Geophys. J. R. astr. Soc. (1984) 79, 779-798

Anomalous seismic crustal structure of oceanic fracture zones

Robert s. White Bullard Laboratories, Department of Earth Sciences, Madingley Rise, Madingley Road, Cambridge CB3 OEZ

Robert s. Detrick Graduate School of Oceanography, University of Rhode Island, Kingston, RI 02881, USA

Martin c . Sinha Bullard Laboratories, Department oj Earth Sciences, Madingley Rise, Madingley Road, Cambridge CB3 OEZ

Marie H. COrmier Graduate school of Oceanography, University of Rhode Island, Kingston, RI 02881, USA

Received 1984 April 14; in original form 1983 December 2

Summary. The seismic structure of crust found within fracture zones falls outside the range of velocity structures observed for normal oceanic crust in the North Atlantic. The crust in fracture zones is frequently very thin and is characterized by low crustal velocities and by the conspicuous absence of a refractor with a velocity typical of oceanic layer 3 . Anomalous crust is present in both large- and small-offset fracture zones. Since they are among the most common tectonic features in the ocean basins, and are particularly closely spaced on slow-spreading ridges, fracture zones represent a major source of seismic crustal heterogeneity. We interpret the anomalous crust as a thin, intensely fractured, faulted and hydrothermally altered basaltic and gabbroic section overlying ultramafics that, in places, are extensively serpentinized. The unusually thin crust found within fracture zones and the gradual crustal thinning over a distance of several tens of kilometres on either side of the fracture zones can be explained by two main processes; firstly the cold lithosphere edge opposite the spreading centre at the ridge- transform intersection modifies the normal intrusive and extrusive processes of the spreading centre leading to the accretion of an anomalous and thin igneous section; and secondly each spreading ridge segment is fed from a separate subcrustal magma supply point, so as the magma flows laterally down the spreading centre it generates a crustal section of decreasing thickness, culminating in the very thin crust of the fracture zones at either end of the ridge segment.

780 R. S. White et al. 1 Introduction

A relatively simple picture of the seismic structure of the ocean crust has emerged from refraction studies over the past two decades. Raitt (1963) characterized the oceanic crust by two main layers: layer 2 (the ‘volcanic’ layer) with a velocity of 5.07 f 0.63 km s-l overlying layer 3 (the ‘oceanic’ layer) with a velocity of 6.69 f 0.26 km s-’. Further work refined this interpretation by subdividing these layers into thinner units (e.g. layers 2A, 2B, 2C of Houtz & Ewing 1976; layer 3B of Sutton, Maynard & Hussong 1971). More recently, modelling of seismic amplitudes (e.g. Spudich & Orcutt 1980) has shown that the variation of velocity with depth in the oceanic crust is represented better by velocity gradients than by a few homogeneous layers of contant velocity. The uppermost crust (layer 2 ) displays a rapid increase in compressional wave velocities from typically 3-4 km s - l at the seafloor to 6.5-6.7 km s-l at depths of about 2 km. Velocity gradients within the lower crust (layer 3) are much smaller, typically only 0.1-0.2 s-’. The crust- mantle boundary is of variable thickness, occurring over a depth interval of up to 2 km. Total crustal thicknesses are typically 5-8 km.

Figure 1. Dots show location of North Atlantic seismic velocity-depth profiles compiled in Fig. 2 and of fracture zone?, shown in 13gs 3~-5 and discussed in text.

Seismic structure of fracture zones 78 1

Results from recent geological and geophysical investigations of several north Atlantic fracture zones suggest that the oceanic crust in and near these features is anomalous, hnerally falling well outside the bounds of seismic crustal variability determined from nQrmal oceanic crust. Since fracture zones occur on average once every 40-80 km along the Mid-Atlantic Ridge, their crustal structure is an important source of seismic crustal hetero- geneity on a large scale and may also provide important constraints on crustal accretionary processes along the spreading centre. In this paper we summarize seismic refraction results from North Atlantic fracture zones (Fig. 1 ) and discuss several mechanisms that may explain their unusual crustal structure.

