Gravity modeling of the ocean-continent transition along the South Atlantic margins

Transcription

1 JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 114,, doi: /2008jb006014, 2009 Gravity modeling of the ocean-continent transition along the South Atlantic margins Diana Dragoi-Stavar 1,2 and Stuart Hall 1 Received 14 August 2008; revised 5 March 2009; accepted 4 June 2009; published 9 September [1] Gravity, magnetic, and seismic data have been used to examine changes in crustal structure of conjugate portions of the South Atlantic volcanic margins south of the Walvis Ridge-Rio Grande Rise hot spot tracks. We have constructed 18 seismically constrained crustal-scale gravity models of the ocean-continent transition: 12 across the African margin and 6 across the South American margin. We attribute changes in character of the gravity anomalies to variations in crustal thickness, sedimentation, and magmatic underplating. To investigate variations in the lower crust and upper mantle, we have removed the relatively well-constrained contributions to the gravity field from shallow sources. The resulting residual anomalies delineate distinct crustal domains associated with crustal thinning and regions of underplating that correspond well with those independently identified from limited seismic refraction data. Near the hot spot tracks, we find underplated crust to be symmetrically distributed with a horizontal extent of 240 km for each margin. Further south, however, the underplated region is asymmetric with significantly more beneath the African than the corresponding South American margin. Residual gravity anomalies indicate that the width of underplated crust decreases from 240 km to 150 km at 33 S on the African side and to 100 km at 38 S on the South American side. The wider underplated region on the African side suggests an asymmetric rifting mechanism in which greater melts accumulate beneath an upper plate margin. Our results are consistent with recently proposed asymmetric rifting models that involve strain softening of a coupled, cold frictional upper lithosphere at low rifting velocities. Such models also predict greater symmetry for areas with more heat such as those associated with the hot spot. Citation: Dragoi-Stavar, D., and S. Hall (2009), Gravity modeling of the ocean-continent transition along the South Atlantic margins, J. Geophys. Res., 114,, doi: /2008jb Introduction [2] Continental rifting and breakup involve the complex interactions of tectonic, magmatic, geodynamic, and sedimentary processes [Coffin et al., 2006]. These interactions result in a wide variety of margin styles, ranging from narrow to extended, and from weakly to strongly volcanic. Passive continental margins are classified as volcanic or nonvolcanic on the basis of the amount of volcanism that occurs during rifting [White et al., 1987; Mutter et al., 1988; White and McKenzie, 1989; Holbrook and Kelemen, 1993]. Globally, volcanic margins represent the majority of passive continental margins [Skogseid, 2001] with as much as 70% of the Atlantic margins considered volcanic [Erzinger et al., 1997]. Volcanic margins are characterized by seaward dipping reflector sequences (SDRs) formed by the eruption and subsequent subsidence of volcanic flows during rifting, and by thick lower crustal layers with anomalously high 1 Department of Geosciences, University of Houston, Houston, Texas, USA. 2 Chevron, Houston, Texas, USA. Copyright 2009 by the American Geophysical Union /09/2008JB006014$09.00 seismic P wave velocities (>7.2 km/s) and high densities (3100 kg/m 3 )[Planke et al., 1991; Eldholm et al., 1995; Kelemen and Holbrook, 1995]. High-velocity lower crustal layers mapped beneath the continental slope vary in thickness from 5 to 15 km [Hinz, 1981; Mutter et al., 1982, 1984; Larsen and Jakobsdottir, 1988; White and McKenzie, 1989; Holbrook and Kelemen, 1993; Kimbell et al., 2004]. Both lower crustal layers and SDR complexes are spatially associated with the ocean-continent transition, and are located between stretched continental crust and normal thickness oceanic crust [Kelemen and Holbrook, 1995; Boutilier and Keen, 1999; Korenaga et al., 2000; Trumbull et al., 2002b; Menzies et al., 2002]. Where imaged on seismic profiles, the crust between SDRs and the lower crustal layer is extremely thin [Watts, 2001] with SDRs in some cases approaching the region of the 7.2 km/s material [Talwani and Abreu, 2000]. The lower crustal layer is commonly referred to as underplating, which may be composed of mantle-derived material below original continental crust [White et al., 1987], or igneous material added within or below the crust [Cox, 1980]. Talwani et al. [1995], however, argue that the term underplate is not appropriate considering that, in general, there is no older crust being 1of15

2 Figure 1. Map of the South Atlantic showing major magmatic features (dark gray), regions of seaward dipping reflectors (SDRs; light gray), seafloor spreading isochrons, and location of study areas shown in Figure 2. RGR, Rio Grande Rise; WR, Walvis Ridge; CFB, continental flood basalts. Modified from Bauer et al. [2000]. underplated. Talwani et al. [1995], consider the lower crustal layer to be the lower part of an initial oceanic crust, whereas the SDRs represent the upper part. [3] The nature and location of the boundary between continental and oceanic crust is of fundamental importance to understanding the rifting process. At volcanic margins the ocean-continent transition occurs over a relatively short distance from 50 to 80 km [White et al., 1987; Mutter et al., 1988; White and McKenzie, 1989] to 150 km wide [Gladczenko et al., 1997, 1998; Dahl-Jensen et al., 1997] compared with that at nonvolcanic margins. The more abrupt transition at volcanic margins has been attributed to weakening of the lithosphere by large quantities of hot intruded igneous rocks [White and McKenzie, 1989]. SDRs and large thicknesses of high-velocity, lower crustal layers at volcanic margins have also been used to support mantle plume involvement in the rifting process [White and McKenzie, 1989; Eldholm and Grue, 1994]. Several volcanic margins are associated with hot spots, and their influence on the amount of magmatism, and the margin s subsequent subsidence and thermal history have been examined by numerous authors [e.g., Coffin and Eldholm, 1993; Eldholm and Grue, 1994; Gladczenko et al., 1997, 