5.1. Upper Crust
 Oceanic layer 2A has been studied extensively at the East Pacific Rise, where it has been interpreted as a layer of extrusive lava flows [e.g., Vera and Diebold, 1994; Harding et al., 1993; Christeson et al., 1996; Hussenoeder et al., 2002b] and to a lesser degree at the Juan de Fuca Ridge [e.g., Cudrak and Clowes, 1993; McDonald et al., 1994]. However, since layer 2A rarely generates first arrivals in seismic data from sea surface shots, at slow-spreading ridges, where scattering from the seafloor is stronger, arrivals from this layer are usually obscured. In the absence of such arrivals, regularized tomographic inversions generate a smooth velocity-depth profile in which the velocity rises steeply in the upper 1 km of the crust [e.g., Canales et al., 2000a]. The thickness and seismic velocity of layer 2A determined by our study match well those observed at and near the ridge axis at slow- to intermediate-spreading ridges, and a few kilometers off-axis at the East Pacific Rise (Figure 15a). There is a remarkably good agreement of velocities in the upper 1 km with those obtained by Hussenoeder et al. [2002a] from 0–2 Ma crust on the Mid-Atlantic Ridge at 35°N. At 8–9°S, the mean layer 2A thickness is greater and there is more variation in this thickness, but both differences may be simply a consequence of the greater age range and complex spreading history of our study area. The thickness of the transition zone at the base of layer 2A matches better the on-axis model of Hussenoeder et al. [2002a] than their off-axis model; it is unlikely that the small differences between these models could be resolved by our noisier data. For both these Mid-Atlantic Ridge studies, the velocity at the top of layer 2B is considerably higher than those observed at Reykjanes Ridge and at Mohns Ridge (Figure 15a).
Figure 15. (a) Thick solid line marks representative structure of the uppermost crust, constructed using a layer 2A velocity and thickness and layer 2B velocity based on mean values inferred from sonobuoy data and a transition thickness between the layers based on multichannel data. Other profiles shown are from the ridge axis at the East Pacific Rise at 9°30′S (EPR) [Hussenoeder et al., 2002b, Figure 19], the Juan de Fuca Ridge (JDF) [Cudrak and Clowes, 1993; McDonald et al., 1994], the Mid-Atlantic Ridge at 35°N (MAR) [Hussenoeder et al., 2002a, Figure 16], and the Reykjanes Ridge (RR) [Smallwood and White, 1998], and from 1.5 Ma crust at the Mohns Ridge (MR) [Klingelhoefer et al., 2000]. (b) Upper and lower bounds on porosity inferred from velocity model represented by thick solid line in Figure 15a using the approach of Berge et al.  (see text).
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 The velocity variations in the upper 1 km of the crust may be interpreted in terms of variations in porosity and pore aspect ratio distribution. For the case of porous basalt, where seismic velocities of the matrix and pore space are very different, Hashin-Shtrikman bounds are too broad to be useful. Narrower and more useful bounds were developed by Berge et al.  using an assumed pore aspect ratio distribution. Application of their approach suggests a porosity of 23–30% at the seabed, decreasing rapidly to ∼12–14% at the top of layer 2B (Figure 15b). This porosity decrease may represent the base of the extrusive lavas, or alternatively a porosity change within the extrusive sequence [e.g., Harding et al., 1993; Cudrak and Clowes, 1993]. The layer 2A thickness matches direct observations of a 280–420 m extrusive layer at the south wall of the Vema transform [Auzende et al., 1989], and the appearance of dykes at ∼500 m depth in old Atlantic crust at Deep Sea Drilling Project holes 417D and 418A [Bryan et al., 1979]. The large variations (±300 m) in the thickness of layer 2A (Figure 5), which are similar to those observed at the Juan de Fuca Ridge [Cudrak and Clowes, 1993; McDonald et al., 1994], may then be explained by a combination of the episodicity of magmatism, the preferential accumulation of extrusive lavas in topographic depressions, and the dismemberment of the crust by extensional tectonics.
 Velocity-depth profiles fall generally within White et al.'s  envelope for 0–7 Ma Atlantic oceanic crust away from fracture zones and mantle plumes, though velocities at the top of layer 2B are slightly higher (Figure 16). The difference may be largely an artifact of modeling approaches which assume a smooth velocity gradient in the uppermost crust. The thickness of layer 2 is generally 2–3 km across the survey area, and is also typical of Atlantic oceanic crust. The velocity structure also resembles that determined for 6–7 Ma crust in the vicinity of Ascension Island [Klingelhoefer et al., 2001].
Figure 16. Solid lines mark velocity-depth profiles extracted from the two-dimensional velocity models at each sonobuoy location and terminated at the maximum depth constrained by the data in the region of that sonobuoy. Shaded region marks the range of velocities for 0–7 Ma Atlantic oceanic crust as compiled by White et al. .
