# Seismic Evidence for Velocity Heterogeneity Along ~40 Ma Old Oceanic Crustal Segment Formed at the Slow-Spreading Equatorial Mid-Atlantic Ridge From Full Waveform Inversion of Ocean Bottom Seismometer Data

Peng Guo, Satish C. Singh

► To cite this version:

Peng Guo, Satish C. Singh. Seismic Evidence for Velocity Heterogeneity Along ~40 Ma Old Oceanic Crustal Segment Formed at the Slow-Spreading Equatorial Mid-Atlantic Ridge From Full Waveform Inversion of Ocean Bottom Seismometer Data. Journal of Geophysical Research: Solid Earth, 2024, 129, 10.1029/2024JB029776. insu-05081559

HAL Id: insu-05081559

https://insu.hal.science/insu-05081559v1

Submitted on 24 May 2025

HAL is a multi-disciplinary open access
archive for the deposit and dissemination of sci-
entific research documents, whether they are pub-
lished or not. The documents may come from
teaching and research institutions in France or
abroad, or from public or private research centers.

L'archive ouverte pluridisciplinaire HAL, est
destinée au dépôt et à la diffusion de documents
scientifiques de niveau recherche, publiés ou non,
émanant des établissements d'enseignement et de
recherche français ou étrangers, des laboratoires
publics ou privés.

<figure><img src="image_1.png" /></figure>

Distributed under a Creative Commons Attribution 4.0 International License
# JGR Solid Earth

## RESEARCH ARTICLE

10.1029/2024JB029776

**Key Points:**

* We apply full waveform inversion to the crustal turning waves recorded by ocean bottom seismometers in the equatorial Atlantic Ocean

* The velocity model exhibits strong crustal velocity heterogeneity, containing four distinct segments correlating well with seafloor morphology

* The strong crustal heterogeneity seems to be caused by a combination of magmatic and tectonic processes

**Supporting Information:**

Supporting Information may be found in the online version of this article.

**Correspondence to:**

P. Guo,
peng.guo@csiro.au;
pengguo.geos@gmail.com

**Citation:**

Guo, P., & Singh, S. C. (2024). Seismic evidence for velocity heterogeneity along ~40 Ma old oceanic crustal segment formed at the slow-spreading equatorial Mid-Atlantic Ridge from full waveform inversion of ocean bottom seismometer data. *Journal of Geophysical Research: Solid Earth, 129*, e2024JB029776. https://doi.org/10.1029/2024JB029776

Received 25 JUN 2024
Accepted 10 OCT 2024

**Author Contributions:**

**Conceptualization:** Peng Guo, Satish C. Singh

**Data curation:** Peng Guo, Satish C. Singh

**Formal analysis:** Peng Guo, Satish C. Singh

**Funding acquisition:** Satish C. Singh

**Investigation:** Peng Guo

**Methodology:** Peng Guo, Satish C. Singh

**Project administration:** Satish C. Singh

**Resources:** Satish C. Singh

**Software:** Peng Guo

**Supervision:** Satish C. Singh

**Validation:** Peng Guo

**Visualization:** Peng Guo

© 2024. The Author(s).
This is an open access article under the terms of the Creative Commons Attribution License, which permits use, distribution and reproduction in any medium, provided the original work is properly cited.

<figure><img src="image_2.png" /></figure>

# Seismic Evidence for Velocity Heterogeneity Along ~40 Ma Old Oceanic Crustal Segment Formed at the Slow-Spreading Equatorial Mid-Atlantic Ridge From Full Waveform Inversion of Ocean Bottom Seismometer Data

Peng Guo¹,² ID and Satish C. Singh³ ID

¹Deep Earth Imaging Future Science Platform, The Commonwealth Scientific and Industrial Research Organisation (CSIRO), Kensington, WA, Australia, ²Energy Research Unit, The Commonwealth Scientific and Industrial Research Organisation (CSIRO), Kensington, WA, Australia, ³Institut de Physique du Globe de Paris, Université Paris Cité, CNRS, Paris, France

**Abstract** In slow spreading environments, oceanic crust is formed by a combination of magmatic and tectonic processes. Using full waveform inversion applied to active-source ocean bottom seismometer data, we reveal the presence of a strong lateral variability in the 40–48 Ma old oceanic crust formed at the slow-spreading Mid-Atlantic Ridge in the equatorial Atlantic Ocean. Over a 120 km-long section between the St Paul fracture zone (FZ) and the Romanche transform fault (TF), we observe four distinct 20–30 km long crustal segments. The segment affected by the St Paul FZ consists of three layers, an ~2 km thick layer with a P-wave velocity <6 km/s, a 1.5 km thick middle crust with a velocity of 6–6.5 km/s, and an underlying layer where velocity is ~7 km/s, representing the lower crust. The segment associated with an abyssal hill morphology contains a high velocity of ~7 km/s at 2–2.5 km below the basement, indicating the presence of primitive gabbro or serpenized peridotite. The segment associated with a low basement morphology seems to have 5.5–6.5 km/s velocity starting near the basement extending down to ~4 km depth, indicating chemically distinct crust. The segment close to the Romanche TF, a velocity 4.5–5 km/s near the seafloor increasing to 7 km/s at 4 km depth indicates a magmatic origin. The four distinct crustal segments have a good correlation with the overlying seafloor morphology. These observed strong crustal heterogeneities could result from alternate tectonic and magmatic processes along the ridge axis, possibly modulated by thermal and/or chemical variations in the mantle during their formation along the ridge segment.

**Plain Language Summary** Generally magmatically accreted oceanic crust has a nearly uniform crustal structure. The speed of seismic waves, which is widely used as a proxy for the physical properties of the Earth, is relatively low in the upper crust but increases rapidly with depth, while in the lower crust the velocity is high but increases more slowly. Nevertheless, oceanic crust formed at slow-spreading mid-ocean ridges can be very heterogeneous. However, it is challenging to quantify the nature of this heterogeneity using conventional travel time based analyses. Here we apply a cutting-edge full waveform inversion technique to an active-source ocean bottom seismometer data set from the equatorial Atlantic Ocean. The aim is to create seismic velocity models of oceanic crust with enhanced details. The waveform inversion reveals four distinct zones within a 120 km long crustal section, with boundaries consistent with the seafloor morphology. The section contains two segments of normal oceanic crust, a segment with high velocities at shallow depth, possibly containing primitive gabbro, and another chemically distinct segment with nearly constant velocity in the whole crust. The observed strong crustal heterogeneities seem to result from a combination of magmatic and tectonic accretion processes possibly induced by thermally or chemically heterogeneous mantle.

## 1. Introduction

Oceanic crust is formed at the Mid-Ocean Ridges (MORs) from basaltic melt derived from the decompression of upwelling mantle beneath the ridges. Insights gained from studies of ophiolites (e.g., Boudier et al., 1996; Kelemen et al., 1997, 2020; Nicolas et al., 1988; Salisbury & Christensen, 1978), ocean drilling results (e.g., Alt et al., 1996; Iturrino et al., 2002; Carlson & Jay Miller, 2004; Miller & Christensen, 1997; Swift et al., 2008; Wilson et al., 2006) and geophysical studies (e.g., Canales et al., 2005; Carbotte et al., 2013; Christeson et al., 2019; Grevemeyer, Ranero, & Ivandic, 2018; Guo et al., 2022; Harding et al., 1993; Singh et al., 2006;
Writing – original draft: Peng Guo

Writing – review & editing: Peng Guo,

Satish C. Singh

Spudich & Orcutt, 1980; Toomey et al., 1994; Vaddineni et al., 2021; White et al., 1992) suggest that oceanic crust formed by magmatic process can be divided into two distinct layers, an upper (layer 2) and a lower crust (layer 3), with the sediments above the basement being layer 1. The upper crust is primarily composed of pillow lavas (layer 2A) and the underlying sheeted dikes (layer 2B), and is characterized by a high seismic velocity gradient (1–2 s⁻¹) where the velocity increases from ~3.5 to ~4.5 km/s at top of the layer to 6–6.5 km/s at the bottom. The lower oceanic crust contains mainly gabbroic rocks where the velocity increases more slowly with depth (~0.1 s⁻¹) from ~6.5 km/s to ~7 km/s (Christeson et al., 2019). The thickness of the lower crust is usually about twice of that of the upper crust.

However, the crustal formation mechanisms can vary with the spreading rate (Carbotte et al., 2016; Morgan & Chen, 1993). At fast and intermediate spreading ridges, magmatic processes dominate resulting in a fairly uniform crustal thickness and structure (Carbotte et al., 2013; Han et al., 2016; Kent et al., 1994; Toomey et al., 1994). At slow spreading ridges, magma supply could vary spatially and temporally (Cannat, 1993; Lin & Morgan, 1992), leading to a heterogeneous crust (Canales et al., 2000; Davy et al., 2020; Detrick et al., 1995; Dunn et al., 2017; Escartín et al., 2008; Hooft et al., 2000). For example, when the magma supply is relatively low, tectonic processes would dominate (Cannat, 1993), leading to the development of oceanic core complexes (OCCs) (Cannat, 1993; Dick et al., 2008; Grevemeyer, Hayman, et al., 2018; MacLeod et al., 2002), which are formed through the exhumation of lower crustal gabbroic and mantle ultramafic rocks at shallow depth and on the seafloor along long-lived detachment faults (Blackman & Collins, 2010; Canales, 2010; Cann et al., 1997; Cannat et al., 1997; Dick et al., 2000; Xu et al., 2020). Buck et al. (2005) showed that long-lived OCCs develop when there is enough melt, 0.5 fraction of crustal thickness, weakening the lithosphere, enabling detachment faulting and exhuming deep-rooted rocks.

Many recent seismic surveys have revealed heterogeneous crustal structures in the Atlantic Ocean, for example, at the Rainbow area of the Mid-Atlantic Ridge (MAR) (Dunn et al., 2017), the Kane OCC (Canales, 2010) and southern flank of the Romanche oceanic transform fault (TF) (Gregory et al., 2021; Wang & Singh, 2022), which usually coincide with extensive faulting, long-lived detachment faults or the presence of corrugated massifs associated with OCCs. Thus, the seismic velocity heterogeneity could result from a combination of different factors, such as initial composition of melt, faulting and subsequent alteration. For example, using travel-time tomography in the 60–75 Million years (Ma) old central Atlantic Ocean, Davy et al. (2020) observed both normal oceanic crust and tectonically dominated crust across five ridge segments separated by a TF and three nontransform offsets. In a similar region, Allen et al. (2022) found two phases of hydrous alteration during crustal accretion at the MAR and suggested that the nature of a crustal segment varies over time, indicating that there could be both spatial and time variation along a ridge segment. On the other hand, Wang and Singh (2022) found dominantly magmatically accreted crust along 850 km profile covering 8–70 Ma old lithosphere in the equatorial Atlantic region. A review of oceanic crustal models (Christeson et al., 2019) suggests that crustal thicknesses and velocity profiles have the greatest standard deviations for the slow-spreading crust, which is attributed to the heterogeneity in the crustal accretion process. Furthermore, hydrothermal alteration can also change rock properties of oceanic crust (Alt et al., 1996; Audhkhasi & Singh, 2019; Boulahanis et al., 2022; Carlson, 2011; Christeson et al., 2007). Seismic surveys and oceanic drilling have discovered low velocity anomalies above the roof of axial melt lens (Singh et al., 1999) and at the dike-gabbro transition (Swift et al., 2008), which are interpreted to be due to hydrothermal alteration. During a tectonically dominated crustal accretion, hydrothermal circulation can lead to hydration of crustal rocks and serpentinization of mantle olivine-rich peridotite (Minshull et al., 1998), leading to a strong heterogeneity in the crust.