0-311a

2 Fracture zone crustal structure

Many independent lines of evidence indicate that abnormal crust is associated with oceanic fracture zones. The most basic evidence is that fracture zones leave a distinct bathymetric trace, frequently but not always a deep trough, sometimes with an adjacent ridge, which can be followed thousands of kilometres across ocean basins like the North Atlantic even when no resolvable age-offset is present (e.g. Schouten & Klitgord 1982). It has also been recognized for many years that the walls bordering many large Atlantic fracture zones (e.g. Kane, Oceanographer, Romanche, Vema) expose a diverse suite of gabbroic and ultramafic rocks that are not commonly found elsewhere in the ocean basins. Francheteau et al. (1976) suggested the presence of a thinner than normal volcanic layer in fracture zones in order to explain the exposure of these rocks which are generally associated with the lower crust and upper mantle. They noted that the great relief found in fracture zone escarpments is actually the product of a large number of small throw faults (less than 500 m) linked by terraces and talus ramps, rather than by a single large fault scarp with a throw of several thousands of metres. The outcrop of rocks typical of the lower crust and upper mantle high on these escarpments, documented by recent submersible studies of the Kane (Karson & Dick 1983) and Oceanographer (OTTER 1984a, b) fracture zones, is difficult to explain if normal crustal thicknesses (5-8 km) are present, leading to the suggestion that the crust accreted near transforms is extremely thin (e.g. Fox 1978; Stroup & Fox 1981).

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Compressional wave seismic velocity (km/s) 2 4 8 8 1 0

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Figure 2. Stacked compilations (by crustal age) of seismic velocity versus depth profiles constrained by synthetic seismogram analyses of seismic refraction lines using ocean bottom seismometers and sonobuoys. For sources of data and individual profiles see White (1984). Locations of profiles shown by dots in Fig. 1.

782 R. S. White et al. Geophysical evidence also suggests that the crust within fracture zones differs from that

on either side. Large amplitude gravity anomalies are frequently associated with fracture zones indicating that the mass distribution in and around these features is quite hetero- geneous (Cochran 1973; Robb & Kane 1975; Louden & Forsyth 1982). While these anomalies are in part due to the effect of juxtaposing lithosphere of different ages (Sibuet & Veyrat-Peinet 1980), large variations in crustal thickness and/or density must also be present. The effective magnetization of crustal rocks within fracture zones is very variable, sometimes being very small (Collette et al. 1974: Twigt, Slootweg & Collette 1979), and sometimes strongly enhanced (Twigt et a!. 1983), again indicating the presence of an anomalous crust. However, by far the best evidence for the anomalous crustal structure of Atlantic fracture zones has come from several recent seismic refraction experiments. In the next two sections we review these results and show that seismically anomalous crust is present at both large- and small-offset fracture zones in the North Atlantic.

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a ) KANE F. Z. Figure 3. Compilation of seismic velocity versus depth determinations from refraction lines across large- offset North Atlantic fracture zones. See Fig. 1 for location of profiles. Shaded area shows the maximum bounds of ‘normal’ oceanic crustal structure derived from the compilations of all North Atlantic solutions over crust older than 2 Ma (Fig. 2). (a) Kane fracturc zone, dashed lines from Detrick & Purdy (19801, solid lines from Cormier et al. (1984); (b) Vema fracture zone, dashed lines from Ludwig & Rabinowitz (19801, solid lines from Detrick et al. (1982); (c) Oceanographer fracture zone, dashed lines from Fox et al. (1976), solid lines from Sinha & Louden (1983) and this paper; (d) Northeast Atlantic fracture zone, dashed line, from White & Matthews (1980), solid line from fracture zone I (this paper). Fracture zone results are derived primarily from travel time fitting, so the stepped constant velocity layer solutions reflect the interpretation technique rather than implying that the crust contain$ uniform velocity layers.

Seismic structure of fracture zones 783

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784 R. S. White et al.

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2.1 L A R G E - O F F S E T F R A C T U R E Z O N E S

The velocity-depth profiles determined from seismic refraction experiments along Atlantic fracture zones are shown in Fig. 3, where they are compared with the range of structures determined from more 'normal' oceanic crust. The bounds of the 'normal' profiles were derived from a compilation by White (1984), summarized in Fig. 2, of all available velocity versus depth profiles in the North Atlantic older than 3 Ma constrained by synthetic seismogram modelling. Velocity solutions for the large-offset fracture zones, namely the Kane fracture zone (Detrick & Purdy 1980; Cormier, Detrick & Purdy 1984), the Oceanographer fracture zone (Sinha & Louden 1983), and the Vema fracture zone (Ludwig & Rabinowitz 1980; Detrick et al. 1982) all fall outside the bounds of normal crust in the North Atlantic.

The most remarkable feature of the crust found in fracture zones is that total crustal thicknesses are in places only 2-3 km, about half the typical thickness of crust in the ocean basins. This very thin crust is apparently confined to a relatively narrow zone (less than 10km wide) centred on the fracture zone itself (Fig. 4). However, there is an additional, more gradual crustal thinning that may extend up to several tens of kilometres from the fracture zone. This is best documented for the Oceanographer fracture zone, but is also apparent south of the Kane fracture zone (Fig. 4). At the Kane fracture zone thin crust may also extend beneath the transverse ridge which forms the northern wall of the fracture zone.

