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Geophys. J . R. astr. SOC.(1986) 86,491-513 and Microfiche GJ86ll Crustal structure of Atlantic fracture zones - 11. The Vema fracture zone and transverse ridge C. C. R. S. POttS and White BuHard Laboratories, Department ofEarth Sciences, Madingley Rise, Madingley Road, Cambridge CB3 OEZ K. E. Louden Department of Oceanography, Dalhousie University, Halifax, Nova Scotia B3H 4J1, Canada Accepted 1986January 29. Received 1986January 17; in original form 1985 September 10 Summary. The crustal structure beneath the Vema fracture zone and its flanking transverse ridge was determined from seismic refraction profiles along the fracture zone valley and across the ridge. Relatively normal oceanic crust, but with an upwarped seismic Moho, was found under the transverse ridge. We suggest that the transverse ridge represents a portion of tectonically uplifted crust without a major root or zone of serpentinite diapirism beneath it. A region of anomalous crust associated with the fracture zone itself extends about 20 km to either side of the central fault, gradually decreasing in thickness as the fracture zone is approached. There is evidence to suggest that the thinnest crust is found beneath the edges of the 20 km wide fracture zone valley. Under the fracture zone valley the crust is generally thinner than normal oceanic crust and is also highly anomalous in its velocity structure. Seismic layer 3 is absent, and the seismic velocities are lower than normal. The absence of layer 3 indicates that normal magmatic accretionary processes are considerably modified in the vicinity of the transform fault. The low velocities are probably caused by the accumulation of rubble and talus and by the extensive faulting and fracturing associated with the transform fault. This same fracturing allows water to penetrate through the crust, and the apparently somewhat thicker crust beneath the central part of the fracture zone valley may be explained by the resultant serpentinization having depressed the seismic Moho below its original depth. Key words: explosion seismology, fracture zones, ocean bottom seismometers, oceanic crust, seismic ray-tracing 1 Introduction Fracture zones are a prominent feature of the seafloor of ocean basins. They are characterized by a ridge-and-trough topography, typically 10-20 km wide, which can some- 492 C. G. Potts, R. S. WhiteandK. E. Louden times be followed for thousands of kilometres. The transverse ridges can be particularly spectacular features, and are often associated with large fracture zones (e.g. Romanche, Vema and S t Paul in the equatorial Atlantic). In addition to forming prominent bathymetric features, fracture zones are the major source of crustal seismic heterogeneity (Detrick & Purdy 1980; White et al. 1984). Francheteau et al. (1976) conclude that the high relief across large Atlantic fracture zones does not expose deep sections into the oceanic crust, but is generated by many small-throw faults. Submersible studies in the Kane fracture zone (Karson & Dick 1983) and the Oceanographer fracture zone (OTTER 1984, 1986) confirm the common occurrence of gabbroic and ultramafic rocks in the valley walls, which had led to the proposal that the crust beneath large fracture zones must be thinner than normal crust (Fox, Detrick & Purdy 1980; Fox & Gallo 1984). Several seismic refraction experiments have been shot in the Atlantic to determine the crustal structure in and around transform faults (White et al. 1984; Detrick, White & Purdy 1986). The crust beneath fracture zone valleys is typically thinner than normal oceanic crust, with low crustal velocities and a marked absence of the oceanic seismic layer 3. Little is known about the seismic or geological structure of the transverse ridge, and few experiments have provided good constraints on the width of the anomalous crustal region associated with fracture zones. 'The seismic refraction experiment described here provides seismic control over the crustal and subcrustal structure along the Vema fracture zone and across its associated transverse ridge. The most important results are the constraints on crustal structure beneath the transverse ridge to the south of the Vema fracture zone, and the nature and width of the anomalous crust below the fracture zone itself. 2 The Verna fracture zone The Vema transform (Fig, 1) offsets the Mid-Atlantic Ridge by 320 km with an age contrast of 26 Myr across the end of the spreading centre. Currently the plates on either side of the fault are moving with a relative velocity of 24 mm yr-' . The transform valley is typically 25-40 km wide between the crests of the north and south walls. In places the top of the southern ridge is only 550 m below sea-level, and may recently have been at sea-level (Bonatti & Crane 1982). Dredge samples indicate that the north wall of the Vema, which has a relatively subdued topography, is probably a disrupted section through an apparently representative segment of normal oceanic crust; the southern transverse ridge is interpreted from the samples obtained as being an uplifted section of crust (Bonatti & Honnorez 1976; Bonatti & Crane 1982; Honnorez, Mevel & Montigny 1984). Material believed to be from oceanic layer 3 on the lowest part of the inward facing wall of the transverse ridge appears to have been reworked and altered, in situ. Gravity and magnetic intensity profiles across the active transform of the Vema at around 44"W and 43OW have been modelled by Robb & Kane (1975). They propose a structure which includes dense, ultrarnafic bodies at shallow depths below the walls of the fracture zone. This solution has been challenged by other authors (e.g. Sibuet & Veyrat-Peinet 1980; Louden & Forsyth 1982) but in the absence of other evidence, no more credible structure could be proposed. Ludwig & Rabinowitz (1980) reported the results of the first seismic refraction experiments in the Vema using free-floating sonobuoys (Fig. 2). No detailed geological Vema fracture zone and transverse ridge 493 20'N 10 N 0" 10ys . 