Crustal structure of the Southwest Indian Ridge at 66 E: seismic constraints

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1 Geophys. J. Int. (26) 166, 1 17 doi:.1111/j.16-26x.26.1.x Crustal structure of the Southwest Indian Ridge at 66 E: seismic constraints T. A. Minshull, 1 M. R. Muller 2 and R. S. White 1 School of Ocean and Earth Science, National Oceanography Centre, University of Southampton, European Way, Southampton SO1 ZH, UK. tmin@noc.soton.ac.uk 2 Cube Geophysics CC, 9 th Street, Melville, Johannesburg 292 South Africa Bullard Laboratories, Department of Earth Sciences, University of Cambridge, Madingley Road, Cambridge CB EZ, UK Accepted 26 March 1. Received 26 March 1; in original form 2 December 2 SUMMARY The Southwest Indian Ridge represents a slow-spreading end-member of the global mid-ocean ridge system, and as such its structure places important constraints on models of melt supply and delivery from the mantle at ridges. We present results from a wide-angle seismic experiment conducted at the ridge axis at 66 E, in a region that has comprehensive swath bathymetric, gravity and magnetic data coverage and where the full spreading rate is 12 mm yr 1. Based on these data, the experiment traversed four spreading segments. Crustal thickness and velocity structure were determined along three intersecting profiles each km long using shots from a -gun, 71 L tuned airgun array towed at 1 m depth and fired at s intervals, recorded on three ocean bottom hydrophones on each profile. OBH data show high-amplitude arrivals from oceanic Layer 2, lower-amplitude arrivals from Layer and wide-angle reflections from the Moho. Forward modelling and inversion of traveltime picks from these data show that the crust consists of a km-thick Layer 2 with a high velocity gradient and a..-km-thick Layer with a low velocity gradient, and a crustal thickness of 2.2. km. Additional constraints on the models come from 2-D modelling of gravity data along the profiles, corrected for -D effects of off-line bathymetry. Along-axis, the thickness of Layer 2 varies little, but Layer is thick at segment centres and very thin at segment boundaries. Along a flowline profile, crustal thickness varies by up to 7 per cent from its minimum value in Myr. The reduced crustal thickness is consistent with observations from very slow-spreading ridge axes elsewhere and may be explained by conductive cooling of the upwelling mantle. The large along-axis variations in Layer thickness indicate that magmatic accretion is focused at segment centres and melt is delivered to segment ends perhaps only by lateral dyke propagation. Flowline variations in crustal thickness may result from episodicity of melt supply on timescales of Myr and by tectonic extension during amagmatic periods. Velocities at the top of Layer 2 are poorly correlated with crustal age based on magnetic anomalies, suggesting also that episodicity is decoupled between adjacent segments. Key words: crustal structure, mid-ocean ridge, Southwest Indian Ridge. GJI Marine geoscience INTRODUCTION The Southwest Indian Ridge (SWIR) forms the boundary between the Antarctic plate to the south and the Somalian plate to the north (Fig. 1). It forms a slow-spreading end-member in the midocean ridge system, with full spreading rates of mm yr 1 (Chu & Gordon 1999), and little change in these rates over at least the last Ma (Patriat & Segoufin 1988). At these spreading rates, decompression melting of the upwelling mantle at the ridge axis is substantially reduced by conductive cooling (White et al. 21), and mantle peridotites are frequently exposed at the seafloor (Dick 1989). Crustal thickness estimates have been made for large sections of the ridge axis based on gravity data (e.g. Grindlay et al. 1998; Rommevaux-Jestin et al. 1998; Cannat et al. 2). Such studies provide a valuable overview, but are limited by a lack of knowledge of lateral density variations in the crust, which can be significant (e.g. Minshull 1996). The combination of gravity and seismic constraints gives more reliable crustal models. However, until recently few seismic constraints have been available on the crustal structure anywhere along the 72 km SWIR axis. The SWIR axis may be divided into a series of different domains based on its morphology, with closely spaced, large-offset fracture C 26 The Authors 1 Downloaded from

2 16 T. A. Minshull, M. R. Muller and R. S. White 26 S A6 Survey Area RTJ 28 S A18 A6 S CIR 2 S Melville FZ SOMALIAN PLATE AIIFZ SEIR S A18 SWIR ANTARCTIC PLATE 6 E 62 E 6 E 66 E 68 E 7 E mgal Figure 1. Satellite gravity image (Sandwell & Smith 1997) of the Southwest Indian Ridge (SWIR) near 66 E, contoured at a 2 MGal interval. Box marks detailed study area. Dashed lines show the traces on the Somalian and Antarctic plates of the Rodrigues triple junction (RTJ), which bound crust formed at the SWIR from crust formed at the Central Indian Ridge (CIR) and Southeast Indian Ridge (SEIR). Solid lines mark magnetic isochrons 6 ( 21 Ma) and 18 ( 2 Ma) (Patriat & Segoufin 1988). Inset shows plate tectonic setting, with thick solid lines marking plate boundaries. The MGal gravity contour, magnetic isochrons 6 and 18, and the triple junction traces (dashed lines on inset) are also marked. AIIFZ marks location of Atlantic II Fracture Zone at7 E. zones in some domains and elsewhere long lengths of ridge axis which lack fracture zone offsets (e.g. Mendel et al. 1997). These domains also exhibit variations in axial depth and relief, axial gravity anomaly, and the abundance of seamount volcanism (Mendel & Sauter 1997; Rommevaux-Jestin et al. 1998; Cannat et al. 1999). Seismic data from 7 1 Ma crust at the Atlantis II Fracture Zone at 7 E, within a domain of