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 The velocity at the top of layer 2B varies systematically across our survey area (Figure 17). First there is a decrease in velocity from south to north, which likely reflects the presence of less altered, higher porosity material to the north, consistent with inference from gravity anomalies that the spreading center may have been abandoned later in the north [Bruguier et al., 2003]. Velocities are lowest near the pseudofaults, where extensive large-scale porosity due to tectonic strain may be expected [e.g., Kleinrock and Hey, 1989]. A similar velocity reduction was observed at the pseudofault marking the southern limit of ridge segment OH-1 at 35°N by Hosford et al. . Secondly, on lines CAM83 and 84, there is a small but significant increase in this velocity with crustal age, at a rate of ∼0.1 km/s per Ma (Figure 18). On line CAM88 this variation with age, if present, is obscured by the velocity decrease toward the pseudofaults. Such variations in layer 2B velocity near ridge axes often can be obscured by wide-angle data analyses which require smooth variations of velocity with depth since large changes in the velocity and thickness of layer 2A are also present [e.g., Carlson, 1998]. For example, at 35°N the mean velocity of layer 2 increases by 0.7–0.8 km/s between zero-age and 2 Ma crust, but this increase may be largely due to changes in layer 2A velocity [Hosford et al., 2001]. A similar, but more rapid increase with age has been observed at Reykjanes Ridge [Smallwood and White, 1998], while no such systematic variation is apparent at the East Pacific Rise [Grevemeyer et al., 1998].
Figure 17. Spatial variation of the seismic velocity at the top of layer 2B, illustrated by contouring a smooth surface fit to velocity values at model nodes using the minimum curvature algorithm of Smith and Wessel . Thin solid line marks shooting track and thick solid line marks the segment boundary. Dashed lines mark active and abandoned spreading centers as in Figure 1. Regions with velocity below 4.7 km/s are shaded.
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Figure 18. Filled circles mark values of seismic velocity nodes at the top of layer 2B, with their estimated uncertainty of ±0.1 km/s. Dashed lines mark best fit linear regressions of velocity as a function of crustal age.
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5.2. Lower Crust
 Layer 3 velocities are remarkably uniform across the survey area, with good fits achieved to both travel times and amplitudes using a constant velocity on all lines except for the northern end of line CAM90. A similar uniformity of layer 3 velocities has been observed in 0–2 Ma crust at 35°N on the Mid-Atlantic Ridge [Hosford et al., 2001]. As elsewhere on the Mid-Atlantic Ridge and Reykjanes Ridge [e.g., Hooft et al., 2000; Smallwood and White, 1998], wide-angle Moho reflections are observed from beneath the current spreading axis (Figures 8 and 13).
 Reduced velocities at the top of layer 3 in the northern part of line CAM90 (Figure 12) are not detected on either the adjacent profiles or the crossing profiles, so must be limited to a region only a few kilometers wide around the active spreading center. Velocities are reduced by up to ∼0.4 km/s, a change that is too large to be accounted for by plausible changes in porosity. The velocity reduction may result from a thermal anomaly due to a recent intrusive event, or from the presence of a small amount of partial melt. If the P wave velocity of layer 3 gabbros varies with temperature at a rate ∂V/∂T = −0.57 × 10−3 km s−1 K−1 [Christensen, 1979], the maximum velocity anomaly corresponds to a temperature anomaly of around 700 K. If the off-axis temperature at the top of layer 3 is ∼200–300°C. as suggested by thermal models incorporating hydrothermal cooling [e.g., Henstock et al., 1993], no partial melt is required to explain the anomaly, though the presence of partial melt could contribute to it.
 The crustal thickness increases systematically from north to south (Figure 19), toward the center of the abandoned spreading segment, the southern half of which remains active [Bruguier et al., 2003]. The total variation is ∼4 km, from 6 km to 10 km. The magnitude of the variation is similar to that seen along the magmatically robust OH-1 segment at 35°N [Hooft et al., 2000], but the mean crustal thickness is about 2 km greater due to the presence of a significant mantle melting anomaly at 9–10°S [Bruguier et al., 2003]. The layer 2 thickness also increases southward, from 2.0–2.5 km on line CAM88 to ∼3 km on lines CAM83 and CAM84, but as elsewhere [Mutter and Mutter, 1993], changes in crustal thickness appear to be taken up dominantly by changes in layer 3 thickness. The southward increase in crustal thickness coincides with increasing enrichment in incompatible trace elements of basalts dredged at the ridge axis [Schilling et al., 1985]. We interpret the thickness variations as due to the presence of a more fertile mantle source to the south of our survey area [Bruguier et al., 2003], as has been suggested for the increased crustal thickness of the OH-1 segment [Niu et al., 2001].