1998; Holbrook and Kelemen, 1993; White and McKenzie, 1989]. [4] The South Atlantic passive continental margins formed as a result of the opening of the South Atlantic and the separation of the African and South American continents during Early Cretaceous [Rabinowitz and LaBrecque, 1979; Austin and Uchupi, 1982; Nurnberg and Muller, 1991]. Significant extension took place prior to opening with rifting propagating northward [Peate et al., 1990], beginning in Late Jurassic (Kimmeridgian Portlandian) in the south along the Namibian margin and ending in the Hauterivian in the north [Rabinowitz and LaBrecque, 1979; Sibuet et al., 1984; Maslanyj et al., 1992; Light et al., 1993]. Seafloor spreading is also thought to have begun in the southernmost South Atlantic at chron M9 (130 Ma) [Austin and Uchupi, 1982] and propagated northward. Continental breakup and initial seafloor spreading appear to have been accompanied by large-scale volcanism that produced the Parana-Etendeka Continental Flood Basalts at 132 ± 1 Ma [Renne et al., 1996]. The Parana-Etendeka subaerial volcanism has been linked to the impact of the Tristan plume on the base of the continental lithosphere [White and McKenzie, 1989]. The flood basalt province was subsequently split in two during opening of the South Atlantic, but this separation was accompanied by continued igneous activity offshore that is reflected in both the Rio Grande Rise and the Walvis Ridge tracks, and SDRs that appear limited to the margins south of the hot spot tracks [Bauer et al., 2000; Talwani and Abreu, 2000]. [5] The flood basalts, the voluminous extrusive constructions on the conjugate margins of West Africa and South America, and the existence of paired hot spots tracks on the seafloor all lend strong support to a South Atlantic flanked by plume-related volcanic margins. Here we examine the influence of the Tristan da Cunha hot spot on the development of the South Atlantic margins by using geophysical data to investigate changes in crustal structure with alongstrike distance from where the hot spot trace intersects the coast (Figure 1). The relative abundance of high-quality multichannel seismic reflection data available for the margins, especially the African margin, makes it possible to constrain upper crustal structure well. With these constraints we have attempted to obtain a better understanding of the deep structure of the margins, and the ocean-continent transition using crustal models obtained from an analysis of satellite-derived free air gravity anomalies. The width, thickness and nature of the ocean-continent transition constitute important parameters that may be used to better understand processes leading to continental breakup and 2of15

3 Figure 2. Study areas along (a) South American and (b) African margins showing locations of modeled profiles overlain on satellite free air gravity anomaly maps. Portions of profiles shown in black coincide with available seismic data. PB, Pelotas Basin; SB, Salado Basin; CB, Colorado Basin; WB, Walvis Basin; LB, Luderitz Basin; OB, Orange Basin. Blue box in Figure 2b encloses area of enhanced satellite derived gravity data provided by GETECH. (c) Reconstruction of the South Atlantic at chron C-34 (shown in red) modified after Moulin [2003] using pole of rotation 61, 66 N, 34, 37 W, 33, 55 [Campan, 1995]. Base map shows satellite gravity anomalies [Sandwell and Smith, 1997]. White outlines indicate SDR distribution at each margin after Gladczenko et al. [1997]. Onshore areas (shown with red) indicate Cretaceous age continental flood basalts. Locations of the modeled profiles indicated with black lines. seafloor spreading. Because both margins were influenced by the same hot spot, an additional objective of the study was compare the ocean-continent transition on conjugate segments of the margins and use the crustal parameters to cast light upon rifting models responsible for the breakup. 2. Data Sources 2.1. Gravity Data [6] Free air anomalies derived from satellite altimetry data [Sandwell and Smith, 1997; also digital file, version 15.2, available at 2min] were the major source of gravity data (Figure 2). These data are available with a grid spacing of 2 arc minutes ( 4 km in the study area) and a resolution of 5 mgal over 20 km wavelengths. High-resolution satellite-derived free air data with a resolution of 4 mgal over km wavelengths [Fairhead et al., 2004, 2001] were provided by GETECH for a portion of the African margin (Figure 2b). Satellite data were augmented in places with shipboard marine free air gravity data from the NOAA/NGDC Marine Trackline Geophysics data set, version 4.1. Ship track gravity and associated bathymetry were selected from 58 and 46 surveys along the African and South American margins, respectively. Although the resolution of shipboard surveys (±1 mgal at 5 10 km wavelength) is generally better than satellite data, coverage is sparse over much of the South Atlantic Bathymetric Data [7] Seafloor depths range from 0 m (coastline) to >5000 m in basins on both margins. Bathymetry data were obtained 3of15

4 from (1) marine ship track data from NGDC, (2) global digital bathymetric data set (ETOPO2 version 8.2) gridded at 2-min intervals, and (3) Shuttle Radar Topography Mission (SRTM30) data. These were supplemented with bathymetry data digitized along published seismic reflection/refraction profiles. ETOPO2 provides predicted seafloor topography inferred from satellite gravity data that are combined with observed water depth measurements [Smith and Sandwell, 1997]. Although several authors have discussed the accuracy of the predicted seafloor topography [Smith and Sandwell, 1997, 2004; Sandwell and Smith, 2001; Smith, 1993; Dixon et al., 1983], it is generally considered reliable over intermediate wavelengths ( km) in areas where sediment cover is thin [Sandwell and Smith, 2001]. Sediment thicknesses in the study area are very variable from a few hundred meters to >5 km. Because the topography/gravity ratio varies from one region to another, estimation of topography from gravity requires accurate depth soundings within a region for calibration. The SRTM digital topographic data set has a grid resolution of 30 s (1 km) and includes land data. Detailed information regarding the SRTM data set may be found at topex.ucsd.edu/www_html/srtm30_plus.html Seismic Data [8] Seismic reflection and refraction studies have been carried out