediment thickness beneath each of the five instruments are given in the caption for Table 3. An important question is whether any part of the delay time differences shown in Table 2 can be accounted for by the sediment thickness differences. Table 3 compares the sediment thickness differences observed on the seismic reflection profiles with those computed from the delay time differences: that is, the computed thickness differences are those that would, by themselves, result in the observed delay time differences. It can be seen that, within their errors, the observed and computed differences agree. Thus it can be concluded that the only resolvable difference in shallow crustal structure beneath the five OBH instruments is the variation in sediment thickness, which in this region is determined almost exclusively by the basement topography. Discussion and conclusions The results described in the previous sections of this paper provide conclusions in three general areas: (1) Crustal vezocity structure: the principle components, beginning with a 5 km s-' velocity at the top of layer 2, are a 0.4 km thickness with gradient 1.1 s-l, a 1.9 km thickness with gradient 0.64s-', a 1.7km thickness with a gradient 0.1 s-l, a 3.2km thickness with zero gradient and constant velocity of 7 km s-' and 0.5 km thick moho transition zone (see Fig. 7). The uppermost few kilpmetres of the upper mantle has little or no gradient. At the very top of layer 2 a - loom thick transition zone with a very steep velocity gradient (10-20 s-') is required to reduce the efficiency of compressional to shear wave conversion at this boundary. This could also be achieved by introducing a thin zone of considerably greater Poisson's ratio. Classically the uppermost two linear gradient layers listed above would be attributed to layer 2 yielding a total thickness of 2.3 km and the lowermost two to layer 3 which would then be 4.9km thick: both are comfortably within the limits of Raitt's average structure (Table 1). The total thickness of this 140Myr old igneous crust is thus 7.2km, this being overlain by an average 700m of sediment. The two-way reflection time through this model agrees to better than 0.1 s with the location of the moho reflection seen as the IPOD/USGS multichannel reflection profile (Fig. 8). The travel times of the converted shear wave arrivals constrain Poissnn's ratio in the ieneous crust to f (2) Homogeneity ojcrustal Structure: it was shown in Fig. 6 that the travel times from all shots along the lookm long explosives line as received by all five receivers could be com-

19 Seismic structure of old Atlantic crust 133 bhed within their errors to give a single velocity-depth function. This by itself does not prove the existence of a laterally homogeneous or ordered structure. It does show, however, that such a simple solution is valid. An equivalent statement would be that the travel-time data do not prove the existence of structural heterogeneity. Secondly, listed in the text are five characteristics of the amplitude distribution in travel time and range that are common to all ten record sections. One velocity-depth function is determined from which synthetic seismograms are calculated that agree well with all amplitude and travel-time observations. Thus the amplitude data do not support the existence of structural inhomogeneity. Thirdly, the circle experiment shows that the differences in shallow crustal delay time between five receivers located within a 6 km square can be wholly attributed to differences in sediment thickness beneath the receivers. The above three statements are valid conclusions only because of the extensive and precise nature of the dataset used in this paper, for example, only the deep-towed hydrophone reflection profiles gave sufficiently accurate measurements of sediment thickness beneath the receivers to provide an explanation for the delay time differences. If we had relied upon conventional surface reflection data or even upon poorly navigated deep-tow data then our conclusions would have been different. Similarly, if only two or threerecord sections were available instead of the ten presented here, then there would have been little justification in doing other than interpreting them separately to produce discrete, if arguably different, velocity-depth functions. Laterally homogeneous, though in common usage, is a poor adjectival phrase to apply to the seismic velocity structure of oceanic crust under any circumstances: there always exists the primary inhomogeneity that is basement topography. On 140Myr old crust this rugged interface, across which exists the largest velocity contrast in all of the crustal column, is neatly buried beneath a deceptively flat and sedimented seafloor. Purdy & Gove (1982) show that the basement topography in this region has amplitudes comparable with that of the flanks of the present day mid-atlantic Ridge. Thus both the amplitude and travel-time effects of this interface must be carefully considered. Purdy (1982a, b) suggested that this topography plays an important role in the focusing and defocusing of energy from sea surface shots and we propose here that the shot to shot amplitude