Crustal structure of Ascension Island from wide-angle seismic data: implications for the formation of near-ridge volcanic islands

Earth and Planetary Science Letters 190 (2001) 41^56 www.elsevier.com/locate/epsl

Crustal structure of Ascension Island from wide-angle seismic data: implications for the formation of near-ridge volcanic islands F. Klingelho«fer a; *, T.A. Minshull b , D.K. Blackman c , P. Harben d , V. Childers e b

a IFREMER, Centre de Brest, P.O. Box 70, F-29280 Plouzane¨, France Southampton Oceanography Centre, School of Ocean and Earth Science, European Way, Southampton SO14 3ZH, UK c IGPP, Scripps Institution of Oceanography, 9500 Gilman Dr. Dept 0225, La Jolla, CA 92093-0225, USA d Lawrence Livermore National Laboratory, 817 South G. St., Livermore, CA 94550, USA e Naval Research Laboratory, 4555 Overlook Ave SW, Washington, DC 20375-5350, USA

Received 1 November 2000; received in revised form 6 April 2001; accepted 3 May 2001

Abstract The study of the internal structure of volcanic islands is important for understanding how such islands form and how the lithosphere deforms beneath them. Studies to date have focused on very large volcanic edifices (e.g., Hawaiian Islands, Marquesas), but less attention has been paid to smaller islands, which are more common. Ascension Island, a 4-km high volcanic edifice with a basal diameter of 60 km, is located in the equatorial Atlantic (8³S), 90 km west of the Mid-Atlantic Ridge on 7 Ma oceanic lithosphere. We present results of a wide-angle seismic profile crossing the island revealing a crustal thickness of 12^13 km, an overthickened layer 3 (7 km thick) and little evidence of lithospheric flexure. Together these results suggest Ascension Island may be older than previously assumed and may have begun forming at an on-axis position around 6^7 Ma. This hypothesis is further supported by the presence of a young 1.4-km high edifice directly adjacent to the Mid-Atlantic Ridge with a volume about 1/7 that of Ascension Island, possibly representing the earliest stages of seamount formation. Excess magmatism appears to be related to the tectonic setting at the ridge^fracture zone intersection. ß 2001 Elsevier Science B.V. All rights reserved. Keywords: Atlantic Ocean; volcanic centers; Ascension Island; refraction methods; £exure; crust

1. Introduction/objectives The structure of the Earth's crust at oceanic

* Corresponding author. Present address: Bullard Laboratories, Department of Earth Sciences, University of Cambridge, Madingley Road, Cambridge CB3 0EZ, UK; E-mail: [email protected]

islands and seamounts is of interest for the information it provides both about the formation processes of large volcanoes and about the rheology of the underlying oceanic lithosphere, which deforms in response to an applied surface load. The ensuing lithospheric £exure has been the subject of intense study for several decades [1^6]. These £exural studies have typically been carried out on very large edi¢ces such as the Hawaiian chain,

0012-821X / 01 / $ ^ see front matter ß 2001 Elsevier Science B.V. All rights reserved. PII: S 0 0 1 2 - 8 2 1 X ( 0 1 ) 0 0 3 6 2 - 4

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Marquesas, Galapagos, Tuamotu and Canary archipelagos, where total crustal thicknesses exceed 15 km. Previous studies of the crustal structure of large volcanic islands have yielded contrasting results. While massive underplating of 5^7 km has been reported beneath Oahu [2], this has been disputed by a reinterpretation of the same OBS data [7]. A very thick underplated body has been inferred beneath the Marquesas islands [8], and it has been suggested its buoyancy may isostatically support the swell in this region [9]. On the other hand, recent work from Reunion and Tenerife indicate little to no underplating (at most a 1^2 km thick layer) [6,10,11]. Seismic experiments suggest that oceanic islands typically contain a central high-velocity core [6,11,12], which probably consists mostly of intrusive rocks, and an outer shell of lower-velocity material consisting primarily of lava £ows, extrusive deposits and the products of mass wasting. Seismic data can distinguish between high-velocity intrusive regions and pyroclastic material (e.g., Jasper Seamount [13]). Good constraints on crustal structure have been obtained for a few oceanic intraplate volcanoes which are entirely submarine, such as Jasper Seamount in the northeast Paci¢c [13] and Great Meteor Seamount in the central Atlantic [12]. These data suggest that small volcanic edi¢ces contain a much larger proportion of low-velocity materials than their larger counterparts. To understand the systematics of seamount formation we need to studies of a range of sizes of objects. For oceanic islands, even when high quality wide-angle seismic data are available, they seldom allow an accurate determination of the deep-crustal structure down to the upper mantle due to the limited depth of penetration with respect to largescale structures [2,8]. Furthermore, large islands represent extreme end-member structures on the exponential seamount size distribution scale and commonly present complex morphologies with multiple cones, radial ridges and £ank rift zones. Since increasing size appears to be associated with increasing complexity [14] the study of smaller, simpler structures can help improve our understanding of the development of the entire family of seamounts and volcanic islands. Lastly, be-

