# Crustal structure and magmatism of the Limpopo transform margin at the Southern Mozambique using wide-angle seismic data

Wei Wang, Satish C. Singh, Zhikai Wang, Aiguo Ruan, Yong Tang, Jérôme

Dyment, Sylvie Leroy, Zhaocai Wu, He Li, Chongzhi Dong

► To cite this version:

Wei Wang, Satish C. Singh, Zhikai Wang, Aiguo Ruan, Yong Tang, et al.. Crustal structure and magmatism of the Limpopo transform margin at the Southern Mozambique using wide-angle seismic data. Geophysical Journal International, 2025, 243 (1), pp.ggaf285. 10.1093/gji/ggaf285. hal-05347088

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# Crustal structure and magmatism of the Limpopo transform margin at the Southern Mozambique using wide-angle seismic data

Wei Wang,<sup>1,2,3</sup> Satish C. Singh,<sup>3</sup> Zhikai Wang,<sup>4</sup> Aiguo Ruan,<sup>1,2</sup> Yong Tang,<sup>1,2</sup> Jérôme Dyment,<sup>1,3</sup> Sylvie Leroy,<sup>5</sup> Zhaocai Wu,<sup>2</sup> He Li<sup>2</sup> and Chongzhi Dong<sup>2</sup>

Accepted 2025 July 22. Received 2025 July 20; in original form 2025 January 10

## SUMMARY

The Limpopo transform margin offshore southern Mozambique results from the separation of Gondwana along the East Africa continental margin. Over the last three decades, more than thirty different reconstruction models have been proposed, sometimes contradicting each other. Here, we present results from the traveltime tomography of wide-angle seismic data acquired during the second China–Mozambique Joint Cruise, allowing the interpretation of the crustal structure and magmatism in the Limpopo Corridor and the Mozambique Basin. Using these results, we determine the extent of the Continent Ocean Transition and the location of the Continent Ocean Boundary on the southern Mozambique margin. The seismic profile is 442 km long, extending from the eastern part of the North Natal Valley in the west and crossing the Limpopo Corridor and the Mozambique Basin to the east. Based on the tomographic velocity model, we delineated three distinct domains from west to east along the profile: (1) a western transitional domain with anomalous or mixed crust, bounded by the Mozambique Fracture Zone to the east, where the crust gradually thins eastward from ~14 km at distance 45 km to ~10.8 km at distance 140 km; (2) a domain of thickened oceanic crust resulting from enhanced magmatism, where the crust thins eastward, from ~10.8 km to ~8.5 km over ~100 km distance; and (3) an eastern domain of normal oceanic crust, where the average crustal thickness is ~8 km. We suggest that (1) the western transitional domain roughly corresponds to the Limpopo Corridor and is of continental crustal origin but was affected and modified by strike-slip motion and magmatic activity, resulting in anomalous or mixed crust. The eastern Continent Ocean Boundary of the Limpopo Margin is close to the Mozambique Fracture Zone. (2) The thickened oceanic domain thins eastwards, and the crustal velocity and thickness change dramatically compared to the oceanic domain. This domain seems to have strongly interacted and contaminated by the Limpopo Corridor during the opening of Mozambique Basin and seafloor spreading. (3) The eastern oceanic domain shows a relatively uniform oceanic crust of ~8 km and high velocity up to 7.4 km s<sup>-1</sup> in the lower crust, suggestive of a hotter mantle that produces more MgO-rich melts probably due to the influence of a thermal mantle anomaly.

**Key words:** Tomography; Indian Ocean; Continental margins: transform; Crustal structure.

## 1 INTRODUCTION

The definition of transform continental margin can be traced as far back as the discovery of transform faults in the ocean (Wilson 1965). This kind of margin is primarily formed by strike-slip motions of the continental margin along transform faults (e.g. Mascle &

Blarez 1987; Basile et al. 1998; Bird 2001) and sometimes referred to as strike-slip or shear continental margin (e.g. Mascle 1976; Rabinowitz & Labrecque 1979; Scrutton 1979; Nagel et al. 1986). Based on studies of the Côte d'Ivoire–Ghana transform margin, morphologically, a transform margin was traditionally characterized by a sharp necking zone, a steep continental slope, a marginal

<footnote n="1"><sup>1</sup>School of Oceanography, Shanghai Jiao Tong University, Shanghai 200030, China. E-mail: ruanag@163.com, tangyong@sio.org.cn</footnote>

<footnote n="2"><sup>2</sup>State Key Laboratory of Submarine Geoscience, Second Institute of Oceanography, Ministry of Natural Resources, Hangzhou 310012, China.</footnote>

<footnote n="3"><sup>3</sup>Université Paris Cité, Institut de Physique du Globe de Paris, CNRS, Paris 75005, France</footnote>

<footnote n="4"><sup>4</sup>School of Ocean and Earth Science, University of Southampton, Southampton SO14 3ZH, UK</footnote>

<footnote n="5"><sup>5</sup>Sorbonne Université, CNRS-INSU, UMR 7193, Institut des Sciences de la Terre de Paris, Paris 75005, France</footnote>
plateau and/or a marginal ridge (e.g. Sage *et al.* 2000; Basile 2015).

The traditional model of the evolution of the transform continental margin includes three main stages: (a) the intracontinental transform stage; (b) the active transform stage; and (c) the passive transform stage (e.g. Mascle & Blarez 1987; Lorenzo 1997; Bird 2001; Basile 2015; Le Pourhiet *et al.* 2017). This traditional model is mainly derived from studies of transform continental margins in magma-poor or non-volcanic settings, ignoring the transform margins in magma-rich scenarios (e.g. Basile 2015; Mercier de Lépinay *et al.* 2016). The structure and tectonics of transform margins show complexity and diversity due to their different formation and evolution. For example, transform margins are rarely found in areas affected by thermal events (hotspots or back-arc basins, for instances, e.g. Basile 2015; Mercier de Lépinay *et al.* 2016). However, the Limpopo transform margin located on the southern Mozambique margin is a magma-rich transform margin and may have been influenced by a thermal mantle anomaly or plume (e.g. Li *et al.* 2021; Roche *et al.* 2021; Watremez *et al.* 2021).

The crustal nature of the Limpopo transform margin is crucial in the reconstruction of the East Gondwana (e.g. Goodlad *et al.* 1982; Martin & Hartnady 1986; Cox 1992; König & Jokat 2010; Leinweber & Jokat 2012; Nguyen *et al.* 2016; Reeves *et al.* 2016; Mueller & Jokat 2019; Thompson *et al.* 2019), but it remains poorly known before the PAMELA (Passive Margins Exploration Laboratories) project (e.g. Moulin & Aslanian 2016; Moulin & Evain 2016) and the second China–Mozambique Joint Cruise. In the past, the seismic data were relatively sparse, especially lacking wide-angle seismic data recorded by oceanic bottom instruments extending the complete domain. Based on the potential field data, such as gravity and magnetics, and the results from plate kinematic studies, previous researchers proposed that the Limpopo Margin and the adjacent Mozambique Coastal Plain are either made of thickened oceanic crust (e.g. Ludwig *et al.* 1968; Green 1972; Watts 2001; Tikku *et al.* 2002; Klausen 2009; Leinweber & Jokat 2011; Mueller & Jokat 2019), thinned continental crust with high-magmatic intrusion content (e.g. Dingle & Scrutton 1974; Lafourcade 1984; Domingues *et al.* 2016; Hanyu *et al.* 2017), or hybrid crust consisting of highly thinned continental crust and large amounts of magmatic additions (e.g. Tomasi *et al.* 2021). The first crustal model was supported by weak linear magnetic spreading anomalies in the North Natal Valley and by continuous regional gravity anomalies (e.g. Tikku *et al.* 2002; Leinweber & Jokat 2011). The second crustal model was mainly supported by the crustal structure of the North Natal Valley, which was derived from the inversion of vector geomagnetic anomaly data and satellite gravity data (e.g. Hanyu *et al.* 2017). The third crustal model was based on a plate kinematic study, where the crustal thickness and lithosphere thinning in this region were estimated using 3-D gravity inversion and subsidence analysis (e.g. Tomasi *et al.* 2021).

