# Link between crustal thickness and Moho transition zone at 9°N East Pacific Rise

Zhikai Wang, Satish C. Singh, J. Pablo Canales

► To cite this version:

Zhikai Wang, Satish C. Singh, J. Pablo Canales. Link between crustal thickness and Moho transition zone at 9°N East Pacific Rise. Earth and Planetary Science Letters, 2025, 658, 396, pp. 28-41. 10.1016/j.epsl.2025.119309. insu-05080897

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Submitted on 23 May 2025

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# Link between crustal thickness and Moho transition zone at 9°N East Pacific Rise

Zhikai Wang<sup>a,b,*</sup> <sup>Id</sup>, Satish C. Singh<sup>a</sup> <sup>Id</sup>, J. Pablo Canales<sup>c</sup> <sup>Id</sup>

## ARTICLE INFO

Editor: Dr H Thybo

Keywords:

Mid-ocean ridge

East Pacific Rise

Oceanic crust

Crustal thickness

Moho transition zone

Full waveform inversion

## ABSTRACT

Oceanic crust is formed from basaltic melt produced by decompression melting due to mantle upwelling at mid-ocean ridges. This crust is separated from the underlying mantle either by a sharp Mohorovičić (Moho) discontinuity or a thick Moho transition zone (MTZ). Determining the relationship between the oceanic crustal structure and the MTZ is critical for understanding the crustal accretion processes at mid-ocean ridges. However, this relationship remains elusive due to the lack of high-resolution velocity model of the oceanic crust and MTZ. Here, we present result from the application of full waveform inversion to wide-angle seismic data acquired over a young oceanic crust near the 9°N East Pacific Rise, allowing us to obtain the crustal and MTZ thicknesses along a ~70 km-long segment. We find that the crustal thickness and the MTZ thickness vary along the segment and they are inversely correlated, although the total cumulative thickness does not vary much along the profile. These variations could be attributed to the different melt migration efficiency or the variations in mantle thermal/chemical structure, indicating mantle heterogeneity along the ridge.

## 1. Introduction

Oceanic crust is formed at mid-ocean ridges (MORs) from basaltic melt derived from decompression of upwelling mantle as two lithospheric plates move apart (Cann, 1970; Searle, 2013). In a fast- and intermediate-spreading environment, the melt rises towards the surface and accumulates within an axial magma chamber (AMC) at mid-crustal depths (Detrick et al., 1987; Mutter et al., 1988; Arnulf et al., 2014a). A portion of the accumulated melt erupts to form lava flows on the seafloor and dike underneath, making up the upper crust (Cann, 1970; Sinton and Detrick, 1992). The remainder of the melt crystallises within the AMC, forming the gabbroic lower crust (Cann, 1970; Sinton and Detrick, 1992; Guo et al., 2022). The crust is separated from the underlying mantle by the Mohorovičić (Moho) transition zone (MTZ), across which significant variations in physical properties and chemical compositions are expected (Cann, 1970; White et al., 1992; Christeson et al., 2019). Therefore, determining the relationship between the oceanic crustal structure and the MTZ is critical for understanding the crustal accretion processes at MORs.

Drilling is the most direct approach for studying the structure of oceanic crust, but drilling down to lower crust and MTZ is challenging

(Teagle and Ildefonse, 2011). Surface exposures of the Oman and Bay of Islands ophiolites, which are thought to be formed at ancient intermediate- to fast-spreading centres, have allowed direct observations of the MTZ (Casey and Karson, 1981; Boudier and Nicolas, 1995). Structural mapping of these ophiolites indicates that the oceanic MTZ is primarily composed of dunites with gabbroic sills and lenses, marking a gradually downward change from layered gabbro in the lower crust to ultramafic mantle consisting dominantly of harzburgites (Karson et al., 1984; Benn et al., 1988; Boudier and Nicolas, 1995). The MTZ can be as thin as 5–10 m (sharp boundary) or as thick as 1–3 km (Karson et al., 1984; Benn et al., 1988; Juteau et al., 1988; Boudier and Nicolas, 1995; Jousselin et al., 2021), where the thickness of gabbroic sills and lenses can reach hundreds of meters in a thick MTZ (Karson et al., 1984; Benn et al., 1988; Boudier and Nicolas, 1995). The strong magmatic flow structures within the MTZ indicate that these gabbroic sills were emplaced at the ridge axis, implying that a large amount of melt was trapped within the MTZ during crustal accretion (Boudier and Nicolas, 1995). However, in the absence of drilling results (Teagle and Ildefonse, 2011), we do not have any *in situ* information about the MTZ.