large-offset, closely spaced fracture zones (Fig. 1 inset), indicate a crustal thickness of km, with the crust locally underlain by up to 2 km of partially serpentinized mantle (Muller et al. 1997, 2). Whilst the morphology in the region of the Atlantis II Fracture Zone is typical of many regions of the SWIR, mantle melting in this region may have been influenced substantially by the cold edge effect of the transform (Muller et al. 2), and the observed structure may be influenced by processes which post-date crustal formation at the ridge axis. In order to study crustal accretion as a function of spreading rate alone, it is valuable to study the crust on-axis and away from the cooling effect of such transforms. Here we present seismic data from the axial region at 66 E, away from large transform offsets. GEOLOGICAL SETTING Between the Melville Fracture Zone at 6 E and the Rodrigues Triple Junction at 7 E, the ridge axis lacks fracture zone offsets and the ridge flank topography is relatively smooth for the SWIR, though there are still several kilometres of relief. Axial valley depth is up to 2 km deeper than that predicted by the plate cooling model of Parsons & Sclater (1977), seamount abundance is lower than to the west of the Melville Fracture Zone and mantle S-wave velocities are higher (Debayle & Leveque 1997; Mendel et al. 1997; Mendel & Sauter 1997). These observations have been used to infer that subaxial crustal and mantle temperatures are lower than to the west of the Melville Fracture Zone, leading to lower degrees of partial melting, less magma supply to the ridge axis, and thinner crust (Cannat et al. 1999). To the west of the Melville Fracture Zone, gravity lineations corresponding to both transform and non-transform discontinuities (NTDs) may be traced all the way from the ridge axis to the triple junction traces bounding crust formed at the SWIR. One nontransform offset just east of the Melville Fracture Zone at 61 6 E on the ridge axis may also be followed out to the triple junction trace (Fig. 1), and intersects it close to magnetic anomaly 18. All of the crust to the east of this non-transform offset is younger than anomaly 18, and thus was created during a period of approximately constant spreading rate on all three ridge axis around the triple junction (Patriat & Segoufin 1988). During this time, it appears that no persistent ridge offsets were generated at the triple junction, since lineations perpendicular to the ridge axis are absent. Hence while the segmentation to the west of 62 E is probably controlled by the tectonics of the triple junction and its response to changes in spreading rate on the three ridge axes, the segmentation east of 62 E is likely to be a function of spreading rate alone. The ridge axis in this region is defined clearly by a linear free air gravity low, and is offset by a series of NTDs or deviations from axial linearity which define segment boundaries (Fig. 1). Along-axis, variations in the mantle Bouguer C 26 The Authors, GJI, 166, 1 17 Downloaded from

3 Crustal structure of the SWIR 17 gravity anomaly (e.g. Kuo & Forsyth 1988) are poorly correlated with morphological segment boundaries (Cannat et al. 1999). Our survey area around 66 E(Fig.2)wasselected to coincide with pre-existing swath bathymetric, shipboard gravity and magnetic data extending km off-axis in both directions (Munschy 1987). This area has been subsequently the focus of further detailed geophysical analysis using swath bathymetric, gravity and magnetic data (Cannat et al. 2; Sauter et al. 2). Based on regional plate reconstructions, the full spreading rate in this area since Chron 2A (.2 Ma) is 12 mm yr 1 (Chu & Gordon 1999). Based on fitting simple block models to profiles between 61 and 69 E, Cannat et al. (2) infer similar rates of 1 mm yr 1 since Chron 2A and 1. 1 mm yr 1 since Chron ( Ma). Our seismic profiles extended beyond the region of swath coverage, so it was necessary to merge single-beam echosounder data acquired during the seismic experiment and other historical data with the swath data to complete the grid displayed in Fig. 2. Since most of the swath data were acquired prior to widespread availability of the Global Positioning System (GPS), positional discrepancies of 1 km are present between these data and our GPS-navigated single-beam data; to smooth out these discrepancies, the combined data set was gridded at m interval. The 2-D seismic and gravity models presented below use our original along-track bathymetry rather than the smoothed grid. The area exhibits marked variations in axial depth from less than m to more than m, a feature which is characteristic of the axial valley east of 62 E. Two prominent highs at 6 6 E and E represent the highest axial elevations between the Jourdanne Mountains at 6 E and the triple junction (Mendel et al. 1997). Our experiment was centred on an intermediate axial high at 66 E interpreted as a segment centre, though no significant along-axis gravity high is observed (Fig. 2). In the hourglass shaped segment around this point, labelled B in Fig. 2, the rift valley floor is almost flat, with a poorly developed neovolcanic ridge and only a few scattered seamounts (Mendel et al. 1997). Dredges in this segment recovered only basalt (Robinson et al. 1996). The m high, east west striking ridge between 66 and 66 2, which coincides with an along-axis gravity high, is studded with at least five scattered seamounts and is interpreted as a sigmoidal axial volcanic ridge within a 2 km spreading segment labelled C in Fig. 2, which is bounded by NTDs around 66 8 and (Mendel et al. 1997). The eastern NTD is also marked by a clear 17 km offset in the axial magnetic anomaly pattern (Munschy 1987; Munschy & Schlich 1989). At