The PAMELA project MOZ3/5 cruises acquired a large amount of wide-angle seismic data and other geological and geophysical information on the southern Mozambique margin (Fig. 1), providing the main active source seismic constraints on the crustal properties and tectonic attributes of the region to date (e.g. Moulin & Aslanian 2016; Moulin & Evain 2016). The studies using the PAMELA seismic data concluded that both the Mozambican Coastal Plain and the North Natal Valley are floored by thick continental crust, which is consistent with the plate-kinematic model proposed by Thompson *et al.* (2019), that is, thinning eastwards and southwards, and that the entire area is affected by abundant magmatism (e.g. Moulin *et al.* 2020; Aslanian *et al.* 2021; Evain *et al.* 2021; Leprêtre *et al.* 2021; Moulin *et al.* 2023; Schnürle *et al.* 2023a), possibly induced

by a mantle plume (e.g. Roche et al. 2021, 2022; Watremez et al. 2021). Overall, the project has solved the controversy about the crustal nature of majority of the region. However, there is still a controversy about the nature and evolution of the N–S distributed Limpopo Corridor in the eastern part of the margin (e.g. Li et al. 2021; Roche et al. 2021; Watremez et al. 2021; Moulin et al. 2023), which hinders the delineation of the Continent Ocean Transition and the definition of the Continent Ocean Boundary between the Limpopo Margin and the Mozambique Basin.

In this study, we present the seismic refraction profile DZ02 (Fig. 1) that provides new evidence on the crustal structure and magmatism along the Limpopo transform margin. The profile DZ02 transects the Limpopo Corridor and extends eastward into the Mozambique Basin farther than the PAMELA profiles (e.g. Moulin et al. 2020; Evain et al. 2021; Roche et al. 2021; Watremez et al. 2021; Moulin et al. 2023), which helps to solve the controversy surrounding the Limpopo Corridor and sheds new light on the crustal structure and magmatism of the Mozambique Basin.

## 2 GEOLOGICAL AND TECTONIC SETTING

From the Early Jurassic period (~180 Ma), the Gondwana Continent progressively rifted and broke-up from north to south, forming the northern, central and southern continental margins along Mozambique, and producing a series of divergent and strike-slip margins (e.g. König & Jokat 2006; Mahanjane 2012; Klimke & Franke 2016; Evain et al. 2021). During the first phase, from the initial rift to ~120 Ma, the southward motion of the Madagascar–India plate along the Davie Fracture Zone led to the opening of the Somalia Ocean, while the Antarctica–Australia plate drifted southwards relative to Africa along the Mozambique Fracture Zone, resulting in the formation of the Mozambique Channel. In the Second phase, beginning in the Valanginian, the Patagonia plate escaped southwestwards from the coast of South Africa along the Agulhas–Falklands Fracture Zone, coinciding with the seafloor spreading in the South Natal Valley (e.g. Evain et al. 2021; Roche et al. 2023). These margins are crucial areas for understanding the evolution of the Gondwana as their crustal nature and geometry strongly impact the plate kinematic reconstruction of Gondwana. Especially, the debate about continental or oceanic crust for the Mozambique Coastal Plain and the North Natal Valley on the southern Mozambique margin has led to tens of different plate kinematic reconstruction models for Gondwana (e.g. Sahabi 1993; König & Jokat 2006; Eagles & König 2008; Gaina et al. 2013; Reeves et al. 2016; Mueller & Jokat 2019; Thompson et al. 2019).

Based on the results of studies using ocean bottom seismometers (OBSs) and multichannel streamer (MCS) data from the PAMELA MOZ3/5 Cruises, the Mozambique Coastal Plain and the North Natal Valley were identified as continental crust domains (e.g. Moulin et al. 2020, 2023; Aslanian et al. 2021; Evain et al. 2021; Leprêtre et al. 2021; Watremez et al. 2021; Schnürle et al. 2023a). In this model, the Mozambique Coastal Plain and the North Natal Valley are bounded to the west by the Kaapval Craton via the Lebombo Monocline, and to the northwest by the Zimbabwe Craton via the Mateke-Sabi Monocline (Fig. 1b; e.g. Salman & Abdulla 1995; Watkeys 2002; Jourdan et al. 2005). To the south they are roughly separated from the South Natal Valley and the Mozambique Ridge by the Continent Ocean Boundary at the Ariel Graben (Fig. 1b; e.g. Moulin et al. 2020; Leprêtre et al. 2021; Roche et al. 2022; Schnürle et al. 2023a). The Limpopo Margin
<figure><img src="image_2.png" /><figcaption>**Figure 1.** (a) Bathymetry map showing the study area and the location of the wide-angle seismic profile DZ02. OBSs used in this work are marked by red dots, except those presented in Figs 2–4 that are highlighted in yellow, and other OBSs are marked by grey dots. The bathymetry data is from GEBCO_2020. Dashed lines (from MZ1 to MZ7) are the seismic profiles of PAMELA MOZ3/5 cruises (e.g. Moulin & Aslanian 2016; Moulin & Evain 2016), dashed line S01 is the seismic profile of the first China–Mozambique Joint Cruise (e.g. Li et al. 2021), and dashed lines (20 070 201, 20 070 202 and 20 140 010) are the seismic profiles from the Central Mozambique Margin (Leinweber et al. 2013; Mueller et al. 2016). The global inset shows the location of the study area. (b) Free-air gravity anomaly (ICGEM, International Centre for Global Earth Models, Ince et al. 2019) in the study area. The main features known in the area are from Mueller & Jokat (2019), which are based on Levinweber & Jokat (2012) and Muller & Jokat (2017) for the magnetic spreading anomalies in the Mozambique Basin. NNV: North Natal Valley, SNV: South Natal Valley, MozR: Mozambique Ridge, AG: Ariel Graben, DP: Dana Plateau, GP: Galathea Plateau, VP: Vauban Plateau, LM: Limpopo Margin, LF: Limpopo Fault, LC: Limpopo Corridor, MFZ: Mozambique Fracture Zone, MB: Mozambique Basin.</figcaption></figure>

is located east of the Mozambique Coastal Plain and the North Natal Valley, and a nearly N–S striped-area fringes its easternmost edge, known as the Limpopo Corridor. The Limpopo Corridor is bounded by the Limpopo Fault to the west and the Mozambique Fracture Zone to the east, respectively, with the Mozambique Basin lying to the east of the Mozambique Fracture Zone (Fig. 1; e.g. Moulin *et al. 2020*; Evain *et al. 2021*; Li *et al. 2021*; Moulin *et al. 2023*).