Traveltime modelling and tomography of wide-angle seismic data are commonly used to determine the crustal and upper mantle structures

<footnote><sup>a</sup> Institut de Physique du Globe de Paris, Université Paris Cité, 1 Rue Jussieu, 75005 Paris, France</footnote>

<footnote><sup>b</sup> Presently at School of Ocean and Earth Science, University of Southampton, European Way, SO14 3ZH Southampton, UK</footnote>

<footnote><sup>c</sup> Department of Geology and Geophysics, Woods Hole Oceanographic Institution, Woods Hole MA 02543, USA</footnote>

<footnote>* Corresponding author.</footnote>

<footnote>E-mail address: zw1m23@soton.ac.uk (Z. Wang).</footnote>
and crustal thickness, where the Moho is assumed as a sharp boundary. These results have revealed that the magmatic crust consists of two layers: an upper crust characterized by low P-wave velocities (Vp: 3.0–6.5 km/s) but high velocity gradients and a lower crust exhibiting high Vp (6.5–7.1 km/s) and significantly reduced velocity gradient (White et al., 1992; Christeson et al., 2019). The mantle underneath has a Vp>7.7 km/s and consists primarily of peridotite (White et al., 1992; Christeson et al., 2019). Although the velocity in the upper crust is well constrained by the crustal refraction arrivals (Pg), the velocity in the lower crust, especially the lower portion of the lower crust, is determined using the wide-angle reflections from the Moho (PmP), resulting in a trade-off between the velocity in the lower crust and the Moho depth (Van Avendonk et al., 2004; Vaddineni et al., 2021). Furthermore, traveltime data alone cannot distinguish between a sharp Moho and a thick MTZ (Ingale and Singh, 2021), hence it is difficult to determine the relationship between the crustal structure and the MTZ.

Multi-channel seismic (MCS) data provide reflection images of the oceanic Moho formed at fast- and intermediate-spreading ridges with seismic characters ranging from impulsive, shingled and diffusive (Kent et al., 1994; Nedimović et al., 2005; Aghaei et al., 2014). The impulsive and shingled Moho are characterized by a single-phase reflection while the diffusive Moho shows sub-horizontally multi-phase reflection events (Kent et al., 1994; Barth and Mutter, 1996; Aghaei et al., 2014). Seismic waveform modelling suggests that the impulsive and shingled Moho reflections are probably produced by a thin MTZ and the diffusive Moho indicates a relatively thick MTZ (Brocher et al., 1985; Collins et al., 1986). Nedimović et al. (2005) suggested that the multi-phase Moho reflection events might be caused by the frozen melt lenses within a thick MTZ. However, MCS data provide the Moho structure in two-way traveltime, which needs to be converted to depth using a velocity model to get the crustal thickness. Due to the lack of accurate velocity model of the subsurface, the uncertainty in the inferred Moho depth can be hundreds metres (Barth and Mutter, 1996; Aghaei et al., 2014), and even >1 km in some areas (Barth and Mutter, 1996). Therefore, our understanding of MTZ from seismic methods remain elusive.

Here, we present result from the application of two-dimensional (2-D) full waveform inversion (FWI) (Shipp and Singh, 2002; Wang et al., 2025) to wide-angle seismic data to constrain the Vp of the crust and MTZ along a segment near the 9°N East Pacific Rise (EPR). Theoretically, FWI can provide high-resolution velocity model of the subsurface at a vertical resolution of half a wavelength (Virieux and Operto, 2009), i.e. hundreds meter, important for studying the fine-scale structures of oceanic crust and MTZ (Guo et al., 2022; Dhabaria and Singh, 2024a; Wang et al., 2025).

## 2. Data and methods

Controlled-source wide-angle seismic data were acquired at the fast-spreading (11 cm/yr) EPR between 8°15'N and 10°05'N during the 1997 Undershoot Seismic Experiment (Toomey et al., 1997). Although the Undershoot Seismic Experiment was performed covering the whole 9°N EPR segment on both ridge flanks (Toomey et al., 1997), here we used only six ocean bottom instruments deployed at dominantly ~8–14 km intervals along a 92 km-long profile on the eastern flank of the EPR (Fig. 1). This seismic profile sampled the ~300 kyr-old crust between the Clipperton transform fault (TF; first-order discontinuity) and the 9°03'N overlapping spreading centre (OSC; second-order discontinuity). The EPR between the 9°03'N OSC and the Clipperton TF is further offset by third-order discontinuities at 9°12'N, 9°20'N, 9°37'N, 9°51.5'N and 9°58'N, respectively (black rectangles in Fig. 1) (White et al., 2006; Aghaei et al., 2014). The choice of seismic profile is mainly based on the number of instruments along a straight line in order to be able to apply 2-D FWI. The source used in data acquisition was an airgun array with a total volume of 8503 in³, towed at 10 m depth and fired at ~460 m interval (Toomey et al., 1997).