the western NTD, a 1 km offset of anomalies 2 and 2A is observed, but no offset in anomaly 1 is evident (Fig. 2). The bathymetric and gravity high at 6 E may represent the midpoint of another segment, labelled A in Fig. 2, while a further bathymetric high at 66 7 E may be interpreted as the centre of a fourth segment D (Mendel et al. 1997). The spreading direction at this longitude is poorly constrained due to a lack of clear fracture zone traces, but based on regional models of plate motion it is very close to north south (Chu & Gordon 1999), so spreading is orthogonal in segments B and C, but may have an obliquity of up to 1 in segments A and D. 27 2'S 8 CAM12 D 27 'S CAM11 A 12 B C 16 CAM 'S 6 'E 66 'E 66 2'E 66 'E Depth (m) Figure 2. Bathymetry of survey area. Thin solid lines mark free air gravity anomalies (Sandwell & Smith 1997) contoured at 2 MGal interval. Thick solid lines mark seismic shooting tracks and black squares mark ocean bottom hydrophone (OBH) positions. Open square marks OBH which failed to record data within km range due to a tape failure, so does not contribute to this study. Thick dashed lines mark segment boundaries inferred from bathymetric, magnetic and seismic data, and red symbols mark magnetic isochron picks from Cannat et al. (2): circles mark central magnetic anomaly, diamonds mark anomaly 2A and triangles mark anomaly. White triangles mark dredge sites (Robinson et al. 1996, 21). C 26 The Authors, GJI, 166, 1 17 Downloaded from

4 18 T. A. Minshull, M. R. Muller and R. S. White SEISMIC DATA Wide-angle seismic data were acquired using a -gun, 71 L tuned airgun array towed at 1 m depth and fired at s intervals to give a nominal shot spacing of m, and nine ocean bottom hydrophones (OBHs), one of which failed to record useful data. Airgun shots were recorded simultaneously on an 8-channel, 8 m hydrophone streamer. Navigation and timing were from GPS. OBH clock offets were checked before and after deployment and a linear drift assumed between these times, leading to timing uncertainties < ms except in the case of OBH6, where because of a clock jump during deployment the clock drift had to be estimated from previous deployments and the offset estimated from examination of the direct water wave arrival. OBH positions on the seafloor were estimated from direct water wave arrival times. The process of repositioning the OBHs is subject to a trade-off between instrument depth, which is poorly known due to the rough topography and the lack of GPS-navigated swath bathymetric data, and the off-line distance of the instrument. For simplicity, off-line distances were assumed to be negligible except where they were constrained by a crossing profile. Typically, a root-mean-square (rms) misfit between observed and predicted direct water wave arrivals of ms remained after OBH relocation, which may be taken as representative of the traveltime uncertainty of crustal arrivals due to positional uncertainties. Data were acquired along a grid of profiles approximately parallel to and perpendicular to the ridge axis. Results from one of the axis-parallel profiles, labelled CAM116 in Fig. 2, were presented by Muller et al. (1999). Here we synthesize these data with those of a further 11 km axis-parallel profile, CAM11, and a km flowline profile CAM12. The combined dataset allows us to study both along-axis variations in crustal structure and their evolution over time. Other profiles were shorter, had fewer OBHs and gave poorer constraints on crustal structure, so are omitted here. Seismic reflection data show no evidence for pockets of sediment overlying oceanic basement, so any sediment cover is likely to be everywhere thinner than a few tens of metres. OBH data are noisy due to scattering from a hard, rough seabed, and arrivals are extremely complex due to the large seafloor topography (Figs and ). Phase identification was only possible after some initial traveltime modelling. High amplitude turning rays in the upper crust, labelled, emerge as first arrivals at offsets of 2. to. km and persist, with decreasing amplitude, to 8 12 km offset. Turning rays from the lower crust, P, are weaker and rarely observed, indicating lower velocity gradients. More commonly, a gap in range and traveltime is observed between the arrivals and highamplitude, long-range arrivals interpreted as Moho reflections and labelled (e.g. Fig. a). These reflections first appear at ranges of 6 12 km and are observed to maximum offsets of 12 km. In some cases the energy gap between and is absent (e.g. Fig c, positive ranges), indicating the absence of a low velocity-gradient lower crust. Elsewhere, direct reflections are absent altogether (e.g. Figs a and b, negative ranges; Fig c, negative ranges), potentially accounted for by a locally steeply dipping or disrupted Moho near segment boundaries. Only OBH19 on line CAM116 showed convincing Pn arrivals turning in the upper mantle, perhaps because here the Moho is horizontal and the crust is thin; these data are shown by Muller et al. (1999). Strong water-column multiples are observed in all record sections; in some cases, crustal and arrivals were clearer in the first water-column multiple than in the primary, and traveltime picks of these phases provided valuable additional constraints on crustal structure. Multiple reflections can appear even where the corresponding primary reflections are absent; the difference probably arises from the steep seafloor and Moho topography, that leads to significant differences in ray paths between primaries and multiples. SEISMIC MODELLING Primary phases were picked at the onset of the first positive amplitude peak, while multiples were picked at the first negative peak to account for phase reversal