Most of the studies of oceanic crust are based on travel-time tomography of crustal and mantle arrivals (e.g., Canales et al., 2005; Escartín et al., 2017; Gregory et al., 2021; Grevemeyer, Hayman, et al., 2018; Korenaga et al., 2000; Minshull et al., 1991; Toomey et al., 1994; White et al., 1992). This technique is based on a high-frequency approximation and ray assumption for computing travel times of seismic arrivals (Van Avendonk et al., 1998), resulting in subsurface velocity models that contain mainly large-scale structures. On the other hand, full waveform inversion (FWI) (Shipp & Singh, 2002; Tarantola, 1984; Virieux & Operto, 2009) that utilizes complete waveform has the potential for inferring crustal models with enhanced accuracy and higher resolution (e.g., Arnoux et al., 2017; Arnulf et al., 2021; Canales, 2010; Christeson et al., 2012; Davy et al., 2021; Górszczyk et al., 2017; Guo et al., 2022; Qin & Singh, 2017; Singh et al., 1999). Here, we apply elastic FWI to the crustal turning arrivals of an active-source OBS data set for estimating a high-resolution P-wave velocity model of the oceanic crust in the equatorial Atlantic Ocean.
<figure><img src="image_3.png" /></figure>

**Figure 1.** Map of the study area and active-source seismic survey in the equatorial Atlantic Ocean. (a) The equatorial Atlantic Ocean, with the double red lines indicating the Mid-Atlantic Ridge and the dashed black lines showing the St Paul, the Romanche and the Chain transform faults (TFs) and FZs from north to south. The blue rectangle shows the map area in panel (b). (b) Bathymetry of the ocean bottom seismic survey (zoom-in of the blue rectangle in panel (a)) with the major oceanic TFs (dashed white lines) and FZs (solid white lines). The black-white beachball indicates the 2016 Mw7.1 Romanche earthquake (Hicks et al., 2020). The black circles show the locations of the ocean bottom seismometers (OBSs) with every 10 OBSs plotted for a total of 50 OBSs from the ILAB-SPARC experiment, and the pink line shows the shot profile for the survey (Wang & Singh, 2022). The light gray lines suggest plate ages labeled every 10 Myr (Müller et al., 2008). (c) Close-up of the area (the dashed rectangle in panel (b)) for the 9 OBSs (12–20) used in this study. The bathymetry was plotted by superimposing the high-resolution bathymetry data collected during the ILAB-SPARC seismic experiment on a relatively low-resolution global bathymetry grid. The pink line passing through the OBSs shows the shot profiles used for inversion. The plate ages were plotted using the light gray lines and labeled every 5 Myr (Müller et al., 2008). The distance along the profile is also marked.

## 2. Study Area

Our study area lies in the equatorial Atlantic Ocean where the slow spreading MAR is offset by ~1,800 km by a set of closely spaced, large-offset (>300 km) transform faults (TFs) (Bonatti et al., 1994), from north to south encompassing the St. Paul, the Romanche, and the Chain TFs, with cumulative age variation of 90 Ma over a distance of 400 km (Figures 1a and 1b). Large-offset TFs can have profound effects on the thermal structure and the dynamics of the mantle upwelling, and hence the melt budget and chemical composition of melt (Bonatti et al., 1994). The St. Paul TF, which contains four strike-slip faults and three short intra-transform ridge segments, offsets the MAR by ~600 km. The sub-areal St. Paul Island consists of a transpressional ridge containing peridotite with various degrees of serpentinization and deformation (Maia et al., 2016). The Romanche TF offsets the MAR by an offset ~880 km, the largest offset on Earth with a 45 Ma lithospheric age contrast. It exhibits a ~20 km wide transform valley and a thick hanging sedimentary basin within the north flank. Seafloor dredgings at the Romanche TF show an extensive presence of basalt and gabbro on the north flank, but predominantly mantle-derived serpentinized and mylonitic peridotites on the valley floor and in the south of the fault zone (Bonatti, 1968; Bonatti & Honnorez, 1976; Seyler & Bonatti, 1997). This observation is consistent with the heterogeneous tomographic velocity model obtained across the Romanche transform valley, which is characterized by a mafic crustal structure of average thickness ~6 km in the north of the TF and a magma-reduced crust with variable thicknesses of 4–6 km in the south (Gregory et al., 2021).

Recent travel-time tomographic studies using wide-angle ocean bottom seismometer (OBS) data have revealed a nearly constant crustal thickness of 5.6–6 km beneath the St Paul fracture zone (FZ) (Growe et al., 2021), Chain FZ (Marjanović et al., 2020) and the Romanche TF (Gregory et al., 2021), resembling a normal oceanic crust.
<figure><img src="image_4.png" /><figcaption>**Figure 2.** Observed data from ocean bottom seismometer 14 after data processing. The blue, red and cyan lines refer to the Pg (crustal turning waves), PmP (Moho reflection) and Pn (Mantle refraction) arrivals. The time axis is reduced with a reduction velocity of 7 km/s.</figcaption></figure>

# 3. Data and Methods

## 3.1. The Active-Source Wide-Angle Ocean Bottom Seismic Data and Processing

The field data was acquired during the ILAB-SPARC cruise in October-November 2018 onboard the French R/V *Pourquoi Pas*? A total of 50 OBSs were deployed along a 700 km long section across the St Paul FZ, the Romanche TF and the Chain FZ from north to south, with an average OBS spacing of 14.2 km to record the wide-angle refraction and reflection data (Figure 1b). The shooting line was extended by 50 km in the south, and 100 km in the north, for a total of 850 km long. All OBSs were equipped with a hydrophone (measuring pressure) and three geophones (measuring vertical and horizontal displacements), with natural frequency of 1 and 4.5 Hz, respectively. The sampling frequency was 250 Hz. The seismic source used in the survey comprised of two sub-arrays with 8 guns on each array, producing a total volume of 4,990 cubic inch, which was fired every ~300 m along the profile.

In this study we used a portion of these data (9 OBSs, the hydrophone component) between the St. Paul FZ and the Romanche TF, covering 40–48 Ma old oceanic crust along a 120 km long section (Figure 1c). We limited the data processing to a minimum prior to FWI to preserve the original waveform information. A predictive gapped deconvolution (gapped Wiener filtering) was applied to the OBS data to suppress the bubble effects (repetitive events) from the air-gun sources while keeping the waveform of the primary arrivals unchanged. We set the prediction gap (repetition time) to 0.2 s and the filter length to 0.8 s. Then we applied a zero-phase bandpass filtering of 3–10 Hz to the processed OBS data. We used zero-phase filtering to preserve the waveform and avoid phase shift (advances or delays in the signal). We also performed 3D to 2D transformation (Forbriger et al., 2014; Pica et al., 1990) to the data, including multiplying the amplitudes of field data by $\sqrt{t}$, where t is the two-way travel-time, and convolving with $1/\sqrt{t}$, because 2D elastic wave equation modeling is used for simulating seismic data recorded in a 3D Earth. Clear crustal turning waves (Pg) with high signal-to-noise ratio can be observed in Figure 2, starting from offsets (source-receiver distance) ± 6–8 km, up to offsets of ±20–25 km, followed by wide-angle Moho reflection (PmP) and upper mantle refraction (Pn) at the farther offsets. We observe large variations in the amplitude-versus-offset behavior, suggesting strong heterogeneities in the oceanic crust.

## 3.2. Travel Time Tomography

Travel time tomography was used to estimate seismic velocity model of the oceanic crust, using crustal turning waves (Pg) and the wide-angle Moho reflection (PmP) from the boundary between crust and mantle. The travel times of the Pg and PmP arrivals were hand picked on the OBS data after bandpass filtering of 4–20 Hz, with picking uncertainties of 30–50 ms and 50–70 ms, respectively (Gregory et al., 2021; Wang & Singh, 2022). We used the velocity model from Wang and Singh (2022) as a starting velocity model for FWI.
### 3.3. FWI

FWI is the current state-of-the-art technique for obtaining high-resolution quantitative subsurface models from seismic data (Shipp & Singh, 2002; Tarantola, 1984; Tromp et al., 2005; Virieux & Operto, 2009). Unlike travel-time tomography which only uses the travel time information (Van Avendonk et al., 1998), the FWI is based on minimizing the waveform difference between the observed and synthetically computed seismic data, incorporating the full physics of seismic wave propagation in the solid Earth (Shipp & Singh, 2002). Resolution studies suggest that the travel-time data provide information on a scale greater than approximately five times the dominant wavelength (Jannane et al., 1989). For FWI, the vertical resolution is usually between a quarter of wavelength to a wavelength (Jannane et al., 1989; Neves & Singh, 1996), while the horizontal resolution is primarily determined by the Fresnel zone. The OBS spacing also plays an important role (Jian et al., 2021). As a rule of thumb, the lateral resolution is about half of the OBS spacing (Dhabaria & Singh, 2024), for example, ~7 km in our study area.

We adopt a multi-stage strategy for inverting the Pg arrivals in the OBS data for obtaining the oceanic crustal model. In the first stage of FWI, where the tomographic model is used as the starting model, a trace-normalized FWI (Guo et al., 2024; Shen, 2010) is applied, with the misfit (waveform differences) function (J) defined as

$$J = \sum_{i=1}^{\text{Ns}} \sum_{j=1}^{\text{Nr}} \left\| \frac{\mathbf{s}_{i,j}}{\|\mathbf{s}_{i,j}\|} - \frac{\mathbf{d}_{i,j}}{\|\mathbf{d}_{i,j}\|} \right\|^2, \quad (1)$$

where $\mathbf{s}_{i,j}$ and $\mathbf{d}_{i,j}$ are seismic traces (1-D time-series vectors) from the synthetic and field data respectively, $\|\|$ is the l-2 norm, Ns and Nr are the number of sources and receivers, i and j are the indexes for the sources and receivers, respectively.

The gradient of the misfit function J with respect to model parameter vector m is

$$\frac{\partial J}{\partial \mathbf{m}} = \sum_{i=1}^{\text{Ns}} \sum_{j=1}^{\text{Nr}} \frac{\partial \mathbf{s}_{i,j}}{\partial \mathbf{m}} \frac{\partial J}{\partial \mathbf{s}_{i,j}}, \quad (2)$$

which can be efficiently calculated using the adjoint-state method (Tarantola, 1984; Tromp et al., 2005).

The adjoint source for the misfit defined in Equation 1, which is the data residual source for wavefield propagated backwards in time, is

$$\frac{\partial J}{\partial \mathbf{s}_{i,j}} = \left( \frac{\delta \mathbf{d}_{i,j}}{\|\mathbf{s}_{i,j}\|} \right) - \left( \frac{\delta \mathbf{d}_{i,j}^T \cdot \mathbf{d}_{i,j}}{\|\mathbf{s}_{i,j}\|^2 \|\mathbf{s}_{i,j}\|} \frac{\mathbf{s}_{i,j}}{\|\mathbf{s}_{i,j}\|} \right), \quad (3)$$

where $\delta \mathbf{d}_{i,j} = \frac{\mathbf{s}_{i,j}}{\|\mathbf{s}_{i,j}\|} - \frac{\mathbf{d}_{i,j}}{\|\mathbf{d}_{i,j}\|}$.

This intermediate step is used to bridge the gap between travel-time tomography and the classic FWI using “true amplitude” waveforms: the influence of amplitude-versus-offset is eliminated in the misfit function and the adjoint source through the trace-by-trace normalization, while the amplitude variation along the time axis within each trace remains. This ensures that the waveform inversion will first focus more on the phase comparison before moving to fit the waveforms and amplitudes. In addition, an important aspect of the trace-normalized inversion involves determining the amplitude normalization factors between the synthetic and real data for the OBS gathers as each OBS may have different local amplification due to site or instrument effects.