We inverted the pressure data recorded by ocean bottom hydro-

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**Fig. 1.** Bathymetry map of the study area. Red curves show the East Pacific Rise between the Clipperton transform fault (TF) and the 9°03'N overlapping spreading centre (OSC). The black rectangles show the locations of third-order discontinuities at 9°12'N, 9°20'N, 9°37'N, 9°51.5'N and 9°58'N from south to north, respectively (White et al., 2006; Aghaei et al., 2014). The black line indicates the seismic profile. Brown and purple triangles represent the locations of ocean bottom hydrophones (OBHs) and ocean bottom seismometers (OBSs), respectively. The blue box in the inset shows the location of the study area. The black scale shows the distance along the profile.

phones (OBHs) and the vertical component data of ocean bottom seismometers (OBSs) to constrain the subsurface velocity along the profile. We band-pass filtered the data between 3 and 30 Hz, and applied a predictive gapped deconvolution with minimum and maximum lags of 0.14 s and 0.35 s to suppress the seismic bubbles (Guo et al., 2022). The deconvolved data show clear crustal and mantle refractions (Pg and Pn) and PmP arrivals (Fig. 2). In this study, we performed 2-D FWI of seismic waves that have propagated in three-dimensional (3-D) space. Therefore, we transformed the seismic data after deconvolution from 3-D to 2-D by multiplying the amplitude of the data by $\sqrt{t}$ (where t is the traveltime) and convolving the seismic data with $1/\sqrt{t}$ (Pica et al., 1990). A 1.0 s-wide time windowing was applied to extract the Pg, PmP and Pn arrivals between 6 and 60 km offsets (Fig. 2).

Starting from an initial velocity model, FWI updates the velocity of subsurface by gradually reducing the mismatch between the observed and modelled seismic data. The initial Vp model for FWI (Fig. 3A) was built using a crustal model from traveltime tomography (Canales et al., 2003), which was extended down by 5 km by hanging a one-dimensional (1-D) mantle velocity below the tomographically estimated Moho where the velocity increases linearly from 7.9 to 8.2 km/s. The velocity around the tomographically constrained Moho interface (thick black curve in Fig. 3A) was smoothed to avoid a sharp boundary between the crust and mantle, leading to a starting MTZ thickness of ~1 km characterized by high velocity gradient (Fig. 3B). We only inverted the Vp of the subsurface. The S-wave velocity and density were calculated from the initial Vp model using the empirical relations (Brocher, 2005) and were fixed in FWI. The synthetic seismic data were modelled using a 2-D time-domain staggered-grid finite-difference scheme (Text S1) (Levander, 1988). The source wavelets used in synthetic modelling were estimated by stacking the aligned near-offset water arrivals (Text S2 and Fig. S1) (Jian et al., 2021; Dhabaria and Singh, 2024a; Wang et al., 2025).

FWI of PmP arrivals is highly nonlinear (Guo et al., 2020; Dhabaria
<figure><img src="image_5.png" /><figcaption>**Fig. 2.** Time window overlapping on seismic section after deconvolution. (A) for OBH25 and (B) for OBS64. The oranges boxes show the 1.0 s-wide time window used in the FWI of 3–10 Hz data. The top of the time window is 0.1 s prior the picked first arrival traveltime. The waveforms of the crustal and mantle refractions (Pg and Pn) and the Moho reflections (PmP) are included in the time window. The traveltime of seismic data is reduced using a reduction velocity of 7.0 km/s.</figcaption></figure>

and Singh, 2024b) and requires a misfit function that can correctly model the nonlinear part of the critically reflected PmP arrivals and also can handle the triplication between Pg, PmP and Pn arrivals around critical angles. In order to overcome this non-linearity, the FWI workflow was implemented in two stages. In the first stage, we applied a trace-normalised FWI which compares the seismic data after trace-by-trace normalisation to primarily fit seismic traveltime (or phase) information (Text S3) (Liu et al., 2016; Tao et al., 2017). The trace-normalised FWI allows to decouple the complex waveforms associated with the critically reflected PmP arrivals and the triplication and helps to recover the velocity of the MTZ (Tao et al., 2017; Dhabaria and Singh, 2024a). We first inverted the seismic data between 3 and 5 Hz (Fig. S2A), and the output was used as starting model for the inversion of 3–10 Hz data (Fig. S2B). Taking the inverted result of trace-normalised FWI as starting model, we performed seismic-gather-normalised FWI for the 3–10 Hz data to further improve the velocity model (Text S3) (Louboutin et al., 2017). The seismic-gather-normalised FWI compares the modelled and observed data normalised by the L₂-norm of the whole seismic gather (Louboutin et al., 2017), which can further improve the velocity model as it tries to fit both seismic amplitude and phase information.