variability on the circle data (Fig. 10) and the small variations from record section to record section on the lookm line data (Fig, 4) can be attributed to this effect. A best attempt is made to correct for the travel-time effects of this topography but the available methods are inadequate, our knowledge of basement topography beneath the shots using conventional surface reflection profiles is inadequate and thus as described in Purdy (1982a, b) the accuracy of the corrections is poor and impossible to estimate quantitatively. Thus we attribute a large part, if not practically all of the travel-time scatter observed in Fig. 6(b) to residual topographic effects which the poor (but best available) correction method failed to remove. The circle experiment results are travel-time differences, and in consequence are insensitive to basement or seafloor topography beneath the shots, and thus do not exhibit this same scatter. The conclusion that follows from these arguments is that this dataset reveals no evidence for lateral structural inhomogeneity other than that specifically related to basement (layer 2) topography. A further reason why laterally (in)homogeneous is a poor term to apply to oceanic crust that it contains no statement of the scale of the structure to which the phrase refers. There exists incontrovertible proof from submersible observations and drilling results that on a scale of centimetres to metres the uppermost section of oceanic crust is grossly inhomogeneous (ward et el. 1975; Aumento et al. 1977). The ophiolite analogy (though of questionable

20 134 G. M. Purdy validity in this case due to the possibility of complexities introduced during and after emplacement) suggests the existence of inhomogeneities on the scale of tens to hundreds of metres throughout the crustal column (e.g. Casey et al. 1981). The data and conclusions presented here do not contradict these geological observations. Our data contain information only on the homogeneity or inhomogeneity of structure 'on the seismic scale'. The traveltime and amplitude interpretations described in this paper are based on the propagation of km wavelength seismic energy along travel paths tens of kilometres in length. For a structural inhomogeneity within the crust to be detectable and definable with these data it must have sufficient vertical extent compared with a wavelength so that when it exists at the turning point of the seismic energy, the propagation of the seismic energy is distorted sufficiently to modify the amplitude pattern; the inhomogeneity must have sufficient lateral extent compared with the seismogram spacing that this resulting amplitude anomaly can be detected and defined; the inhomogeneity must cause a travel-time anomaly large compared with the data bandwidth and the reduction errors. Thus the seismic scale in this case is a scale of kilometres: a different world from that observable from submersibles or with drilling and thus we justify the statement that these results do not contradict known geological fact. The ophiolite analogy is more difficult to completely reconcile with our findings due to the presence of larger-scale inhomogeneities in some of these units (e.g. Karson & Casey 1981 ; Casey et al. 1981). However, this is not necessarily considered a contradiction of our findings due to the unavoidable uncertainties in the details of the tectonic environment of the ophiolite crust at the time of its formation. (3) Towards a simpler view of the velocity structure of oceanic crust: the conclusion that the crustal structure in this region can be described by a single velocity-depth function is due to three factors: (a) Experiment location: because of the dense aeromagnetic data coverage and precise tectonic framework available for this region (Schouten & Klitgord 1982) the experiment could be located away from the structural changes related to fracture zones (Detrick & Purdy 1980; Schouten & White 1980) or the possibility of structural anomalies associated with reorganizations in relative plate motions. Secondly, because the crust is old (- 140Myr) the uppermost section of layer 2 is well sealed and thus exhibits high velocities (5 km s-') and low velocity gradients (0.6-1.I s-l) compared with young crust (Houtz & Ewing 1976; Whitmarsh 1978; Ewing & Purdy 1982). If isovelocity contours within the shallow crust are conformal with the basement topography (Purdy 1982a) then these low gradients result in negligibly small lateral changes in velocity along the ray path. This was shown to be the case for young Pacific crust with high velocity gradients (- 3 s-') but low amplitude (- l00m) topography (Purdy 1982a, fig. 18) and similarly may be shown to be the case here with high amplitude topography (- 400 m) but low velocity gradients (< 1 s-l). The point here is that if we carried out an identical study over the young crust of the mid-atlantic Ridge flanks where the basement topography is of comparable amplitude but the shallow crustal velocity gradients are significantly greater, then our result may apparently contradict the conclusions here (that the crustal structure can be described by a single velocity-depth function). We speculate, however, that this apparent disagreement would be due only to the lateral velocity changes that, because of a combination of high velocity gradient and large amplitude topography, would be sufficiently large to invalidate any solution in terms of laterally homogeneous layers: in truth the crustal structure could still be simple and be completely described by a single velocity-depth function (with its origin at the seafloor) and the basement topography profile (Purdy 1982a). (b) Experiment design : apparent structural inhomogeneities are quickly and easily introduced by an interpreter into solutions based on sparse data, poorly navigated relative to the