Thin crust is apparently a relatively common feature in these fracture zones. It has been found in the transform portion of the Oceanographer fracture zone and along most of the 200 km long segment of the Kane fracture zone trough that has been investigated.

Seismic strucfure offracture zones 785

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Distance (km) Flgure 4. Seismic cross-sections at vertical exaggeration of 2.5 : 1 perpendicular to: (a) Vema fracture zone (Detrick ef al. 1982) from separate reversed seismic refraction lines; (b) Oceanographer fracture Zone line B across the active transform (Sinha & Louden 1983) from ray tracing; ( c ) Oceanographer fracture zone and fracture zone I line C1 across the inactive portions of the fracture zones (from Fig. 5) from ray tracing; (d) Oceanographer fracture zone line C2 across the inactive portion of the fracture Zone (Sinha & Louden 1983) from ray tracing; ( 3 ) Kane fracture zone (Detrick & Purdy 1980) from delay time analysis.

786

Unusually thin crust is also present beneath the nodal deep at the intersection of the Kane fracture zone and the Mid-Atlantic Ridge (Cormier et af. 1984). While very thin crust is common in these large-offset, slow-slipping transforms, more normal crustal thicknesses of 6-7 km (albeit often with anomalous velocities) have also been observed along parts of the Oceanographer, Kane and Vema fracture zones.

A second distinctive feature of fracture zones is that they are characterized by relatively low crustal velocities. Thus despite an elevated Moho, mantle delay times beneath fracture zones may be only slightly less than that of normal oceanic crust. Detrick ef af. (1982) have estimated that the average crustal velocity in fracture zones is about 20 per cent less than the average velocity (about 6 km s- l ) of a normal oceanic crustal section. The presence of relatively low crustal velocities and densities beneath transform valleys together with greater than normal water depths explains why gravity profiles across fracture zones do not show a large mass anomaly despite the presence, in some cases, of an anomalously thin crustal section (Louden & Forsyth 1982; Detrick et al. 1982).

Thirdly, fracture zone crust is unusual in that a refractor with a velocity typical of oceanic layer 3 is rarely observed, although a well-developed crust-mantle boundary is usually present. This observation is particularly significant in that layer 3 is the most consistently observed feature of the seismic structure of normal oceanic crust. Velocity gradients within layer 3 are very small, typically only 0.1-0.2 s- ' , resulting in a relatively narrow range of observed velocities (Fig. 2). The absence of a typical layer 3 refractor strongly suggests that the accretionary processes in fracture zones are fundamentally different from those which form most oceanic crust.

R. S. White et al.

2.2 SMALL-0I :FSET F R A C T U R E Z O N E S

Although there is substantial evidence that anomalous crust is found in large-offset fracture zones, there have been fewer experiments on fracture zones with small offsets or on fast- spreading ridges. The evidence which exists, however, indicates that the crust within even small Atlantic fracture zones is different from that which is typical of most of the North Atlantic in similar ways to the crust within large-offset fracture zones. This is an important observation since it suggests that this anomalous fracture zone crust is much more common than previously suspected. It also indicates that the magnitude of the thermal edge effect at ridge-transform intersections which would be expected to have little influence across small- offset transforms, is not the only factor responsible for the accretion of unusually thin crust in fracture zones.

One of the best available seismic refraction profiles for investigating the structure of a small-offset fracture zone was observed by Sinha & Louden (1983) along their line C1. This profile, shot parallel to the ridge axis, was intended primarily to study the large (about 125 km) offset Oceanographer fracture zone. However, the profile also crossed the much smaller (less than 35 km offset) fracture zone I which lies some 90 km to the south. We have re-analysed data from this profile paying particular attention to the structure of fracture zone I.

Our interpretation of the velocity structure beneath fracture zone 1 is based on arrivals from 25 explosive shots received at four internally recording sonobuoys deployed along the 107 km long profile (see fig. 2 of Sinha & Louden 1983 for map of shot and receiver locations). Charge sizes were 23, 102 or 160 kg fired at depths of between 76 and 162 m. Sonobuoy 1 was located at the northern end of the line, 10 km south of the Oceanographer fracture zone valley; sonobuoy 2 was over normal crust about one-third of the distance south along the line; sonobuoy 3 was positioned two-thirds of the way along the line; and

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Seismic structure oj,fracture zones 79 1

sonobuoy 6 was located at the southern end, 15-20 km south of fracture zone I. Reduced record sections from the sonobuoys are shown in Fig. 5(a-d).