In each stage of FWI, the misfit function gradually decreases with increasing iteration (Figs. S3A,B). The misfits for most seismic gathers were reduced to 25–50% after FWI (Fig. S3C). Compared with the initial Vp model (Fig. 3A), the inverted Vp model from FWI shows fine-scale structures in the crust (Fig. 4A). We compared the observed data and the modelled data before and after FWI to assess the improvement in data fitting for two criteria used during the inversion: (1) normalising the seismic data trace by trace to mainly compare the traveltime information (Fig. 5 and S4) and (2) normalising the seismic data using the L₂-norm of the whole seismic gather to compare the seismic amplitude with offset (Fig. 6 and S5), related to the FWI misfit functions in the first and

second stages, respectively. The synthetic data after FWI match the observed data well as compared to those modelled using the initial model, demonstrating that the inverted velocity model from FWI is closer to the real subsurface velocity structure than the initial model.

We conducted checkerboard tests (Text S4) to assess the resolution of the FWI velocity model. The results suggest that FWI can resolve minimum structures of 0.5 × 10 km size (vertical × horizontal) with ±5% velocity anomaly between 10 and 80 km horizontal distance (Fig. 7). We also performed synthetic tests (Text S5) to assess the resolvability of the FWI method for a thin or thick MTZ. These synthetic tests indicate that FWI can recover a MTZ as thin as 500 m or as thick as 3.0 km between 10 and 80 km horizontal distance (Fig. 8).

## 3. Results

### 3.1. Crustal structure

The final inverted velocity model (Fig. 4A) and the velocity anomaly (difference between inverted and initial velocity models) (Fig. 4C) show a lot more structures in the crust and MTZ, but here we only discuss the large-scale structures that have been well resolved. Compared with the initial model (Fig. 3A), the upper crust in the inverted model has lower velocity and shows more lateral velocity variations (Fig. 4A and C). The upper crust is characterized by high vertical velocity gradients, where the Vp increases rapidly from 3.0±0.1 km/s at the basement to 6.5 km/s at 1.8±0.2 km depth below the seafloor (Fig. 4A, B and 9). The inverted velocity model reveals a heterogeneous lower crust, where alternate high-and-low-velocity layers are observed (Figs. 4A and 9). We used the contour of vertical velocity gradient of 0 s⁻¹ to represent the boundaries between the high- and low-velocity layers (Fig. 4B). The thickness of these layers varies from ~500 m to ~1 km, and the maximum velocity reduction within the low-velocity layers is ~500 m/s. In contrast, the initial model has a dominantly positive vertical velocity gradient and does not show large velocity reduction in the lower crust (Figs. 3 and 9). Therefore, the high-and-low-velocity layering in the lower crust in the inverted model is not inherited from the initial model and is required by the seismic data.

At the base of the model at 8.0–9.5 km depth, a positive high velocity gradient zone is observed (Fig. 4B), separating the typical lower crustal velocity above with the mantle velocity below. A positive high velocity gradient zone is present around the tomographic Moho in the initial model (Fig. 3B) but is thinner than that in the inverted model. We picked the depth of the top of the high velocity gradient zone in the inverted model marked by a vertical velocity gradient of 0 s⁻¹ and smoothed it over a horizontal distance of 10 km, which is interpreted as the base of the crust (red dashed curves in Fig. 4A,B and D). This base of the crust shallows from a depth of ~9.5 km between 10 and 20 km distance to ~8.0 km between 45 and 55 km distance, and it lies between 8.2 and 8.5 km depth further north. The mean velocity at the base of the crust is ~7.0±0.2 km/s, consistent with the global average velocity (7.1±0.1 km/s) at the base of crust formed at fast-spreading ridges (Christeson et al., 2019).

The interpreted crustal base derived from the FWI is shallower than the tomographic Moho from Canales et al. (2003) along the entire profile (Fig. 4A). The crustal thickness varies between 5.1 and 6.5 km (red curve in Fig. 4E) with an average thickness of ~5.6 km, thinner than the average crustal thickness (~6.8 km) obtained from the traveltime tomography (Canales et al., 2003), but close to that (~5.8 km) estimated from the MCS studies in the neighbouring region (Aghaei et al., 2014). The crust is thicker south of the 30–40 km horizontal distance at ~9°36′–9°41′N than to the north, and the thickest crust is observed between 10 and 20 km horizontal distance (red curve in Fig. 4E), which is consistent with the traveltime tomography study (Canales et al., 2003). The thinnest crust is observed at 45–55 km horizontal distance, and the crust gradually thickens by ~500–700 m further north (red curve in Fig. 4E).
<figure><img src="image_6.png" /><figcaption>**Fig. 3.** (A) Initial P-wave velocity (Vp) model for FWI. The thick black curve is the tomographic Moho from Canales et al. (2003). The brown and purple triangles show the locations of OBHs and OBSs, respectively. The 4.5, 5.5, 6.5 and 7.85 km/s velocity contours are shown as black dashed curves from top to bottom. (B) Vertical velocity gradient of the initial Vp. The black dashed curves are the 0 s⁻¹ velocity gradient contour. The vertical to horizontal exaggeration is 3:1.</figcaption></figure>