200s 60'W SOUTH AMERICA 50'W 40"W 30°W 20aw 1oow Figure 1. The location of the Vema fracture zone and other major features of the Mid-Atlantic Ridge between Africa, and South America. The seismic refraction experiment was near the western ridgetransform intersection. interpretation was presented, but it was concluded that the seismic structure along the Vema fracture zone is not the same as that below apparently normal oceanic crust nearby, and is not easily reconciled with the velocity structure of deep-ocean crust observed elsewhere. 120 11'30 1' 0'30 00 Figure 2. The location of the shooting tracks from this experiment superimposed on the simplified bathymetry of Prince & Forsyth (1986). Lines A, B and C are described in this paper. Also shown by broken lines are Detrick et al. (1982) OBS line, the six sonobuoy experiments (SLF3-6 and R45, R48) of Ludwig & Rabinowitz (1980), and Line E2, the OBS refraction line along the spreading centre of the Mid-Atlantic Ridge shot o n the same cruise (White et a/. 1986). 494 C. G. Potts, R. S. White and K. E. Louden Velocity (km s-'1 2 4 6 8 10 2 4 n E Y v 6 c +Ir P 10 12 14 Figure 3. Velocity-depth functions determined for the Vema fracture zone by Detrick et QI. (1982) superimposed on the bounds for normal ocean crust greater than 3 M y r (from White 1984). Tau-p solutions from two OBSs are shown, with a plane layer solution from one instrument. Vertical scale is depth into igneous crust. (See Detrick ef al. 1982 for location of shooting track and receivers.) In 1978 a single explosive refraction line (Fig. 2) was fired into three OBSs (Detrick et al. 1982). The data were topographically corrected using sediment thicknesses estimated from single channel reflection profiles. Detrick et al. conclude that the crust beneath the fracture zone has lower seismic velocities than normal oceanic crust, and is probably thinner. From the velocity-depth functions (Fig. 3 ) it is also apparent that a typical 'layer-3' type region (with seismic velocity of approximately 6.7 km s-', and low velocity gradients) is missing, or extensively altered. The tau-p solutions from this experiment were used as a starting point for interpreting the data from the OBS refraction experiment described in this paper. 3 The seismic experiment The experiment discussed in this paper comprised three wide-angle seismic lines, A, B and C, located on a regional map in Fig. 2 and in detail in Fig. 4. They were shot from RRS Vema fracture zone and transverse ridge 44'30'w 495 43'30'W 44'W I 11" Shot LINE B l k n t 7n ~ . 10'30 N Figure 4. Detailed layout of the seismic experiment superimposed o n a simplified bathymetry. Cambridge OBSs are prefixed by 'C', Bedford Institute of Oceanography OBSs are prefixed by 'R'. Shot locations are small dots. Shot sizes were alternated along the lines to provide short range data for shallow structure and long-range arrivals for deep structure. Also marked by a heavy Line is the location of the airgun reflection profile along line B (Fig. 5 ) . A complete table of shot locations, sizes and detonation times is included in Fig. C of the microfiche appendix (GJ86/1) . Shackleton (leg 9/82). A fourth line along the spreading centre (labelled MAR on Fig. 2) is described elsewhere (Louden &White 1984; White, Louden & Potts 1986). Nine OBSs were laid. Four were Cambridge instruments (prefix 'C' on Fig. 4), used with a hydrophone and a three-component deployed geophone package (Duschenes, Potts & Rayner 1986). Five were Bedford Institute of Oceanography (BIO) instruments (prefix 'R' on Fig. 4) containing two-component geophones inside the main pressure vessel (Heffler & Barrett 1979). Three of the Cambridge OBSs recorded good quality data. The fourth, C5, recorded only a single channel of seismic data and was not used in the analysis. One Bedford Institute of Oceanography OBS (R14) failed t o release from the seafloor and was lost. Clock drifts and other details for all the instruments are listed in Fig. D of the Appendix (see Microfiche GJ 8611). A single-channel reflection profile, using a 2.6 1 (160 in3) airgun, was recorded along the shooting track in the fracture zone to provide sediment thickness control. This profile shows clearly the rugged nature of the igneous basement (Fig. 5). Shot depths and times were determined from the bubble-pulse periods (Spudich & Orcutt 1980a, b) and are listed in the Appendix (Fig. C). Accuracy is estimated as 21.5 ni in depth and 20.01 s in time. Ranges from shots t o each receiver were calculated by ray-tracing both direct and multiple water paths through a model of the water sound velocity versus depth structure (illustrated in the Appendix, Fig. B). A consistent set of shot positions (listed in the Appendix, Fig. C) and receiver locations (Appendix, Fig. D), was then derived by integrating the calculated ranges with the ship's position derived from satellite fixes and log and heading information. The absolute position of each receiver is known to only +200 m, although the relative ranges between shots into any given receiver are considerably better defined than this. Figs 6-8 show three typical record sections. Other record sections used in this analysis are illustrated in Figs E to Z of the microfiche Appendix. The record sections shown are not 496 Figure 5. Single channel 2.6 1 (160 in3) airgun reflection profile recorded along shooting track B. The three panels join to form a continuous east-west profile (see Fig. 4 for location of profile). The rugged basement topography is easily identified, except beneath the intersection deep. Vema fracture zone and transverse ridge Shotnumber r l 7475 497 OBS 1 LINE B 70 0001 0304 0607 0990 9293 9596 9899 101 8 h v) -4 1 I I I E ,w I 80 60 40 20 0 Range (Km) Figure6. Record section from OBS C1, line B, vertical (Z) geophone (low gain). Bandpass filtered between 2.0 and 20.0 Hz. Each trace has equal maximum plotting amplitude. First arrivals can be picked out to 80 k m range. sediment or water path corrected, and the effect of the rugged basement topography along the lines can be seen on Fig. 7, and more dramatically on Fig. 8, which is a section from line A across the transverse ridge. The first break on each hydrophone or vertical geophone trace is interpreted as the refracted P-wave arrival. N o S-waves were identified, except for P to S conversions at the basement just below some of the instruments. Picking accuracy varied with instrument, Shotnumber ? 