cause the crustal structure, lithology and densities can likely be better constrained by seismic experiments, this information can help shed light on the style(s) of isostatic compensation in such geodynamic settings. With these objectives in mind a geophysical survey of Ascension Island, a modest sized volcanic edi¢ce, was conducted in May 1999. 2. Ascension Island Ascension Island is a 4-km high volcanic edi¢ce with a basal diameter of 60 km, located in the equatorial Atlantic (8³S), 90 km west of the Mid-Atlantic Ridge between the Ascension F.Z. and the Bode Verde F.Z. (Fig. 1). It lies between magnetic anomalies 3P and 4 on 7 Ma oceanic lithosphere [15]. Radiometric dates of lava £ows exposed at the surface range from 0.61 to 1.5 Ma [16,17], while the presence of fresh £ows suggests the volcano has likely been active in the last few hundred years [18]. While these ages represent minimum ages of activity, it is thought the edi¢ce was emplaced fairly rapidly and that the age of the lithosphere at the time of loading was 6 Ma ( þ 1 Ma) [19]. Thus according to thermal cooling^lithospheric thickness relations [20,21] a mechanical plate thickness of 12 km should be expected [19]. The results of three-dimensional (3-D) gravity modelling indicate a 2-km depression of the Moho beneath the island which, if due to £exure, would indicate a much lower e¡ective elastic thickness (Te ) of 3 þ 1 km [19]. The modelled 2-km subsidence due to £exure is consistent with hyaloclastite (shallow water volcanic £ows) observed in the borehole Ascension #1, a geothermal exploration well, at depths down to 1790 m below sea level [22] (see Fig. 2). A further result of this modelling was that the strong gravity gradient observed across the island cannot be explained by a simple £exural model and must be due to density variations inside the volcanic edi¢ce itself [19]. Ascension Island is a promising study area because of good geological data available in the form of surface mapping and from a deep geothermal exploration well drilled to a depth of

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Fig. 1. Bathymetry of the study area around Ascension Island. Stars mark OBH, triangles mark portable seismometers deployed on land, circles mark sonobuoy deployment locations, and black line marks the pro¢le presented in this paper. Open circle marks the location of the ASC#1 deep borehole. Inset shows regional setting near the Mid-Atlantic Ridge (MAR). Asc marks Ascension Island and A^D mark large seamounts [15]. Black arrows show absolute plate motion vectors since 30 Ma in the `¢xed' hotspot reference frame for the African and South American plates [53] and white arrows show the spreading direction [54].

3126 m [22]. Among the rocks recovered were subaerial basalts, hyaloclastites and shallow- to deeper-water basaltic rocks with hydrothermal deposits. Seismic velocities range between 3.9 and 4.7 km/s at a depth from 2011 to 2218 m below the surface, the limited interval that was logged. The volcanic rocks found exposed on the island form an alkaline suite with a compositional range of basalt^hawaiite^mugearite^benmorite^ trachyte^rhyolite, with the more basaltic £ows found in the SW of the island and the rhyolitic and trachytic domes and lava £ows found in the eastern part of the island. The existence of rhyolitic rocks suggests that felsic magma chambers have formed which could provide heat for a convective hydrothermal system [16]. It has been suggested that Ascension Island has a plume/hotspot origin [23,15]. The anomalously

shallow bathymetry of the nearby segment of the MAR [15,24] as well as trace element signatures and isotopic ratios of basalts from the nearby ridge axis, which are intermediate between typical MORB and enriched basalts from hotspot islands [25^27], have been cited as supporting this model. A plume location near Circe seamount (9³S, 11.7³W) has been proposed, with asthenospheric £ow channeled towards the ridge axis [28]. Subsequently it was suggested that a plume is situated beneath two large on-axis seamounts at 9³50'S (Fig. 1 labelled C and D), with channeling of upwelling asthenosphere towards Ascension Island [15]. Alternatively, it has been suggested that the observed geophysical and geochemical anomalies could result from melting of small mantle heterogeneities that have risen into the melting region at or close to the ridge axis [29,24].