## 3.2. Moho transition zone (MTZ)

FWI does not provide a sharp boundary for the Moho, but an increase in velocity over a certain depth range, which we interpret as the MTZ. The base of the crust with zero velocity gradient marks the top of the MTZ. We used two approaches to define the bottom of the MTZ: (I) 7.85 km/s velocity contour, the global average velocity at the top of the mantle (<7.5 Myr) formed at fast-spreading ridges (Christeson et al., 2019) and (II) the base of the high velocity gradient zone. If we pick the depth of the 7.85 km/s velocity contour (purple dashed curves in Fig. 4A, B and D) the thickness of the MTZ would be between 1.1 and 2.4 km (blue dashed curve in Fig. 4E). If we take the base of the large positive velocity gradient zone as the bottom of the MTZ (purple solid curves in Fig. 4A and B), the thickness of the MTZ would be between 1.6 and 2.9 km (blue solid curve in Fig. 4E) where the average mantle velocity is ~7.97±0.13 km/s. In both cases, the MTZ is relatively thin south of the 30–40 km horizontal distance (Fig. 4E). The thickness of the MTZ shows a negative correlation with the crustal thickness along the profile, i.e., where the MTZ is thick the crust is thin, and vice versa (Fig. 4E and 9).

## 4. Discussion

Our results show (I) the presence of layered structures in the oceanic lower crust and (II) an inverse correlation between the thicknesses of crust and MTZ, where the crust is thin in the north and thick in the south whereas the MTZ is thick in the north and thin in the south.

Seismic reflection studies of the 9°N EPR have shown the presence of axial melt lens (AML) at ~1.4–1.9 km depth in the mid-crust (Detrick et al., 1987; Kent et al., 1993) whereas tomographic studies indicate the presence of low velocity zone down to 6–7 km depth below the seafloor (Dunn et al., 2000), indicating the existence of partial melt in the lower

crust. Furthermore, Marjanović et al. (2014) and Arnulf et al. (2014b) show the presence of secondary melt sills within 1.65 km depth below the AML. Studies of the Oman ophiolite suggest that the melt can intrude and crystalize at different depths in the lower crust (Boudier et al., 1996; Kelemen et al., 1997). The observed alternate high-and-low-velocity layering in the lower crust could be due to melt of different compositions injected and crystallised at different depths within the lower crust (Fig. 10; e.g. Guo et al. (2022)). The gabbroic rocks drilled from the Hess Deep in the equatorial Pacific are mainly composed of olivine, clino-pyroxene and plagioclase (Carlson and Jay Miller, 2004; Lissenberg et al., 2013). A small increase (by 5%) of the olivine content can lead up to 600 m/s increase in Vp of the gabbroic rocks (Carlson and Jay Miller, 2004; Guo et al., 2022). Therefore, the low-velocity layers within the lower crust could be formed by melt with relatively low olivine concentration while the high-velocity layers could represent olivine-rich gabbroic sills. This interpretation supports the ‘sheeted sill’ model (Boudier et al., 1996; Kelemen et al., 1997) where *in situ* melt intrusion and crystallization form the oceanic lower crust. Moreover, the off-axis melt sills are observed up to a distance of ~12 km from the ridge crest (Canales et al., 2012; Han et al., 2014; Aghaei et al., 2017) and could form gabbroic sills with different compositions from those formed at the ridge axis, contributing to the formation of a heterogeneous lower crust.

Some early studies using 1-D velocity analysis found that the MTZ at 12°N EPR is ~1.0 km thick within 50 km off-axis distance (Lewis and Garmany, 1982) and the MTZ at 9°35'N EPR is ~1.7 km thick at 10 km off-axis distance (Vera et al., 1990). A MCS study from the intermediate-spreading Juan de Fuca Ridge observed that the MTZ could be up to 2.0 km thick (Nedimović et al., 2005). These estimates fall in the ranges of MTZ thickness obtained using FWI, but our results provide a 2-D view continuous over 70 km distance along the profile and its relationship to the crustal structure.