The crustal architecture and properties of the Limpopo Corridor are crucial for delineating the Continent Ocean Transition on the eastern edge of the southern Mozambique margin. The previous studies have led to two models of the origin of the Limpopo Corridor. The first model considers that the Limpopo Corridor originated from a continental crust which was thinned and affected by strike-slip motion or highly oblique rifting accompanied by magmatic intrusions (e.g. Moulin *et al. 2020, 2023*; Evain *et al. 2021*; Li *et al. 2021*). This eventually led to the formation of a mixed or anomalous crust, and its velocity structure is neither extended continental nor oceanic, just like a buffer zone of the Limpopo transform margin (e.g. Aslanian *et al. 2023*). In this model, the Continent Ocean Boundary roughly coincides with the Mozambique Fracture Zone on the eastern edge of the Limpopo Corridor (e.g. Moulin *et al. 2020, 2023*; Evain *et al. 2021*; Li *et al. 2021*). The second group of authors argue that the Limpopo Margin first underwent a

NW–SE rifting and then transitioned clockwise to N–S strike-slip motion during the post-rift stage (e.g. Watremez *et al. 2021*; Roche *et al. 2021, 2022, 2023*). The Limpopo Fault is a major fault of the Limpopo Margin and is the eastern limit of the thinned continental crust of the Mozambique Coastal Plain and North Natal Valley. It abuts a narrow transitional domain (~10–20 km) to the east, where the crust was possibly subjected to intensive faulting and magmatism, and then reached the thickened oceanic domain or the normal oceanic domain. In this case, the Continent Ocean Boundary would be located on the eastern side of the Limpopo Fault and close to the Limpopo Fault, and most of the Limpopo Corridor would belong to the thickened oceanic domain (e.g. Watremez *et al. 2021*; Roche *et al. 2021, 2022*).

# 3 DATA

## 3.1 Data acquisition and processing

The second China–Mozambique Joint Cruise in June 2021 onboard R/V ‘Dayang Hao’ acquired three control-source wide-angle seismic profiles offshore Mozambique (DZ01, DZ02 and DZ04 in Fig. 1). The profiles DZ02 and DZ04 were shot in a nearly W–E direction whereas the profile DZ01 was nearly in N–S direction (Fig. 1). A total of 70 OBS sites were deployed during the cruise
using 28 OBSs in three deployments, and each OBS was equipped
with a three-component geophone and a hydrophone. The OBS
spacing was ~20 km along all profiles. The acquisition source was
an airgun array consisting of four Bolt air guns, producing a total
volume of 8000 cubic inches (~131.1 L). The airgun array was
towed ~100 m behind the vessel at ~10 m depth below sea surface
and was fired at 200 m interval (Wang et al. 2024). The MCS seis-
mic data and multibeam data were also collected along the three
profiles.

In this study, we performed traveltime tomography to constrain
the velocity structure along the DZ02 profile. We corrected the in-
ternal clock drift of OBS data by comparing the differences between
the GPS time and the OBS internal clock before deployment and
after recovery. Then we relocated the OBSs by comparing the ob-
served and calculated arrival times of direct water wave near the
zero offset (approximately ±3 km), following a least-squares inver-
sion algorithm (e.g. Du et al. 2018). To identify different seismic
arrivals, we filtered the seismic data between 3 and 15 Hz using a
fifth-order minimum-phase Butterworth bandpass filter.

## 3.2 Phase identification

The *P*-wave refractions from sedimentary layer (Ps) were identified outside of the direct water wave on six OBSs (01–06) (see example in Fig. 2). All the OBS gathers show clear *P*-wave refractions from the crust (Pg), *P*-wave reflections from the basement (PbP) and the Moho boundary (PmP) (Figs 2–4). The Pg arrivals have high amplitude at near offsets for all OBSs but they have weak amplitudes and highly variable apparent velocities for some OBSs, possibly because of the complex basement topography. The PmP arrivals were identified according to clear triplications, increased seismic amplitude, or both (e.g. Wang *et al. 2020*; Ingale & Singh 2021). Only OBSs deployed east of 200 km horizontal distance (except OBS13, Fig. 5a) recorded *P*-wave refractions from the mantle (Pn) (see example in Fig. 4).

A total of 18 964 arrivals were manually picked, including 513 Ps, 3149 PbP, 7525 Pg, 5471 PmP and 2306 Pn. We assigned the picking uncertainties based on the signal quality and visual inspection, and the uncertainties for all sets of picks are presented in Table 1.

# 4 MODELLING

## 4.1 Initial model

We applied traveltime tomography to constrain the velocity structure of the subsurface along the seismic profile. The starting model is 442 km long and 30 km deep, discretized at 200 m horizontally and 50 m vertically. It contains four layers, from top to bottom the water column, sedimentary layer, crust and mantle. The seafloor depth was obtained from the high-resolution multibeam bathymetry data, and the depth of basement was constrained by the MCS data (Fig. S1, Supporting Information; Wang et al. 2024). The initial depth of the Moho boundary was taken from the tomography result along the MZ5 seismic profile from the PAMELA project (Watremez et al. 2021), which is in close vicinity to our seismic profile.

The velocity of the water column was set to 1.5 km s⁻¹ and was fixed in tomography. The starting velocity of other layers was constructed by including a 1-D linear velocity to generate a 2-D velocity model. The velocity of the sedimentary layer increases from 1.6 km s⁻¹ at the top to 4.0 km s⁻¹ at 7 km depth. The crustal velocity increases from 4.5 km s⁻¹ at the top of the crust to 7.0 km

s⁻¹ at 20 km depth. The velocity is 8.0 km s⁻¹ at the top of the mantle and increases to 8.2 km s⁻¹ at the base of the model.

## 4.2 Traveltime tomographic inversion

Starting from the initial model, we performed a ray-based traveltime tomography (Van Avendonk et al. 1998, 2004) to iteratively update the velocity structure of the subsurface. The shortest-path algorithm (Moser 1991) was used to calculate the ray paths of different seismic phases and their arrival times. The velocity model was refined by gradually minimizing the difference between the modelled and manually picked traveltimes using a least-squares approach (Van Avendonk et al. 1998, 2004). The regularization and smoothing parameters were tested and selected in each iteration to ensure a smooth model is obtained after traveltime tomography (e.g. Wang et al. 2022; Wang & Singh 2022).

We adopted a layer stripping strategy in the traveltime tomogra-
phy. We first inverted the Ps and PbP arrivals simultaneously to con-
strain the sediment velocity and the depth of the basement. Given
that the starting depth of the basement was determined from the
MCS data and the seismic phases from the sedimentary layer were
relatively sparse, we designed the inversion parameters to domi-
nantly update the sediment velocity and only slightly modified the
depth of basement. Then, we fixed the structure of the sedimentary
layer and inverted the traveltime of Pg arrivals to constrain the upper
crust velocity, followed by a joint inversion of Pg and PmP arrivals
to constrain both crustal velocity and Moho depth. Finally, we fixed
the structure above the Moho and inverted the traveltime of Pn ar-
rivals to constrain the upper mantle velocity. The final tomographic
model is shown in Fig. 5(a).

The standard chi-square ($\chi^2$) value was used to assess the differ-
ence between the modelled and manually picked traveltimes. The
final $\chi^2$ values and root-mean-square (RMS) misfits for all types of
seismic arrivals are presented in Table 1. The spatial distribution of
residuals is presented in Fig. S2 (Supporting Information).

## 4.3 Uncertainty analysis and resolution test

4.3.1 Ray coverage

We used the derivative weight sum (DWS; Toomey & Foulger 1989) to represent the ray density throughout the final tomographic model (Fig. 5b). In the sedimentary layer, the ray coverage is relatively poor, especially in areas of thin sedimentary cover where the number of picked turning rays Ps and reflective phases PbP are small. The crust is well covered by dense rays, with the upper crust being well constrained by Pg rays whereas the lower crust and Moho are constrained by dense PmP rays. However, the Pn rays are relatively sparse and unevenly distributed, constraining only a very limited portion of the uppermost mantle, which shows a velocity of approximately 8.0 km s⁻¹ (Fig. 5a).