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3. Seismic data A wide-angle seismic study of the island was conducted during a four-day geophysical survey aboard the RRS James Clark Ross. Speci¢c objectives included mapping the shape of the Moho beneath the island, detecting and quantifying any magmatic underplating, and determining the origin of the strong local gravity gradient across the island. During the experiment a 6186-cubic inch (101-l)-tuned airgun array was ¢red at 1-min intervals for about 2 days along lines extending up to 45 km from the coast of the island. A total of 526 shots were ¢red on the 86-km long pro¢le presented in this paper. The airgun source was towed at a depth of 20 m and a speed of 5 knots leading to a mean trace spacing of 130 m. These shots were received by two ocean bottom hydrophones (OBH), ¢ve sonobuoys o¡shore and six landstations on the island (Fig. 1). The two OBHs were moored in mid-water in the sound channel to allow calibration of the source and for use in a separate investigation of acoustic T-phase propagation and coupling. The data quality is generally good (Figs. 3a, 4a and 5) and arrivals recorded by the OBHs and most of the landstations could be traced to ranges of 50^55 km, including rays undershooting the island. During the experiment, P2 and P3 phases (originating from turning rays in the upper and lower crust) were registered by all landstations and all OBH instruments. On recordings from most landstations and from the two OBHs, Pm P re£ections from the Moho could be identi¢ed. Weak Pn arrivals from the upper mantle are found in the data of one OBH and three landstations. All data were acquired using a sampling rate of 0.004 s and subsequently bandpass-¢ltered with gain ramps at 3^5 Hz and 24^36 Hz. For some of the landstations a narrower bandpass ¢lter (3^5 Hz, 18^24 Hz or 3^5 Hz, 15^18 Hz) showed better results for large o¡sets. The OBH positions were determined by ¢tting the water primary arrivals using the least-squares method. The sonobuoys generally do not show arrivals beyond the 15^25 km range, and thus were used only to constrain upper crustal velocities. Source-receiver o¡sets

Fig. 2. Lithologic data and sonic log from deep drill hole Ascension 1 (after Nielson and Stiger [22]). Dashed line shows velocities derived from the seismic model of this study.

were determined by ray tracing through the water column and subsequently the shots were divided into several gathers with the sonobuoy position assumed ¢xed within each gather [30]. The data were modelled using the inversion and ray tracing algorithm of Zelt and Smith [31]. Estimated picking uncertainties were 50 ms for the upper crustal arrivals recorded by the OBH, 70^80 ms for lower crustal arrivals, and 90 ms for re£ected phases. Sonobuoy data were generally noisier and had additional uncertainties caused by drift; travel times picked from sonobuoy records were generally assigned a picking uncertainty of 70 ms, but an uncertainty of 200 ms was assigned to SB12, which was o¡ the line. The uncertainty estimated for the landstations was between 80 ms (SBC, APS) and 100 ms (ACH) depending on the data quality and the signal to noise ratio. Modelling was performed using a layer-strip-

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Fig. 3. (A) Bandpass ¢ltered (30^5 Hz, 24^36 Hz) data from landstation SBC. The data are displayed with gain proportional to source-receiver o¡set and are reduced at a velocity of 8 km/s. Inset shows the same data at a higher gain to show ¢rst arrivals more clearly. (B) Synthetic seismograms calculated from the model for the same station using asymptotic ray theory [31]. The synthetic seismograms are spaced at a 500-m interval and the same o¡set-dependent gain has been applied as in (A). The source wavelet consists of a 29-point low-pass ¢ltered Ricker wavelet. P-wave quality factors were chosen to be 400 for the basaltic basement, 700 for the lower crust and 1000 for the upper mantle. Inset shows a blow-up of the un¢ltered data using a higher gain. Thin lines are travel times predicted by the ¢nal velocity model.

ping approach, proceeding from the top of the structure towards the bottom. As far as possible, lateral variations were modelled by varying layer thickness; lateral velocity variations were introduced only when required by the data. We used a 2-D iterative damped least-squares inversion of travel times [31]. Upper layers, where not directly constrained by arrivals from within this layer, were adjusted to improve the ¢t of lower layers. Velocity gradients and the phase identi¢cation in the velocity model were further constrained by synthetic seismogram modelling using asymptotic ray theory [32].