1 72 73 74 75 76 40 OBS 12 LINE 6 78 79 8 0 8 1 82 8384 85 86878889 102 103 90 91 9293 94 95Bg 97 9899 lOOIOli I 104 20 Figure 7. Unfiltered record section from OBS R12, line B, hydrophone. Note ‘ringy’ character of seismograms which obscure amplitudes and second arrival data. Each trace has equal maximum plotth& amplitude. 498 C G. Potts, R. S. White and K. E. Louden OBS 15 LINE A Shotnumber Range (Km) Figure 8. Unfiltered record section from OBS R 1 5 , line A, hydrophone. Each trace has equal maximum plotting amplitude. Note the effect of the transverse ridge topography. Data quality from shots over the ridge was consistently good. range and shot size, and so estimated errors were assigned to each pick, and used throughout the interpretation. Almost all the seismograms had a ‘ringy’ character. Fig. 7 shows an example of this, on a section recorded by the hydrophone of OBS R12 on line B, along the fracture zone. This phenomenon is often observed in OBS data, and there is some evidence that it may be attributable to resonance of the instruments themselves (Loncarevic 1981) or is a result of reverberations in shallow layers. It has, however, been observed particularly in experiments in fracture zones using OBSs of different designs, and also surface receivers such as sonobuoys (e.g Sinha & Louden 1983) and multichannel hydrophone streamers (Potts, Calvert & White 1986). The ringy character may thus be a feature of fracture zone structure, caused by dispersion, attenuation and reverberation of the seismic signals in the highly fractured crust which is inevitably present, especially in the active transform region. Whatever the detailed cause, the effect is important, since in the Vema experiment the ringing distorted the amplitudes of early arrivals and obscured second arrivals. It proved impossible to model the amplitudes of the data sensibly. 4 Forward modelling using seismic ray-tracing 4.1 LINE B, ALONG THE FRACTURE ZONE Line B, which consisted of 35 shots and five instruments along the fracture zone, was raytraced first. The final model provided a starting point for ray-tracing line A which crossed the transverse ridge. Ray-tracing was performed using S E I S I 1 , with models incorporating vertically and horizontally varying seismic structures (Cerveny, Molotkov & PZen&‘k 1979). Fig. 9 shows the velocity-depth structure of the starting model for line B, derived from the results of Detrick et al. (1982). The structure was applied initially as a uniform lateral model and then modified to produce the final solution. Ray-tracing was stopped when it become apparent that the only method remaining to improve the fit of the calculated arrival Vemafracture zone and transverse ridge 499 Velocity (Km/s) Figure 9. The starting structure for forward modelling by ray-tracing along the fracture zone, line B. The structure was derived directly from the tau-p results of Detrick et al. (1982) (Fig. 3). Vertical scale is depth into igneous crust. times to the observed data was artificially to alter sediment thicknesses or local upper crustal structure beneath individual shot-points. Fig. 10 shows all the rays generating arrivals into the five OBSs along line B with the associated calculated and observed travel times for three of the instruments. Rays traced into individual receivers can be identified more clearly on the Appendix (Figs CC and EE), with corresponding large-scale plots of the observed and calculated travel times in Figs DD and FF. Water depths were picked to *I0 m from the echo-sounder profile recorded while firing. Sediment thicknesses were determined from the airgun reflection line, shot approximately along the shooting track (Figs 4 and 5). The inaccuracy of control over sediment thicknesses from the reflection profile is the major source of uncertainty along the whole of line B, and dictates that the minimum time-error is k0.075 s for all arrivals, however sharp their apparent onset may be. Weaker arrivals were sometimes assigned larger picking errors. 4.2 COMMENTS ON THE FINAL MODEL FOR LINE B An isovelocity plot of the final model is shown in Fig. 1 1. The major features are: (a) Crustal structure. Crustal velocities along the fracture zone are low compared with normal oceanic crust of the same age. Fig. 12 shows the velocity-depth functions from three sites along line B plotted against the bounds for normal oceanic crust greater than 3 Myr (from White 1984). Along nearly 60 km of the line, from OBS R13 to within 10 km of R1 1 , the crustal structure is fairly constant. At the western end of the line the crustal structure changes relatively abruptly, between OBS R11 and C1. The velocity-depth function (number 3 ) between the instruments indicates that a different velocity profile has emerged. (b) Moho depth and subcrustal structure. The seismic Moho beneath the fracture zone is marked by a large velocity increase and along most of the .line is shallower than under 7 - I I I I I I I I I 6 - 8t o m a 5 - ,q * p a&! c1 - a u - 0 4 -v) ? 7 t- A I I I I I I I I 1 I I I I I I I I I I c2 - 6 - @n mu ;u 8 e- 5: 4 - 7 - I I I I I I I 1 Y I I I I 1 I I I I 6 - 1 9 a a B c a d 'j 5 - 8 Ll u,, n 0 on R12 - " ' a?I 4 I I I I A I 1 I 1 1 15 I I I I I I I I I 10 20 SO 40 50 60 70 80 90 h E Y Y 0 W Distance (Km) I E Figure 10. (Top) Calculated arrival times (dots) generated by ray-tracing through the final model along line B, for OBS C 2 , R12 and C1. Shown for comparison are the observed arrivals (rectangles). The dimensions of the rectangles are estimates of the time (vertical) and distance (horizontal) errors o n the data. The distance uncertainty is assumed t o be + 2 0 0 m, the estimated uncertainty of the receiver location..The location of each OBS is marked with a solid triangle. (Bottom) All the raypaths constraining the final model for line B, along the fracture zone. Similar plots for OBS R11 and R 1 3 are shown in Figs EE and FF of the Appendix. 