at the sea surface. Traveltime uncertainties were estimated as ms for arrivals within 6 km range, for which picking uncertainties are small and the uncertainty is dominated by the positional uncertainties. Observed crustal phases at 6 12 km and >12 km range were assigned uncertainties of 6 and 7 ms, respectively, and arrivals were assigned an uncertainty of ms to account for the interference effects of preceding first arrivals. P arrivals were assumed to be present immediately before beyond the Moho triplication point and to contribute to the observed arrival by constructive interference; these arrivals were assigned the same traveltime and uncertainty as. All multiples were assigned uncertainties of ms to account for the uncertainty in water depth where rays emerge, since where the OBH is off the shooting track, the ray emergence point is also off the track. Despite this relatively large uncertainty, the inclusion of multiple picks greatly enhances the resolution of the final velocity models. Velocity models were constructed using the traveltime modelling and inversion approach of Zelt & Smith (1992). Models were parametrized with two crustal layers, representing oceanic Layers 2 and, with linear gradients within these layers and no velocity discontinuity between them. Horizontal node spacings were chosen so that, as far as possible, regions of the model with dense ray coverage had parameter resolutions, as defined by Zelt & Smith (1992), of greater than.. Velocity node spacings were. km at the seabed, 6. km at the Layer 2/ boundary, and 9. km at the Moho. Depth node spacings were 1. km at the seabed and. km elsewhere. The starting model, based on initial forward modelling, consisted of a 2-km-thick Layer 2 with velocities increasing from. to 6. km s 1 and a 2-km-thick Layer with velocities increasing from 6. to 7. km s 1. In regions of the model where ray coverage was sufficiently high, Zelt & Smith s (1992) inversion scheme was applied to three to four parameters at a time; elsewhere, the inversion was unstable and traveltime fits were achieved by forward modelling. In general, only the velocity at the seabed and the thickness of the layers were allowed to vary. The velocity close to the seabed was well constrained by arrivals at km range ( Figs 6). The depth to the Layer 2/Layer boundary was then adjusted to match longer range arrivals, and then the Moho depth was adjusted to match P and arrivals. Except on line CAM116, mantle velocities were constrained only crudely by critical distances. Models were adjusted until the rms misfit of each phase was less than the picking uncertainty (Table 1) and rays were traced to the maximum number of picks; where unconstrained by seismic data, models were also adjusted to fit gravity data (see below). The final models also match the strong constraints provided by the presence and absence of seismic phases. The relatively low values of χ 2, the normalized traveltime misfit, suggest that models may be mildly over-parametrized. The node spacing chosen represents a compromise between under-parametrizing well-constrained regions of the models and over-parametrizing regions that are sampled by few rays. Model parameter uncertainties were estimated using the perturbation approach of Zelt & Smith (1992) in which parameters are varied until C 26 The Authors, GJI, 166, 1 17 Downloaded from

5 Crustal structure of the SWIR 19 West (a) P CAM11 OBH12 East (b) (c) 8 9 P P -2-2 Range (km) P CAM11 OBH1 CAM11 OBH16 Figure. Seismic data from OBHs on line CAM11, bandpass filtered at 6 16 Hz and displayed with a reduction velocity of 8 km s 1 and gain proportional to range. White dots mark predicted traveltimes for every fifth traveltime pick. In each case the upper panel shows the primary arrivals and the lower panel shows the first water-column multiples (a) OBH12 (b) OBH1 (c) OBH16. rms misfits are no longer acceptable. This approach suggests velocity uncertainties of.2. km s 1 in well-constrained regions of the models and a Moho depth uncertainty of around. km in such regions. The perturbation approach assumes that variations in individual model parameters are independent of each other. If velocity and depth errors are cumulative, the overall Moho depth uncertainty is doubled to around.8 km even in well-constrained regions of the models. GRAVITY MODELLING Gravity data were acquired throughout the survey using a LaCoste- Romberg shipboard gravimeter and corrected for modest meter drift using base ties in Mauritius. A cross-over analysis gave a root-meansquare crossover error of. mgal. Gravity data were used as an additional constraint on crustal thickness variations, particularly in regions of the models where the crustal thickness was poorly constrained by seismic data. The observed data contain significant variations due to out-of-plane effects from the rough seabed topography. Prior to creating 2-D density models along our profiles, these -D effects were removed using the approach described by Muller et al. (2) with assumed densities of 2, 286 and kg m for sea water, crust and mantle respectively, and an assumed crustal thickness of km. The magnitude of the correction exceeds 2 mgal in places. However, the component of the correction due to assumed Moho relief does not exceed mgal. Since the off-line C 26 The Authors, GJI, 166, 1 17 Downloaded from