4. The velocity model The ¢nal velocity model consists of two water layers, three crustal layers and a layer of typical upper mantle velocities (Fig. 6), and ¢ts the picked travel times well (Table 1). Each layer is de¢ned by depth and velocity nodes. The velocity in the water column, estimated from historical measurements of temperature and salinity, was de¢ned by two layers, with a velocity of 1.54 km/s at the surface, a velocity inversion at 900 m depth and then a gradual increase to 1.51 km/s. The sea£oor was parameterized with depth

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Fig. 4. (A) Data from OBH 2 for the pro¢le with same gain, ¢lter and scaling applied as in Fig. 3. (B) the corresponding synthetic seismograms. Pn arrivals are not detected in the data from this OBH.

nodes at a spacing of 1 km derived from bathymetric data, in order to account for its high rugosity. A layer with velocities typical for layer 2A of oceanic crust is de¢ned with depth nodes at intervals of 2.5 km and velocity nodes at a spacing not less than 5 km. The velocities at the top of this layer vary from 2.5 km/s at the sea£oor to 1.9 km/s on the island, where this layer consists of poorly consolidated subaerial volcanics. The velocity at the base of Layer 2A is constant over the model at 3 km/s. The spacing of depth nodes increases to 5 km for the Layer 2/3 boundary, which in our parameterization coincides with the 6.25 km/s velocity contour. This boundary has been included because the data require a change in velocity gradient. A small increase in velocity at this boundary enhances the model ¢t signi¢cantly.

At the Moho discontinuity the depth node spacing is 10 km where necessary. Velocity variations at the base of Layer 3 are constrained by the moveout of Moho re£ections. The model includes a velocity jump from 6.95^7.4 km/s at the base of Layer 3 to 8.0 km/s in the uppermost mantle. Rays penetrate the upper mantle only between 5 and 40 km and between 70 and 80 km model distance, and the upper mantle velocity is assumed to be constant throughout the remainder of the model. At the southwest end of the pro¢le, between 0 and 20 km o¡set, the model exhibits normal oceanic crust with a thickness of 6 km. The Moho geometry is poorly constrained here since there are no PmP arrivals and Pn arrivals are unreversed, but a horizontal Moho is consistent with

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Fig. 5. Data from OBH 4 for o¡sets between 29 and 53 km with same gain, ¢lter and scaling applied as in Fig. 3. Thin lines are travel times predicted by the ¢nal velocity model.

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a mantle velocity of 8.0 km/s. The average thickness of oceanic Layer 2 here is 2.6 km and that of Layer 3 is 3.2 km. The velocity^depth relationship here (Fig. 7, left) falls within the range of 0^ 127 Ma Atlantic oceanic crust [33], except for the very low upper crustal velocities. On the basis of sonobuoy data [34], the Mid-Atlantic Ridge close to Ascension has a similar upper crustal velocity structure including two layers, though velocities appear to be somewhat higher (Fig. 7, left). The crustal thickness at the ridge is V8 km and thus 2 km thicker than the crust near Ascension Island, mainly due to an increased Layer 3 thickness of 5 km.

Fig. 6. (A) Final velocity model including the model boundaries used during inversion (solid lines) and isovelocity contours every 0.25 km/s. Positions of OBHs (stars), landstations (triangles) and sonobuoys (circles) are indicated. Bold dashed line indicates the Moho depth found from 3-D gravity modelling [19]. (B) Resolution for the depth nodes at the base of Layer 2A, Layer 2 and the Moho and for the velocity nodes at the top of Layer 2A and the base of Layer 3.