501 Vema fracture zone and transverse ridge 0 Sea level I 0 WEST I 7 - T I 10 20 30 40 50 Distance (Km) 60 70 80 I I I I 90 EAST Figure 11. Isovelocity contour plot of final model for line B, d o n g the fracture zone. The structure is constrained by the raypaths shown in Pig. 10. Contour interval is 0.2 km SKI.The numbered vertical lines indicate the locations of the velocity-depth functions shown in Fig. 12. normal crust of the Same age. Sub-Moho velocities start a t approximately 8.0 km s-l. The crust is thinnest (4.3 km) below the intersection deep (between OBSs C2 and R13), but remains thin for approximately 50 km westwards along the fracture zone. The thickening of the crust at the ends of the line is poorly constrained, compared with the control over Moho depth from the dense network of intersecting raypaths in the middle of the section (see Fig. 10). Thickening at the eastern end of the line, below the edge of the intersection deep is accompanied by lowering of the seismic velocities throughout the crust. However, at the western end of the line the velocities near the top of the crust increase as the seismic Moho deepens. Crusta: structure along the Vema fracture zone is similar to Sinha & Louden’s (1983) ‘type A’, since it exhibits low seismic velocities and is moderately thin, and to the crust found by Whitmarsh & Calvert (1986) in the Charlie-Gibbs fracture zone. 4.3 L I N E A , ACROSS THE T R A N S V E R S E R I D G E Line A, consisting of 21 shots and three OBS across the transverse ridge, was ray-traced from three different starting points. The structure of line B at its intersection with line A was the first starting model tried (Fig. 13, profile number 1). After 1 0 iterations, it became apparent that n o realistic adjustment of this model would reproduce the observed arrival times. Even though this attempt to model line A was unsuccessful, two features of the modelling were useful in establishing a new starting structure: (a) The shallow crustal structure over the ridge was definitely not the same as the shallow structure at the fracture zone line intersection. (b) To reproduce the observed arrivals, a region of high seismic velocity with low velocity gradients must be present under the ridge, 2-3 km beneath the seafloor. Layer 3 exhibits consistent high velocities (6.7-7.2 km s - ’ ) and low gradients (less than 0.15 s-’) (Spudich & Orcutt 1980b, White 1984), and its introduction into the structure 502 C. G. Potts, R. S. White and K . E. Louden Velocity (km s-’1 2 4 6 8 1 2 4 n € Y 6 Y r -P a Q) n 10 12 14 Figure 12. Velocity-depth profiles below three sites on line B, from the final model after ray-tracing (see Fig. l l ) , superimposed on the velocity bounds for normal crust greater than 3 Myr (from White 1984). ‘M’ marks the seismic Moho interface. beneath the ridge is sensible, considering point (b) above. The inclusion of an intermediate layer allowed more flexibility in the ray-tracing modelling, and the first reasonable fit of calculated arrivals t o the observed data was achieved. Fig. 13 shows the ray-paths through the model, called ‘solution one’, with the calculated arrival times, and Fig. 15(a) is an isovelocity contour plot. Due to the rough topography over the ridge, water depths from the wide-angle echosounder cannot be determined better than +I00 m. This inaccuracy, combined with the lack of off-line bathymetry, was the largest consistent source of error, and causes a minimum uncertainty in the travel times of 20.1 s (rectangular symbols on Fig. 13). There was no reflection profile along line A, so sediment thicknesses in the fracture zone valley at the north end of the line were extrapolated from the line B data and the isopach map of Prince & Forsyth (1986). Only long-range shots into OBS R13 (> 15 km) were used. This was because R13 did not lie exactly on the shooting track for line A, but was approximately 4.6 km to the east, and at short ranges the deviation from the plane of the experiment becomes more serious. The 5 4- l l a I D U P O L a 3- 5 - I I I - RED R13 - I I I I I I I I fP E 4 - Q 0 A - - U h cn Y 00 \ T II I I I 5 - un 0 4- a ‘ I - g o o R15 g u m u Kl I - a U E 3- A n I I 30 40 E 15 20 0 S 10 20 Distance ( k m ) 50 N Figure 13. The first model, ‘solution one’, for line A, across the transverse ridge, assuming a fixed velocity structure at the line intersection, and a layer-3 region beneath the ridge. Shown (top) are the arrivals generated by the model (dots) compared with the observed values (rectangles). The dimensions of the rectangles are determined by estimates of the errors o n the observed data. Bottom are all the ray-paths constraining the model into the three OBSs. C.G. Potts, R. S. White and K . E. Louden 504 A R R I V A L TIMES L I N E A (SOLUTION T W O ) I 5 - I I I I 1 I I I 0 0 a 4 0 O - - t - s 3- p a u o IJ I 1 1 1 0 5 n E Y Y 5P 10 0 P c1 p-w I 18 l5 L 0 10 - R11 A 30 Distance ( k m ) 20 40 50 Figure 14, The second model, ‘solution two’, for line A, across the transverse ridge, with n o imposed velocity structure. Top are the observed (rectangles) and computed (dots) travel times, and bottom are the rays constraining the structure. The dimensions of the rectangles are determined by estimates of the errors on the observed data. Vema fracture zone and transverse ridge 01 I I TRANSVERSE RIDGE I I I 505 I c4 TRANSITION ZONE I 5 c c $ a 10 15 LINE A SOLUTION ONE 0 10 20 30 40 N Distance (Km) S 50 I I I I 10 20 30 40 1 5 n E Y Y 5Q 10 al n 15 0 S Distance (Km) 50 N Figure 15. Isovelocity contour plots of the two solutions for line A. Solution one has very thin crust beneath the abrupt ‘transition’ zone, whereas solution two allows the normal crust to alter more gradually, as layer 3 thins out. Contour interval 0.2 k m s - l . fracture zone structure from line B at the line intersection was kept fixed throughout the modelling. To maintain this fixed structure, and fit the observed arrival times, a ‘transition’ zone evolved beneath the edge of the transform valley in which crustal layers thinned to only 2.5 km and seismic velocities were reduced. The nature of the transition zone was dictated by the choice of the line B structure as a constraint at the line intersection. To test the sensitivity of the modelling process, this C. G. Potts, R. S. White and K . E. Louden 506 constraint was removed. Once again, a reasonable fit to the observed data was obtained quickly after adjustment of the structure (Fig. 14). The isovelocity contour plot of this model, called 'solution two', is shown in Fig. 15(b). 