Figure 19. (a) Contour plot of crustal thicknesses across the survey area. The crustal thickness grid was generated by applying the minimum curvature algorithm of Smith and Wessel  to values at all model nodes, and therefore incorporates constraints from synthetic seismogram and two-dimensional gravity modeling as well as those from PmP travel times. However, parts of the grid lying more than 10 km from a PmP travel time constraint are masked. Contour interval is 1 km. Solid lines mark shooting track; lines are thickened where thickness is constrained by PmP travel times. Dashed lines mark active and abandoned spreading center as in Figure 1, and thick solid line marks segment boundary [Bruguier et al., 2003]. (b) Difference between crustal thickness shown in Figure 19a and crustal thickness inferred from residual mantle Bouguer anomalies [Bruguier et al., 2003], contoured at 1 km intervals. A positive value indicates that the gravity-derived crustal thickness is greater. Areas where the difference is greater than 1 km are shaded.
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 If a thick, hot mantle plume head were present to the south of our survey area, we would expect to see a greater southward increase in crustal thickness derived from mantle Bouguer gravity anomalies that inferred from seismic data. No such trend is inferred (Figure 19b). However, there is a large and well constrained discrepancy in crustal thickness around the abandoned spreading center on lines CAM84 and CAM88, where the gravity inversion predicts a crust 2–3 km thicker than suggested by the seismic models. A difference of 1–2 km is also present around the segment center of the active spreading center. These differences are unlikely to be due to changes in mean crustal density since only small variations in mean crustal density are expected for these crustal thicknesses [Minshull, 1996], and a large change in density would be expected to be accompanied by a corresponding change in seismic velocity. Therefore we attribute them to reduced mantle densities and higher temperatures than those predicted by the thermal model, which neglects the effects of ridge jumps and ridge propagation. Although the ridge jump likely occurred at depth before its surface expression [Bruguier et al., 2003], hot, low-density asthenospheric mantle may remain at relatively shallow depths beneath the recently abandoned spreading center, particularly to the north. Hot asthenosphere may be present also beneath the active segment center.
 The thinnest crust is observed beneath the pseudofault trough at the northern end of line CAM86 (Figure 12), beneath the new pseudofault at the northern end of line CAM91, and beneath the propagating rift tip at the northern end of line CAM96 [Bruguier et al., 2003]. A reduced crustal thickness in these locations might be expected due to the juxtaposition of the spreading center against older lithosphere, but the age discontinuity involved is in all cases relatively small (generally <1 myr). The reduced thickness is therefore analogous to that which occurs at small-offset fracture zones on slow-spreading ridges [e.g., Minshull et al., 1991; Tolstoy et al., 1993], and may be due to focused magmatic accretion at the segment center. Our observations contrast with those of Hasselgran et al. , who infer from seismic reflection data the presence of thickened crust beneath pseudofaults with similar age differences on the flanks of the Juan de Fuca ridge.
 The inferred termination points of the pseudofaults (Figure 1) suggest that the abandoned spreading center began propagating at ∼2.5 Ma. This age coincides with the age of the crust where the Moho steps down to the west on line CAM83. Perhaps, then, the onset of excess melt production caused ridge propagation to initiate as a result of excess gravitational-spreading stresses [Phipps Morgan and Parmentier, 1985]. However, it is clear from the northward decrease in crustal thickness that such stresses were not required to maintain ridge propagation once it had been initiated.
 Our crustal thickness measurements also allow an examination of the relationship between crustal thickness and axial morphology at constant spreading rate. In the south of our survey area, where seismic data indicate that the axial crust is 10 km thick and gravity data indicate a similar thickness beneath the abandoned spreading center, both the active and the abandoned spreading centers have rifted axial highs (Figure 20). Where the crustal thickness is reduced to 7 km, there is a deep axial valley. However, at intermediate thicknesses of 8–8.5 km, the morphology varies widely. Where line CAM84 crosses the abandoned axis, there is a clear axial high, though it is broad and subdued in amplitude (Figures 1 and 20). An axial high is also present at 33°S on the Mid-Atlantic Ridge, where the spreading rate and crustal thickness are similar [Tolstoy et al., 1993]. Conversely, where this line crosses the new axis, and the crust is as thick or perhaps slightly thicker (Figures 7 and 19), there is a clear axial valley (Figures 1 and 20). The presence of an abandoned axial high on line CAM84 favors a flexural origin for the high such as that suggested by Buck . The presence of an axial valley at the new spreading centre may also be consistent with such an origin since the release of flexural stresses generated by the abandoned axial high will favor the formation of a valley until sufficient flexural stress has been accumulated from new accretion.
Figure 20. Bathymetric profiles across (a) the abandoned spreading center and (b) the active spreading center. Solid profiles are from line CAM83, dashed profiles are from line CAM84, and dotted profiles are from line CAM88. Each profile is labeled with the crustal thickness in the vicinity of the axis, rounded to the nearest 0.5 km, and for clarity profiles for lines CAM84 and CAM88 have been shifted vertically by 1000 and 2000 m, respectively.
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