6 T. A. Minshull, M. R. Muller and R. S. White West East (a) CAM116 OBH18 9 South -2-2 North (b) CAM12 OBH2 8 9 South 2 (c) -2-2 North CAM12 OBH Range (km) Figure. OBH data displayed as in Fig. : (a) OBH18 on line CAM116 (b) OBH2 on CAM12 (c) OBH6 on CAM12. seabed topography is well known and likely variations in seabed density are not large enough to change the correction by more than 1 2 mgal, the uncertainty in the corrected gravity data may be assumed to be in the region of mgal. At a typical Moho depth of about 7 km below sea level, a mgal anomaly corresponds to a Moho depth variation of around.9 km over a scale length of km. Lithospheric age varies little along lines CAM11 and CAM116, so any correction for mantle density variation due to lithospheric thermal structure would be smaller than its associated uncertainties. However, line CAM12 samples crust from zero age out to approximately 6 Ma to the north and 8 Ma to the south. Therefore, lithospheric thermal structure is likely to contribute significantly to the gravity anomaly along this line. A thermal anomaly due to lithospheric thickening was computed using the thermal model of Phipps Morgan & Forsyth (1988) and the approach described by Muller et al. (2). In this calculation we assumed that line CAM12 was sufficiently far from significant fracture zone offsets for its thermal structure to be unaffected by such offsets. The computed thermal anomaly had an amplitude of about 2 mgal, and was removed from the observed gravity anomaly prior to modelling. Velocity models were converted to density models using the empirical velocity density relationship for oceanic crust of Carlson & Herrick (199). Density models consisted of constant density polygons with corners corresponding to seismic velocity nodes and mean densities derived from the mean seismic velocity within these polygons. Constant densities of 2 and kg m were used for sea water and the mantle, respectively. Gravity anomalies were calculated using the method of Talwani et al. (199). The density structure was varied to improve the fit to the observed gravity data C 26 The Authors, GJI, 166, 1 17 Downloaded from

7 Crustal structure of the SWIR CAM Depth (km) T-D/6 (s) Depth (km) T-D/6 (s) Depth (km) T-D/6 (s) 9 Multiple Figure. Observed and modelled traveltimes reduced at 6 km s 1 (lower panels) and corresponding ray diagrams (upper panels) for line CAM11. For each phase, rays are traced to every shotpoint for which a pick was made. In the lower panels, vertical lines mark picks, with the line length corresponding to the pick uncertainty, and thin lines mark traveltimes computed by ray tracing through the final model. primarily by varying the Moho depth in regions poorly constrained by seismic data, and any changes were incorporated back into seismic velocity models such that the final density and seismic velocity models satisfy Carlson & Herrick s (199) relationship. RESULTS AND DISCUSSION Comparison with gravity models The final velocity models (Fig. 7) show a crust that is almost everywhere thinner than the 6 7 km typical of normal oceanic crust at faster spreading rates (White et al. 21), and that also varies dramatically along-axis at a segment scale. These profiles fit well the corrected gravity data (Fig. 8). Crustal thickness variations have been determined independently for the 66 E region using swath bathymetric data and a gravity grid constructed from closely spaced shipboard profiles (Cannat et al. 2; 26). The gravity-based approach requires simplifying assumptions to be made about crust and mantle densities, but has the advantage of giving a -D perspective on crustal thickness variations that is unavailable from our sparse seismic grid. The gravity-derived crustal thickness variations, which are low-pass filtered, match well those derived from the seismic data, though the amplitude of gravity-derived variations is generally smaller (Fig. 9). The very short-wavelength thickness variations in the seismic models are artefacts of the parametrization (detailed bathymetry overlying a smooth Moho) and should not be interpreted, but the difference in amplitude may reflect systematically lower densities at segment ends, as suggested by the seismic velocity models (Fig. 7; Minshull 1996). The good agreement between the two approaches suggests either that any systematic deviations in the density structure from that assumed in the gravity analysis are small, or that the gravity contributions of such deviations from the crust and the mantle cancel each other. The poorest fit occurs along line CAM11, where the gravity-derived crustal thickness is systematically smaller, locally by up to 2 km, than that derived from seismic data. Crustal thickness and melt supply The mean crustal thickness for all three profiles is.2 km, broadly consistent with the melt thickness determined by forward modelling of the rare earth element concentrations of basalts dredged from the region (Fig. 2; Robinson et al. 21). Despite their differences in Moho shape, segments B and C on line CAM11 have similar mean C 26 The Authors, GJI, 166, 1 17 Downloaded from

8 12 T. A. Minshull, M. R. Muller and R. S. White CAM Depth (km) T-D/6 (s) Depth (km) T-D/6 (s) Depth (km) T-D/6 (s) Figure 6. Observed and modelled traveltimes and corresponding ray diagrams for line CAM12. Format as for Fig.. crustal thicknesses of. km and. km, respectively, although on the off-axis profile CAM116, segment B is significantly thinner (.2 ±.8 km) than segments A and C (. ±.8 km). Rare earth element inversions performed on Robinson et al. s dataset are consistent with the crustal thickness results in confirming that, during the period that these data sample, melt production was lower in segment B than in segments A and C (White et al. 21). On each of the three profiles (Fig. 7), the seismic velocities are typical of oceanic crustal structure, with velocities rising steeply from 2.. km s 1 to 6. km s 1 and then more gently to 7. km s 1 ; this twolayer structure generates characteristic and distinct phases in the data (Figs ). The Layer 2 thickness of km is typical of that observed in most ocean