over each margin. Along the African margin, depth-converted multichannel reflection data from Gladczenko et al. [1998] and Brown et al. [1995] are available for the northern part of the study area (Figure 2b). Reflection profiles from Young [1992] cover the southern portion. Two major reflection and refraction transects made near 24 S image the crust from top to bottom [Bauer et al., 2000]. In addition, there are onshore refraction observations in the Damara Orogen and Namibian craton [Green, 1983] and deep seismic soundings in the area of Damara Orogen [Baier et al., 1983]. [9] Along the South American margin, seismic studies include multichannel reflection and wide-angle refraction profiles across the Colorado Basin near 40 S [Franke et al., 2002, 2006] and the Argentine continental margin near 43 S [Hinz et al., 1999; Franke et al., 2002; Neben et al., 2002]. Further north there are reflection profiles S [Abreu, 1998; Talwani and Abreu, 2000], S [Bassetto et al., 2000; Abreu, 1998; Talwani and Abreu, 2000], and near 38 S [Hinz et al., 1999] (Figure 2a). Refraction data are available for 34 S 35 S [Leyden et al., 1971], and other parts of the Colorado Basin [Ludwig et al., 1979] Sediment Thickness and Distribution [10] Sediment thickness maps include a regional total sediment thickness grid for the entire South Atlantic [Divins and Rabinowitz, 1991], and local isopach maps at different stratigraphic intervals on either side. For the African side, there are sediment thickness maps along the Namibian margin, [Stewart et al., 2000], and the Orange [Jungslager, 1999] and Walvis [Clemson et al., 1999] basins. For the South American side, there are sediment thickness maps for the Colorado and Salado basins [Franke et al., 2002, 2006; Urien, 2001], the Pelotas Basin [Milliman, 1978], and a prerift unconformity time structure map for the Colorado Basin [Bushnell et al., 2000]. Basin sediments have been subdivided in sedimentary sequences based upon seismic stratigraphic interpretations of Bagguley [1997] and Stewart et al. [2000] along the African margin and of Hinz et al. [1999] and Abreu [1998] for the South American margin Well Data [11] Except for the uppermost sediments, published well control is sparse. Along the African margin, data released for the Kudu wells in the Orange Basin have been included in the work of Bolli et al. [1978], Bagguley [1997], Stewart et al. [2000], Young [1992], and Gladczenko et al. [1998]. Ocean drilling programs have sampled parts of the sedimentary section on the outer portion of the African margin (DSDP Legs 40 and 75, and ODP Leg 175). Igneous basement has been sampled at only one site on the Walvis Ridge (Site 530 during DSDP Leg 75). On the South American margin, well data were obtained from Abreu [1998], Ludwig et al. [1968, 1979], Kowsmann et al. [1977], Bushnell et al. [2000], Talwani and Abreu [2000], and Jackson et al. [2000] Magnetic Data [12] Magnetic anomaly profile data were obtained from the Marine Trackline Geophysical Database (NGDC). Ship track data over the African margin were selected from 77 surveys between latitudes 18 S to35 Sand longitudes 0 E to20 E. Along the South American margin, ship track data were selected from 67 surveys, between latitudes 25 S and 45 S and longitudes 70 W and 30 W. [13] Seafloor age data were extracted from the digital age grid of the ocean floor isochrons [Muller et al., 1997] and integrated with ship track magnetic data to help correlate and interpret magnetic anomalies over the margins. 3. Methods [14] We have used 2-D forward gravity models along profiles approximately perpendicular to each margin to investigate changes in their crustal and upper mantle structure. Gravity anomalies south of Walvis Ridge-Rio Grande tracks form linear bands of highs and lows generally parallel to the margins (Figure 2) and therefore a 2-D approach is appropriate. Models have been constrained with available seismic data, and geologic information [Gerrard and Smith, 1982; Maslanyj et al., 1992; Light et al., 1992; Light et al., 1993; Stewart et al., 2000; Clemson et al., 1997; Urien, 2001; Mohriak et al., 2002; Bushnell et al., 2000]. Models were constructed along 12 profiles across the African margin and 6 across the South American margin (Figure 2). Wherever possible, profiles were selected to exactly coincide with available seismic reflection and/or refraction transects. Gravity values were extracted from the satellitederived free air gravity grid using a 5-km spacing along the profiles. Where possible, bathymetric data digitized along individual seismic lines were used. Elsewhere, we used SRTM 30 bathymetric grids and bathymetry from ship tracks that intersected the modeled profiles Model Geometry [15] Models were constructed using a multilayer geometry that included water, sediments, SDRs, upper crust, lower 4of15

5 crust, and magmatic underplating, and an underlying halfspace (upper mantle) that extended to a depth of 50 km. On the basis of seismic and geologic evidence, the sedimentary and underplating portions were subdivided in additional layers. Three sedimentary layers were used to represent postrift basin fill, in accordance with sequence stratigraphic interpretations and geologic data [Bagguley, 1997; Stewart et al., 2000]. Underplating was divided into an upper and lower layer based upon seismic velocity data along seismic transects BGR 1 and BGR 2 that showed the presence of a transitional layer situated between the lowermost underplating and the lower crustal layer [Bauer et al., 2000]. Wherever possible, interfaces between individual layers were obtained from depth-converted seismic reflection and refraction data. [16] Seismic reflection data provide reasonably good constraints on the shallow structure but imaging of the deep structure is limited to a small area of each margin. Moho depths are available for the Walvis Basin near 22 S [Gladczenko et al., 1998] and between 22 S to 25 S [Bauer et al., 2000], and the Colorado Basin near 40 S [Franke et al., 2006] and 43 S[Hinz et al., 1999]. Estimated uncertainties in P wave velocities are about ± km/s in the upper and middle crust (<20 km) and ±0.2 km/s, in the lower crust and uppermost mantle, implying uncertainties in the Moho depth determinations of 3 5 km [Bauer et al., 2000] Layer Densities [17] Each layer was assigned a density, which for consistency was kept constant for all models. Reflection and/or