The along-strike variations in the MTZ thickness could be due to the
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**Fig. 4.** FWI results. (A) Inverted P-wave velocity (Vp) model. The thick black curve is the tomographic Moho from Canales et al. (2003). The red dashed curve is the interpreted crustal base corresponding to the top of the large positive velocity gradient zone beneath the crust. The dashed purple curve is the bottom of the MTZ interpreted using a smooth version of the 7.85 km/s velocity contour (Christeson et al., 2019). The solid purple curve is the bottom of the MTZ interpreted using the base of the large positive velocity gradient zone. The 4.5, 5.5, 6.5 and 7.85 km/s velocity contours are shown as black dashed curves from top to bottom. The brown and purple triangles show the locations of OBHs and OBSs, respectively. The vertical to horizontal exaggeration is 3:1. (B) Vertical velocity gradient of the inverted Vp. The black dashed curves are the $0 \text{ s}^{-1}$ velocity gradient contour. (C) Velocity anomaly: difference between inverted and initial velocity models. (D) Zoomed-in plot of inverted velocity around the MTZ. The white dashed, black dashed and black solid curves are the 7.0, 7.5 and 8.0 km/s velocity contours, respectively. (E) The crustal thickness (in red) and the MTZ thickness (in blue) variations along the profile. The blue dashed and solid curves are the MTZ thickness calculated using a smooth version of the 7.85 km/s velocity contour (purple dashed curves in A) and using the base of the large positive velocity gradient zone (purple solid curves in A) as the bottom of the MTZ, respectively.

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**Fig. 5.** Comparisons of trace-by-trace normalised seismic data (A) before and (B) after FWI for OBH25. The observed data (in red) is filtered to 3–10 Hz and the modelled data (in black) are calculated using the 3–10 Hz source wavelet shown in Fig. S1C. The seismic data are plotted at every five traces. Traveltime (T) of the seismic data is reduced using a reduction velocity of 7.0 km/s. X is the offset.

<figure><img src="image_13.png" /></figure>

**Fig. 6.** Comparisons of seismic-gather-normalised data (A) before and (B) after FWI for OBH25. The seismic data are normalised by the $L_2$-norm of the whole seismic gather after windowing. The observed data (in red) is filtered to 3–10 Hz and the modelled data (in black) are calculated using the 3–10 Hz source wavelet shown in Fig. S1C. The seismic data are plotted at every five traces. Traveltime (T) of the seismic data is reduced using a reduction velocity of 7.0 km/s. For better visibility, a scalar weighting factor $(1 + 0.1 \times X)$ was applied for each trace to enhance the amplitude at large offsets, where X is the offset.

changes in thermal structure along the ridge. Thermal structure plays an important role in controlling the vertical depth of melt introduction and crystallization at fast-spreading ridges (Maclennan et al., 2004). For a relatively hot ridge segment, melt will migrate rapidly, pool and crystallize at shallower depth in the crust with little melt accumulate within

the MTZ. In contrast, for a relatively cold ridge segment, some melt could accumulate at deeper depths in the lower crust or at Moho depth, forming a thick MTZ. The along-strike variations in the MTZ thickness could also reflect changes in the efficiency of melt migration. A thin MTZ would indicate a rapid percolation of melt from the upwelling mantle to the accreting crust. The formation of a thick MTZ could be due to less efficient melt extraction from mantle to crust leading to the accumulation and crystallization of a large amount of melt around the Moho depth (Fig. 10). Small amount of melt may crystallize in the thin MTZ as well in the study region, as gabbroic sills and lenses were observed in the hundreds-meter thick MTZ in ophilotes (Benn et al., 1988; Boudier and Nicolas, 1995). The along-axis change in melt migration efficiency could be due to the different viscosities of the compositionally different melt between ridge segments, which could be attributed to the heterogeneity of the parental mantle source. The compositions of lava collected along the northern EPR between 5°30′N and 14°30′N show significant geochemical diversity between third-order ridge segments (Langmuir et al., 1986). Chemical variations between ridge segments bounded primarily by third-order discontinuities are also observed in the study of major and minor elements of volcanic glass samples collected along southern EPR between 13 and 23°S, representing melting of mantle containing discrete heterogeneities (Sinton et al., 1991). Therefore, we suggest that the transition between relatively thick and thin MTZ at the horizontal distance of 30–40 km at ~9°35′–9°41′N could be associated with the third-order tectonic (and chemical) discontinuity at 9°37′N (Fig. 1).

These interpretations are supported by the presence of melt around Moho depth beneath the 9–10°N EPR revealed by the seafloor compliance (Crawford et al., 1999; Crawford and Webb, 2002). Our interpretations are further supported by the negative correlation between the thicknesses of the crust and MTZ. A relatively thick MTZ underlying a relatively thin crust suggests that a significant part of melt was crystallized in the MTZ. However, the total cumulative thickness of the crust and MTZ does not vary much along the profile, albeit the total melt supply from the mantle to crust might be uniform along the entire ridge segment at a fast-spreading rate (Parmentier and Morgan, 1990). Similarly, Nicolas et al. (1996) found that the thin lower crust in the Oman ophiolite is generally associated with a thick MTZ while the thick lower crust is associated with a thin MTZ. These observations indicate that there is an anti-correlation between the ophiolite's crustal thickness and MTZ thickness, assuming the combined thickness of the extrusive basalt and sheeted dike is constant (Nicolas et al., 1996). The extensive presence of thick gabbro sills observed in the relatively thick MTZ in the Oman ophiolite demonstrate that a large amount of magma has ponded within the MTZ (Boudier and Nicolas, 1995), supporting our interpretation.