In the second stage, the output velocity model from the trace-normalized FWI is used as the starting model. We perform the classic true-amplitude FWI (Shipp & Singh, 2002; Tarantola, 1984), with the misfit function defined as

$$J = \sum_{i=1}^{\text{Ns}} \sum_{j=1}^{\text{Nr}} \|\mathbf{s}_{i,j} - \mathbf{d}_{i,j}\|^2, \quad (4)$$
and the adjoint source as

$$ \frac{\partial J}{\partial \mathbf{s}_{i,j}} = (\mathbf{s}_{i,j} - \mathbf{d}_{i,j}). \qquad (5) $$

The waveform misfit functions of FWI in Equations 1 and 4 contains the waveform differences of the observed and synthetically computed crustal around the P-wave turning waves. Despite the different misfit functions, the only difference in the two FWI workflows is the adjoint source (Tromp et al., 2005). For the P-wave velocity, the gradient of the misfit function (Equation 2) is (Shipp & Singh, 2002)

$$ \frac{\partial J}{\partial \mathbf{v}_p} = \frac{\rho v_p}{2(\lambda + \mu)^2} \sum_{N_s} (\sigma_{xx} + \sigma_{zz})(\tau_{xx} + \tau_{zz}), \qquad (6) $$

where $v_p$ is the P-wave velocity, $\rho$ is the density, $\lambda$ and $\mu$ are the Lamé parameters. $\sigma_{xx}$ and $\sigma_{zz}$ are the normal stress components of the source wavefield from forward-time propagation by solving the elastic wave equation in the stress and particle-velocity formulation (Virieux, 1986) with the source wavelet, and $\tau_{xx}$ and $\tau_{zz}$ are the normal stress components of the adjoint wavefield from backward-time propagation using the adjoint source. The implementation of the cross-correlation operation in Equation 6 requires simultaneous access to the stress wavefields at the same propagation time steps between the forward-time source and backward-time adjoint wavefields. We used the boundary value wavefield reconstruction method (Nguyen & McMechan, 2015) to avoid storing the entire history (time-space) of the source wavefield. The price is one additional wavefield propagation simulation that is feasible with modern computing facilities.

The S-wave velocity model was derived from the P-wave velocity using Equation 6 in Brocher (2005). The Brocher (2005)'s model slightly overestimates S-wave velocity from that of P-wave for oceanic crustal lithologies (Christensen, 1996). Considering that we only used Pg arrivals to invert for the P-wave velocity and the impact of S-wave velocity on the Pg waveform is small (Guo et al., 2021), the influence of the S-wave velocity on the estimated P-wave velocity model is minimal. However, the elastic modeling used here properly takes the effect of full elastic reflection/transmission partitioning at the sediment/basement interface and accurately models the amplitudes (Arnulf et al., 2014; Marjanović et al., 2019). Density was linked to the P-wave velocity based on the empirical relation of Hamilton (1978) for velocity smaller than 2.2 km/s, and the Gardner et al. (1974)'s relation for higher velocities. As the oceanic crust is 40–50 Myr old, the effect of attenuation would be small and therefore, we didn't include attenuation during the inversion, especially as the narrow frequency band (3–10 Hz) is used here. We used the time-domain staggered-grid finite-difference method (Virieux, 1986) with fourth-order spatial and second-order temporal accuracy for solving the elastic wave equation. The grid spacing for the finite-difference was 20 m in both the horizontal (distance) and vertical (depth) directions. The time step was 0.0012 s. We applied the convolutional perfectly matched layer absorbing boundary condition (Komatitsch & Martin, 2007) at the model boundaries. Considering the sparse distribution of OBSs, we applied a Gaussian smoothing operator to the gradient model for regularization. The vertical smoothing length was 0.4 km, about a quarter of the dominant wavelength. The lateral resolution was about half of the OBS spacing (Dhabaria & Singh, 2024). After extensive tests, we set the horizontal smoothing length to 3 km. The model update was further conditioned using the conjugate-gradient method for defining the search direction for minimizing the misfit functions. The step length for model update was estimated from a line search along the gradient direction using a parabolic fit (Vigh & Starr, 2008).

## 3.4. Data Selection

For the inversion, we used the pressure component of the crustal turning waves (Pg), because the Pg arrivals have the most linear behavior with regard to the P-wave velocities of the oceanic crust. Moho reflections (PmP) were not used in the FWI for constraining the velocity model, because of its strong nonlinearity around the critical angles (Guo et al., 2021). A time-window of 0.8 s was applied to the OBS gathers, by muting the data before 0.3 s and after 0.5 s of the picked Pg travel times, to reduce the influence of noise and to isolate the Pg arrivals from the other seismic events. Throughout the FWI, we excluded the Pg arrivals with offsets smaller than 5 km because of the potential interferences with strong seafloor reflections.
<figure><img src="image_5.png" /><figcaption>**Figure 3.** Source wavelet estimation from near-offset Pg arrivals and free-surface multiples. (a) Shows the near-offset Pg arrivals before (dashed gray lines) and after (solid black lines) aligning, in the frequency range 3–10 Hz. Solid green curve shows the stacking result. (b) Shows the free-surface multiples between −1 and 1 km offsets before (dashed gray lines) and after (solid black lines) alignment, in the frequency range 3–40 Hz. Solid green curve shows the stacking result.</figcaption></figure>

The inversion was carried out by first inverting the near-to-intermediate source-receiver offset data (from ±5–8 km up to ±15 km offsets) followed by including the Pg waveforms from the farther offsets (up to ±20–25 km offsets). The workflow effectively updates the P-wave crustal model from shallower to deeper depth, because turning waves of smaller offsets propagate in the shallower oceanic crust. The updated velocity model from the prior stage was used as the starting model for the next stage of inversion. Since the number of shots is considerably larger than the number of receivers for an OBS seismic survey, we used reciprocity for switching the locations of the sources and receivers for an efficient seismic wave modeling while maintaining the same accuracy (Operto et al., 2006). The sources were applied to the OBS locations at the seafloor for the normal stress components. The simulated data of pressure wavefield can be obtained from the average of the normal stress components at the original shot positions.

### 3.5. Source Wavelet Estimation

Solving the elastic wave equation requires a source wavelet as an input. The direct water waves from the air-gun array at the sea surface tend to overlap with the strong scatterings/reflections from the rugged seafloor and reflections from the basement, therefore it can be challenging to obtain a reliable source signature directly from the water waves. Here we compare two methods for estimating the source wavelet. First, we estimate the source wavelet by aligning and stacking the near-offset crustal Pg arrivals from 5 to 7 km. Pg waves turn rapidly in the upper crust and propagate nearly vertically near the source and receivers, rendering few waveform alterations. As Pg rays travel through the crust as turning rays, they are not much affected by seafloor scattering, which have diffraction move-out of water velocity. Furthermore, as these are transmitted waves, the effect of small-scale heterogeneity in the crust is healed. Figure 3a shows the selected crustal Pg waves from the near offsets (dashed gray lines and solid black lines, before and after aligning) of OBS 15 in the frequency band of 3–10 Hz. We stacked the aligned waveforms to enhance the signal-to-noise ratio, and the output is the estimated source wavelet (solid green line in Figure 3a).

Alternatively, we can extract the wavelet signature from the near-offset free-surface multiples. We observe a good isolation of the direct waves in their free-surface related multiples (Guo et al., 2021) in the frequency band of 3–40 Hz, possibly because the strong seafloor-related scatterings in the primary arrivals have been dispersed during the two-way wave propagation in the seawater. We filtered the near-offset data (between −1 and 1 km) at higher frequencies, effectively isolating the water multiples from other events. Figure 3b shows the free-surface multiples (dashed gray lines), from 10.3 to 10.7 s. The waves have been aligned, as shown by the solid black lines. Similarly, we stacked the aligned waves to obtain the source wavelet (solid green line in Figure 3b).

Finally, we filtered the two estimated source wavelets, from the Pg arrivals and the free-surface multiples, respectively, into the same frequency range of 3–10 Hz. Figure 4 illustrates the comparison of the wavelets from the two different parts of the data, which reveals remarkably consistent source signatures in both the time (Figure 4a) and frequency domains (Figure 4b). The starting time in Figure 4a for the source wavelet during wavefield simulation was determined by matching with the water waves. We used the extracted source wavelets for forward modeling. The good match (Figure S1 in Supporting Information S1) on the shape of the waveforms between synthetic and field data suggests that the estimated source wavelet in Figure 4 provides a reliable approximation to the original source signature during the OBS data acquisition. We used the source wavelet from near-offset multiples for FWI.
<figure><img src="image_6.png" /><figcaption>**Figure 4.** Comparison between (a) the extracted source wavelets and (b) the corresponding frequency spectra from Pg and free-surface multiples of water waves, in the frequency range of 3–10 Hz.</figcaption></figure>

# 4. Results

## 4.1. Oceanic Crustal Models From FWI

Figure 5a shows the P-wave velocity model from travel-time tomography, which provides a very good travel-time fit for the first arrivals (Wang & Singh, 2022). Since the travel time is mainly sensitive to large-scale velocity structures, the tomographic velocity model contains few details in the oceanic crust but clearly shows velocity of 5–6.5 km/s in the upper crust and 6.6–7 km/s in the lower crust. Travel-time tomography suggests that the crustal thickness is nearly uniform (5.6 ± 0.2 km) along the whole profile (Wang & Singh, 2022).

Starting from the travel-time tomographic velocity model (Figure 5a), we proceeded with a sequential application of the trace-normalized FWI and the classic true-amplitude FWI methods. The crustal velocity model from trace-normalized FWI contains intermediate-scale features (Figure 5b). Figure 5c presents the final velocity model with

<figure><img src="image_7.png" /><figcaption>**Figure 5.** Seismic P-wave velocity models over a 120 km-long section between the St Paul fracture zone and the Romanche transform fault (TF). (a) The velocity model from travel-time tomography, (b) the velocity model from the trace-normalized full waveform inversion (FWI), and (c) the final velocity model from the classic true-amplitude FWI. Black triangles at the seafloor mark the ocean bottom seismometer (OBS) locations, with the associated OBS IDs labeled in panel (a). The bold black line around 11 km depth in panel (a) is the Moho, the boundary of crust and mantle, from travel-time tomography (Wang & Singh, 2022). The velocity contours of dashed black lines are from 5 to 7 km/s with an increment of 0.5 km/s. Distance 0 along the profile is at the location of extended 6 km from the lower end of the shot profile in Figure 1b, with the longitude and latitude of (−16.9855, −3.65597). The labels of “N” and “S” at top of (a) indicate the north and south directions.</figcaption></figure>
<figure><img src="image_8.png" /><figcaption>**Figure 6.** Data misfits from full waveform inversion (FWI). (a) The misfits as a function of iteration indexes come from the trace-normalized FWI of the near-to-intermediate (iterations 1–20) and full (iterations 21–40) offsets, and true-amplitude FWI of the near-to-intermediate (iterations 41–60) and full (iterations 61–80) offsets. Each inversion, as indicated by different background colors, contains 20 iterations. (1) and (3) use near-to-intermediate offset Pg waves and (2) and (4) use all (full offset range) the Pg arrivals in the ocean bottom seismometer (OBS) data. Note that the misfit was normalized to 1 at the beginning of each inversion step for plotting purpose. (b) Data misfit for each of the 9 OBS gathers from the tomographic and the final FWI models. The waveform data misfits were calculated using Equation 4.</figcaption></figure>

fine-scale geological features emerging from the true-amplitude FWI. The FWI-derived velocity models (Figures 5b and 5c) delineate the oceanic crust with a higher resolution and more structures than that from the travel-time tomography (Figure 5a). We also used the source wavelet estimated from near-offset Pg arrivals for FWI; the inverted velocity model is shown in Figure S2 in Supporting Information S1, which is almost identical to Figure 5c with minimal difference.