The travel times of the arrivals were modelled using a ray-tracing computer program for a laterally varying medium (cerveny & Ravindra 1971). The theoretical travel times through our best-fitting model are shown as ticks superimposed on the data in Fig. 5(a-d). The ray-paths used in the modelling are shown in Fig. 6(a) and our final velocity model is contoured in Fig. 6(b). The structure beneath fracture I is well constrained by many crossing ray-paths returned both from the crust and from the upper mantle (Fig. 6a).

Many of the features of the anomalous crust present in fracture zone I are similar to those found in large offset fracture zones discussed in the previous section, and indeed are similar to those of the Oceanographer fracture zone at the northern end of the profile.

First, the crustal thickness beneath fracture zone I is only about 4.5 km, significantly less than the 5-6 km thick normal crust found between fracture zone I and the Oceano- grapher fracture zone to the north (Fig. 6b). Secondly, the lower crustal section beneath fracture zone I, like that beneath the Oceanographer fracture zone, displays abnormally low seismic velocities. This is particularly well shown by the velocity contours in Fig. 6(b), where it can be seen that the velocities beneath fracture zone I (6.25 km s-l increasing with depth to 6.55 km s-'), fall completely outside the normal layer 3 velocities (6.8- 7.2 km s-'), in the intervening, unperturbed crust. Thirdly, the abnormal crust extends for about 20 km on either side of the trace of the fracture zone fault. This is evident both from the width of crustal thinning and, more dramatically, from the lateral extent of the anomalously low crustal velocities (Fig. 6b). The edges of the anomalous region are characterized by horizontal velocity gradients comparable in magnitude to the vertical velocity gradients of 0.1-0.2 s - l found in the lower crust. The lateral extent of the anomalous crust is similar for both the small-offset fracture zone I and for the large-offset Oceanographer fracture zone. Lastly, we note that although they are less well constrained than the crustal velocities, there are velocity variations in the upper mantle that are apparently related to the broad region of abnormal crust associated with the fracture zones: beneath the fracture zones the mantle velocity is about 7.7 km s-*, rising to 8.0 km s-.' beneath the normal crust. Although there is a deep trough marking the position of fracture zone I along most of its length, at the particular point crossed by our profile it has no topographic expression despite its anomalous crust.

A two-ship, wide-aperture multichannel reflection profile obtained across the small- offset (about 20 km) Blake-Spur fracture zone, SW of Bermuda, provides independent evidence for the existence of anomalously thin crust at small North Atlantic fracture zones (Mutter, Detrick & NAT Study Group 1984). Reflections from the crust-mantle boundary, typically observed at about 2 s two-way time below the basement reflection in much of the western North Atlantic, rise to only 0.8 s beneath the trough of the Blake-Spur fracture zone, corresponding to a crustal thickness of only 2-2.5 km. This is the first time that Moho reflections have been observed across an oceanic fracture zone and this profile clearly shows that the very thin crust is present over only an approximately 15 km wide zone centred on the fracture zone. The transition between this thin crust and more normal crustal thicknesses O q either side of the fracture zone is quite abrupt, although a more gradual crustal thinning extends up to several tens of kilometres from the fracture zone. The magnitude and lateral extent of the anomalously thin crust at the Blake-Spur fracture zone are comparable to that at larger-offset fracture zones like the Kane and Oceanographer.

Further evidence for anomalous crust being associated with very small offset fracture 'zones comes from a detailed seismic experiment reported by White & Matthews (1980), across an unnamed fracture zone in the north-eastern Atlantic (Fig. 1). The offset on this

792 R. S. White e t al.

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794 R. S. White et al. fracture zone, mapped by a closely spaced grid of survey lines over a 50 km square region of 40-50 Ma old crust varies from little more than zero to a maximum of only 15 km. Nevertheless fan shoots across the fracture zone show a marked decrease in delay times of seismic rays turned in the lower crust over a 7 km wide swathe along the fracture zone, even where the offset is negligible. Unusually high refraction velocities of 7.3-7.6 km s f l

were found for the lower crust from a seismic refraction line shot along the fracture zone, in strong contrast to the normal layer 3 velocity of 6.87 * 0.07 km s - l determined from the adjacent oceanic crust. The only other seismic refraction evidence for abnormal fracture zone crust comes from the Kurchatov fracture zone, a small-offset oblique transform north of the Azores (Fig. 1). Long-range mantle arrivals from shots above the fracture zone are late by as much as 0.45 s (Steinmetz, Whitmarsh & Moreira 1977), suggesting lowered crustal velocities in the fracture zone crust.