Along our profile, the change from a relatively thin to thick MTZ occurs over a short distance of ~10 km (Fig. 4E). A similar pattern has been observed in the Oman ophiolite where the transition from a thin to thick MTZ occurs over <5 km distance (Jousselin and Nicolas, 2000). The seismic reflection study at 9°N EPR (Aghaei et al., 2014) also found that the character of the Moho reflection varies over 3–4 km spatial distance. Given different lateral resolutions of these methods, these observations indicate that the thermal structure and/or melt migration efficiency can vary quickly along the axis at fast-spreading ridge.
<figure><img src="image_14.png" /><figcaption>**Fig. 7.** Checkerboard test using 0.5 × 10 km (vertical × horizontal) checkerboard pattern. Panels on the top and bottom show the input checkerboard pattern and the recovered anomaly, respectively. The maximum velocity perturbation is 5%. The black curves are the ±2.5% contours. The brown and purple triangles show the locations of OBHs and OBSs, respectively.</figcaption></figure>

Laterally abrupt changes in the thermal structure and melt migration efficiency will influence ridge segmentation, possibly governing the distributions of third-order ridge discontinuities (Aghaei et al., 2014).

The average crustal thickness estimated from MCS data is ~5.8 km in the 9°N EPR region (Aghaei et al., 2014). However, seismic refraction study in the neighbouring region suggests ~1 km thicker crust (Canales et al., 2003). This indicates a discrepancy in the oceanic crustal thickness obtained using seismic refraction and reflection methods, though these study areas are not the same. The Moho depths estimated from reflection and refraction studies appear to have good consistency at some regions close to subduction trenches in the Pacific Ocean (Ivandic et al., 2008; Kodaira et al., 2014). However, in these studies, the Moho depths estimated from OBS data has large uncertainties of the order of 1 km. Dhabaria and Singh (2024a) applied FWI to OBS data to constrain the crustal structure along a 70-Ma segment in the equatorial Atlantic Ocean, and their result indicates the crustal thickness has been ~1.3 km over-estimated in the previous traveltime tomography studies (Growe et al., 2021; Wang and Singh, 2022). As FWI of wide-angle seismic data can provide fine-scale velocity of the crust and upper mantle and constrain the thickness of the MTZ (Jian et al., 2021; Dhabaria and Singh, 2024a; Wang et al., 2025), it can reconcile the discrepancy between the seismic reflection and the refraction methods.

The quality of the FWI result is affected by the noise within the seismic data (Liu et al., 2016), the equation used for wave propagation in the subsurface (Virieux and Operto, 2009) and the aliasing effects associated with the sparsity of receivers and shots (Brenders and Pratt, 2007). The seismic data used in this study have high quality where clear Pg, PmP and Pn arrivals can be identified. Meanwhile, a proper pre-conditioning of the FWI gradient, such as the low-pass wavenumber filtering used here, can effectively mitigate the noise and spatial aliasing effects by removing the high-wavenumber artifacts (Kamei et al., 2013; Jian et al., 2021; Dhabaria and Singh, 2024a). We simulated the synthetic waveform by solving an elastic wave equation, which can accurately model both the transmitted P-wave and any mode-conversions in the subsurface (Levander, 1988; Shipp and Singh, 2002). The azimuthal seismic anisotropy of the oceanic crust (Dunn and Toomey, 2001) is not considered in the FWI because we implemented FWI in 2-D space. Applying 3-D FWI (Boddupalli et al., 2021) involving anisotropy could provide better inversion result, and this will be performed in the future when high-quality 3-D wide-angle seismic data are collected from the study area. Collectively, our results demonstrate that FWI is a powerful tool for understanding the structures of crust and MTZ and crustal accretion processes at MORs.
<figure><img src="image_15.png" /></figure>

<figure><img src="image_16.png" /></figure>

<figure><img src="image_17.png" /></figure>

<figure><img src="image_18.png" /></figure>

(caption on next page)
**Fig. 8.** Synthetic tests for the recovery of a (A) 0.5 km, (B) 1.0 km, (C) 2.0 km and (D) 3.0 km thick Moho transition zone (MTZ). Only the portion of the model around MTZ is shown, and the full models are shown in Fig. S7. The synthetic true model for seismic data generation is modified by inserting a 0.5, 1.0, 2.0 or 3.0 km thick MTZ into the final model from FWI of field data. The velocity of the inserted MTZ increases linearly with depth from 7.0 to 7.85 km/s. The 1-D profiles overlapping the synthetic true model show vertical velocity profiles every 5 km between 5 and 80 km horizontal distance. The starting model for synthetic test is the same as that used for FWI of field data as shown in Fig. 3A. The 1-D profiles overlapping the starting model compare the synthetic true (solid curves) and starting (dashed curves) velocities every 5 km between 5 and 80 km horizontal distance. The 1-D profiles overlapping the inverted model compare the synthetic true (solid curves) and inverted (dashed curves) velocities every 5 km between 5 and 80 km horizontal distance. The red and magenta curves in all panels represent the top and bottom of the inserted MTZ, respectively.