## 4.2. Data Misfit and Waveform Comparison

For each of the trace-normalized and true-amplitude FWI stages, we first inverted the near-to-intermediate offsets data, followed by inversion of the full range of Pg arrivals. Each inversion step consists of 20 iterations, leading to 80 iterations in total. Figure 6a shows the data misfits for all the 80 iterations from the trace-normalized and the true-amplitude FWI. Figure 6b shows the data misfits for each of the 9 OBS gathers from the tomographic and the final FWI models. The data misfits have been reduced substantially after the FWI for each of the OBS gather.
<figure><img src="image_9.png" /><figcaption>**Figure 7.** Observed and simulated seismic waveform data from ocean bottom seismometer 15. (a) The field data (black) and the synthetic waveform (red) using the tomographic model. (b) The field data (black) and the synthetic waveform (red) after 80 iterations of the full waveform inversion. The travel time was reduced with a velocity of 7 km/s for both the field and simulated data.</figcaption></figure>

Figure 7 shows the observed and simulated seismic waveform data for OBS 15. The simulated seismic data after the FWI show a much better match with the field data than those from the tomographic model, suggesting that the FWI method has mapped the waveform information from the data into model space and provided higher resolution crustal velocity model.

### 4.3. Sensitivity and Resolution Tests

To demonstrate that the fine-scale structures from FWI of wide-angle Pg waveforms are required by the data, we removed some of the detailed velocity structures in the FWI model (Figure 5c) to see if the waveform match deteriorates. Additionally, we also ran a second FWI to see if the removed structures can be recovered. Furthermore, we performed checkerboard tests for estimating the size of the anomaly that can be recovered by the FWI of crustal turning waves.

We removed some of the interesting features in the FWI model and computed synthetic data. We observe that the waveform fit between the field data and the synthetic data deteriorates (Figure 8) when we remove the high and low velocity structures A and B introduced by the FWI. This observation suggests that these structures are required to explain the strong amplitude variations with offsets. At distance between 510 and 550 km, where there are high velocities (5.5–6 km/s) near the basement in region with low basement morphology, we removed the positive velocity update C from the FWI model (Figure 9). The computed synthetic waveforms using the modified model do not fit the data at all.

Additionally, we also performed a second FWI after removing these features and using modified models as the starting models to see whether a second FWI can recover these omitted structures. Figures 10a–10c show the comparison between the original FWI velocity model (Figure 5c), the modified FWI model with structures A and B removed, and the velocity model from the second FWI. Both structures A and B have been recovered in Figure 10c. Similar tests were performed for the high velocities near the basement (Figure 9), with results shown in Figure 10f. These tests provide confidence that these local features from FWI is required by the data, and are not artifact of inversion.
<figure><img src="image_10.png" /><figcaption>**Figure 8.** Data comparison after removing the high and low velocity structures (A and B) introduced by full waveform inversion (FWI) near ocean bottom seismometer (OBS) 14. (a) and (c) show part of the FWI model (Figure 5c) and the model after removing these structures near OBS 14. (b) Shows waveform comparison of field and synthetic data using the FWI model. (d) Shows waveform comparison of field and synthetic data using the modified FWI model after removing these structures.</figcaption></figure>

In the next step, we conducted checkerboard tests. We generated synthetic models by adding positive and negative Gaussian-shaped velocity anomalies with a magnitude of 0.4 km/s to the tomographic velocity model (Figure 5a). The anomalies have sizes of 7 km in length and 1 km in thickness for one case (Figure 11a), and 7 km in length and 0.5 km in thickness for the other case (Figure 12a). We computed the “observed” data using the perturbed velocity model. We used the same source and receiver positions and frequency ranges as for the field data, and used the same inversion parameters. The FWI started from the unperturbed tomographic model to see how well these

<figure><img src="image_11.png" /><figcaption>**Figure 9.** Data comparison after removing the high velocity structures introduced by full waveform inversion (FWI). (a) and (c) show part of the FWI model (Figure 5c) and the model after removing the positive velocity structures. (b) Shows waveform comparison of field and synthetic data using the FWI model for ocean bottom seismometer 19. (d) Shows waveform comparison of field and synthetic data using the modified FWI model after removing structure C.</figcaption></figure>
<figure><img src="image_12.png" /><figcaption>**Figure 10.** Model recovery test: we removed some velocity features (A, B, and C) from the full waveform inversion (FWI) model in Figure 5c, then conducted another FWI to see whether these omitted features can be recovered. (a) and (d) are velocity models from the original FWI, (b) and (e) are models with certain features removed, and (c) and (f) are velocity models from a second FWI showing recovered omitted features.</figcaption></figure>

perturbations can be recovered. After the FWI, the anomalies are well resolved and exhibit a good match to the true models (Figures 11 and 12). In Figure 12 where the anomalies are smaller, the resolution becomes less favorable however still provide a reasonable match to the majority part of the true model. These studies indicate that FWI of OBS data can resolve structures on the scale of 7 km (half OBS spacing) laterally, and 500–1,000 m vertically. With the available Pg offsets, FWI provide constraints down to ~4–5 km depth from basement (~1 km deeper than travel-time tomography, Figures S3–S7 in Supporting Information S1). The estimated errors from the

<figure><img src="image_13.png" /><figcaption>**Figure 11.** Checkerboard test 1. Each of the anomalies has 7 km lateral extension and 1 km thickness. (a) The true checkerboard velocity perturbation added to the tomographic model, and (b) the recovered velocity anomalies from the full waveform inversion. The dashed and solid black lines at top indicate the seafloor and the basement (the top of layer 2), respectively.</figcaption></figure>
<figure><img src="image_14.png" /><figcaption>**Figure 12.** Checkerboard test 2. Each of the anomalies has 7 km lateral extension and 0.5 km thickness. (a) The true checkerboard velocity perturbation added to the tomographic model, and (b) the recovered velocity anomalies from the full waveform inversion. The dashed and solid black lines at top indicate the seafloor and the basement (the top of layer 2), respectively.</figcaption></figure>

checkerboard tests are about 100–300 m/s. The velocity becomes less well constrained with more pronounced errors in the area of rough basement morphology. The data also has less resolution for the anomalies in the lower part of the model. It is rather difficult to estimate uncertainties in the velocity values, but we expect it to be of the order of ~200 m/s. One can use Marchov chain Monte Carlo type of methods (Gebraad et al., 2020; Guo et al., 2020; Zhang & Curtis, 2020) or the covariance matrix approach (Izzatullah et al., 2019) to provide more robust and comprehensive uncertainty analysis, however limited by computing capabilities these methods become intractable for the problem at the regional scale. We also performed checkerboard tests for travel-time tomography (Figures S4–S7 in Supporting Information S1), which show that travel time tomography can resolve anomalies of 14 km laterally and 1.5 km thickness, down to 3 km depth from basement but failed for smaller ones. The results suggest FWI has much better resolution than the travel-time tomography; the resolution can be further improved when using a smaller OBS spacing.

**5. Discussion**

We plot the high-resolution seafloor bathymetry data surrounding the OBS profile in Figure 13a. To highlight the salient features of the final crustal models from FWI, in addition to the velocity model in Figure 13b (the same with Figure 5c), we also show the velocity difference between the FWI and tomographic models (Figure 13c), the vertical velocity gradient (Figure 13d), and the horizontal velocity gradient (Figure 13e). Compared with the tomographic model (Figure 5a), we observe generally an increase in the velocity in the shallow crust and an overall decrease in velocity at the greater depths (Figure 5d), with boundaries becoming sharper. The vertical velocity gradient in Figure 13d in the shallow crust (1–2.5 s⁻¹) is in general larger than that at the greater depths (0.1–0.2 s⁻¹), and highlights layered structures in the crust. On the other hand, the complicated positive and negative patterns in the horizontal velocity gradient (Figure 13e) bring to light the strong lateral crustal heterogeneities (up to 0.3 s⁻¹ horizontal gradient) in the region.

From the high-resolution seafloor bathymetry in Figure 13a, we can discern four distinct features, each spanning 20–30 km in distance. From north to south, first, we observe linear features in the proximity to OBS 14 and OBS 12, which are parts of the St Paul FZ. Moving further south, we encounter an abyssal hill morphology, followed by a notable sediment layer with a thickness reaching up to 1.5 km and a remarkable basement topography. Lastly, on the southmost side, a landslide surface becomes evident, which is associated with the northern flank of the Romanche TF. These four zones are defined as zone 1 to 4.
<figure><img src="image_15.png" /><figcaption>**Figure 13.** FWI-derived oceanic crustal models along the seafloor bathymetry. (a) The high-resolution seafloor bathymetry along the ocean bottom seismometer (OBS) profile between the St. Paul fracture zone and the Romanche transform fault. Black triangles at the seafloor mark the OBS locations, with the associated OBS IDs labeled. The pink line shows the source locations for the survey. The plate ages were plotted using black lines and labeled every 5 Myr (Müller et al., 2008). (b) The final P-wave velocity from full waveform inversion (FWI) (the same with Figure 5c). Black triangles at the seafloor mark the OBS locations, with the OBS IDs labeled. The velocity contours of dashed black lines are from 5 to 7 km/s with an increment of 0.5 km/s. The dashed white lines in zones 1 and 4 mark the vertical velocity gradient changes. The solid white line generally traces the ~6.5 km/s contour. (c) The velocity update from FWI, that is, the difference between the final model from FWI and travel-time tomography. (d) The vertical velocity gradient, the derivative of the velocity model with respect to depth (z direction). (e) The horizontal velocity gradient, the derivative of the velocity model with respect to the distance (x direction). The dashed red arrows mark the locations for the 1D profiles in Figure 14. The labels of “N” and “S” at top of (a) indicate the north and south directions. The dashed lines and rectangles in panels (d) and (e) are discussed in the discussion section.</figcaption></figure>

Correspondingly, we discover four distinct crustal segments in the velocity model within the ~120 km long section (Figure 13b) associated with the overlying seafloor bathymetry. Although the tomographic velocity model (Figure 5a) indicates four different crustal features, the FWI results enhance these features. The velocity difference plot (Figure 13c) indicates that there has been up to ±600 m/s change in velocity from the tomography results. The velocity update from FWI also shows the presence of some sub-horizontal features, especially in zone 1, which becomes more pronounced in the vertical velocity gradient (Figure 13d), as indicated by the dashed rectangles. Although the horizontal velocity gradient is more complex, it does delineate the boundaries between different zones, as marked by the dashed green lines in Figure 13e. Depending on the velocity, the horizontal velocity gradients at the zone boundaries can be positive or negative. Here, FWI of Pg waves provides constraints for the oceanic crust down to 4 km depth from basement.

Figure 14 shows four selected one-dimensional (1-D) velocity-depth profiles below the basement from the four zones from the tomographic (Figure 5a) and FWI (Figure 5c) models. We also plotted the compiled velocity structures (Christeson et al., 2019) for mature oceanic crust older than 7.5 Ma at slow spreading rates from travel-time tomography. Although the velocity profiles from tomographic and FWI models fall within the ranges for
<figure><img src="image_16.png" /><figcaption>**Figure 14.** Velocity-depth (1-D) profiles from different distances along the profile in the seismic P-wave velocity models for the oceanic crust. The solid blue curves are from the travel-time tomographic model (Figure 5a), and the solid black curves are from the full waveform inversion model (Figure 5c). The yellow curves were from the compilation of velocity structure (Christeson et al., 2019) for the slow-spreading rate mature crust older than 7.5 Ma. The dashed horizontal lines indicate the boundary for vertical velocity gradient changes, and the solid horizontal lines indicate the top boundary of high velocities.</figcaption></figure>

mature oceanic crust at slow spreading environments, there are significant variations in the 1-D velocity profiles across different locations, with boundaries well defined.

In summary, the four segments have distinct velocity structures: (a) Segment 1 consists of three layers where velocity increases from 4.8 to 7 km/s from seafloor down to 4 km depth. (b) Segment 2 has two layers with a high velocity (7 km/s) at 2 km below the seafloor. (c) Segment 3 consists of a single layer with a high velocity increasing from 5 to 6 km from the seafloor down to 4 km depth. (d) Segment 4 consists of three layers, similar to segment 1. Below, we will discuss the velocity structure in detail and the underlying accretion mechanisms for the four distinct segments.