basins for similar aged crust. The Layer 2 velocity is controlled dominantly by the porosity of the basaltic layer, with the velocity-gradient reflecting a decrease of pore space with depth caused primarily by crack closure with increasing confining pressure. So the Layer 2 thickness reflects more the physical conditions of the upper crust than the melt supply. However, the mean Layer thickness is much less than the km normally observed (e.g. White et al. 1992). The observed crustal thickness represents material lying above a relatively abrupt velocity discontinuity, and the simplest interpretation is that this discontinuity represents a boundary between mafic crustal rocks and ultramafic mantle rocks. However, serpentinized mantle rocks have been recovered from the seafloor at several locations on the axis of the SWIR east of 62 E, including close to, though not within, our survey area (Seyler et al. 2). Therefore, the crustal material lying above the Moho in this area is likely to consist at least partly of serpentinized mantle rocks, as inferred at mid-ocean ridges elsewhere (Cannat 199). Such rocks can have P wave velocities typical of Layer (i.e km s 1 ), or as low as kms 1 (e.g. Miller & Christensen 1997). As such, they can be difficult to distinguish seismically from basaltic and gabbroic rocks. Based on observations of the SWIR and of the Arctic midocean ridges, Dick et al. (2) have suggested that oceanic crust at very slow spreading rates has an entirely different structure than that of normal oceanic crust, instead consisting of isolated basaltic volcanic edifices lying directly on serpentinized mantle. However, whilst such a model may apply elsewhere on the SWIR, the presence of a typical, though thin, oceanic Layer and the variation of its thickness with proximity to segment ends suggests that serpentinite is not the dominant lithology in the lower crust. Conversely, melt may also be trapped in the mantle beneath slowspreading ridges (e.g. Cannat 1996), so that the crustal thickness may underestimate the volume of melt generated in the mantle. Support for this idea has come from an observed correlation between spreading rate and upper mantle seismic velocity in the western North Atlantic (Lizzaralde et al. 2). Our data do not allow us C 26 The Authors, GJI, 166, 1 17 Downloaded from

9 Crustal structure of the SWIR 1 Table 1. Summary of rms misfits and χ 2 errors between observed and modelled traveltimes for lines CAM11 and CAM12. Data for line CAM116 are given by Muller et al. (1999) but repeated here for completeness. Line Phase Number of picks rms misfit (ms) χ 2 error CAM P Multiple, P Multiple All phases CAM P Pn 6.1 Multiple, P Multiple All phases CAM P Multiple, P Multiple All phases to tell whether a significant volume of melt is trapped in the mantle or whether there is a significant volume of mantle rocks in the crust. However, we note that the good match between seismically determined crustal thicknesses and melt thicknesses inferred from basalt geochemistry, both at the SWIR and elsewhere at very slowspreading ridges (White et al. 21), suggests that these volumes are small and/or of similar magnitude to each other, so that the apparent crustal thickening due to the presence of serpentinized ultramafic rocks in the crust balances the apparent thinning due to retention of some melt in the upper mantle. Therefore, in the discussion that follows we assume that the seismic crustal thickness represents well the melt thickness. The origin of the reduced melt thickness remains controversial. Results from forward and inverse modelling of basalt rare earth element compositions suggest that the melt supply to the ridge axis matches that expected due to conductive cooling of a melting zone lying above normal temperature asthenosphere (Robinson et al. 21; White et al. 21). However, based on isostatic arguments and on the isotopic composition of SWIR basalts, Cannat et al. (1999) and Meyzen et al. (2) have suggested instead that the eastern end of the SWIR lies above locally anomalously cold asthenosphere. Axial Segmentation On-axis (line CAM11), the crustal thickness varies from 6 km at segment midpoints to km at segment ends. This pattern is clear in segment B, in the eastern part of segment A and in the western part of segment C, where the profile remains on-axis. Towards the ends of the profile, thickness variations may be influenced also by variations in crustal age (see below). These axial thickness variations are just as large as those observed across the large-offset Atlantis II Fracture Zone at 7 E (Muller et al. 2). Three axial discontinuities are constrained reliably on line CAM116 and two on line CAM11. Segment B is 7 km long on both profiles and segment C is 26 km long on line CAM11. The Moho dips by in the vicinity of axial discontinuities, and crustal thickness V p2 (km s 1 ) Depth (km) V p2 (km s 1 ) Depth (km) V p2 (km s 1 ) Depth (km) CAM116 CAM12 CAM11 Figure 7. Final seismic velocity models for the three profiles modelled, with lower panels showing the seismic velocity at the top of the crust. The velocity contour interval is. km s 1. The dashed lines just below the 6. km s 1 contour mark the base of Layer 2. Ticks annotated along the tops of the profiles mark OBH positions and line crossings. Spreading segments are labelled on lines CAM11 and CAM116 and approximate lithospheric ages are labelled on line CAM12. variations appear to be dominated by variations in oceanic Layer, with constant Layer 2 thickness at the discontinuities. A pattern of thinner crust at magnetically defined segment ends is also consistent with the gravity-derived crustal thickness variations of Cannat et al. (2). If these crustal thickness variations are interpreted as variations in melt supply, they cannot be explained by the cold edge effect of the small axial discontinuities, but rather may indicate focused magmatic accretion at segment centres (Dick 1989; Muller et al. 1999). It is difficult to believe that mantle melting, which extends to depths of 6 km (e.g. White et al. 21) is variable over horizontal length scales as short as 26 km. However, such focused accretion may be explained by -D models of mantle upwelling that include the effects of magma migration along the base of the C 26 The Authors, GJI, 166, 1 17 Downloaded from