4.4 COMMENTS O N T H E F I N A L M O D E L S FOR L I N E A Solution one (Fig. 15a) is the best constrained structure modelled for line A since it incorporates the structural tie with line B. The main difference between the two final solutions is the inclusion in solution one of the transition zone. Fig. 16 shows velocity-depth functions from selected sites along both models. Summarized below are observations on the nature of the two structures: (a) Both models indicate that a layer-3 type region exists beneath the transverse ridge and that, as the fracture zone is approached, this region thins and alters (Fig. 15). Both models LINE A Solution One Velocity ( k m 6-9 Velocity (km 9.') 2 4 6 8 1 2 0 S a P 20 30 35 43.4 4 6 8 1 0 ti 8 0 10 12 LINE A Solution Two Velocity 4 6 Vstoclty (km (km s - 0 8 1 2 0 4 6 8 8.') 1 0 S Figure 16. Velocity-depth profiles from the two solutions to line A , across the transverse ridge. Shown, top, are profiles from solution one, and bottom, from solution two. The common trends observed are crustal thinning, disappearance of layer 3 and lowering of crustal velocities. The main difference is the inclusion of very thin crust in solution one beneath the southern flank of the fracture zone valley. Vema fracture zone and transverse ridge 507 also suggest that the crust beneath the transverse ridge is normal in character, with velocitydepth curves falling within the bounds for normal ocean crust (Fig. 15). (b) There is evidence in both models that crustal thinning commences u i e r 20 km away froin the fracture zone: in solution one, thinning from approximately 4 t o 2 km; in solution two, thinning less dramatically from approximately 5 to 3.5 km. This overall crustal thickness change is superimposed on the disappearance of layer 3, which is vertically thinned and laterally altered in solution one, or simply thinned out in solution two. I n both cases the net crustal velocities decrease while velocity gradients increase. (c) In solution one the thinnest crust, approximately 2 km thick, lies beneath the flanks of the fracture zone, with a thickening of the seismic crust beneath the middle of the fracture zone valley. Solution two is more homogeneous, including thin, but not exceptionally so, crust. The results of forward modelling using real data are not unique, and the final models are influenced by the initial assumptions of the starting structure. The two models presented here satisfy the observed arrival times, and are both plausible geological structures. Three-dimensional gravity modelling over the Vema fracture zone conducted independently by Prince & Forsyth (1986) produced a structure which indicated very thin crust beneath both flanks of the fracture zone with rather thicker crust under the central part of the valley. This result is in general agreement with the structure under the fracture zone valley suggested by our solution one. Prince & Forsyth’s results also, however, suggest the presence of a thick crustal root beneath the transverse ridge on the south wall of the fracture zone which we definitely do not find. Since the gravity dataset is rather limited in this region the discrepancy between seismic and gravity interpretations of the transverse ridge may be explained by poor gravity data coverage. 5 Discussion 5.1 F R A C T U R E ZONE T R A N S V E R S E RIDGES Line A, across the Vema fracture zone, is the first detailed seismic refraction experiment t o investigate directly the structure of a large transverse ridge. Both solutions modelled (Fig. 15) indicate that the ridge is a section of oceanic crust which has been uplifted and warped by an unknown process. The oceanic crust, however, was itself originally formed in the vicinity of a ridge-transform intersection. Thus the crustal deformation is superimposed on the lateral variations in crustal structure due to the proximity of the fracture zone. In both models the crust under the ridge contains a region analogous to layer 3, albeit with somewhat higher velocity gradients (0.2-0.25 s-‘) than are normally observed (0.15 s - l , Spudich & Orcutt 1980b), and a recognizable upper crustal layer 2. Overall crustal thickness decreases northwards and layer 3 disappears, as the structure changes from nearly normal to fracture zone ‘type A’ crust (as defined by Sinha & Louden 1983). Our model of the structure across the transverse ridge supports the earlier conclusion from rock sampling that the ridge consists of a section of uplifted oceanic crust formed near a fracture zone (Honnorez et al. 1984; Bonatti 1978). Two main mechanisms to explain crustal uplift of fracture zone transverse ridges have been proposed. The first is that the extensive fracturing in the region of an active transform allows water to penetrate through the crust to the upper mantle, where serpentinization occurs (Francis 1981). Serpentinization is the reaction between water and olivine below 5OO0C, which results, amongst other things, in a rock volume increase. The low density serpentinite is buoyant, rises diapirically under the flanks of the fracture zone valley, and 508 C. G. Potts, R . S. White and K . E. Louden uplifts the crust. Serpentinized upper mantle material exhibits reduced seismic velocities (Coleman 1971). The second suggested mechanism (Bonatti 1978) is that compressional stresses caused by the re-orientation of the transform fault following changes in spreading direction produce vertical tectonic movement of blocks of crustal material adjacent to the fracture zone. Bonatti & Hamlyn (1 978), for example, attribute the considerable vertical movements suggested by the chemistry of rocks dredged from the transverse ridge of the Owen fracture zone to mantle upwelling below the spreading ridge at the time of crustal formation; the rocks are later exposed tectonically by uplift along the fracture zone. Large transverse ridges are often found where slow-spreading mid-ocean ridges change direction, as for example near the Vema and Romanche in the central Atlantic, the Charlie-Gibbs in the North Atlantic and along the south-west Indian Ridge in the south Indian Ocean. At such locations tectonic compression is likely to occur. The seismic results reported here cannot, alone, discriminate