## 5.1. Zone 1

At the north side of the model section (Figure 13b), where the oceanic crust is affected by the St Paul FZ, the observed velocity model shows three-layered oceanic crustal velocities, containing a ~2-km thick layer where the velocity increases from ~4.8 to ~6 km/s (separated by a dashed white line), a ~1.5-km thick middle crust with reduced vertical velocity gradient with a velocity of 6–6.5 km/s, which is underlain by high velocity ~6.6–7.0 km/s (top boundary indicated by the solid white line) in the lower crust. The three-layer crust suggests a magmatic process for the crustal formation, consistent with the travel-time tomographic results of Growe et al. (2021), Gregory et al. (2021), and Marjanović et al. (2020), for the St. Paul FZ, Romanche TF, and the Chain FZ, respectively. Our results, however, show more detailed structures in the crust.

The rugged velocity contours in the upper crust could be introduced by the tectonic processes (faulting and hydrothermal alterations) associated with the St Paul FZ. For example, the velocity anomaly around OBS 12 corresponds to the linear feature in the FZ illustrated in Figure 13a. The relatively slow seismic velocity suggests highly fractured upper crust affected by the St Paul FZ, as P-wave velocities in the upper oceanic crust are largely controlled by porosity (e.g., Carlson, 2001, 2010, 2014; Fox et al., 1973; Hyndman & Drury, 1976; Salisbury et al., 1996; Spudich & Orcutt, 1980b; Wilkens et al., 1991). The base of the uppermost crust, marked by the vertical gradient changes (Figure 13d), could be related to the transition from highly fractured crust to less fractured crust. This gradient change could also represent the layer 2/3 boundary.

Considering the uncertainty in this part the model from the checkerboard tests (Figures 12 and 13), we need to be cautious in over-interpreting some of the small-scale features, but it would be worth pointed out some of these features. For example, the low velocity between 2.8 and 3.4 km depths from the basement (Figure 14a) might be caused by the introduction of hydrous phases such as chlorite, clay minerals and even serpentine (if the unaltered rock was rich in olivine) from enhanced hydrothermal alteration (Carlson & Jay Miller, 2004; Singh et al., 1999; Wilson et al., 2006), in addition to fracturing. IODP Hole U1309D at the Atlantis Massif (Blackman et al., 2014,
2019) reveals a strong correlation between seismic velocity reduction and mineral alteration, relating to past hydrothermal activities. We also observed layered structures between 1.8 – 4 km depth (indicated by the dashed rectangle in Figure 13d). Guo et al. (2022) found similar structures for 7–12 Ma old crust further south, and interpreted them to be due to the presence of high and low velocity layers associated with in situ melt injection and crystallization in the lower crust. The presence of the layered structure, along a seismic profile parallel to the ridge axis, perpendicular to those in Guo et al. (2022), could be due to 3-D nature of crustal accretion; What is interesting here is that such a layering is observed in a zone affected by the St Paul FZ. However, if this sub-segment was formed at a small ridge segment within the transform zone (Maia et al., 2016), and if the lithosphere beneath the TF is thin (Wang et al., 2022), there might be enhanced magma supply at inter-transform ridge segment with an increase in melt production due to water-induced melting (Wang et al., 2022).

## 5.2. Zone 2

Underlying the abyssal hill morphology at 550–570 km distance along the profile, the oceanic crust consists of two layers, ~2–2.5 km thick upper crust and a high velocity ~7 km/s zone with a thickness at least ~3.5 km. The change of vertical velocity gradient, marked by the dashed line in zone 2 (Figure 13d), is consistent with the top boundary of high velocities (solid white line, Figure 13b). A comparison of a velocity profile with compiled velocity structures (Christeson et al., 2019) in Figure 14b suggests that although the high velocities ~7 km/s starting from 2 to 2.5 km depth falls within the range observed for old oceanic crust at slow-spreading rate, they are at the higher end. The P-wave velocity in the lower crust comprised of gabbro is usually between 6.6 and 7.2 km/s (Carlson & Jay Miller, 2004; Christeson et al., 2019). The gabbroic rocks are mainly composed of olivine (Ol), clinopyroxene (Cpx) and plagioclase (Pl), with seismic velocities $V_{Ol} > V_{Cpx} > V_{Pl}$. P-wave velocities increase with increasing pyroxene content, with gabbros having velocities of 7 km/s containing about 25% modal pyroxene (Carlson & Jay Miller, 2004). On the other hand, increasing velocity seems to have little or no variations in Ol. Carlson and Jay Miller (2004) observed that increasing modal Ol content in oceanic gabbros recovered by drilling in Holes 735B, 894G, and 923A does not increase the P-wave velocity due to the offset by the associated serpentine content (Carlson et al., 2009). The observed high velocity could also come from partially serpentinized peridotite, which can be uplifted from the mantle to the shallower crust through tectonic activities such as long-lived detachment faulting. For example, seismic surveys in the Kane OCC of the Atlantic Ocean have imaged the distribution of gabbro, serpentinized peridotite and basalt (Canales, 2010). However, the seafloor morphology of ridge-parallel abyssal hills is typical of volcanic terrain, suggesting a robust magmatic supply. Therefore we suggest the velocity of ~7 km/s starting from 2 to 2.5 km depth more likely comes from primitive gabbros rich in Cpx. We observe alternating high and low velocities between 575 and 585 km distance (Figure 13b), which could be a trace of fault, indicating segment boundary. Underneath the linear feature related to St Paul FZ, the high velocity ~6.6 km/s about 1 km from seafloor could suggest gabbro or serpentinized peridotite associated with an OCC.

## 5.3. Zone 3

The segment between ~520 and 545 km is associated with the low basement morphology and contains a nearly constant high velocity ranging from 5.5 to ~6.5 km/s commencing near the basement and extending down to a depth of 4 km. Although the velocity is in general within the ranges observed at slow spreading rates (Figure 14c), it does not resemble a magmatically accreted two layer oceanic crust. The dashed green lines in Figure 13e mark the zone boundaries, suggesting the boundary changes with depths. We interpret the lower crust below 6.5 km/s contour, starting from ~2–3 km depth below the basement. The appearance of ~6 km/s near basement and the subtle velocity variations at depth suggest a chemically different composition, compared to the normal oceanic crust. Although we acknowledge that wide-angle OBS data may have limited resolution in the vicinity of the basement due to the lack of turning wave propagation, particularly in areas with rough topography, resolution study implies that the estimated velocities in the upper 3–4 km depth range are reliable (Figures 12 and 13). The nearly constant velocity down to a depth of 4 km suggests it may be hydrated crust containing hydrothermally altered gabbro or can contain altered mantle rocks such as highly serpentinized peridotite. The altered mafic and ultramafic rocks can be emplaced from greater depths through tectonic activities. Similarly, Davy et al. (2020) observed reduced contrast in the velocity gradient between the upper and lower crust, suggesting a low magma budget and tectonically controlled crustal accretion.
5.4. Zone 4

Similar to zone 1, the P-wave velocity show features of a normal oceanic crust with three distinct layers. The top layer has a high vertical velocity gradient, underlain by a layer with velocity from ~6 km/s and intermediate velocity gradient, and a bottom layer with velocity of ~6.5–7 km/s and a low velocity gradient interpreted as the lower crust (Figures 13 and 14d). The dashed red line in Figure 14d could be the boundary of layer 2A/2B, and the solid red line marks the top of lower crust. The velocity of 4.5–5 km/s near the basement could be basalt, gabbro, or serpentinized peridotite. Considering previous seafloor dredgings have discovered an extensive occurrence of basaltic and gabbroic rocks (Bonatti, 1968; Bonatti & Honnorez, 1976; Seyler & Bonatti, 1997), the velocity near the basement is more likely from basaltic rocks. The velocity increases to 6.5–7 km/s in the lower crust and changes more slowly with depth, suggesting the presence of gabbroic rocks. There are some layered structures in the lower crust (the dashed rectangle in Figure 13d), similar to zone 1, which might be related to layering in the lower crust (Guo et al., 2022).

We interpret the significant heterogeneities in the oceanic crust arising from the complex interplay of melt generation and emplacement, hydrothermal alteration, and tectonic forces. The magma generated in the mantle rises through the mantle in the lower crust, moving along fissures and melt channels. The melt is eventually emplaced in the crust by intrusion (dike) and extrusion (forming the basaltic lava flow), while it may cool and crystallize (in-situ) forming the lower (gabbroic) crust. The newly formed crust undergoes hydrothermal alteration, which introduces heterogeneities by affecting rock density, porosity, and the presence of mineral precipitates. At slow-spreading ridges, tectonic activities such as large faulting may expose deeper sections of the crust or even mantle rocks at the surface, which can result in a highly variable crustal structure with different types of rocks juxtaposed against one another due to faulting and deformation. On the other hand, ubiquitous heterogeneity in the mantle (Hart, 1988; Lambart et al., 2019; Ritsema & Lekić, 2020) can play a vital role in melt generation and crustal composition (Shorttle & Maclennan, 2011). On the MAR north of the Atlantis TF where mafic and ultramafic plutonic rocks were exposed by detachment faulting, drill core of gabbros from Hole U1309D rendered a detailed, 1,415 m record of the lower crust. Result of Nd (¹⁴³Nd/¹⁴⁴Nd) isotopic analysis of Pl and Cpx of drill core gabbroic cumulate reveals substantial heterogeneity in the composition of mantle melt delivered to the lower crust (Lambart et al., 2019), exceeding the range from mid-ocean ridge basalt (MORB) (Hart, 1988) by a factor of seven. These results also suggest that the mantle melt has limited mixing before being delivered to the crust, supporting in-situ melt crystallization for lower crustal accretion (Boudier et al., 1996; Guo et al., 2022; Kelemen et al., 1997), albeit the composition of MORB suggests that the melt has been homogenized through efficient mixing at the shallow crustal melt lens (Lambart et al., 2019). In north Iceland, also a slow-spreading environment, petrological analysis using basaltic rocks indicates the compositions of enriched and depleted major elements that existed within similar age eruptions from a single volcanic system, indicating a fine-scale mantle source heterogeneity (Shorttle & Maclennan, 2011). In addition, lava samples from off-axis seamount in the fast-spreading East Pacific Rise contain significant compositional variations ranging from depleted MORB to high enriched MORB at small (~5 km) scales (Anderson et al., 2021), exhibiting extreme mantle melt heterogeneity. The crustal and upper mantle heterogeneities could also be related to thermal anomaly associated with small-scale convention along the ridge (Ballmer et al., 2011).

6. Conclusions

Applying FWI to the crustal turning waves observed in the OBS data from the equatorial Atlantic Ocean, we have discovered strong heterogeneities in the oceanic crust between the St Paul FZ and the north flank of the Romanch TF. Together with the high-resolution seafloor bathymetry data, we define four distinct zones in the seismic velocity model of the oceanic crust. We observe crustal velocities exhibiting features of a magmatic oceanic crust associated with the St Paul FZ and the north flank of Romanch TF. The segment underlying the abyssal hills contains a high velocity (~7 km/s) crust starting from 2–2.5 km depths, possibly containing unaltered gabbro and serpentinized peridotite. We also observe a segment with nearly constant velocity of 5.5–6.5 km/s from the basement down to 4 km depth, suggesting a hydrated crust. The oceanic crustal heterogeneity along a single segment of ~120 km distance offers a rare window to gain insights for the many factors that contribute to crustal accretion, including magmatic, tectonic and hydrothermal activities. The variations in the mantle temperature and composition might also play a role in making crust heterogeneous.
## Data Availability Statement

The raw OBS data from the OBS12 to OBS15 are available online at the PANGAEA website (https://doi.pangaea.de/10.1594/PANGAEA.937195). The raw OBS data from the OBS16 to OBS20 are available online at https://doi.org/10.1594/PANGAEA.946565. The travel time pickings for the Pg arrivals and the derived velocity models from FWI can be accessed at Guo and Singh (2023).