10 1 T. A. Minshull, M. R. Muller and R. S. White Gravity Anomaly (mgal) A B C D A B C CAM CAM Ma Ma.6 Ma CAM Figure 8. Results from gravity modelling. Solid lines mark the gravity anomalies predicted by converting the seismic velocity models shown in Fig. 7 to corresponding density models. Dashed lines mark observed gravity and vertical bars mark observed gravity after correction for -D effects. The height of the bars indicates the uncertainty in the correction. For line CAM12, this correction includes also the effects of lithospheric thermal structure. lithosphere (e.g. Magde & Sparks 1997). A more detailed discussion of the application of such models to the SWIR is given by Muller et al. (2). Oceanic seismic Layer is commonly identified with the gabbroic section of the crust and hence with the presence of a magma chamber (e.g. Salisbury & Christensen 1978), though it has been demonstrated that in some regions the top of Layer occurs within the sheeted dyke section (Detrick et al. 199), and at the SWIR the gabbro section may extend locally into seismic Layer 2 (Muller et al. 1997). Drilling through the gabbro section at Ocean Drilling Program Site 7B has shown that, at least at the Atlantis II Fracture Zone, the gabbroic section has been built up from a series of magma intrusions, with magma redistributed by compaction and deformation processes prior to solidification (Dick et al. 2). The virtual absence of Layer at segment ends (Fig. 7), therefore, indicates that the formation of such magma chambers is limited to segment centres. However, the length scale of the segments is such that magma may be distributed from a central injection point, where Layer is thickest, by dyke intrusion to the segment ends, and then extruded intermittently to the seafloor (e.g. Rubin & Pollard 1988). For example, dykes in Iceland have been observed to propagate subsurface Thickness (km) Thickness (km) Thickness (km) CAM11 W CAM116 W CAM12 7. Ma Ma.6 Ma N A B C D A B C Figure 9. Comparison of crustal thickness variations determined from our seismic profiles (solid lines) with the crustal thickness determined along the same profiles by Cannat et al. (2) using gravity data only (dashed lines). The ridge axis is located at km distance in the flowline profile CAM12. Profiles CAM11 and CAM116 are axis-parallel. laterally several tens of kilometres (e.g. Sigurdsson 1987). Such a model would be consistent with limited exposure of serpentinized mantle at the seafloor in regions where less than per cent of the extension has been taken up by dyke intrusion. Variations in melt supply with time Although some of the variations in crustal thickness along the flowline profile CAM12 might be attributed to sampling away from the segment centre in older crust (Fig. 2), it is unlikely that this profile strays more than a few kilometres from the segment centre. Therefore, we interpret the observed thickness variations as corresponding to time variations in the melt supply. The mean crustal thickness along this profile of. km is interpreted as a time-averaged melt thickness for the centre of segment B, and should, therefore, overestimate the mean crustal thickness for the whole of this segment. This mean is however lower than the. km mean thickness for segment B on-axis, suggesting that the current melt production of segment B is higher than typical values during the last 7 Myr Significantly thinner crust ( km) is observed on line CAM12 at 2 and 2 km model distance (Fig. 7c; Fig. 9), or a lithospheric age of Ma, than the zero-age thickness of. km, so at the segment centre the melt production has increased by 7 per cent in the last Myr This reduced crustal thickness for segment B is confirmed by the crossing profile CAM116. Such temporal variations in crustal thickness have previously been inferred from gravity data in our survey area (Cannat et al. 2), and also further west on the SWIR flanks (Mendel et al. 2). The wavelength of crustal thickness variations is up to E S E C 26 The Authors, GJI, 166, 1 17 Downloaded from

11 Crustal structure of the SWIR 1 km, corresponding to 6 Myr, though some shorterwavelength fluctuations are observed west of 7 E (Mendel et al. 2). Similar fluctuations in magma supply, with a period of 6 Myr, have been inferred from gravity data from the Mid-Atlantic Ridge (e.g. Tucholke & Lin 199; Tucholke et al. 1997). Seismic studies in the Atlantic have shown that large-scale melt production can vary significantly over a 2 Myr time interval (e.g. Morris et al. 199; Hosford et al. 21). More detailed studies have shown that there are also local fluctuations in melt supply with wavelengths as short as km, corresponding to periods of less than 1 Myr (Canales et al. 2). Since none of the other profiles traverse Ma crust, we cannot tell whether melt production was also reduced in other segments at Ma. However, the crustal thickness of segment D on line CAM11 is also about km. This crust has a similar age of Ma to the crust in segment C on line CAM116, where the mid-segment crustal thickness is over 6 km (Fig. 9). Hence it seems that temporal variations in crustal production are decoupled between adjacent