refraction velocities obtained along both margins [Bauer et al., 2000; Gladczenko et al., 1998; Franke et al., 2006; Hinz et al., 1999; Abreu, 1998; Ludwig et al., 1979] were converted to densities using published velocity-density relations. Densities for sedimentary rocks were calculated using Ludwig et al. [1970] Postrift sediments were subdivided into three separate layers: Lower Cretaceous (2450 kg/m 3 ), Upper Cretaceous (2110 kg/m 3 ), and Tertiary (2000 kg/m 3 ). Along the South American margin, the Upper Cretaceous sediments were assigned a density of 2300 kg/m 3. These densities compare well with published values for sediments in cores from Site 362 (DSDP Leg 40) [Bolli et al., 1978] and those obtained for basic lavas and sediments from Kudu 9-A2 and 9-A3 wells [Bagguley, 1997]. Densities for igneous and metamorphic rocks were obtained from the velocity-density relations of Christensen and Mooney [1995]. The upper and lower crustal layers were given densities of 2700 and 2900 kg/m 3, respectively. [18] Published SDR densities vary considerably reflecting P wave velocity variations observed at different volcanic margins. Values range from 2550 kg/m 3 to 2680 kg/m 3 on the Norwegian margin [Mjelde et al., 2005], to 2750 kg/m 3 at the southeastern Brazilian margin [Bassetto et al., 2000] and to 2850 kg/m 3 for U.S. east coast margin [Holbrook et al., 1994]. The SDR density used in this study (2600 kg/m 3 ) is an average value derived from seismic velocities along seismic refraction transects BGR 1 and 2 for Namibian margin [Bauer et al., 2000]. An SDR density of 2600 kg/m 3 has been used in gravity studies of the Namibian [Gladczenko et al., 1998], and Norwegian [Tsikalas et al., 2005] margins, and the Hatton-Rockall Basin [Edwards, 2002]. [19] Underplating and upper mantle densities are consistent with those for other continental margins [Watts, 2001]. All gravity models included a magmatic underplated lower crustal layer with density 3100 kg/m 3 and an overlying transitional layer with density 3000 kg/m 3 that was identified on seismic data by Bauer et al. [2000]. A density of 3300 kg/m 3 was used for the mantle Modeling Procedure [20] We used a commercial gravity-modeling program [GM-SYS, 2004], based upon the method described by Talwani et al. [1959], to calculate the gravity response of each model. Models were constructed by dividing each cross section into several polygonal layers with variable geometries and by assigning densities as described above. Edge effects were avoided by extending the models several hundred kilometers, east and west of the gravity profile. The seaward end of each profile extends over what is considered normal oceanic crust. Because free air gravity anomalies generally approach zero over normal oceanic crust, seaward ends of the profiles were used to balance and calibrate each model. For regions of normal oceanic or continental crust, the ratio between upper and lower crustal layer thicknesses was kept constant. For stretched continental crust associated with crustal thinning and/or initiation of rifting, the upper crustal layer was thinned preferentially; that is, the ratio of upper and lower crustal thickness was not held constant. In our models we made no attempt to fit anomalies with wavelength <10 km because the satellite gravity data do not contain meaningful signals below km, and because the focus of our study was regional variations in the crust and upper mantle rather than localized features such as volcanic intrusions and seamounts. 4. Results [21] Gravity responses for each model are generally consistent with observed anomalies, and only minor adjustments to initial geometries derived from reflection and refraction data were necessary to satisfactorily fit the data within the accuracy limits (namely, 2 4 mgals). Poorer fits were obtained over sedimentary basin portions of profiles 4, 6, 8, 10 and 11, and over the steeper gradients associated with the ocean-continent transition. Where model responses do not match the observed data well, the differences may be due to sediment density variations, a factor not included in our uniform density models. Sediment thickness maps indicate that southern basins (e.g., Orange Basin) along the African margin are deeper than those further north (i.e., Luderitz and Walvis basins) with sediment thicknesses reaching 8 9 km. Along the South American margin, sediment thicknesses of 7 8 km in the Pelotas Basin are greater than those in Argentine basins further south. If sediment density increases with depth, then the density used for the deeper sediments will be too low, and the calculated gravity too negative. This is the case for both the Orange Basin, where there is 6 mgal difference between observed and calculated anomalies and the Pelotas Basin where there is a 3 4 mgal difference. [22] Moho depths vary considerably along the modeled cross sections, from roughly 12 km at the seaward end to 35 km at the coastline. Crustal thickness was measured 5of15

6 Figure 3. Two-dimensional gravity models for three conjugate pairs of profiles across the African and South American margins: (a) northern A-2, (b) central D-8, and (c) southern F-12 parts of the study area. Profile locations shown in Figure 2. Portions of the models constrained by seismic data are indicated by areas within the dashed white outlines. Each model shows a multilayer crustal cross section, a comparison of observed and calculated free air gravity anomalies, and the residual gravity anomaly, which is attributed to deep crustal and upper mantle sources. Residual anomalies were obtained by subtracting the gravity effects produced by water depth variations, sedimentary layers, SDRs, and crystalline basement from observed free air data (see section 5). Horizontal bars (shown with red) over the residual anomaly graph delineate boundaries of linear segment 3 (see text). M0, M4, and G show locations of distinctive magnetic anomalies identified by Rabinowitz and LaBrecque [1979]. 6of15