4.3.2 Monte carlo uncertainty analysis

We applied a Monte Carlo analysis (Korenaga et al. 2000) to ex-
amine the dependence of the tomographic inversion on the starting
model and to quantify the uncertainty of the tomographic model.
Because this study mainly interprets the crustal structure, we only
performed the Monte Carlo analysis for crustal velocity and Moho
depth.
<figure><img src="image_3.png" /><figcaption>**Figure 2.** Seismic gather recorded by OBS04 on the DZ02 profile. (a) Seismic record after bandpass filtering between 3 and 15 Hz. The traveltime (*T*) is reduced using a reduction velocity of 8.0 km s⁻¹. *X* is the offset. (b) Seismic record with the traveltime picks superimposed. *P*-wave in water, Pw: red; *P*-wave in sediments, Ps: orange; *P*-wave reflected from the basement, PbP: cyan; Crustal turning rays, Pg: green; Wide-angle reflection from Moho, PmP: blue; Mantle turning rays, Pn: Pink. (c) Comparisons of the picked traveltimes (coloured dots) and modelled traveltimes (black dots).</figcaption></figure>

Fifty starting models were created by modifying the initial 1-D crustal structure hung below the sedimentary layer. We randomly perturbed three parameters, the velocities at the top and bottom of the crust by ±5 per cent, and the Moho depth by ±1 km. We used these starting models to invert the picked traveltimes, using the same tomography algorithm and inversion scheme as described before. We obtained fifty tomographic models showing similar velocity structure as that shown in Fig. 5(a), indicating our tomographic inversion is robust with respect to the data used. We calculated the standard deviation of the fifty inverted models and used it as a measure of uncertainty (variance from the mean model) of the tomographic result (Fig. 6).

The variance in the final crustal model is ≤0.1 km s⁻¹ in the upper crust and ≤0.15 km s⁻¹ in the lower crust (Fig. 6a). Large variance up to 0.2 km s⁻¹ is only observed around ~35 km along profile distance where the ray coverage is low. The maximum standard

deviation of the Moho depth is ~500 m, and the preferred Moho falls in the standard deviation of the average Moho depth from the Monte Carlo analysis (Fig. 6b).

### 4.3.3 Checkerboard testing

To evaluate the resolution of the tomographic velocity model, we conducted checkerboard tests using the same source-receiver geometry as that of the field OBS data (e.g. Zelt & barton 1998; Canales et al. 2003). Because the sedimentary layer has relatively sparse rays and this study mainly focuses on the crustal structure, we only performed the checkerboard resolution tests for crustal velocity structure. The checkerboard models were designed by adding 2-D sinusoidal velocity perturbation with the maximum amplitude of ±10 per cent into the previous final velocity model. The size of the anomaly patterns in the checkerboard models were 25 km
<figure><img src="image_4.png" /><figcaption>**Figure 4.** Seismic gather recorded by OBS15 on the DZ02 profile. Same legend and colour code as Fig. 2.</figcaption></figure>

× 2 km (horizontal × vertical; Fig. 7a), 20 km × 2 km (Fig. 8a) and 15 km × 1 km (Fig. 9a), respectively. We simulated the synthetic travel times using the checkerboard models and the shortest-path ray tracing algorithm (Moser 1991), and Gaussian random noise was added to these synthetic data to mimic the picking uncertainty.

We inverted these synthetic traveltimes using the same workflow and inversion parameters as those used for inversion of field OBS data set. The anomalies of size 25 km × 2 km and 20 km × 2 km are well recovered after tomography with better recovery in the upper crust (Figs 7b and 8b) due to the denser ray coverage in the upper crust. The anomalies of size 15 km × 1 km are not recovered at all (Fig. 9b). Collectively, the checkerboard tests show that traveltime tomography of our OBS data can recover velocity anomalies of size 25 km × 2 km and 20 km × 2 km with ± 10 per cent velocity perturbation.

## 5 RESULTS AND INTERPRETATION

The wide-angle seismic profile DZ02 starts from the eastern limit of the North Natal Valley in the west, passes through the Limpopo Corridor and crosses through the Mozambique Basin in the east (Fig. 1). According to the different features of the tomographic velocity model along the profile, we divided the profile into three domains (Fig. 5a): the transitional domain, the thickened oceanic crust domain and the normal oceanic crust domain. The westernmost part of the profile, from the start of the profile to the ~140 km horizontal distance, is interpreted as the transitional domain, roughly corresponding to the Limpopo Corridor. Further eastward up to the ~240 km horizontal distance, the central part is the thickened oceanic domain. The easternmost part, from the ~240 km horizontal distance to the eastern end of the profile is the normal oceanic domain. We describe the velocity structure and nature of crust in these domains in detail below.
<figure><img src="image_5.png" /><figcaption>**Figure 5.** (a) Final *P*-wave velocity model of DZ02 profile. OBS locations are indicated by red inverted triangles. Thick blue lines indicate the Moho constrained by PmP. (b) Ray coverage through the final velocity model, represented by the derivative weight sum (DWS). (c) Vertical velocity gradient of crust in the final model. The dashed line indicates the upper/lower crust boundary, which is determined by the 6.5 km s⁻¹ contour. b.s.l.: below sea level, LF: Limpopo Fault, MFZ: Mozambique Fracture Zone.</figcaption></figure>

## 5.1 The transitional domain

The transitional domain delineated in the DZ02 profile nearly coincides with the Limpopo Corridor, which is bounded by the Limpopo Fault to the west and the Mozambique Fracture Zone to the east (Fig. 1b). The velocity of the uppermost crust in the transitional domain is ~5 km s⁻¹ (Fig. 5a). The normal crust shows a high-velocity gradient in the upper part and a significantly reduced velocity gradient in the lower part. Here, we took the 6.5 km s⁻¹ velocity contour, the global mean velocity at the base of the upper oceanic crust (White *et al.* 1992; Christeson *et al.* 2019), as a proxy of the upper

and lower crustal boundary along the whole profile. The upper crust in this domain shows a large vertical velocity gradient of ~0.36 s⁻¹ (Fig. 5c). The lower crustal velocity increases with depth to ~7.2 km s⁻¹ at the base of crust, showing a significantly reduced vertical velocity gradient of ~0.08 s⁻¹ (Fig. 5c). In this domain, the crust gradually thins eastward from ~14 km at distance 45 km to ~10.8 km at distance 140 km along the profile (Figs 5a and 10a). The eastward crustal thinning is mainly accommodated by the changes in lower crustal thickness, as the upper crustal thickness is nearly constant in this domain (Figs 5a and 10a).
<figure><img src="image_6.png" /><figcaption>**Figure 3.** Seismic gather recorded by OBS09 on the DZ02 profile. Same legend and colour code as Fig. 2.</figcaption></figure>

<table><caption>**Table 1.** Picking and modelling statics for all observed compressional phases.</caption><thead><tr><th>Phases</th><th>Number of picks</th><th>Pick uncertainty (ms)</th><th>RMS (ms)</th><th>χ<sup>2</sup></th></tr></thead><tbody><tr><td>Ps</td><td>513</td><td>40</td><td>62</td><td>2.4</td></tr><tr><td>PbP</td><td>3149</td><td>80</td><td>68</td><td>0.7</td></tr><tr><td>Pg</td><td>7525</td><td>60</td><td>68</td><td>1.3</td></tr><tr><td>PmP</td><td>5471</td><td>80</td><td>64</td><td>0.6</td></tr><tr><td>Pn</td><td>2306</td><td>100</td><td>76</td><td>0.6</td></tr><tr><td>Total</td><td>18 964</td><td>40–100</td><td></td><td></td></tr></tbody></table>