3 Models for the formation of anomalous fracture zone crust

Any model for the formation of crust in fracture zones must explain the characteristic features of its seismic velocity structure which set it apart from the crust generated else- where along the spreadiqg centre. The main characteristics are: (1) variable, but frequently quite thin crust; (2) anomalously low crustal velocities and the absence of a refractor with a velocity typical of seismic layer 3 ; (3) the occurrence of anomalous crust along both the active (transform) and the passive portions of the fracture zone, as well as beneath the nodal deep a t the ridge-transform intersections; (4) very thin crust confined to a relatively narrow zone (generally less than 10 km wide) centred on the fracture zone trough, with a more gradual crustal thinning extending up to 20 km away; (5) anomalous crust present regardless of the magnitude of the offset across the fracture zone.

There are two different types of mechanisms that have been suggested to explain the abnormal crust found in Atlantic fracture zones. The first seeks to explain the thin crust by tectonic thinning while the second postulates that the anomalous crust reflects a modification of normal accretionary processes near ridge-transform intersections.

Karson & Dick (1983) have suggested that crustal thinning may occur at ridge-transform intersections by back-tilting away from and down-tilting into the transform valley. Other possible mechanisms for mechanically thinning the crust are necking by extension across the fracture zone valley or tectonic thinning associated with the emplacement of large serpentinite bodies in fracture zones (Bonatti 1976, 1978). Although tectonic thinning may be an important process along parts of the fractures zone, it does not easily explain why very thin crust is found at ridge-transform intersections, and within the rift valley, why the crustal thinning extends several kilometres away from the fracture zone or why crustal velocities in fracture zones are usually very different from those of normal oceanic crust. The relatively common occurrence of anomalously thin crust at fracture zones along the slow-spreading Mid-Atlantic hdge suggests that this anomalous crust is a primary feature which reflects a modification of normal accretionary processes near ridge-transform intersections.

There are two main processes which may be responsible for the accretion of anomalously thin crust in fracture zones. A number of authors have emphasized the importance of the thermal effect of juxtaposing a cold, thick edge of lithosphere against the accreting plate boundary at ridge-transform intersections. This may result in less partial melting at depth (Stroup & Fox 1981; Fox & Gallo 1984), or in viscous head loss of rising asthenosphere (Sleep & Biehler 1970), and may impair the ability of melt already segregated from the mantle to reach the surface. The reduction in magma supply to the surface near the

Seismic structure o f fracture zones 795 transform will result in the accretion of a correspondingly thinner crust than o n a normal ridge segment. This thermal edge effect will depend on the age contrast across the transform. It will be greatest for large-offset fracture zones on slow-spreading ridges (e.g. Kane, Oceanographer and Vema fracture zones) and least for small-offset transforms.

Fox & Gallo (1984) have shown that the depth of the intersection deep is directly related to the age offset across the transform, suggesting that this thermal edge effect exerts a strong control on magmatic and tectonic processes near ridge-transform intersections. However, Wilson & Forsyth ( 1 982), in modelling the 3-D temperature structure of a ridge-transform- ridge system have demonstrated that even for a relatively large-offset transform like the Kane or Oceanographer (10 Ma age offset), conductive cooling across the transform extends only a distance of about 10 km from the fault. While this is consistent with the width of the narrow zone of extremely thin crust found along these fracture zones, lateral heat conduc- tion across the transform cannot explain the more gradual crustal thinning which extends several tens of kilometres on either side of the fracture zone. This mechanism alone also does not appear to provide a satisfactory explanation for the occurrence of anomalously thin crust at small-offset transforms since the thermal edge effect should be relatively small in such cases.

A second process which may influence crustal formation near transforms is the nature of the magma ‘plumbing’ system feeding individual spreading ridge segments bounded by fracture zones. Results from a number of recent investigations suggest that adjacent ridge segments spread independently of one another regardless of the age-offset across the inter- vening fracture zone (Schouten & White 1980; Schouten & Klitgord 1982; Le Douaran, Needham & Francheteau 1982) and that a significant fraction of the upwelling material required to form new oceanic crust flows along the ridge axis fed by a single, sub-crustal magma injection zone centrally located along the ridge segment (Francheteau & Ballard 1983; Parmentier & Forsyth 1984; White 1984). Sinha & Louden (1983), explained the gradual crustal thinning observed to extend several tens of kilometres from the Oceano- grapher fracture zone by a progressive reduction in the crustal production rate along an individual spreading ridge segment towards the transform. In Iceland, magma movement within the crust has been detected up to 70 km laterally from a single intrusive event (Sigurdsson & Sparks 1978), so it may be that the typical spacing between fracture zones of 40-80 km in the Atlantic reflects the distance over which magma from a single subcrustal supply can feed a spreading centre segment. Thus the magma supply may be restricted near fracture zones simply because they are at the distant ends of the intrusive systems feeding independently spreading ridge segments. This explains why anomalously thin crust is observed approaching both small- and large-offset fracture zones, because the lateral reduc- tion in magma supply occurs independently of the narrow zone of conductive cooling caused by the cold lithosphere edge at the transform itself.