7 from the basement surface to the observed Moho (where seismic data are available) or to the gravity derived Moho. In those parts of the profiles where SDRs are present, the top of the extrusives was used as the basement surface. Profiles with good seismic control (e.g., profile F on Figure 3), display Moho depths and crustal thicknesses at the seaward and landward ends that are typical for normal oceanic and continental crust [Gladczenko et al., 1998; Franke et al., 2002]. We have defined the region between these crustal types (i.e., continental crust with thickness 30 km and oceanic crust with thickness km) as the ocean-continent transition. [23] The seaward termination of the ocean-continent transition is here defined by the seaward extent of the SDRs wedge [Mutter et al., 1982; Eldholm et al., 1995]. We have used the SDR boundary mapped by Gladczenko et al. [1998], Bauer et al. [2000], and Schreckenberger et al. [2001] along the African margin, and that mapped by Hinz et al. [1999], Bassetto et al. [2000], and Schreckenberger et al. [2001] along the South American margin. The landward termination of the ocean-continent transition is delineated by the transition from thick igneous to continental crust with a normal velocity structure, which is located near the landward limit of SDRs. Trumbull et al. [2002a] identify the continent-ocean boundary along the African margin as that coinciding with the landward edge of SDRs in the upper crust and by prominent gravity and magnetic anomalies. The thickness of underplating (density 3100 kg/m 3 ) estimated from the gravity models is quite variable (i.e., 0 7 km) but corresponds well with thicknesses constrained by seismic refraction data. [24] In all 18 models that we constructed the shallow structure is well constrained by extensive seismic data. Adequate constraints on the deep structure, however, are only available for 6 models (namely, profiles 1, 2, 3, 4, E, and F). We have based much of our gravity modeling for other profiles on the geometries and overall dimensions derived from these more well-constrained profiles. Here we focus on the results from 6 models that constitute three conjugate pairs: Profiles 2 and A, 8 and D, and F and 12. These profile pairs have been selected to provide representative examples from the northern, central and southern portions of the study area (Figure 2). A detailed description of all 18 gravity models is given by Dragoi-Stavar [2007]. [25] Profile 2 passes through the middle portion of the Walvis Basin (Figure 2b) following seismic reflectionrefraction transect MCS 4 [Gladczenko et al., 1998]. Moho depths are well constrained by seismic data along its entire length from 34 km near the coast to 14 km at the seaward end. At the SW end the crust is somewhat thicker (7.5 8 km) than normal oceanic crust possibly reflecting proximity to the Walvis Ridge. At the NE end, the crust thickens rapidly by 8 km to 34 km over a horizontal distance of only 35 km (Figure 3). [26] The SDRs form a single 165 km wide layer that varies in thickness from 5.6 km in the center to 1.5 km (at km 175) (Figure 3). The underplating layer extends over 200 km, has an average thickness of 5 km but reaches a maximum of 6.7 km at km 162 (Figure 3). [27] Previous gravity modeling along MCS 4 [Gladczenko et al., 1998] included an intrusion with density of 2950 kg/m 3 to explain the 40 mgal high located above the inner shelf area. The limited spatial extent of the high on a detailed free air map indicates a 3-D rather than 2-D source. Consequently we have modeled the high using a vertical cylinder with a density contrast of +100 kg/m 3 (relative to the lower crust density), a depth to top of 9 km, depth to bottom of 30 km, and radius of 14 km. The calculated anomaly (45 mgals) compares well with observed high. [28] Total thickness of the postrift sedimentary succession varies along the profile but reaches a maximum of 4 km in the center of the basin. Profile 2 also crosses an old rift unconformity metasedimentary basin beneath the shelf edge that reaches depths of approximately 13 km. [29] An RMS misfit of 1.48 mgals was obtained without modification to the crustal structure derived from seismic data. Portions of the profile where the residual displays a somewhat higher value (3 mgals) are those over the old rift basin, and in the vicinity of the most rapid changes in Moho geometry (Figure 3). [30] Profile A extends NW SE across the northern edge of the Pelotas Basin extending offshore for a distance of 1000 km (Figure 2a). The model was constructed using 2 seismic reflection profiles: line PD [Abreu, 1998] (km 38 to km 355), and line 48 [Bassetto et al., 2000] (km 387 to km 652). No seismic reflection or refraction data were available to constrain deep crustal structure along the profile. Moho depth decreases from 34 km at the coastline (constrained from the crustal model of Mooney et al. [1998]), to 25 km at km 53. Between km 53 and the shelf break (km 159) Moho depths remain roughly constant (24 km) before decreasing to 19 km at 405 km and finally to 15 km at the seaward end of the line. On the basis of our gravity model we estimate that the oceancontinent transition is approximately 425 km wide in this area. [31] SDRs imaged between km 177 and km 327 (line PD) [Abreu, 1998] and between km 403 and km 440 (line 48) [Bassetto et al., 2000] have been included in our gravity model as a single 266 km wide layer with a thickness of 3 4 km. A distal wedge of SDRs in the oceanic crust further seaward discussed by Bassetto et al. [2000], has not been included. The model also includes a 246 km wide (km 159 to km 405) lower crustal body (magmatic underplating) with a thickness that averages 4 5 km (Figure 3). [32] On the basis of the sequence stratigraphic interpretation developed by Hinz et al. [1999] for areas further south, sedimentary layers were divided into 3 units: seabed to Upper Eocene, Upper Eocene to Turonian and Turonian to basement and their geometry was obtained from depth converted seismic reflection profiles. [33] Near km 536, profile A crosses the Porto Allegre Lineament which is a major crustal lineament characterized by normal faults with large offsets and a concentration of volcanic highs, seamounts, and guyots. No attempt was made to fit the short wavelength anomalies associated with these features. [34] An overall good fit (RMS misfit 1.8 mgals) was obtained between observed and calculated gravity. A somewhat higher residual (4 mgals) was obtained at km 307, where a gravity high marks the seaward limit of SDRs. [35] Profile 8 is 547 km long and crosses the northern part of the Orange Basin roughly 60 km north of the Kudu High (Figure 2b). Upper portions of the gravity model 7of15