To investigate the lateral variations in crustal velocity along the profile, we extracted the velocities at 0.5 and 1.0 km depth below the basement, and at 0.5, 1.0 and 2.0 km depth above the Moho, respectively (Fig. 10b). According to this result, we can further determine that the velocity gradient of the upper part of the crust is significantly greater than that of the lower part of the crust above the Moho, and so are the lateral velocity variation trends of upper and lower crusts. The upper crustal velocity in the easternmost part of the transitional domain increases sharply, roughly 0.6 km s⁻¹

over 40 km distance and then decreases even more sharply after 140 km (Figs 5a and 10b). We set the extreme velocity area in the uppermost crust as the eastern boundary of the transitional domain, which roughly coincides with the Mozambique Fracture Zone delineated in previous studies (e.g. Moulin et al. 2020). Therefore, the velocity increase is probably induced by the strike-slip processes and crustal interaction, which may have occurred between the transitional domain and the thickened oceanic domain along the Mozambique Fracture Zone. The lower crustal velocity varies dramatically laterally, which may relate to the widespread and intense magmatism occurring during the formation of the region, such as the episodic magmatic intrusions (from Karoo to Miocene events; e.g. Moulin et al. 2020; Evain et al. 2021; Roche et al. 2021; Watremez et al. 2021; Moulin et al. 2023). The volcanosedimentary sequence, identified by tectono-stratigraphic analysis using the MCS data, also supports the presence of magmatism of this region (e.g. Li et al. 2021; Evain et al. 2021; Roche et al. 2021; Moulin et al. 2023; Schnürle et al. 2023a;). Additionally, the strike-slip or extensional faulting along the Mozambique Fracture Zone
<figure><img src="image_7.png" /><figcaption>**Figure 6.** Monte Carlo analysis results. (a) Uncertainty in the crustal velocity. (b) Moho depth and uncertainty in the Moho depth. The maximum standard deviation of the Moho depth is ~500 m. b.s.l.: below sea level.</figcaption></figure>

may also have influenced the lower crustal velocity structure (e.g. Evain *et al.* 2021).

We extracted 1-D velocity-depth (1-D-Vz) profiles of crust and compared them with those of the 125–143 Ma Atlantic oceanic crust (White *et al.* 1992) and the Limpopo Corridor (Fig. 11). Overall, the 1-D velocity curves of the transitional domain are in good agreement with those of the Limpopo Corridor from the previous results (PAMELA project; Moulin *et al.* 2020; Evain *et al.* 2021; Watremez *et al.* 2021; Moulin *et al.* 2023) but are less consistent with the 125–143 Ma Atlantic oceanic crust in terms of crustal velocity and thickness (Fig. 11b).

## 5.2 The thickened oceanic domain

We interpret the thickened oceanic domain between 140 and 240 km horizontal distance along the profile. In this domain, the crustal velocity is ~5.2 km s⁻¹ at the uppermost crust and increases to 6.5 km s⁻¹ at ~9 km depth (vertical velocity gradient of ~0.36 s⁻¹; Figs 5a and 5c). Further below, the lower crustal velocity increases slowly to ~7.2 km s⁻¹, yielding a gradient of ~0.13 s⁻¹. Pronounced lateral velocity variation is observed in the upper crust where the velocity increases eastward (Figs 5a and 10b). The 1-D velocity-depth profiles of this domain (Figs 11a and c), although characterized by a slightly thicker crust, are more closely resemble the 125–143 Ma Atlantic oceanic crust (White *et al.* 1992) than those of the Limpopo Corridor. The crust in this domain thins eastward, from ~10.8 to ~8.5 km over ~100 km distance. However, the upper and

lower crustal thicknesses in this domain are quite peculiar, with the upper crust first thickening from ~3.5 to ~4.0 km, and then slowly thinning down to about 3.0 km (Fig. 10a). Conversely the thickness of the lower crust thins rapidly by ~2.0 km over the first 60 km lateral distance and then remains at ~5.5 km. Although the thinning of the crust is gentler than in the transitional domain, the lower crustal thickness of the thickened oceanic domain changes more dramatically (Fig. 10a).

The top of the crust in the thickened oceanic domain is characterized by laterally undulating velocity profiles below the basement (Fig. 10b). The lateral variations of the velocity profiles above the Moho indicate that the vertical velocity gradient in the lower part of this domain is larger than in the transitional domain. The eastern part of this domain is marked by an abrupt increase of the top crustal velocity. The thickened oceanic domain is significantly different from its neighbouring domains in terms of crustal thickness, velocity, and their variations, indicating differences in terms of crustal nature and formation, and magmatism, between this and the adjacent domains.

## 5.3 The normal oceanic domain

The easternmost part of the profile, from distance ~240 km to the eastern end of profile, is interpreted as the normal oceanic domain. In this domain, the uppermost crustal velocity is the largest along the whole DZ02 profile, up to ~5.8 km s⁻¹ (Figs 5a and 11d). The upper crustal velocity increases rapidly at a ~0.32 s⁻¹ vertical velocity
<figure><img src="image_8.png" /><figcaption>**Figure 7.** Checkerboard result showing the recovery of velocity anomaly checks that have dimensions of 25 km × 2 km with a maximum velocity perturbation of 10 per cent. (a) The input checkerboard pattern. (b) The recovered velocity anomalies. b.s.l.: below sea level.</figcaption></figure>

gradient to 6.5 km s⁻¹ over ~2 km depth (Figs 5a and c). Laterally, the upper crust shows little variation in velocity (Fig. 10b). Below the upper crust, the velocity further increases with depth to 7.3–7.4 km s⁻¹ at the base of the crust, giving a vertical velocity gradient of ~0.17 s⁻¹ which is higher than those of the lower crust in the other two domains (Figs 5a and c). Due to the presence of extensive high-velocity bodies, the lateral velocity changes significantly in the lower crust, and the vertical velocity gradient in the lower crust is more substantial than in the thickened oceanic domain (Figs 5a, c and 10b).

The crustal thickness in this domain is uniform. The average crustal thickness is ~8 km, which is ~2 km thicker than the global average of oceanic crust (Fig. 10a; White et al. 1992; Christeson et al. 2019). The 1-D velocity-depth profiles extracted from this domain mainly fall into the velocity-depth envelope for 125–143 Ma crust in the Atlantic Ocean (Fig. 11d; White et al. 1992), further supporting that the crust is of oceanic origin.

# 6 DISCUSSION

## 6.1 Nature of the anomalous crust below the Limpopo Corridor and location of the Continent Ocean Boundary

The Limpopo Corridor varies along the N–S direction (Fig. 1b), narrowing to the north and widening to the south, and its crust thins to the east and thickens to the south. Its crustal velocity neither

corresponds to a thinned continental crust nor oceanic crust, and hence is defined as anomalous crust (e.g. Evain et al. 2021; Moulin et al. 2023).