## Acknowledgments

The OBS data were acquired during the ILAB-SPARC experiment onboard the French R/V *Pourquoi Pas?* October to November 2018. This project was funded by the European Research Council Advanced Grant 339442 TransAtlanticILAB. This research was funded by the Deep Earth Imaging Future Science Platform, CSIRO. The authors thank Z. Wang and E. Gregory for providing the travel-time picks (used for defining time window of data for FWI). P. Guo thanks Z. Wang for helpful discussion on the raw OBS data. We thank Rick Carlson, Tim Henstock, and two anonymous reviewers for their constructive comments that improve the paper significantly. This is an IPGP contribution number 4290.

## References

Allen, R. W., Collier, J. S., & Henstock, T. J., & VoiLA Consortium. (2022). The role of crustal accretion variations in determining slab hydration at an Atlantic subduction zone. *Journal of Geophysical Research: Solid Earth*, 127(8), e2022JB024349. https://doi.org/10.1029/2022JB024349

Alt, J. C., Laverne, C., Vanko, D. A., Tartarotti, P., Teagle, D. A., Bach, W., et al. (1996). Hydrothermal alteration of a section of upper oceanic crust in the Eastern Equatorial Pacific: A synthesis of results from site 504 (DSDP LEGS 69, 70, and 83, and ODP LEGS 111, 137,140, and 148). In *Proceedings of the Ocean Drilling Program, scientific results* (Vol. 148, pp. 417–434).

Anderson, M., Wanless, V. D., Perfit, M., Conrad, E., Gregg, P., Fornari, D., & Ridley, W. I. (2021). Extreme heterogeneity in mid-ocean ridge mantle revealed in lavas from the 8° 20' N near-axis seamount chain. *Geochemistry, Geophysics, Geosystems*, 22(1), e2020GC009322. https://doi.org/10.1029/2020GC009322

Arnoux, G. M., Toomey, D. R., Hooft, E. E., Wilcock, W. S., Morgan, J., Warner, M., & VanderBeek, B. P. (2017). Seismic evidence that black smoker heat flux is influenced by localized magma replenishment and associated increases in crustal permeability. *Geophysical Research Letters*, 44(4), 1687–1695. https://doi.org/10.1002/2016GL071990

Arnulf, A. F., Biemiller, J., Lavier, L., Wallace, L. M., Bassett, D., Henrys, S., et al. (2021). Physical conditions and frictional properties in the source region of a slow-slip event. *Nature Geoscience*, 14(5), 334–340. https://doi.org/10.1038/s41561-021-00741-0

Arnulf, A. F., Harding, A. J., Singh, S. C., Kent, G. M., & Crawford, W. C. (2014). Nature of upper crust beneath the Lucky Strike volcano using elastic full waveform inversion of streamer data. *Geophysical Journal International*, 196(3), 1471–1491. https://doi.org/10.1093/gji/ggt461

Audhkhasi, P., & Singh, S. C. (2019). Seismic structure of the upper crust from 0–75 Ma in the equatorial Atlantic Ocean on the African plate using ultralong offset seismic data. *Geochemistry, Geophysics, Geosystems*, 20(12), 6140–6162. https://doi.org/10.1029/2019GC008577

Ballmer, M. D., Ito, G., Van Hunen, J., & Tackley, P. J. (2011). Spatial and temporal variability in Hawaiian hotspot volcanism induced by small-scale convection. *Nature Geoscience*, 4(7), 457–460. https://doi.org/10.1038/ngeo1187

Bird, P. (2003). An updated digital model of plate boundaries. *Geochemistry, Geophysics, Geosystems*, 4(3). https://doi.org/10.1029/2001GC000252

Blackman, D. K., Abe, N., Carlson, R. L., Guerin, G., Ildefonse, B., & Kumpf, A. (2019). Seismic properties of gabbroic sections in oceanic core complexes: Constraints from seafloor drilling. *Marine Geophysical Research*, 40(4), 557–569. https://doi.org/10.1007/s11001-019-09385-7

Blackman, D. K., & Collins, J. A. (2010). Lower crustal variability and the crust/mantle transition at the Atlantis Massif oceanic core complex. *Geophysical Research Letters*, 37(24). https://doi.org/10.1029/2010GL045165

Blackman, D. K., Slagle, A., Guerin, G., & Harding, A. (2014). Geophysical signatures of past and present hydration within a young oceanic core complex. *Geophysical Research Letters*, 41(4), 1179–1186. https://doi.org/10.1002/2013gl058111

Bonatti, E. (1968). Ultramafic rocks from the mid-Atlantic ridge. *Nature*, 219(5152), 363–364. https://doi.org/10.1038/219363a0

Bonatti, E., & Honnorez, J. (1976). Sections of the Earth's crust in the equatorial Atlantic. *Journal of Geophysical Research*, 81(23), 4104–4116. https://doi.org/10.1029/JB081i023p04104

Bonatti, E., Ligi, M., Gasperini, L., Peyve, A., Raznitsin, Y., & Chen, Y. J. (1994). Transform migration and vertical tectonics at the Romanche fracture zone, equatorial Atlantic. *Journal of Geophysical Research*, 99(B11), 21779–21802. https://doi.org/10.1029/94JB01178

Boudier, F., Nicolas, A., & Ildefonse, B. (1996). Magma chambers in the Oman ophiolite: Fed from the top and the bottom. *Earth and Planetary Science Letters*, 144(1–2), 239–250. https://doi.org/10.1016/0012-821X(96)00167-7

Boulahanis, B., Carbotte, S. M., Canales, J. P., Han, S., & Nedimović, M. R. (2022). Structure and evolution of northern Juan de Fuca crust and uppermost mantle over the last 8 Ma from an active-source seismic tomography study. *Journal of Geophysical Research: Solid Earth*, 127(9), e2022JB023987. https://doi.org/10.1029/2022JB023987

Brocher, T. M. (2005). Empirical relations between elastic wavespeeds and density in the Earth's crust. *Bulletin of the Seismological Society of America*, 95(6), 2081–2092. https://doi.org/10.1785/0120050077

Buck, W. R., Lavier, L. L., & Poliakov, A. N. (2005). Modes of faulting at mid-ocean ridges. *Nature*, 434(7034), 719–723. https://doi.org/10.1038/nature03358

Canales, J. P. (2010). Small-scale structure of the Kane oceanic core complex, Mid-Atlantic Ridge 23 30' N, from waveform tomography of multichannel seismic data. *Geophysical Research Letters*, 37(21). https://doi.org/10.1029/2010GL044412

Canales, J. P., Collins, J. A., Escartín, J., & Detrick, R. S. (2000). Seismic structure across the rift valley of the Mid-Atlantic Ridge at 23 20' (MARK area): Implications for crustal accretion processes at slow spreading ridges. *Journal of Geophysical Research*, 105(B12), 28411–28425. https://doi.org/10.1029/2000JB900301

Canales, J. P., Detrick, R. S., Carbotte, S. M., Kent, G. M., Diebold, J. B., Harding, A., et al. (2005). Upper crustal structure and axial topography at intermediate spreading ridges: Seismic constraints from the southern Juan de Fuca Ridge. *Journal of Geophysical Research*, 110(B12). https://doi.org/10.1029/2005JB003630

Cann, J. R., Blackman, D. K., Smith, D. K., McAllister, E., Janssen, B., Mello, S., et al. (1997). Corrugated slip surfaces formed at ridge-transform intersections on the Mid-Atlantic Ridge. *Nature*, 385(6614), 329–332. https://doi.org/10.1038/385329a0

Cannat, M. (1993). Emplacement of mantle rocks in the seafloor at mid-ocean ridges. *Journal of Geophysical Research*, 98(B3), 4163–4172. https://doi.org/10.1029/92JB02221

Cannat, M., Lagabrielle, Y., Bougault, H., Casey, J., de Coutures, N., Dmitriev, L., & Fouquet, Y. (1997). Ultramafic and gabbroic exposures at the Mid-Atlantic Ridge: Geological mapping in the 15 N region. *Tectonophysics*, 279(1–4), 193–213. https://doi.org/10.1016/S0040-1951(97)00113-3

Carbotte, S. M., Marjanović, M., Carton, H., Mutter, J. C., Canales, J. P., Nedimović, M. R., et al. (2013). Fine-scale segmentation of the crustal magma reservoir beneath the East Pacific Rise. *Nature Geoscience*, 6(10), 866–870. https://doi.org/10.1038/ngeo1933

Carbotte, S. M., Smith, D. K., Cannat, M., & Klein, E. M. (2016). Tectonic and magmatic segmentation of the Global Ocean Ridge System: A synthesis of observations. *Geological Society, London, Special Publications*, 420(1), 249–295. https://doi.org/10.1144/SP420.5
Carlson, R. L. (2001). The effects of temperature, pressure, and alteration on seismic properties of diabase dike rocks from DSDP/ODP Hole 504B. *Geophysical Research Letters*, 28(20), 3979–3982. https://doi.org/10.1029/2001GL013426

Carlson, R. L. (2010). How crack porosity and shape control seismic velocities in the upper oceanic crust: Modeling downhole logs from Holes 504B and 1256D. *Geochemistry, Geophysics, Geosystems*, 11(4). https://doi.org/10.1029/2009GC002955

Carlson, R. L. (2011). The effect of hydrothermal alteration on the seismic structure of the upper oceanic crust: Evidence from Holes 504B and 1256D. *Geochemistry, Geophysics, Geosystems*, 12(9). https://doi.org/10.1029/2011GC003624

Carlson, R. L. (2014). The effects of alteration and porosity on seismic velocities in oceanic basalts and diabases. *Geochemistry, Geophysics, Geosystems*, 15(12), 4589–4598. https://doi.org/10.1002/2014GC005537

Carlson, R. L., & Jay Miller, D. (2004). Influence of pressure and mineralogy on seismic velocities in oceanic gabbros: Implications for the composition and state of the lower oceanic crust. *Journal of Geophysical Research*, 109(B9). https://doi.org/10.1029/2003JB002699

Carlson, R. L., Miller, D. J., & Newman, J. (2009). Olivine enigma: Why alteration controls the seismic properties of oceanic gabbros. *Geochemistry, Geophysics, Geosystems*, 10(3). https://doi.org/10.1029/2008GC002263

Christensen, N. I. (1996). Poisson's ratio and crustal seismology. *Journal of Geophysical Research*, 101(B2), 3139–3156. https://doi.org/10.1029/95JB03446

Christeson, G. L., Goff, J. A., & Reece, R. S. (2019). Synthesis of oceanic crustal structure from two-dimensional seismic profiles. *Reviews of Geophysics*, 57(2), 504–529. https://doi.org/10.1029/2019RG000641

Christeson, G. L., McIntosh, K. D., & Karson, J. A. (2007). Inconsistent correlation of seismic layer 2a and lava layer thickness in oceanic crust. *Nature*, 445(7126), 418–421. https://doi.org/10.1038/nature05517

Christeson, G. L., Morgan, J. V., & Warner, M. R. (2012). Shallow oceanic crust: Full waveform tomographic images of the seismic layer 2A/2B boundary. *Journal of Geophysical Research*, 117(B5). https://doi.org/10.1029/2011JB008972

Davy, R. G., Collier, J. S., Henstock, T. J., VoiLA, C., Rietbrock, A., Goes, S., & Labahn, E. (2020). Wide-angle seismic imaging of two modes of crustal accretion in mature Atlantic Ocean crust. *Journal of Geophysical Research: Solid Earth*, 125(6), e2019JB019100. https://doi.org/10.1029/2019JB019100