8 between km 277 and km 550 were constrained by seismic reflection data along profile 1 of Bagguley [1997] and Stewart et al. [2000]. Crustal thickness variations indicate a decrease from normal continental crust (30 km) at km 523 beneath the inner shelf area, to 20 km in the vicinity of the shelf break (km 463), to 14 km at km 388 in the center of the basin, then to 9 km at km 265 at the seaward limit of the SDRs and finally to normal oceanic crust (6.5 km) between km 278 and the seaward end of the profile. There is a rapid change in crustal thickness at the NE end where Moho depths increase by 7 km over a horizontal distance of 70 km (Figure 3). We estimate the ocean-continent transition is about 260 km wide along this profile. [36] The model includes a 208 km wide SDR zone (km 279 to km 487) that reaches a maximum thickness 4 km at km 427 and then decreases both seaward and landward to 2 3 km. The magmatic underplating layer used in the model has the same width as the SDR layer and is positioned at the same horizontal position as the SDRs. The variation in thickness of the magmatic underplating layer suggests an average thickness of 4 to 5 km. [37] The postrift sedimentary succession has a maximum thickness of 4.5 km and is composed of several seismically defined stratigraphic units. The shallowest unit records the Tertiary evolution of the sedimentary prograding wedge, whereas the deepest unit (basement to Aptian) is equivalent to megasequences MS20 and MS30 of Bagguley [1997]. The seismic facies of megasequence MS30 have been interpreted by Gladczenko et al. [1997] as SDRs. [38] A RMS misfit of 1.55 mgals was obtained between the observed and calculated free air gravity anomalies. A higher misfit (5 mgals) was obtained near the coastline where a gravity low separating the shelf break gravity high from an onshore gravity high was only partially reproduced by the gravity model. [39] Profile D is 760 km long and crosses through the Salado Basin at 38 S extending offshore (Figure 2a). The model is constructed so as to coincide with multichannel seismic reflection profile BGR [Hinz et al., 1999] that extends from km 110 to km 652. Seismic interval velocities [Hinz et al., 1999] for identified horizons were used to depth convert the seismic time section, which is approximately 12 s (TWT) deep. Moho was imaged on parts of the seismic section but was not observed under the continental slope (i.e., beneath the SDRs). From the coastline the depth to Moho decreases from 30 km at km 131, to 27 km on the outer shelf area, to 19 km at km 350 beneath the slope and the SDRs wedge, and decreases to 12 km at the seaward end on the modeled line. We estimate that the ocean-continent transition is approximately 320 km wide along this profile. [40] In our gravity model the SDRs wedge is located beneath the slope extending over a distance of 75 km (km 320 to km 395) with a variable thickness that averages 3 km. The magmatic underplating layer is substantially wider (134 km) than the SDR layer extending further seaward with a thickness of 3 4 km. [41] The Salado Basin is a relatively symmetrical narrow basin with up to 5 km of continental and overlying marine sediments. Elsewhere geometries for the sedimentary layers were constructed from Line BGR using regional unconformities interpreted by Hinz et al. [1999]. A good overall fit (RMS misfit = 1.56 mgals) between the observed and calculated free air gravity anomalies was obtained. [42] Profile 12 crosses the southern portion of the Orange Basin toward the Agulhas-Columbine Arch near the coastline and is approximately 740 km long (Figure 2b). The upper 11 km (i.e., sedimentary cover and top of basement) are constrained between km 525 and km 711 by seismic reflection data from profile of Brown et al. [1995]. Crustal thickness derived from gravity modeling decreases from 33 km at the coastline (km 740) to 24 km at the hinge zone (km 640), 16 km at km 537 and approximately 10 km at km 410. At the seaward end of the model Moho depth is 11.5 km and crustal thickness (6 km) corresponds to normal oceanic crust. We estimate that the ocean-continent transition zone is 225 km wide (km 430 to km 655). [43] The SDRs layer, which is constrained by seismic reflection data between km 525 and km 635, was extended laterally in the gravity model producing a 185 km wide zone. SDR thicknesses reach a maximum of 4 km (km 537) and then decrease both seaward and landward. The zone of magmatic underplating extends for about 198 km (km 440 to km 638) with a maximum thickness of 7 km at km 561. [44] Major postrift sequences defining the drift sedimentary succession of Tertiary, Upper Cretaceous and Lower Cretaceous ages [Brown et al., 1995; Muntingh and Brown, 1993] were digitized to produce an initial geometry for sedimentary layers. This geometry was adjusted slightly until a good fit was obtained. Eastward from km 700 toward the coastline, the gravity model includes a basement high in the area where Precambrian/Paleozoic basement was imaged [Brown et al., 1995] along the flank of the Agulhas- Columbine Arch. A good fit (RMS residual = 2.0 mgals) between calculated and observed gravity anomalies was obtained over the whole length of profile 12. [45] Profile F crosses the Rawson Basin on the southern Argentine margin, and extends 200 km further seaward. The upper 30 km is well constrained by refraction data between km 545 and km 855 (Figure 3). Initial model geometry was derived from the reflection and seismic refraction data of Hinz et al. [1999], Neben et al. [2002], and Franke et al. [2002]. Depths to Tertiary and Cretaceous sedimentary layers, SDRs, basement, Moho, underplating, and intracrustal variations were obtained from the velocity model of Franke et al. [2002]. Moho depths vary from 30 km beneath continental crust at km 548 to 23 km under the slope and the area covered by SDRs, to 14 km beneath oceanic crust at the seaward end (Figure 3). [46] SDRs were modeled as a 80 km wide, 3 km thick layer. Neben et al. [2002] and Schreckenberger et al. [2001] describe three separate sequences of SDRs at 43 S, bounded by strong unconformities, constituting a km wide and several thousand meters thick volcanic wedge seaward of the prerift and synrift basins beneath the outer shelf. The SDR geometry along profile F was obtained by digitizing the two superposed SDRs wedges from a refraction line at 43 S[Franke et al., 2002]. These were then represented as a single density layer with uniform density 2.6 g/cm 3.Inthe model an underplating layer with a width of 195 km and an average thickness of 3 4 km was used. [47] A good fit (RMS misfit = 1.37 mgals) between observed and calculated anomalies was obtained over most 8of15