The crustal nature of the Limpopo Corridor is a key to locate the Continent Ocean Boundary on the eastern side of the southern Mozambique margin. As mentioned above, two models have been proposed for the origin of the Limpopo Corridor: the first one suggests that the Continent Ocean Boundary is located at the eastern end of the Limpopo Corridor, roughly coinciding with the Mozambique Fracture Zone (e.g. Moulin et al. 2020, 2023; Evain et al. 2021; Li et al. 2021). The bathymetry around the Mozambique Fracture Zone sharply deepens, while showing a clear transition between positive and negative free air gravity anomalies (Fig. 1b). The positive gravity anomalies of the Limpopo Corridor are interpreted as reflecting the presence of large and thick contourite structures (e.g. Thompson et al. 2019; Li et al. 2021; Babonneau et al. 2022). Along with the opening of the Mozambique Basin, a broad zone of strike-slip or highly oblique rifting developed on the Limpopo Margin, resulting in brittle deformation of the upper crust, with the Limpopo Fault as a major strike-slip fault in this area. The Limpopo Corridor, bounded by the Limpopo Fault and Mozambique Fracture Zone, is most strongly affected by the N–S strike-slip or highly oblique rifting, and its lower crust is speculated to have been formed by oceanward flow and exhumation of the eastern fringes of the Mozambique Coastal Plain and North Natal Valley lower continental crust, accompanied by intense magmatism (intrusion or underplating; e.g. Evain et al. 2021; Moulin et al.
<figure><img src="image_9.png" /><figcaption>**Figure 8.** Checkerboard result showing the recovery of velocity anomaly checks that have dimensions of 20 km × 2 km with a maximum velocity perturbation of 10 per cent. (a) The input checkerboard pattern. (b) The recovered velocity anomalies. b.s.l.: below sea level.</figcaption></figure>

2023). It behaves just like a buffer zone of the Limpopo transform margin (e.g. Aslanian *et al.* 2023). The second hypothesis is that the Continent Ocean Boundary is located on the western part of the Limpopo Corridor, immediately adjacent to the Limpopo Fault. In this case, the Limpopo Fault, a major strike-slip fault, serves as the eastern limit of the thinned continental crust of the Mozambique Coastal Plain and North Natal Valley, and is immediately adjacent to an extremely narrow transitional domain (~10–20 km) to the east, before reaching the thickened oceanic domain or the oceanic domain. In other words, the Limpopo Corridor is primarily composed of thickened oceanic crust. The crustal nature of the narrow transitional domain is unclear, which may have been affected and modified by intense faulting and magmatism. This hypothesis also indicates that the Limpopo Margin first underwent a NW–SE riffting, followed by a clockwise transition to N–S strike-slip motion during the post-rift, the transform stage (e.g. Watremez *et al.* 2021; Roche *et al.* 2021, 2022, 2023).

The transitional domain on profile DZ02 almost coincides with the Limpopo Corridor identified in previous studies (e.g. Moulin *et al.* 2020, 2023; Evain *et al.* 2021; Li *et al.* 2021; Fig. 1b and Fig. 5a). From section 5.1, it is clear that the 1-D velocity-depth profiles extracted from the transitional domain of the DZ02 profile fit well with the Limpopo Corridor crust derived from the PAMELA project, indicating that the velocity model of DZ02 in the Limpopo Corridor is consistent with previous results (Figs 11a and b; Moulin *et al.* 2020, 2023; Evain *et al.* 2021; Watremez *et al.* 2021). Moreover, the crustal thickness and Moho depth of this part constrained

by DZ02 and MZ5, two profiles located in proximity, are very similar (Fig. 1a; Watremez *et al.* 2021). The Continent Ocean Boundary defined by Watremez *et al.* (2021) was primarily supported by the identification of the Limpopo Transform Fault Zone, based on a set of commercial seismic reflection profiles that are coincident with or located near MZ4 and MZ5 (Roche *et al.* 2021). However, Evain *et al.* (2021) mainly identified the Mozambique Fracture Zone using a part of the same commercial seismic reflection profiles. In parallel, Li *et al.* (2021) also identified the Mozambique Fracture Zone using the academic seismic reflection profiles MZ5, S01 and MZ1, and defined it as the Continent Ocean Boundary (Fig. 1). The extreme velocity area in the uppermost crust of our DZ02 profile matches well the Mozambique Fracture Zone.

In addition, there are significant differences in crustal thickness and velocity structure between the transitional domain and the thickened oceanic domain, especially the different variational trends of the upper and lower crustal thicknesses and the sharp increase in velocity in the uppermost crust near the MFZ (Figs 5a and 10). As shown in Fig. 5(a), the 7 km s⁻¹ velocity contour of the lower crust in the transitional domain varies drastically, presumably due to strong magmatic activity within the thinned lower continental crust. In summary, although the crust of Limpopo Corridor exhibits a certain degree of oceanic crustal properties, it is fundamentally different from the thickened oceanic crust or oceanic crust. It originated from the continental crust, underwent strike-slip or highly oblique riffting, and was also subjected to intense magmatism, and ultimately formed into the anomalous or mixed crust (e.g. Evain
<figure><img src="image_10.png" /><figcaption>**Figure 9.** Checkerboard result showing the recovery of velocity anomaly checks that have dimensions of 15 km × 1 km with a maximum velocity perturbation of 10 per cent. (a) The input checkerboard pattern. (b) The recovered velocity anomalies. b.s.l.: below sea level.</figcaption></figure>

et al. 2021; Moulin et al. 2023). Therefore, we conclude that the Continent Ocean Boundary on the eastern edge of the Limpopo Margin is in the vicinity of the Mozambique Fracture Zone.

As for the early tectonic motions in this area, we cannot conclude on the basis of one profile whether the Limpopo Margin first experienced NW-SE oblique extension followed by a clockwise transition to N-S strike-slip motion (e.g. Watremez et al. 2021; Roche et al. 2021, 2022, 2023), or was only subjected to N-S strike-slip movement or highly oblique rifting (e.g. Moulin et al. 2020, 2023; Evain et al. 2021). This important issue for the geodynamic evolution of the area and Gondwana plate reconstructions is beyond the scope of this paper.

## 6.2 Difference in crustal properties between thickened oceanic domain and normal oceanic domain

To the east, the DZ02 profile crosses the Mozambique Fracture Zone to abut in the Mozambique Ocean Basin. The nearly W-E magnetic anomalies in the Mozambique Basin are offset by several N-S fractures, meaning that the crust becomes younger eastwards as it crosses these fracture zones (Fig. 1b; e.g. Mueller & Jokat 2017; Muller & Jokat 2019). As mentioned in Section 5, the portion of DZ02 in the Mozambique Basin is subdivided into the thickened and normal oceanic domains. The differences in the nature of the crust between the two domains are reflected both in the variation of crustal thickness and in the different velocity structures (Figs 5a and 10).

On the other hand, there are marked differences between the 1-D velocity-depth profiles of the two domains (Figs 11c and d).

Various factors may have been responsible for these differences. For example, the thickened oceanic domain is separated from the Limpopo Margin by the Mozambique Fracture Zone, and the strike-slip motion along the Limpopo Corridor might have affected the crustal structures and properties of both sides of the Mozambique Fracture Zone. The interaction between the thickened oceanic domain and its western transitional domain may lead to crustal contamination (e.g. Evain et al. 2021). This differs from the case of the Côte d'Ivoire-Ghana marginal ridge (Sage et al. 1997). The uplift of the uppermost crust in the thickened domain at distance ~160 km may have been caused by the strike-slip motion during the oceanic crust formation in the immediate vicinity of the Mozambique Fracture Zone. The peculiar thickness variations of the upper and lower crusts and the dramatic fluctuations in the uppermost crust velocity seem to be related to this interaction (Figs 5a and 10). In contrast, the oceanic domain far from the Mozambique Fracture Zone is unaffected by the strike-slip motion.