The relatively low crustal velocities found in fracture zones can be attributed, in large part, to the intense fracturing and brecciation of crustal rocks caused by faulting within the active transform (Fox er al. 1976; Detrick et al. 1982). In addition to the major zone of faulting which cuts through the lithosphere at the principle transform deformation zone (eg. ARCYANA 1975), there are faults surrounding the ridge-transform intersection deep which also break up the crust in the fracture zone valley (Bowen 1984). Low seismic velocities in the upper crust may also, in part, be due to the accumulation of locally thick sections of loosely consolidated basaltic rubble from the steep fracture zone walls. Recent submersible studies in the Kane (Karson & Dick 1983) and Oceanographer (OTTER 1984a, b) transforms indicate that mass wasting is an important process in fracture zones and that significant thicknesses of basaltic rubble can accumulate in the transform

796 R. S. White et al.

valley. The relatively thin, highly fractured crust within the transform will also provide a particularly favourable environment for seawater to penetrate deeper than usual, thus serpentinizing the ultramafic components of the lower crust and upper mantle.

Thus the lower crustal velocities observed above the seismic Moho in fracture zones, which are quite unlike the typical layer 3 velocities found in normal oceanic crust, may represent either a highly fractured mafic crust or partially serpentinized ultramafics under- lying an extremely thin basaltic carapace. The large apparent variations in depth of the Moho observed along parts of both the Kane and Oceanographer fracture zones may reflect either changes in the thickness of the mafic crust due to small variations in the magma supply, or differences in the depth and extent of in situ serpentinization of the ultramafics lying beneath a very thin mafic crust.

The occurrence of highly variable and anomalous crust within fracture zones is all the more marked when it is compared with the remarkably uniform crust found over the strips of crust between adjacent fracture zones. The random variability has been shown to be very small within individual detailed survey areas carefully positioned away from fracture zone effects (Purdy 1983; White & Purdy 1983). On the broader scale represented by the collected velocity-depth determinations from widely spaced seismic profiles in Fig. 2, there is also considerably less variability over normal crust than that found within fracture zones. Such heterogeneity as is present can largely be explained by a few, systematic causes. The most important amongst these are age-dependent changes as the crust moves away from the spreading centre (Fowler 1976, 1978; Bunch & Kennett 1980), and as the cracks and fissures of the upper crust become progressively filled with hydrothermal precipitates over a period of 10-40 Ma (Houtz & Ewing 1976). Smaller alterations in velocities are caused by weathering of the exposed basement and by horizontal velocity anisotropy in the upper crust (Stephen 1981; White & Whitmarsh 1984), and in the mantle (Raitt et al. 1969).

4 Conclusions The seismic refraction data summarized in this paper show that the crust within both large and small Atlantic fracture zones falls well outside the range of velocity structures typically observed for oceanic crust in the North Atlantic. The crust varies in thickness along fracture zones, but is frequently very thin and is typically characterized by low crustal velocities and by the absence of a normal layer 3 refractor. We interpret this anomalous crust as a thin, intensely fractured and hydrothermally altered basaltic and gabbroic section overlying ultramafics that in places may be extensively serpentinized. The unusually thin crust present in fracture zones can be explained by an alteration in normal accretionary processes near transform faults caused, in part, by the thermal effect of juxtaposing a cold plate edge at the ridge-transform intersection and, in part, because transforms mark the ends of separately spreading ridge segments where the magma supply may be restricted simply because of its lateral distance from the centre of the intrusive system. Locally, tectonic processes may also be important in producing the very thin crust found along parts of many large Atlantic fracture zones. The variability in structure along individual fracture zones reflects small temporal changes in the magma supply, the degree of serpentinization and the local tectonic processes, which all contribute to the generation of abnormal crust.

Acknowledgments We thank Marion Jeffery and Ann Rayner for typing the manuscript. This work was supported by grants from the Natural Environment Research Council (UK) and the National Science Foundation (USA). Department of Earth Sciences, Cambridge: contribution number ES442.