between these two mechanisms. Both would result in the upwarped crustal section which is observed. If extensive serpentinization has occurred, a large body of low velocity, low density material may be emplaced under the ridge; unfortunately, this could prove to have a similar seismic character to oceanic layer 3 (Francis 1981). The simplicity of the final crustal cross-section produced by modelling the seismic data, however, betrays no trace of the structural disruption which extensive intrusion of serpentinite must cause. The only aspect of the crustal model that might indicate the presence of a large serpentinized body is the anomalously high velocity gradient found in layer 3. This value could, however, be explained as being symptomatic of a change in the process of formation of the crust in the vicinity of fracture zones (White e t al. 1984). The suggestion that the uplift was due to tectonic compression in the Vema is considered preferable. 5.2 C R U S T A L S T R U C T U R E BENEATH THE F R A C T U R E ZONE V A L L E Y Fig. 11 shows an isovelocity contour plot of the final model along line B, with velocitydepth profiles for three sites shown on Fig. 12. Crustal velocities are low compared with normal oceanic crust of the same age, and the crust along most of the line is significantly thinner than normal crust (4-4.5 km compared with 5-8 km). There is an absence of a layer-3 type region along most of the line. The observed sub-Moho velocities (7.9-8.0 km s-l) and gradients (less than 0.1 s-') are not unusual, compared to normally observed values. Low crustal velocities and thinner than normal crust are commonly observed in seismic refraction surveys over fracture zones (e.g. Detrick et al. 1982; Cormier, Detrick & Purdy 1984; Sinha & Louden 1983; White et al. 1984; Whitmarsh & Calvert 1986). The results from the Vema confirm these observations. The low crustal velocities are attributed mainly to the intense fracturing and brecciation caused by faulting in the active transform. Bowen & White (1986) observed faulting around the intersection deep of the Vema fracture zone, which would also produce a final phase of disruption of the crust as it passed into the inactive domain. Karson & Dick (1983), reported significant thicknesses of basaltic rubble in the Kane fracture zone, presumably debris from the fracture zone walls. Similar rubble was observed during submarine investigations of the Oceanographer transform valley (OTTER 1984, 1986) and also encountered during drilling in the Vema by DSDP (Perch-Nielson et al. 1977). The accumulation of' this low velocity material on top of fractured and faulted crust would further reduce the apparent velocity of the crust. The almost universal absence of a layer-3 type region in fracture zone crust is harder to explain. Even after 60-77 Myr, in the Tydeman fracture zone (Calvert & Potts 1985), no Vemofracture zone and transverse ridge 509 typical layer 3 is observed. Its absence, similarly, from most young sections of fracture zones studied (e.g. Vema, Charlie-Gibbs, Oceanographer) indicates that the effect must be a primary feature of crust formed in, or near, fracture zones. One of the most interesting features of the results from line A, across the fracture zone arid transverse ridge, is the possibility, from the correspondence between the seismic solution one (Fig. 15a) and the gravity profile of Prince & Forsyth (1986), that the thinnest crust, or shallowest seismic Moho, occurs beneath the flanks of the fracture zone valley. Results from line MAR, along the spreading centre, support this hypothesis (Louden &White 1984; White et aE. 1986). Other authors (e.g. Sinha & Louden 1983; Cormier et al. 1984) have reported limited regions of apparently very thin crust along fracture zones (i.e. Sinha & Louden 1983, type ‘B’ crust). It is conceivable that such results represent sampling by the seismic experiments of the flanking thin crust, or excursions of thin crust across the fracture zone valley, or both. It is not yet possible t o decide whether these crustal variations represent real shortterm changes in the ridge-transform crustal accretion processes, or later variations in hydrothermal alteration of the crust and mantle or, again, both. It can no longer be doubted, however, that fracture zones create significant lateral variability on the scale of a few kilometres as well as the regional effects, discussed below in Section 5.3, on a scale of a few tens of kilometres. 5.3 THE W I D T H O F A N O M A L O U S F R A C T U R E Z O N E C R U S T Estimates of the full width of’ fracture zones from geophysical methods are scattered between 7 k m (White & Matthews 1980) and 30-40 k m (White et al. 1984). The dataset is limited, and as yet n o obvious correlation has been found between the apparent width of anomalous crust and the age, offset, spreading rate or any other variable feature of a fracture zone. The estimate of 10-20 km total width for the low velocity crustal and subcrustal slab from the Tydeman fracture zone (Potts et al. 1986) is in agreement with the results reported by several other authors (e.g. Sinha & Louden 1983; Cormier et al. 1984). The width of anomalously thin crust under the Blake-Spur fracture zone, determined by reflection profiling across an old section (135 Myr), is approximately I 5 km (Mutter et al. 1984). Several authors (e.g. White et al. 1984; Sinha & Louden 1983; Corniier e f al. 1984; Ambos 1984) suggest that anomalous crust along fracture zones is a combination of (i> largescale gradual crustal thinning, (ii) smaller scale lowering of crustal velocity, in conjunction with more marked thinning of the crust (i.e. Sinha & Louden’s 1983, type A crust) and (iii) occasional sections of very thin crust (i.e. Sinha & Louden’s 1983, type B crust). The results from the Vema strongly support these proposals; the structure modelled under the ridge shows that crustal thinning is effective u p t o 20 km away from the axis of the transform, and low crustal velocities become important approximately 1 0 kni from the axis. This indicates that the overall width of affected crust could be 40 kni (crustal thinning) but the highly anomalous region may overall be only 20 km wide or less. Surveys which d o not sample oceanic crustal structure over significant distances (>20 km) clear of a fracture zone may detect only the region of low velocity crust, and not resolve the larger-scale variation in crustal thickness. 