In addition to the gradual uplift of Moho and it's flattening eastward, another notable feature is that the velocity contours of the thickened oceanic domain rise towards the oceanic domain, indicating that the overall velocity of the latter is greater than that of the former. Also, high velocity lower crustal bodies are widely distributed in the oceanic domain (>7.4 km s⁻¹; Figs 5a, 10 and 11). It is possible that both the differences in crustal velocities and crustal
<figure><img src="image_11.png" /><figcaption>**Figure 10.** (a) The thickness of the full crust (red), upper crust (blue) and lower crust (green), using the 6.5 km s⁻¹ contour to roughly separate the upper and lower crustal layers. (b) Upper crustal velocity variation along profile at 0.5 and 1.0 km depth below the acoustic basement, lower crustal velocity variation along profile at 0.5, 1.0 and 2.0 km above the Moho. The blank regions on both ends of the profile with poor or no ray coverage are not included in the discussion.</figcaption></figure>

thickness between the two domains result from the same causes, related to the major strike-slip zone that is the Mozambique Fracture Zone. There may also be differences in magma supply between the two domains during the formation of the oceanic crust related to either the distance to the speculative Karoo plume, then located in the North Natal Valley (e.g. Cox 1989; White & McKenzie 1989) or to the younger age of the oceanic crust eastward along the profile (Fig. 1b).

Specifically, unlike passive margins dominated by divergent motion, the identification of the Continent Ocean Boundary at transform margins, which are characterized by strike-slip motion, primarily depends on the position of the fracture zone or transform faults, as there is typically a sharp contrast in crustal composition and velocity structure across the boundary (e.g. Sage et al. 2000; Basile 2015). Therefore, we can rule out the possibility that the observed crustal thinning from the thickened to the normal oceanic domain results from a decreasing contribution of inherited continental crust. Alternatively, the eastward thinning might reflect cooler or more refractory mantle conditions farther from the influence of the speculative plume (e.g. Cox 1989; White & McKenzie 1989); however, this interpretation conflicts with the higher lower-crustal velocities in the normal oceanic domain. To summarize, the differences between the thickened oceanic domain and the normal oceanic domain may primarily be attributed to the effects of strike-slip motion and crustal contamination along the Mozambique Fracture Zone

(e.g. Evain et al. 2021), and/or to spatial and temporal variations in magmatic supply, rather than to differences in the properties of the primary melts that formed the two domains. Of course, as the magma-rich transform margins are rarely documented (e.g. Basile 2015; Mercier de Lépinay et al. 2016), further observations and evidence are still required to thoroughly clarify this issue.

## 6.3 Magmatism and mantle temperature

The typical transform continental margin model is mainly derived from case studies of magma-poor transform margins, which are generally subjected to strike-slip or highly oblique rifting with little oblique extension. In this type of margins, the Continent Ocean Transition is pretty narrow, and the crustal structure and nature on both sides of the transform fault change abruptly, with a steep Moho in the transform zone and crustal thinning within a short horizontal distance (e.g. Basile 2015; Mercier de Lépinay et al. 2016). However, according to previous works, the southern Mozambique margin is a magma-rich transform continental margin (e.g. Li et al. 2021; Roche et al. 2021; Watremez et al. 2021). In addition, the southern Mozambique margin may not only have been subjected to strike-slip or highly oblique rifting, but may also have undergone NW-SE oblique extension before slowly transitioning to N-S strike-slip motion, as the central and northern Mozambique margins (e.g. Leinweber et al. 2013; Mueller et al. 2016; Mueller &
<figure><img src="image_12.png" /><figcaption>**Figure 11.** Comparison of the 1-D velocity-depth profiles extracted from the final P-waves velocity model for different domains along the DZ02 line with compilations from the literature, the oceanic crusts of similar ages in the Atlantic (White *et al.* 1992) and the crusts of Limpopo corridor (Moulin *et al.* 2020, 2023; Evain *et al.* 2021; Watremez *et al.* 2021). (a) Distribution of the 1-D velocity-depth profiles below the acoustic basement extracted and presented in panels (b), (c) and (d). The colour-code is according to the segmentation along DZ02. (b) Comparison of the 1-D velocity profiles extracted in the transitional domain with the compilation. (c) Comparison of the 1-D velocity profiles extracted in the thickened oceanic domain with the compilation. (d) Comparison of the 1-D velocity profiles extracted in the oceanic domain with the compilation. b.s.l.: below sea level, o.c.: oceanic crust.</figcaption></figure>

Jokat 2017; Senkans *et al.* 2019; Vormann *et al.* 2020; Vormann & Jokat 2021). In any case, these factors may have caused the southern Mozambique transform margin to differ from the typical or classical magma-poor transform margins in terms of crustal structures and properties, and one of the key factors is magmatism. However, the tectonic motion of the Mozambique margin is still a matter of debate (e.g. Evain *et al.* 2021; Moulin *et al.* 2023). The variability in crustal structures and properties of transform margins has been widely documented in various case studies, such as the St Paul Fracture Zone in the Equatorial Atlantic Ocean and the Mediterranean buffer zones (e.g. Pellen *et al.* 2016; Schnürle *et al.* 2023b).

The magmatism seems to have influenced the structure along the DZ02 profile (Fig. 5a). First, the Moho and crustal thickness change gently from the transitional domain to the thickened oceanic domain, and without any abrupt changes in the vicinity of the Mozambique Fracture Zone. Secondly, the magmatism and the strike-slip motion along the transform margin have profoundly influenced the

crustal structure and properties of the transitional domain and the thickened oceanic domain. The thickness of the upper crust in the transitional domain remains nearly uniform, whereas the uppermost crustal velocity increases sharply near the Mozambique Fracture Zone, and the lower crust seems to be intruded by magmatic activity. The irregular variations in thickness of the upper and lower crust and the uppermost crustal velocity in the thickened oceanic domain may be due to the strike-slip motion along the Limpopo Corridor and Mozambique Fracture Zone during the formation of the oceanic crust. There may be interaction and contamination between the thickened oceanic domain and the adjacent Limpopo Margin to the west (e.g. Evain *et al.* 2021). The crustal thickness of the oceanic domain is around 8 km, thicker than regular oceanic crust (6.1 ± 0.9 km; Christeson *et al.* 2019), suggesting an increased melt production possibly related to a thermal mantle anomaly of this region (e.g. Moulin *et al.* 2020, 2023; Evain *et al.* 2021; Watremez *et al.* 2021; Roche *et al.* 2022). The crustal velocity of the oceanic
domain is relatively high, especially in the lower crust (>7.4 km s⁻¹), likely reflecting a hotter mantle and producing MgO-rich melts (e.g. White & McKenzie 1989; Korenaga et al. 2002; Grevemeyer et al. 2021).

According to the empirical relationship between the mantle temperature at the upper mantle solidus and the crustal thickness (e.g. Klein & Langmuir 1987; Dhabaria & Singh 2024), eq. (1), we can estimate the upper mantle temperature of Mozambique Basin during the crust formation.