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Table 1 Number of picks, picking error, rms-mis¢t and M2 parameter for some phases identi¢ed on the pro¢le

Complete pro¢le

Phase

Number of picks

water upper crust lower crust PmP Pn

1747 667 756 726 108 55

The upper crustal layers thicken considerably along the slopes of the island to a thickness of about 5.6 km directly beneath the island. The ¢rst upper crustal layer with velocities between 2.3 and 3 km/s is thickest below both island £anks, with a thickness of about 1.4 km but thins to 0.7 km directly below the island surface (Fig. 6). For comparison, the sonic log deployed in the borehole Ascension #1 shows an increase from 3.5 to 5 km/s in the depth range between 2000 and 2225 m [22] which agrees well with the velocity model (Fig. 2). At the location of landstation APP in the SW part of the island this layer thins to nearly zero thickness and much higher velocities are found at shallow depth in the records of this station than of any other landstation. The depth of the Layer 2/3 boundary is relatively constant, varying only by 0.9 km along the pro¢le, generally shallower to the NE of the pro¢le than the SW and rising by about 0.5 km below the island. Moho depth increases below the island by about 3 km, from 9.4 km to 12.5 km depth. The depth of the Moho below the island is broadly consistent with the Moho depth derived from gravity and £exural modelling [19] (Fig. 6). Velocities at the Moho increase from 6.95 km/s in normal oceanic crust to 7.4 km/s below the NE part of the island. Velocities of 7.2^7.8 km/s are commonly found beneath volcanic islands and seamounts at Moho depth (Fig. 7). The velocity contours within Layer 3 diverge beneath the island, rather than being £exed consistently downward. Velocity gradients and the phase identi¢cation were constrained by synthetic seismogram modelling [32] (Figs. 3b and 4b). The high-velocity gradients of the upper crust generate high-am-

Picking error (s)

Rms mis¢t (s)

M2

0.05 0.05 0.08 0.10 0.10

0.091 0.065 0.090 0.074 0.117 0.197

1.054 0.735 1.211 0.673 1.382 3.968

plitude arrivals, as observed particularly in the data from OBH2 (Fig. 4). In order to estimate the velocity and depth uncertainty of the ¢nal velocity model we performed a perturbation analysis. To determine the robustness of our conclusions concerning the shape of the Layer 2/3 boundary, we constructed 10 di¡erent velocity models with di¡erent shapes for this boundary, varying between a £exed model with the boundary following the Moho and a model with slightly greater shallowing of the boundary beneath the island than our best-¢tting ¢nal model, and calculated the M2 error for each model (Fig. 8). The M2 is de¢ned as the root-mean-square (rms) travel time mis¢t between observed and calculated arrivals normalized to the picking uncertainty. An F-test applied to the M2 values of the resulting models [31] suggests that smooth models with a signi¢cantly di¡erent Layer 2/3 boundary shape can be rejected at the 95% con¢dence level (Fig. 8). Additional information about the quality of the velocity model can be gained from the resolution parameter (Fig. 6) [31]. Resolution is a measure of the number of rays passing through a region of the model constrained by a particular velocity or depth node and is therefore dependent on the node spacing. Values greater than 0.5 are considered well resolved [31]. The velocities at the base of Layer 2A, at the base of Layer 2 and at the top of Layer 3 are single parameters in the model and therefore their resolution is s 0.98. The resolution for the velocity at the top of Layer 2A is greater than 0.5 from 25 to 70 km along the model, except for two nodes. The resolution for depth nodes of the base of Layer 2A is greater than 0.5

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Fig. 7. (Left) Velocity^depth pro¢le of the undisturbed oceanic crust at 10 km model distance and of the Mid-Atlantic Ridge close to Ascension Island [34]. Grey-shaded area marks velocity range for 0^127 Ma Atlantic oceanic crust [33]. (Right) Velocity^ depth pro¢le below the island at 55 model km. Also shown are velocity^depth relationships for Great Meteor Seamount at OBH 24 slightly NW of the center of the seamount [12], Jasper Seamount below the NW ridge [13], Hawaii [2], Marquesas [8] and Tenerife at the center of the edi¢ces [6]. In each case, w is the width of the base of the edi¢ce and h is the approximate height of the edi¢ce.

Fig. 8. The variation of the M2 parameter for 10 di¡erent models between a £exed model and a model with a slightly greater shallowing of the Layer 2/3 boundary beneath the island than the ¢nal model. Grey-shaded area represents the statistical 95% con¢dence limit of the F-test.

between 5 and 85 km. The model would be better resolved with wider depth node spacing, but the close spacing is required to allow this boundary to follow the overlying seabed/land surface. Two nodes at 47 and 49 km show a very low resolution, because their node spacing is smaller than for the other nodes of this layer to ¢t the high-velocity gradients around landstation APP, a feature which is supported by the gravity modelling. The depth of the base of Layer 2 shows good resolution between 30 and 70 km, where a downward de£ection would be expected for a purely £exed model. Constraints on the Moho depth are good between 10 and 30 km and 45 to 65 km. The number of rays re£ected or crossing the Moho between 30 and 45 km is comparatively small and here the shape of the Moho is constrained mainly by gravity modelling. The resolution of the velocity at the base of Layer 3 is good between 20 and 60 km model distance. Outside this region, the resolution values for all layers