6 Conclusions (a) The crust beneath the flanking transverse ridge of the Vema fracture zone is essentially normal in its seismic velocity structure and its overall thickness. The seismic Moho is upwarped under the transverse ridge. No major zone of serpentinite diapirism was detected 510 C.G. Ports, R. S. White and K . E, Louden under the ridge. This suggests that the ridge represents a relatively normal section of oceanic crust uplifted by tectonic forces associated with the Vema transform, rather than by diapirism. (b) Gradual crustal thinning occurs over a region of up to 20 km on either side of the transform fault. (c) The thinnest crust is found under the flanks of the transform valley, and near the ridge-transform intersection. The crust appears to be somewhat thicker (though still thinner than normal oceanic crust) beneath the central part of the valley. This effect could be attributed to depression of the seismic Moho due to serpentinization by water penetrating down through the crust in the extensively faulted region along the middle of the transform valley (i.e. in the vicinity of the transform fault). (d) Under the transform valley the crust is not only somewhat thinner than normal oceanic crust, but is also highly anomalous in its velocity structure. It exhibits a marked absence of a seismic layer 3 , which is one of the most characteristic features of normal oceanic crust, and generally has rather low seismic velocities. The absence of seismic layer 3 indicates that accretionary magmatic processes in the transform are far from normal, whilst the low seismic velocities are attributed to the extensive faulting and fracturing in the crust of the transform zone, together with the occurrence of talus and rubble layers in the upper section. Acknowledgments The success of the field-work at sea was due in large part to the support and hard work of our colleagues on both sides of the Atlantic, the technical staff from Research Vessel Services of the NERC and the officers and crew of RRS Shackleton. We thank them all. Beverley Smith processed the words and Phyl Fisher drafted most of the diagrams. The Atlantic Geoscience Center of the Bedford Institute of Oceanography kindly loaned six of their ocean bottom seismometers for use on this experiment and we are especially indebted to Brian Nichols for his efforts and indispensable contributions whilst digitizing the BIO OBS data, and to Don Locke for running the BIO instruments at sea. Roger Prince and Don Forsyth provided bathymetric information which we gratefully acknowledge. Funding was from the Natural Environment Research Council under grants GR3/4157 and GR3/4403 to RSW and from the National Science and Engineering Research Council grant A8459 to KEL. Department of Earth Sciences, Cambridge, contribution number ES 638. References Ambos, E. L., 1984. Oceanographer transform fault structure compared to that of surrounding oceanic crust - results from seismic refraction data analysis, unpublished PhD thesis, University of Hawaii. Bonatti, E., 1978. Vertical tectonism in oceanic fracture zones, Earth planet. Sci. Lett., 37, 369-379. Bonatti, E. & Crane, K., 1982. Oscillatory spreading explanation of anomalously old uplifted crust near oceanic transforms, Nature, 300, 343-345. Bonatti, E. & Hamlyn, P. R., 1978. Mantle uplifted block in the Western Indian Ocean, Science, 201, 1982. Bonatti, E. & Honnorez, J., 1976. Sections of the Earth’s crust in the equatorial Atlantic, J. geophys. Res., 81,4104-4116. Bowen, A . N. & White, R. S., 1986. Deep-tow profiles from the Vema transform and ridge-transform intersection, J. geol. SOC. Lond., in press. Calvert, A. J. & Potts, C. G., 1985. Seismic evidence for hydrothermally altered mantle beneath the old crust of the Tydeman Fracture Zone, Earth planet. Sci. Lett., 75,439-449. Vema fracture zone and transverse ridge 511 Eervenp, V., Molotkov, I. A. & PSenhk, J., 1979. Ray Method in Seismology, Univerzita Karlova, Praha. Coleman, R. G., 1971. Petrological and geophysical nature of serpentinites, Bull. p o l . SOC.Am.. 82, 897 -918. Carmier, M., Detrick, R. S. & Purdy, G. M., 1984. Anomalously thin crust in oceanic fracture zones new seismic constraints from the Kane fracture zone, J. geophys. Res., 89, 1 0 249---lo266. Detrick, R. S., Cormier, M. H., Prince, R. A , Forsyth, D. W. & Ambos, E. L., 1982. Seismic constraints on the crustal structure within the Vema Fracture Zone, J. geophys. Res., 87,10599- 10612. Detrick, R. S. & Purdy, G. M., 1980. The crustal structure of the Kane fracture zone from seismic refraction studies, J. geophys. R e x , 85, 3759-3777. Detrick, R. S., White, R. S. & Purdy, G. M., 1986. Crustal structure of oceanic t'racture zones from seismic reflection and refraction studies, in The Tectonics of Fracture Zones, ed. Fox, P. J., Academic Press, London, in press. Duschene, J. D., Ports, C. G. & Rayner, M., 1985. The Cambridge Deep Ocean Geophone,Mar. geophys. Res., 7,455-466. Fox, P. J., Detrick, R. S. & Purdy, G. M., 1980. Evidence for crustal thinning near fracture zones: implications for ophiolites, in Proc. int. Ophiolite Syrnp., pp. 161-168, ed. Panayiotou, A.. Ministry of Agriculture and Natural Resources, Cyprus Geological Survey Department. Fox, P. J. & Gallo, D. G . , 1984. A tectonic model for ridge-transform-ridge plate tectonics: implications for the structure of oceanic lithosphere, Tectonophys., 104,205 -242. Francheteau, J., Choukroune, P., Hekinian, R., Le Pichon, X. & Needham, H. D., 1976. Oceanic fracture zones do not provide deep sections in the crust, Can. J. Earth Sci., 13, 1223-1235. Francis, T. J. G., 1981. Serpentinisation faults and their role in the tectonics of slow spreading ridges, J. geophys. Res., 86,11616-11 622. Heffler, D. E. & Barrett, D. L., 1979. OBS development at Bedford Institute of Oceanography, Mar. geophys. Res., 4,227-245. Honnorez, J., Mevel, C. & Montigny, R., 1984. Geotectonic significance of gneissic amphibolites from the Vema fracture zone, equatorial Mid-Atlantic Ridge, J. geophys. Res., 