21 Seismic structure of old Atlantic crust 135 sement topography and inaccurately corrected for its travel-time effects. The critically iportant components of the experiment were the accurate relative navigation provided by ioothed 1 min interval Loran C positions; five receivers fixed in position throughout the ooting over a sediment column the thickness of which was precisely determined by the deep #wed hydrophone reflection profile (Fig. 12); conventional reflection profiling along the,ooting line to provide reasonable estimates of sediment thickness; constant shot charge size as r as possible to improve reliability of synthetic seismogram modelling (Fig. 9); a total of most 200 data points in the range window 0-35 km reliably to define the travel-time curve 3g. 6) in spite of the unavoidable errors involved with correcting for the basement toporaphy ; and lastly the presence of moho reflections on the IPOD/USGS multichannel reflecon profile (Fig. 8) added greatly to the confidence in the correctness of the final velocityepth function (Fig. 7). (c) Scale: as was mentioned previously because of the physical limitations of the convenional seismic method we can only be concerned with features on the seismic scale. On this cale of a few kilometres the crustal structure in this region appears laterally homogeneous. f the resolution and bandwidth of the seismic experiment were increased sufficiently then he known (drilling-submersible- ophiolite) small-scale inhomogeneities would eventually be letectable. However, this may prove to be a specific disadvantage if the primary interest lies n the definition of the longer-term systematic changes in crustal velocity structure that may )e related to fundamental processes (crustal age, spreading rate, tectonic environment, etc.). fie seismic scale may in fact prove to be an extremely helpful low-pass filter that allows ;tudy of the crust on a scale large enough to permit the definition of the important longterm changes whilst filtering out the effects of smaller-scale structures related to the random volcanic and tectonic processes of crustal emplacement. We propose that the results presented in this paper are in agreement with a model of the ocean basins that consists of ribbons of normal crust formed at discrete ridge segments and separated by fracture zone structures. The velocity-depth structure of the ribbon crust on the seismic scale can be described by a single function that is dependent in a systematic way upon age (and other primary processes). This simple model is not supported by the historical seismic refraction dataset because a large proportion of the old experiments were poorly located relative to fracture zones and other structure anomalies (unknown at the time) and the density and sediment thickness information available for each experiment was insufficient to overcome the errors introduced by the basement topography. Acknowledgments The majority of the data reduction for this paper was expertly carried out by Nina Lian. Ralph Stephen provided the synthetic seismogram programs and careful advice on their correct application. The ocean bottom hydrophone instruments were designed, built and operated by Donald Koelsch and Carleton Grant. Successful data acquisition was due to the officers, crew and scientific complement of R/V Knorr during Cruise KN77-2. Fig. 1 was compiled by Kristin Rohr. This manuscript benefited from critical reviews by John Ewing, Ralph Stephen, and Charles Denham. The research described in this paper was supported by the National Science Foundation under grant OCE Woods Hole Oceanographic Institution Contribution Number References Aumento, F. et al., Init. Rep. Deep Sea drill. Proj., 37, Washington, US Government Printing Office.

22 136 G. M. Purdy Ballard, R. D., Bryan, W. B., Heirtzler, J. R., Keller, G., Moore, J. G. & Van Andel, Tj., Manned submersible operations in the FAMOUS area: Mid Atlantic Ridge, Science, 190, Casey, J. F., Dewey, J. F., Fox, P. J., Karson, J. A. & Rosencrantz, E., Heterogeneous nature of oceanic crust and upper mantle: a perspective from the Bay of Islands ophiolite complex, in The Sea, 7, , ed. Emiliani, C., Wiley (Interscience), New York. Chapman, C. H., A new method for computing synthetic seismograms, Geophys. J. R. asb. SOC., 36, Christensen, N. I. & Salisbury, M. H., Structure and composition of the lower oceanic crust, Rev. Geophys. Space Phys., 13, Detrick, R. S. & F'urdy, G. M., The crustal structure of the Kane Fracture Zone from seismic refraction studies, J. geophys. Res., 