9 Figure 4. Residual gravity anomaly profiles over the South Atlantic margins. Residual anomalies represent those from deep crustal and upper mantle sources (see section 5 for more details). Approximately linear segments shown as different colors corresponding to gradients 1 through 5 (see text). Profile locations shown in Figure 2. of the profile. The misfit was higher at the landward end where no data are available to constrain the thickness of the sedimentary layer. 5. Analysis [48] Free air gravity anomalies along the margins contain the combined effects of (1) water depth variations, (2) changes in sedimentary layer thickness and/or lithology, (3) intracrustal sources including intrusions and magmatic underplating, (4) crustal thinning, and (5) upper mantle thermal and/or compositional variations. Of the 18 models constructed, only 6 have sufficient constraints on the deep structure to allow contributions from effects 3 5 to be adequately evaluated. However, in all 18 models the shallow structure is relatively well constrained by seismic data and contributions from effects 1 and 2 and shallower portions of effect 3 can be estimated reasonably well. We have estimated contributions from deep sources by subtracting the gravity effects produced by water depth variations, sedimentary layers, SDRs, and crystalline basement from the observed free air data. Because the deepest sedimentary layers along the African margin reach to approximately 9 km (i.e., Orange Basin), the gravity effects due to shallow sources were calculated for the upper 10 km. On the South American margin the deepest sedimentary layers reach to >8 km (i.e., southern Pelotas Basin, line B). Because of this thickness together with an average SDR thickness of 4 km, we calculated the gravity effects for the upper 12 km along the South American margin. [49] The resulting residual anomalies for all 18 profiles are shown in Figure 4. The total change between the seaward and landward ends ( mgals, SW Africa, and mgals, South America) is attributed to the change in crustal thickness from continental to oceanic, and the high density contrast (i.e., 400 kg/m 3 ) between crust and upper mantle. The residual profiles comprise a series of approximately linear segments with varying gradient that can be observed over both margins (Figure 4). To identify key crustal elements, we have compared the residual anomalies with the deep structure along those profiles where seismic control is available (namely, profiles 1,2,3,4, E and F). Correspondence between residual anomalies and deep structure suggests that changes in gradient may be used to approximately delineate crustal domain boundaries. Using these gradient changes, we identify five crustal domains: [50] 1. A roughly uniform thickness continental crust associated with a relatively horizontal gradient at the 9of15

10 Table 1. Residual Gravity Gradient for Modeled Profiles Residual Gradient (mgal/km) Approximate Width (km) Underplating Profile Gradient 2 Gradient 3 Gradient 4 Gradient 3 Gravity A B C D E F onshore end (e.g., profiles 4 and E on Figure 4; profile F on Figures 3 and 4). [51] 2. A region of rapid crustal thinning associated with a steep gradient (1.2 to 2.3 mgals/km) located landward of the shelf break gravity high (e.g., near km 500, profile 8, and near km 300, profile D on Figure 3). [52] 3. A region of underplated crust associated with an intermediate gradient (0.5 to 1.1 mgals/km). The change from steep to more gradual slope coincides roughly with the seaward limit of continental crustal thinning and the underplated region (e.g., km 410 on profile 8, Figure 3). The landward limit of underplating is not as well constrained as the seaward extent but can be identified along profiles 2, 3, and 4 on the basis of reflection-refraction lines MCS 4 [Gladczenko et al., 1998] and BGR 1 and 2 [Bauer et al., 2000]. A similar change in gradient along profile F is also associated with the onshore extent of magmatic underplating observed on refraction data [Franke et al., 2002]. [53] 4. A region of gradual crustal thinning from thicker than normal to normal oceanic crust associated with a more gentle gradient (0.1 to 0.3 mgals/km). Refraction lines MCS 4 [Gladczenko et al., 1998], and BGR 1 and 2 [Bauer et al., 2000] indicate Moho depths of km that correlate with the landward limit of this gentle gradient. Further seaward, normal oceanic crustal thickness (6.5 7 km) and Moho depths of 12 km are observed. This spatial correlation between the seaward extent of magmatic underplating and the change in residual anomaly gradient is also observed along profile F (Figure 3). The change in slope along profile 3 is shifted slightly (30 km) seaward of the offshore limit of underplating observed on refraction transect BGR 1 [Bauer et al., 2000]. [54] 5. A normal oceanic crust with uniform thickness at the seaward end of the profiles associated with a more or less horizontal gradient (e.g., profiles A and F on Figures 3 and 4). [55] All profiles extend over the ocean-continent transition and therefore gradients 3 and 4 are observed on all profiles except profile E (Figure 4 and Table 1). A region of abrupt crustal thinning on profile E (near 0 km, Figure 4) produces a steep gradient 2 but the gentler gradient 3 is not observed on this particular residual gravity profile (Figure 4). Gradient 2 is clearly observed on profiles 5, 12, C, D, E and F (Figures 3 and 4). The roughly horizontal gradient 1 at the landward end is only clearly observed on profiles 3, 4, 10, D and F (Figures 3 and 4). However, the landward ends of profiles 1 3 are located km from the coastline and therefore may not reach normal thickness continental crust. Deviations in this horizontal gradient may also be the result of increasing crustal thickness onshore or due to horizontal density changes produced by igneous intrusions in the lower crust. Because the profiles were extended to reach normal oceanic crust, the more or less horizontal gradient 5 associated with normal oceanic crust is expected at the seaward termination of the profiles. [56] We have used seismic refraction data, our 2-D gravity models, and the spatial extent of gradient 3 to estimate alongstrike variations in the width of