$$[(T_o - 1150)/12 - P_{min}]^2 = D_c * 2.6212 * [(T_o - 1150)/12]^{0.038} / (10.2 * 0.006), \quad (1)$$

where $T_o$ is the mantle temperature at upper mantle solidus in °C; $P_{min}$ is the minimum pressure of melting in kbar, equal to the pressure at the base of the crust when assuming the upwelling mantle accumulates at the base of the crust; $D_c$ is the crustal thickness in km. $P_{min}$ can be calculated using eq. (2):

$$P_{min} = \rho_{water}gh_{water} + \rho_{crust}gh_{crust}. \quad (2)$$

The pressure unit should be converted from N m⁻² to kbar (1 bar ≡ 1020 g cm⁻²) when substituting eq. (2) in eq. (1). The left-hand side of eq. (1) is set to $L(T)$ to get eq. (3), and the right side is set to $R(T)$ to get eq. (4):

$$L(T) = [(T - 1150)/12 - P_{min}]^2, \quad (3)$$

$$R(T) = D_c * 2.6212 * [(T - 1150)/12]^{0.038} / (10.2 * 0.006). \quad (4)$$

We can numerically solve eq. (1) from the intersection of the $L(T)$ and $R(T)$, that is, $L(T) = R(T)$. Using an average water and crustal density of 1.025 and 2.86 g cm⁻³ (e.g. Carlson & Herrick 1990), we can compute the minimum pressure beneath the mature crust (6.1 ± 0.9 km; Christeson et al. 2019) and the crust affected by a thermal mantle anomaly from Mozambique basin underlying a 5 km thick water column. Plotting eq. (1) using these values of minimum pressure gives us the upper mantle temperature of 1383 and 1418 °C for the crustal thickness of 6.1 and 8 km, respectively, indicating an increase in mantle temperature by ~35 °C in the area affected by a thermal mantle anomaly (Fig. 12). It is comparable to that of the crystallization temperatures (1400–1500 °C) of the most Mg-rich magmas of the Karoo large igneous province based on Al-in-olivine thermometry (e.g. Heinonen et al. 2015). Naturally, it is temporally more appropriate to compare our calculations with the magmas of the Bombeni–Movene event (~140 Ma) than with those of the Karoo event (~183 Ma, e.g. Schnürle et al. 2023a). However, there is currently a lack of reported crystallisation temperatures for the former. This increase in mantle temperature increases the melt fraction in the ascending mantle column from 10.3 per cent for the mature crust to 11.8 per cent for the crust of the Mozambique basin affected by a thermal mantle anomaly (e.g. Klein & Langmuir 1987; Dhabaria & Singh 2024).

## 7 CONCLUSIONS

This study focuses on the latest wide-angle seismic profile DZ02 acquired on the southern Mozambique margin, which extends from the eastern part of the North Natal Valley in the west, through the Limpopo Corridor, and deeper into the Mozambique Basin to the east. The P-wave velocity model of DZ02 profile was obtained by picking and fitting the seismic phases of all stations. Based on our results and analysis, as well as a large number of previous works in this study area, we conclude that:

<figure><img src="image_13.png" /><figcaption>**Figure 12.** Variation of temperature functionals $F(T)$ following the empirical relationship of pressure and upper mantle solidus temperature with crustal thickness (Klein & Langmuir 1987). The temperature functions of the left $L(T)$ (solid line) and right $R(T)$ (dashed line) side of eq. (1) are plotted for the minimum pressure threshold at the base of extremely thin crust from Equatorial Atlantic Ocean (EAO, red, 4 km; Dhabaria & Singh 2024), mature crust from global compilations (global, black, 6.1 km; Christeson et al. 2019), crust affected by a thermal mantle anomaly from Mozambique Basin (MB, 8 km, blue; this study), giving the estimated upper mantle temperatures of 1333, 1383 and 1418 °C, respectively.</figcaption></figure>

(1) The profile can be clearly divided into three domains based on differences in crustal thickness and velocity structure: (a) a westernmost transitional domain with anomalous or mixed crust, bounded by the Mozambique Fracture Zone to the east; (b) a central thickened oceanic domain resulting from enhanced magmatism; and (c) an easternmost normal oceanic domain with a crustal thickness of ~8 km and a high velocity lower crust (up to 7.4 km s⁻¹).

(2) The Limpopo Corridor, which is bounded by the Limpopo Fault to the west and the Mozambique Fracture Zone to the east, originated from the thinned continental crust of the Limpopo Margin. It was subjected to strike-slip motion along the Mozambique Fracture Zone during the rifting, break-up and seafloor spreading of Gondwana, accompanied by the intense magmatic intrusion and underplating, and eventually formed anomalous or mixed crust. The Continent Ocean Boundary between the Limpopo Margin and Mozambique Basin is close to the Mozambique Fracture Zone.

(3) The upper and lower crustal thicknesses and the uppermost crustal velocity of the thickened oceanic domain vary dramatically, and the overall crustal velocity is significantly lower than that of the oceanic domain to the east. These results indicate that the formation of the thickened oceanic domain was also influenced by strike-slip motion along the Mozambique Fracture Zone, and there may have been interaction with and contamination from the western transition domain.

(4) The crustal thickness of the oceanic domain remains uniform. The differences in crustal thickness and velocity between this domain and the thickened oceanic domain may be related to its greater distance to the transform margin strike-slip zone and the speculative Karoo plume, and temporal variation in magma supply.

(5) The oceanic domain with a crustal thickness of ~8 km and a high-velocity lower crust (up to 7.4 km s⁻¹) is suggestive of a hotter mantle (upper mantle temperature: ~1418 °C) that produced
enhanced MgO-rich melts (melt fraction: ~11.8 per cent) probably
due to the influence of a thermal mantle anomaly.

DATA AVAILABILITY

Bathymetric data are available at http://www.gebco.net. Gravity data are available at https://doi.org/10.5194/essd-11-647-2019. The data including original seismic images and traveltime picks used in this study are available at https://doi.org/10.6084/m9 figshare. 28120961. Figures in this manuscript were plotted using GMT (Wes- sel et al. 2013) or Seismic Unix (Stockwell 1999). The traveltime tomography code used in this work is propriety code but will be available on request from S. C. Singh (singh@ipgp.fr).

ACKNOWLEDGMENTS

We thank the captain, crew and scientific team of the second China-
Mozambique Joint Cruise onboard R/V ‘Dayang Hao’ for data col-
lection. W. Wang’s contributions to this work were supported by
the National Natural Science Foundation of China (Grant Nos.
42376052 and 42076047), the National Key Research and De-
velopment Program of China (Grant Nos. 2023YFE0126100 and
2023YFC2808803), the Leading Talents in Scientific and Techno-
logical Innovation Project of Zhejiang Province (No. 2021R52058),
the China Scholarship Council (No. 202006230254). Z. Wang’s
contributions to this study were supported by the Royal Society-
Newton International Fellowship 2022 (NIF\R1\221849). WW
thanks Louise Watremez for her support and assistance during his
visiting in France. We thank Daniel Aslanian and another anony-
mous reviewer, as well as the editor and the assistant editor, for their
constructive comments that helped to improve the manuscript.

SUPPORTING INFORMATION

Supplementary data are available at *GJIRAS* online.

**Figure S1.** Migrated image in depth domain along the seismic profile DZ02. The locations of the OBSs are marked by red inverted triangles. The blue dots indicate the manually picked seafloor. The red dots indicate the manually picked acoustic basement. The green line represents the acoustic basement obtained by tomography in this study. The multichannel seismic data were processed by the China National Offshore Oil Corporation using the commercial software CGG (e.g. Li *et al.* 2021). All major processing steps were applied in the pre-stack domain, including amplitude compensation, static correction, bandpass frequency filtering, gain and mute analysis, predictive deconvolution, multiple suppression using Surface-Related Multiple Elimination and a filter in the Radon domain, velocity analysis and residual static corrections. Pre-stack Kirchhoff depth migration was performed on the common-offset gathers.

**Figure S2.** (a) Histogram showing the initial residuals of Ps, PbP, Pg and PmP arrivals. (b) Histogram showing the final residuals of Ps, PbP, Pg and PmP arrivals. (c) Normalized traveltime residuals as a function of offset before (blue dots) and after tomography (red dots).

Please note: Oxford University Press is not responsible for the
content or functionality of any supporting materials supplied by
the authors. Any queries (other than missing material) should be
directed to the corresponding author for the paper.

CONFLICT OF INTEREST

The authors declare that they have no conflict of interest.

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