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References ARCYANA, 1975. Transform faults and rift valley from Bathyscaph and Diving Saucer, Science, 190,

Bonatti, E., 1976. Serpentinite protrusion in the oceanic crust, Earth planet. Sci. Lett., 32, 107-1 13. Bonatti, E., 1978. Vertical tectonism in oceanic fracture zones, Earth planet. Sci. Lett., 37, 369-379. Bowen, A. N., 1984. A near-bottom seismic reflection profiler, PhD dissertation, University of Cambridge. Bunch, A. W. H. & Kennett, B. L. N., 1980. The crustal structure of the Reykjanes Ridge at 59" 30'N,

Cerveng, V. & Ravindra, R., 1971. Theory ofSeismic Head Waves, University of' Toronto Press, Canada. Cochran, J . R., 1973. Gravity and magnetic investigations of Guiana basin, western equatorial Atlantic.

Collette, €3. J.. Schouten, J . A., Rutten, K. & Slootweg, A. P., 1974. Structure of the mid Atlantic Ridge

Cormier, M., Detrick, R. S. & Purdy, G. M., 1984. Anomalously thin crust in oceanic fracture zones,

Detrick, R. S., Cormier, M. H., Prince, R. A. & Forsyth, D. W., 1982. Seismic constraints on the crustal

Detrick, R. S. & Purdy, G. M., 1980. The crustal structure of the Kane fracture zone from seismic

Fowler, C. M. R., 1976. Crustal structure of the Mid-Atlantic Ridge crest at 37" N, Geophys. J. R. astr.

Fowler, C. M. R., 1978. The Mid-Atlantic Ridge: structure at 45" N , Geophys. J. R. astr. Soc., 54,

Fox, P. J., 1978. The effect of transform faults on the character of the oceanic crust (abstract), Abstr. Progr. geol. SOC. Am., 7,403.

Fox, P. J. & Gallo, D. G., 1984. A tectonic model for ridge-transform-ridge plate boundaries: implications for the structure of oceanic lithosphere, Tertonophys., 104, 205-242.

Fox, P. J . , Schreiber, E., Rowlett, H. & McCamy, K., 1976. The geology of the Oceanographer I'racture Zone: a model for fracture zones, J . geophys. Res., 81,4117-4128.

Francheteau, J. & Ballard, Ti. D., 1983. The East Pacific Rise near 21" N, 13" N and 20" S: inferences for along strike variability of axial processes of the Mid-Ocean Ridge, Earth planet. Sci. Lett.,

Francheteau, J., Choukroune, P., Hekinian, R., Le Pichon, Y. & Needham, H. D., 1976. Oceanic fracture

Houtz, R. E. & Ewing, J. I . , 1976. Upper crustal structure as a function of plate age, J. geophys. Res.,

Karson, J. A. & Dick, H. J. B . , 1983. Tectonics of ridge-transform intersections at the Kane fracture zone, Mar. geophys. Res., 6 , 5 1-98.

Le Douaran, S. , Needham, H. D. & Francheteau, J., 1982. Pattern of opening rates along the axis of the mid-Atlantic Ridge,Nature, 300, 254-257.

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 . ash. Soc., 68,725-750.

108- 11 7.

Geophys. J. R. astr. Soc., 6 1, 14 1 ~ 166.

Bull. geol. Soc. Am., 84,3249-3268.

province between 12" Nand 18" N , Mar. geophys. Res., 2, 143-179.

new seismic constraints from the Kane fracture zone, J. geophys. Res., in press.

structure within the Vema fracture zone, J. geophys. Res., 87, 10 599- 10 6 12.

refraction studies, J. geophys. Res., 85, 3759-3777.

SOC., 47,459-491.

167-183.

64,93-116.

zones do not provide deep sections in the crust, Can. J. Earth Sci., 13, 1223-1235.

81,2490-2498.

Ludwig, W. J . & Rabinowitz, P. D., 1980. Structure of Vema Fracture Zone,Mar. Geol., 35,99-110. Mutter, J. C., Detrick, R. S. & NAT Study Group, 1984. Multichannel seismic evidence for anomalously

thin crust of the Blake Spur Fracture Zone,J. geophys. Res., in press. OTTER Scientific Team, 1984a. The geology of the Oceanographer transform: the transform domain,

Mar. geophys. Res., in press. OTTER Scientific Team, 1984b. The geology of the oceanographer transform: the ridge-transform

intersection,Mar. geophys. Res., 6, 109-141. Parmentier, E. M. & I'orsyth, D. W., 1984. Three-dimensional flow beneath a slowspreading ridge axis:

a dynamic origin for the deepening of the median valley towards fracture zones, J. geophys. Res., in press.

h r d y , G. M., 1983. The seismic structure of 140my old crust in the western central Atlantic Ocean, Geophys. J. R. astr. Soc., 12,115-138.

b i t t , R. W., 1963. The crustal rocks, in The Sea, vol. 3, pp. 85-102, ed. Hill, M. N., Wiley, New York. Raitt, R. W., Shor, G. G., [:rancis, T. J . G. & Morris, G. B . , 1969. Anisotropy of the Pacific upper

mantle, J . geophys. Res., 74,3095.