underplating. In particular we have determined how this width varies with distance from where the hot spot trace intersects the continental margin. Estimates show a southward decrease in underplating width on both margins. Similar maximum widths ( km) are observed near the hot spot traces on both margins. The width then decreases along the African margin to 195 km at 23 S and to 155 km at 24.5 S, south of which the width appears to be more or less constant. Along the South American margin the width decreases from 230 km to 90 km at 38 S before increasing to 130 km on profile F. The 130 km width along profile F, however, is based upon the velocity model of Franke et al. [2002] rather than the width of the SDRs wedge which is 60 km. [57] Limits of gradient 3 give a somewhat smaller estimate of the width of underplating than that obtained from gravity modeling. In our gravity modeling the limits of underplating are generally based on the assumption that the SDRs constitute the extrusive component of the magmatic underplating. The validity of this assumption is strengthened by the spatial association of inner and outer SDR wedges with the lower crustal high velocity layer seen in reflection [Gladczenko et al., 1997, 1998] and wide-angle refraction data (lines BGR 1 and 2) [Bauer et al., 2000]. Where there is seismic control, the residual gradient 3 estimate of the width agrees reasonably well with that indicated by the seismic data (e.g., profile 4). However, gravity models along profiles without deep seismic control (e.g., lines 5 to 12 along the SW African margin) are based upon along-strike extrapolation of SDR wedges. Constraints on the width of underplating from our gravity models are quite variable. The close correlation of the landward edge of the underplating with both the shelf break gravity anomaly high and the hinge zone [Clemson et al., 1997] suggest that this boundary can be constrained within km. In contrast, the seaward boundary is less well defined and variations of ±30 50 km are possible without significantly altering the fit with the gravity data. 6. Discussion [58] The deep structure of the South Atlantic margins has been previously interpreted in terms of both symmetric pure shear [Fontana, 1987], and asymmetric simple shear [Ussami et al., 1986; Castro, 1987] kinematic rifting 10 of 15

11 Figure 5. Comparison of reconstructed continental margins for (a) SE Greenland-Hatton Bank [after Hopper et al., 2003] and (b) Africa-South America (profiles 11 and E, this study). Crustal structure in Figure 5a is based upon seismic velocity analysis and in Figure 5b upon both gravity and seismic data. In Figure 5a the conjugate margins are reconstructed at magnetic anomaly C22 time. In Figure 5b the conjugate margins are reconstructed at the seaward edge of the interpreted magmatic underplating. Locations of profiles 11 and E are shown in Figure 2. models. Extensive underplating and associated long-lived uplift of mountains along the Namibian margin, however, are more consistent with asymmetric extensional models [Lister et al., 1986] in which the African margin south of the Walvis Ridge is considered an upper plate margin [Etheridge et al., 1989]. Our gravity modeling and analysis of residual gradients are also consistent with an asymmetric development that is expressed in both upper crustal structure and magmatic underplating. In particular, magmatic underplating is consistently better developed in extent, and possibly thickness, along the African margin. We have compared the amounts of underplating on conjugate portions of the margins using the rotation poles of Muller and Roest [1992] and Hall and Bird [2007], both of which give similar results. The larger extent of underplating in models 1 and A (Table 1) suggests that hot spot proximity may play a role in the amount of this material. However, the more or less constant extent of underplating along the African margin south of 24 S suggests that the hot spot influence is regionally quite limited. Similarly the hot spot influence does not appear to extend south of 31 S along the South American margin. The width of underplating along the African margin is similar to that observed at other volcanic passive margins including those not associated with a hot spot (e.g., NW Australia, U.S. east coast, and the Norwegian volcanic margin) where the underplated zone is 200 km wide [Eldholm and Grue, 1994; Tsikalas et al., 2005]. [59] Except for profiles 1 and A the extent of underplating for each conjugate pair is larger on the African side but the amount of asymmetry varies significantly along strike. In Figure 5 we compare the asymmetry displayed by conjugate pair 11-E with that of reconstructed continental margins of Greenland and Hatton Bank near 59 N [Hopper et al., 2003]. Both the extent of underplating and degree of asymmetry (2:1) are remarkably similar. Although passage of the Iceland plume beneath Greenland margin has been proposed as a possible explanation for the extra igneous material, Hopper et al. [2003] consider the lack of disruption in the magnetic anomaly pattern due to offaxis volcanism, and the large distance between the region and the site of major plume activity as evidence that the hot spot is not directly responsible for the asymmetry. Instead they propose that the asymmetry is inherent in the rifting process [Hopper et al., 2003]. Along strike further south, over the Edoras Bank, a more symmetrical development is observed [Holbrook et al., 2001]. We observe a similar change in the degree of asymmetry along strike both to the north (conjugate pair 5-B) and possibly to the south (12-F). [60] Clearly the nature of the rifting process is critical to understanding this asymmetry. Many dynamical rifting models [e.g., Braun and Beaumont, 1987; Bassi et al., 1993] are either intrinsically symmetric or predict symmetric extension. More recent rifting models that examine the role of frictional plastic strain softening suggest that the different extensional styles depend on the relative contributions from the frictional plastic and ductile layers, which promote asymmetry and symmetry, respectively [Huismans and Beaumont, 2002, 2003]. Rifting models in which symmetrical lithospheric extension takes place in the absence of strain softening [Huismans and Beaumont, 2002] or where the crust and upper mantle are coupled at high rifting velocities of 30 cm/a [Huismans and 11 of 15