Davy, R. G., Frahm, L., Bell, R., Arai, R., Barker, D. H., Henrys, S., et al. (2021). Generating high-fidelity reflection images directly from full-waveform inversion: Hikurangi subduction zone case study. *Geophysical Research Letters*, 48(19), e2021GL094981. https://doi.org/10.1029/2021GL094981

Detrick, R. S., Needham, H. D., & Renard, V. (1995). Gravity anomalies and crustal thickness variations along the Mid-Atlantic Ridge between 33° N and 40° N. *Journal of Geophysical Research*, 100(B3), 3767–3787. https://doi.org/10.1029/94JB02649

Dhabaria, N., & Singh, S. C. (2024). Seismic evidence of an extremely thin crust along a 70 Ma slow-spreading crustal segment in the equatorial Atlantic Ocean. *Geology*, in press. https://doi.org/10.1130/g52480.1

Dick, H. J., Natland, J. H., Alt, J. C., Bach, W., Bideau, D., Gee, J. S., et al. (2000). A long in situ section of the lower ocean crust: Results of ODP Leg 176 drilling at the Southwest Indian Ridge. *Earth and Planetary Science Letters*, 179(1), 31–51. https://doi.org/10.1016/S0012-821X(00)00102-3

Dick, H. J., Tivey, M. A., & Tucholke, B. E. (2008). Plutonic foundation of a slow-spreading ridge segment: Oceanic core complex at Kane Megamullion, 23 30' N, 45 20' W. *Geochemistry, Geophysics, Geosystems*, 9(5). https://doi.org/10.1029/2007GC001645

Dunn, R. A., Arai, R., Eason, D. E., Canales, J. P., & Sohn, R. A. (2017). Three-dimensional seismic structure of the Mid-Atlantic Ridge: An investigation of tectonic, magmatic, and hydrothermal processes in the rainbow area. *Journal of Geophysical Research: Solid Earth*, 122(12), 9580–9602. https://doi.org/10.1002/2017JB015051

Escartín, J., Mevel, C., Petersen, S., Bonnemains, D., Cannat, M., Andreani, M., et al. (2017). Tectonic structure, evolution, and the nature of oceanic core complexes and their detachment fault zones (13 20' N and 13 30' N, Mid Atlantic Ridge). *Geochemistry, Geophysics, Geosystems*, 18(4), 1451–1482. https://doi.org/10.1002/2016GC006775

Escartín, J., Smith, D. K., Cann, J., Schouten, H., Langmuir, C. H., & Escrig, S. (2008). Central role of detachment faults in accretion of slow-spreading oceanic lithosphere. *Nature*, 455(7214), 790–794. https://doi.org/10.1038/nature07333

Forbriger, T., Groos, L., & Schäfer, M. (2014). Line-source simulation for shallow-seismic data. Part 1: Theoretical background. *Geophysical Journal International*, 198(3), 1387–1404. https://doi.org/10.1093/gji/ggu199

Fox, P. J., Schreiber, E., & Peterson, J. J. (1973). The geology of the oceanic crust: Compressional wave velocities of oceanic rocks. *Journal of Geophysical Research*, 78(23), 5155–5172. https://doi.org/10.1029/JB078i023p05155

Gardner, G. H. F., Gardner, L. W., & Gregory, A. (1974). Formation velocity and density—The diagnostic basics for stratigraphic traps. *Geophysics*, 39(6), 770–780. https://doi.org/10.1190/1.1440465

Gebraad, L., Boehm, C., & Fichtner, A. (2020). Bayesian elastic full-waveform inversion using Hamiltonian Monte Carlo. *Journal of Geophysical Research: Solid Earth*, 125(3), e2019JB018428. https://doi.org/10.1029/2019JB018428

Górszczyk, A., Operto, S., & Malinowski, M. (2017). Toward a robust workflow for deep crustal imaging by FWI of OBS data: The eastern Nankai Trough revisited. *Journal of Geophysical Research: Solid Earth*, 122(6), 4601–4630. https://doi.org/10.1002/2016JB013891

Gregory, E. P. M., Singh, S. C., Marjanović, M., & Wang, Z. (2021). Evidence for thick mafic crust at the slow-slipping Romanche oceanic transform fault. *Geology*, 49(907). https://doi.org/10.1130/G49097.1

Grevemeyer, I., Hayman, N. W., Peirce, C., Schwardt, M., Van Avendonk, H. J., Dannowski, A., & Papenberg, C. (2018). Episodic magmatism and serpentinized mantle exhumation at an ultraslow-spreading centre. *Nature Geoscience*, 11(6), 444–448. https://doi.org/10.1038/s41561-018-0124-6

Grevemeyer, I., Ranero, C. R., & Ivandic, M. (2018). Structure of oceanic crust and serpentinization at subduction trenches. *Geosphere*, 14(2), 395–418. https://doi.org/10.1130/GES01537.1

Growe, K., Grevemeyer, I., Singh, S. C., Marjanović, M., Gregory, E. P. M., Papenberg, C., et al. (2021). Seismic structure of the St. Paul Fracture Zone and Late Cretaceous to Mid Eocene oceanic crust in the equatorial Atlantic Ocean near 18° W. *Journal of Geophysical Research: Solid Earth*, 126(11), e2021JB022456. https://doi.org/10.1029/2021JB022456

Guo, P., Saygin, E., & Kennett, B. (2024). Space-time monitoring of seafloor velocity changes using seismic ambient noise. *Journal of Geophysical Research: Solid Earth*, 129(6), e2023JB027953. https://doi.org/10.1029/2023jb027953

Guo, P., & Singh, S. C. (2023). Seismic evidence for velocity heterogeneity along ~40 Ma old oceanic crustal segment formed at the slow-spreading Mid-Atlantic Ridge in the equatorial Atlantic Ocean from full waveform inversion of ocean bottom seismic data (version 1) [Dataset]. *Zenodo*. https://doi.org/10.5281/zenodo.8283195

Guo, P., Singh, S. C., Vaddineni, V. A., Grevemeyer, I., & Saygin, E. (2022). Lower oceanic crust formed by in situ melt crystallization revealed by seismic layering. *Nature Geoscience*, 15(7), 591–596. https://doi.org/10.1038/s41561-022-00963-w
Guo, P., Singh, S. C., Vaddineni, V. A., Visser, G., Grevetmeyer, I., & Saygin, E. (2021). Nonlinear full waveform inversion of wide-aperture OBS data for Moho structure using a trans-dimensional Bayesian method. *Geophysical Journal International*, 224(2), 1056–1078. https://doi.org/10.1093/gji/ggaa505

Guo, P., Visser, G., & Saygin, E. (2020). Bayesian trans-dimensional full waveform inversion: Synthetic and field data application. *Geophysical Journal International*, 222(1), 610–627. https://doi.org/10.1093/gji/ggaa201

Hamilton, E. L. (1978). Sound velocity–density relations in sea-floor sediments and rocks. *Journal of the Acoustical Society of America*, 63(2), 366–377. https://doi.org/10.1121/1.381747

Han, S., Carbotte, S. M., Canales, J. P., Nedimović, M. R., Carton, H., Gibson, J. C., & Horning, G. W. (2016). Seismic reflection imaging of the Juan de Fuca plate from ridge to trench: New constraints on the distribution of faulting and evolution of the crust prior to subduction. *Journal of Geophysical Research: Solid Earth*, 121(3), 1849–1872. https://doi.org/10.1002/2015JB012416

Harding, A. J., Kent, G. M., & Orcutt, J. A. (1993). A multichannel seismic investigation of upper crustal structure at 9 N on the East Pacific Rise: Implications for crustal accretion. *Journal of Geophysical Research*, 98(B8), 13925–13944. https://doi.org/10.1029/93JB00886

Hart, S. R. (1988). Heterogeneous mantle domains: Signatures, genesis and mixing chronologies. *Earth and Planetary Science Letters*, 90(3), 273–296. https://doi.org/10.1016/0012-821X(88)90131-8

Hicks, S. P., Okuwaki, R., Steinberg, A., Rychert, C. A., Harmon, N., Abercrombie, R. E., et al. (2020). Back-propagating supershear rupture in the 2016 M w 7.1 Romanche transform fault earthquake. *Nature Geoscience*, 13(9), 647–653. https://doi.org/10.1038/s41561-020-0619-9

Hooft, E. E. E., Detrick, R. S., Toomey, D. R., Collins, J. A., & Lin, J. (2000). Crustal thickness and structure along three contrasting spreading segments of the Mid-Atlantic Ridge, 33.5–35 N. *Journal of Geophysical Research*, 105(B4), 8205–8226. https://doi.org/10.1029/1999JB900442

Hyndman, R. D., & Drury, M. J. (1976). The physical properties of oceanic basement rocks from deep drilling on the Mid-Atlantic Ridge. *Journal of Geophysical Research*, 81(23), 4042–4052. https://doi.org/10.1029/JB081i023p04042

Iturrino, G. J., Ildefonse, B., & Boitnott, G. (2002). Velocity structure of the lower oceanic crust: Results from Hole 735B, Atlantis II Fracture Zone. In *Proc. ODP, Sci. Results* (Vol. 176, pp. 1–71). Ocean Drilling Program. https://doi.org/10.2973/odp.proc.sr.176.018.2002

Izzatullah, M., van Leeuwen, T., & Peter, D. (2019). Bayesian uncertainty estimation for full waveform inversion: A numerical study. In *SEG International Exposition and Annual Meeting, SEG*. https://doi.org/10.1190/segam2019-3216008.1.D043S131R001.

Jannane, M., Beydoun, W., Crase, E., Cao, D., Koren, Z., Landa, E., et al. (1989). Wavelengths of earth structures that can be resolved from seismic reflection data. *Geophysics*, 54(7), 906–910. https://doi.org/10.1190/1.1442719

Jian, H., Nedimović, M. R., Canales, J. P., & Lau, K. H. (2021). New insights into the rift to drift transition across the northeastern Nova Scotian margin from wide-angle seismic waveform inversion and reflection imaging. *Journal of Geophysical Research: Solid Earth*, 126(12), e2021JB022201. https://doi.org/10.1029/2021jb022201

Kelemen, P. B., Koga, K., & Shimizu, N. (1997). Geochemistry of gabbro sills in the crust-mantle transition zone of the Oman ophiolite: Implications for the origin of the oceanic lower crust. *Earth and Planetary Science Letters*, 146(3–4), 475–488. https://doi.org/10.1016/S0012-821X(96)00235-X

Kelemen, P. B., Matter, J. M., Teagle, D. A. H., & Coggon, J. A., & Oman Drilling Project Science Team, & Oman Drilling Project Science Team. (2020). Site GT3: Sheeted dike to gabbro transition. The Oman drilling Project science team. In *Proceedings of the Oman drilling Project*. https://doi.org/10.14379/OmanDP.proc.108.2020

Kent, G. M., Harding, A. J., Orcutt, J. A., Detrick, R. S., Mutter, J. C., & Buhl, P. (1994). Uniform accretion of oceanic crust south of the Garrett transform at 14 15' S on the East Pacific Rise. *Journal of Geophysical Research*, 99(B5), 9097–9116. https://doi.org/10.1029/93JB02872

Komatitsch, D., & Martin, R. (2007). An unsplit convolution perfectly matched layer improved at grazing incidence for the seismic wave equation. *Geophysics*, 72(5), SM155-SM167. https://doi.org/10.1190/1.2757586

Korenaga, J., Holbrook, W. S., Kent, G. M., Kelemen, P. B., Detrick, R. S., Larsen, H. C., et al. (2000). Crustal structure of the southeast Greenland margin from joint refraction and reflection seismic tomography. *Journal of Geophysical Research*, 105(B9), 21591–21614. https://doi.org/10.1029/2000JB900188