segments. These variations might be explained by the effects of magma injection along the base of the lithosphere, as discussed above, if the lithosphere geometry varies with time. Alternatively, there may be instabilities in melt extraction from the mantle (e.g. Scott & Stevenson 1986). Further, we note that both segment B on line CAM116 and segment D on line CAM11 have a reduced Layer 2 thickness of 1. km (Fig. 7), and nowhere else on these profiles is the Layer 2 thickness less than 2 km. The regions of very thin crust on line CAM12 are also associated with a Layer 2 thickness of less than 2 km, both at its crossing with line CAM116 and 2 km north of the ridge axis. The reduced Layer 2 thickness may have resulted from enhanced tectonic extension that has affected these segments during a period of reduced magmatism. Upper crustal structure The upper m of the crust is not sampled by turning rays in our experiment, and the 1 2 km seismometer spacing used means that oceanic Layer 2 is not sampled continuously along the profiles. The seismic velocity of the uppermost crust increases systematically with age as voids are filled with hydrothermal minerals (e.g. Grevemeyer & Weigel 1996). A systematic variation of seismic velocity with age based on magnetic anomalies is not observed within our survey area. However, distinct changes in upper crustal velocity are observed across most of the segment boundaries, related perhaps to the different stage of tectonomagmatic cycling within each segment and suggesting that the velocity may depend on the time elapsed since the most recent period of volcanic activity within a segment. The lowest velocities of 2. ±. km s 1 are found in segment A on both lines CAM11 and CAM116, suggesting that the most recent volcanism has occurred here. Recent volcanism also is indicated by the very large topographic swell within the axial valley of segment A (Fig. 2) and by high inferred magnetizations (Sauter et al. 2). Based on magnetic anomalies, however, the crust sampled in segment A is older (..8 Ma) than the zero-age crust in the adjacent segment B where upper crustal velocities of. ±. km s 1 are observed. These observations may be explained if segment B is currently in a phase of amagmatic extension that has lasted at least..8 Myr. In both segment B on line CAM116 and segment D on line CAM11, there is an association between high velocities at the top of the crust and a thin Layer 2; both effects may result from amagmatic extension, which can lead to both enhanced porefilling alteration and exhumation of deeper crustal levels. Velocities reach a maximum of. ±. km s 1, and the maximum crustal age where velocities are constrained is. ±. Ma. The range of seismic velocities is broadly compatible with the range found elsewhere at ridge axes for. Ma crust (Grevemeyer & Weigel 1996). Some anisotropy of uppermost Layer 2 velocities may be present, with values at OBH2 of. km s 1 on line CAM12 and. km s 1 on line CAM116, a difference of 1 per cent in these perpendicularly orientated lines. Although the difference in velocities is barely larger than their uncertainties, such velocities are observed on both sides of OBH2 on the respective profiles, and the higher velocities on line CAM116 are also seen in data from OBH19 (Fig. 7b). The slow direction is perpendicular to the strong tectonic fabric seen in sidescan images from the area (Searle et al. 199), and this fabric provides a plausible explanation for the observed apparent anisotropy, although with our limited experimental geometry we cannot rule out that the possibility that velocity differences are due to lateral heterogeneity. CONCLUSIONS From our study we draw the following conclusions: (1) The seismic velocity structure of the axial crust of the SWIR at 66 E consists of a km-thick high velocity-gradient Layer 2 with velocities increasing with depth from 2.. km s 1 to 6. km s 1. This layer is underlain by a..-km-thick low velocity-gradient Layer with seismic velocities increasing with depth from 6. to 7. km s 1. (2) The mean crustal thickness at 66 E is.2 km. The reduced crustal thickness with respect to that observed at faster spreading ridges may be explained by conductive cooling of normal upwelling mantle. () Crustal thickness varies along-axis by up to a factor of two. The crust is systematically thicker at segment centres and thinner at segment ends. These variations occur primarily in oceanic Layer. Such variations may be explained by delivery of melt from the mantle to segment centres and distribution to segment ends by lateral dyke propagation. () Crustal thickness varies also with age, with up to 7 per cent variation from its minimum value over a period of Myr Again, thickness variations occur dominantly within Layer. () The seismic velocity at the top of Layer 2 varies significantly across our survey area, but this variation is weakly correlated with crustal age based on magnetic anomalies, and may instead be linked to the age of the most recent magmatic episode. ACKNOWLEDGMENTS Seismic data acquisition was funded by UK Natural Environment Research Council grant GR/888. TAM was supported by a Royal Society University Research Fellowship and MRM by a Carl and Emily Fuchs Foundation Overseas Scholarship and a CVCP Overseas Research Studentship. We thank all who sailed with us on RRS Discovery cruise 28 for their assistance, and Bob Fisher, Marc Munschy, Daniel Sauter and Mathilde Cannat for allowing us access to charts and unpublished data during this work. We thank two anonymous reviewers for constructive comments. Figures were produced using the GMT package of Wessel & Smith (1998). C 26 The Authors, GJI, 166, 1 17 Downloaded from

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