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Fig. 9. (Upper panel) Ray coverage of the model, with every third ray from point-to-point ray tracing plotted. (Lower panel) Observed travel time picks and calculated travel times (line) for all receivers along the model. Positions of the receivers are indicated by black inverted triangles.

are below 0.5, because of their poor ray coverage (Fig. 9). The number of picks, picking error, the values for the M2 = 1.5 parameter and the rms mis¢t for the most important phases of the model are listed in Table 1. 5. Gravity modelling Since seismic velocities and densities for oceanic crust are well-correlated, gravity modelling provides an important additional constraint on the seismic model. Crustal P-wave velocities from the seismic model were converted to densities using the relationship of Carlson and Raskin [35] with upper mantle densities set to a constant 3.30 g/cm3 . Although the seismic data are 2-D, it is important to incorporate 3-D e¡ects of the sea£oor bathymetry into the gravity model. To

estimate the magnitude of these e¡ects, we divided the model into eight layers, each of constant density. We assumed that this density structure is constant perpendicular to the pro¢le and extrapolated these layers laterally, pinching out against the 3-D bathymetry on the £anks of the island. The resulting model incorporates only the 3-D e¡ects of the known geometry of the sea£oor; 3-D e¡ects of underlying structures, of unknown geometry, are not accounted for. We then calculated the gravity contribution of each layer using the method of [36] using the Taylor series coe¤cients up to an order of ¢ve. We also calculated the 2-D gravity anomaly along the pro¢le using the same constant density layers as for the 3-D model (Fig. 10). The di¡erence between the anomaly of the 2-D model anomaly and that of the 3-D model represents an estimate of the in£uence of the o¡-line bathymetry on the gravity anomaly. We subtracted this di¡erence from the anomaly

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Fig. 10. Comparison between measured gravity data (black dots) and the gravity anomaly computed for two di¡erent density models derived from the velocity model. The dashed line represents a 2-D model and the solid line a 2-D model corrected for 3-D e¡ects of o¡-line bathymetry. Dots represent the gravity anomaly [19]. Error bars represent the estimated þ 3 mGal uncertainty of the marine dataset [19]. A linear trend has been removed from the data. Dotted line shows the bathymetry.

due to a more complex 2-D model which included lateral density changes (Fig. 10). The corrected anomaly can be compared with the observed gravity anomaly [19] (Fig. 10). On land, we removed in£uence of the island by using a Bouguer anomaly with a Bouguer correction density of 2300 kg/m3 . In the data, there is a large-scale trend of increasing gravity to the west due to the in£uence of thermal cooling. Since this e¡ect occurs on a much larger scale than the seismic model, and its precise magnitude is di¤cult to compute due to the thermal e¡ect of the Ascension Fracture Zone, the model anomalies are compared with the data after subtraction of a straight-line ¢t to the residuals. The results of this modelling show that the mis¢t between the anomaly of the 2-D model and the observed data can be explained by 3-D topographic e¡ects, and that the gravity gradient across the island is well explained by the shallow density variations inferred from the velocity variations required to ¢t the seismic data. 6. Discussion The seismic data presented here indicate the following variations in crustal structure. The uppermost crustal layer, with velocities between 2.2 and 4.0 km/s, thickens considerably on the submarine £anks of the island, where it probably consists primarily of slope breccia and lava £ows.

This layer is thicker on the SW £ank of the island, which is older and has been more strongly eroded. The southern part of the island displays the highest near-surface velocities and densities; perhaps here the lower-velocity material has been removed by erosion, as suggested by the presence of a broad, submerged platform further west. Denser ma¢c rocks crop out here, in contrast to the latestage rhyolitic £ows found on the NE side of the island [16]. Directly beneath the island, velocities increase to between 7.1 and 7.4 km/s at the base of the crust, compared with 6.9 km/s in the neighboring oceanic crust. These velocities correspond neither to mantle material, which has a velocity of 8.0 km/s adjacent to the island, nor to normal oceanic gabbro which has velocities of 6.8^ 7.2 km/s. Underplated gabbros derived from melting of asthenosphere with unusually high potential temperature can have increased velocities [37]. Alternatively, the high-velocity material might correspond to gabbro sills alternating with dunites [38] or to olivine-rich gabbros resulting from fractionation of basaltic melt in crustal magma chambers [39]. The last interpretation is plausible since the basalts sampled at the surface of the island have undergone signi¢cant degrees of fractionation [40]. Given the morphologic variation between endmembers in the seamount population [14] one might expect a similar systematic variation in crustal structure. For Ascension Island one would thus predict a structure intermediate between that