89, 1 1 379-11 400. Karson, J. A. & Dick, H. J. B., 1983. Tectonics of ridge transform intersections at the Kane fracture zone, Mar. geophys. Res., 6,51-98. Loncarevic, B. D., 1981. The Lopez Island OBS intercomparison experiment. Special issue, Mar. geophys. Res., 5 , (1). Louden, K. E. & Forsyth, D. W., 1982. Crustal structure and isostatic compensation near the Kane fracture zone from topography and gravity measurements - I. Spectral analysis approach, Geophys. J . R. astr. Soc., 68, 725-750. Louden, K. E. & White, R. S., 1984. Changes in crustal structure along the Mid-Atlantic Ridge at its western intersection with the Vema Fracture Zone, EOS, 6 5 , 1009. Ludwig, W. J. & Rabinowitz, P. D., 1980. Structure of the Vema fracture zone,Mar. Geol., 35,99-110. Mutter, J. C., Detrick, R. S. & NAT study group, 1984. Multichannel seismic evidencc for anomalously thin crust at the Blake Spur fracture zone, Geology, 12,534-537. OTTER scientific team, 1984. The geology of the Oceanographer transform: the ridge transform intersection,Mar. geophys. Res., 6,109-141. OTTER scientific team, 1986. The geology of the Oceanographer transform: the transform domain, Mar. geophys. Res., 7, No. 3, 329-358. Perch-Nielson, K., Supko, P. R., Boersma, A, Bonatti, E., Carlson, R. L., McCoy, I;., Neprochnov, Y. P., Zimmerman, H. B. & Fisher, D. E., 1977. Site 353: Vema Fracture Zone, in Inir. Rep. Deep Sea drill. Proj., 39, 27-44. eds Supko, P. R., Perch-Nielson, K. et al., Government Printing Office, Washington, DC. Ports, C. G., Calvert, A. J. & White, R. S., 1986. Crustal structure of Atlantic fracture zones - 111. The Tydeman fracture zone, Geophys. J. R. astr. Soc., in press. Prince, R. A. & Forsyth, D. W., 1986. Three-dimensional modelling of gravity anomalies at the Vema fracture zone mid-Atlantic Ridge intersection: crustal thickness and the extent of local isostatic compensation, J. geophys. Res., submitted. Robb, J. M . & Kane, M. F., 1975. Structure of the Vema Fracture Zone from gravity and magnetic intensity profiles, J. geophys. Res., 80,4441 -4445. Sibuet, J. C. & Veyrat-Peinet, B., 1980. Gravimetric model of the equatorial fracture zones, J. geophys. Res., 85, 943-954. Sinha, M. C. & Louden, K. E., 1983. The Oceanographer fracrure zone - I. Crustal structure from seismic refraction studies, Geophys. J. R. astr. Soc., 75,713-736. 512 C. G. Potts, R. S. White and K. E. Louden Spudich, P. & Orcutt, J., 1980a. Petrology and porosity of an oceanic crustal site: results from wave form modelling of seismic refraction data, J. geophys. Res., 85, 1409-1433. Spudich, P. & Orcutt, J,, 1980b. A new look at the seismic velocity structure of the oceanic crust, Rev. Geophys. 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Appendix on microfiche A microfiche accompanies this paper (GJ 86/1). It contains copies of the seismic record sections with normalized trace amplitudes for each ocean bottom seismometer from each of the three shooting lines used in the interpretations reported here. Tabulations of the shot and receiver data are included, as are enlarged versions of the ray-tracing fits to the line B arrivals. The contents of the Appendix on microfiche are as follows. A. Map showing positions of all shots and receivers. Cambridge OBSs are prefixed by ‘C’, Bedford Institute of Oceanography OBSs are prefixed by ‘R’. Shot locations are small dots. B. Water sound-velocity structure in the Vema fracture zone, used in range calculations. Structure down to 2500 m (shown by solid line) is from direct measurement using a soundvelocimeter lowered on a conducting cable. Structure below 2500 m (broken line) is from existing nearby oceanographic measurements. C. Shot data showing shot size, depth, time, latitude, longitude, and seafloor depth. D. Receiver data showing latitude, longitude, water depth and clock drift. E. Record section from line A, receiver C1, horizontal geophone. F . Record section from line A, receiver C2, vertical geophone. G. Record section from line A, receiver C4, vertical geophone. H. Record section from line A, receiver C4, horizontal geophone. I. Record section from line A, receiver R13, hydrophone. J. Record section from line A, receiver R15, hydrophone. K. Record section from line B, receiver C1, vertical geophone. L. Record section from line B, receiver C1, vertical geophone (charge scaled). M. Record section from line B, receiver C l , hydrophone. N. Record section from line B, receiver C2, hydrophone. 0. Record section from line B, receiver C4, vertical geophone. P. Record section from line B, receiver R11, vertical geophone. Q. Record section from line B, receiver R12, hydrophone. R. Record section from line B, receiver R13, vertical geophone. S. Record section from line B, receiver R15, horizontal geophone. T. Record section from line C, receiver C1, hydrophone. U. Record section from line C, receiver C2, vertical geophone. V. Record section from line C, receiver C4, vertical geophone. W. Record section from line C, receiver R l l , horizontal geophone. X. Record section from line C, receiver R12, hydrophone. Y. Record section from line C, receiver R13, horizontal geophone. Verna fracture zone and transverse ridge 513 Z. Record section from line C, receiver R I S , horizontal geophone. AA. True scale plot of seafloor topography and sediment depths across the transverse ridge (line A) and along the fracture zone valley (line B). BB. Last-square fits of straight lines to uncorrected first arrivals along line B. Vertical scale is arrival time reduced at 8.0 km s-' .Horizontal scale is range in kilometres. CC. Raypaths through the final model for line B, along the fracture zone, into OBS C1, C2 and R12. DD. Calculated arrival times (ctosses) generated by final model for line B shown in Fig. CC. Boxes show observed arrival times. Solid circles are location of OBS C1, C2 and R12. The dimensions of the rectangular boxes correspond to the estimated errors on the data. EE. Raypaths through the final model for line B, along the fracture zone, into OBS RI 1 and R13. FF. Calculated arrival times (crosses) generated by final model for line B shown in Fig. EE. Boxes show observed travel times. The dimensions of the rectangles correspond to the estimated errors on the data.