798 R. S. CVFzrte et al. RoDb, J. M. & Kane, M. I?., 1975. Structure of the Vema Fracture Zone from gravity and magnetic

Schouten, I t . & Klitgord. K. I)., 1982. The memory of the accreting plate boundary and the continuity

Schouten, G. & Klitgord, K. [I., 1983. Overlapping spreading centres on East Pacific Rise. Nature, 303,

Schouten, H . & White, R. S., 1980. Zero-offset fracture zones, Geology, 8, 175-179. Sibuet, J. C. & Veyrat-Peinet, B., 1980. Gravimetric model o f the Atlantic equatorial fracture zones,

Sigurdsson, ti. ti Sparks, S. K. J.. 1978. Lateral magma flow within rifted Icclnndic crust. Nature, 274,

Sinha, M. C. & Louden, K. E., 1983. The Oceanographer 1:racture Zone - I. Crustal structure from seismic refraction studies, Geophys. J. R. astr. Soc., 75, 713-736.

Sleep, N. H. & Biehler, S. , 1970. Topography and tectonics at the intersection of fracture zones with central rifts. J. geophys. Res., 75, 2748-2752.

Spudich, 1'. & Orcutt, J . , 1980. A new look at the seismic velocity structure of the oceanic crust, Rev. Geophys. Space Phys., 18,627-645.

Steinmetz, L., Whitmarsh, R. B. & Moreira, J . S., 1977. Upper mantle structure beneath the mid-Atlantic Ridge north of the Azores based on observations of compressional waves, Geophys. J. R. astr.

Stephen, R. A, 1981. Seismic anisotropy observed in upper oceanic crust, Geophys. Rex Lett., 8,

Stroup, J . 13. & [:ox, P. J . , 1981. Ckologic investigations in thc Cayman Trough: evidence for thin crust along the midCayman Rise, J. Geol., 89, 395-420.

Sutton, G. H.. Maynard, G. L. & Hussong, D. M., 1971. Widespread occurrence of a high velocity basal layer in the Pacific crust found with repetitive sources and sonobuoya, in The Srructure and Physical Properties o f ' the Earth's Crust, ed. Heacock. J . G., Geophys. Monogr. A m . geophys. Un., 14, I93--209. Washington DC.

T w i g , W., Slootweg, A. 1'. & Collette, H. J. , 1979. Topography and a magnetic analysis of an area south-east of the Azores (36" N, 23" W).Mar. geophys. R e x , 4, 91-104.

Twigt, W., Verhoef, J., Rohr, K., Mulder. Th. F. A. & Collette, B. J . , 1983. Topography, magnetics and gravity over the Kane fracture zone in the Cretaceous Magnetic Quiet Zone (African Plate), Proc. K . ned. Akad. Wet. B, 86, 181-210.

White, R. S., 1984. Atlantic oceanic crust: seismic structure o f a slow spreading ridge, in Ophiolitesand Oceanic Lithosphere, Spec. Publ. geol. Soc. Lond., 13, 101-1 11 , eds Gass, 1. G., Lippard, S. J . & Shelton, A. W., Blackwell Scientific Publications, Oxford.

White, R. S. & Matthews, D. H., 1980. Variations in oceanic upper crustal structure in a small area of the northeastern Atlantic, Geophys. J. R. astr. Soc., 61, 401-436.

White, R. S. & Purdy, G. M., 1983. Crustal velocity structure on the flanks of the mid-Atlantic Ridge at 24", Geuphys. J. R. astr. Soc., 75,361 -386.

White, R. S. & Whitmarsh, R. B., 1984. An investigation of seismic anisotropy due to cracks in the upper oceanic crust at 45" N, mid-Atlantic Ridge, Geophys. J. R. a s k Soc., 79,439-467.

Wilson, B. & Forsyth, D., 1982. Three dimensional temperature structure of a ridge-transform-ridge system (abstract), EOS, 63, 1100.

intensity prot'iles.X Eeophys. Rex , 80, 444 1-4445.

of fracture zoncs, Earth platier. Sci. Lett., 59, 255-266.

549-550.

J. Keophys. Res., 85,943- 954.

126- 130.

Soc., 50,353-380.

86 5 -- 86 8.