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of large emergent islands (e.g., Hawaiian Islands, Reunion, Canary Islands) and that of smaller submerged seamounts (e.g., Jasper Seamount). A comparison of the velocity pro¢les to that of six other volcanic edi¢ces supports this assessment (Fig. 7). The velocity gradients are highest in the shallowest layers, similar to the smaller (3.8 km high, 25 km basal diameter) Jasper Seamount [41,13]. Most of the larger structures exhibit a 2^4-km region of velocities s 7.0 km/s at depths of 10^14 km beneath the top of the edi¢ce (Fig. 7), in good agreement with the velocity structure from this study. While such velocities have commonly been cited as evidence for magmatic underplating [2,8], the Ascension model lacks a velocity step between Layer 3 velocities and the possible underplate, and no re£ection from the top of the underplate is seen in the seismic data. Although some of the larger structures have thick underplates (Hawaii, Marquesas), the 7-km high Tenerife edi¢ce shows a similar velocity structure to Ascension, with a velocity jump from 7.2 to 8.0 km/s at the base of the crust, and no clear signs of underplating [6]. The estimated e¡ective elastic thickness for Ascension Island of 3 km is much less than the predicted mechanical thickness of 12 km based on the thermal age [19]. Though the model on which the 3-km value is based assumes that the depression of the Moho beneath the island is due to £exure, and our new seismic work has shown that £exure is not the main cause of the depression, the diameter of the shallow moat around the island is still consistent with this value [19]. Several e¡ects were proposed to explain the low Te [19]: (1) the extreme curvature of the £exed plate causes bending stresses exceeding the yield strength and thus invalidating the simple £exed elastic plate theory [42]; (2) the thermal age of the lithosphere was reset at the time of loading by hotspot heating [43]; (3) the lithosphere beneath the island was locally reheated by the emplacement of the volcanic edi¢ce itself [19]. To these we add and discuss the following hypothesis: (4) the construction of Ascension Island may have taken longer than previously assumed. If the constructional history were longer, (e.g., 5^6 Myr) then the foundation of the edi¢ce was lain on 1^2-

Ma lithosphere with a much lower elastic thickness, close to a state of Airy isostasy. The ¢rst and third e¡ects may contribute to the low Te ; the second e¡ect is unlikely to be signi¢cant because there is no evidence for a hotspot swell around Ascension Island [24]. Our results support hypothesis 4. The crustal structure as revealed by the velocity model (Fig. 6) does not show the typical features of an intraplate volcanic island, e.g., a prominent £exural bulge, with accompanying deep water- or sediment-loaded moats, and a £exed Layer 2/3 boundary (Fig. 11a). Rather, the Layer 2/3 boundary is sub-horizontal (Fig. 11b) and Layer 3 is overthickened, suggesting that it formed in this con¢guration near the ridge axis. An alternative interpretation would be that the high velocities at mid-crustal depths beneath the island represent the presence of an intrusive core, e.g., as interpreted at Great Meteor Seamount [12]. However, such a core would have to be within, rather than above, the pre-existing oceanic crust, which is perhaps more likely with a near-ridge origin. The observation of a thickened Layer 3 is consistent with global relationships between Layer 3 thickness and overall crustal thickness, based heavily on observations

Fig. 11. Cartoon illustrating the idealized structure of volcanic islands. (A) Intraplate volcanic island with strong £exure and regional isostasy, (B) Simpli¢ed structure of Ascension Island showing overthickened Layer 3 formed near-axis, close to Airy isostatic equilibrium.