85, Dorman, L. M. & Jacobson, R. S., Linear inversion of the body wave data - part 1 ; velocity structure from travel times and ranges, Geophysics, 46, Ewing, J. & Ewing, M., Seismic refraction measurements in the atlantic ocean basins, in the Mediterranean Sea, on the Mid-Atlantic Ridge, and in the Norwegian, Bull. geol. SOC. Am., 70, Ewing, J. & Purdy, G. M., Upper crustal velocity structure in the ROSE area of the East Pacific Rise, J. geophys., Res., in press. Ewing, M., Sutton, G. H. & Officer, C. B., Jr, Seismic refraction measurements in the Atlantic Ocean, part VI: typical deep stations, North America Basins, Bull. seism. SOC. Am., 44, Fox, P. J., Schreiber, E. & Peterson, J. J., The geology of the oceanic crust: compressional wave velocities of oceanic rocks, J. geophys. Res., 78, Fuchs, K. & Muller, G Computation of synthetic seismograms with the reflectivity method and comparison with observations, Geophys. f. R. astr. SOC., 23, Goslin, J., Beuzart, P., Francheteau, J. & Le Pichon, X., Thickening of the oceanic layer in the Pacific Ocean,Mar. geophys. Res., 1, Grow, J. A. & Markl, R., IPOD-USGS multichannel seismic reflection profile from Cape Hatteras to Mid-Atlantic Ridge, Geology, 5, Hersey, J. B., Bunce, E. T., Wyrick, R. F. & Dietz, F. T., Geophysical investigation of the continental margin between Cape Henry, Virginia, and Jacksonville, Florida, BuZ1. geol. SOC. Am., 70, Houtz, R. & Ewing, J. I., Detailed sedimentary velocities of the Western North Atlantic margin, Bull. seism. SOC. Am., 68, Houtz, R. & Ewing, J. I., Sedimentary velocities of the Western North Atlantic margin, Bull. seism. SOC. Am., 54, Houtz, R. & Ewing, J. I., Upper crustal structure as a function of plate age, J. geophys. Res., 81, Karson, J. A. & Casey, J. F., Lateral variations in the internal structure of the Bay of Islands ophiolite complex (abstract), Eos, 62, Katz, S. & Ewing, M., Seismic refraction measurements in the Atlantic ocean part VII: Atlantic ocean basin, west of Bermuda, Bull. geol. SOC. Am., 67, Kennett, B. L. N. & Orcutt, J. A., A comparison of travel time inversions - Marine refraction profiles, J. geophys. Res., 81, Koelsch, D. E. & Purdy, G. M., An ocean bottom hydrophone instrument for seismic refraction experiments in the deep ocean,mar. geophys. Res., 4, Larson, R. L. & Hilde, T. W. C., A revised time scale of magnetic reversals for the Early Cretaceous and Late Jurassic, J. geophys. Res., 80, Ludwig, W. B., Seismic refraction, in The Sea, 4, ed. Maxwell, A. E., Wiley, London. Officer, C. B. Jr, Ewing, M. & Wuenschel, P. C., Seismic refraction measurements in the Atlantic ocean, Part IV: Bermuda, Bermuda Rise, and Nares Basin, BUZZ. geol. SOC. Am., 63, F'urdy, G. M., 1982a. The variability in seismic structure of Layer 2 near the East Pacific Rise at 12"N, J. geophys. Res., in press. Purdy, G. M., 1982b. The correction for the travel time effects of seafloor topography in the interpretation of marine seismic data, J. geophys. Res., in press. Purdy, G. M. & Detrick, R. S., A seismic refraction experiment in the Central Banda Sea, J. geophys. Res., 83, Purdy, G. M., Ewing, J. I. & Bryan, G. M., A deep towed hydrophone seismic reflection survey around IPOD Sites 417 and 418,Mar. Geol., 35, 1-19.

23 Seismic structure of old Atlantic crust 137 Purdy, G. M. & Gove, L. A., Reflection profiling in the deep ocean using a near bottom hydrophone, Mar. geophys. Res., in press. Purdy, G. M.. Gove, L. A., Koelsch, D. E. Power, G. & Allison, M. D., The Woods Hole Oceanographic Institution system for reduction and handling of marine seismic refraction data, unpublished WHOI Tech. Rep., in preparation. Purdy, G. M. & Rohr, K., A geophysical survey within the Mesozoic magnetic anomaly sequence south of Bermuda, J. geophys. Res., 84, Raitt, R. W., The crustal rocks, in, The Sea, 3, , ed. Hill, M. N., Wiley (Interscience), New York. Rohr, K. & Purdy, G. M., Sediment and upper basement velocities from wide angle reflection data, Eos, 62, 1035 (abstract). Schouten, H. & Klitgord, K. D., Map showing Mesozoic magnetic anomalies, Western North Atlantic, MAPMF 915, United States Geological Survey, Reston, Virginia. Schouten, H. & Klitgord, K. D., The memory of the accreting plate boundary and the continuity of fracture zones, Earth planet. Sci. Lett., 59, Schouten, H. &White, R. S., Zero-offset fracture zones, Geology, 8, Sheridan, R. E., Drake, C. L., Nafe, J. E. & Hennion, J., Seismic-refraction study of Continental margin east of Florida, Bull. Am. Ass. Petrol. Geol.. 50, Spudich, P. & Orcutt, J., A new look at the seismic velocity structure of the oceanic crust, Rev. Geophys. Space Phys., 18, Tucholke, B. E. et al., hit. Rep. Deep Sea drill. Proj., XLIII, Washington, US Government Printing Office. White, R. S. & Matthews, D. H., Variations in oceanic upper crustal structure in a small area of the northeastern Atlantic, Geophys. J. R. astr. Soc., 61, White, R. S. & Stephen, R. A., Compressional to shear wave conversion in oceanic crust, Geophys. J. R. astr. Soc., 63, Whitmarsh, R. B., Seismic refraction studies of the upper igenous crust in the north Atlantic and porosity estimates for layer 2, Earth planet. Sci. Lett., 37,