<figure><img src="image_19.png" /></figure>

**Fig. 9.** Comparisons of 1-D velocity-depth profiles. (A) Comparison of the initial and inverted velocity profiles averaged between 16 and 24 km horizontal distance where the crust is thick and the MTZ is thin. (B) Comparison of the initial and final inverted velocity profiles averaged between 46 and 54 km horizontal distance where the crust is thin and the MTZ is thick. The black dashed lines show the depths of upper and lower crustal boundary. The red and purple dashed lines represent the top of the MTZ and the MTZ bottom defined by 7.85 km/s velocity contour, respectively.

<figure><img src="image_20.png" /></figure>

**Fig. 10.** Schematic diagram showing structures of oceanic crust and Moho transition zone (MTZ). The oceanic crust is separated into an upper crust (~1.8 km thick) and a layered lower crust. The dark brown blocks in the lower crust refer to the low-velocity layers from FWI. The layered lower crust indicates the oceanic lower crust is formed by *in situ* melt injection and crystallization at different depths. The thickness of MTZ is inversely correlated with the crustal thickness. The red horizontal elongated ellipsoids represent the frozen gabbro sills, which is accumulated and crystalized during its ascending from the mantle to the crust.

## 5. Conclusions

We have obtained high-resolution velocity model for the oceanic crust and MTZ formed along the fast-spreading EPR at 9°N by applying FWI to wide-angle seismic data recorded by ocean bottom instruments.

The FWI model shows alternate high-and-low-velocity layers in the lower crust, indicating *in situ* melt injection and crystallization at different depths within the lower crust. The Moho beneath the crust is not sharp but is a transition zone, and its thickness is inversely related with the crustal thickness. The variations in thicknesses of crust and MTZ could be due to changes in mantle thermal structure along the ridge segment, where melt will accumulate at shallower depth within the crust for relatively hot segments while a significant amount of melt could freeze at Moho depth for relatively cold segments thickening the MTZ. The variations in crustal thickness and MTZ thickness could also reflect different efficiencies of melt migration through the MTZ, where the thin MTZ is due to rapid migration of melt from the mantle to the crust while the thick MTZ is due to less efficient melt migration leading to a significant part of melt crystalizing at the Moho depth. The changes in melt migration efficiency along axis results from the compositional difference of melt produced by a chemically heterogeneous mantle at the 9°N EPR.

## Data availability

Seismic data are available at the Institut de Physique du Globe de Paris Research Collection (https://dataverse.ipgp.fr/privateurl.xhtml?token=6eca5518-ffd6-4737-a016-32f5d28f5fe4).

## CRediT authorship contribution statement

**Zhkai Wang:** Writing – original draft, Visualization, Methodology, Investigation, Formal analysis, Funding acquisition, Software, Writing – review & editing.
**Satish C. Singh:** Writing – review & editing, Super-vision, Software, Project administration, Funding acquisition, Conceptualization.
**J. Pablo Canales:** Data curation, Writing – review & editing.
Declaration of competing interest

The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

Acknowledgements

We thank the captain, crew and scientific team of the R/V Maurice Ewing Leg 97-08 for the data collection. Acquisition of the data used in this study was supported by NSF Award OCE-9634132. Z. Wang's contributions to this work were supported by the postdoctoral fellowship from the GPX project of Institut de Physique du Globe de Paris (IPGP) and by the Royal Society-Newton International Fellowships 2022 (NIF \R1\221849). J.P. Canales' contributions to this study were supported by NSF Award OCE-0118383 and by the Independent Research & Development Program at WHOI. Results presented in this paper were performed either on the IPGP S-CAPAD platform or on the IRIDIS High Performance Computing Facility of the University of Southampton, and we acknowledge the associated support services in the completion of this work. Some of the idea about Moho developed during the preparation of ERC Advanced Grant 101141564 MohoLAB. We thank Dr. A. Stracke and other two anonymous reviewers and the editor Dr. Hans Thybo for their constructive comments that helped improve the manuscript. IPGP contribution number 4297.

Supplementary materials

Supplementary material associated with this article can be found, in the online version, at doi:10.1016/j.epsl.2025.119309.

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