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from oceanic plateaus formed on or near the ridge axis [44]. In this hypothesis the depression of the Moho beneath the island does not represent 2^3km £exure, but rather is a feature inherited from the time of crustal accretion. Several other volcanic edi¢ces lining the south side of the Ascension Fracture zone (Fig. 1, inset) may have formed in a similar manner. Approximately 600 and 300 km WSW of Ascension Island are two unnamed seamounts (B, A) of nearly identical dimensions, which may possibly represent drowned ancestral Ascension Islands [15,45]. These rest on 36- and 20-Ma lithosphere and seamount A rises to only 262 m below sea level [15]. Their alignment is inconsistent with the plate motion vector (Fig. 1) and suggests edi¢ce

53

construction may be related to processes at the fracture zone-spreading center junction. Furthermore, a mere 15 km west of the ridge axis, directly adjacent to the axial valley is a young edi¢ce with a 60-km N^S length, 24-km E^W width, and 1.4-km height, representing a volume about 15% that of Ascension Island (Fig. 12). This seamount is an active constructional volcanic edi¢ce as demonstrated by the Brunhes magnetic anomaly traversing it and by a caldera structure observed in swath bathymetric data [46] (Fig. 12). The associated free-air gravity anomaly (80 mGal) is about half that observed for Ascension (150 mGal) [19]. The current height above the sea£oor base of 2.8 km water depth, indicates that such a substantial edi¢ce can grow in 1 Myr

Fig. 12. Shaded sea£oor relief map from swath bathymetric survey of [46], the dataset of [19] and other unpublished data showing the dimensions of Ascension Island and a 1.4-km high on-axis seamount and their positions with respect to the Mid-Atlantic Ridge and adjacent fracture zone. The seamount possibly represents an early stage of growth of a volcanic edi¢ce which may reach the size of Ascension Island. Where no shipboard data are available, interpolated bathymetric data from [55] are contoured.

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time to cover half the distance to the surface. As smaller seamounts have steeper slopes than larger volcanic islands [14,41], continued growth at current extrusion rates would result in the young seamount emerging within the next 2^3 Myr and reaching the present volume of Ascension Island in 6 Myr. Such longevity of volcanism is not unusual for major volcanic features such as the Hawaiian chain, where volcanoes show typically 4 Myr of activity [47,48] and the Canary Islands, where volcanism on some islands has continued for well over 10 Myr (Gran Canaria 15 Ma to present, [49]; Tenerife 12 Ma to present, [50]). The presence of long-lived volcanism in a chain of seamounts which does not align with absolute plate motions is di¤cult to explain by conventional plume theories, but is common in regions such as the South Paci¢c and may arise from the interaction of partially molten asthenosphere with tectonic stresses generated near mid-ocean ridges and fracture zones [51]. Our results, together with the data for Jasper Seamount, which has a 7 Myr long eruptive history [13] suggest that even for relatively small volcanic edi¢ces the ages of surface rocks may not be representative of the age of the main shield-building phase, and that hotspot tracks based on dating such rocks may have considerable uncertainties. 7. Conclusions Our wide-angle seismic survey constrains the crustal structure beneath Ascension Island and for the ¢rst time was able to undershoot an entire volcanic island, revealing a maximum crustal thickness of 13 km. Only minor amounts of underplating may be present. Crustal thickening appears to be divided equally between an overthickened Layer 3 (7 km rather than the usual 4 km) and the 3^4 km height of the volcanic island. The absence of a £exed Layer 2/3 boundary is atypical for regionally compensated intraplate volcanic islands. Gravity modelling based on the seismic velocity structure con¢rms the minor degree of lithospheric £exure beneath the island. One interpretation for the observed structure is that a substantial portion of the crustal thickening

occurred on axis. Therefore the volcanic edi¢ce capped by Ascension Island may be older than previously thought. Dating of lava £ows at the bottom of the 3-km Ascension #1 borehole may serve to test the hypothesis of the age of Ascension Island. Acknowledgements We are indebted to the o¤cers, crew and technicians aboard the RRS James Clark Ross for their professional work during the cruise. The experiment was funded by the US Department of Energy (through Lawrence Livermore National Laboratory), the US O¤ce of Naval Research, the UK Natural Environment Research Council, the United Nations (through the Comprehensive Test Ban Treaty O¤ce in Vienna), and IFREMER's Centre de Brest. Funding for F.K. was provided by the European Community within the scope of a Marie-Curie fellowship and T.A.M. was supported by a Royal Society University Research Fellowship. The GMT [52] software package was used in the preparation of this paper. We thank Neil Mitchell for access to unpublished bathymetric data and Dennis Nielson for providing the sonic log in digital form. We are grateful to I. Grevemeyer, P. Charvis and A. Bonneville for suggestions which helped improve the original manuscript.[AC]

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