Lambart, S., Koornneef, J. M., Millet, M. A., Davies, G. R., Cook, M., & Lissenberg, C. J. (2019). Highly heterogeneous depleted mantle recorded in the lower oceanic crust. *Nature Geoscience*, 12(6), 482–486. https://doi.org/10.1038/s41561-019-0368-9

Lin, J., & Morgan, J. P. (1992). The spreading rate dependence of three-dimensional mid-ocean ridge gravity structure. *Geophysical Research Letters*, 19(1), 13–16. https://doi.org/10.1029/91GL03041

MacLeod, C. J., Escartin, J., Banerji, D., Banks, G. J., Gleeson, M., Irving, D. H. B., et al. (2002). Direct geological evidence for oceanic detachment faulting: The Mid-Atlantic Ridge, 15 45' N. *Geology*, 30(10), 879–882. https://doi.org/10.1130/0091-7613(2002)030<0879: dgefod>2.0.co;2

Maia, M., Sichel, S., Briais, A., Brunelli, D., Ligi, M., Ferreira, N., et al. (2016). Extreme mantle uplift and exhumation along a transpressive transform fault. *Nature Geoscience*, 9(8), 619–623. https://doi.org/10.1038/ngeo2759

Marjanović, M., Plessix, R. É., Stopin, A., & Singh, S. C. (2019). Elastic versus acoustic 3-D full waveform inversion at the East Pacific rise 9° 50' N. *Geophysical Journal International*, 216(3), 1497–1506. https://doi.org/10.1093/gji/ggy503

Marjanović, M., Singh, S. C., Gregory, E. P., Grevetmeyer, I., Growe, K., Wang, Z., et al. (2020). Seismic crustal structure and morphotectonic features associated with the Chain Fracture Zone and their role in the evolution of the equatorial Atlantic region. *Journal of Geophysical Research: Solid Earth*, 125(10), e2020JB020275. https://doi.org/10.1029/2020JB020275

Miller, D. J., & Christensen, N. I. (1997). Seismic velocities of lower crustal and upper mantle rocks from the slow-spreading Mid-Atlantic Ridge, south of the Kane Transform Zone (MARK). In *Proc. Ocean Drill. Program* (Vol. 153, pp. 437–456).

Minshull, T. A., Muller, M. R., Robinson, C. J., White, R. S., & Bickle, M. J. (1998). Is the oceanic Moho a serpentinization front? *Geological Society, London, Special Publications*, 148(1), 71–80. https://doi.org/10.1144/GSL.SP.1998.148.01.05

Minshull, T. A., White, R. S., Mutter, J. C., Buhl, P., Detrick, R. S., Williams, C. A., & Morris, E. (1991). Crustal structure at the Blake Spur fracture zone from expanding spread profiles. *Journal of Geophysical Research*, 96(B6), 9955–9984. https://doi.org/10.1029/91JB00431

Morgan, J. P., & Chen, Y. J. (1993). The genesis of oceanic crust: Magma injection, hydrothermal circulation, and crustal flow. *Journal of Geophysical Research*, 98(B4), 6283–6297. https://doi.org/10.1029/92JB02650

Müller, R. D., Sdrolias, M., Gana, C., & Roest, W. R. (2008). Age, spreading rates, and spreading asymmetry of the world's ocean crust. *Geochemistry, Geophysics, Geosystems*, 9(4). https://doi.org/10.1029/2007GC001743

Neves, F. A., & Singh, S. C. (1996). Sensitivity study of seismic reflection/refraction data. *Geophysical Journal International*, 126(2), 470–476. https://doi.org/10.1111/j.1365-246x.1996.tb05303.x

Nguyen, B. D., & McMechan, G. A. (2015). Five ways to avoid storing source wavefield snapshots in 2D elastic prestack reverse time migration. *Geophysics*, 80(1), S1–S18. https://doi.org/10.1190/geo2014-0014.1
Nicolas, A., Reuber, I., & Benn, K. (1988). A new magma chamber model based on structural studies in the Oman ophiolite. *Tectonophysics*, **151**(1-4), 87-105. https://doi.org/10.1016/0040-1951(88)90242-9

Operto, S., Virieux, J., Dessa, J. X., & Pascal, G. (2006). Crustal seismic imaging from multifold ocean bottom seismometer data by frequency domain full waveform tomography: Application to the eastern Nankai trough. *Journal of Geophysical Research*, **111**(B9). https://doi.org/10.1029/2005JB003835

Pica, A., Diet, J. P., & Tarantola, A. (1990). Nonlinear inversion of seismic reflection data in a laterally invariant medium. *Geophysics*, **55**(3), 284-292. https://doi.org/10.1190/1.1442836

Qin, Y., & Singh, S. C. (2017). Detailed seismic velocity of the incoming subducting sediments in the 2004 great Sumatra earthquake rupture zone from full waveform inversion of long offset seismic data. *Geophysical Research Letters*, **44**(7), 3090-3099. https://doi.org/10.1002/2016GL072175

Ritsema, J., & Lekić, V. (2020). Heterogeneity of seismic wave velocity in Earth's mantle. *Annual Review of Earth and Planetary Sciences*, **48**(1), 377-401. https://doi.org/10.1146/annurev-earth-082119-065909

Salisbury, M. H., & Christensen, N. I. (1978). The seismic velocity structure of a traverse through the Bay of Islands ophiolite complex, Newfoundland, an exposure of oceanic crust and upper mantle. *Journal of Geophysical Research*, **83**(B2), 805-817. https://doi.org/10.1029/JB083iB02p00805

Salisbury, M. H., Christensen, N. I., & Wilkens, R. H. (1996). Nature of the layer 2/3 transition from a comparison of laboratory and logging velocities and petrology at the base of Hole 504B. In *Proceedings-ocean drilling program scientific results* (pp. 409-416). National Science Foundation.

Seyler, M., & Bonatti, E. (1997). Regional-scale melt-rock interaction in lherzolitic mantle in the Romanche Fracture Zone (Atlantic Ocean). *Earth and Planetary Science Letters*, **146**(1-2), 273-287. https://doi.org/10.1016/S0012-821X(96)00220-8

Shen, X. (2010). Near-surface velocity estimation by weighted early-arrival waveform inversion. In *SEG International Exposition and Annual Meeting*. SEG. SEG-2010.

Shipp, R. M., & Singh, S. C. (2002). Two-dimensional full wavefield inversion of wide-aperture marine seismic streamer data. *Geophysical Journal International*, **151**(2), 325-344. https://doi.org/10.1046/j.1365-246X.2002.01645.x

Shorttle, O., & MacLennan, J. (2011). Compositional trends of Icelandic basalts: Implications for short-length scale lithological heterogeneity in mantle plumes. *Geochemistry, Geophysics, Geosystems*, **12**(11). https://doi.org/10.1029/2011GC003748

Singh, S. C., Collier, J. S., Harding, A. J., Kent, G. M., & Orcutt, J. A. (1999). Seismic evidence for a hydrothermal layer above the solid roof of the axial magma chamber at the southern East Pacific Rise. *Geology*, **27**(3), 219-222. https://doi.org/10.1130/0091-7613(1999)027%3C0219:SEFAHL%3E2.3.CO;2

Singh, S. C., Crawford, W. C., Carton, H., Seher, T., Combier, V., Cannat, M., et al. (2006). Discovery of a magma chamber and faults beneath a Mid-Atlantic Ridge hydrothermal field. *Nature*, **442**(7106), 1029-1032. https://doi.org/10.1038/nature05105

Spudich, P., & Orcutt, J. (1980). A new look at the seismic velocity structure of the oceanic crust. *Reviews of Geophysics*, **18**(3), 627-645. https://doi.org/10.1029/RG018i003p00627

Spudich, P., & Orcutt, J. (1980b). Petrology and porosity of an oceanic crustal site: Results from wave form modeling of seismic refraction data. *Journal of Geophysical Research*, **85**(B3), 1409-1433. https://doi.org/10.1029/JB085iB03p01409

Swift, S., Reichow, M., Tikku, A., Tominaga, M., & Gilbert, L. (2008). Velocity structure of upper ocean crust at Ocean Drilling Program Site 1256. *Geochemistry, Geophysics, Geosystems*, **9**(10). https://doi.org/10.1029/2008GC002188

Tarantola, A. (1984). Inversion of seismic reflection data in the acoustic approximation. *Geophysics*, **49**(8), 1259-1266. https://doi.org/10.1190/1.1441754

Toomey, D. R., Solomon, S. C., & Purdy, G. M. (1994). Tomographic imaging of the shallow crustal structure of the East Pacific Rise at 9° 30' N. *Journal of Geophysical Research*, **99**(B12), 24135-24157. https://doi.org/10.1029/94JB01942

Tromp, J., Tape, C., & Liu, Q. (2005). Seismic tomography, adjoint methods, time reversal and banana-doughnut kernels. *Geophysical Journal International*, **160**(1), 195-216. https://doi.org/10.1111/j.1365-246X.2004.02453.x

Vaddineni, V. A., Singh, S. C., Grevemeyer, I., Audhkhasi, P., & Papenberg, C. (2021). Evolution of the crustal and upper mantle seismic structure from 0-27 Ma in the equatorial Atlantic Ocean at 2 43' S. *Journal of Geophysical Research: Solid Earth*, **126**(6), e2020JB021390. https://doi.org/10.1029/2020JB021390

Van Avendonk, H. J., Harding, A. J., Orcutt, J. A., & McClain, J. S. (1998). A two-dimensional tomographic study of the Clipperton transform fault. *Journal of Geophysical Research*, **103**(B8), 17885-17899. https://doi.org/10.1029/98JB00904

Vigh, D., & Starr, E. W. (2008). 3D prestack plane-wave, full-waveform inversion. *Geophysics*, **73**(5), VE135-VE144. https://doi.org/10.1190/1.2952623

Virieux, J. (1986). P-SV wave propagation in heterogeneous media: Velocity-stress finite-difference method. *Geophysics*, **51**(4), 889-901. https://doi.org/10.1190/1.1442147

Virieux, J., & Operto, S. (2009). An overview of full-waveform inversion in exploration geophysics. *Geophysics*, **74**(6), WCC1-WCC26. https://doi.org/10.1190/1.3238367

Wang, Z., & Singh, S. C. (2022). Seismic evidence for uniform crustal accretion along slow-spreading ridges in the equatorial Atlantic Ocean. *Nature Communications*, **13**(1), 7809. https://doi.org/10.1038/s41467-022-35459-z

Wang, Z., Singh, S. C., Prigent, C., Gregory, E. P., & Marjanović, M. (2022). Deep hydration and lithospheric thinning at oceanic transform plate boundaries. *Nature Geoscience*, **15**(9), 741-746. https://doi.org/10.1038/s41561-022-01003-3

White, R. S., McKenzie, D., & O'Nions, R. K. (1992). Oceanic crustal thickness from seismic measurements and rare earth element inversions. *Journal of Geophysical Research*, **97**(B13), 19683-19715. https://doi.org/10.1029/92JB01749

Wilkens, R. H., Fryer, G. J., & Karsten, J. (1991). Evolution of porosity and seismic structure of upper oceanic crust: Importance of aspect ratios. *Journal of Geophysical Research*, **96**(B11), 17981-17995. https://doi.org/10.1029/91JB01454

Wilson, D. S., Teagle, D. A., Alt, J. C., Banerjee, N. R., Umino, S., Miyashita, S., et al. (2006). Drilling to gabbro in intact ocean crust. *Science*, **312**(5776), 1016-1020. https://doi.org/10.1126/science.1126090

Xu, M., Zhao, X., & Canales, J. P. (2020). Structural variability within the Kane oceanic core complex from full waveform inversion and reverse time migration of streamer data. *Geophysical Research Letters*, **47**(7), e2020GL087405. https://doi.org/10.1029/2020GL087405

Zhang, X., & Curtis, A. (2020). Variational full-waveform inversion. *Geophysical Journal International*, **222**(1), 406-411. https://doi.org/10.1093/gji/ggaa170